Constraining the Cause of the End-Guadalupian Extinction with Coupled Records of Carbon and Calcium Isotopes ∗ Adam B

Home , Capitanian, Chaotian (geology), Guadalupian, Lopingian, Wuchiapingian

Earth and Planetary Science Letters 396 (2014) 201–212

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Earth and Planetary Science Letters

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Constraining the cause of the end-Guadalupian extinction with coupled records of carbon and calcium isotopes ∗ Adam B. Jost a, ,RolandMundilb,BinHec,ShaunT.Brownd,e, Demir Altiner f, Yadong Sun g, Donald J. DePaolo d,e, Jonathan L. Payne a a Department of Geological and Environmental Sciences, Stanford University, Stanford, CA 94305, USA b Berkeley Geochronology Center, Berkeley, CA 94709, USA c State Key Lab of Isotope Geochemistry, Chinese Academy of Sciences, Guangzhou, Guangdong 510640, PR China d Department of Earth and Planetary Science, University of California, Berkeley, CA 94720, USA e Geochemistry Department, Earth Science Division, Lawrence Berkeley National Laboratory, Berkeley, CA 94270, USA f Department of Geological Engineering, Middle East Technical University, Ankara 06531, Turkey g State Key Laboratory of Geobiology and Environmental Geology, China University of Geosciences, Wuhan, Hubei 430074, PR China article info abstract

Article history: Anegativeδ13C excursion in carbonate sediments near the Guadalupian/Lopingian (Middle/Late Permian) Received 26 October 2013 boundary has been interpreted to have resulted from a large carbon cycle disturbance during the end- Received in revised form 18 March 2014 Guadalupian extinction event (ca. 260 Ma). However, the carbon isotope data alone are insufficient Accepted 6 April 2014 to uniquely determine the type and magnitude of perturbation to the global carbon cycle. Calcium Available online 3 May 2014 isotopes can be used to further constrain the cause of a carbon isotope excursion because the carbon Editor: J. Lynch-Stieglitz and calcium cycles are coupled via CaCO3 burial. In this study, we present coupled carbon and calcium Keywords: isotope records from three Guadalupian–Lopingian (G/L) sections in China (Penglaitan and Chaotian) and 13 44/40 Capitanian Turkey (Köserelik Tepe). The δ Candδ Ca records differ among our studied sections and do not Emeishan co-vary in the same manner. No section shows δ13Candδ44/40Ca changes consistent with massive, Middle Permian rapid volcanic CO2 emissions or methane clathrate destabilization. Additionally, many sections with large diagenesis (>3h) changes in δ13C exhibit δ18O evidence for diagenetic alteration. Only one section exhibits a large large igneous province excursion in the δ44/40Ca of limestone but the absence of a similar excursion in the δ44/40Ca of conodont apatite suggests that the limestone excursion reflects a mineralogical control rather than a perturbation to the global calcium cycle. Hence, we interpret the large isotopic changes observed in some sections to have resulted from local burial conditions or diagenetic effects, rather than from a large carbon and calcium cycle disturbance. Perturbations to the global carbon and calcium cycles across the G/L transition were much less intense than the disturbances that occurred across the subsequent Permian–Triassic boundary. This finding is consistent with the much smaller magnitude of the end-Guadalupian extinction relative to the end-Permian. © 2014 Elsevier B.V. All rights reserved.

1. Introduction Clapham, 2012; Stanley and Yang, 1994). Given strong evidence for a causal link between Siberian Traps volcanism and the end- 1.1. Overview Permian extinction (Payne and Clapham, 2012), the emplacement of the Emeishan large igneous province in the latest Guadalupian One of the least understood extinction events of the Phanero- raises the possibility of a similar trigger for the end-Guadalupian zoic occurred at the end of the Guadalupian Epoch (ca. 260 Ma) extinction (Courtillot and Renne, 2003). In the case of the end- (Fig. 1)(Bambach, 2002; Raup and Sepkoski, 1982). This extinc- Permian extinction, calcium and carbon isotope records provide tion event was temporally distinct, and causally independent of the evidence that the extinction event was linked to ocean acidiﬁcation subsequent end-Permian mass extinction (ca. 252 Ma) (Clapham following massive CO2 release caused by Siberian Traps volcan- and Payne, 2011; Clapham et al., 2009; Jin et al., 1994; Payne and ism (Hinojosa et al., 2012; Payne et al., 2010). Because no calcium isotope records exist for the end-Guadalupian, it is currently un- * Corresponding author. known whether a similar change in ocean chemistry occurred dur- E-mail address: [email protected] (A.B. Jost). ing this earlier biodiversity crisis. http://dx.doi.org/10.1016/j.epsl.2014.04.014 0012-821X/© 2014 Elsevier B.V. All rights reserved. 202 A.B. Jost et al. / Earth and Planetary Science Letters 396 (2014) 201–212

1.2. Calcium isotopes and the ocean calcium cycle

We express calcium isotope ratios as the relative abundance of 44Ca versus 40Ca using standard δ notation: 44Ca 40 sample δ44/40Ca = Ca − 1 × 1000 44Ca 40Ca BSE We report our δ44/40Ca data relative to bulk silicate earth (BSE) rather than seawater because the calcium isotope composition of seawater has been shown to vary over the Phanerozoic (Nielsen et al., 2011). On the BSE scale, modern seawater is approximately 0.9h, and SRM-915a is −1.0h (Nielsen et al., 2011). Calcium isotopes in CaCO3 sediments are fractionated relative to seawater by h Fig. 1. Middle–Late Permian timescale illustrating the different epoch and stage 1.0 to 1.9 during the formation of CaCO3 sediments (Blättler et nomenclature, as well as the conodont zonation for the Capitanian and early Wuchi- al., 2012; DePaolo, 2004; Gussone et al., 2005); in part, this range apingian. Ages are from the International Chronostratigraphic Chart (v2013/01) of reflects the ∼0.9h difference in fractionation between calcite and the International Commission on Stratigraphy (ICS). aragonite (Blättler et al., 2012; Gussone et al., 2005; Marriott et al., 2004; Nielsen et al., 2011). Because calcium influxes have similar Geochemical proxy data from marine limestones suggest the isotopic compositions and the ocean is well-mixed with respect end-Guadalupian extinction was associated with an environmen- to calcium due to its long residence time (Hippler et al., 2003; 44/40 tal crisis. The redox chemistry of deep-sea cherts indicates that Schmitt et al., 2001), substantial changes in the δ Ca of sea- the end-Guadalupian was the start of long-lasting and wide-spread water are likely either due to mass flux imbalances or changes in ocean anoxia, which later increased in intensity at the end of the fractionation. Changes in fractionation associated with shifts in the primary CaCO polymorph being deposited in the oceans appear to Permian (Isozaki, 1997). A negative shift in δ13C occurs in several 3 explain long-term changes in the δ44/40Ca of seawater (Blättler et stratigraphic sections near the Guadalupian/Lopingian (G/L) bound- al., 2012; Farkaš et al., 2007). ary (Bond et al., 2010b; Isozaki et al., 2007; Kaiho et al., 2005; We can distinguish whether δ44/40Ca variation in limestone Lai et al., 2008; Shen et al., 2013; Tierney, 2010; Wang et al., 2004; resulted from changing mineralogy or from changing seawater Wei et al., 2012; Wignall et al., 2009a; Yan et al., 2013). These δ13C δ44/40Ca by using other Ca-bearing minerals (Fantle, 2010). This fluctuations are consistent with some perturbation to the carbon approach has been applied to the Cenozoic calcium isotope record cycle, and have been variously attributed specifically to sea-level using marine barite (Griffith et al., 2008a, 2008b, 2011) and to change (Wang et al., 2004), ocean stratification and anoxia (Isozaki the Permian–Triassic boundary interval using conodont microfos- sils (Hinojosa et al., 2012). Here, we use the δ44/40Ca of conodonts et al., 2007; Tierney, 2010), methane release (Bond et al., 2010a; (hydroxyapatite) to differentiate changing mineralogy from global Retallack and Jahren, 2008; Retallack et al., 2006), and Emeishan changes in seawater δ44/40Ca. In the case of a global flux imbal- volcanism (Wei et al., 2012). ance, the δ44/40Ca records from carbonates and conodonts should 13 Because any negative δ C excursion may be explained through mirror each other. If the local sediment mineralogy shifted from several mechanisms, identifying the precise cause requires addi- aragonite to calcite, the δ44/40Ca of carbonates would increase but tional geochemical proxies. Calcium isotopes can provide further the δ44/40Ca of conodonts would not. However, if there was a constraint on the possible causes of δ13C excursions because the global change from aragonite to calcite deposition, we expect to 44/40 carbon and calcium cycles are linked through weathering and find a positive excursion in carbonate δ Ca, and a negative shift in conodont δ44/40Ca as seawater δ44/40Ca adjusts to the new car- the burial of CaCO (Blättler et al., 2011; De La Rocha, 2000; 3 bonate fractionation factor (Fig. 2). Holmden et al., 2012a; Payne et al., 2010). For example, the large The δ44/40Ca composition of conodonts may be influenced by 13 negative δ C excursion associated with the end-Permian event has vital effects. We attempt to minimize this issue by analyzing 13 been attributed to volcanic input of C-depleted CO2 (Payne and monogeneric samples, and assume that any vital effects are con- Kump, 2007), ocean stagnation and overturn (Grice et al., 2005; stant during our study interval. Isozaki, 1997; Kump et al., 2005; Nielsen and Shen, 2004), and collapse of the biological pump (Rampino and Caldeira, 2005). How- 2. Geologic setting ever, the associated excursions in both δ13C and δ44/40Ca can only 2.1. Penglaitan, Guangxi Province, China be accounted for by volcanic CO2 input and subsequent ocean acidification (Hinojosa et al., 2012; Payne et al., 2010). We collected limestone rock and conodont samples from the To better constrain the nature of the end-Guadalupian event, G/L global stratotype section and point (GSSP) at Penglaitan, south here we present coupled calcium and carbon isotope records for China (Fig. 3). Limestone samples were collected over 150 m of the G/L boundary interval at three different stratigraphic sections: section exposed along the Hongshui River, beginning near the base Penglaitan (Guangxi, China), Chaotian (Sichuan, China), and Kösere- of Member IV of the Maokou Formation (interbedded radiolar- lik Tepe (Adana, Turkey). We compare these new data with existing ian chert, micritic limestone, and sandstone), through Member V δ13C and δ18O data from sections in China, Iran, and Japan to il- (a.k.a., the Laibin Limestone, a massive, fossiliferous limestone), and terminating in the lower Heshan Formation (interbedded ra- lustrate the variations in these isotope proxy records. We show diolarian chert and micritic limestone) (Jin et al., 2006). The G/L that the geochemical changes during this interval are not compat- boundary lies within the uppermost Laibin Limestone, at the first ible with significant global acidification in response to Emeishan appearance datum (FAD) of Clarkina postbitteri postbitteri (Jin et al., volcanism or methane release, and that only a small carbon cycle 2006). Conodont samples were collected beginning 30 m below the perturbation, if any, could have taken place. G/L boundary and through the remainder of the section. A.B. Jost et al. / Earth and Planetary Science Letters 396 (2014) 201–212 203

Fig. 2. Predicted calcium isotope records for seawater, carbonates, and conodonts in four hypothetical scenarios: In steady state, the carbonates and conodonts are offset from seawater based on their fractionation factor. The global flux imbalance represents a temporary change in carbonate mass balance; in this case, the calcium outflux exceeds influx for a brief interval. A local mineralogy shift (here from aragonite to calcite) affects the fractionation factor for carbonate, but does not haveaneffectontheglobal calcium isotope compositions of seawater. A global mineralogy (aragonite to calcite) shift affects the fractionation factor for carbonate permanently, and changes the isotopic composition of seawater and conodonts.

nous limestone followed by massively-bedded, algal-rich wackestone and packstone containing chert beds and nodules. Conodonts are not abundant at this section, so biostratigraphic constraints are poor. The upper limestone portion of the Maokou Formation belongs to Jinogondolella postserrata zone, and the lowermost unit of the late Capitanian mudstone of the Maokou Fm. to the J. shannoni zone (Isozaki et al., 2008; Lai et al., 2008). No conodonts were recovered from the lower Wuchiaping Formation at Chaotian, nor the nearby Shangsi section (Sun et al., 2008).

2.3. Köserelik Tepe, Adana Province, Turkey Fig. 3. Late Permian paleogeographic reconstruction (Ron Blakey, NAU Geology) with the localities featured in our study: (A) Penglaitan, Guangxi, China, (B) Chaotian, The Köserelik Tepe section is located in the Adana Province of Sichuan, China, and (C) Köserelik Tepe, Adana, Turkey; as well as the localities from previous studies discussed here: (1) Tieqiao, Guangxi, China, (2) Kamuru, Miyazaki, Turkey (Fig. 3), 80 km southeast of the city of Kayseri. We sam- Japan, (3) Abadeh, Fars, Iran, and (4) Xiong Jia Chang and Gouchang, Guizhou, China. pled limestone rock over 350 m of exposed outcrop, all belonging to the Yıgılıtepe˘ Formation (Demirtaslı, ¸ 1967; Özgul et al., 1973), The Capitanian/Wuchiapingian boundary was located in the with our section beginning in the lower Capitanian (just above field according to Jin et al. (2006). The position of the Wor- the youngest quartz arenite deposit), and ending in the upper dian/Capitanian boundary was estimated based on conodont bios- Wuchiapingian. The Yıgılıtepe˘ Formation contains shallow-water tratigraphy from the nearby Tieqiao section (Mei et al., 1998). shelf limestone deposits, with abundant algae and foraminifera, along with corals, brachiopods, and trilobites in lesser abundances. The section exhibits four transgressive–regressive cycles, which are 2.2. Chaotian, Sichuan Province, China reflected in sedimentary transitions from lime mudstone to fossiliferous packstone to wackestone with some cycles capped with The Chaotian section is located in northern Sichuan Province dolomite. (Fig. 3), approximately 30 km north of the city of Guangyuan, and At Köserelik Tepe, the last appearance of fusuline foraminifera has been the focus of several sedimentological (He et al., 2010; (schwagerinid) occurs at 131.8 m. Samples between 196 to 218 m Isozaki et al., 2008, 2004; Saitoh et al., 2013b) and geochemical are rich in Shanita foraminifera, which places this part of the studies (Lai et al., 2008; Saitoh et al., 2013a). We sampled lime- section in the upper Capitanian. The level at 223 m marks the stone over approximately 100 m of exposed outcrop along the Jial- beginning of the Paraglobivalvulina mira-Reichelina Zone, which ing River, beginning in the lower Capitanian portion of the Maokou corresponds to the lower Wuchiapingian (Altiner et al., 2000; Formation, and ending in the mid-upper Wuchiapingian portion of Gaillot and Vachard, 2007). Therefore, we tentatively place the the Wuchiaping Formation. The lower 45 m of Maokou Formation Capitanian/Wuchiapingian boundary at 223 m. we sampled consists of thickly-bedded, fossiliferous packstones and wackestones, in some parts containing thick chert beds and 3. Methods large chert nodules. The upper 15 m of the Maokou Formation (up- 3.1. Sample preparation per Capitanian) consists of dark gray lime mudstones with abundant interbedded chert, phosphate nodules, and dolomite (Isozaki We cut limestone samples using a tile saw and prepared pow- et al., 2008; Saitoh et al., 2013a, 2013b). The top of the Maokou ders using a bench-top drill press equipped with a 0.8 mm dental and the G/L boundary at this location is at the base of a 1.5 m- drill bit. We collected only micrite where possible, and avoided thick volcanic ash bed, referred to as the Wangpo Tuff, although skeletal grains and sparry calcite cements. We extracted conodont this layer is also argued to be erosional products of Emeishan elements from bulk rock by dissolution with 10% acetic acid and flood basalts (He et al., 2010). The overlying Wuchiaping Formation heavy liquid separation (sodium polytungstate) at China University begins with sandy limestone, which is overlain by dark, bitumi- of Geosciences in Wuhan, China. 204 A.B. Jost et al. / Earth and Planetary Science Letters 396 (2014) 201–212

3.2. Trace element analysis 2σ < 0.1h based on duplicate measurements, we assigned them a 2σ of 0.1h, which is our long-term reproducibility of NIST-915a. Limestone powders were digested in 1 M acetic acid overnight, centrifuged, and decanted to separate the dissolved fraction from 4. Results the insoluble fraction and avoid contamination from non-carbonate phases such as clays and organic matter. Elemental concentrations 4.1. Penglaitan of Ca, Mg, Sr, and Mn were determined by inductively-coupled plasma optical emission spectrophotometry (ICP-OES) at Stanford At Penglaitan, δ13C values start near +5.3h in the upper Wor- University. dian, then decrease to values between +3.2 and +4.0h for most of the Capitanian. In the uppermost Capitanian, δ13C values in- 3.3. Carbon and oxygen isotope analysis crease slightly to just over +5.0h, and then steadily decrease to near +3.3h at the end of the section (Fig. 4). Carbon isotope val- All carbon and oxygen isotope analyses were performed at Stan- ues are typical for the Middle Permian (+3to+5h)(Korte et ford University. Three hundred and sixty-five samples were ana- al., 2005), and agree with previous measurements from Penglaitan lyzed using a Finnigan MAT252 coupled to a Kiel III carbonate de- (Chen et al., 2011; Wang et al., 2004). vice. Between 60 and 250 μg of limestone powder were reacted in The largest change in δ13C at Penglaitan is a small positive ◦ phosphoric acid (H3PO4)at70 C for 600 s. Precision on SRM-8544 shift of approximately 1h below the G/L boundary, followed by (NBS-19) and internal standard M2 were 0.06h and 0.035h (1σ ) an approximately 2h decrease until the end of the measured sec- for oxygen and carbon, respectively. tion. Our records do not extend far enough into the Wuchiapingian The remaining 105 samples were analyzed using a Thermo to indicate at what value the shift ends, though data from the Finnigan Gasbench interfaced with a Thermo Finnigan Delta Plus nearby Tieqiao section suggest δ13C returned to Middle Permian XL mass spectrometer via a Thermo Finnigan ConFlo III unit. Be- values (∼+5h)(Tierney, 2010). Carbon isotope values are signif- tween 200 and 700 μg of limestone powder were weighed into icantly positively correlated with oxygen isotope values (Table 1), sealed vials that were flushed with He gas and reacted with 0.5 but not with any other geochemical parameter (δ44/40Ca, Mn/Sr, ◦ mL of phosphoric acid for 1 h at 72 C. External precision of oxygen Sr/Ca, Mg/Ca) at Penglaitan. and carbon isotope data is <0.1h (1σ ), based upon repeated mea- Oxygen isotope values at Penglaitan vary between −8.8 and surements of two internal lab standards (calibrated against NBS 18, −1.9h, with the majority between −7.0 and −3.0h. There is no NBS 19, and LSVEC). All δ13C and δ18O values are reported relative apparent stratigraphic trend in δ18O at this section, except for a to V-PDB. small (∼1h) shift towards lower values above the G/L boundary. Oxygen isotopes at Penglaitan are not correlated with any other 3.4. Calcium isotope analysis geochemical parameter except δ13C. The bulk limestone samples from the first 128 m have no ob- All calcium isotope analyses were performed at the Center for vious trend in δ44/40Ca and have a mean value of −0.38h.Inthe Isotope Geochemistry, University of California, Berkeley. Bulk lime- uppermost Capitanian, in the Laibin Limestone, δ44/40Ca reaches stone samples were digested following the same method used for amaximumof+0.01h before returning to a mean of −0.33h the trace element analysis. After separation from the insoluble above 133 m, which is similar to the values found before the G/L fraction, limestone samples were dried and redissolved in 3.0 M boundary excursion (Fig. 4). Background values in this study are HNO3. For conodont samples, we constructed monogeneric sam- comparable to those presented in the lower-resolution Phanerozoic ples (Jinogondolella, Clarkina) using 1–5 individual elements with compilation by Farkaš et al. (2007), and also to those presented in little or no residual carbonate coating. We cleaned samples in the high-resolution end-Permian study by Hinojosa et al. (2012). three 10-min sonic baths each of acetone, 1.0 M ammonium ac- δ44/40Ca at Penglaitan is not correlated with any other geochemi- etate, and 1.0 M acetic acid (9 washes total), and then digested cal parameter. them in 6 M HNO3. The conodont data begin approximately 20 m below the bound- After complete dissolution, we spiked the dissolved samples ary and continue 10 m above the base of the Lopingian. The with a 42Ca- and 48Ca-enriched tracer to allow for correction of δ44/40Ca record from the conodonts lacks a clear positive or neg- mass-dependent fractionation occurring during the thermal ion- ative shift through the sampled interval (Fig. 4). Conodont values ization process. We loaded the spiked solution onto an element- reach a minimum of −0.92h, and a maximum of −0.51h.These specific resin column (DGA Eichrom Technologies) to isolate and values are similar to those of Late Permian and Early Triassic con- + purify Ca2 from the sample matrix. Samples were run in a odonts (Hinojosa et al., 2012). Thermo-Fisher Triton multi-collector thermal ionization mass spec- Mn/Sr values vary between 1 and 3 through most of the section. trometer (TIMS). Calcium isotope values are reported relative to Mn/Sr values are highest near the G/L boundary interval, but with Bulk Silicate Earth (BSE) (40Ca/44Ca = 47.162 and 42Ca/44Ca = most values <5. 0.31221). 4.2. Chaotian 3.5. Data analysis At Chaotian, δ13C values vary mostly between +4.0 and +4.5h Co-variation among all geochemical data was calculated using for the first 45 m of the section before beginning to drop steadily Spearman’s Rank test, and significance (p < 0.05) was assigned af- to almost +3h at 48 m. The interval between 48 m and 70 m ter a Bonferroni correction was applied to account for the number is minimally sampled due to very poor outcrop exposure in an of correlations being assessed (N = 50). organic-rich mixed chert and lime mudstone interval. Carbon iso- We fitted a trend line and a 2σ error envelope to our cal- tope values starting at 60 m suggest a 9h negative shift to −6.1h cium isotope records by resampling our record 200 times using at 62 m in the latest Capitanian. At 67 m, δ13C vary between randomly selected values for each sample level from a normal dis- −2.2h and +1.5h and rebound to pre-G/L values mostly between tribution defined by the mean and standard deviation of each sam- +4.5 and +5.0h by 74 m (Fig. 5). ple. Mean resampled values were plotted using the smooth.spline To fill in the chemostratigraphic record where we were un- function in R (smooth parameter, spar = 0.50). For samples with able to sample outcrop, we used data from Saitoh et al. (2013a), A.B. Jost et al. / Earth and Planetary Science Letters 396 (2014) 201–212 205

Fig. 4. The δ13C, δ18O, and δ44/40Ca record from Penglaitan. The ﬁlled data points are bulk carbonate data, and the x-shapes are conodont data. Conodont biozones are from Mei et al. (1998), using the stratigraphic heights from Chen et al. (2011). Error bars on the δ44/40Ca data are 2σ from replicate measurements. Numbers represent FAD of conodont species: 1 – J. shannoni;2–J. altudaensis;3–J. prexuanhaenensis;4–J. xuanhaenensis;5–J. granti;6–Clarkina postbitteri hongshuiensis;7–C. postbitteri postbitteri; 8–C. dukouensis;9–C. asymmetrica.

Table 1 Oxygen isotope values from our samples above and below the Spearman’s rank test results for δ13Cversusδ18O for each section. Significance is muddy G/L boundary interval vary between −7.3 and −3.9h. designated to p-values less than a Bonferroni-corrected 95% confidence threshold, 18 = 13 18 The poorly-exposed mudstone interval has δ O values between which is 0.05/50 0.001. Correlations between δ Candδ Oaresignificantat − h − h Penglaitan, Tieqiao (Tierney, 2010 only), Kamura, and Iran. 12.6 and 3.5 (Saitoh et al., 2013a). Where data from both studies overlap stratigraphically, δ18O values are within 1h of Section ρ p-value each other (Fig. 5). Penglaitan 0.41 <0.001 − Calcium isotope values of the limestone are constant within er- Chaotian 0.25 0.03 − h Köserelik Tepe −0.20 0.01 ror for the first 50 m of the section and average around 0.39 . Tieqiao Values above the G/L boundary begin slightly less than the pre- (Tierney, 2010)0.70 <0.001 boundary average at −0.54h,decreaseto−0.70h, and then in- Chaotian crease to −0.32h by just above 70 m. Values then decrease to (Saitoh et al., 2013a) −0.08 0.47 − h Kamura around 0.50 for the remainder of the section (Fig. 5). Most of (Isozaki et al., 2007) −0.64 <0.001 these changes are within error of each other. Iran Mn/Sr values for our Chaotian samples were all very low (Liu et al., 2012)0.53 <0.001 (<0.13). There are no published Mn/Sr data for the muddy in- Tieqiao terval. There is no significant correlation between any of the geo- (Yan et al., 2013)0.38 0.03 Guizhou chemical proxies we measured at Chaotian, nor is there between (Wignall et al., 2009a; Bond et al., 2010b)NANA the δ13C and δ18O values reported in Saitoh et al. (2013a) (Ta- ble 1). who utilized a core from the same section to measure δ13Cinthe 4.3. Köserelik Tepe poorly-exposed muddy interval in high detail. We fit the data from Saitoh et al. (2013a) into our record based on the position of the At Köserelik Tepe, δ13C values start at +4.1h and increase to G/L boundary, and made minor adjustments to match levels with values between +4.5 and +5.5h by 10 m. At 50 m, there is a similar δ13C values from both datasets (Fig. 5). The combined data 22 m stratigraphic gap due to poor exposure, and samples con- 13 show that the decline in δ C beginning at around 45 m acceler- tinue to have δ13C values between +4.0 and +5.8h range, mostly ates significantly just below 50 m (where our section ends, around >+4.8h. At 130 m, δ13C values begin to decrease to a minimum the FAD of Jinogondolella shannoni, Saitoh et al., 2013a) and reaches of +3.3h at 162 m. After a 14 m gap, values continue between a minimum of −3h. Values between 50 and 70 m vary between +4.0and+5.5h between 176 and 204 m, and again begin de- −6h and +3h, with the majority between −1h and +3h. creasing to a minimum of +1.1h. At this minimum, though, δ13C 206 A.B. Jost et al. / Earth and Planetary Science Letters 396 (2014) 201–212

Fig. 5. The δ13C, δ18O, and δ44/40Ca record from Chaotian. Error bars on the δ44/40Ca data are 2σ from replicate measurements. Open circles are data from Saitoh et al. (2013a), correlated with our section based on lithology and matches with our δ13C data. Conodont biozones are from Saitoh et al. (2013a). Numbers represent FAD of conodont species: 1 – Jinogondolella shannoni. values vary between +1.1 and +4.0h, not including one dolomi- major (>3h) δ13C excursions (Chaotian, Guizhou, and Tieqiao), tized sample (sample KST-246, Mg/Ca = 0.58) with a value of while others contain more minor excursions (1–3h)(Kamura, 13 −6.6h, located stratigraphically just above the δ C minimum. Köserelik Tepe), or no discernable excursion (0–1h) (Penglaitan, 13 After a 4.5 m gap above this dolomitized horizon, δ Cbegins Abadeh). The timing and duration of the largest excursions are not at +4.3h and gradually increases to a maximum of +7.0h at consistent across sections. Based on conodont biostratigraphy, the 307 m, and then gradually decreases to values between +4.4 and start of the excursion varies from as early as the J. postserrata zone +4.9h by 338 m, the end of the section (Fig. 6). Carbon isotopes (Chaotian, Saitoh et al., 2013a)toaslateasC. postbitteri postbitteri here are positively correlated with Sr/Ca (p < 0.001, ρ = 0.31) and negatively correlated with Mn/Sr (p < 0.001, ρ =−0.37). (Tieqiao, Yan et al., 2013), while the end varies from as early as Oxygen isotope values vary between −2.5 and −10.7h J. xuanhanensis (Guizhou, Bond et al., 2010b)toaslateasC. duk- throughout the entire section. Samples below 50 m and above ouensis (Tieqiao, Yan et al., 2013). The implied duration of this 200 m are most variable in their δ18O composition. Samples be- excursion varies by section, ranging from <1Myupto∼4My tween 50 and 200 m appear to cycle between relatively high values (Sun et al., 2010). − − h − − h ( 4.0to 2.5 ) at 50 m, to lower values ( 9.3to 5.6 )at There are significant positive correlations between δ13C and − 100 m, and then back to higher values (mostly between 5.0 and δ18O at Tieqiao (in Tierney, 2010 only) and Iran, and a negative −3.3h)(Fig. 6). Oxygen isotopes at Köserelik Tepe are significantly correlation at Kamura (Table 1). There are no δ18Odatareported positively correlated with Mg/Ca (p < 0.001, ρ = 0.49). from Guizhou. To explore which records may have been affected Calcium isotope values at Köserelik Tepe vary between −0 70h . 18 13 and −0.23h for the Capitanian. At the positive δ13Cexcursionin by diagenesis, we plotted δ Oversusδ C for every section, 18 − h the early Wuchiapingian, δ44/40Ca values decrease to −0.87h, and and identified samples with δ O values less than 7 (Figs. 7 18 then return to an average of −0.56h, though all data are within and 8). Limestone with δ O < −7h would imply precipitation ◦ error of each other (Fig. 6). Calcium isotopes do not co-vary signif- at >45 C(Hays and Grossman, 1991), assuming a δ18Oofsea- icantly with any other geochemical parameter. water equal to −1h (relative to V-SMOW) for an ice-free world Mn/Sr values at this section are mostly <0.2 and only three (Savin, 1977). While samples with δ18O greater than −7h may 13 samples are greater than 0.8. In addition to δ C, Mn/Sr also nega- have experienced diagenetic alteration of their elemental and iso- =− tively co-varies with Sr/Ca (p < 0.001, ρ 0.65). topic chemistry, we reason that samples with δ18O less than −7h

13 must have experienced such alteration because temperatures above 4.4. Comparison with previously published δ Crecords ◦ 45 C do not reﬂect seawater conditions, but formation or alteration at depth. This process highlights data points from all sec- We compiled published δ13C and δ18O records spanning the G/L interval from Tieqiao, Guangxi province (Tierney, 2010; Yan tions (except Guizhou), most severely at Tieqiao and Kamura, and et al., 2013), western Guizhou province (Bond et al., 2010a; moderately severely at Chaotian and Köserelik Tepe (Fig. 7). If we Wignall et al., 2009a), Kamura, Japan (Isozaki et al., 2007), and removed these data points, δ13C excursions at Tieqiao and Kamura Abadeh, Iran (Liuetal.,2012)(Fig. 7). Some of the sections exhibit would be absent. A.B. Jost et al. / Earth and Planetary Science Letters 396 (2014) 201–212 207

Fig. 6. The δ13C, δ18O, and δ44/40Ca record from Köserelik Tepe. Error bars on the δ44/40Ca data are 2σ from replicate measurements.

Fig. 7. Comparison of δ13C across all the sections. Red data points are samples with δ18Ovalues<−7h. Data from Guizhou are grey because no δ18O data are available. Numbers represent FAD of conodont species: 1 – Jinogondolella postserrata;2–J. shannoni;3–J. altudaensis;4–J. prexuanhaenensis;5–J. xuanhaenensis;6–J. granti; 7–Clarkina postbitteri hongshuiensis;8–C. postbitteri postbitteri;9–C. dukouensis;10–C. asymmetrica. Conodont records are as reported in Chen et al. (2011) (Penglaitan and Tieqiao), Saitoh et al. (2013a) (Chaotian), and Bond et al. (2010b) (Guizhou). 208 A.B. Jost et al. / Earth and Planetary Science Letters 396 (2014) 201–212

Fig. 9. Comparison of δ44/40Ca at all the sections in this study. Hollow circles are 13 18 Fig. 8. Scatter plot of δ Cversusδ O for the studied sections and comparison bulk limestone δ44/40Ca, and x-shapes are conodont δ44/40Ca (only at Penglaitan). sections from the published literature. Data from Guizhou are not included because The dotted horizontal line is the G/L boundary. The black ﬁt lines are made using 18 18 no δ O data are available. The dashed line is −7h for δ O. All points to the left a Monte Carlo resampling scheme and the spline.smooth function in R.Erroren- ofthedashedline(<−7h) are highlighted in red in Fig. 7. velopes are 2σ . There is no consistent pattern in δ44/40Ca among the three sections in this study, indicating that changes at each section are not related to ﬂuctua- 44/40 44/40 5. Discussion tions in global seawater δ Ca. The positive δ Caexcursioninbulklimestone with no corresponding change in conodont δ44/40Ca below the G/L boundary at Penglaitan is most consistent with local mineralogical control on δ44/40Ca in bulk 5.1. Comparison of isotopic records limestone.

Although many have argued for a global negative excursion in should be consistent across sections. However, conodont biostratig- 13 δ C at the G/L boundary (Bond et al., 2010b; Isozaki et al., 2007; raphy indicates widely varying starting and ending times for the Kaiho et al., 2005; Lai et al., 2008; Liu et al., 2012; Saitoh et al., δ13C isotope excursion, where present. Many of the sections that 2013a; Wang et al., 2004; Wei et al., 2012; Wignall et al., 2009a; have significant δ13C changes have excursions spanning at least Yan et al., 2013), our compilation of new data and data from two conodont zones, implying the event lasted more than one mil- existing sections illustrates that the negative δ13Cexcursionis lion years. not always present. Furthermore, where the negative excursion is A global shift in the δ13C composition of seawater lasting for present, the stratigraphic position and magnitude of the excursion is inconsistent. This situation is unlike the end-Permian record this long is unlikely for two reasons. Sections that do not exhibit 13 of δ13C, where many sections across the globe show large and large δ C shifts have no evidence for long depositional hiatuses 13 biostratigraphically-synchronous δ13Cexcursions(Korte and Kozur, spanning these zones. For example, there is no large δ Cshift 2010; Payne and Clapham, 2012; Shen et al., 2013). at Penglaitan, but all Capitanian conodont zones from J. shannoni The disagreement in the δ44/40Ca records among the sections to J. postbitteri postbitteri are present (Mei et al., 1998), meaning we sampled is similar to what we observed in the δ13C records the maximum length of a single time gap, if any, is less than 1 (Fig. 9). We observe a positive excursion at Penglaitan, and no ex- million years. Wignall et al. (2009b) observed a sequence boundary cursion at Chaotian or Köserelik Tepe given the error on our mea- at the base of the Laibin Limestone at Penglaitan [this has also 44/40 surements. There is no positive excursion in the conodont δ Ca been argued to be a storm deposit feature (Qiu et al., 2013)], but it 44/40 data at Penglaitan, implying that seawater δ Ca composi- could not have lasted the >1 million years required to exclude the tion was unchanging through this interval and that the observed δ13C excursion. Thus, the absence of a large excursion at Penglaitan 44/40 changes in the bulk limestone δ Ca record are diagenetically- is unlikely to be from an incomplete stratigraphic record. or facies-controlled (Fig. 2). The long duration of the possible δ13C excursion is also prob- The variability in isotopic records among sections could reflect lematic because it is extremely difficult to keep seawater DIC at a global environmental event at the G/L that was not recorded in ∼−4h for longer than its residence time (∼150 ky) without in- some areas due to stratigraphic hiatuses, or it could indicate that voking a voluminous source of 13C-depleted carbon such as vol- there was, at most, a small carbon cycle perturbation at the G/L canic CO or methane (Halverson et al., 2002). However, these pos- and that the larger changes in δ13C and δ44/44Ca at some sections 2 reflect local effects or diagenetic alteration. We discuss below why sible sources are unlikely. Numerical models suggest that volcanic 13 we prefer the latter scenario. degassing should result in a negative δ C excursion followed by 13 a positive excursion caused by input of C-depleted CO2 and sub- 5.1.1. No evidence for a global carbon/calcium cycle event sequent anoxia and high organic carbon burial (Payne and Kump, 13 If there was a global event at the G/L that was recorded in δ C 2007). Unbuffered CO2 input also decreases CaCO3 saturation (acid- + and δ44/40Ca of limestone, the timing and duration of the event ification) and increases weathering of Ca2 , thereby decreasing the A.B. Jost et al. / Earth and Planetary Science Letters 396 (2014) 201–212 209

+ relative outflux of Ca2 from the ocean (Payne et al., 2010), forcing negatively correlated, which is consistent with dolomitization of 44/40 adecreaseintheδ Ca of CaCO3 sediments. part of the section. Some have argued that the emplacement of Emeishan was a The mudstone interval at Chaotian is also rich in organic car- large source of CO2 (Ganino and Arndt, 2009; Ganino et al., 2008), bon [up to 16 wt% in some horizons (Saitoh et al., 2013b)], and has and a trigger for the extinction event (Courtillot and Renne, 2003). widely varying δ13Cvalues.Thelowδ13Cvaluesmayresultfrom However, the pattern of a linked negative and positive excursion in remineralization of 13C-depleted organic matter during deposition δ13C expected from modeling is not present in any section. Addi- or after burial. Given a lack of a negative correlation between δ13C tionally, we do not observe a negative δ44/40Ca excursion at the G/L and δ18Ointhissection(Table 1), remineralization of organic mat- boundary. This is unlike at the Permian/Triassic boundary where a ter is more likely the cause of anomalously low δ13Cwithoutlow negative excursion occurs in δ44/40Ca of bulk limestone (Payne et δ18O. al., 2010) and conodonts (Hinojosa et al., 2012). Based on these ob- Elevated Mn/Sr ratios are common in limestones with high pro- 13 servations, it is unlikely that Emeishan CO2 release impacted δ C portions of non-marine, diagenetic carbonate (Veizer, 1983). Mn/Sr or δ44/40Ca. is correlated with δ13C at Köserelik Tepe, most clearly in the sam- It is possible to generate a negative δ13C excursion of several ples below the dolomitized horizons. These samples were either permil without a detectable positive excursion by introducing a altered during exposure and meteoric alteration or during dolomi- small amount of highly 13C-depleted (δ13C −30h) carbon, such tization after burial. High Mn/Sr can also be found in samples de- as from methane clathrate destabilization (Retallack and Jahren, posited under reducing conditions. We differentiate between these 2008; Retallack et al., 2006). However, small volumes of CO2 or possibilities based on correlation with other evidence for reducing 5 CH4 need to be released extremely quickly (10 yr) to alter conditions or diagenesis. For example, Penglaitan has very elevated ocean and atmospheric chemistry. Such a scenario fails to account Mn/Sr values in some levels, but the presence of glauconite and for the long duration of the δ13C excursion at many sections. pyrite framboids and lack of correlation between Mn/Sr and any Is it possible that the error of our δ44/40Ca measurements other geochemical proxies suggest the high Mn/Sr values reflect re- obscures an important isotope excursion? Mass balance suggests ducing conditions during deposition. There are no elevated Mn/Sr that this possibility is unlikely. If we assume that the smallest values in our samples from Chaotian, but the mudstone interval δ44/40Ca fluctuation detectable is 0.2h, we can approximate what may be selectively altered. we would expect to find in δ13C. Assuming a seawater [Ca] = 2+ × 12 15 mM (Horita et al., 2002), and a Ca influx of 32 10 mol/yr 5.1.3. Potential mineralogical controls on δ44/40Ca − h at an average of 0.5 and a fractionation between seawater and Decoupling of the conodont and bulk limestone δ44/40Ca = h 2+ CaCO3 1.4 (DePaolo, 2004), the influx of Ca wouldneedto records at Penglaitan suggests a local control, potentially by a more than double for 100 kyr to yield an excursion greater than change in the dominant calcium carbonate polymorph (Fig. 2). h 2+ 0.2 . Assuming that every mole of excess Ca was a result of The oceans in the Late Permian favored aragonite as an inor- excess CO2 input to the atmosphere, a starting concentration of ganic precipitate (Stanley and Hardie, 1998); however, given that = dissolved inorganic carbon (DIC) 2 mM, and an instantaneous crinoids and brachiopods (common Paleozoic reef components) × 18 − h − h excess CO2 input of 3.5 10 mol at 15 , we expect a 7.7 construct calcite skeletons, and that inorganic calcite precipitates 13 excursion in δ C. An excursion of this magnitude should be clearly have been demonstrated to fractionate less than aragonite from recorded in many sections, but we only vaguely observe changes solution (Blättler et al., 2012; Gussone et al., 2005), the δ44/40Ca on that order at Chaotian and Guizhou, the former showing ev- excursion could be attributable to a local increase in calcite fossil idence for alteration and local effects. More depleted sources of grains in an otherwise aragonite-dominated system. 13 carbon would produce even larger excursions in δ C. The positive δ44/40Ca excursion at Penglaitan coincides with a facies change, from deep-water mudstones, to grainstones rich 5.1.2. Evidence for local effects and alteration of δ13C in crinoid, brachiopod, bivalve, and bryozoan fragments, which Changes in δ13C at the various sections may be due to local we interpret to reflect deposition in a shallow slope environment effects. Oxygen isotopes in limestone can be readily altered dur- (see supplemental Fig. A1). Modern brachiopods have a δ44/40Ca ing diagenesis (Given and Lohmann, 1985). Positive co-variation of composition similar to inorganic calcite, and echinoderms have δ13C and δ18O in some sections (Penglaitan, Tieqiao, Iran) indi- a δ44/40Ca composition between inorganic calcite and aragonite cates that many samples with lower δ13C have been recrystallized (Blättler et al., 2012), the latter possibly a result of vital effects or contain diagenetic cement (Fig. 8)(Banner and Hanson, 1990; from stronger biological control on mineralization (Böhm et al., Given and Lohmann, 1985). If we remove all δ13C values from sam- 2006; Gussone et al., 2006). Therefore, the positive excursion in ples with δ18Ovalues<−7h, assuming those samples are least δ44/40Ca of limestone at Penglaitan may have been generated by a likely to reflect seawater conditions, the negative carbon isotope proportional increase of these fossil grains in the sediment. If the excursions at Kamura and Tieqiao disappear, and Chaotian is left average aragonite grain is −0.6h, and the average calcite grain is as the only section with a significant δ13Cexcursion(Fig. 7). There 0h (Blättler et al., 2012), a mixture of 90% aragonite grains and are no δ18O data published for Guizhou, so we cannot apply this 10% calcite grains will yield an average sediment composition of method for those data. −0.33h. Increasing the proportion of calcite grains to 60% yields There is other evidence that the δ13C excursion in the bound- an average sediment composition of −0.03h. Precise sampling of ary interval at Chaotian does not reflect the δ13Cofseawater.The textures within thick-sections using a microdrill would reveal the mudstone interval, which hosts the negative δ13C excursion, con- relative contributions from these different carbonate components, tains dolomite (Saitoh et al., 2013a), which tends to have higher but this is beyond the scope of this study. δ18O values than normal marine limestone (Degens and Epstein, The Laibin Limestone contains a small volume of pore-filling 1964). Mudstone samples with higher δ18O may reflect higher lev- cement, which may be a source of the elevated δ44/40Ca values els of dolomitization. This is probably also true for the Kamura sec- (see Fig. A1). However, given that the volume of this cement is tion, where the Lopingian-age Mitai Formation is dolomitized and <25%, its δ44/40Ca composition would need to be much greater the underlying Guadalupian-age Iwato Formation is not (Isozaki than 0h to cause a 0.3h excursion in bulk limestone, which is et al., 2007), and at the Köserelik Tepe section, where δ18O and not likely given the range of modern carbonate compositions (see Mg/Ca are positively correlated. Also, δ13C and δ18OatKamuraare Blättler et al., 2012). 210 A.B. Jost et al. / Earth and Planetary Science Letters 396 (2014) 201–212

5.1.4. Potential diagenetic controls on δ44/40Ca 6. Conclusions Preferential diagenesis could also account for stratigraphic variation in δ44/40Ca. During diagenesis, equilibrium fractionation fac- New carbon and calcium isotope data from Penglaitan, Chao- 2+ tor (α) between pore fluid Ca and CaCO3 is very close to 1.000 tian, and Köserelik Tepe sections highlight the spatial and tem- (DePaolo, 2011; Fantle and DePaolo, 2007; Jacobson and Holmden, poral variability in the Guadalupian–Lopingian chemostratigraphic 2008), which makes it only possible to change the δ44/40Ca value record. The absence of any consistent and reproducible pattern in of carbonate sediment soon after burial, when the pore fluid and δ13C and δ44/40Ca among these sections and previously studied sediment are not yet equilibrated (Fantle and DePaolo, 2007). Sea- sections indicates that the excursions present reflect local condi- water and near-surface marine pore waters are enriched in 44Ca tions or diagenetic effects, and that any changes to the isotope relative to carbonate sediment, and therefore pore water may in- compositions of carbon and calcium in the global ocean at this crease the δ44/40Ca of carbonate sediment before chemical equi- time were minor and/or below detection. This does not neces- librium is reached (Fantle and DePaolo, 2007). Assuming any pore sarily exclude the possibility of an environmental perturbation at fluid has much less calcium in it than a host limestone, the max- the end-Guadalupian, but implies that the intensity of any event imum possible change in the δ44/40Ca of limestone due to equi- was small. These geochemical findings are compatible with emerg- libration is likely less than 0.1h (Fantle and DePaolo, 2007), and ing evidence from paleobiological records that the intensity of the cannot fully explain the ∼0.3h excursion we observe at Penglai- end-Guadalupian extinction was much less severe than previously tan. thought. Low sea-level and increased mixing with meteoric waters are other proposed mechanisms for changing δ44/40Ca. For example, Acknowledgements modern groundwater flowing into Florida Bay has lower δ44/40Ca values than seawater, which affects the δ44/40Ca composition of This study was supported by NSF grants 0923669 (R.M.) locally precipitated limestone (Holmden et al., 2012b). This phe- and 0923620 (J.L.P.), and by NASA grant NNX09AN6767 (J.L.P.). nomenon has been proposed as an explanation for changes to S.T. Brown and D.J. DePaolo were supported by the Director, Office δ44/40Ca in the geologic record (Holmden et al., 2012b). We expect, of Science, Office of Basic Energy Sciences, of the U.S. Department then, that lower sea-level would increase the influence of ground- of Energy under Contract No. DE-AC02-05CH11231. The authors water calcium in shallow facies. For Penglaitan, we find the most thank D.A. Fike and 2 anonymous reviewers for their constructive positive δ44/40Ca values in the shallowest marine facies. At Chao- comments; S. Denyszyn and Y. Zhong for assistance in the field; tian, we find the most negative δ44/40Ca values in shallow facies and P. Blisniuk, J.L. Hinojosa, A. Jin, G. Li, D. Mucciarone, P. Srini- above in the lowermost Lopingian (Saitoh et al., 2013b), but other vasan, and T. Owens for assistance in the laboratory. samples from similar facies below this level are more positive. As- suming all carbonate platforms have similar groundwater δ44/40Ca Appendix A. Supplementary material as Florida Bay, mixing between seawater and groundwater cannot account for stratigraphic variation in limestone δ44/40Ca. Supplementary material related to this article can be found online at http://dx.doi.org/10.1016/j.epsl.2014.04.014. 5.2. Reconciling the isotopic and fossil records References The end-Guadalupian extinction has long been considered to Algeo, T.J., Twitchett, R.J., 2010. Anomalous Early Triassic sediment fluxes due be one of the major extinctions of the Phanerozoic (Bambach, to elevated weathering rates and their biological consequences. Geology 38, 2002; Stanley and Yang, 1994). However, recent compilation of 1023–1026. taxonomically standardized Late Permian fossil occurrence data, Altiner, D., Özkan-Altiner, S., Koçyigit,˘ A., 2000. Later Permian foraminiferal biofa- which includes extensive new data from south China, reveals that cies belts in Turkey: paleogeographic and tectonic implications. 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