Earth and Planetary Science Letters 396 (2014) 201–212 Contents lists available at ScienceDirect Earth and Planetary Science Letters www.elsevier.com/locate/epsl 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 iso- topes 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 col- lapse 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 acid- ification (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.