Global and Planetary Change 105 (2013) 7–20

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Large vertical δ13CDIC gradients in Early Triassic seas of the South China craton:Implications for oceanographic changes related to Siberian Traps volcanism

Huyue Song a,b, Jinnan Tong a,⁎, Thomas J. Algeo b,⁎, Micha Horacek c, Haiou Qiu d,Haijun Song a, Li Tian a, Zhong-Qiang Chen a

a State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences, Wuhan 430074, Chinab Department of Geology, University of Cincinnati, Cincinnati, OH 45221-0013, USAc BLT Wieselburg, Research Center Francisco-Josephinum, Rottenhauser Str. 1, 3250 Wieselburg, Austriad School of Material Science and Chemical Engineering, China University of Geosciences, Wuhan 430074, China

⁎ Corresponding authors.E-mail addresses: jntong@cug.edu.cn (J. Tong), thom

0921-8181/$ – see front matter © 2012 Elsevier B.V. Allhttp://dx.doi.org/10.1016/j.gloplacha.2012.10.023

a b s t r a c t

a r t i c l e i n f o

Article history:Received 14 December 2011Received in revised form 18 September 2012Accepted 18 October 2012Available online 2 November 2012

Keywords:carbon isotopesanoxiaoxygen minimum zonechemoclinemass extinctionbiological pumpbiotic recovery

Vertical gradients in the δ13C of seawater dissolved inorganic carbon (Δδ13CDIC) can be estimated forpaleomarine systems based on δ13Ccarb data from sections representing a range of depositional water depths.An analysis of eight Lower Triassic sections from the northern Yangtze Platform and Nanpanjiang Basin,representing water depths of ~50 to 500 m, allowed reconstruction of Δδ13CDIC in Early Triassic seas of theSouth China craton for seven time slices representing four negative (N) and three positive (P) carbon-isotope ex-cursions: 8.5‰ (N1), 5.8‰ (P1), 3.5‰ (N2), 6.5‰ (P2), 7.8‰ (N3),−1.9‰ (P3), and 2.2‰ (N4). These values aremuch larger than vertical δ13CDIC gradients in the modern ocean (~1–3‰) due to intensified stratification andreduced vertical mixing in Early Triassic seas. Peaks in Δδ13CDIC around the PTB (N1) and in the early tomid-Smithian (P2–N3) coincided with episodes of strong climatic warming, reduced marine productivity, andexpanded ocean anoxia. The Dienerian–Smithian boundarymarks the onset of amajor mid-Early Triassic distur-bance, commencing ~1 Myr after the latest Permian mass extinction, that we link to a second eruptive stage ofthe Siberian Traps. Inhospitable oceanic conditions generally persisted until the early Spathian, when strong cli-matic cooling caused re-invigoration of global-ocean circulation, leading to an interval of negative Δδ13CDICvalues and a sharp increase in δ13Ccarb driven by upwelling of nutrient-rich deepwaters. These developmentsmarked the end of the main eruptive stage of the Siberian Traps.

© 2012 Elsevier B.V. All rights reserved.

1. Introduction

The end-Permian mass extinction, the largest biotic crisis of thePhanerozoic (Erwin et al., 2002; Alroy et al., 2008), was accompaniedby a distinct negative shift in marine carbonate carbon isotopes(δ13Ccarb) in Permian–Triassic boundary (PTB) sections globally (Fig. 1;Korte and Kozur, 2010; Luo et al., 2011). Subsequently during the EarlyTriassic, the exogenic (Earth-surface) carbon cycle experienced multipleperturbations that were recorded by large (>3‰) negative and positiveexcursions in δ13Ccarb profiles (Fig. 1; Tong et al., 2002, 2007a; Payne etal., 2004; Corsetti et al., 2005; Zuo et al., 2006; Galfetti et al., 2007a,2007b; Horacek et al., 2007a, 2007b, 2009; Brühwiler et al., 2009;Meyer et al., 2011). These C-isotope excursions have been attributed toa variety of primary controls, including massive volcanism (Payne andKump, 2007), mixing of stratified oceanic watermasses (Horacek et al.,2007b, 2009), and variations in marine productivity (Algeo et al.,2011a), although their causes remain contentious.

as.algeo@uc.edu (T.J. Algeo).

rights reserved.

Insights regarding Early Triassic δ13Ccarb excursions may be gainedthrough analysis of changes in the δ13C of dissolved inorganic carbon(DIC) in contemporaneous seawater. In the modern ocean, the“biological pump” establishes a vertical isotopic gradient in seawaterDIC (Δδ13CDIC) that ranges up to ~3‰ (Fig. 2; Kroopnick, 1985; Hodellet al., 2003), although restricted brackish basins can develop gradi-ents to ~20‰ (Dyrssen et al., 1996; Dyrssen, 1999). For paleomarinesystems, Δδ13CDIC must be estimated from Δδ13Ccarb, i.e., the differ-ence in the carbon isotopic compositions of shallow and deep carbon-ate facies from a single region or basin (Brand et al., 2003). Thisapproach has been used to study seawater DIC in the Ordovician(Munnecke et al., 2003), Silurian (Azmy et al., 1998), Late Cretaceous(Fisher and Arthur, 2002), and at the Cretaceous–Tertiary boundary(Kump, 1991). Recently, Meyer et al. (2011) reported the existenceof a large δ13CDIC gradient in PTB sections of the Nanpanjiang Basinin South China (Fig. 3A), which they attributed to high levels of pri-mary productivity in Early Triassic marine systems.

In this study, we measured δ13Ccarb through extant Lower Triassicstrata from eight sections in South China (Fig. 3A) representing shal-low (b~50 m), intermediate (~50–200 m), and deep (~200–500 m)settings (Fig. 4). Five sections were situated on the northern margin

Fig. 1. Reference δ13Ccarb profiles for Lower Triassic sections from three different regions. Data sources: South China (Payne et al., 2004), Italy (Horacek et al., 2007a), Iran (Horaceket al., 2007b). Abbreviations: Gr. = Griesbachian, Di. = Dienerian, C. = Clarkina, H. = Hindeodus, I. = Isarcicella, Ns. = Neospathodus, LPME = latest Permian mass extinction. N1–N4 and P1–P3 represent negative and positive C-isotope excursions, respectively, that are discussed in the text; N1 is the negative excursion at the Permian–Triassic boundary.

8 H. Song et al. / Global and Planetary Change 105 (2013) 7–20

of the Yangtze Platform, including (in order of increasing waterdepths) Meishan, Daxiakou, Hushan, West Pingdingshan, andSouth Majiashan. Three sections were situated in the NanpanjiangBasin on the southern margin of the South China craton, including(in order of increasing water depths) Dajiang, Mingtang, and Bianyang.The study sections thus provide shallow-to-deep transects across twolargely discrete watermasses of the South China craton during theEarly Triassic (Fig. 3A). We then reconstructed Δδ13CDIC for seven dis-crete time slices within the Early Triassic in order to investigate secularchanges inmarine productivity and carbon cyclingduring and followingthe PTB crisis.

2. Paleogeographic setting and study sections

2.1. General setting

The South China craton was located at low northern paleolatitudes(~15–30°N) in the eastern Paleotethys Ocean during the Early Triassic

Fig. 2. Model of biological pump and vertical δ13CDIC gradient in the modern ocean. Inthis example the vertical gradient, Δδ13C(shallow–deep), is ~3‰. δ13CDIC profile fromKroopnick (1985). Abbreviations: DIC, dissolved inorganic carbon; OC, organic carbon;OMZ, oxygen minimum zone.

(Fig. 3B). The Yangtze Platform formed the central part of this cratonand accumulated shallow-marine carbonate sediments. A rampexisted on its western paleomargin (present-day north), with waterdepths increasing gradually from the margin of the Yangtze Platformwestward into the Qingling Sea, an embayment of the Paleotethys(Fig. 3A, Feng et al., 1997). The Meishan section was located on theramp's upper slope, the Daxiakou and Hushan sections on the ramp'slower slope, and the West Pingdingshan and South Majiashan sec-tions in the deep-water basin further oceanward. On the easternpaleomargin (present-day south) of the South China craton, theNanpanjiang Basin accumulated mainly deep-water radiolariancherts and claystones (He et al., 2005a, 2007; Yin et al., 2007). Isolat-ed carbonate platforms within the Nanpanjiang Basin, of which thelargest was the Great Bank of Guizhou, accumulated shallow-marinefossiliferous wackestones and packstones (Lehrmann et al., 1998;Algeo et al., 2007). Clastic facies were deposited adjacent to exposedland areas on the margins of the South China craton (i.e., the Kamdianand Cathaysian old lands).

2.2. Yangtze Platform ramp sections

The Meishan D section is located in Changxing County, ZhejiangProvince, in east-central China (Fig. 3A). It is the Global Stratotype Sec-tion and Point (GSSP) for the PTB, which is defined by the first appear-ance of the conodontHindeodus parvus at the base of Bed 27c (Yin et al.,2001). Meishan has a well-exposed Lower Triassic succession of earlyGriesbachian through middle Smithian age, as shown by conodont,bivalve, and ammonoid zonation (Tong and Yang, 1998; Chen et al.,2007; Zhang et al., 2007). Griesbachian strata consist mainly black togreenish-gray shale interbedded with marlstone, whereas Dienerianto Smithian strata are composed of thin- to thick-bedded limestone.Water depths at Meishan were mostly in the range of ~100–200 mthroughout the Late Permian–Early Triassic interval, as inferred fromlithology, faunal assemblages, and paleogeographic reconstructions oframp morphology (Fig. 4; Feng et al., 1997; Yin et al., 2001; He et al.,2005b; Chen et al., 2010; Kaiho et al., 2012).

The Daxiakou section is located in the Three Gorges area of HubeiProvince, in central China (Fig. 3A). The Lower Triassic succession

Fig. 3. (A) Early Triassic palaeogeography of South China (modified from Feng et al., 1997). Study sections: 1, Meishan; 2, Daxiakou; 3, Hushan; 4, West Pingdingshan; 5, SouthMajiashan; 6, Dajiang; 7, Mingtang; 8, Bianyang. B) Early Triassic global paleogeography; base map courtesy of R. Blakey (http://jan.ucc.nau.edu/~rcb7/).

9H. Song et al. / Global and Planetary Change 105 (2013) 7–20

comprises beds of Griesbachian through Smithian age, as shown byconodont zonation (Zhao et al., 2005). Here, the lower part of theLower Triassic is dominated by mudstone and thin-bedded limestoneand the upper part by dolostone. These strata were deposited on ashallow carbonate platform that was >500 km distant from sourcesof detrital siliciclastics (Feng et al., 1997; Tong et al., 2002). AtDaxiakou, dominance of the Late Permian conodont fauna by Clarkinais indicative of a deeper carbonate ramp setting with water depths of~200–300 m (Fig. 4; Feng et al., 1997; Lai et al., 2001; Zhao et al.,2005, 2013).

The Hushan section is located about 19 km east of Nanjing, JiangsuProvince, in east-central China (Fig. 3A). The section comprises a con-tinuous succession from the uppermost Permian (Changhsingian)through the lower Olenekian, as established on the basis of conodontand ammonoid biostratigraphy (Chen et al., 1988; Cao and Wang,1993). Here, Lower Triassic strata accumulated on the distal part of

a carbonate ramp at water depths of ~200–300 m (Fig. 4; Feng etal., 1997; Tong and Yin, 2002; Chen et al., 2010).

The West Pingdingshan (WPDS) and South Majiashan (SMJS) sec-tions are located in the northeastern part of Chaohu County in AnhuiProvince, in east-central China (Fig. 3A). The ages of these sectionshave been determined on the basis of conodont and ammonoid biostra-tigraphy (Tong et al., 2005). TheWPDS section, which is a candidate forthe Induan–Olenekian boundary GSSP (Tong et al., 2003), comprises acontinuous succession of basal Griesbachian to lower Smithian age; itconsists of gray shales intercalatedwith thin-beddedmicritic limestonein a rhythmic pattern of meter-scale cyclicity (Li et al., 2007; Guo et al.,2008). The SMJS section comprises most of the Smithian and Spathian,consisting of mudstone with thin limestone interbeds (Tong et al.,2003). In both sections, we sampled only the thin limestone beds withhigh carbonate concentrations. The Lower Triassic strata of this areawere deposited on the distal part of a carbonate ramp or a deep shelf

Fig. 4. Environmental cross-section showing interpretative water depths for study sections on the northern Yangtze Platform (blue arrows) and Nanpanjiang Basin (red arrows)(modified from Chen et al., 2010). Water depth estimates for the Yangtze Platform sections were taken from published sources (as given in text), and those for the NanpanjiangBasin sections are original to this study and based upon sedimentologic and faunal characteristics of these units (see Feng et al., 1997; He et al., 2005a, 2007 for detailed unitdescriptions). Relative differences in water depth within a single basin can be estimated more reliably than differences between basins, so some uncertainty adheres to comparisonsbetween the Yangtze Platform and Nanpanjiang Basin sections.

10 H. Song et al. / Global and Planetary Change 105 (2013) 7–20

and represent water depths of ~300–500 m (Feng et al., 1997; Tong etal., 2002; Chen et al., 2010).

2.3. Nanpanjiang Basin sections

The Dajiang, Mingtang, and Bianyang sections are located on or inproximity to the Great Bank of Guizhou, a carbonate buildup in theNanpanjiang Basin of south-central China (Fig. 3A). The ages of allthree sections have been determined on the basis of conodont and fora-miniferal biostratigraphy (Lehrmann et al., 1998; Payne et al., 2004;Meyer et al., 2011). At Dajiang, Lower Triassic strata consist of micriticlimestone, microbialite, and dolostone deposited in a platform-interiorsetting, representing water depths of meters to, at most, a few tens ofmeters (Fig. 4; cf. Meyer et al., 2011). At Mingtang and Bianyang, thelower Griesbachian is composed of carbonate mudstone and the restof the Lower Triassic succession of medium- to thick-bedded fossilifer-ous limestone with interbeds of shale and carbonate breccia. These sec-tions are located on the platform slope of the Great Bank of Guizhou andrepresent deeper environments. The Mingtang section is located on theupper platform slope, closer to the platform margin than Bianyang,which is located on the mid-platform slope (cf. Meyer et al., 2011).Herein, we estimate water depths as ~50–100 m for Mingtang and~100–200 m for Bianyang (Fig. 4).

3. Methods

Carbon and oxygen isotopic analyses were performed at the StateKey Laboratory of Biogeology and Environmental Geology of theChina University of Geosciences (Wuhan). For each rock sample,~1–2 g of powder was collected using a dental drill, with care takento avoid calcite veins, fossils, and visible diagenetic features. About150–400 μg of powder was placed in a 10 mL Na-glass vial, sealedwith a butyl rubber septum, and reacted with 100% phosphoric acidat 72 °C after flushing with helium. The evolved CO2 gas was analyzedfor δ13C and δ18O using a MAT 253 mass-spectrometer coupled di-rectly to a Finnigan Gasbench II interface (Thermo Scientific). Isotopicvalues are reported as per mille relative to the Vienna Pee Dee belem-nite (V-PDB) standard. Analytical precision was better than ±0.1for δ13C and ±0.1 for δ18O based on replicate analyses of two labora-tory standards (GBW 04416 and GBW 04417). Isotopic data for the

Meishan, Hushan, and Daxiakou sections are from Tong et al. (2007a),and all other data are original to this study.

4. Results

4.1. Standard reference δ13Ccarb curve

Carbonate C-isotope profiles have been generated for Lower Trias-sic sections in many regions, including South China (Tong et al., 2002,2007a; Payne et al., 2004; Zuo et al., 2006; Galfetti et al., 2007a,2007b), Tibet (Galfetti et al., 2007a), Italy (Horacek et al., 2007a),Iran (Horacek et al., 2007b) and Japan (Horacek et al., 2009). We se-lected three records having well-defined profiles and good biostrati-graphic control to serve as reference sections for the present study(Fig. 1). These profiles document four major negative excursions(N1–N4) and three major positive excursions (P1–P3) in δ13Ccarb

through the Lower Triassic. The peak (i.e., minimum or maximumδ13C value) of each excursion is approximately correlative amongthe three reference sections, demonstrating that the excursions rep-resent global changes in the δ13C of seawater DIC. We have also sum-marized C-isotope excursion patterns for 23 Lower Triassic sectionsglobally (Table 1). Most sections record all excursions where strataof the appropriate age are present, and the few instances of “missing”excursions may be due to low sample density or local stratigraphicincompleteness.

The seven excursions utilized in this study are globally correla-tive within existing biostratigraphic constraints. N1 represents a~3–4‰ negative shift that commenced in the late ChanghsingianHindeodus latidentatus/Clarkina yini Zone and peaked rapidly in theearliest Triassic H. parvus Zone (Korte and Kozur, 2010; Luo et al.,2011). P1 represents a ~2–3‰ positive shift that commenced inthe H. parvus Zone and peaked slowly in the latest GriesbachianClarkina carinata Zone. N2 represents a ~1–3‰ negative shift thatcommenced around the Griesbachian–Dienerian boundary and peakedrapidly in the Neospathodus kummeli Zone. P2 represents a ~6–8‰positive shift that commenced in the Ns. kummeli Zone but acceleratedin the Neospathodus dieneri Zone before peaking around the Dienerian–Smithian boundary; it is associatedwith the highest δ13Ccarb values of theLower Triassic (Fig. 1). N3 represents a ~8–11‰ negative shift that com-menced at or just above the Dienerian–Smithian boundary and peaked

Table 1Comparisons of carbon isotopic negative (N) and positive (P) excursions across theTethyan and equatorial Panthalassic oceans during the Early Triassic.

Sections N1 P1 N2 P2 N3 P3 N4 references

NPDS1 Y Y Y Y Y Y ? Tong et al. (2007b)WPDS1 Y Y Y Y Y Y – This studySMJS1 – – – – Y Y Y This studyMeishan2 Y Y Y Y – – – This studyHushan3 Y Y Y Y Y – – This studyDaxiakou4 Y Y Y Y Y – – This studyDawen5 Y Y Y Y Y – – Payne et al. (2004)Dajiang5 Y Y Y Y Y ? – Payne et al. (2004) and this studyGuandao5 – ? ? Y Y Y Y Payne et al. (2004)Laolaicao5 Y Y ? Y Y Y ? Meyer et al. (2011)Mingtang5 – ? Y Y Y Y Y This studyBianyang5 – ? ? Y Y Y Y This studyLekang6 Y Y ? Y Y Y ? Tong et al. (2007b)Zuodeng7 Y Y Y Y Y Y Y Tong et al. (2007b)Jinya8 Y Y Y Y Y Y Y Galfetti et al. (2007a)Tulong9 Y – – – Y Y Y Brühwiler et al. (2009)Pufels10 Y Y Y Y Y – – Horacek et al. (2007a)Uomo10 Y Y Y Y Y Y Y Horacek et al. (2007a)Abadeh11 Y Y Y Y Y Y – Horacek et al. (2007b)Amol12 Y Y ? Y Y Y Y Horacek et al. (2007b)Zal13 Y Y Y Y Y Y Y Horacek et al. (2007b)Losar14 Y Y – Y Y Y Y Galfetti et al. (2007b)Kamura15 Y Y ? Y Y Y – Horacek et al. (2009)

“?” means “not significant”, “–” means “no available data”. Section location: 1. Chaohu,South China (NPDS = North Pingdingshan; WPDS= West Pingdingshan; SMJS = SouthMajiashan); 2. Changxing, South China; 3. Nanjing, South China; 4. Xingshan, SouthChina; 5. Luodian, South China; 6. Wangmo, South China; 7. Tiandong, South China; 8.Hechi, South China; 9. Nielamu, South Tibet; 10. Southern Tyrol/Alto Adige, Italy; 11.Isfahan, Iran; 12. Alborz Mountain Range, Iran; 13. Tabriz, Iran; 14. Himachal Pradesh,India; 15. Kyushu, Japan.

11H. Song et al. / Global and Planetary Change 105 (2013) 7–20

slowly, reaching aminimumvalue at or just below the Smithian–Spathianboundary; it is generally the largest C-isotopic shift in Lower Triassic sec-tions (Fig. 1). P3 represents a ~3 to>5‰positive shift that commenced inthe latest SmithianNeospathodus waageni Zone and peaked rapidly in theearliest Spathian Neospathodus crassatus Zone. N4 represents a ~3–6‰negative shift that commenced in the early Spathian N. crassatus Zoneand peaked slowly in the mid-Spathian Neospathodus homeri Zone.Above N4, most sections show a shift toward somewhat heavierδ13Ccarb values (potentially allowing recognition of a P4 horizon) beforethe onset of relatively uniform values in theMiddle Triassic (Tong et al.,2007a).

Recent radiometric dating studies (e.g., Mundil et al., 2010; Shen etal., 2011) have provided the basis for refinement of the Late Permian–Early Triassic timescale. The durations of the four substages of the EarlyTriassic are estimated as follows: Griesbachian, 730 kyr; Dienerian,270 kyr; Smithian, 600 kyr; Spathian, 3.4 Myr (Algeo et al., 2013).With-in this age framework, the N1 to P3 shifts each occurred rapidly, over in-tervals ranging from ~200 to 500 kyr (Fig. 1). The N4 excursion occurredmore slowly, over an interval of ~1.5–2.0 Myr, and thus appears to rep-resent a slowing of C-cycle perturbations and an incipient stabilizationfollowing the PTB crisis. Differences in the shapes of the C-isotope pro-files of the reference sections (and of the study sections below) arestrongly influenced by section-specific, non-linear variation in sedimen-tation rates and, hence, probably do not yield meaningful informationregarding spatial non-uniformity in rates of change of δ13CDIC.

4.2. Study section δ13Ccarb profiles

Sections from the northern Yangtze Platform (Fig. 5) exhibit all ofthe excursions (N1–N4, P1–P3) defined in the reference sections(Fig. 1), although section-specific differences exist. N1, P1, and N2are well-defined in the shallow-water sections (Meishan andDaxiakou), but N2 is weakly developed or almost non-existent inthe deep-water sections (Hushan and WPDS). N1, P1, and N2 are as-sociated with distinctly lighter δ13Ccarb values in the deep-water

sections relative to the shallow-water sections. This depth-relatedpattern does not hold for P2 because Hushan records a heavierδ13Ccarb composition than either Meishan or Daxiakou. However, N3yields distinctly lighter δ13Ccarb in the deep-water SMJS section thanin the shallower Daxiakou section. P3 and N4 are present only atSMJS, so depth-related patterns cannot be assessed. The deep-watersections are characterized by a larger overall range of C-isotopicvalues (based on P2 minus N1) than the shallow-water sections:~13‰ at Hushan and ~10‰ at WPDS versus ~7‰ at Meishan and~4‰ at Daxiakou (Fig. 5). SMJS, a deep-water section representing ayounger stratigraphic interval, also shows a large C-isotope range(~12‰), although this section lacks a shallow-water equivalent inthe northern Yangtze Platform region.

Sections from the Nanpanjiang Basin (Fig. 6) also exhibit all of theexcursions (N1–N4, P1–P3) defined in the reference sections (Fig. 1),and Mingtang and possibly Bianyang record an additional positiveexcursion (“P4”) around the Lower/Middle Triassic boundary. N1was not obtained at Mingtang or Bianyang due to the dominance ofmudstone facies around the base of the Griesbachian; at Dajiang,N1 is characterized by a relatively heavy δ13C value (~0‰). N2 iswell-defined in the shallow-water Dajiang and intermediate-waterMingtang sections but appears to be poorly developed in thedeep-water Bianyang section. P2, N3, P3, and N4 are associated withdistinctly lighter δ13Ccarb values in the deep-water section relativeto the shallower sections, although the depth-related difference inδ13C is sharply reduced for N4 relative to the earlier excursions.Unlike for the northern Yangtze Platform region, the overall rangeof C-isotopic values (based on P2 minus N3) is similar for all threeNanpanjiang Basin sections regardless of water depth: ~8‰ atDajiang, ~6‰ at Mingtang, and ~7‰ at Bianyang (Fig. 6). Since N3and N1 have nearly identical δ13Ccarb values at Dajiang, these rangesmay be directly comparable to estimates for the northern YangtzePlatform region (see above).

We regard the C-isotope profiles of the 8 study sections asrepresenting in-situ, primary or near-primary marine carbonate δ13Ccompositions. Small sample size precluded detailed petrographic analy-sis, but we note that (1) most samples consisted of micritic limestone,(2) the study sections are overwhelming carbonate in composition(averaging 70 to 95% in different sections; see Supplement), so only afraction of the carbonate can be cement, and (3) there are substantialdifferences in the C-isotopic composition of the shallow-water versusdeep-water sections of our study, so the carbonate in the deepwatersections cannot be derived mainly from shallow facies via hemipelagicsedimentation. Although the δ13C composition of marine carbonatescan be altered during diagenetic stabilization, e.g., when water/rock ra-tios are high (Algeo et al., 1992) or when large quantities of organicmatter or methane have been oxidized (Irwin et al., 1977), many car-bonate sediments undergo stabilization in a relatively closed diageneticsystem (i.e., lowwater/rock ratios) and in the (near-)absence of DIC de-rived from other sources, resulting in little or no change in bulk-rockδ13Ccarb values (Scholle and Arthur, 1980; Marshall, 1992). The primarycharacter of the δ13Ccarb profiles of the 8 study sections is shown by(1) their similarity to other age-equivalent C-isotope profiles (Fig. 1);(2) lack of covariation between C- and O-isotopes such as that devel-oped during meteoric or burial diagenesis (Algeo et al., 1992; Wenzel,2000; see Supplement); (3) low Mn/Sr ratios (b10; Tong et al., 2002),which are indicative of early diagenetic stabilization of the host carbon-ate (Barnaby and Rimstidt, 1989); and (4) relatively heavy δ18O values(between −4‰ and −8‰ for >95% of study samples; see Supple-ment), which are also indicative of early host carbonate stabilization(cf. Algeo et al., 1992; Kaufman and Knoll, 1995). The only obviouslyaltered samples are a subset (n=10) from the Mingtang section thatexhibit δ18O valuesb−11‰ (Fig. 6). However, the δ13C values ofthese samples are similar to the remainder of the Mingtang section, soprobably represent a near-primary marine composition despite alter-ation of their O-isotopic signature during late burial recrystallization.

Fig. 5. Carbon isotope profiles for shallow-water (left) to deep-water sections (right) from the northern Yangtze Platform. Abbreviations: U.P. = Upper Permian. C. = Clarkina,H. =Hindeodus, I. = Isarcicella,Ns. =Neospathodus. The C-isotope profiles for theMeishan, Hushan, andDaxiakou sectionswere previously published in Tong et al. (2007a); all other dataare original to this study.

12 H. Song et al. / Global and Planetary Change 105 (2013) 7–20

4.3. Reconstruction of vertical seawater δ13CDIC gradients (Δδ13CDIC)

The C-isotope profiles of the study sections (Figs. 5 and 6)were usedto reconstruct vertical gradients in seawater δ13CDIC in Early Triassicseas of the South China craton. For each time slice (N1–N4, P1–P3),the δ13Ccarb values of all study sections containing that horizon wereplotted together as a function of depositional water depth (Fig. 7A).TheΔδ13CDIC for a given time slice was then calculated as themaximumdifference in δ13Ccarb between shallow and deep study sections (i.e.,Δδ13C(shallow–deep)). Δδ13CDIC was largest for N1 (8.5‰), diminishedsomewhat for P1 (5.8‰) and N2 (3.5‰), increased again for P2(6.5‰) and N3 (7.8‰), and then was sharply reduced for P3 (−1.9‰)and N4 (2.2‰). Meyer et al. (2011) reported Δδ13CDIC for N1 (~7‰),P2 (~5‰), N3 (~4‰), andN4 (≤~2‰), values that are somewhat small-er than ours because their analysis was limited to relatively shallow(≤200 m) sections from the Nanpanjiang Basin whereas ours includesdeeper sections from the northern Yangtze Platform (Fig. 4). Combiningthe Δδ13CDIC profiles for the seven time slices in a single graph (Fig. 7B)reveals that (1) the profiles for positive and negative excursions aresimilar in shape other than being offset by 4 to 8‰, and (2) both setsof profiles show limited vertical δ13CDIC variation above a water depthof 250 m and a relatively abrupt shift toward lighter δ13CDIC valuesbelow that level. We interpret the abrupt shift in δ13CDIC at a waterdepth of ~250 m as the top of the permanent thermocline in EarlyTriassic seas of the South China craton.

Our ability to reconstruct Early Triassic seawater δ13CDIC gradientsis spatially constrained. The foregoing analysis provides estimates of

δ13CDIC over a range of water depths from ~50 to 500 m (Fig. 4),reflecting conditions in the ocean-surface (~b100 m) and thermo-cline watermasses (~100–1000 m). Assessment of Early Triassicδ13CDIC at water depths >500 m is likely to be difficult or impossible,because (1) the few surviving deep-ocean sections contain no car-bonate (e.g., Algeo et al., 2010, 2011b), and (2) the carbonate com-pensation depth, below which no carbonate sediment is preserved,may have been relatively shallow prior to the diversification ofcoccoliths in the late Triassic to Jurassic (de Vargas et al., 2007).Also, our analysis ideally would have been undertaken separatelyfor the two study regions, but the limited number of sections foreach region (3 to 5) and the limited range of water depths represent-ed by the Nanpanjiang Basin sections (b200 m; Fig. 4) precludedreconstruction of separate regional Δδ13CDIC gradients. However,where there is data overlap for the two study regions, as at waterdepths of ~150 m for the P1 and P2 excursions (Fig. 7A), the observedδ13Ccarb values are similar—validating our pooling of δ13Ccarb data forthe two study regions.

One point should be emphasized: Δδ13CDIC is a separate proxyfrom δ13Ccarb and can potentially vary independently of the latter.We illustrate this point by comparing secular variation in Δδ13CDIC(Fig. 8A) with shallow- and deep-water δ13Ccarb curves for EarlyTriassic seas of the South China craton (Fig. 8B). Δδ13CDIC (1) risessharply around the PTB before gradually declining through theGriesbachian–Dienerian, (2) rises sharply a second time around theDienerian–Smithian boundary, peaking in the late Smithian, (3) fallsrapidly to negative values in the early Spathian, and (4) then shifts

Fig. 6. Carbon isotope profiles for shallow-water (left) to deep-water sections (right) of the Nanpanjiang Basin, southwest China. Mingtang samples shown in red may be diage-netically altered, based on highly negative δ18O values (see Fig. 7). Abbreviations: M. = Middle, U. = Upper, H. = Hindeodus, I. = Isarcicella, Ns. = Neospathodus.

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slowly to small positive values (b2.5‰) during the mid- to lateSpathian. Themajor changes in Δδ13CDIC are associated with extremesin δ13Ccarb but do not show a consistent relationship to minimum andmaximum δ13Ccarb values. Thus, the first increase in Δδ13CDIC (at thePTB) is associated with a major negative δ13Ccarb excursion (N1),whereas the second increase in Δδ13CDIC (at the Dienerian–Smithianboundary) is associated with a major positive δ13Ccarb excursion(P2; Fig. 8). However, our datasetmay illustrate a tendency forΔδ13CDICto reach peak values in association with negative δ13Ccarb excursions(e.g., for N1, N3, and N4). We conclude that these two proxies, whiletheoretically independent, may in fact tend to covary.

5. Discussion

5.1. Controls on vertical δ13CDIC gradients

The magnitude of vertical gradients in seawater δ13CDIC is con-trolled by an interplay between rates of marine productivity and up-welling. The “biological pump” creates vertical δ13CDIC gradientsthrough the preferential export of isotopically light carbon (12C)from the ocean surface layer in the form of sinking organic matter(Falkowski et al., 1998; Thomas et al., 2004). Upwelling (or verticalmixing) returns this isotopically light carbon to the ocean-surface

Fig. 7. Vertical δ13C gradients in Early Triassic seas of the South China craton. δ13C values are plotted against the mid-range estimate of water depth for each study section (cf. Fig. 4).(A) For each time slice (N1–N4, P1–P3), the maximum difference between shallow and deep sections is given as Δδ13C(shallow–deep) at the bottom of the profile. Note the pattern ofsecular variation in Δδ13C(shallow–deep), with maximum values at N1 (earliest Griesbachian) and N3 (mid-Smithian). (B) Summary of all seven Δδ13Ccarb profiles in A. This plot revealsan abrupt decrease in δ13Ccarb at ~250 m water depth, which we interpret to represent the top of the permanent pycnocline in the South China region. Note that the study sectionsof Meyer et al. (2011) were located mostly above 250 m and, hence, do not capture the full range of vertical variation in δ13Ccarb present in Early Triassic seas of the South Chinacraton.

14 H. Song et al. / Global and Planetary Change 105 (2013) 7–20

layer and thus reduces the vertical δ13CDIC gradient created by the biolog-ical pump. Because rates of upwelling and primary productivity are oftenclosely linked in the global ocean, their effects on Δδ13CDIC partiallycancel out, although upwelling regions generally exhibit somewhathigher Δδ13CDIC than the rest of the ocean on average (Kroopnick,

Fig. 8. Secular variation in (A) vertical δ13CDIC gradient in Early Triassic seas of South Chinabased on Fig. 7 with additional data points interpolated between the N1–N4 and P1–P3 nand maxima in the study sections (Figs. 5 and 6) and serve here to demonstrate the scomposite of the profiles for Dajiang and the nearby Guandao section (Tong et al., 2007acurve minus Δδ13CDIC (from A). Note the generally large values of Δδ13CDIC throughout the Grslow changes in Δδ13CDIC with time, and the early Spathian interval during which Δδ13CDIC is(from Algeo et al., 2011a). (D) Summary of interpretations and hypotheses regarding oceaniabbreviations as in Fig. 1.

1985). However, even large changes in primary productivity have onlylimited effects on Δδ13CDIC values in the open ocean: Δδ13CDIC rangesfrom b1‰ in low-productivity oceanic gyres (b25 g m−2 yr−1) to amaximum of ~3‰ in some high-productivity upwelling zones such asthe western Arabian Sea (~200–400 g m−2 yr−1; Fig. 9A; McCreary et

craton, and (B) modeled shallow- and deep-water δ13CDIC. In A, the Δδ13CDIC curve isodes. The additional points were estimated from correlations of local δ13Ccarb minimalowness of secular variation in Δδ13CDIC. In B, the shallow-water δ13Ccarb curve is a, their figure 9); the deep-water δ13Ccarb curve was calculated as the shallow-wateriesbachian to early Spathian, Δδ13CDIC maxima in association with N1 and N3, relativelynegative (i.e., δ13Cdeep>δ13Cshallow; hachured field). (C) Chemical weathering intensityc conditions during the Early Triassic; see text for discussion. ST = Siberian Traps; other

Fig. 9. (A) Productivity vs.Δδ13CDIC, and (B)watermass stratification vs.Δδ13CDIC formod-ern marine systems. Productivity exhibits no systematic relationship to Δδ13CDIC and ex-erts a weaker influence on Δδ13CDIC than water-column stratification (as proxied byΔσt(deep–shallow)). Water-column stratification exhibits a strong relationship to Δδ13CDICand is the only parameter that can produce the large Δδ13CDIC values (3.5–8.5‰) charac-teristic of Early Triassic seas.

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al., 1996; Peeters et al., 2002). In contrast, variation in rates of verticalmixing can have a much larger influence on Δδ13CDIC. In basins withlimited vertical mixing, Δδ13CDIC can range up to ~7‰ (as in the2200-m-deep Black Sea) or ~20‰ (as in the 180-m-deep FramvarenFjord; Deuser, 1970; Dyrssen et al., 1996; Dyrssen, 1999). These steepΔδ13CDIC gradients exist despite low levels of primary productivity(and, hence, a weak biological pump): ~1–10 g m−2 yr−1 in the BlackSea and ~12–24 g m−2 yr−1 in Framvaren Fjord (Algeo and Lyons,2006). The critical factor is the degree of water-column stratification,which can be expressed as Δσt(deep–shallow), or the difference in densitiesbetween the deep and surface watermasses (n.b., σt is a function ofwatermass salinity and temperature; Chester, 1990).Watermass stratifi-cation is much stronger for the Black Sea (Δσt=4) and Framvaren Fjord(Δσt=11; Algeo et al., 2008a) than for open-ocean settings (Δσt~1–3;Fig. 9B).

Marine productivity is inferred to have increased during the EarlyTriassic as a consequence of (after) effects of the PTB crisis (Algeo etal., 2011a; Meyer et al., 2011). Increases in marine productivity re-quire a source of excess nutrients, either via oceanic upwelling or en-hanced riverine delivery from continental areas. Because EarlyTriassic seas were characterized by intensified water-column stratifi-cation, upwelling is an unlikely source of excess nutrients, but en-hanced subaerial weathering and riverine delivery of nutrients is

not. A major soil erosion event at the PTB has been recognized fromincreased concentrations of dibenzofuran and related soil-derivedorganic compounds just below the marine mass extinction horizonin sections in Italy (Sephton et al., 2005), Greenland (Fenton et al.,2007), Australia (Grice et al., 2007) and South China (Grice et al.,2005; Xie et al., 2005, 2007; Wang and Visscher, 2007). Subsequently,nutrient fluxes to the marine environment may have depended on el-evated rates of bedrock erosion (Algeo and Twitchett, 2010; Algeo etal., 2011a), which developed during the Early Triassic as a conse-quence of (1) faster chemical reaction rates catalyzed by climaticwarming and increased soil acidity related to volcanic emissions ofSO2 (Wignall, 2001; cf. Self et al., 2008), and (2) landscape distur-bance related to replacement of mature conifer forests by disastervegetation (Looy et al., 1999, 2001; Hermann et al., 2011).

The degree of stratification of Early Triassic oceans is inferred tohave increased as a consequence of strong climate warming(Horacek et al., 2007b, 2009). Warming across the PTB, possibly byas much as 8–12 °C, has been demonstrated by studies of δ18O in bio-genic calcite (Korte et al., 2005a, 2005b; Kearsey et al., 2009) and co-nodont apatite (Joachimski et al., 2012), and recent work indicatesthat strong warming persisted through the Smithian (Sun et al.,2012). Early Triassic warming was due to release of volcanic CO2

from the Siberian Traps and thermogenic methane related to mag-matic intrusions into West Siberian coal basins (Retallack, 1999;Retallack and Jahren, 2008). Increases in both factors (i.e., marine pro-ductivity andwater-column stratification) contributed to the expansionof oceanic oxygen-minimum zones (Algeo et al., 2010, 2011b), to thewidespread development of suboxic to anoxic conditions on EarlyTriassic seafloors (Isozaki, 1997; Wignall and Twitchett, 2002; Bondand Wignall, 2010), and to the demise of progressively shallower-dwelling planktic taxa (Isozaki et al., 2007; Shen et al., 2012).

Relationships in modern marine systems discussed above providekey constraints on interpretations ofΔδ13CDIC gradients in Early Triassicmarine systems. Δδ13CDIC estimates for the Griesbachian–Smithian(Figs. 7 and 8) are large (up to 8.5‰) compared to typical Δδ13CDIC gra-dients in themodern ocean (~1–3‰; Fig. 9; Kroopnick, 1985). This con-trast is even more pronounced given that the Δδ13CDIC gradient in themodern ocean is steepest at 500–1000 m,within the core of the oxygenminimum zone (OMZ) (Kroopnick, 1985), whereas our Δδ13CDIC esti-mates for Early Triassic seas represent a depth range of 0–500 m orless (Fig. 7). Although the Δδ13CDIC of Early Triassic seas may have in-creased in response to changes in rates of both marine productivityand vertical mixing, it is clear from the examples above that the latterfactor must have been the dominant influence, and that the Early Trias-sic must have been an interval of extraordinarily reduced verticalmixing in the global ocean. We tentatively infer that, during intervalsof peak water-column stratification around the PTB and Dienerian–Smithian boundary, Early Triassic seas of the South China cratonexhibited a degree of stratification similar to that of the modern BlackSea (Δσt~4) based on nearly equivalent Δδ13CDIC gradients (Fig. 9B).

The reversal in Δδ13CDIC observed during the early to mid-Spathian(Fig. 8) deserves special comment. In the modern ocean, downwellingregions such as the North Atlantic can exhibit higher δ13CDIC values inthe deep watermass relative to the ocean-surface layer, although thiseffect is generally small (b0.5‰; Kroopnick, 1985; Labeyrie et al.,1992). Epicontinental seas can exhibit substantially greater variationin δ13CDIC at short lateral and vertical scales owing to local variabilityin rates of productivity or remineralization of organic matter or tolocal watermass circulation patterns (Holmden et al., 1998; Panchuket al., 2005; Immenhauser et al., 2008). Upwelling of watermassescontaining 13C-depleted DIC (Immenhauser et al., 2003, 2008) or highdiffusive fluxes of 13C-depleted DIC from remineralized organic matterin the sediment (as in modern Florida Bay; Patterson and Walter,1994) can potentially lead to an inversion of marineΔδ13CDIC gradients.Owing to paleogeographic factors, the western paleomargin of theYangtze Platform was a region of upwelling through most of the

16 H. Song et al. / Global and Planetary Change 105 (2013) 7–20

Paleozoic (e.g., Cai et al., 2007; Jiang et al., 2007; Zhou and Jiang, 2009;Luo and He, 2011). We hypothesize that a general re-invigoration ofglobal-ocean circulation during the early Spathian caused intensifica-tion of upwelling on this margin and, probably, in other upwelling re-gions globally, and that the evidence for this re-invigorated circulationis inversion of Δδ13CDIC gradients in Early Triassic seas.

5.2. Interpretation of Early Triassic δ13Ccarb excursions

Carbonate C-isotope records are subject to many influences(Kump and Arthur, 1999; Berner, 2002). When considered from anocean-surface perspective, positive δ13Ccarb excursions are commonlyassociated with increased productivity or organic carbon burial(Arthur et al., 1988; Kump and Arthur, 1999), or with increased stor-age of DIC in the deep ocean related to intensified stratification(Wenzel and Joachimski, 1996; Horacek et al., 2007b, 2009). Negativeδ13Ccarb excursions are commonly associated with reduced marineproductivity or organic carbon burial, as during mass extinctionevents such as those at the PTB (Corsetti et al., 2005; Luo et al.,2011) and the Cretaceous–Paleogene boundary (Zachos et al.,1989). However, reduced primary productivity (Kump, 1991) is notthe only mechanism for such excursions, and concurrent increasesin volcanogenic or seafloor-hydrate methane release can generatelarge quantities of 12C-enriched DIC (Dickens et al., 1997; Berner,2002; McElwain et al., 2005). One other possibility to generate a neg-ative excursion is overturn of a stratified ocean, bringing 13C-depletedDIC back into the ocean-surface layer (Wenzel and Joachimski, 1996;Horacek et al., 2009). With regard to the PTB, the N1 excursion hasbeen attributed to methane release, with secondary contributionsfrom biomass reduction and volcanic gas emissions (Berner, 2002),as well as to oxidation of large amounts of soil organic matter(Algeo et al., 2011a), while the large positive δ13Ccarb excursions ofthe Early Triassic (especially P2) have been attributed to volcanicCO2 emissions (Payne and Kump, 2007). The large Early Triassicδ13Ccarb excursions have also been attributed to variations in marineproductivity (Algeo et al., 2011a; Meyer et al., 2011) and rates of ver-tical overturn linked to water-column stratification (Horacek et al.,2007b, 2009).

Our reconstructedΔδ13CDIC records for the Early Triassic (Fig. 8) pro-vide constraints on and insights into the range of possible causes forlarge δ13Ccarb excursions in Early Triassic marine systems. Given theapparent close relationship of δ13Ccarb to Δδ13CDIC (see Section 4.3),viable mechanisms must account for concurrent changes in both pa-rameters (per Occam's razor). Several proposed explanations for largeexcursions in δ13Ccarb are inadequate as they cannot produce observedrelationships to Δδ13CDIC. This conclusion applies to any mechanismthat relies primarily on the addition of isotopically light DIC to theocean-surface layer, because such processes will concurrently reduceδ13Ccarb and Δδ13CDIC (n.b., adding 12C to the ocean surface would re-duce the vertical δ13CDIC gradient), a pattern at odds with observationsof negative excursions being associated with Δδ13CDIC maxima (seeSection 4.3). This consideration casts doubt upon methane release, soilorganic matter oxidation, and volcanic CO2 emissions as the primarycauses of large δ13Ccarb excursions during the Early Triassic (n.b., thisargument is not evidence against the occurrence of such processes,just their importance as agents of large δ13Ccarb excursions). Further,any mechanism that relies primarily on additions (or removals) ofisotopically distinct DIC to the ocean would have to operate for in-tervals as long as the durations of Δδ13CDIC excursions (i.e., up to1 Myr), because geologically instantaneous injections of isotopicallydistinct DIC would have been mixed through the whole ocean at time-scales of a few millennia, resulting in a shift in δ13Ccarb but no changein Δδ13CDIC (since the isotopic composition of the shallow and deepwatermasses would have been affected equally by such infusions atlonger timescales).

The remaining mechanisms for generating large δ13Ccarb excursionsin Early Triassic seas are related to changes in primary productivity andwater-column stratification (see Section 5.1). As a general mechanism,primary productivity changes fail on two counts. First, increases in pri-mary productivity should result in a positive shift in δ13Ccarb and a con-current steepening of the vertical δ13CDIC gradient, a pattern that isconsistent with the P2 horizon but not with the N1 and N3 horizonswhere Δδ13CDIC is even larger than for P2 (Fig. 8). Second, as discussedin Section 5.1, productivity changes alone cannot produce Δδ13CDICvalues >~3‰. As in the case of Δδ13CDIC (Section 5.1), it appears thatchanges in water-column stratification are the most likely mechanismto account for large excursions in δ13Ccarb during the Early Triassic.Intensified stratification should result in an increase in Δδ13CDIC and aconcurrent shift toward lower δ13Ccarb (owing to reduced productivityrates in the absence of nutrient resupply through upwelling), as ob-served for horizons N1, N3, and N4 (Fig. 8). Reduced stratificationwould have the opposite effect, leading to upwelling, enhanced primaryproductivity, lower Δδ13CDIC, and higher δ13Ccarb values, as observed forhorizon P3 (see also Section 5.1). The principal exception to this scenar-io is horizon P2, which is characterized by high δ13Ccarb in conjunctionwith a Δδ13CDIC maximum (Fig. 8). Although tentative, the implicationof this pattern is that strongwater-column stratification (due to climat-ic warming) developed simultaneously with high marine productivityand/or major injections of 13C-enriched volcanic CO2 (e.g., Payne andKump, 2007; Meyer et al., 2011). We would like to emphasize that theinterpretations above do not preclude other processes (e.g., methanerelease or soil organic matter oxidation) that might have had a limitedspatial or temporal influence on δ13Ccarb.

5.3. Global significance of Early Triassic Δδ13CDIC and δ13Ccarb records

The foregoing analysis of controls on Δδ13CDIC and δ13Ccarb allowsus to develop an integrated model for oceanographic changes duringthe Early Triassic. Whereas Meyer et al. (2011) identified only a singleΔδ13CDIC maximum (N1) because of the limited number of time slicesused in their study, we can document twomaxima, the first at the PTB(N1) and the second within the Smithian (P2–N3; Fig. 8). The N1 ho-rizon is associated with low δ13Ccarb and a spike in chemicalweathering rates (Algeo and Twitchett, 2010; Algeo et al., 2011a),consistent with strong warming, enhanced water-column stratifica-tion, and a pronounced decline in marine productivity (cf. Algeo etal., 2013). It has been widely linked to the onset of major SiberianTraps volcanism (e.g., Reichow et al., 2009; Korte and Kozur, 2010).The P1–N2 horizons are associated with a gradual decline in the in-tensity of stratification; they represent minor events of uncertainorigin. The P2–N3 horizons are associated with an initial increase inδ13Ccarb at the Dienerian–Smithian boundary, followed by a longdecline toward a δ13Ccarb minimum in the late Smithian (Fig. 8).This pattern is consistent with either an initial increase in marine pro-ductivity or large-scale input of 13C-enriched DIC from volcanic(?)sources (e.g., Payne and Kump, 2007), followed by a protracted declinein marine productivity. The productivity decline during the Smithianwas probably the result of persistently intense water-column stratifica-tion, as reflected in high Δδ13CDIC values for both P2 and N3. An earlySmithian increase in chemical weathering rates (Fig. 8) suggests a coe-val disturbance of terrestrial environments, possibly due to a sharedcausality with the N1 event. The cause of the Dienerian–Smithianboundary crisis has not been identified to date, but we hypothesizethat it records a second major eruptive stage of the Siberian Traps.Both crises (N1 and P2–N3) coincided with episodes of expansion ofoceanic anoxia: (1) in the latest Changhsingian and early Griesbachian(Grice et al., 2005; Algeo et al., 2007, 2008b), and (2) in the lateDienerian to mid-Smithian interval (Galfetti et al., 2007a, 2007b; Li etal., 2013).

A marked change in the character of all records is evident in theearly Spathian (Fig. 8). Δδ13CDIC abruptly shifts to negative values,

17H. Song et al. / Global and Planetary Change 105 (2013) 7–20

δ13Ccarb shows a rapid trend toward higher values, and chemicalweathering rates decline (Fig. 8). We interpret this pattern to indicatethat ocean circulation was re-invigorated, bringing 13C-depleted DICinto the ocean-surface layer (hence negative Δδ13CDIC), and stimulat-ing marine productivity through upwelling of nutrients (hence higherδ13Ccarb). Lower chemical weathering rates are an indication ofcooling, reduced acid rainfall, and/or fewer disturbances of terrestrialenvironments. Collectively, these records suggest that a strong globalcooling took place during the early Spathian, a climatic event that hasbeen independently documented from biogeographic data (Brayardet al., 2007, 2009a; Galfetti et al., 2007a, 2007c). We hypothesizethat these far-reaching climatic and oceanographic changes reflecttermination of the main stage of Siberian Traps volcanism (Fig. 8).However, the pattern of Δδ13CDIC and δ13Ccarb variation in the mid-lateSpathian (N4) might be consistent with an additional minor eruptivestage (Fig. 8).

5.4. Relationship of oceanographic changes to the Early Triassic marinebiotic recovery

The marine biotic recovery after the latest Permian mass extinc-tion took up to several million years, although the exact definitionof “recovery” is debated (Hallam, 1991; Twitchett et al., 2004;Fraiser and Bottjer, 2007; Tong et al., 2007b; Bottjer et al., 2008;Chen et al., 2009). Some clades recovered relatively rapidly, e.g.,conodonts, ammonoids, and foraminifera, all of which exhibit a diver-sity peak around the Dienerian/Smithian boundary or during theearly Smithian followed by a second major decline (Orchard, 2007;Brayard et al., 2009b; Stanley, 2009; Song et al., 2011). Other marineinvertebrate groups exhibit little recovery during the Griesbachian–Smithian interval, a modest recovery during the Spathian, and astronger recovery during the early Middle Triassic (Erwin et al.,2002; Bottjer et al., 2008). The duration of the recovery of marine eco-systems following the end of the Permian is in marked contrast to therelatively rapid recovery that occurred after mass extinction eventssuch as that at the Cretaceous–Paleogene boundary (Harries, 1999;Beerling et al., 2001). A number of hypotheses have been offered toaccount for the long duration of the post-Permian recovery, includinga protracted reintegration of decimated marine ecosystems (Erwin,1994; Erwin et al., 2002; Chen and Benton, 2012), persistently inhos-pitable conditions in marine environments (Hallam, 1991; Twitchettet al., 2004; Takahashi et al., 2009), or repeated environmental distur-bances from episodic volcanism (Payne and Kump, 2007; Retallack etal., 2011), episodic expansion of the oceanic oxygen-minimum zone(Algeo et al., 2010, 2011b), or episodic mixing of stratified oceans(Horacek et al., 2007b, 2009). The results of the present study(Fig. 8) are consistent with inhospitable environmental conditionsof either an episodic or protracted nature, and with the hypothesisthat Siberian Traps volcanic activity induced strong climatic warmingthat triggered oxygen-minimum zone expansion. However, we alsoinfer mixing of stratified ocean watermasses in the early Spathian,which represented a different type of environmental perturbationthat had effects on contemporaneous marine biotas (Brayard et al.,2007, 2009a; Orchard, 2007; Stanley, 2009; Song et al., 2011).

6. Conclusions

Reconstruction of vertical δ13CDIC gradients (Δδ13CDIC) for sevenEarly Triassic time slices yielded values of 8.5‰ (N1), 5.8‰ (P1), 3.5‰(N2), 6.5‰ (P2), 7.8‰ (N3),−1.9‰ (P3), and 2.2‰ (N4). These valuesare much larger than for Δδ13CDIC gradients in the modern ocean (~1–3‰) because of intensified water-column stratification and reducedvertical mixing in Early Triassic seas. Peaks in Δδ13CDIC around the PTB(N1) and in the early tomid-Smithian (P2–N3) coincidedwith episodesof strong climatic warming, reduced marine productivity, and expand-ed ocean anoxia. The Dienerian–Smithian boundary marks the onset

of a major mid-Early Triassic disturbance, commencing ~1 Myr afterthe latest Permian mass extinction, that we link to a second eruptivestage of the Siberian Traps. Inhospitable oceanic conditions generallypersisted until the early Spathian, when strong climatic cooling causedre-invigoration of global-ocean circulation, leading to an interval ofnegative Δδ13CDIC values and a sharp increase in δ13Ccarb driven byupwelling of nutrient-rich deepwaters. These developments markedthe end of the main eruptive stage of the Siberian Traps.

Acknowledgments

This study was supported by the 973 Program (grant no.2011CB808800), the Natural Science Foundation of China (grantsnos. 40830212 and 41172312), the “111” project (grant no. B08030),and the Open Research Program of BGEG (grant no. BGEG1016). Weare grateful to Jinxun Zuo, Junhua Huang, Chaoyong Hu, Yong Du, HuiShi, and Yanlin Xiong for laboratory and field assistance. Research byTJA was supported by the Sedimentary Geology and Paleobiology pro-gram of the U.S. National Science Foundation. This paper is a contribu-tion to IGCP Project 572 “Permian–Triassic Ecosystems”.

Appendix A. Supplementary data

Supplementary data to this article can be found online at http://dx.doi.org/10.1016/j.gloplacha.2012.10.023.

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