Earth and Planetary Science Letters 419 (2015) 81–92

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Earth and Planetary Science Letters

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Effects of ocean acidification on the marine isotope record at the Paleocene–Eocene Thermal Maximum ∗ Elizabeth M. Griffith a, , Matthew S. Fantle b, Anton Eisenhauer c, Adina Paytan d, Thomas D. Bullen e a Department of Earth and Environmental Sciences, University of Texas at Arlington, TX 76019, USA b Department of Geosciences, Pennsylvania State University, University Park, PA 16802, USA c GEOMAR Helmholtz Centre for Ocean Research Kiel, Wischhofstr. 1-3, 24148 Kiel, Germany d Institute of Marine Sciences, University of California, Santa Cruz, CA 95064, USA e Branch of Regional Research, Water Resources Division, U.S. Geological Survey, Menlo Park, CA 94025, USA a r t i c l e i n f o a b s t r a c t

Article history: Carbonates are used extensively to reconstruct paleoclimate and paleoceanographic conditions over Received 8 August 2014 geologic time scales. However, these archives are susceptible to diagenetic alteration via dissolution, Received in revised form 11 February 2015 recrystallization and secondary precipitation, particularly during ocean acidification events when intense Accepted 5 March 2015 dissolution can occur. Despite the possible effects of diagenesis on proxy fidelity, the impacts of Available online xxxx diagenesis on the calcium isotopic composition (δ44Ca) of carbonates are unclear. To shed light on this Editor: G.M. Henderson issue, bulk carbonate δ44Ca was measured at high resolution in two Pacific deep sea cores Keywords: (ODP Sites 1212 and 1221) with considerably different dissolution histories over the Paleocene–Eocene 44 calcium isotopes Thermal Maximum (PETM, ∼55 Ma). The δ Ca of marine barite was also measured at the deeper Site diagenetic effects 1221, which experienced severe carbonate dissolution during the PETM. Large variations (∼0.8h) in Paleocene–Eocene Thermal Maximum bulk carbonate δ44Ca occur in the deeper of the two sites at depths corresponding to the peak carbon marine barite isotope excursion, which correlate with a large drop in carbonate weight percent. Such an effect is not observed in either the 1221 barite record or the bulk carbonate record at the shallower Site 1212, which is also less affected by dissolution. We contend that ocean chemical changes associated with abrupt and massive carbon release into the ocean–atmosphere system and subsequent ocean acidification at the PETM affected the bulk carbonate δ44Ca record via diagenesis in the sedimentary column. Such effects are considerable, and need to be taken into account when interpreting Ca isotope data and, potentially, other geochemical proxies over extreme climatic events that drive sediment dissolution. © 2015 Elsevier B.V. All rights reserved.

1. Introduction glaciations (Kasemann et al., 2005, 2014). Calcium isotopes can also shed light on recrystallization rates during marine burial dia- The Ca isotopic composition (δ44Ca) of marine of- genesis (Fantle and DePaolo, 2007; Fantle et al., 2010; Turchyn and fers the potential to reconstruct the global cycle of Ca over ge- DePaolo, 2011; Fantle, 2015). ologic time scales. Previous Ca isotope work has provided in- One advantage of using Ca isotopes to reconstruct the past is ∼ sight into seawater chemical evolution over million-year time the long residence time of Ca ( 1Ma; Berner and Berner, 1996), scales (De la Rocha and DePaolo, 2000; Fantle and DePaolo, 2005; relative to the ocean mixing time, in the global ocean. The re- sulting assumption of a homogeneous oceanic Ca reservoir sug- Heuser et al., 2005; Farkas et al., 2007; Griffith et al., 2008a, 2011; gests that the δ44Ca of synchronous minerals at various locations Fantle, 2010), the triggers of mass extinctions (Payne et al., 2010; in the ocean should record a value reflective of global seawa- Hinojosa et al., 2012), the behavior of the Earth system during ter δ44Ca, and not local or regional δ44Ca. Recent work, however, Ocean Anoxic Events (Blättler et al., 2011) and Neoproterozoic suggests complications arising from the assumption of homoge- neous oceanic δ44Ca, primarily in terms of spatial heterogeneities in nearshore settings (Holmden et al., 2012) and the impact of Corresponding author. * spatially variable fractionation factors associated with Ca removal E-mail addresses: [email protected] (E.M. Griffith), [email protected] (M.S. Fantle), [email protected] (A. Eisenhauer), [email protected] (A. Paytan), in carbonates that precipitate in different marine settings (Fantle, [email protected] (T.D. Bullen). 2010; Kasemann et al., 2014). http://dx.doi.org/10.1016/j.epsl.2015.03.010 0012-821X/© 2015 Elsevier B.V. All rights reserved. 82 E.M. Griffith et al. / Earth and Planetary Science Letters 419 (2015) 81–92

Carbonate minerals are ubiquitous in Phanerozoic marine sed- The calcareous nannofossil chalk with clay recovered from Leg iments and thus are often used as the archive for various paleo- 199 Site 1221C abruptly transitions to zeolitic clay with dramatic climatic and paleoceanographic proxies. While it has long been color variations at the Paleocene–Eocene boundary (Shipboard Sci- recognized that carbonates are susceptible to post-depositional al- entific Party, 2002b; Fig. 1). Very pale brown to brown strongly teration (e.g., Epstein et al., 1951; Urey et al., 1951), the diagenetic bioturbated chalk occurs at the base of the sampled interval. The effects on Ca isotopes are unclear. It has previously been suggested last occurrence of Paleocene benthic foraminifers is at the base of that carbonate-rich sections are generally well buffered against di- the zeolitic clay layer (1221C-11X-3 90 cm), where the chalk below agenetic alteration of δ44Ca, however the effects of extreme disso- gradually grades into zeolitic clay over a span of several centime- lution events have not been considered (Fantle and DePaolo, 2007; ters. The first occurrence of Eocene benthic foraminifers is found at Fantle, 2010; Fantle et al., 2010). It is imperative that any isotopic 1221C-11X-3 50 cm, within the chalk where CaCO3 weight percent effects associated with such events are elucidated, and their gene- reaches ∼80% (Shipboard Scientific Party, 2002b). sis understood, as such events are often the foci of isotopic proxy Nannofossils are present in some samples from Site 1221 in studies (Payne et al., 2010; Blättler et al., 2011). the interval between the last and first occurrences of Paleocene The current study investigates the effects of large-scale carbon- and Eocene benthic foraminifers (154.3 to 153.9 mbsf, respectively; ate dissolution during the Paleocene–Eocene Thermal Maximum Lyle et al., 2002). is interspersed throughout the core, but (PETM; ∼55 Ma) on deep-sea carbonate and marine barite δ44Ca never constitutes >15% of those particles documented in smear records. The PETM event was caused by a massive carbon release slides (Lyle et al., 2002). Dolomite is generally more abundant out- into the ocean–atmosphere system, resulting in ocean acidification side the carbon isotope excursion (CIE) than within the CIE. 44 and consequent widespread dissolution of marine calcium carbon- High-resolution δ Ca records derived from bulk CaCO3 (both ate (e.g., Zachos et al., 2005). The understanding of diagenetic ef- sites) and coeval barite (Site 1221C only) were determined for fects on proxy records during the PETM has a broader relevance to the PETM (Tables 1 and 2). Bulk carbonate samples were dis- similar episodes in Earth history over a range of time scales, and solved using weak (1 M) acetic acid, as previously described by offers critical insight into the expression of diagenetic effects in Fantle and DePaolo (2005). Barite was extracted using a sequential the rock record. leaching process (Paytan et al., 1993), screened for purity using a Assuming that seawater δ44Ca does not change as a result of scanning electron microscope with energy dispersive X-ray spec- perturbations that persist over time scales substantially less than troscopy, and prepared for isotopic analysis following Griffith et al. 44 40 the residence time of Ca (Komar and Zeebe, 2011), comparing syn- (2008c). Calcium isotopic composition ( Ca/ Ca) was measured chronous Ca isotope records at different open ocean locations (or by thermal ionization mass spectrometry (TIMS) at the U.S.G.S. derived from different minerals at the same locations and depth (Menlo Park) and the Helmholtz Centre for Ocean Research (Kiel, interval) can elucidate the effects of post-depositional alteration Germany) using double-spike techniques on a Finnigan MAT 261 on the proxy archives of interest. Differences in contemporaneous adjustable collector instrument and Thermo Scientific TRITON, re- 44 δ44Ca records may then be attributed to diagenetic effects, min- spectively. All data are reported in delta notation (i.e., δ Ca or 44/40 h eralogical variations, or changes in archive-specific isotopic frac- δ Ca, in permil, , units) relative to modern seawater (Griffith tionation factors (e.g., Fantle, 2010; Fantle and Tipper, 2014). By et al., 2008c). External reproducibility, determined as the aver- h contrast, similar temporal trends at multiple locations suggest with age 2σ mean of replicate analyses of individual samples is 0.15 h greater confidence that measured δ44Ca records reflect primary en- and 0.17 for Menlo Park and Kiel, respectively. External pre- vironmental signals that are globally interpretable. cision was determined using reference materials measured over The current study presents records of δ44Ca derived from bulk several years including seawater, NIST SRM 915a, La Jolla Ca, and carbonate and marine barite, 87Sr/86Sr from marine barite, carbon CaF2 (Griffith et al., 2008b, 2008c). The Sr isotopic composition (87Sr/86Sr) of barite was measured on a VG354-S TIMS at the Uni- (δ13C) and oxygen (δ18O) isotopic composition of bulk carbonate, versity of California-Santa Cruz. External reproducibility, expressed trace element composition and weight percent (wt%) CaCO3 from as 2σ , on duplicate runs is <0.00006 and measurement of two Pacific Ocean locations – Ocean Drilling Program (ODP) Leg mean NIST SRM 987 yielded 0.710741 ± 7(within run standard error). 198 Site 1212B (Shatsky Rise) and Leg 199 Site 1221C (Mahi Mahi Trace elemental ratios were determined on both a 1 M acetic Fracture Zone) – that are used to shed light on the role of dia- acid leach (Sites 1212 and 1221) and a 0.1 M ammonium acetate– genesis in influencing sedimentary δ44Ca values. The ODP sites acetic acid buffer leach (Site 1221) (Table 1), following previously investigated have considerably different dissolution histories over published methods (Apitz, 1991; Delaney and Linn, 1993). Site the PETM: Site 1212B remained above the compensation 1221 sediments were sequentially leached by ultra-pure water depth (CCD) throughout the event and Site 1221C was within the (18.2M cm) and 1 M NH OH (ion exchange step), and subse- lysocline before, below the CCD during, and above the CCD after 4 quently dissolved in the buffer. Both Sr/Ca and Mg/Ca ratios were the PETM (Zeebe et al., 2009). The data suggest that the dissolution determined on the 1 M acetic acid and 0.1 M ammonium acetate– history over the PETM systematically (and predictably) influences acetic acid buffer leaches of sediments from Site 1221, which rep- the δ44Ca of deep-sea carbonates and, possibly, marine barite. resent the carbonate component. All leaches were analyzed by in- ductively coupled plasma–atomic emission spectrometry (ICP-AES) 2. Methods at either the University of California-Santa Cruz (1221C 1 M acetic acid leach) or University of Texas at Arlington (1212B 1 M acetic Bulk carbonate and marine barite were sampled from ODP sed- acid leach and 1221C 0.1 M ammonium acetate–acetic acid buffer iment cores at Site 1212B (paleowater depth ∼2200 m; Takeda leach). External precision of the Sr/Ca and Mg/Ca ratios deter- and Kaiho, 2007) and Site 1221C (paleowater depth ∼3200 m; mined by duplicate dissolutions of samples in the 1 M acetic acid Murphy et al., 2006). The Leg 198 Site 1212B Paleocene–Eocene leach is <3% (RSD). External precision of the 0.1 M ammonium boundary interval consists predominantly of pale yellowish brown acetate–acetic acid buffer trace metal ratios determined by dupli- nannofossil ooze with clay alternating with pale orange to gray- cate sample analyses were within 8% (RSD) for Sr/Ca, but were ish orange nannofossil ooze (Shipboard Scientific Party, 2002a). An poorly replicated for Mg/Ca. abrupt contact at the base of the PETM event is thought to be an The δ18O and δ13C values of Site 1212B bulk carbonates (Ta- unconformity, however, “the significance. . . is currently not under- ble 3) were measured using a Finnigan-MAT 252 mass spectrom- stood” (Shipboard Scientific Party, 2002a). eter with a Kiel III carbonate device at the University of Florida. E.M. Griffith et al. / Earth and Planetary Science Letters 419 (2015) 81–92 83

Fig. 1. ODP Site 1221C core photograph and SEM of bulk sample (left) and barite residues in backscatter electron mode (right). weight percent of bulk samples from Murphy et al. (2006) and barite weight percent from Faul and Paytan (2005).

All C and O isotope data are reported in delta notation relative age. The resulting Site 1221 age model indicates large variations to Vienna Pee Dee Belemnite (VPDB). External reproducibility, cal- in apparent rate from 0.25 cm/ka within the CIE to culated as twice the standard deviation of replicate analyses of 1.5 cm/ka at the end of the recovery phase, similar to patterns seen NBS-19 (n = 8), is 0.04h and 0.12h for δ13C and δ18O, respec- at other sites (e.g., Röhl et al., 2007). tively. Total inorganic carbon was measured at the University of Florida using a UIC Coulometrics 5011 CO2 coulometer coupled 3. Results with an AutoMate carbonate device. Ages were determined relative to the carbon isotope excursion The CIE is well expressed at 1212B, with a pre-excursion av- (CIE) by correlating the bulk carbonate δ13C record at Site 1212B erage δ13C value of 3.21h, an approximate 3h decrease in bulk to Site 690 δ13C (Röhl et al., 2007; Table 4). Using updated 3He carbonate δ13C ∼ 50 ka after the identified onset of the excursion, ages for Site 690 (Murphy et al., 2010)results in 1212B sedimen- and a minimum value of 0.30h (Fig. 2). Over the core section of tation rates (between 0.15 and 0.22 cm/ka) that are within the 1212B analyzed, the average CaCO3 content is 94% (2RSD = 8.5%, range of previous estimates (Shipboard Scientific Party, 2002a). Site n = 30), although there is a brief drop to ∼80 wt % at 79.92 mbsf 13 1221 δ C record is significantly affected by dissolution and is not (∼15 ka after the onset of the CIE). The drop in CaCO3 content is dated in the same manner. Instead, the previously identified ben- followed by the minimum δ18O (−1.57h) at 79.90 mbsf (∼15 ka thic extinction event at Site 1221 (Nomura and Takata, 2005) and a later); pre-CIE δ18O values are −0.86 ± 0.19h (2SD, n = 14) while 3He age model for Site 1221 (Ma et al., 2014)were used to assign the latter part of the recovery period has an average δ18O of 84 E.M. Griffith et al. / Earth and Planetary Science Letters 419 (2015) 81–92

Table 1 Bulk carbonate Site 1212B and 1221C δ44Ca and trace elemental ratio data.

44 44 Sample Depth Age from δ Ca δ Ca 2σ mean n Sr/Ca Mg/Ca Sr/Ca Mg/Ca b b c c (mbsf) base CIE CaCO3 CaCO3 (mmol/mol) (mmol/mol) (mmol/mol) (mmol/mol) (ka) (rel sw) (rel NIST915a)a Leg 198 1212B, 9-5, 45–46 79.65 198 −0.85 1.03 0.05 2 1.0 6.4 1212B, 9-5, 48–49 79.68 179 −0.78 1.1 0.09 2 1.0 6.6 1212B, 9-5, 51–52 79.71 161 −0.83 1.05 0.18 2 0.9 7.0 1212B, 9-5, 54–55 79.74 139 −0.94 0.94 0.05 2 1.0 8.4 1212B, 9-5, 58–59 79.78 117 −0.97 0.91 0.28 2 0.9 9.2 1212B, 9-5, 62–63 79.82 90 −0.83 1.05 0 2 1.0 9.2 1212B, 9-5, 64–65 79.84 80 −0.69 1.19 0 2 1.0 8.5 1212B, 9-5, 66–67 79.86 70 −1.08 0.8 0.14 2 1.1 9.4 1212B, 9-5, 72–73 79.92 15 −0.94 0.94 0.14 2 1.5 7.0 1212B, 9-5, 76–77 79.96 −9 −1.17 0.71 0.05 2 1.4 5.8 1212B, 9-5, 83–84 80.03 −41 −1.04 0.84 0.05 2 1.3 6.4 1212B, 9-5, 87–88 80.07 −59 −0.97 0.91 0.18 2 1.3 5.1 1212B, 9-5, 93–94 80.13 −86 −1.27 0.61 0.23 2 1.3 5.9 1212B, 9-5, 96–97 80.16 −99 −0.92 0.96 0.09 2 1.3 5.1 1212B, 9-5, 100–101 80.2 −117 −0.71 1.17 0.05 2 1.3 6.2 Duplicate 1.3 5.9 Leg 199 1221C, 11-3, 0–3 153.4 195.9 −1.24 0.64 1 2.3 1221C, 11-3, 25–27 153.65 179.41 −1.29 0.59 1 1.3 1221C, 11-3, 35–37 153.75 172.79 −0.74 1.14 0.18 2 1.2 1221C, 11-3, 45–47 153.85 166.18 −0.83 1.05 0.23 2 1.3 1.4 5.1 1221C, 11-3, 50–52 153.9 155.62 −0.92 0.96 0.23 2 1.3 1.4 8.5 1221C, 11-3, 54–56 153.94 140.06 −0.23 1.65 1 4.6 1.5 12.7 1221C, 11-3, 58–60 153.98 124.49 −0.41 1.47 0.28 2 1.4 1.9 31.4 1221C, 11-3, 62–64 154.02 108.93 −0.51 1.37 1 2.0 5.4 162.9 1221C, 11-3, 66–68 154.06 93.37 −0.32 1.56 0.23 2 2.1 2.5 71.9 1221C, 11-3, 70–72 154.177.81 −0.37 1.51 0.23 2 3.6 9.2 164.9 1221C, 11-3, 80–82 154.238.90 −0.60 1.28 1 2.1 6.5 133.8 1221C, 11-3, 82–84 154.22 31.12 −1.20 0.68 0.18 2 2.4 7.4 401.8 1221C, 11-3, 84–86 154.24 23.34 −1.01 0.87 1 2.4 5.2 228.2 Duplicate 23.34 5.2 95.0 1221C, 11-3, 86–88 154.26 15.56 −1.06 0.82 0.05 2 2.4 6.3 294.9 1221C, 11-3, 88–90 154.28 7.77 −1.2 0.68 0.28 2 1.8 5.7 274.3 1221C, 11-3, 92–94 154.32 −1.51 −1.29 0.59 0.05 2 1.5 1.7 15.1 1221C, 11-3, 96–98 154.36 −4.29 −1.38 0.5 1 1.4 1.8 17.5 Duplicate 154.36 −4.29 −1.38 0.5 0.14 2 1.5 1221C, 11-3, 105–107 154.45 −10.56 −1.52 0.36 0.18 2 1.5 1.5 9.8 1221C, 11-3, 110–112 154.5 −14.05 −1.43 0.45 0.05 2 1.4 1.6 10.1 1221C, 11-3, 130–132 154.7 −30.25 −1.06 0.82 0.09 2 1.3 1.5 8.7 1221C, 11-3, 135–137 154.75 −33.95 −1.15 0.73 0 2 1.3 1.6 10.0

Note: all samples were measured for δ44/40Ca at USGS-Menlo Park (Bullen). a Conversion following Hippler et al. (2003). b 1 M acetic acid leach. c 1 M ammonium acetate–acetic acid buffer leach.

−0.60 ± 0.13h (2SD, n = 6). Trace element data in the bulk car- while the bulk carbonate δ44Ca record at 1221C increases system- bonate show a pattern of decreasing Sr/Ca and increasing Mg/Ca atically by up to 0.8h. beginning at 79.94 mbsf and the initiation of the CIE. Bulk carbon- By comparison, while bulk carbonate δ44Ca changes abruptly 44 ate in 1212B has an average δ44Ca of −0.93 ± 0.32h (2SD, n = 15) and substantially over the PETM at 1221C, bulk carbonate δ Ca at (Fig. 2). 1212B is relatively invariant during the CIE of the PETM but shows − − h The δ13C and δ18O bulk carbonate and species-specific benthic an apparent increase from 1.2 to 0.8 over the PETM (from ∼ foraminiferal records previously published for 1221C (Fig. 3) do just before the CIE to 170 ka after the boundary). not exhibit the expected large CIE or decrease in δ18O (Nunes and Norris, 2005). Trace element data in the acetic acid leaches from 4. Discussion 1221C show large but variable increases in Sr/Ca values from 1.2 The lack of agreement in absolute values or trends in the bulk to 4.6 mmol/mol (Fig. 3) opposite to that seen in 1212B. Analysis carbonate δ44Ca data from Sites 1221 and 1212 indicate that one of the ammonium acetate–acetic acid buffer leach at 1221C also or both of the Ca isotopic records are affected to some degree exhibit elevated Mg/Ca values from 5 to ∼400 mmol/mol (Fig. 3). by diagenesis; the depth interval over which the records differ Mean bulk carbonate δ44Ca at Site 1221C is identical to Site suggests that this effect is associated with carbonate dissolution 1212B, but exhibits high variability considering the time scale of induced by ocean acidification. The 1221C bulk carbonate δ44Ca − ± h = <0.5Maover the interval studied ( 0.96 0.80 2SD, n 22; record exhibits a large, abrupt, and relatively long-lasting positive 44 Fig. 3). Marine barite δ Ca from Site 1221C has an average value excursion (on the order of +0.8h close to the onset of the CIE) of −1.99 ± 0.70h (2SD, n = 27) over the same time period. Site that is markedly different from the 1221C barite and the 1212B 44 1221C carbonate and barite exhibit different δ Ca trends over the bulk carbonate records (Fig. 3). 44 PETM; most notably, the records diverge during the CIE as CaCO3 We contend that the shift to heavier δ Ca values in 1221C decreases from ∼80% to a few wt%. The barite δ44Ca record shows is not an original environmental signal that reflects temperature significant variations at the boundaries of the dissolution event, or species dependent changes in isotopic fractionation or a real E.M. Griffith et al. / Earth and Planetary Science Letters 419 (2015) 81–92 85

Table 2 Marine barite Site 1221C.

44 44 44 b e 87 86 Sample Depth Age from δ Ca δ Ca δ Ca 2σ mean n Sr/Ba Sr/ Sr 2σ mean n (m) base CIE barite ‘seawater’ ‘seawater’ (mmol/mol) (mbsf) (ka) (rel sw) (rel sw) (rel NIST915a)a Leg 199 1221C, 11-3, 0–3 153.40 195.9 −1.95 0.06 1.94 0.18 2 0.70776 1 1221C, 11-3, 25–28 153.65 179.41 −1.76 0.25 2.13 0.30 2 37.9 0.70776 1 1221C, 11-3, 35–38 153.75 172.79 −1.99 0.02 1.90 0.24 2 35.5 0.70776 1 1221C, 11-3, 45–48 153.85 166.18 −1.95 0.06 1.94 0.20 2 34.4 0.70775 1 1221C, 11-3, 50–54 153.90 155.62 −1.84 0.17 2.05 0.18 2 37.4 0.70779 1 1221C, 11-3, 54–58d 153.94 140.06 −1.09 0.92 2.80 0.20 2 36.7 0.70775 1 Repeatd 153.94 140.06 −1.36 0.65 2.53 0.04 2 1221C, 11-3, 58–62 153.98 124.49 −2.25 −0.24 1.64 0.35 3 32.8 1221C, 11-3, 62–66d 154.02 108.93 −1.37 0.64 2.52 0.71 (2) 38.4 1221C, 11-3, 66–70 154.06 93.37 −1.93 0.08 1.96 0.16 2 33.6 0.70775 1 1221C, 11-3, 70–72 154.10 77.81 −1.95 0.06 1.94 0.18 3 41.2 0.70776 1 1221C, 11-3, 80–82 154.20 38.90 −2.07 −0.06 1.82 0.01 2 36.1 0.70775 1 1221C, 11-3, 82–84 154.22 31.12 −2.03 −0.02 1.86 0.28 3 37.8 0.70775 0.00006 3 Repeat 154.22 31.12 −2.08 −0.07 1.81 0.18 (2) 1221C, 11-3, 84–86d 154.24 23.34 −2.44 −0.43 1.45 0.20 2 27.3 0.70776 0.00004 4 Repeatd 154.24 23.34 −2.59 −0.58 1.30 0.18 (3) 1221C, 11-3, 86–88d 154.26 15.56 −2.37 −0.36 1.52 0.02 2 37.4 0.70777 0.00001 3 Repeatd 154.26 15.56 −2.09 −0.08 1.80 (1) 1221C, 11-3, 88–90 154.28 7.77 −2.20 −0.19 1.69 0.20 (2) 37.4 0.70780 0.00000 3 1221C, 11-3, 92–94 154.32 −1.51 −2.35 −0.34 1.54 0.03 2 37.6 0.70776 0.00002 2 1221C, 11-3, 96–98 154.36 −4.29 −2.26 −0.25 1.63 0.16 3 34.3 0.70778 0.00004 2 1221C, 11-3, 105–108 154.45 −10.56 −1.89 0.12 2.00 0.03 2 39.3 0.70775 1 1221C, 11-3, 110–113 154.50 −14.05 −2.03 −0.02 1.86 0.09 2 40.0 1221C, 11-3, 130–133d 154.70 −30.25 −1.55 0.46 2.34 0.40 2 38.9 1221C, 11-3, 135–138d 154.75 −33.95 −2.42 −0.41 1.47 0.17 2 35.4 0.70775 1 1221C, 11cc, 0–3c 154.82 −39.14 −1.79 0.22 2.10 0.11 2 38.6 0.70777 0.00001 3 1221C, 11cc, 9–12c 154.91 −45.81 −2.03 −0.02 1.86 0.17 2 37.9 0.70773 0.00001 2 a Conversion following Hippler et al. (2003). b Samples in parentheses were measured at USGS-Menlo Park; others were analyzed in Kiel, Germany. c Indicates samples with ages extrapolated from sedimentation rates established by He ages in samples further upcore (Ma et al., 2014). d Indicates samples that are considered outliers or had high errors. e Paytan et al. (2007).

Table 3 13 18 Site 1212B data: carbonate (wt% CaCO3, δ C, δ O).

13 18 Sample Depth Age from base CaCO3 δ C δ O (mbsf) CIE (ka) (wt%) CaCO3 CaCO3 (rel VPDB) (rel VPDB) Leg 198 1212B, 9-5, 40–41 79.60 229 97.85 1.80 −0.64 1212B, 9-5, 43–44 79.63 210 94.58 1.66 −0.68 1212B, 9-5, 45–46 79.65 198 96.77 1.55 −0.64 1212B, 9-5, 48–49 79.68 179 96.51 1.58 −0.50 1212B, 9-5, 50–51 79.70 167 95.80 1.52 −0.58 1212B, 9-5, 51–52 79.71 161 95.44 1.40 −0.57 1212B, 9-5, 54–55 79.74 142 94.46 0.95 −0.75 1212B, 9-5, 56–67 79.76 130 91.11 1.04 −0.95 1212B, 9-5, 58–59 79.78 117 93.15 0.64 −1.04 1212B, 9-5, 60–61 79.80 103 94.15 0.69 −1.04 1212B, 9-5, 62–63 79.82 90 94.50 0.44 −1.21 1212B, 9-5, 64–65 79.84 83 91.46 0.43 −1.37 1212B, 9-5, 66–67 79.86 70 90.70 0.32 −1.42 1212B, 9-5, 68–69 79.88 50 89.33 0.30 −1.55 1212B, 9-5, 70–71 79.90 30 85.34 0.81 −1.57 1212B, 9-5, 72–73 79.92 15 80.10 2.27 −1.31 1212B, 9-5, 74–75 79.94 0 92.45 2.89 −0.97 1212B, 9-5, 76–77 79.96 −9 93.19 3.08 −0.96 1212B, 9-5, 78–79 79.98 −18 94.90 3.01 −0.99 1212B, 9-5, 80–81 80.00 −27 94.58 3.11 −0.94 1212B, 9-5, 83–84 80.03 −41 95.62 3.18 −0.77 1212B, 9-5, 85–86 80.05 −50 95.54 3.14 −0.89 1212B, 9-5, 87–88 80.07 −59 95.91 3.30 −0.89 1212B, 9-5, 89–90 80.09 −68 95.46 3.36 −0.84 1212B, 9-5, 91–92 80.11 −77 97.21 3.22 −0.82 1212B, 9-5, 93–94 80.13 −86 96.26 3.26 −0.80 1212B, 9-5, 96–97 80.16 −99 95.23 3.25 −0.84 1212B, 9-5, 100–101 80.20 −117 96.08 3.33 −0.71 1212B, 9-5, 104–105 80.24 −135 95.90 3.30 −0.73 1212B, 9-5, 110–111 80.30 −162 96.28 3.22 −0.96 86 E.M. Griffith et al. / Earth and Planetary Science Letters 419 (2015) 81–92

Table 4 Carbon isotope tie points from ODP Site 690 and ages for correlation and dating ODP Leg 198 Site 1212B following Röhl et al. (2007). Tie 690 mbsf 1263 mcd 1212B mbsf Age from base CIE Age from base CIE pointsa (Zachos et al., 2005) (Zachos et al., 2005) This study (ka) (ka) (Röhl et al., 2007) (Murphy et al., 2010) H 166.13 333.14 196.87 G 167.12 333.73 153.50 217 F 169.05 334.56 79.70 94.23 167 E 169.39 334.75 79.76 81.17 130 D 169.56 335.10 79.82 71.25 90 C 170.02 335.22 79.86 42.38 70 B 170.33 335.39 79.90 21.90 30 A 170.63 335.68 79.92 0.75 P/E 170.64 335.69 79.94 0.00 0 A- 171.24 336.00 79.98 −30.80 −18 B- 172.81 80.09 −102.53 Note ages were linearly extrapolated between points using ages from Murphy et al. (2010). a Letters after Bains et al. (1999) as indicated by Röhl et al. (2007).

13 18 Fig. 2. ODP Site 1212B (a) δ C of bulk CaCO3 relative to VPDB, (b) δ O of bulk CaCO3 relative to VPDB, (c) Sr/Ca (blue circles) and Mg/Ca (red triangles) of bulk CaCO3 44 in mmol/mol, weight percent (wt%) CaCO3 (black), and (d) bulk carbonate δ Ca relative to modern seawater. Error bars are the precision of each sample calculated as the 44 2σ mean of replicate analyses or the average 2σ mean if only one analysis was made. δ Ca curve is cubic smoothing spline of data. vertical line is an average value of modern bulk carbonate (Fantle and DePaolo, 2005). Grey shaded area represents the “core” of the δ13C excursion during the PETM interval (Röhl et al., 2007). “Recovery Phase I” is indicated (green shading) as defined by Röhl et al. (2007) as a rapid rise in δ13C but low carbonate. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.) change in seawater δ44Ca due to changes in input sources. If the which is reflected in the rapid drop in carbonate content in the observed shift is attributed to surface ocean warming this would sedimentary section (Fig. 3). Benthic δ18O values at ∼154 mbsf ◦ imply an unlikely 40 C increase in temperature (e.g., Gussone et increase by about 1h, inconsistent with calcification of benthic al., 2005). The time scale of the ocean acidification event suggests foraminifera in warming bottom waters and the inferred rela- that the global ocean δ44Ca would not have changed significantly tionship between δ44Ca and temperature. Mineralogical analyses over the PETM (Komar and Zeebe, 2011); however, it is possible document the presence of diagenetic minerals such as dolomite that local differences in the δ44Ca of surface waters could explain and authigenic apatite over the PETM interval at Site 1221, which our observations. While attractive, there is no clear mechanism is consistent with the formation of authigenic minerals in the for explaining such a large shift to higher values, particularly in sedimentary section. Perhaps most convincingly, the sedimentary the open ocean. The signal clearly did not result from mixing of record at Site 1212 exhibits a minimal loss of carbonate over the mineral sources with different isotopic compositions, as there are PETM and no evidence of a comparable offset in δ44Ca. Because no source materials on Earth capable of increasing the δ44Ca of diagenetic alteration appears a feasible explanation for the obser- surface waters (Fantle and Tipper, 2014). A local Ca isotopic distil- vations at 1221, we explore this hypothesis in detail below. lation effect, on the other hand, due to rapid and significant up- take of Ca by marine calcifiers contradicts the widely held opinion 4.1. Potential influence of carbonate dissolution on sedimentary δ44Ca that biological calcification in the surface ocean decreases during ocean acidification events (e.g., Hönisch et al., 2012 and references Carbonate dissolution alone is not a feasible mechanism for therein). generating sizeable Ca isotopic effects in the solid phase. Partial Accordingly, there is considerable evidence supporting the con- dissolution of foraminiferal tests of G. sacculifer and N. pachyderma tention that the δ44Ca offset observed at 1221C reflects post burial has demonstrated the relative insensitivity of foraminiferal δ44Ca diagenetic effects related to ocean acidification during the PETM. to partial dissolution (Gussone et al., 2009). No Ca isotope mea- Site 1221 was below the CCD during the CIE (Zeebe et al., 2009), surements of partially dissolved coccoliths or nannofossil ooze E.M. Griffith et al. / Earth and Planetary Science Letters 419 (2015) 81–92 87

13 18 Fig. 3. ODP Site 1221C (a) δ C of bulk CaCO3 (b) δ O of bulk CaCO3 (black) and benthic foraminifera (red and blue) relative to VPDB (Nunes and Norris, 2005), (c) Sr/Ca (mmol/mol) of bulk CaCO3 (open blue circles), Sr/Ca (mmol/mol) of calcite fraction (blue), Mg/Ca (mmol/mol) of calcite fraction (red triangles), and weight percent (wt%) 44 44 44 CaCO3 in black (Murphy et al., 2006), and (d) barite δ Ca relative to modern seawater (gray) and bulk carbonate δ Ca (black). Error bars are plotted as in Fig. 1. δ Ca curve is a cubic smoothing spline of data (Table 2). Dashed vertical line is average modern coretop barite (Griffith et al., 2008c). Solid vertical line is an average value of modern bulk carbonate (Fantle and DePaolo, 2005). See Fig. 2 for description of shading. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.) exist, but dissolution of carbonates has been hypothesized to im- the PETM (Fig. 1). Apatite is present (≤ 5% in smear slides) from pact Ca isotopes minimally (e.g., Langer, 2005; Fantle and DePaolo, 154.10 to 154.00 mbsf, which corresponds to the end of the CIE. 2007). It also stands to reason that if dissolution did fractionate Ca Diagenetic textures and overprinting of δ18O, noted previously isotopically, any fractionated surface generated will be removed as (Nomura and Takata, 2005; Nunes and Norris, 2005), support the the dissolution front advances into the particle. So while a dissolu- hypothesis of substantial recrystallization of, or secondary calcite tion effect on δ44Ca might be noticeable in small particles (highest precipitation associated with, foraminiferal tests. In addition, the surface area to volume) during minor dissolution events, such a bulk carbonate δ18O values at 1212, relative to 1221, indicate sub- signal is likely to be removed entirely from the record during the stantial diagenetic overprinting of δ18O to lower values at 1221 sort of aggressive dissolution associated with the PETM ocean acid- (Fig. 4). Both molar Sr/Ca and Mg/Ca ratios in the acetic-acid solu- ification event. ble fraction are high within the section of core affected by carbon- Relatedly, though distinct from dissolution as a fractionating ate dissolution (i.e., chemical “burndown”; Fig. 3). At face value, mechanism, it is feasible that carbonate dissolution preferentially these data imply recrystallization or precipitation of secondary car- accesses isotopically distinct, and differentially soluble, carbonates bonates in the presence of pore fluids with elevated Sr and Mg in the sediment column (i.e., selective dissolution). Such a hy- concentrations, perhaps influenced by both carbonate and clay dis- pothesis requires that there is substantial isotopic heterogeneity solution. However, as there is no clear correlation between the in foraminiferal and/or nannofossil carbonate, a result that is not trace metal ratios and δ44Ca values, we contend that while clays consistent with currently available data. While some work sug- may have influenced pore water Sr/Ca and Mg/Ca they do not con- gests significant Ca isotopic differences between the outer (ga- trol the δ44Ca values of the carbonate. metogenic) calcite layer and the inner (ontogenic) calcite lay- We suggest that the high δ44Ca values in the acetic acid leach ers of two planktonic foraminifers (G. truncatulinoides and G. tu- reflect either re-equilibration of existing carbonates with high mida) of up to 3.7h (Kasemann et al., 2008), dissolution studies δ44Ca pore fluids and/or precipitation of authigenic minerals (car- of planktonic foraminiferal tests indicate that partial dissolution bonates or other phases) from high δ44Ca pore fluids. The mass of does not necessarily access such heterogeneity, if it is widespread authigenic mineral formed, in the latter case, would have to be (Gussone et al., 2009). Thus, the bulk of published work to date comparable to the mass of pre-existing carbonate and the frac- does not support the expression of large Ca isotopic heterogeneity tionation factor associated with authigenic precipitation assumed in carbonates due to partial dissolution (e.g., Gussone et al., 2003, to be close to one to explain the δ44Ca data (i.e., Fantle and De- 2007; Kisakürek et al., 2011; Müller et al., 2011). Paolo, 2007; Fantle, 2015). There are two options for the timing of authigenic mineral precipitation: during and/or after the period 4.2. Potential influence of secondary carbonate precipitation on of chemical burndown. If carbonates are the primary authigenic sedimentary δ44Ca phase, then it is most likely that they formed after chemical burn- down ceased, perhaps associated with the observed and modeled There is both mineralogical (e.g., authigenic minerals iden- global overshoot of the CCD during the recovery phase (e.g., Zeebe tified from smear slides including ‘sugary coatings’ noted on et al., 2009; Zeebe and Zachos, 2013). foraminiferal tests) and geochemical (e.g., carbonate δ18O and trace The geochemical composition of an authigenic mineral depends element composition) evidence at 1221C for secondary mineral on the pore fluid composition and the partition coefficient or frac- precipitation in the sedimentary column over the PETM interval. tionation factor associated with precipitation (e.g., Fantle, 2010; Both authigenic apatite and dolomite occur in the depth intervals Fantle et al., 2010; Fantle and Higgins, 2014). In general, the pore near the CIE (Lyle et al., 2002). Dolomite is most abundant outside fluid composition is dictated by the mass fluxes into and out of of the CIE interval at 1221 (∼15% in smear slides at 154.34 and the pore fluid (i.e., the rates of dissolution, ion exchange, and pre- 153.7 mbsf), yet is often absent or constitutes <5% of the parti- cipitation) balanced against the rates at which transport processes cles found in smear slides within the clay layers associated with operate (i.e., diffusion and advection). When transport processes 88 E.M. Griffith et al. / Earth and Planetary Science Letters 419 (2015) 81–92

13 Fig. 4. Results versus age from the base of the Carbon Isotope Excursion (a) ODP Sites 1212B (black) and 1221C (gray) δ C of bulk CaCO3 relative to VPDB, (b) 1212B (black) 13 and 1221C (gray) δ C of bulk CaCO3 relative to VPDB, (c) Sr/Ca (mmol/mol) of bulk CaCO3 (open blue circles), Sr/Ca (mmol/mol) of calcite fraction (blue), Mg/Ca (mmol/mol) 44 of calcite fraction (red triangles) from 1221C, (d) weight percent CaCO3 1212B (black) and 1221C (gray), (e) bulk carbonate δ Ca relative to modern seawater from 1212B (solid black circles and long dashed line) and 1221C (open black circles and short dashed line) and 1221C ‘seawater’ δ44Ca (gray) calculated from barite data relative to modern seawater. Error bars and curves are plotted as in Fig. 2. Dashed vertical line is average modern seawater value based on the barite data (Griffith et al., 2008c). Solid vertical line is an average value of modern bulk carbonate (Fantle and DePaolo, 2005). See Fig. 2 for description of shading. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

+ Fig. 5. Schematic trends in pore fluid (a) Ca concentration [Ca2 ] and (b) Ca isotopic composition (δ44Ca) versus depth in sediments (over approximately 1 m) resulting from a carbonate dissolution perturbation associated with ocean acidification during the PETM. Schematic pore fluid profiles are displayed for three different times: a pre-carbon isotope excursion (CIE) steady state, the burndown associated perturbation, and a return to steady state. Note that the return to steady state isotopic gradient depends on whether or not the sediment and pore fluid are equilibrated at ∼1m; in the schematic shown, isotopic equilibration is assumed at ∼1m. Black dashed line in b is modern seawater δ44Ca for reference. The fractionation factor associated with authigenic precipitation is assumed to be close to one. are fast relative to reaction (i.e., in an open system), the system to understand the evolution of pore fluid δ44Ca during dissolution is said to have a large reactive length scale and the fluid can be events similar to the one associated with the PETM. maintain considerably out of equilibrium with respect to the solid. In nannofossil ooze sections, there is a strong δ44Ca gradient When reaction rate is fast compared to transport, then the sys- in the uppermost ten to fifteen meters of the sedimentary sec- tem behaves more like a closed system and has a smaller reactive tion (Fig. 5b ‘Pre-CIE Steady State’), which reflects the balance length scale. If there is isotopic fractionation or a strong elemen- between diffusive communication with seawater (defined as 0h tal partitioning behavior during mineral precipitation, the fluid can in this study) and the dissolution/recrystallization of carbonates be modified or distilled. Because pore fluid δ44Ca ultimately deter- (δ44Ca ≈−1.3h on the seawater scale; Fantle and DePaolo, 2007; mines the isotopic composition of authigenic minerals, it is critical Fantle et al., 2010; Fantle, 2015). Such a boundary layer will al- E.M. Griffith et al. / Earth and Planetary Science Letters 419 (2015) 81–92 89 ways exist in the presence of recrystallizing sediments, though the of the seawater–sediment interface. This hypothesis does not re- depth range over which it extends can change to reflect (primarily) quire mineral formation during a transient state, though such a changes in the reactive flux of solute to the pore fluid. In a recrys- scenario is plausible. In either case, the measured bulk carbonate 44 tallizing column, for instance, the boundary layer thickness (LCa) δ Ca values are consistent with secondary mineral formation, and 1/2 can be approximated by LCa ≈ (D/RMK) , where D is the diffu- possibly extensive recrystallization, in the sedimentary column af- sion coefficient of Ca appropriate for the sedimentary column and ter the cessation of carbonate burndown. RMK is a reactive term with units of inverse time (Fantle et al., Above the clay-rich PETM interval, the return to typical carbon- 2010; Fantle and Higgins, 2014). Accordingly, during ocean acid- ate δ44Ca values reflects the reestablishment (and dominance) of ification associated with the PETM, this gradient may have been the flux of surface ocean-derived carbonate particles. This hypoth- compressed due to massive and rapid carbonate dissolution, shift- esis accounts for the entire apparent bulk carbonate δ44Ca “ex- ing pore fluid δ44Ca values in the upper meter of the sedimentary cursion”, explains why the highest sedimentary δ44Ca values are column towards (but not necessarily identical to) bulk carbonate offset in the depth realm from the onset of the PETM, and even values (Fig. 5b ‘Perturbation’). It should be clearly stated, however, accounts for the few outlier barite samples, including both anoma- that while burndown can compress the boundary layer, there will lously high and low δ44Ca values in marine barites at either end of always be a δ44Ca gradient in pore fluids at the top of the sedi- the CIE (see discussion below). Further, while it is unclear whether mentary section. elevated Sr/Ca ratios represent a pure carbonate fraction, such a Given the corrosive nature of pore fluids during chemical burn- mechanism (in addition to increases in partition coefficients) can down, it is unlikely that authigenic carbonates formed during the also explain high authigenic trace element ratios. burndown. Instead, as the CCD deepened and dissolution in the At 1221C, benthic foraminiferal δ18O values increase by ∼0.9h, sedimentary column slowed, pore fluid δ44Ca evolved towards though this is not pervasive throughout the PETM interval. This steady state, generating a transient pore fluid profile in the process signal is only clearly observed over a small depth interval co- (Fig. 5b ‘Return To Steady State’). During this interval, corrosive incident with the maximum bulk carbonate and marine barite conditions would have diminished, making it possible to form sec- δ44Ca, and thus is inferred to have occurred close to the seawater– ondary carbonates at various depths in the sedimentary column. sediment interface. Benthic foraminiferal δ18O values at 1221C are Assuming a fractionation factor close to one for inorganic carbon- inconsistent with δ18O values predicted for tests that formed, or ate precipitation at slow rates (Fantle and DePaolo, 2007; Tang et underwent diagenesis, at temperatures that were warmer than al., 2008; Fantle, 2015), the resulting authigenic carbonates would those during initial formation (i.e., “typical diagenesis”; Schrag et have δ44Ca values reflective of the pore fluid δ44Ca. In the diffu- al., 1995). Taken at face value, benthic foraminiferal δ18O suggest sive boundary layer near the top of the section, pore fluid δ44Ca diagenesis under colder conditions than those of formation, a sce- would have been seawater-like (Fig. 5b ‘Return To Steady State’). nario not consistent with deep water warming associated with the With increasing depth (and an increasing extent of equilibration PETM. between solid and pore fluid), pore fluid δ44Ca values would tran- An alternative explanation is that the benthic foraminifera δ18O sition to those of the bulk carbonate (Fantle and DePaolo, 2007; reflects interaction with pore fluids with δ18O values slightly Fantle et al., 2010). higher than those of the bottom waters, or either increased bot- The Ca isotopic composition of 1221 sediments over the PETM, tom water δ18O or lower pH at the end of the PETM. While an therefore, are most simply interpreted as reflecting precipitation of alteration of pore fluid δ18O can technically occur due to carbon- authigenic carbonate within the diffusive boundary layer after the ate dissolution, mass balance suggests that it is highly unlikely cessation of chemical burndown. Assuming a typical meter-scale that carbonate dissolution could impact pore fluid δ18O on the or- reactive boundary layer in the sedimentary column, this hypoth- der of a permil. This constraint, plus the inferred proximity to the esis requires authigenic precipitation of CaCO3 to have occurred seawater–sediment interface, suggests that it is most likely that an over a limited depth range (order 10 cm) at the percent level and increase in fractionation factor associated with a decrease in pH not over the entire interval impacted by burndown. There is little (Zeebe, 2007)or an increase in bottom water δ18O are responsible evidence that other authigenic phases that may have formed dur- for the benthic foraminiferal δ18O trend. While it is also possi- ing burndown, such as dolomite or apatite, control the observed ble that the benthic fraction analyzed contained dolomite, which trends in δ44Ca. has a larger fractionation factor than calcite at a given temperature Alternatively, the data can be explained by enhanced recrys- (Horita, 2014), no dolomite rhombs were observed in the benthic tallization rates in sediments affected by burndown, effectively fraction analyzed (Nunes and Norris, 2005). decreasing the reactive length scale. If the latter scenario is con- In contrast to the benthic foraminiferal tests, bulk carbonate sidered, and the measured 1221 δ44Ca values assumed to reflect δ18O values within the recovery interval at 1221 are isotopically accurately the depth over which the diffusive boundary layer ex- lighter than contemporaneous carbonates at 1212 (Fig. 4). This tends (∼30 cm; Fig. 3), then this would suggest recrystallization suggests that bulk sediments at 1221 are likely also impacted by − − rates >10 4 a 1, or more than 1000 times faster than “typical” re- diagenetic alteration at temperatures warmer than those associ- crystallization rates in bulk nannofossil ooze (Richter and DePaolo, ated with formation (e.g., Schrag et al., 1995). Regardless of the 1987, 1988; Richter and Liang, 1993; Richter, 1996; Fantle and De- timing, both bulk carbonate and benthic foraminiferal calcite are Paolo, 2007; Fantle, 2015). Though this scenario seems improbable, clearly affected by secondary calcite precipitation, in contrast to such an interpretation is generally supported by the rather large bulk carbonates at 1212. This supports the contention that the increase in benthic δ18O at ∼154 mbsf. δ44Ca values observed at 1221 reflect secondary carbonate precip- Accordingly, the bulk carbonate δ44Ca at Site 1221 can be in- itation and not a primary signal. terpreted in terms of a depth, and not time, variant signal related to the saturation state of ocean bottom waters that is not reflec- 4.3. Implications for Ca isotope proxy development tive of the global Ca cycle. Bulk carbonate δ44Ca at the base of the PETM interval may be relatively unaltered, or the result of The large differences in δ44Ca measured in the bulk carbon- secondary carbonate formation from a pore fluid impacted by re- ate records from 1212 and 1221 suggest that a δ44Ca carbonate action (be it dissolution or recrystallization) with bulk carbonate. record from a single site may not be suitable for global reconstruc- The transition to higher δ44Ca would then reflect the formation of tions, especially during climatic or environmental perturbations authigenic minerals in the diffusive boundary later within ∼10 cm such as ocean acidification events. Certainly, as suggested previ- 90 E.M. Griffith et al. / Earth and Planetary Science Letters 419 (2015) 81–92 ously (Fantle, 2010), this issue should be resolved by measuring fluid δ44Ca over the course of the PETM. Such an interpretation δ44Ca records from multiple sites and proxy archives in order to agrees with the explanation of a depth, and not time, variant sig- reconstruct the past accurately. nal recorded in the bulk carbonate δ44Ca put forth above. At the In a similar fashion, the data presented demonstrate how the base of the CIE, assuming an equilibrium fractionation factor of analysis of coeval minerals (e.g., carbonate and barite from the −2.01h (Griffith et al., 2008c), authigenic barite appear to reflect same sample) can be used to identify diagenetically altered sam- isotopically light pore fluid δ44Ca values, which have been influ- ples. Given that barite is less susceptible to post depositional di- enced by the dissolution of (or isotopic equilibration with) bulk agenesis in the absence of sulfate reduction (Paytan et al., 1993) carbonate. and is more similar to the 1212B carbonate record, we conclude The relatively high δ44Ca values in the barites at ∼154.0mbsf that the bulk carbonate δ44Ca at 1221C cannot represent a global are somewhat more difficult to explain. We hypothesize that these signal and is therefore impacted significantly by diagenesis. barites precipitated in the sedimentary column from pore fluids Likewise, the increase of 0.4h in bulk carbonate δ44Ca at 1212 that were isotopically similar to seawater (i.e., within the diffu- cannot represent a global seawater signal, as the magnitude is too sive boundary layer), following the interpretation of bulk carbonate large and the time scale too short relative to the residence time δ44Ca discussed above. If a fractionation factor of −2.01h is as- of Ca in the Paleogene ocean to represent the evolution of global sumed, this implies that the fluids from which the barites formed seawater. If, as suggested for Site 1221, an increase in δ44Ca of bulk were at least +1h relative to seawater. Such a high value for pore carbonate over time scales much less than the residence time of Ca fluid and/or seawater δ44Ca seems improbable. Thus, the data ap- in the oceans is indicative of diagenesis, then δ44Ca at 1212 is also pear to require a fractionation factor significantly less than the core likely impacted by diagenesis. The implication of this observation top-derived −2.01h. The observation that barite δ44Ca is signifi- is that even those records that appear pristine, based on % CaCO3, cantly affected by the variations in pore fluid chemistry over the can be significantly overprinted by authigenic mineral precipitation PETM while Sr/Ba is not can be explained by the relative sensi- associated with ocean acidification. tivity of Ca and Sr to alteration. Calcium concentrations in barite Relatedly, the data presented beg the question of whether or are more than ten times lower than Sr (∼2.4 mmol Ca/mol Ba; not marine barite is a robust archive of seawater δ44Ca over the Averyt and Paytan, 2003), and Ca concentrations in seawater about entire PETM event. Before and after the CIE, the calculated δ44Ca of 100 times higher than Sr concentrations, such that Ca is potentially Paleogene seawater (0.10h ± 0.20h 2SD, n = 9) is statistically in- more easily perturbed in barite than Sr during diagenesis. distinguishable from Holocene seawater values (Fig. 3d). This value Ultimately, identification of two populations of authigenically- is determined assuming a constant isotopic offset between barite influenced barite δ44Ca is important, as the data suggest that the and seawater of −2.01h (Griffith et al., 2008c), after removing the barite Ca isotopic fractionation factor during diagenesis can vary outliers identified in Table 2. At the boundaries of the CIE at 1221, in the sedimentary column. Not only has this not been observed however, there appear to be two populations of barites with δ44Ca previously, but it has the potential to complicate the use of ma- significantly different from the barite before and after the CIE. The rine barite as a paleo-seawater Ca isotope proxy over intervals of first occurs at the base of the CIE, close to the depth at which bulk abrupt and extreme change in ocean chemistry. More work has to carbonate δ44Ca begins to increase strongly. These marine barites be done to confirm these interpretations by using barite from other have substantially lower δ44Ca values (up to −0.6h) than barites sections or time intervals. before and after (and within) the CIE. The second population of barites that differ from the pre CIE values occurs at the beginning 5. Conclusions of the recovery phase of the bulk carbonate record (and end of the CIE), at the same depth horizon in which benthic δ18O increases Previous modeling results of global changes in Ca input to the by ∼1h and bulk carbonate δ44Ca peaks. oceans due to enhanced weathering and carbonate dissolution dur- While these barite samples could be primary and reflective of ing the PETM (Komar and Zeebe, 2011)suggest that these pro- changes in ocean chemistry, or changes in isotopic fractionation cesses will result in little change to the global seawater Ca isotope such as those related to changes in the saturation state of barite ratio (<0.05h); much less than the changes measured in carbon- and thus barite precipitation rates, the rather consistent Sr/Ba in ate and marine barite δ44Ca records from two sites in the Pacific these samples (Table 2; 36.6 ± 5.9 mmol/mol, average ± 2SD, with vastly different dissolution histories. The lack of agreement n = 22) suggest a constant barite saturation state across the PETM among synchronous carbonate records indicate that dissolution, (Paytan et al., 2007). In addition, the average 87Sr/86Sr of these recrystallization and secondary precipitation (i.e., post burial dia- outlier barites (0.70776 ± 0.00003 2SD) is similar to that expected genetic effects) affected the bulk carbonate δ44Ca record at 1221, for contemporaneous seawater (Hodell et al., 2007). However this which was below the CCD during the core interval of the PETM; does not negate the possibility of an authigenic origin for these 1212 was likely affected to a lesser extent. The high values of the outlier barites. This is due to the fact that the 87Sr/86Sr of pore bulk carbonate δ44Ca at 1221 reflect either re-equilibration of ex- fluids from which authigenic barite would have precipitated would isting carbonates with high δ44Ca pore fluids or precipitation of have been isotopically identical to seawater within the upper tens authigenic carbonate from high δ44Ca pore fluids. High δ44Ca pore to hundreds of centimeters of the sedimentary column (e.g., Fantle fluids recorded in carbonates at the top of the clay-rich PETM and DePaolo, 2006). In such a case, there is little isotopic leverage interval are reflecting their proximity to the sediment–seawater in- to produce authigenic barite with substantially different 87Sr/86Sr terface during secondary mineral formation after the cessation of than seawater in the upper part of the sedimentary column, even carbonate burndown. Thus the secondary minerals record a depth in the presence of rapidly dissolving carbonate (e.g., Fantle, 2015). dependent (not time dependent) signal. Above the clay-rich PETM Consequently, we propose that both intervals of outlier barites interval, the return to typical carbonate δ44Ca values reflects the include a proportion of authigenic barite affected by the evolution re-establishment of the flux of carbonate particles derived from the of pore fluid δ44Ca over the course of the PETM. It should be noted surface ocean. This hypothesis is in agreement with bulk carbonate that outside of the extreme dissolution event of the PETM, authi- and benthic foraminiferal oxygen isotopes and trace metal ratios genic barite records seawater chemistry if it precipitates within the (Sr/Ca and Mg/Ca) at 1221. Marine barite δ44Ca is mostly unaf- upper sediment mixed layer (Paytan et al., 1996). Interestingly, the fected by post burial diagenetic effects seen in the bulk carbonate two populations of outlier barite exhibit completely opposite Ca record at 1221 within the core of the event but may be affected at isotopic effects, suggesting that they reflect the evolution of pore the base of the event due to short-lived perturbations to the pore E.M. Griffith et al. / Earth and Planetary Science Letters 419 (2015) 81–92 91

fluid Ca isotopic composition which should be better preserved in Leg 199 Site 1221. In: Proceedings of Ocean Drilling Project. Scientific Results, the barite record which is thought to behave as a closed system vol. 199. below the sediment mixed layer (Paytan et al., 1996). The barite Griffith, E.M., Paytan, A., Caldeira, K., Bullen, T.D., Thomas, E., 2008a. A dynamic marine calcium cycle during the past 28 million years. Science 322, 1671–1674. record, where not affected by changes in pore fluid chemistry at Griffith, E.M., Paytan, A., Kozdon, R., Eisenhauer, A., Ravelo, A.C., 2008b. Influences the sediment-seawater interface, suggests seawater values at the on the fractionation of calcium isotopes in planktonic foraminifera. Earth Planet. PETM were generally similar to present day values. A coherent re- Sci. Lett. 268, 124–136. construction of seawater Ca isotope values over extreme climatic Griffith, E.M., Schauble, E.A., Bullen, T.D., Paytan, A., 2008c. 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Gussone, N., Böhm, F., Eisenhauer, A., Dietzel, M., Heuser, A., Teichert, B.M.A., Reitner, The authors acknowledge samples and data provided by the In- J., Worheide, G., Dullo, W.-C., 2005. Calcium isotope fractionation in calcite and tegrated Ocean Drilling Program. This work was supported in part aragonite. Geochim. Cosmochim. Acta 69, 4485–4494. Gussone, N., Langer, G., Geisen, M., Steel, B.A., Riebesell, U., 2007. Calcium isotope by NSF CAREER Grant OCE-0449732 (A.P.). This is a contribution to fractionation in coccoliths of cultured Calcidiscus leptoporus, Helicosphaera car- EUROclimate project 04 ECLIM FP08 CASIOPEIA (DFG Ei272/29-1). teri, Syracosphaera pulchra and Umbilicosphaera foliosa. Earth Planet. Sci. Lett. 260, The authors thank Joe Street for helpful comments on an earlier 505–515. version of this manuscript, and are grateful to P. 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