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Carbon and sulfur isotopic fluctuations associated with the end-Guadalupian mass extinction in South China
Gondwana Research 24 (2013) 1276–1282
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Carbon and sulfur isotopic fluctuations associated with the end-Guadalupian mass extinction in South China Yan Detian a, b,⁎, Zhang Liqin c, Qiu Zhen b a b c
Key Laboratory of Tectonics and Petroleum Resources of Ministry of Education, China University of Geosciences, Wuhan 430074, China Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China Key Laboratory of Geospace Environment and Geodesy of Ministry of Education, Wuhan University, Wuhan 430079, China
a r t i c l e
i n f o
Article history: Received 22 March 2012 Received in revised form 1 November 2012 Accepted 15 February 2013 Available online 13 March 2013 Handling Editor: J.G. Meert Keywords: Stable isotopes Mass extinctions End-Guadalupian South China
a b s t r a c t Concentrations of total organic matter (TOC), carbon isotopic compositions of carbonate and organic matter (δ13Ccarb, δ13Corg), and sulfur isotopic compositions of carbonate associated sulfate (δ34Ssulfate) across the Guadalupian–Lopingian (G–L) boundary were analyzed from identical samples of Tieqiao section, Laibin, Guangxi province, South China. The δ13Ccarb values show a positive excursion from −0.45‰ to the peak of 3.80‰ in the Laibin limestone member of the Maokou Formation, followed by a drastic drop to −2.60‰ in the lowest Heshan formation, then returned to about 1.58‰. Similar to the trends of the δ13Ccarb values, Δ13Ccarb–org values also show a positive excursion followed by a sharp negative shift. The onset of a major negative carbon isotope excursion postdates the end Guadalupian extinction that indicates subsequent severe disturbance of the ocean–atmosphere carbon cycle. The first biostratigraphic δ34Ssulfate values during the G–L transition exhibit a remarkable fluctuation: a dramatic negative shift followed by a rapid positive shift, ranging from 36.88‰ to − 37.41‰. These sulfate isotopic records suggest that the ocean during the G–L transition was strongly stratified, forming an unstable chemocline separating oxic shallow water from anoxic/euxinic deep water. Chemocline excursions, together with subsequent rapid transgression and oceanic anoxia, were likely responsible for the massive diversity decline of the G–L biotic crisis. © 2013 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.
1. Introduction The Guadalupian–Lopingian (G–L) (Middle–Late Permian) transition was a critical interval in geological history, during which dramatic climatic, oceanic, and biological changes occurred (Jin et al., 1994; He et al., 2003; Isozaki et al., 2007a). In particular, a mass extinction coinciding with the end Guadalupian global regression, was identified (Jin et al., 1994; Stanley and Yang, 1994; Wang and Sugiyama, 2000; Clapham et al., 2009). Numerous hypotheses, including global regression (Jin et al., 1994; Hallam and Wignall, 1999; Wang and Jin, 2000), oceanic anoxia (Isozaki, 1997; Weidlich, 2002), Emeishan flood basalt eruptions (Sephton et al., 2002; Zhou et al., 2002; Wignall et al., 2009a), climate cooling (Isozaki et al., 2007a; Lai et al., 2008) and catastrophic methane outburst (Retallack et al., 2006; Retallack and Jahren, 2008), have been proposed as the cause of this particular geologic event. The ultimate cause, however, is still highly controversial. Some studies have revealed a large-magnitude negative carbon excursion, associated with the mass extinction, implying an abrupt and global change in the carbon cycle (Wang et al., 2004; Isozaki et al., 2007b; Lai et al., 2008). However, the cause of the δ 13C shift and ⁎ Corresponding author at: Key Laboratory of Tectonics and Petroleum Resources of Ministry of Education, China University of Geosciences, Wuhan 430074, China. Tel.: +86 13986195196.
its relationship to the extinction has been a matter of debate. By analogy to carbon cycles, sulfur reservoir sizes and redox variations within the sedimentary sulfur cycle account for the observed variations in δ 34S values (Burdett et al., 1989; Strauss, 1997). Until now, few comparable sulfur isotopic studies across this critical boundary have been conducted. In this paper, we present new carbon and sulfur isotopic chemostratigraphic data across the G–L boundary that provides useful clues to understand the tremendous environmental, oceanographic changes and relevant biotic crisis during this critical transition.
2. Geological setting Marine sections recording continuous deposition across the G–L boundary are particularly rare because the end Guadalupian global regression tends to produce an unconformity. However, the Laibin area of Guangxi Province, South China is exceptional and records the complete G–L succession of pelagic biozones. The Global Stratotype Section and Point (GSSP) for the G–L boundary is located at the Penglaitan section, east of Laibin, which has well-constrained biostratigratigraphic resolution for this interval (Jin et al., 2006). The well-preserved Tieqiao section, ~25 km from the classic Penglaitan section, was selected for the chemostratigraphic study because it shows the same depositional sequence (Fig. 1).
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Y. Detian et al. / Gondwana Research 24 (2013) 1276–1282
The Laibin area is located in the Jiangnan Basin and lies between the Yangtze and Cathaysian cratons throughout the Late Palaeozoic– early Triassic interval (Wang and Jin, 2000). The rapid change between the end Guadalupian regression and the quick transgression during the early Lopingian drastically changed the geography of the entire South China block (Wang and Jin, 2000). The Guadalupian– Lopingian boundary succession, in ascending order, includes the Maokou (Laibin limestone member), and Heshan Formations. At the Tieqiao section, the Laibin limestone Member, 11 m thick, consists of massive pale-gray slope debris and mound limestone. The overlying Heshan Formation, 150 m thick, is mainly composed of thin-bedded cherty limestone and biogenic limestone. Details of the biostratigraphy and other relevant information were provided previously (Mei et al., 1998; Jin et al., 2006). Some studies suggested that mass biotic extinctions occurred in the base of the J. granti zone, considerably below the Guadalupian–Lopingian boundary (Sun and Xia, 2006). However, Chen et al. (2009) indicated that the potential extinction horizon is ~ 30 cm above the G/L boundary at the Tieqiao section.
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3. Analytical methods Systematic analyses for δ 13Ccarb, δ 13Corg and δ 34Ssulfate were conducted in the G–L boundary succession at the Tieqiao section. Diagenetic influences on the carbonate samples for isotopic analyses were first evaluated by petrographic examination; only homogeneous micritic carbonates were selected for isotopic analyses. The total organic carbon (TOC) content was measured by High TOC II analyzer and the analytical precision is better than ±0.05‰. Sample splits (~5 mg) for inorganic carbon and oxygen analysis were reacted with 100% phosphoric acid at 50 °C for 24 h; isotopic ratios were measured on a Finnigan MAT-251 mass-spectrometer. Analytical precision is better than ±0.05‰ both for δ 13Ccarb and δ 18O. Sample splits (300 mg to 1.5 g) for δ13Corg analysis were first dissolved with 5 N HCl in a centrifuge beaker to remove carbonates through multiple acidifications (at least two times) and subsequent drying in the heating oven, and repeatedly rinsed with deionized water to neutrality. The decalcified samples (30 to 110 mg) + CuO wire (1 g) were added to a quartz tube, and combusted at 500 °C for 1 h and 850 °C for
Fig. 1. Location and geological map for Tieqiao section, Laibin, Guangxi Province, South China (after Wang et al., 2004).
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another 3 h. Isotopic ratios were analyzed using cryogenically purified CO2 on a Finnigan MAT-253 mass spectrometer. Analytical precision for δ13Corg is better than ±0.06‰. All stable carbon isotope data were reported in standard δ-notation relative to Vienna Peedee Belmnite (VPDB) standard. The pulverized samples for δ 34Ssulfate analysis were treated with NaOCl for 24 h and repeatedly rinsed with deionized water. The bleached residues were treated with dilute 5 N HCl to decompose the carbonate. The acid-soluble sulfate hosted in the carbonate was dissolved into the solution and precipitated as BaSO4 by addition of BaCl2 solution. BaSO4 was mixed with V2O5 + pure quartz sands and combusted up to 1050 °C in the presence of copper turnings under vacuum to convert to SO2 gases. The purified SO2 was used to analyze the sulfur isotopic ratio in a Finnigan Delta-S mass spectrometer. Sulfur isotopic ratios are expressed as standard δ-notation relative to Canon Diablo Troilite (VCDT) standard. The analytical precision is better than ± 0.2‰.
values, Δ 13Ccarb–org values show a positive excursion starting from the lower J. xuanhanensis zone and reach the maximum (27.92‰) in the lowest C. postbitteri postbitteri zone (Fig. 2). Going upwards, the Δ 13Ccarb–org values decrease rapidly in the mid C. postbitteri postbitteri zone, and then recover to 24.90‰ in the mid C. dukouensis zone and generally persist upwards. The δ34Ssulfate values vary between 36.88‰ and −37.41‰ showing a negative excursion from 32.43‰ in the lower J. xuanhanensis zone to the trough at −37.41‰ in the lowest C. postbitteri postbitteri zone in the Upper Laibin limestone member, which bounce back rapidly to about 16‰ in the uppermost Laibin limestone member. This is followed by a continuous increase in δ34Ssulfate values to reach the maximum (36.88‰) in the upper C. postbitteri postbitteri zone. Going upwards, the δ 34Ssulfate values decrease and display relatively consistent values between 13.54‰ and 17.29‰ in the C. dukouensis zone.
4. Results
5.1. δ 13Ccarb, δ 13Corg and Δ 13Ccarb–org isotopic systematics
Systematic analytical results for δ 13Ccarb, δ 13Corg and δ 34Ssulfate are illustrated in Table 1 and Fig. 2. The δ 13Ccarb values show a positive excursion from − 0.45‰ in the lower J. xuanhanensis zone to the zenith of 3.80‰ near the lowest C. postbitteri postbitteri zone, this is followed by a drastic drop in δ 13Ccarb values of about 6.4‰ to the trough at − 2.60‰ in the mid C. postbitteri postbitteri zone, which bounce back slightly to about 1.58‰ in the C. dukouensis zone. The δ 13Corg values generally vary between − 25.12‰ and −23.14‰ (avg. − 24.11‰) across the G–L boundary at this section, from which one major negative excursion in the mid C. postbitteri postbitteri zone is identified (Fig. 2). The Δ 13Ccarb–org values calculated for the Tieqiao section vary between 22.32‰ and 27.92‰. Similar with the trends of the δ 13Ccarb
The low organic content in carbonates can reduce the possible diagenetic effects of TOC on δ 13Ccarb signatures through organic decomposition and bacterial sulfate reduction during burial, but it may enhance the diagenetic influences upon the δ 13Corg values (Knoll et al., 1986). No apparent covariance is present between the organic carbon abundance and δ 13Corg value (Fig. 3A, Table 1). This suggests that the primary signatures of δ 13Corg values are minimally altered during burial (Hayes et al., 1989). The organic matter at Tieqiao is exclusively of marine biomass, which is overwhelmingly composed of shortchain n-alkanes (with no apparent odd-over-even carbon number predominance) contributed by phytoplankton, and rare isoprenoids of bacterial/algal origin. This precludes other possible biomass sources (i.e., terrestrial plants richer in 13C) being responsible for
5. Discussion
Table 1 Carbon and sulfur isotopic compositions from Tieqiao section, Laibin, Guangxi Province, South China. Sample Log height Conodont no. (cm) zone
Lithology
CAS δ13Ccarb δ18Ocarb δ13Corg Δ13Ccarb–org δ34Ssulfate TOC Pyrite sulfur Fe (‰, VPDB) (‰, VPDB) (‰, VPDB) (‰, CDT) (wt. %) (wt. %) concentration concentration (ppm) (ppm)
P-125 P-127 P-128 P-130 P-131 P-133 P-134 P-135 P-136 P-137 P-138 P-139 P-140 P-142 P-144 P-145 P-147b P-148 P-149 P-150 P-151 P-152 P-153 P-154 P-155 P-156 P-157 P-158 P-159 P-160 P-161 P-163
Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone
−0.45 −0.11 0.34 0.75 2.34 2.19 1.69 1.28 2.06 0.99 3.20 2.65 2.35 2.78 2.72 3.25 3.80 1.36 2.36 1.16 1.92 0.45 0.73 −1.72 −1.16 −2.60 −2.46 0.04 −0.72 1.58 1.28 0.96
50 155 210 315 365 475 520 570 600 630 660 685 705 775 805 820 850 855 860 865 873 881 891 921 961 1006 1041 1101 1151 1216 1296 1411
J.x. J.x. J.x. J. granti J. granti J. granti J. granti J. granti J. granti J. granti J. granti J. granti J. granti C.p.h. C.p.h. C.p.h. C.p.p C.p.p C.p.p C.p.p C.p.p C.p.p C.p.p C.p.p C.p.p C.p.p C.p.p C.p.p C. dukouensis C. dukouensis C. dukouensis C. dukouensis
−9.294 −7.174 −6.294 −7.896 −8.476 −8.727 −7.679 −6.147 −9.365 −5.282 −6.245 −5.976 −6.237 −6.768 −6.324 −5.176 −6.073 −7.561 −7.797 −6.494 −6.238 −7.81 −6.545 −8.307 −8.678 −7.121 −7.089 −8.075 −8.308 −9.699 −8.547 −8.734
−25.12 −24.91 / −25.07 / / −24.52 / / / −24.00 −23.97 −23.58 −23.59 −23.67 −24.11 −24.11 −23.86 −23.79 −23.94 −23.45 −24.14 −24.23 −24.07 −24.50 −24.96 −24.78 −23.14 −23.87 −23.32 / −24.05
24.67 24.80 / 25.82 / / 26.21 / / / 27.20 26.62 25.94 26.38 26.38 27.36 27.92 25.22 26.15 25.10 25.37 24.59 24.96 22.35 23.34 22.36 22.32 23.18 23.15 24.90 / 25.02
32.43 28.01 27.77 25.00 21.72 26.39 19.56 / / / −23.16 −29.76 −21.68 −2.95 −31.50 −37.41 −24.14 16.13 16.25 16.07 13.87 18.47 11.42 / / 35.37 36.88 16.92 13.54 17.29 15.21
0.14 / / 0.02 / / 0.02 / / / 0.22 0.15 0.11 0.38 0.38 0.30 0.06 0.02 0.04 0.03 0.04 0.07 0.04 0.03 0.03 0.02 0.03 0.04 0.03 0.03 / 0.02
0.007 / / 0.005 / // 0.009 / / / 0.082 0.117 0.009 0.099 0.115 0.289 0.007 0.006 0.009 0.006 0.039 0.006 0.007 0.006 0.006 0.009 0.009 0.008 0.006 0.009 / 0.006
1890 / 3112 / 2730 / 2934 2590 / 2021 / 1635 / 2106 / 2936 3430 / / 1400 / / 1890 / / 3220 / 630 2248 1190 / 910
1133 / 749 / 1395 / 1275 / / / 4122 1172 / 2339 701 1107 / 805 / 2746 / / 1854 / / / 79 132 / 3220 / 1215
Y. Detian et al. / Gondwana Research 24 (2013) 1276–1282 Fig. 2. Carbon–sulfur isotopic chemostratigraphic profiles across the G–L boundary at Tieqiao section. Arrow indicates the possible G–L mass extinction event (Chen et al., 2009). Conodont biostratigraphy and bed numbers are based on Jin et al. (2006). J.x.—Jinogondolella xuanhanensis, J. granti—Jinogondolella granti, C.p.h.—Codonofusiella postbitteri hongshuiensis, C.p.p.—Codonofusiella postbitteri postbitteri, and C.d.—Codonofusiella dukouensis.
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Fig. 3. Cross-plots between different values at Tieqiao. (A) δ13Corg vs. TOC; (B) δ13Ccarb vs. δ13Corg; (C) δ13Ccarb vs. δ18Ocarb; (D) δ34Ssulfate vs. Pyrite sulfur concentration; (E) δ34Ssulfate vs. Fe concentration; (F) δ34Ssulfate vs. CAS concentration; (G) δ18Ocarb vs. CAS concentration; and (G)δ34Ssulfate vs. δ18Ocarb.
the positive δ 13C excursion. In addition, the studied section is thinner than 15 m thick, so they should have experienced generally the same burial temperature during burial, exerting the same diagenetic influences on the overall samples of this study. The relatively concomitant variations between δ 13Ccarb and δ 13Corg values are also considered as primary carbon isotopic signatures (Figs. 2; 3B), because diagenetic
effects on the isotopic composition of carbonate and organic carbon are generally different (Knoll et al., 1986). The data of δ 13Ccarb and δ 18Ocarb, showing no statistical relationship by means of simple linear regression (R 2 = 0.1) (Fig. 3C), further indicate that the primary carbon isotopic signatures are basically well-preserved (Banner and Hanson, 1990; Marshall, 1992; Banner, 1995).
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Our data derived from the Tieqiao section as described above clearly show a positive δ 13Ccarb excursion, which is followed by a rapid drop and subsequent increase across the G–L boundary. A parallel variation in Δ 13Ccarb–org values is well demonstrated in this section as well (Fig. 2). The positive δ 13Ccarb excursion and subsequent variations across the G–L boundary reported here are unlikely from local signatures, because this excursion in δ 13Ccarb values associated with the G–L boundary has been reported from some other sections around the Tethyan domain and the Far East (Wang et al., 2004; Kaiho et al., 2005; Isozaki et al., 2007b; Lai et al., 2008). The systematic fluctuations of the δ 13Ccarb, δ 13Corg and Δ 13Ccarb–org across the G–L boundary further indicate global perturbations of carbon cycle. Hayes et al. (1989) suggested that if Δ 13Ccarb–org remains constant while δ 13Ccarb varies, the δ 13Ccarb shifts can be attributed to global changes, such as changes in atmosphere CO2 and/or weathering rate. In contrast, variations in Δ 13Ccarb–org can be interpreted as changes in the origin of organic matter, fractionation processes or diagenesis. At Tieqiao, the gradual increase in δ 13Ccarb and δ 13Corg from the J. xuanhanensis zone through the J. granti and C. postbitteri hongshuiensis zones to the C. postbitteri postbitteri zone suggests the fraction of depleted-carbon in oceanic dissolved carbon decreased. The rise of Δ 13Ccarb–org indicates that the rate of increase of δ13Corg is lower than that of δ 13Ccarb, which reflects a change in the relative carbon fluxes between ocean water, organic matter and sediments. The progressive increase in burial of organic matter could have led to substantial removal of 12C from the carbon reservoir, resulting in the low atmospheric CO2 concentration (or pCO2) level and a positive excursion of δ 13C during this period (Kump and Arthur, 1999). The low pCO2 level, on the other hand, could have decreased photosynthetic carbon fixation of marine phytoplankton, through which less 12C would be incorporated into the organic carbon, leading to the heavy δ13Corg values (Freeman and Hayes, 1992; Rau, 1997). The relatively high abundance of organic carbon within this excursion horizon indicates the possibility of high productivity during its deposition (Figs. 2). From the lower C. postbitteri postbitteri zone, δ 13Ccarb and δ 13Corg both decrease, but with a difference in the timing and magnitude of the δ 13C variations that creates a remarkable decline in Δ 13Ccarb–org. This negative shift represents the most prominent isotopic signal in the Tieqiao section, as Δ 13Ccarb–org values never come back to levels over 26‰ above this segment. This negative δ 13C excursion reflects major changes in the carbon cycle around the time of the biotic crisis (Fig. 2), and its cause is still debated (e.g., Erwin, 1993; Wang et al., 2004; Retallack and Jahren, 2008). Considering that this negative shift postdates the extinction boundary and about 50% of all marine species become extinct (Jin et al., 2006; Sun and Xia, 2006), our data support the observation of Wang et al. (2004), that the great magnitude of Δ 13Ccarb–org shift might result from a lower abundance of marine biomass after the end-Guadalupian mass extinction. 5.2. Sulfur isotope variations and stratified ocean The isotopic composition of sedimentary sulfides and sulfates is sensitive indicators for changes of the geological, geochemical or biological environments (Shen et al., 2001, 2011). Presently, few comparable sulfur isotope studies across G–L boundaries exist. To the best of our knowledge, this is the first, systematic sulfur isotope documentation from a continuous section tied to well-established biostratigraphy, and it allows us to add a new geochemical constraint to the current arguments on the cause and consequence of the end-Guadalupian mass extinction. At Tieqiao, the absence of apparent covariance between high pyrite concentration, high apparent Fe concentration, high carbonateassociated sulfate (CAS) concentration and low values of δ34Ssulfate indicates that pyrite oxidation during the CAS extraction process does not contribute to CAS isotopic composition (Fig. 3D; E; F; Table 1). The
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covariance between δ18Ocarb and CAS concentration is also not clear; however, the slight correlation between the δ18Ocarb and δ34Ssulfate suggests a likely diagenetic modification of δ34Ssulfate values in these samples (Fig. 3G; H). This proposed diagenetic alteration may provide a clue to the environmental changes occurring (oxic-anoxic cycles) at this time (Riccardi et al., 2006). The δ 34Ssulfate values at Tieqiao exhibit a remarkable fluctuation: a dramatic negative shift followed by a rapid positive shift during the G–L transition. Burdett et al. (1989) and Strauss (1997) suggested that the δ 34Ssulfate values are mainly controlled by the balance of bacterial sulfate reduction and sulfide oxidation. The sulfur isotopic composition of Late Permian seawater sulfate, with respect to CDT, is assumed to have been around +13‰ (Kampschulte and Strauss, 2004). In contrast, the J. xuanhanensis zone and mid-early J. granti zone deposits have relatively high δ34Ssulfate values (19.56‰–32.43‰, avg. 25.87‰). The sulfate values of mid-early J. granti zone are statistically ~12‰ heavier than the Late Permian seawater sulfate values. This isotopic difference indicates that the Guadalupian ocean was strongly stratified and the deep water may have been predominantly anoxic when the mid-early J. granti zone sediments were deposited. Under this scenario, deepwater anoxia would have sequestered isotopically light sulfur of 32S into deep anoxic parts (Zhang et al., 2010). As a result, seawater sulfate in the shallow waters (e.g., the Laibin area) would probably have been enriched in 34S in particular when oceanic sulfate input was limited, producing sulfate with heavy δ34S values as seen in the mid-early J. granti zone sediments (Zhang et al., 2010). The dramatic negative shift of δ34Ssulfate values in the Upper Laibin member (from the upper J. granti zone through the C. postbitteri hongshuiensis zone to the lowest C. postbitteri postbitteri Zone) further strongly supported the deepwater anoxia condition. During the interval of negative excursion, 32S would have been transported from anoxic deep water into oxic shallow water, producing a chemical event that perturbed the sulfur cycle, as observed at Tieqiao (Fig. 2). This process is evidenced by the relatively high pyrite contents in the Upper Laibin limestone member (Table 1). This large negative δ 34S shift of ~50‰ is difficult to explain without invoking predominant deep-water anoxia (Zhang et al., 2010). From the uppermost Laibin member, the short-lived, large positive δ34Ssulfate excursion suggests a strongly stratified ocean formed again, i.e. a chemocline separating oxic shallow water from anoxic/euxinic deep water. This anoxic environment has been demonstrated by pyrite petrography and benthic diversity (Wignall et al., 2009b). The extremely positive δ 34Ssulfate values might reflect a regional environmental setting superimposed on the global seawater signature. During the G–L transition, the Laibin area lay within the Jiangnan Basin, which likely had limited access to the open ocean and led to extreme 34S enrichment via bacterial sulfate reduction (Strauss, 1997). 6. Implications for the Late Guadalupian extinction Isotopic data at Tieqiao indicate large anomalies in the carbon and sulfur cycle accompanying the mass extinction. The onset of a major negative carbon isotope excursion was nearly coincident with the mass extinction level, which was around the G/L boundary, implying severe destabilization of the ocean–atmosphere carbon cycle. The sulfur isotopic evidence suggests that the Guadalupian ocean was strongly stratified. The oxic bottom waters predominated over the Tieqiao area during the Laibin limestone intervals, but the redox environment may have changed due to chemocline instability. During the uppermost Laibin and lower Heshan intervals, an anoxic environment predominated during rapid transgression. An upward chemocline excursion led to significant amounts of H2S gases being released into the photic zone and even atmosphere, resulting in photic zone euxinia (Kump et al., 2005); this scenario would have been very harmful, both for the benthic and planktonic fauna. The subsequent rapid transgression resulted in oceanic anoxia, expelling the benthos from their previous ecologic niches, thereby leading to their significant
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decline. Therefore, chemocline excursions, together with subsequent rapid transgression and oceanic anoxia, may account for the massive diversity decline during the G–L transition. Acknowledgments This work was supported by the NSFC projects (41103018 and 41273001). We thank Fusong Zhang and Lianjun Feng (IGGCAS) for their assistance in lab work. References Banner, J.L., 1995. Application of the trace element and isotope geochemistry of strontium to studies of carbonate diagenesis. Sedimentology 42, 805–824. Banner, J.L., Hanson, G.N., 1990. Calculation of simultaneous isotopic and trace element variations during water–rock interaction with applications to carbonate diagenesis. Geochimica et Cosmochimica Acta 54, 3123–3137. Burdett, J.W., Arthur, M.A., Richardson, M., 1989. A Neogene seawater sulfur isotope age curve from calcareous pelagic microfossils. Earth and Planetary Science Letters 94, 189–198. Chen, Z.Q., George, A.D., Yang, W.R., 2009. Effects of Middle–Late Permian sea-level changes and mass extinction on the formation of the Tieqiao skeletal mound in the Laibin area, South China. Australian Journal of Earth Sciences 56, 745–763. Clapham, M.E., Shen, S.Z., Bottjer, D.J., 2009. The double mass extinction revisited: reassessing the severity, selectivity, and causes of the end-Guadalupian biotic crisis (Late Permian). Paleobiology 35, 32–50. Erwin, D.H., 1993. The Great Phanerozoic Crisis: Life and Death in the Permian. Columbia University Press, New York. Freeman, K.H., Hayes, J.M., 1992. Fractionation of carbon isotopes by phytoplankton and estimates of ancient CO2 levels. Global Biogeochemical Cycles 6, 185–198. Hallam, A., Wignall, P.B., 1999. Mass extinctions and sea-level changes. Earth-Science Reviews 48, 217–250. Hayes, J.M., Popp, B.N., Takigiku, R., Johnson, M., 1989. An isotopic study of biogeochemical relationships between carbonates and organic carbon in the Greenhorn Formation. Geochimica et Cosmochimica Acta 53, 2961–2972. He, B., Xu, Y.G., Chung, S.L., Xiao, L., Wang, Y., 2003. Sedimentary evidence for a rapid, kilometer-scale crustal doming prior to the eruption of the Emeishan flood basalts. Earth and Planetary Science Letters 213, 391–405. Isozaki, Y., 1997. Permo-triassic boundary superanoxia and stratified superocean: records from lost deep sea. Science 276, 235–238. Isozaki, Y., Kawahata, H., Minoshima, K., 2007a. The Capitanian (Permian) Kamura cooling event: the beginning of the Paleozoic–Mesozoic transition. Palaeoworld 16, 16–30. Isozaki, Y., Kawahata, H., Ota, A., 2007b. A unique carbon isotope record across the Guadalupian–Lopingian (Middle–Upper Permian) boundary in mid-oceanic paleoatoll carbonates: the high-productivity “Kamura event” and its collapse in Panthalassa. Global and Planetary Change 55, 21–38. Jin, Y.G., Zhang, J., Shang, Q.H., 1994. Two phases of the end-Permian mass extinction. In: Embry, A.F., Beauchamp, B., Glass, D.J. (Eds.), Pangea: global environments and resources: Canadian Society of Petroleum Geologists, Memoir, 17, pp. 818–822. Jin, Y.G., Shen, S.Z., Henderson, C.M., 2006. The Global Stratotype Section and Point (GSSP) for the boundary between the Capitanian and Wuchiapingian stage (Permian). Episodes 29, 253–262. Kaiho, K., Chen, Z.Q., Ohashi, T., Arinobu, T., Sawada, K., Cramer, B.S., 2005. A negative carbon isotope anomaly associated with the earliest Lopingian (Late Permian) mass extinction. Palaeogeography, Palaeoclimatology, Palaeoecology 223, 172–180. Kampschulte, A., Strauss, H., 2004. The sulfur isotopic evolution of Phanerozoic seawater based on the analysis of structurally substituted sulfate in carbonates. Chemical Geology 204, 255–286.
Knoll, A.H., Hayes, J.M., Kaufman, A.J., Sweet, K., Lambert, I.B., 1986. Secular variation in carbon isotope ratios from upper Proterozoic successions of Svalbard and East Greenland. Nature 321, 832–838. Kump, L.R., Arthur, M.A., 1999. Interpreting carbon-isotope excursions: carbonates and organic matter. Chemical Geology 161, 181–198. Kump, L.R., Pavlov, A., Arthur, M.A., 2005. Massive release of hydrogen sulfide to the surface ocean and atmosphere during intervals of oceanic anoxia. Geology 33, 397–400. Lai, X.L., Wang, W., Wignall, P.B., Bond, D.G., Jiang, H.S., Ali, J.R., John, E.H., Sun, Y.D., 2008. Paleoenvironmental change during the end-Guadalupian (Permian) mass extinction in Sichuan, China. Palaeogeography, Palaeoclimatology, Palaeoecology 269, 78–93. Marshall, J.D., 1992. Climatic and oceanographic isotopic signals from the carbonate rocks record and their preservation. Geological Magazine 129, 143–160. Mei, S.L., Jin, Y.G., Wardlaw, B.R., 1998. Conodont succession of the Guadalupian– Wuchiapingian boundary strata, Laibin, Guangxi, South China and Texas, USA. Palaeoworld 9, 53–76. Rau, G.H., 1997. CO2aq-dependent photosynthetic 13C fractionation in the ocean: a model versus measurements. Global Biogeochemical Cycles 11, 267–278. Retallack, G.J., Jahren, A.H., 2008. Methane release from igneous intrusion of coal during Late Permian extinction events. Journal of Geology 116, 1–20. Retallack, G.J., Metzger, C.A., Greaver, T., Jahren, A.H., Smith, R.M.H., Sheldon, N.D., 2006. Middle–Late Permian mass extinction on land. Geological Society of America Bulletin 118, 1398–1411. Riccardi, A.L., Arthur, M.A., Kump, L.R., 2006. Sulfur isotopic evidence for chemocline upward excursions during the end-Permian mass extinction. Geochimica et Cosmochimica Acta 70, 5740–5752. Sephton, M.A., Looy, C.V., Veefkind, R.J., Visscher, H., Brinkhuis, H., De Leeuw, J.W., 2002. Synchronous record of δ13C shifts in the oceans and atmosphere at the end of the Permian. Geological Society of America Special Paper 356, 455–462. Shen, Y.A., Buick, R., Canfield, D.E., 2001. Isotopic evidence for microbial sulphate reduction in the early Archaean era. Nature 410, 77–81. Shen, Y.A., Farquhar, J., Zhang, H., Masterson, A., Zhang, T.G., Wing, B.A., 2011. Multiple S-isotopic evidence for episodic shoaling of anoxic water during Late Permian mass extinction. Nature Communications 2, 1–5. Stanley, S.M., Yang, X., 1994. A double mass extinction at the end of the Paleozoic era. Science 266, 1340–1344. Strauss, H., 1997. The isotopic composition of sedimentary sulfur through time. Palaeogeography, Palaeoclimatology, Palaeoecology 132, 97–118. Sun, D.Y., Xia, W.C., 2006. Identification of the Guadalupian–Lopingian boundary in the Permian in a bedded chert sequence, South China. Palaeogeography, Palaeoclimatology, Palaeoecology 236, 272–289. Wang, Y., Jin, Y.G., 2000. Permian palaeogeographic evolution of the Jiangnan Basin, South China. Palaeogeography, Palaeoclimatology, Palaeoecology 160, 35–44. Wang, X.D., Sugiyama, T., 2000. Diversity and extinction patterns of Permian coral faunas of China. Lethaia 33, 285–294. Wang, W., Cao, C.Q., Wang, Y., 2004. The carbon isotope excursion on GSSP candidate section of Lopingian–Guadalupian boundary. Earth and Planetary Science Letters 220, 57–67. Weidlich, O., 2002. Permian reefs re-examined: extrinsic control mechanisms of gradual and abrupt changes during 40 my of reef evolution. Geobios 35, 287–294. Wignall, P.B., Sun, Y.D., Bond, D.P.G., Izon, G., Newton, R.J., Vedrine, S., Widdowson, M., Ali, J.R., Lai, X.L., Jiang, H.S., Cope, H., Bottrell, S.H., 2009a. Volcanism, mass extinction, and carbon isotope fluctuations in the Middle Permian of China. Science 324, 1179–1182. Wignall, P.B., Vedrine, S., Bond, D.P.G., Wang, W., Lai, X.L., Ali, J.R., Jiang, H.S., 2009b. Facies analysis and sea-level change at the Guadalupian–Lopingian Global Stratotype (Laibin, South China), and its bearing on the end-Guadalupian mass extinction. Journal of the Geological Society 166, 655–666. Zhang, T.G., Shen, Y.A., Zhan, R.B., Shen, S.Z., Chen, X., 2010. Large perturbations of the carbon and sulfur cycle associated with the Late Ordovician mass extinction in South China. Geology 37, 299–302. Zhou, M.F., Malpas, J., Song, X.Y., Robinson, P.T., Sun, M., Kennedy, A.K., Lesher, C.M., Keays, R.R., 2002. A temporal link between the Emeishan large igneous province (SW China) and the end-Guadalupian mass extinction. Earth and Planetary Science Letters 196, 113–122.
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Carbon and sulfur isotopic fluctuations associated with the end-Guadalupian mass extinction in South China Yan Detian a, b,⁎, Zhang Liqin c, Qiu Zhen b a b c
Key Laboratory of Tectonics and Petroleum Resources of Ministry of Education, China University of Geosciences, Wuhan 430074, China Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China Key Laboratory of Geospace Environment and Geodesy of Ministry of Education, Wuhan University, Wuhan 430079, China
a r t i c l e
i n f o
Article history: Received 22 March 2012 Received in revised form 1 November 2012 Accepted 15 February 2013 Available online 13 March 2013 Handling Editor: J.G. Meert Keywords: Stable isotopes Mass extinctions End-Guadalupian South China
a b s t r a c t Concentrations of total organic matter (TOC), carbon isotopic compositions of carbonate and organic matter (δ13Ccarb, δ13Corg), and sulfur isotopic compositions of carbonate associated sulfate (δ34Ssulfate) across the Guadalupian–Lopingian (G–L) boundary were analyzed from identical samples of Tieqiao section, Laibin, Guangxi province, South China. The δ13Ccarb values show a positive excursion from −0.45‰ to the peak of 3.80‰ in the Laibin limestone member of the Maokou Formation, followed by a drastic drop to −2.60‰ in the lowest Heshan formation, then returned to about 1.58‰. Similar to the trends of the δ13Ccarb values, Δ13Ccarb–org values also show a positive excursion followed by a sharp negative shift. The onset of a major negative carbon isotope excursion postdates the end Guadalupian extinction that indicates subsequent severe disturbance of the ocean–atmosphere carbon cycle. The first biostratigraphic δ34Ssulfate values during the G–L transition exhibit a remarkable fluctuation: a dramatic negative shift followed by a rapid positive shift, ranging from 36.88‰ to − 37.41‰. These sulfate isotopic records suggest that the ocean during the G–L transition was strongly stratified, forming an unstable chemocline separating oxic shallow water from anoxic/euxinic deep water. Chemocline excursions, together with subsequent rapid transgression and oceanic anoxia, were likely responsible for the massive diversity decline of the G–L biotic crisis. © 2013 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.
1. Introduction The Guadalupian–Lopingian (G–L) (Middle–Late Permian) transition was a critical interval in geological history, during which dramatic climatic, oceanic, and biological changes occurred (Jin et al., 1994; He et al., 2003; Isozaki et al., 2007a). In particular, a mass extinction coinciding with the end Guadalupian global regression, was identified (Jin et al., 1994; Stanley and Yang, 1994; Wang and Sugiyama, 2000; Clapham et al., 2009). Numerous hypotheses, including global regression (Jin et al., 1994; Hallam and Wignall, 1999; Wang and Jin, 2000), oceanic anoxia (Isozaki, 1997; Weidlich, 2002), Emeishan flood basalt eruptions (Sephton et al., 2002; Zhou et al., 2002; Wignall et al., 2009a), climate cooling (Isozaki et al., 2007a; Lai et al., 2008) and catastrophic methane outburst (Retallack et al., 2006; Retallack and Jahren, 2008), have been proposed as the cause of this particular geologic event. The ultimate cause, however, is still highly controversial. Some studies have revealed a large-magnitude negative carbon excursion, associated with the mass extinction, implying an abrupt and global change in the carbon cycle (Wang et al., 2004; Isozaki et al., 2007b; Lai et al., 2008). However, the cause of the δ 13C shift and ⁎ Corresponding author at: Key Laboratory of Tectonics and Petroleum Resources of Ministry of Education, China University of Geosciences, Wuhan 430074, China. Tel.: +86 13986195196.
its relationship to the extinction has been a matter of debate. By analogy to carbon cycles, sulfur reservoir sizes and redox variations within the sedimentary sulfur cycle account for the observed variations in δ 34S values (Burdett et al., 1989; Strauss, 1997). Until now, few comparable sulfur isotopic studies across this critical boundary have been conducted. In this paper, we present new carbon and sulfur isotopic chemostratigraphic data across the G–L boundary that provides useful clues to understand the tremendous environmental, oceanographic changes and relevant biotic crisis during this critical transition.
2. Geological setting Marine sections recording continuous deposition across the G–L boundary are particularly rare because the end Guadalupian global regression tends to produce an unconformity. However, the Laibin area of Guangxi Province, South China is exceptional and records the complete G–L succession of pelagic biozones. The Global Stratotype Section and Point (GSSP) for the G–L boundary is located at the Penglaitan section, east of Laibin, which has well-constrained biostratigratigraphic resolution for this interval (Jin et al., 2006). The well-preserved Tieqiao section, ~25 km from the classic Penglaitan section, was selected for the chemostratigraphic study because it shows the same depositional sequence (Fig. 1).
1342-937X/$ – see front matter © 2013 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.gr.2013.02.008
Y. Detian et al. / Gondwana Research 24 (2013) 1276–1282
The Laibin area is located in the Jiangnan Basin and lies between the Yangtze and Cathaysian cratons throughout the Late Palaeozoic– early Triassic interval (Wang and Jin, 2000). The rapid change between the end Guadalupian regression and the quick transgression during the early Lopingian drastically changed the geography of the entire South China block (Wang and Jin, 2000). The Guadalupian– Lopingian boundary succession, in ascending order, includes the Maokou (Laibin limestone member), and Heshan Formations. At the Tieqiao section, the Laibin limestone Member, 11 m thick, consists of massive pale-gray slope debris and mound limestone. The overlying Heshan Formation, 150 m thick, is mainly composed of thin-bedded cherty limestone and biogenic limestone. Details of the biostratigraphy and other relevant information were provided previously (Mei et al., 1998; Jin et al., 2006). Some studies suggested that mass biotic extinctions occurred in the base of the J. granti zone, considerably below the Guadalupian–Lopingian boundary (Sun and Xia, 2006). However, Chen et al. (2009) indicated that the potential extinction horizon is ~ 30 cm above the G/L boundary at the Tieqiao section.
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3. Analytical methods Systematic analyses for δ 13Ccarb, δ 13Corg and δ 34Ssulfate were conducted in the G–L boundary succession at the Tieqiao section. Diagenetic influences on the carbonate samples for isotopic analyses were first evaluated by petrographic examination; only homogeneous micritic carbonates were selected for isotopic analyses. The total organic carbon (TOC) content was measured by High TOC II analyzer and the analytical precision is better than ±0.05‰. Sample splits (~5 mg) for inorganic carbon and oxygen analysis were reacted with 100% phosphoric acid at 50 °C for 24 h; isotopic ratios were measured on a Finnigan MAT-251 mass-spectrometer. Analytical precision is better than ±0.05‰ both for δ 13Ccarb and δ 18O. Sample splits (300 mg to 1.5 g) for δ13Corg analysis were first dissolved with 5 N HCl in a centrifuge beaker to remove carbonates through multiple acidifications (at least two times) and subsequent drying in the heating oven, and repeatedly rinsed with deionized water to neutrality. The decalcified samples (30 to 110 mg) + CuO wire (1 g) were added to a quartz tube, and combusted at 500 °C for 1 h and 850 °C for
Fig. 1. Location and geological map for Tieqiao section, Laibin, Guangxi Province, South China (after Wang et al., 2004).
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another 3 h. Isotopic ratios were analyzed using cryogenically purified CO2 on a Finnigan MAT-253 mass spectrometer. Analytical precision for δ13Corg is better than ±0.06‰. All stable carbon isotope data were reported in standard δ-notation relative to Vienna Peedee Belmnite (VPDB) standard. The pulverized samples for δ 34Ssulfate analysis were treated with NaOCl for 24 h and repeatedly rinsed with deionized water. The bleached residues were treated with dilute 5 N HCl to decompose the carbonate. The acid-soluble sulfate hosted in the carbonate was dissolved into the solution and precipitated as BaSO4 by addition of BaCl2 solution. BaSO4 was mixed with V2O5 + pure quartz sands and combusted up to 1050 °C in the presence of copper turnings under vacuum to convert to SO2 gases. The purified SO2 was used to analyze the sulfur isotopic ratio in a Finnigan Delta-S mass spectrometer. Sulfur isotopic ratios are expressed as standard δ-notation relative to Canon Diablo Troilite (VCDT) standard. The analytical precision is better than ± 0.2‰.
values, Δ 13Ccarb–org values show a positive excursion starting from the lower J. xuanhanensis zone and reach the maximum (27.92‰) in the lowest C. postbitteri postbitteri zone (Fig. 2). Going upwards, the Δ 13Ccarb–org values decrease rapidly in the mid C. postbitteri postbitteri zone, and then recover to 24.90‰ in the mid C. dukouensis zone and generally persist upwards. The δ34Ssulfate values vary between 36.88‰ and −37.41‰ showing a negative excursion from 32.43‰ in the lower J. xuanhanensis zone to the trough at −37.41‰ in the lowest C. postbitteri postbitteri zone in the Upper Laibin limestone member, which bounce back rapidly to about 16‰ in the uppermost Laibin limestone member. This is followed by a continuous increase in δ34Ssulfate values to reach the maximum (36.88‰) in the upper C. postbitteri postbitteri zone. Going upwards, the δ 34Ssulfate values decrease and display relatively consistent values between 13.54‰ and 17.29‰ in the C. dukouensis zone.
4. Results
5.1. δ 13Ccarb, δ 13Corg and Δ 13Ccarb–org isotopic systematics
Systematic analytical results for δ 13Ccarb, δ 13Corg and δ 34Ssulfate are illustrated in Table 1 and Fig. 2. The δ 13Ccarb values show a positive excursion from − 0.45‰ in the lower J. xuanhanensis zone to the zenith of 3.80‰ near the lowest C. postbitteri postbitteri zone, this is followed by a drastic drop in δ 13Ccarb values of about 6.4‰ to the trough at − 2.60‰ in the mid C. postbitteri postbitteri zone, which bounce back slightly to about 1.58‰ in the C. dukouensis zone. The δ 13Corg values generally vary between − 25.12‰ and −23.14‰ (avg. − 24.11‰) across the G–L boundary at this section, from which one major negative excursion in the mid C. postbitteri postbitteri zone is identified (Fig. 2). The Δ 13Ccarb–org values calculated for the Tieqiao section vary between 22.32‰ and 27.92‰. Similar with the trends of the δ 13Ccarb
The low organic content in carbonates can reduce the possible diagenetic effects of TOC on δ 13Ccarb signatures through organic decomposition and bacterial sulfate reduction during burial, but it may enhance the diagenetic influences upon the δ 13Corg values (Knoll et al., 1986). No apparent covariance is present between the organic carbon abundance and δ 13Corg value (Fig. 3A, Table 1). This suggests that the primary signatures of δ 13Corg values are minimally altered during burial (Hayes et al., 1989). The organic matter at Tieqiao is exclusively of marine biomass, which is overwhelmingly composed of shortchain n-alkanes (with no apparent odd-over-even carbon number predominance) contributed by phytoplankton, and rare isoprenoids of bacterial/algal origin. This precludes other possible biomass sources (i.e., terrestrial plants richer in 13C) being responsible for
5. Discussion
Table 1 Carbon and sulfur isotopic compositions from Tieqiao section, Laibin, Guangxi Province, South China. Sample Log height Conodont no. (cm) zone
Lithology
CAS δ13Ccarb δ18Ocarb δ13Corg Δ13Ccarb–org δ34Ssulfate TOC Pyrite sulfur Fe (‰, VPDB) (‰, VPDB) (‰, VPDB) (‰, CDT) (wt. %) (wt. %) concentration concentration (ppm) (ppm)
P-125 P-127 P-128 P-130 P-131 P-133 P-134 P-135 P-136 P-137 P-138 P-139 P-140 P-142 P-144 P-145 P-147b P-148 P-149 P-150 P-151 P-152 P-153 P-154 P-155 P-156 P-157 P-158 P-159 P-160 P-161 P-163
Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone Cherty limestone
−0.45 −0.11 0.34 0.75 2.34 2.19 1.69 1.28 2.06 0.99 3.20 2.65 2.35 2.78 2.72 3.25 3.80 1.36 2.36 1.16 1.92 0.45 0.73 −1.72 −1.16 −2.60 −2.46 0.04 −0.72 1.58 1.28 0.96
50 155 210 315 365 475 520 570 600 630 660 685 705 775 805 820 850 855 860 865 873 881 891 921 961 1006 1041 1101 1151 1216 1296 1411
J.x. J.x. J.x. J. granti J. granti J. granti J. granti J. granti J. granti J. granti J. granti J. granti J. granti C.p.h. C.p.h. C.p.h. C.p.p C.p.p C.p.p C.p.p C.p.p C.p.p C.p.p C.p.p C.p.p C.p.p C.p.p C.p.p C. dukouensis C. dukouensis C. dukouensis C. dukouensis
−9.294 −7.174 −6.294 −7.896 −8.476 −8.727 −7.679 −6.147 −9.365 −5.282 −6.245 −5.976 −6.237 −6.768 −6.324 −5.176 −6.073 −7.561 −7.797 −6.494 −6.238 −7.81 −6.545 −8.307 −8.678 −7.121 −7.089 −8.075 −8.308 −9.699 −8.547 −8.734
−25.12 −24.91 / −25.07 / / −24.52 / / / −24.00 −23.97 −23.58 −23.59 −23.67 −24.11 −24.11 −23.86 −23.79 −23.94 −23.45 −24.14 −24.23 −24.07 −24.50 −24.96 −24.78 −23.14 −23.87 −23.32 / −24.05
24.67 24.80 / 25.82 / / 26.21 / / / 27.20 26.62 25.94 26.38 26.38 27.36 27.92 25.22 26.15 25.10 25.37 24.59 24.96 22.35 23.34 22.36 22.32 23.18 23.15 24.90 / 25.02
32.43 28.01 27.77 25.00 21.72 26.39 19.56 / / / −23.16 −29.76 −21.68 −2.95 −31.50 −37.41 −24.14 16.13 16.25 16.07 13.87 18.47 11.42 / / 35.37 36.88 16.92 13.54 17.29 15.21
0.14 / / 0.02 / / 0.02 / / / 0.22 0.15 0.11 0.38 0.38 0.30 0.06 0.02 0.04 0.03 0.04 0.07 0.04 0.03 0.03 0.02 0.03 0.04 0.03 0.03 / 0.02
0.007 / / 0.005 / // 0.009 / / / 0.082 0.117 0.009 0.099 0.115 0.289 0.007 0.006 0.009 0.006 0.039 0.006 0.007 0.006 0.006 0.009 0.009 0.008 0.006 0.009 / 0.006
1890 / 3112 / 2730 / 2934 2590 / 2021 / 1635 / 2106 / 2936 3430 / / 1400 / / 1890 / / 3220 / 630 2248 1190 / 910
1133 / 749 / 1395 / 1275 / / / 4122 1172 / 2339 701 1107 / 805 / 2746 / / 1854 / / / 79 132 / 3220 / 1215
Y. Detian et al. / Gondwana Research 24 (2013) 1276–1282 Fig. 2. Carbon–sulfur isotopic chemostratigraphic profiles across the G–L boundary at Tieqiao section. Arrow indicates the possible G–L mass extinction event (Chen et al., 2009). Conodont biostratigraphy and bed numbers are based on Jin et al. (2006). J.x.—Jinogondolella xuanhanensis, J. granti—Jinogondolella granti, C.p.h.—Codonofusiella postbitteri hongshuiensis, C.p.p.—Codonofusiella postbitteri postbitteri, and C.d.—Codonofusiella dukouensis.
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Fig. 3. Cross-plots between different values at Tieqiao. (A) δ13Corg vs. TOC; (B) δ13Ccarb vs. δ13Corg; (C) δ13Ccarb vs. δ18Ocarb; (D) δ34Ssulfate vs. Pyrite sulfur concentration; (E) δ34Ssulfate vs. Fe concentration; (F) δ34Ssulfate vs. CAS concentration; (G) δ18Ocarb vs. CAS concentration; and (G)δ34Ssulfate vs. δ18Ocarb.
the positive δ 13C excursion. In addition, the studied section is thinner than 15 m thick, so they should have experienced generally the same burial temperature during burial, exerting the same diagenetic influences on the overall samples of this study. The relatively concomitant variations between δ 13Ccarb and δ 13Corg values are also considered as primary carbon isotopic signatures (Figs. 2; 3B), because diagenetic
effects on the isotopic composition of carbonate and organic carbon are generally different (Knoll et al., 1986). The data of δ 13Ccarb and δ 18Ocarb, showing no statistical relationship by means of simple linear regression (R 2 = 0.1) (Fig. 3C), further indicate that the primary carbon isotopic signatures are basically well-preserved (Banner and Hanson, 1990; Marshall, 1992; Banner, 1995).
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Our data derived from the Tieqiao section as described above clearly show a positive δ 13Ccarb excursion, which is followed by a rapid drop and subsequent increase across the G–L boundary. A parallel variation in Δ 13Ccarb–org values is well demonstrated in this section as well (Fig. 2). The positive δ 13Ccarb excursion and subsequent variations across the G–L boundary reported here are unlikely from local signatures, because this excursion in δ 13Ccarb values associated with the G–L boundary has been reported from some other sections around the Tethyan domain and the Far East (Wang et al., 2004; Kaiho et al., 2005; Isozaki et al., 2007b; Lai et al., 2008). The systematic fluctuations of the δ 13Ccarb, δ 13Corg and Δ 13Ccarb–org across the G–L boundary further indicate global perturbations of carbon cycle. Hayes et al. (1989) suggested that if Δ 13Ccarb–org remains constant while δ 13Ccarb varies, the δ 13Ccarb shifts can be attributed to global changes, such as changes in atmosphere CO2 and/or weathering rate. In contrast, variations in Δ 13Ccarb–org can be interpreted as changes in the origin of organic matter, fractionation processes or diagenesis. At Tieqiao, the gradual increase in δ 13Ccarb and δ 13Corg from the J. xuanhanensis zone through the J. granti and C. postbitteri hongshuiensis zones to the C. postbitteri postbitteri zone suggests the fraction of depleted-carbon in oceanic dissolved carbon decreased. The rise of Δ 13Ccarb–org indicates that the rate of increase of δ13Corg is lower than that of δ 13Ccarb, which reflects a change in the relative carbon fluxes between ocean water, organic matter and sediments. The progressive increase in burial of organic matter could have led to substantial removal of 12C from the carbon reservoir, resulting in the low atmospheric CO2 concentration (or pCO2) level and a positive excursion of δ 13C during this period (Kump and Arthur, 1999). The low pCO2 level, on the other hand, could have decreased photosynthetic carbon fixation of marine phytoplankton, through which less 12C would be incorporated into the organic carbon, leading to the heavy δ13Corg values (Freeman and Hayes, 1992; Rau, 1997). The relatively high abundance of organic carbon within this excursion horizon indicates the possibility of high productivity during its deposition (Figs. 2). From the lower C. postbitteri postbitteri zone, δ 13Ccarb and δ 13Corg both decrease, but with a difference in the timing and magnitude of the δ 13C variations that creates a remarkable decline in Δ 13Ccarb–org. This negative shift represents the most prominent isotopic signal in the Tieqiao section, as Δ 13Ccarb–org values never come back to levels over 26‰ above this segment. This negative δ 13C excursion reflects major changes in the carbon cycle around the time of the biotic crisis (Fig. 2), and its cause is still debated (e.g., Erwin, 1993; Wang et al., 2004; Retallack and Jahren, 2008). Considering that this negative shift postdates the extinction boundary and about 50% of all marine species become extinct (Jin et al., 2006; Sun and Xia, 2006), our data support the observation of Wang et al. (2004), that the great magnitude of Δ 13Ccarb–org shift might result from a lower abundance of marine biomass after the end-Guadalupian mass extinction. 5.2. Sulfur isotope variations and stratified ocean The isotopic composition of sedimentary sulfides and sulfates is sensitive indicators for changes of the geological, geochemical or biological environments (Shen et al., 2001, 2011). Presently, few comparable sulfur isotope studies across G–L boundaries exist. To the best of our knowledge, this is the first, systematic sulfur isotope documentation from a continuous section tied to well-established biostratigraphy, and it allows us to add a new geochemical constraint to the current arguments on the cause and consequence of the end-Guadalupian mass extinction. At Tieqiao, the absence of apparent covariance between high pyrite concentration, high apparent Fe concentration, high carbonateassociated sulfate (CAS) concentration and low values of δ34Ssulfate indicates that pyrite oxidation during the CAS extraction process does not contribute to CAS isotopic composition (Fig. 3D; E; F; Table 1). The
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covariance between δ18Ocarb and CAS concentration is also not clear; however, the slight correlation between the δ18Ocarb and δ34Ssulfate suggests a likely diagenetic modification of δ34Ssulfate values in these samples (Fig. 3G; H). This proposed diagenetic alteration may provide a clue to the environmental changes occurring (oxic-anoxic cycles) at this time (Riccardi et al., 2006). The δ 34Ssulfate values at Tieqiao exhibit a remarkable fluctuation: a dramatic negative shift followed by a rapid positive shift during the G–L transition. Burdett et al. (1989) and Strauss (1997) suggested that the δ 34Ssulfate values are mainly controlled by the balance of bacterial sulfate reduction and sulfide oxidation. The sulfur isotopic composition of Late Permian seawater sulfate, with respect to CDT, is assumed to have been around +13‰ (Kampschulte and Strauss, 2004). In contrast, the J. xuanhanensis zone and mid-early J. granti zone deposits have relatively high δ34Ssulfate values (19.56‰–32.43‰, avg. 25.87‰). The sulfate values of mid-early J. granti zone are statistically ~12‰ heavier than the Late Permian seawater sulfate values. This isotopic difference indicates that the Guadalupian ocean was strongly stratified and the deep water may have been predominantly anoxic when the mid-early J. granti zone sediments were deposited. Under this scenario, deepwater anoxia would have sequestered isotopically light sulfur of 32S into deep anoxic parts (Zhang et al., 2010). As a result, seawater sulfate in the shallow waters (e.g., the Laibin area) would probably have been enriched in 34S in particular when oceanic sulfate input was limited, producing sulfate with heavy δ34S values as seen in the mid-early J. granti zone sediments (Zhang et al., 2010). The dramatic negative shift of δ34Ssulfate values in the Upper Laibin member (from the upper J. granti zone through the C. postbitteri hongshuiensis zone to the lowest C. postbitteri postbitteri Zone) further strongly supported the deepwater anoxia condition. During the interval of negative excursion, 32S would have been transported from anoxic deep water into oxic shallow water, producing a chemical event that perturbed the sulfur cycle, as observed at Tieqiao (Fig. 2). This process is evidenced by the relatively high pyrite contents in the Upper Laibin limestone member (Table 1). This large negative δ 34S shift of ~50‰ is difficult to explain without invoking predominant deep-water anoxia (Zhang et al., 2010). From the uppermost Laibin member, the short-lived, large positive δ34Ssulfate excursion suggests a strongly stratified ocean formed again, i.e. a chemocline separating oxic shallow water from anoxic/euxinic deep water. This anoxic environment has been demonstrated by pyrite petrography and benthic diversity (Wignall et al., 2009b). The extremely positive δ 34Ssulfate values might reflect a regional environmental setting superimposed on the global seawater signature. During the G–L transition, the Laibin area lay within the Jiangnan Basin, which likely had limited access to the open ocean and led to extreme 34S enrichment via bacterial sulfate reduction (Strauss, 1997). 6. Implications for the Late Guadalupian extinction Isotopic data at Tieqiao indicate large anomalies in the carbon and sulfur cycle accompanying the mass extinction. The onset of a major negative carbon isotope excursion was nearly coincident with the mass extinction level, which was around the G/L boundary, implying severe destabilization of the ocean–atmosphere carbon cycle. The sulfur isotopic evidence suggests that the Guadalupian ocean was strongly stratified. The oxic bottom waters predominated over the Tieqiao area during the Laibin limestone intervals, but the redox environment may have changed due to chemocline instability. During the uppermost Laibin and lower Heshan intervals, an anoxic environment predominated during rapid transgression. An upward chemocline excursion led to significant amounts of H2S gases being released into the photic zone and even atmosphere, resulting in photic zone euxinia (Kump et al., 2005); this scenario would have been very harmful, both for the benthic and planktonic fauna. The subsequent rapid transgression resulted in oceanic anoxia, expelling the benthos from their previous ecologic niches, thereby leading to their significant
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