Episodic euxinia in the Changhsingian (late Permian) of South China: Evidence from framboidal pyrite and geochemical data

Sedimentary Geology 319 (2015) 78–97

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Episodic euxinia in the Changhsingian (late Permian) of South China: Evidence from framboidal pyrite and geochemical data Hengye Wei a,b,⁎, Thomas J. Algeo b,⁎⁎, Hao Yu c, Jiangguo Wang c, Chuan Guo c, Guo Shi a a b c

College of Earth Science, East China Institute of Technology, Nanchang, Jiangxi Province, 330013, China Department of Geology, University of Cincinnati, Cincinnati, OH 45221, USA Key Laboratory of Petroleum Resources Research, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China

a r t i c l e

i n f o

Article history: Received 6 September 2014 Received in revised form 23 November 2014 Accepted 24 November 2014 Available online 12 February 2015 Editor: J. Knight Keywords: Trace metals Organic carbon isotopes Sulfur isotopes Redox proxies Mass extinction Permian–Triassic boundary

a b s t r a c t A multiproxy study of a new Upper Permian–Lower Triassic section (Xiaojiaba) in Sichuan Province, China, documents large changes in marine productivity, redox conditions and detrital input prior to the latest Permian mass extinction. Marine productivity, as proxied by total organic carbon content (TOC), biogenic SiO2, and excess barium, displays a long-term decline through most of the Changhsingian stage (late late Permian), culminating in very low values around the Permian–Triassic boundary. Concurrently, redox proxies including pyrite framboid, δ34Spy, Moauth and Uauth, and Corg/P document a shift from suboxic to dysoxic/oxic conditions that was interrupted by several episodes of benthic euxinia, and detrital siliciclastic proxies (Al, Hf, Nb, and REEs) suggest an increased flux of weathered material from land areas. The long-term changes in productivity, redox conditions, and terrigenous detrital fluxes were probably caused by a regional sea-level fall across the South China Craton. On the other hand, the brief euxinic episodes occurring during the late Permian had oceanographic causes, probably related to the transient upward expansion of the chemocline at the top of the oceanic oxygen-minimum zone. These euxinic episodes may have been harbingers of the more widespread anoxia that developed concurrently with the latest Permian mass extinction and that may have played a major role in triggering the largest biotic crisis of the Phanerozoic. © 2015 Elsevier B.V. All rights reserved.

1. Introduction The latest Permian mass extinction (LPME) was the largest biotic crisis in Earth's history, with a ~ 90% species-level loss of diversity among marine invertebrates (Erwin et al., 2002; Alroy et al., 2008). The causes of this extinction event were complex, but an important factor is thought to have been an abrupt expansion of oceanic anoxia and photic zone euxinia (Wignall and Hallam, 1992; Wignall et al., 1995; Knoll et al., 1996, 2007; Wignall and Twitchett, 1996; Hotinski et al., 2001; Grice et al., 2005; Huey and Ward, 2005; Kump et al., 2005; Gorjan et al., 2007; Cao et al., 2009; Brennecka et al., 2011; Kaiho et al., 2012; Dustira et al., 2013; Takahashi et al., 2014). Some studies have claimed that the onset of widespread oceanic anoxia (‘superanoxia’) started during the early Wuchiapingian (Isozaki, 1997; Kato et al., 2002), whereas other studies have proposed a later onset, e.g., during the late Wuchiapingian or early Changhsingian (Nielsen and Shen, 2004; Wignall et al., 2010). The timing and extent of oceanic

anoxia during the late Permian and early Triassic remain uncertain (Dustira et al., 2013) and are in need of further investigation. Here, we present an analysis of marine productivity, redox conditions, and terrigenous detrital fluxes at Xiaojiaba, a new Permian–Triassic (P–Tr) boundary section located in Sichuan Province, South China (Fig. 1a). In this study, we utilize total organic carbon (TOC), biogenic silica (SiO2(bio)), and excess Ba (Baxs) as proxies for marine productivity; the morphology, sulfur concentration (Spy) and isotopic composition (δ34Spy) of pyrite, organic carbon: phosphorus ratio (Corg/P), and authigenic Mo and U concentrations (Moauth and Uauth) as proxies for marine redox conditions; and Al, Hf, Nb and rare earth element (REE) concentrations as proxies for terrigenous detrital inputs. These analyses provide new insights into changes in redox conditions and their underlying controls during the late Permian prior to the LPME. 2. Geological background 2.1. Paleogeographic setting

⁎ Correspondence to: H. Wei, College of Earth Science, East China Institute of Technology, Nanchang, Jiangxi Province, 330013, China. ⁎⁎ Corresponding author. E-mail addresses: [email protected], [email protected] (H. Wei), [email protected] (T.J. Algeo).

http://dx.doi.org/10.1016/j.sedgeo.2014.11.008 0037-0738/© 2015 Elsevier B.V. All rights reserved.

The study area was on the northern (paleo-western) margin of the South China Craton, a microcontinent that was located in the tropics during the late Permian (Fig. 1b). The Xiaojiaba and Chaotian sections described below were located within the Guangwang Basin, one of

H. Wei et al. / Sedimentary Geology 319 (2015) 78–97

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Fig. 1. (a) Location map for the Xiaojiaba study section; also shown is the nearby Chaotian section. (b) Paleogeography of the South China Craton (modified from Wang and Jin, 2000 and Feng et al., 1997; base map is from Algeo et al., 2013). (c) Paleogeography of the Guangwang Basin (modified from Wang et al., 2009). In the paleogeographic maps, the South China Craton is rotated ~90° counter-clockwise relative to its modern orientation. ET = Ekman transport; EBC = eastern boundary current; GY = Guangyuan; JB = Jiangnan Basin.

several deep-water basins that developed on this same cratonic margin (Feng et al., 1997; Wang and Jin, 2000; Ma et al., 2006; Fig. 1c). The Guangwang Basin is located in the area between the cities of Guangyuan and Wangcang and is equivalent to the ‘Guangyuan-Liangping Bay’ of Yin et al. (2014).

2.2. Xiaojiaba study section The Xiaojiaba section is a new P–Tr boundary section located at the Xiaojiaba Bridge on the No. 108 national highway in the Chaotian district, Guangyuan, northwestern Sichuan Province, South China

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(32°38.25′ N, 105°53.35′ E) (see also Wei et al., 2012a) (Fig. 1a). It was deposited in a deep-shelf environment, about 45 km from a slope on the paleowestern margin of the Yangtze Platform (Fig. 1c). The upper Permian Wuchiaping and Dalong formations and the lower Triassic Feixianguan Formation exhibit nearly complete exposure in outcrop (Fig. 2). The 13-m-thick Wuchiaping Formation (Wuchiapingian stage) consists of intercalated gray to dark-gray, thinbedded nodular chert and thin- to medium-bedded lime mudstone. The 15-m-thick Dalong Formation consists of dark, weakly bioturbated, thin-bedded radiolarian muddy chert in the lower and middle parts, and of thin- to medium-bedded mudstone, siliceous limestone, and nodular limestone in the upper part (Fig. 2). The siliceous and nodular limestone succession contains abundant fossils, including brachiopods, gastropods, and ammonoids. Although a few volcanic tuffs are found in the lower and middle parts of this formation, the upper part contains a large number of such layers. The Dalong Formation is thought to span the upper Wuchiapingian to uppermost Changhsingian stages, terminating at the stratigraphic level of the LPME (Fig. 2). The base of the Feixianguan Formation consists about 0.8 m of darkgray calcareous shale with several thin-bedded gray muddy limestone interlayers. This P–Tr boundary interval contains two N5-cm-thick and a number of N0.5-cm-thick volcanic tuff layers (at ~ 15 m, Fig. 2), suggesting frequent, large volcanic eruptions. The two thicker tuff layers are sometimes referred to as the “boundary clays” and are a distinctive feature of the P–Tr boundary in sections across South China (Yin et al., 1992; Peng et al., 2001). Overlying these tuffs are thin-bedded muddy limestone, thin- to medium-bedded lime mudstone, and thick-bedded lime mudstone. No fossils have been found in the lower Feixianguan Formation based on field and thin-section observations (Fig. 2). For ease of reference, we have defined four stratigraphic intervals in the Dalong and lower Feixianguan formations based on lithologic and geochemical characteristics of the study section (Fig. 2; see Section 4). Interval I, at 0 to 8.7 m, is characterized by black thin-bedded radiolarian cherts showing little chemostratigraphic variation, representing a relatively stable deepwater environment. Interval II, at 8.7 to 12.5 m, is characterized by an upward gradation from fossil-poor mudstone with limestone interbeds in the lower part to fossil-rich siliceous limestones in the upper part, and by incipient shifts in the profiles of δ34Spy, δ13Corg, and biogenic SiO2. Interval III, at 12.5 to 14.5 m, is characterized by thin- to medium-bedded nodular limestone, low TOC, and large fluctuations in the profiles of δ34Spy, δ13Corg, and biogenic SiO2. This unit may represent an unstable shallow-water environment. Interval IV, at 14.5 to 16.2 m (spanning the LPME horizon and P–Tr boundary), is characterized by mixed shale-limestone deposition and loss of most benthic faunal components. 2.3. Chaotian P–Tr boundary section We compare our results for Xiaojiaba to those for the previously studied Chaotian section, which is located only 4.5 km southwest of Xiaojiaba (Fig. 1a). Chaotian also represents a deep-shelf environment (Ji et al., 2007; Saitoh et al., 2013; Fig. 1b), although water depths were probably slightly less than for Xiaojiaba. As at Xiaojiaba, frequent ash layers in this section provide evidence of significant volcanic activity well before the LPME. Planktic, nektic, and benthic organisms disappeared at the top of the Dalong Formation (correlative with the LPME), recording the effects of strong environmental stresses (Isozaki et al., 2007). The conodont and ammonoid biostratigraphy of the Changhsingian stage is well constrained at Chaotian (Isozaki et al., 2007; Ji et al., 2007). 2.4. Correlations and biostratigraphic framework Although the Xiaojiaba section has not received any biostratigraphic study to date, its proximity to the Chaotian section permits detailed lithostratigraphic correlations and, thus, transfer of biostratigraphic

information (Fig. 3). These two sections show similar lithologic changes in vertical profile, displaying an abrupt transition from shallowwater limestone in the upper Wuchiaping Formation to deepwater mudstone, or chert in the basal Dalong Formation and an abrupt transition from limestone in the upper Dalong Formation to the dark-gray calcareous shale in the lower Feixianguan Formation (Fig. 3). Earlier biostratigraphic studies have defined ammonoid and conodont zones in the Chaotian section (Isozaki et al., 2007; Ji et al., 2007), in which the stratigraphic positions of the Wuchiapingian– Changhsingian (W–C) and P–Tr boundaries are well-constrained. The W–C boundary is marked by the appearance of the ammonoids of Pseudostephanites and Tapashanites (Isozaki et al., 2007). The correlative level at Xiaojiaba is at ~ 2.8 m in Interval I (Fig. 3). The stratigraphic intervals corresponding to the Changhsingian conodont zones Clarkina yini, C. changxingensis, C. subcarinata, and C. orientalis are thus readily estimated for the Chaotian section (Fig. 3). The P–Tr boundary, which is characterized by the first appearance of Hindeodus parvus (Yin et al., 2001), is located in the ~ 1-m-thick shale at the base of the Feixianguan Formation at both Chaotian and Xiaojiaba (Isozaki et al., 2007) (Fig. 3). 3. Methods Fifty-six samples were collected in the Xiaojiaba section at an average stratigraphic spacing of ~0.3 m. Care was exercised to collect fresh samples without veins, roots, or strongly weathered surfaces. Samples were powdered in an agate ball mill. Twenty-six thin sections were selected and cut for petrographic analysis, which included examination of sediment fabric, bioturbation intensity, and fossil distribution. Pyrite was extracted using the chromium reduction method of Canfield et al. (1986). For each powdered sample, a 2 to 8-g split was treated with 1 M CrCl2 solution and 6 N HCl + 10 ml alcohol. The evolved hydrogen sulfide was immediately purged by N2 flow to a trap containing a zinc acetate solution, to which 2% AgNO3 + 6 N NH4OH were added to precipitate Ag2S. The Ag2S precipitate was then filtered, rinsed, dried, and weighed. The reproducibility of replicate analyses was generally better than 1%. Pyrite sulfur (Spy) is calculated by stoichiometry from the extracted Ag2S. The dried Ag2S was mixed with V2O5 and pure quartz sand and then converted to SO2 by combustion with copper turnings at 1050 °C. The sulfur isotopic composition of the purified SO2 was analyzed using a Finnigan Delta-S mass spectrometer at the Institute of Geology and Geophysics, Chinese Academy of Sciences in Beijing. Sulfur isotopic ratios are expressed using standard δ-notation relative to the Cañon Diablo Troilite (CDT) standard. Analytical precision is better than ±0.2‰. Sample splits (0.3 to 5 g) for bulk δ13Corg analysis were treated with 6 N HCl for 24 h to remove carbonate. The solution was then retreated with excess 6 N HCl and allowed to sit for 6 h to make sure that all carbonate was dissolved. About 30–100 mg of decalcified sample and 1 g of CuO wire were placed in a quartz tube and combusted at 500 °C for 1 h and at 850 °C for 3 h. The carbon isotope ratio of the generated CO2 was measured in a Finnigan MAT-252 mass spectrometer at the Institute of Geology and Geophysics, Chinese Academy of Sciences in Beijing. The carbon isotopic ratio is reported in standard δ notation relative to the Vienna Peedee Belemnite (VPDB) standard. Analytical precision is better than ±0.1‰. Sample splits (200 mg) for total organic carbon (TOC) analysis were firstly treated with 10% HCl at 60 °C to remove carbonate, and then washed with distilled water to remove HCl. Afterwards, the samples were dried overnight at 50 °C and then measured using a LECO CS-400 analyzer in the Research Institute of Petroleum Exploration and Development in Beijing. The standard deviation of TOC measurements is lower than ±0.10%. Sample splits (0.5 g) for major-element analysis were fused to glass pellets and analyzed using a Phillips PW 1500 X-ray fluorescence

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Fig. 2. Lithologic log and fossil distributions at Xiaojiaba. Volcanic tuffs (star symbols) are found mostly just below the P–Tr boundary, between 11 and 16 m (i.e., the interval shown in photo at upper right). The spindle diagrams represent fossil abundance (not species diversity); the horizontal scales show relative variation within each biotic group but are not equivalent between groups. Intervals I–IV are discussed in the text.

spectrometer at the Institute of Geology and Geophysics, Chinese Academy of Sciences in Beijing. The precision of the major-element concentration data is better than ±3% of reported values.

Trace elements and rare earth elements (REEs) were measured by inductively coupled plasma mass spectrometry (ICP-MS) with a Finnigan MAT Element II mass spectrometer at the Institute of Geology

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Fig. 3. Stratigraphic correlation of the Xiaojiaba study section with Chaotian (dashed lines). Note that these sections are only 4.5 km apart (Fig. 1). Data for Chaotian are from Isozaki et al. (2007) and Ji et al. (2007). See Fig. 2 for lithologic key.

and Geophysics, Chinese Academy of Sciences in Beijing. Powdered samples (40 mg) were dissolved and digested in mixed acid (HF + HNO3 + HClO4) in 15 ml Savillex Teflon screw-cap beakers and

high-pressure Teflon bombs at 120 °C for 6 days, dried, and then diluted to 50 ml for analysis. Precision for all trace-element concentrations is estimated to be ±5% of reported values.

H. Wei et al. / Sedimentary Geology 319 (2015) 78–97

The authigenic fraction (Xauth) of the elements Mo and U (cf., Tribovillard et al., 2006) was calculated as: Xauth ¼ ½X–½Al  ½X=Aldetrital

ð1Þ

83

the thin section represents a random cut through an ensemble of spheres of non-uniform sizes, the measured size distributions are only an approximation of actual framboid sizes (Wilkin et al., 1996) although probably within ~10% of actual size (Cashman and Ferry, 1988).

where the detrital X/Al ratio was based on average upper crustal concentrations of Mo (1.5 ppm), U (2.8 ppm), and Al (8.04%) (McLennan, 2001). Biogenic chert (SiO2(bio)) was calculated as:

4. Results

SiO2ðbioÞ ¼ SiO2ðtotalÞ –ðm  K2 O  2:36Þ–ððAl2 O3 –m  K2 OÞ  1:18Þ ð2Þ

TOC content at Xiaojiaba ranges from 0.02% to 2.44%, with an average of 0.52% (Table 1). A relatively steady decline to nearly zero is observed upsection from Interval I to Interval IV (Fig. 5). Biogenic silica (SiO2(bio)) content ranges from 2.8% to 89.1%, with an average of 46.2% (Table 1, Fig. 5). SiO2(bio) is nearly uniform (75–85%) in Interval I, before declining in Interval II to lower values (20–25%) in Interval III. The upper part of Interval III and Interval IV show a modest rebound in SiO2(bio) to 40–50%. Biogenic silica shows a strong exponential relationship to TOC (r2 = 0.54), suggesting a common control related to primary productivity in the study area (Fig. 6a). Excess barium (Baxs) content ranges from 0 to 739 ppm, with an average of 192 ppm (Table 2). Baxs peaks at N600 ppm in the lower part of Interval I and then declines somewhat irregularly upsection, reaching near-zero values in the upper part of Interval II and in Intervals III and IV (Fig. 5). Baxs shows a modest correlation with SiO2(bio) (r 2 = 0.30), suggesting a common control related to primary productivity in the study area (Fig. 6b).

where m is the slope of the Al2O3–K2O regression, and the coefficients represent the weight ratios of SiO2/Al2O3 (2.36) and 2 × SiO2/Al2O3 (1.18). These weight ratios pertain to clay minerals of stoichiometric composition having TOTO and TOT structures (i.e., illite and chlorite), respectively (cf., Algeo et al., 2007). Eq. (2) subtracts the amount of SiO2 present in TOTO clay minerals (based on total K2O concentration) and in TOT clay minerals (based on excess Al2O3 not resident in TOTO clay minerals) from total SiO2. It assumes that non-clay SiO2 was derived from biosilica sources as opposed to detrital quartz grains. Excess barium (Baxs) was calculated as: Baxs ¼ Ba–Al  50

ð3Þ

where the constant 50 represents the estimated Ba/Al ratio of the detrital fraction (units in ppm/% or 10- 4) (Fig. 4). This estimate is based on a regression of Ba versus Al for the Xiaojiaba study section. In contrast to most REEs, which have a fixed 3+ valence, europium (Eu) can switch between valences of 2 + and 3 + (Elderfield and Greaves, 1982). It is therefore sensitive to environmental redox conditions and can become fractionated from other REEs. Deviations of europium relative to its neighbors (the Eu anomaly, or Eu/Eu*) were quantified as: 

Eu=Eu ¼ 3  Eu=ð2  Sm þ GdÞ:

ð4Þ

All REE concentrations were normalized to Post-Archean Average Shale (PAAS) (Taylor and McLennan, 1985; McLennan, 2001). For measurement of pyrite framboid sizes, 25 samples from the limestone of upper Dalong Formation and lower Feixianguan Formation around the P–Tr boundary were selected for preparation as thin sections. Pyrite framboid diameters were measured using a FEI Nova NanoSEM 450 scanning electron microscope (SEM) equipped with secondary electron energy dispersive X-ray (EDX) analysis located at the East China Institute of Technology in Nanchang, China. Only normal pyrite framboids and infilled framboids were measured, with at least 100 size measurements per sample. Because the polished surface of

4.1. TOC, SiO2(bio), and Baxs

4.2. Spy and δ34Spy Pyrite sulfur (Spy) content ranges from 0.03% to 1.07%, with an average of 0.31% (Table 1, Fig. 7). Spy profile shows a general decline upsection interrupted by several peaks that coincide with thicker tuff layers. Dark muddy cherts in the lower part of the Dalong Formation contain higher Spy content, whereas gray limestones in the upper part of the Dalong Formation contain much lower Spy content. δ34Spy ranges from −43.4‰ to +5.0‰, with an average of −26.4‰ (Table 1). Intervals I and II exhibit relatively light values (− 40‰ to − 25‰), although with a trend toward heavier values in the upper part of Interval II (Fig. 7). Intervals III and IV are characterized by much heavier δ34Spy, from −20‰ to +10‰, with occasional extremely negative values from −43.4‰ to −28.7‰ (Fig. 7). Spy shows a strong correlation with δ34Spy (r2 = 0.71; Fig. 8a). This relationship follows an exponential pattern, with a rapid shift toward heavier δ34Spy as Spy approaches 0%, and a slow shift toward lighter δ34Spy as Spy increases above 0.2%. The δ34Spy values of the low- and high-Spy endmembers are approximately +5‰ and −40‰, respectively. These results are very similar to those for the P–Tr boundary section at Nhi Tao, Vietnam, which yielded δ34Spy values of + 10‰ and − 45‰ for low- and high-Spy endmembers, respectively (Algeo et al., 2008; Fig. 8b). 4.3. Moauth,, Uauth, Corg/P ratio, Fe2O3/Al2O3, REE, Hf, Nb and δ13Corg

Fig. 4. Ba versus Al for Xiaojiaba. The line represents the inferred detrital Ba/Al ratio of 50. The sample points above this line contain excess Ba of possible biogenic origin.

Authigenic molybdenum (Moauth) content ranges from 0 to 80.7 ppm, with an average of 13.1 ppm (Fig. 7). However, most Moauth values are lower than 30 ppm, with only a few samples exhibiting pronounced enrichment. The highest values are found in Interval I, and the Moauth profile exhibits an irregular decline upsection through Intervals II to IV. Authigenic uranium (Uauth) content ranges from 0 to 15.9 ppm, with an average of 5.5 ppm (Fig. 9). Uauth values fluctuate between 5 and 15 ppm through Intervals I and II, before declining to values uniformly b5 ppm in Intervals III and IV. Total Mo and U concentrations are only marginally greater than Moauth and Uauth (Table 2), indicating that the detrital fraction of these trace elements is minor. Phosphorus (P) content ranges from 0.02 to 0.78%, with an average of 0.13% (Table 1). The molar Corg/P ratio ranges from b 1 to 253, with

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Table 1 Major element, TOC, pyritic sulfur, and isotopic data for Xiaojiaba, South China. Sample no.

height (m)

TOC (%)

δ13Corg (‰)

Spy (%)

δ34Spy (‰)

SiO2 (%)

TiO2 (%)

Al2O3 (%)

Fe2O3 (%)

MnO (%)

MgO (%)

CaO (%)

Na2O (%)

K2O (%)

P2O5 (%)

LOI (%)

SiO2(bio) (%)

Fe2O3/Al2O3

XJB31 XJB29 XJB28 XJB27 XJB26 XJB25 XJB24 XJB23 XJB22 XJB20 XJB19 XJB18 XJB17 XJB16 XJB15 XJB14 XJB13 XJB12 XJB11 XJB10 XJB09 XJB08 XJB07 XJB06 XJB05 XJB04 XJB03 XJB02 XJB01

15.15 14.94 14.87 14.75 14.35 13.90 13.20 12.80 12.47 11.80 11.45 11.15 10.95 10.28 10.15 9.80 9.30 8.85 8.20 7.65 5.65 5.00 4.00 3.25 2.65 2.25 1.70 1.30 0.30

0.04 0.04 0.02 0.02 0.02 0.06 0.05 0.05 0.28 0.55

−26.57 −26.02 −26.35 −22.09 −26.20 −27.34 −26.74 −25.56 −24.56 −25.87 −26.87 −27.18 −26.87 −27.12 −26.79 −25.93 −26.81 −27.38 −27.34 −26.93 −27.62 −27.67 −28.00 −27.48 −27.85 −27.28 −27.81 −27.95 −26.93

0.12

−28.71

0.033 0.60 0.05 0.05 0.04 0.03 1.07 0.24

1.60 1.60 −43.41 5.03 −12.50 0.75 −20.32 −29.76 −26.93

26.49 70.28 37.89 66.09 25.14 7.04 15.59 5.56 30.06

0.08 0.23 0.16 0.65 0.11 0.05 0.09 0.05 0.13

2.19 6.00 3.32 13.34 2.34 1.05 1.90 1.19 3.78

1.19 1.70 1.79 6.16 1.11 0.66 0.86 0.42 1.21

0.14 0.05 0.10 0.10 0.08 0.10 0.10 0.10 0.04

0.63 0.60 0.68 2.16 1.02 1.01 0.96 0.93 0.70

36.70 9.29 29.00 4.50 36.95 48.10 42.93 49.15 32.78

0.39 0.50 0.48 3.10 0.77 0.45 0.56 0.54 0.66

0.42 1.35 0.72 2.63 0.32 0.10 0.28 0.09 0.69

0.07 0.10 0.06 0.15 0.06 0.06 0.05 0.05 0.16

31.34 9.59 25.21 0.99 31.79 41.07 36.54 41.88 29.72

21.33 56.14 30.07 34.65 19.62 4.57 11.11 2.76 21.15

0.54 0.28 0.54 0.46 0.47 0.63 0.45 0.35 0.32

3.43 2.44 1.88 0.79 1.85 6.45 5.80 6.43 10.35

69.29 51.28 62.56

0.18 0.11 0.13

4.11 2.92 2.83

1.42 0.93 1.08

0.01 0.03 0.02

0.43 0.46 0.66

10.65 21.56 15.81

0.36 0.42 0.42

0.87 0.47 0.49

0.38 0.10 1.19

12.11 21.62 14.27

59.60 44.40 55.89

0.35 0.32 0.38

17.13 31.43 4.33

68.22 32.52 92.21 91.02 91.92

0.11 0.08 0.06 0.04 0.09

2.50 2.04 1.33 0.94 2.00

0.99 0.77 1.22 0.30 0.97

0.01 0.09 0.01 0.01 0.01

0.51 0.59 0.29 0.17 0.35

13.61 33.86 0.56 3.10 0.87

0.49 0.61 0.24 0.23 0.32

0.41 0.20 0.22 0.15 0.35

0.57 0.66 0.05 1.78 0.08

12.07 28.09 3.69 2.12 2.88

62.33 27.71 89.08 88.80 87.21

0.40 0.38 0.92 0.32 0.49

7.16 2.87 102.09 1.43 71.00

91.73

0.08

1.77

1.13

0.01

0.25

0.62

0.25

0.31

0.07

3.59

87.56

0.64

98.61

84.60

0.24

4.69

0.86

0.01

0.43

0.61

0.30

1.43

0.06

6.66

73.55

0.18

253.27

1.10 0.51 0.87

0.69 0.32 0.88 0.43 0.90 1.15

2.44

0.16

−33.18

0.18

−38.74

0.63 0.24 0.29 0.37

−36.98 −35.45 −39.03 −41.19

0.56

−40.27

0.94 0.07

−26.01 −30.34

an average of 33 (Table 1, Fig. 7). The highest Corg/P ratios (N100) are found in the lower part of the Dalong Formation, and Corg/P declines steadily through Interval I before reaching low and nearly uniform values of b20 in Intervals II to IV. Al2O3 content ranges from 0.9% to 13.3%, with an average of 3.2%, and Fe2O3 content ranges from 0.3% to 6.2%, with an average of 1.3% (Table 1). Fe2O3/Al2O3 ratios range from 0.18 to 0.92, with an average of 0.44 (Fig. 7). The Fe2O3/Al2O3 profile suggests that most samples do not deviate greatly from the average shale value of ~ 0.4 (Lyons and Severmann, 2006), although one sample with high Spy content in the upper part of Interval I exhibits a markedly higher Fe2O3/Al2O3 ratio (Fig. 7). Total REE concentration (ΣREE) ranges from 8 to 123 ppm, with an average of 38 ppm (Table 3, Fig. 9). ΣREE fluctuates irregularly between 10 and 60 ppm from Interval I to the lower part of Interval III before rising to 60–120 ppm in the upper part of Interval III and Interval IV. ΣREE exhibits strong positive covariation with Al2O3 (r2 = 0.65; Fig. 10). The europium anomaly (Eu/Eu*) ranges from 0.44 to 5.24, with an average of 1.45 (Table 3, Fig. 9). The lower part of Interval I yields Eu/Eu* N2, but the remainder of the section shows values mostly close to 1.0, although with Eu/Eu* of b0.5 in a few samples (Fig. 9). Hf concentration ranges from 0.22 to 3.66 ppm, with an average of 1.09 ppm (Table 2, Fig. 9). Nb concentration ranges from 0.70 to 7.61 ppm, with an average of 2.58 ppm (Table 2, Fig. 9). Both Hf and Nb show a gradual increase upsection with peaks in the upper part of Interval II and the lower part of Interval IV (Fig. 9). δ13Corg ranges from − 28.0‰ to − 22.1‰, with an average of − 26.7‰ (Table 1). The δ13Corg profile exhibits background values of − 28‰ to − 26‰ throughout the study section. Positive C-isotopic excursions are evident in the lower part of Interval II, upper part of Interval II, and lower part of Interval IV (Fig. 9). δ13Corg covaries negatively with TOC to a modest degree (r2 = 0.36; Fig. 11). This relationship follows an exponential pattern, with a rapid shift toward heavier δ13Corg as TOC approaches 0%, and a slow shift toward lighter δ13Corg as TOC increases above 0.5%. The δ13Corg values of the low- and high-TOC

Corg/P

endmembers are approximately − 22‰ to − 25‰ and − 28‰, respectively. 4.4. Pyrite morphologies and framboid size distributions 4.4.1. Pyrite morphologies The limestone beds of the Dalong and Feixianguan formations contain varying quantities of pyrite (Table 4), with morphologies ranging from normal framboids to infilled framboids and euhedral pyrite crystals. The Dalong Formation limestones contain common to abundant framboids, whereas the limestone interlayers in the dark shales of the lowermost Feixianguan Formation contain few or no framboids (Fig. 12). Normal framboids consist of spherical or elliptical aggregates of microcrystals, the intragranular spaces of which are easily recognized, and which occur both as discrete grains and as grain clusters (Fig. 13a, b, d). Euhedral pyrite crystals have idiomorphic form and are internally structureless (Fig. 13c, e, f). Morphologies intermediate between normal framboids and euhedral crystals are present, including infilled framboids (Fig. 13d, e), massive spherules (Fig. 13e), overgrown framboids (Fig. 13f), and pyrite lumps (Fig. 13f). Infilled framboids are closely packed with barely recognizable boundaries between microcrystals, and small holes indicative of the former framboidal texture (Fig. 13d, e). Massive spherules consist of structureless pyrite in a spherical shape (Fig. 13e), and pyrite lumps consist of anhedral crystals in structureless masses (Fig. 13f). Another intermediate morphology is overgrown framboids, which show a secondary growth of microcrystals with an idiomorphic texture as an exterior rim (Fig. 13f). All of these intermediate morphologies can be produced through continued growth of pyrite microcrystals in a normal framboid, resulting in crystal amalgamation, pore space infilling, and generation of anhedral and massive crystal textures (e.g., Sawlowicz, 1993). Thus, pyrite morphology can evolve from normal framboids to infilled or overgrown framboids to massive or euhedral crystals. Annular framboids, polyframboids, and macroframboids are not common in the study samples. Annular framboids are hollow framboids consisting of a sphere (a ring in cross-section) formed of minute pyrite

H. Wei et al. / Sedimentary Geology 319 (2015) 78–97

85

Fig. 5. Chemostratigraphic profiles of TOC, SiO2(bio), and Baxs for Xiaojiaba.

grains (Sawlowicz, 1993; Fig. 13g). The annular form of framboids can be a first step in the growth of a normal framboid (Kizilshtein and Minaeva, 1972). Their occurrence in the Xiaojiaba section may indicate low availability of dissolved iron or sulfide (e.g., Papunen, 1966; Love,

1967; Kosacz and Sawlowicz, 1983). Polyframboids consist of multiple, conjoined framboids of different sizes (Fig. 13g). Development of polyframboids occurs through similar processes as normal framboids, although the size of polyframboids (up to 90 μm in diameter) reflects

86

H. Wei et al. / Sedimentary Geology 319 (2015) 78–97

Fig. 6. (a) SiO2(bio) versus TOC and (b) SiO2(bio) versus Baxs for Xiaojiaba. The relationship in (a) was fit with an exponential function and that in (b) with a linear equation. SiO2(bio) = biogenic chert; Baxs = excess barium.

formation within the sediments instead of in the water column (Sawlowicz, 1993). Macroframboids are texturally similar to normal framboids but are much larger, reaching 50 μm in diameter (Croxford and Jephcott, 1972; Fig. 13h). Such large sizes require a long residence time at the oxic/anoxic interface during framboid formation, consistent with a diagenetic origin. 4.4.2. Pyrite framboid size distributions Pyrite framboid size distributions show considerable sample-tosample variation in vertical profile (Fig. 14). For most samples, the maximum framboid size is larger than 17 μm, although several samples show maximum values smaller than 13 μm. Mean pyrite framboid sizes range from 5.3 to 22.3 μm (Table 4). Crossplots of mean size vs standard deviation (Fig. 15a) and mean size vs skewness (Fig. 15b) are useful means of discriminating framboids of syngenetic versus diagenetic origin, which are indicative of euxinic and dysoxic/oxic conditions, respectively (cf., Wilkin et al., 1996). Pyrite framboids contained in modern sediments underlying sulfidic water columns are on average smaller and less variable in size than framboids from sediments underlying oxic or dysoxic water columns (Wilkin et al., 1996). 5. Discussion 5.1. Marine productivity changes Paleoproductivity changes can be reconstructed using proxies such as TOC, excess Ba, and biogenic SiO2 (e.g., Schmitz, 1987; Gingele and Dahmke, 1994; Bonn et al., 1998; Algeo et al., 2011a, 2013; Schoepfer et al., in press). Late Paleozoic to early Mesozoic phytoplankton were largely non-mineralizing forms, primarily contributing organic matter to marine sediments (Riegel, 2008). The gradual upsection decrease in TOC at Xiaojiaba (Fig. 5) may reflect a long-term decline in marine primary productivity. Very low TOC values in Intervals III and IV (mostly b0.1%) may be indicative of a severe decline in productivity during the pre-LPME late Changhsingian. This decline appears to predate the primary productivity crash associated with the LPME horizon across the South China Craton (Algeo et al., 2013). Excess barium (Baxs) accumulates in sinking organic matter undergoing decay, forming small barite crystals that are resistant to remineralization under oxic bottomwater conditions (Dehairs et al., 1980, 1990; Bishop, 1988). The proportion of Baxs that survives burial can be far larger (~30–50%; Dymond et al., 1992; Paytan and Kastner, 1996; Paytan and Griffith, 2007) than that for organic carbon and phosphorus (b10%; Canfield, 1994; Hedges and Keil, 1995), making it a useful productivity proxy in many marine systems (Dymond et al., 1992; Bonn et al., 1998; Eagle et al., 2003). At Xiaojiaba, the decline in Baxs through Intervals I and II mirrors the decline in TOC (Fig. 5). Although declining TOC might have been due to increasing oxygenation

and, thus, to greater aerobic decay of organic matter, the same cannot be true of excess Ba, the stability of which increases under more oxic conditions. Weak positive covariation of Baxs with SiO2(bio) (r2 = 0.30, Fig. 6b) provides additional evidence for productivity controls on organic matter accumulation (cf., Schoepfer et al., 2013). In deep-water Permo-Triassic sections, the biogenic SiO2 flux is largely derived from radiolarians (Feng et al., 2007; Isozaki et al., 2007; Algeo et al., 2010, 2011a), which are single-celled primary consumers living mainly in the ocean-surface layer (Caron et al., 1995; Dennett et al., 2002). Because marine phytoplankton are their primary food source, the radiolarian SiO2 flux can be a good proxy for surfacewater productivity. Changes in nutrient availability, which are linked to environmental disturbances or climate change, are known to influence radiolarian abundance (Kakuwa, 1996; Beauchamp and Baud, 2002; Isozaki et al., 2007). At Xiaojiaba, the upsection decrease in biogenic SiO2 concentrations from the base of Interval II through Interval III is consistent with a decrease in primary productivity as inferred from both the TOC and Baxs profiles (Fig. 9). This inference is reinforced by strong positive covariation of SiO2(bio) with TOC (r2 = 0.54, Fig. 6a). The consistent pattern shown by these three proxies suggests that the derived paleoproductivity inferences are robust. The paleogeographic location of the Xiaojiaba section on the paleowestern margin of the South China Craton would have favored relatively high productivity levels owing to offshore Ekman transport of an eastern boundary current, resulting in upwelling of nutrient-rich thermocline waters (Fig. 1b). A similar condition also existed on the western Laurentian margin of Pangea during the Permian (Beauchamp and Baud, 2002; Schoepfer et al., 2012, 2013). At Xiaojiaba, the abundant radiolarian cherts in Interval I of the Dalong Formation as well as in the underlying Wuchiaping Formation (Fig. 2) are evidence of a vigorous upwelling-related high-productivity system in the eastern Paleotethys. Replacement of the biogenic silica factory by carbonate sedimentation during the late Changhsingian at Xiaojiaba and in other western Yangtze platform sections (Nafi et al., 2006; Riccardi et al., 2006, 2007; Isozaki et al., 2007) may have been a consequence of decreased upwelling (cf., Beauchamp and Baud, 2002; Kidder and Worsley, 2004; Schoepfer et al., 2013) and/or regional shallowing in South China (Tong et al., 1999). 5.2. Benthic redox variation 5.2.1. Pyrite framboid size distributions Pyrite morphology and framboid size distribution are promising tools for evaluating bottomwater redox conditions (Wilkin et al., 1997). With regard to pyrite morphology (Fig. 13), secondary pyrite growths such as infilled framboids, overgrown framboids, and pyrite lumps indicate a diagenetic origin. Polyframboids and macroframboids also have a diagenetic origin, indicating formation in the sediment

Table 2 Trace element data for Xiaojiaba, South China. Li (ppm)

Sc (ppm)

V (ppm)

Cr (ppm)

Co (ppm)

Ni (ppm)

Cu (ppm)

Zn (ppm)

Rb (ppm)

Sr (ppm)

Mo (ppm)

Cd (ppm)

Ba (ppm)

Pb (ppm)

Th (ppm)

U (ppm)

Zr (ppm)

Hf (ppm)

Nb (ppm)

Moauth (ppm)

Uauth (ppm)

Baxs (ppm)

XJB31 XJB29 XJB28 XJB27 XJB26 XJB25 XJB24 XJB23 XJB22 XJB18 XJB17 XJB16 XJB13 XJB12 XJB11 XJB10 XJB09 XJB08 XJB07 XJB06 XJB05 XJB04 XJB03 XJB02 XJB01

11.4 35.2 23.9 17.1 6.11 2.7 5.51 3.69 16.1 33.2 19.2 21.6 33.1 9.52 16.1 22.7 23.1 37.8 21.3 31.3 34.5 20.3 24.7 25.6 31.1

2.3 5.02 3.09 18.2 2.22 0.89 2.21 0.91 2.94 3.91 3.11 2.87 2.75 1.76 1.05 0.66 1.11 1.06 1.29 1.31 1.99 0.96 1.3 1.36 4.18

21 42 38 144 20 16 22 21 75 265 85 115 85 47 365 76 225 136 329 232 306 151 259 266 202

12 27 26 28 14 8 13 11 27 53 33 49 46 22 40 28 45 195 69 164 249 118 271 222 147

2.12 5.11 5.53 15.3 3.58 2.12 4.18 2.87 11.3 13.5 5.43 4.78 5.04 2.9 5.09 2.2 4.07 3.91 4.41 3.99 5.8 2.98 4.94 4.15 0.92

16.5 41 30.1 6.3 17.1 16.3 23.1 20.3 29.7 63.8 63 52.6 41.7 48.2 53.7 18.7 39.2 39.3 71.6 42.4 51.3 32.6 46.4 36.1 41.3

16 47 25 18 27 7 17 10 101 138 73 82 51 38 28 24 40 45 42 38 37 28 30 25 31

24 62 48 52 20 13 24 15 49 68 81 129 47 41 33 19 28 21 48 38 33 22 29 28 14

18.8 61.5 31.8 100 11.4 2.2 9.8 2 29.3 35.7 19.6 19.7 16 6.1 6.7 3.8 9.8 8.1 8.7 10.2 9.8 7.5 6.7 9.5 49.9

775 203 335 192 632 883 450 534 1720 1041 572 746 663 727 71 321 81 75 42 133 90 99 85 90 112

0.06 0.66 2.27 0.47 2.15 0.18 0.73 0.41 1.41 9.42 3.09 4.98 6.38 2.81 80.9 3.71 15.7 9.41 71.5 21.1 25.7 15.6 21.8 26.8 5.78

0.52 0.85 0.41 0.15 0.16 0.14 0.44 0.17 0.69 17.2 10.5 6.05 2.87 6.58 3.97 1.12 4.56 6.76 16.9 7.09 8.32 7.86 8.84 7.59 14.4

122 146 81 396 37 20 31 16 74 109 81 85 626 40 131 326 216 291 409 510 599 739 624 416 142

6.31 9.68 6.33 11.8 6.04 2.55 2.6 1.53 7.01 9.44 8.39 8.03 7.36 5.12 4.04 4.61 4.23 1.6 4.59 2.45 2.61 2.7 2.29 2.63 31.5

1.67 5.61 2.46 13.2 1.58 0.7 1.34 0.55 2.31 3.2 2.23 2 1.81 1.32 0.88 0.52 1.25 0.98 0.93 1.03 0.82 0.97 0.72 0.9 3.67

3.45 3.26 2.09 1.79 3.7 4.97 3.08 3.18 8.42 14.1 6.09 16.5 8.11 8.31 8.61 10.8 4.64 1.97 1.38 15.2 1.96 5.94 1.57 3.17 6.04

39 128 50 106 36 10 20 11 62 73 56 52 40 35 25 42 42 29 37 34 41 26 41 31 105

0.86 2.81 1.28 3.66 0.85 0.24 0.48 0.22 1.62 1.71 1.38 1 0.91 0.74 0.47 0.46 0.81 0.71 0.73 0.65 0.82 0.69 0.73 0.72 2.67

3.1 7.61 3.29 6.53 2.79 0.704 1.55 0.723 3.96 3.88 3.76 2.48 2.26 2.22 1.18 1.74 1.47 1.22 1.32 1.4 1.82 1.16 1.55 1.36 5.54

0 0 1.83 0 1.84 0.04 0.48 0.25 0.91 8.88 2.7 4.61 6.05 2.54 80.72 3.59 15.44 9.05 71.27 20.29 24.71 15 20.96 25.77 5.16

3.02 2.07 1.43 0 3.24 4.76 2.7 2.94 7.67 13.28 5.51 15.94 7.61 7.91 8.35 10.61 4.24 1.88 1.03 14.48 1.87 5.66 1.5 3.02 5.11

64 0 0 43 0 0 0 0 0 0 4 10 560 0 96 301 163 291 362 510 599 739 624 416 18

H. Wei et al. / Sedimentary Geology 319 (2015) 78–97

Sample no.

87

88

H. Wei et al. / Sedimentary Geology 319 (2015) 78–97

Fig. 7. Chemostratigraphic profiles of pyrite sulfur (Spy), δ34Spy, authigenic Mo (Moauth) and U (Uauth), Corg/P, and Fe2O3/Al2O3 for Xiaojiaba.

instead of in the water column. Such morphologies tend to form in sediments overlain by dysoxic/oxic watermasses (e.g., Raiswell and Berner, 1985). With regard to framboid size distributions (Figs. 14, 15), larger mean diameters (~ N 6 μm) and wider variation of framboidal

diameters (with maximum diameters ~ N 18 μm) suggest a diagenetic origin and dysoxic/oxic water-column conditions (Wilkin et al., 1996, 1997; Wilkin and Arthur, 2001; Wei et al., 2012b). Samples with smaller mean diameters (~b6 μm), narrower size distributions (standard

Fig. 8. δ34Spy versus Spy for (a) Xiaojiaba (this study) and (b) Nhi Tao, Vietnam (Algeo et al., 2008). Both datasets were fit with 2nd-order polynomial equations; 1:1, 4:1, and 9:1 represent mixing ratios of syngenetic-to-authigenic pyrite, with endmember compositions of ~−45‰ for the syngenetic fraction and ~ +10‰ for the diagenetic fraction.

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Fig. 9. Chemostratigraphic profiles of total REE concentration (ΣREE), europium anomaly (Eu/Eu*), Al2O3, Hf and Nb, and δ13Corg for Xiaojiaba.

deviation ~ b2.25), and smaller maximum diameters (~b 13 μm) are characteristic of a syngenetic origin and euxinic water-column conditions (Wilkin et al., 1996, 1997; Wilkin and Arthur, 2001). Based on pyrite morphologies and framboid size distributions, we infer that paleo-marine redox conditions in the Guangwang Basin during the late Changhsingian (upper Dalong Formation) and early Induan (lower Feixianguan Formation) were mainly dysoxic to oxic but punctuated by several euxinic episodes. Redox proxies exhibit 6 to 7 redox cycles through the study section, with each cycle changing from reducing to oxidizing conditions over intervals ranging in thickness from 0.6 to 1.3 m (Fig. 14). These meter-scale cycles of redox variation may represent astronomically forced climate change and consequent fluctuations in sea-level elevation on the Yangtze platform (e.g., Tong et al., 1999; Yin et al., 2014). Redox conditions became more variable around the LPME. The darkgray shale of the lower Feixianguan Formation (Interval IV) yields 34Sdepleted pyrite S-isotopic compositions (b− 30‰; Fig. 7), suggesting deposition under reducing conditions. Previous studies have reported

anoxia/euxinia around the LPME in the nearby Shangsi section (Riccardi et al., 2006, 2007) as well as at other localities globally (Wignall and Twitchett, 1996, 2002; Isozaki, 1997; Grice et al., 2005; Kump et al., 2005; Riccardi et al., 2006, 2007; Gorjan et al., 2007; Bond and Wignall, 2010; Algeo et al., 2011a; Kaiho et al., 2012; Dustira et al., 2013; Takahashi et al., 2014). However, thin limestone interbeds within the shale at Xiaojiaba contain rare or no framboids, and one of them contains large framboids (mean = 22 μm, max. = 47 μm) (Fig. 14). These features are consistent with dysoxic or oxic conditions, which appear to have episodically interrupted the reducing conditions that generally prevailed around the LPME (cf., Kaiho et al., 2012). 5.2.2. Geochemical data for redox conditions Redox conditions can also be reconstructed using a combination of sedimentary proxies such as δ34Spy, Corg/P, Moauth, and Uauth (cf., Jones and Manning, 1994). The sulfur isotopic composition of diagenetic pyrite is generally heavier than that of syngenetic framboidal pyrite (Raiswell and Berner, 1985). At Xiaojiaba, the gradual upsection

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Table 3 REE data for Xiaojiaba, South China. Sample

La (ppm)

Ce (ppm)

Pr (ppm)

Nd (ppm)

Sm (ppm)

Eu (ppm)

Gd (ppm)

Tb (ppm)

Dy (ppm)

Ho (ppm)

Er (ppm)

Tm (ppm)

Yb (ppm)

Lu (ppm)

Y (ppm)

ΣREE (ppm)

Eu/Eu*

XJB31 XJB29 XJB28 XJB27 XJB26 XJB25 XJB24 XJB23 XJB22 XJB18 XJB17 XJB16 XJB13 XJB12 XJB11 XJB10 XJB09 XJB08 XJB07 XJB06 XJB05 XJB04 XJB03 XJB02 XJB01

18.50 22.20 8.90 25.20 12.40 4.49 4.88 1.66 7.26 8.41 13.00 10.50 7.90 7.26 2.61 3.10 5.75 2.54 2.41 6.96 3.17 4.48 2.45 3.46 12.50

26.50 41.20 16.10 48.00 22.00 6.83 9.50 3.23 13.50 16.80 26.50 17.90 13.60 11.80 4.30 4.11 9.90 4.96 1.82 8.77 5.03 6.19 3.73 5.60 20.50

3.25 4.97 1.66 5.43 2.27 0.68 0.98 0.35 1.43 1.84 3.07 1.92 1.56 1.41 0.56 0.57 1.11 0.62 0.60 1.27 0.67 0.93 0.46 0.71 2.56

12.80 19.70 6.18 20.70 8.55 2.66 3.63 1.27 5.61 6.97 11.90 7.34 6.17 5.49 2.17 2.48 4.18 2.24 2.53 5.87 2.34 3.77 1.95 2.82 9.74

2.95 4.29 1.13 4.28 1.74 0.57 0.73 0.25 1.06 1.30 2.74 1.40 1.27 1.18 0.45 0.67 0.83 0.55 0.54 1.51 0.62 0.86 0.38 0.67 1.66

0.46 0.67 0.24 1.14 0.21 0.09 0.13 0.05 0.20 0.24 0.26 0.30 0.39 0.21 0.12 0.18 0.19 0.23 0.17 0.63 0.45 0.58 0.51 0.38 0.27

3.76 4.72 1.07 4.04 2.17 0.78 0.72 0.20 1.11 1.32 2.79 1.57 1.38 1.36 0.55 0.82 0.81 0.33 0.47 1.89 0.61 0.94 0.53 0.59 1.28

0.82 0.99 0.19 0.81 0.43 0.17 0.14 0.04 0.19 0.25 0.57 0.31 0.27 0.30 0.09 0.17 0.14 0.05 0.09 0.34 0.08 0.17 0.07 0.12 0.25

5.35 5.77 1.20 4.86 2.79 0.92 0.74 0.24 1.23 1.44 3.54 1.94 1.57 1.91 0.51 1.13 0.84 0.35 0.49 2.18 0.50 0.97 0.46 0.64 1.57

1.23 1.18 0.25 1.00 0.61 0.18 0.13 0.05 0.28 0.29 0.68 0.41 0.32 0.40 0.11 0.27 0.15 0.07 0.11 0.46 0.11 0.22 0.10 0.13 0.32

4.05 3.90 0.85 3.20 1.97 0.55 0.43 0.18 0.99 0.96 2.14 1.27 0.97 1.32 0.35 1.05 0.49 0.19 0.38 1.35 0.32 0.57 0.34 0.39 1.21

0.74 0.70 0.13 0.53 0.34 0.07 0.07 0.02 0.18 0.14 0.36 0.18 0.14 0.22 0.06 0.18 0.07 0.04 0.06 0.20 0.07 0.09 0.05 0.07 0.21

4.61 4.59 0.91 3.46 2.26 0.48 0.45 0.17 1.15 0.91 2.49 1.28 0.98 1.42 0.38 1.33 0.43 0.30 0.40 1.17 0.43 0.67 0.39 0.48 1.50

0.78 0.77 0.14 0.55 0.38 0.08 0.07 0.04 0.22 0.15 0.40 0.21 0.15 0.23 0.05 0.22 0.07 0.04 0.08 0.18 0.10 0.11 0.06 0.07 0.26

43.10 37.60 10.60 30.00 20.60 8.26 5.10 1.88 12.10 9.76 19.20 15.80 11.10 15.80 3.74 12.90 5.17 2.15 4.25 19.80 3.72 7.03 4.15 4.71 11.20

85.79 115.65 38.95 123.20 58.11 18.54 22.60 7.74 34.41 41.02 70.44 46.53 36.67 34.50 12.31 16.27 24.97 12.49 10.15 32.79 14.50 20.53 11.48 16.12 53.83

0.63 0.69 1.02 1.29 0.49 0.63 0.81 1.04 0.87 0.84 0.44 0.94 1.37 0.76 1.08 1.11 1.10 2.50 1.60 1.73 3.46 3.00 5.24 2.84 0.86

increase in δ34Spy (Fig. 7) suggests an oxidizing trend. In Interval III, heavy δ34Spy values (N0‰) correlate with an abundance of euhedral pyrite (maximum diameter ~20 μm), which is also indicative of dysoxic/ oxic conditions. The lightest δ34Spy values (~−40‰) correlate with the greatest abundance of framboids, reflecting episodic development of euxinic conditions. Covariation of δ34Spy with Spy exhibits a pattern indicative of mixing of two different pyrite fractions (Fig. 8a). An almost identical pattern was reported for a Permian–Triassic boundary section by Algeo et al. (2008), who inferred mixing of a small but ubiquitous fraction of 34 S-enriched diagenetic pyrite with a variable but generally larger fraction of 34S-depleted syngenetic pyrite (Fig. 8b). In both cases, the endmember compositions are about + 10‰ for the diagenetic endmember and − 45‰ for the syngenetic endmember. The strong correlation of δ34Spy values with framboid abundance provides further evidence that the intervals with low δ34Spy were deposited under reducing conditions. Fe2O3/Al2O3 ratios can serve as a redox proxy because “excess Fe” above background detrital levels is added when syngenetic pyrite forms under euxinic conditions (Lyons et al., 2003; Lyons and Severmann, 2006). The average upper crustal Fe2O3/Al2O3 ratio is 0.38

Fig. 10. ΣREE versus Al2O3 for Xiaojiaba.

(Taylor and McLennan, 1985; McLennan, 2001), but background detrital siliciclastic ratios can vary within the range of ~0.3–0.5 owing to local influences. At Xiaojiaba, most Fe2O3/Al2O3 ratios are close to average crustal values, implying a lack of euxinic conditions, although a few samples in Interval I yield ratios up to 0.92 (Fig. 7), suggesting transient euxinic episodes. Sedimentary Corg/P represents an environmental redox proxy owing to differential rates of remineralization and/or retention of organic carbon and phosphorus under oxic versus anoxic conditions (Algeo and Ingall, 2007). For sediments in which organic matter is derived from marine phytoplankton, the initial molar (C/P)org ratio is ~ 106:1 (Redfield, 1958; Redfield et al., 1963). Sedimentary Corg/P ratios trend lower in oxic facies because organic carbon is lost to aerobic decay whereas remineralized P is adsorbed onto Fe-oxyhydroxides and thus retained in the sediment (Föllmi, 1996; Slomp et al., 1996; Ingall and Jahnke, 1997). In contrast, sedimentary Corg:P ratios trend higher in

Fig. 11. δ13Corg versus TOC for Xiaojiaba. The dataset was fit with a 2nd-order polynomial equation.

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Table 4 Results from SEM pyrite framboid analysis from Xiaojiaba section, South China. Rare = b30 (framboids counted per sample); occasional = 30–49; common 50–98; abundant 99 and above. Sample

Height from the base (m)

Number of framboids

Relative abundance

Standard deviation (μm)

Skewness

Mean (μm)

MFD (μm)

XJB37 XJB36 XJB35 XJB34 XJB33 XJB32 XJB31+ XJB29 XJB28 XJB27 XJB27− XJB26+ XJB26 XJB26− XJB25+ XJB25− XJB24+ XJB24 XJB23+ XJB23− XJB22 XJB20+ XJB20 XJB19 XJB18

17.11 16.97 16.29 15.90 15.46 15.37 15.17 14.94 14.87 14.75 14.59 14.45 14.35 14.10 13.99 13.53 13.47 13.20 13.10 12.75 12.47 12.00 11.80 11.45 11.15

72 80 152 32 1 0 0 0 101 101 100 129 83 81 132 89 101 78 70 208 102 185 185 162 144

Common Common Abundant Occasional Rare No framboid No framboid No framboid Abundant Abundant Abundant Abundant Common Common Abundant Common Abundant Common Common Abundant Abundant Abundant Abundant Abundant Abundant

3.56 2.38 3.48 12.19

3.05 0.66 0.74 0.11

6.57 6.39 8.74 22.28 11

26.91 13.59 22.37 47.38

3.45 4.37 3.64 2.23 4.60 3.49 3.16 3.43 1.93 3.10 3.25 2.19 3.05 3.72 2.51 2.57 3.19

1.24 0.70 1.54 0.33 1.39 2.01 4.40 1.61 0.29 1.41 1.25 0.67 1.48 3.13 1.80 1.68 0.73

7.63 8.30 6.47 5.94 8.22 7.59 6.88 6.50 5.31 6.64 6.92 5.52 6.84 6.10 6.28 6.43 7.86

18.33 19.52 18.75 12.57 22.47 23.52 32.91 20.07 10.67 19.41 17.45 13.78 18.25 24.80 16.86 17.66 20.87

MFD = maximum framboid diameter.

Fig. 12. Relative abundances of framboids near the P–Tr boundary. Note that the pair of volcanic tuffs at the Dalong–Feixianguan formation contact is equivalent to the ‘boundary clay’ at Meishan (Beds 25 and 28) and elsewhere across the South China Craton (cf., Peng et al., 2001).

anoxic facies because P-bearing compounds decay preferentially and the remineralized P is not retained in the sediment. Based on data in Algeo and Ingall (2007), Corg/P ratios of N200 imply anoxia in lower Interval I, ~100 imply suboxic conditions in mid-Interval I, and 30–50

imply oxic conditions in upper Interval I through Interval IV (Fig. 7). This pattern is consistent with other proxies documenting a progressive shift toward less reducing conditions upsection, although the absolute redox levels inferred at a given stratigraphic level are more oxidizing

Fig. 13. Scanning electron microscope (SEM) photomicrographs showing different pyrite morphologies. (a) Abundant pyrite framboids (bright crystals); sample XJB23−. (b) Magnification of detail in (a) showing framboidal pyrite and loose aggregation of pyrite microcrystals. (c) Framboidal and euhedral pyrite (note the magnified image in the white box showing ordered arrangement of pyrite microcrystals); sample XJB25+. (d) Normal framboids and infilled framboids, the latter showing small holes reflecting relict framboidal texture; sample XJB20+. (e) Normal framboids (i), infilled framboids (ii–iii), massive pyrite (iv–v), and euhedral pyrite (vi); sample XJB20+. Pyrite morphologies generally evolve in a sequence from (i) to (vi). (f) Euhedral pyrite, anhedral pyrite, and overgrown framboidal pyrite, the latter showing an outer rim of idiomorphic crystals; sample XJB20. (g) Polyframboids and annular framboids; sample XJB25+. (h) Macroframboids, which are texturally similar to normal framboids but much larger (~50 μm); sample XJB19.

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Fig. 14. Box-and-whisker plots showing framboid size distributions from the upper Dalong Formation to the lower Feixianguan Formation (i.e., the P–Tr boundary transition) at Xiaojiaba. Each box extends from the 25th to the 75th percentile, with median values shown as a vertical line and minimum and maximum values as a horizontal line. The stratigraphic levels of peak euxinia are indicated by arrows.

than for the other proxies considered here. Corg/P ratios tend to integrate redox variation over longer time intervals than other local redox proxies (Algeo and Ingall, 2007), so the Corg/P profile for Xiaojiaba is likely to represent a time-averaged trend. Trace elements such as Mo and U are highly redox sensitive, and reducing bottom waters result in greatly enhanced authigenic uptake by marine sediments (Calvert and Pedersen, 1993; Algeo and Maynard, 2004). Uptake of Uauth usually begins under suboxic conditions corresponding to the Fe(II)–Fe(III) redox boundary, whereas uptake of Moauth requires the presence of H2S and, hence, euxinic conditions (Helz et al., 1996; Zheng et al., 2000; Algeo and Tribovillard, 2009). At Xiaojiaba, the modest enrichments of authigenic Mo and U in Interval I (10–90 ppm and 2–10 ppm, respectively) are indicative of suboxic to episodically euxinic conditions, and the low concentrations in Interval III (b10 ppm and b 2 ppm, respectively) are consistent with generally well-oxygenated conditions (Algeo and Tribovillard, 2009). Synthesizing the observations above, the general pattern is that average redox conditions ameliorated slowly from suboxic (and

occasionally euxinic) in the lower Dalong Formation to dysoxic/oxic interrupted by multiple euxinic episodes in the upper Dalong Formation and lower Feixianguan Formation. We infer that increasing oxygenation through time may have been caused by a gradual shallowing of regional sea level across the South China Craton (e.g., Tong et al., 1999; Yin et al., 2014). On the other hand, the multiple, short euxinic episodes that punctuate the Xiaojiaba record are likely to represent upward movements of the oceanic chemocline (i.e., the upper surface of the oceanic oxygen-minimum zone, or OMZ) during the Changhsingian and early Induan. The frequency of such episodes appears to reflect an increasing destabilization of the chemocline in latest Changhsingian oceanic systems. 5.3. Siliciclastic sediment flux Rare earth elements (REEs) behave coherently during weathering, erosion, and fluvial transportation (McLennan, 1989) and exhibit little or no fractionation during post-depositional processes such as

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Fig. 15. (a) Mean versus standard deviation of pyrite framboid size distributions; and (b) mean versus skewness of framboid size distributions for Xiaojiaba.

diagenesis (Chaudhuri and Cullers, 1979) and metamorphism (Muecke et al., 1979). They are regarded as useful tracers of various geological and oceanographic processes (Piper, 1974; Murray and Leinen, 1993) and are widely used in detrital sediment provenance studies (Piper, 1974; Sholkovitz, 1990; Nath et al., 1997, 2000; Munksgaard et al., 2003; Frimmel, 2009). Total REE concentration (ΣREE) is commonly related to clay-mineral content (Roaldset, 1979; Koporulin et al., 2009; Zhao et al., 2013), which is consistent with the strong relationship between ΣREE and Al2 O 3 at Xiaojiaba (r2 = 0.65, Fig. 10). The Al2O3 profile show peaks in the upper Interval II (at ~12 m) and the lower Interval IV near the ‘boundary clay’ prior to the P–Tr boundary mass extinction (Fig. 9). The trace elements Hf and Nb, which are also indicators of detrital input (Ross and Bustin, 2009), consistently show peaks at the same stratigraphic levels (Fig. 9). Interestingly, the organic-carbon isotope profile also shows two positive excursions at the same positions (Fig. 9), which may reflect the influx of isotopically heavy land plant debris during episodes of intensified subaerial weathering (cf., Kraus et al., 2013). On the basis of ΣREE, Al2O3, Hf and Nb, and organic-carbon isotope data, we infer that both Interval I cherts and Interval III limestones contain only limited clays, but that the upper Interval II and lower Interval IV contain significantly higher clay concentrations (Fig. 9), which may record two pulses of increased terrigenous detrital input just prior to the LPME. Eu/Eu* ratios are indicative of siliciclastic sediment provenance, with higher ratios typically associated with REE contributions from feldspars and, thus, certain types of igneous rocks (Towell et al., 1969; Stolz, 1985; Michard, 1989; Rudnick, 1992; Klinkhammer et al., 1994). Low Eu/Eu* ratios (~0.9–1.0) are typical of shales and, thus, average upper crustal rocks (Taylor and McLennan, 1985; McLennan, 2001). At Xiaojiaba, Eu/Eu* values of ~2–5 in the lower 3 m of Interval I suggest substantial REE contributions from feldspars. These samples are from the Wuchiapingian stage and may represent inputs from eroded Emeishan volcanics (Hou et al., 2012; Shellnutt and Iizuka, 2012). The upper part of Interval I through Interval IV shows Eu/Eu* values of b 1.0, typical of crustal rocks (Fig. 9). The lowest Eu/Eu* values are associated with the

greatest ΣREE concentrations (Fig. 16), demonstrating that Eu/Eu* values b 1.0 are derived from a clay-mineral source. The general upsection increase in detrital siliciclastic content in the Xiaojiaba section was probably related to the latest Changhsingian sea-level fall across the South China Craton (cf., Tong et al., 1999; Yin et al., 2014). It was thus indirectly related to long-term changes in marine productivity and redox conditions within the study section. On the other hand, the abrupt influx of siliciclastic sediment around the LPME horizon (the “boundary clay”) may reflect rapid environmental changes associated with the end-Permian biocrisis. Other studies have documented enhanced subaerial weathering and erosion during and following the LPME (Sephton et al., 2005; Algeo and Twitchett, 2010), probably in response to widespread destruction of terrestrial vegetation (Looy et al., 2001). These terrestrial environmental changes may have had important consequences for the end-Permian marine mass extinction (Algeo et al., 2011b).

Fig. 16. Eu anomaly (Eu/Eu*) versus total REEs (ΣREE) for Xiaojiaba.

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6. Conclusions From evidence at Xiaojiaba, the late Permian paleo-ocean shows decreasing primary productivity, increasing oxygenation, and an elevated terrigenous detrital influx during the latest Changhsingian. These trends may have been caused by regional sea-level shallowing across the South China Craton. However, multiple brief episodes of euxinia, especially just below the latest Permian mass extinction horizon, probably record upward expansion and increased instability of the oceanic chemocline (i.e., the top of the oxygen-minimum zone) at that time. Fluctuations of the oceanic chemocline prior to the LPME and strongly reducing conditions during deposition of the ‘boundary clay’ may have been a major kill-factor in the end-Permian biocrisis. Acknowledgments This research was supported by the National Natural Science Foundation of China (NSFC) through grants 41302021, 41203030 and 41002007, and by the Jiangxi Province Education Project (GJJ13453). Research by TJA is supported by the Sedimentary Geology and Paleobiology program of the U.S. National Science Foundation, the NASA Exobiology program, and by the State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Wuhan (program: GPMR201301). We thank Bryony Caswell and an anonymous reviewer as well as the editor Jasper Knight for their constructive and helpful comments on the original manuscript. We are additionally indebted to Professor Daizhao Chen for providing some data and invaluable suggestions. We thank Fusong Zhang for his help in organic carbon isotopic measurement and Zenghui Guo for his assistance during the lab work. TJA thanks the U.S. National Science Foundation, the NASA Exobiology program, and the State Key Laboratory of Geological Processes and Mineral Resources at the China University of Geosciences-Wuhan (program GPMR201301) for research support. This study is a contribution to IGCP Projects 572 and 630. 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.sedgeo.2014.11.008. These data include Google map of the most important areas described in this article. References Algeo, T.J., Ingall, E.D., 2007. Sedimentary Corg:P ratios, paleocean ventilation, and Phanerozoic atmospheric pO2. Palaeogeography, Palaeoclimatology, Palaeoecology 256, 130–155. Algeo, T.J., Maynard, J.B., 2004. Trace element behavior and redox facies in core shales of Upper Pennsylvanian Kansas-type cyclothems. Chemical Geology 206, 289–318. Algeo, T.J., Tribovillard, N., 2009. Environmental analysis of paleoceanographic systems based on molybdenum–uranium covariation. Chemical Geology 268, 211–225. Algeo, T.J., Twitchett, R.J., 2010. Anomalous Early Triassic sediment fluxes due to elevated weathering rates and their biological consequences. Geology 38, 1023–1026. Algeo, T.J., Hannigan, R., Rowe, H., Brookfield, M., Baud, A., Krystyn, L., Ellwood, B.B., 2007. Sequencing events across the Permian–Triassic boundary, Guryul Ravine (Kashmir, India). Palaeogeography, Palaeoclimatology, Palaeoecology 252, 328–346. Algeo, T.J., Shen, Y., Zhang, T., Lyons, T., Bates, S., Rowe, H., Nguyen, T.K.T., 2008. Association of 34S-depleted pyrite layers with negative carbonate δ13C excursions at the Permian–Triassic boundary: Evidence for upwelling of sulfidic deep–ocean water masses. Geochemistry, Geophysics, Geosystems 9, Q04025. http://dx.doi.org/ 10.1029/2007GC001823. Algeo, T.J., Hinnov, L., Moser, J., Maynard, J.B., Elswick, E., Kuwahara, K., Sano, H., 2010. Changes in productivity and redox conditions in the Panthalassic Ocean during the latest Permian. Geology 38, 187–190. Algeo, T.J., Kuwahara, K., Sano, H., Bates, S., Lyons, T., Elswick, E., Hinnov, L., Ellwood, B., Moser, J., Maynard, J.B., 2011a. Spatial variation in sediment fluxes, redox conditions, and productivity in the Permian–Triassic Panthalassic Ocean. Palaeogeography, Palaeoclimatology, Palaeoecology 308, 65–83. Algeo, T.J., Chen, Z.Q., Fraiser, M.L., Twitchett, R.J., 2011b. Terrestrial-marine teleconnections in the collapse and rebuilding of Early Triassic marine ecosystems. Palaeogeography, Palaeoclimatology, Palaeoecology 308, 1–11.

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