We used boron‐based proxies for reconstructing past variations in seawater carbonate chemistry. In culture experiments, both the boron/calcium ratio (B/Ca) and the boron isotopic composition (δ 11 B) of planktic foraminifer shells have been shown to reflect carbonate chemistry parameters [ Allen et al. , 2012 ; Henehan et al. , 2013 ; Sanyal et al. , 1996 , 2001 ]. These relationships are based on the inference from boron isotope studies that the borate ion (B(OH) 4 − ) is the species predominantly incorporated into calcite [ Hemming and Hanson , 1992 ], and both the aqueous abundance of B(OH) 4 − and its isotopic composition increase with seawater pH [ Kakihana et al. , 1977 ]. δ 11 B in planktic foraminifer shells increases with pH [ Henehan et al. , 2013 ; Hönisch et al. , 2007 ; Sanyal et al. , 1996 , 2001 ] and has previously been used to reconstruct seawater pH on Pleistocene [ Henehan et al. , 2013 ; Hönisch and Hemming , 2005 ] and Cenozoic [ Foster et al. , 2012 ; Pearson et al. , 2009 ] timescales. A full understanding of the controls on B/Ca in planktic foraminifera is still emerging, but culturing experiments have demonstrated that B incorporation is well correlated to the ratio of aqueous B(OH) 4 − to total dissolved inorganic carbon (DIC) or bicarbonate (HCO 3 − ) [ Allen et al. , 2011 , 2012 ]. However, quantification of the carbonate chemistry changes documented by B/Ca remains complicated and cannot be used by itself to quantify pH changes [ Allen et al. , 2012 ]. In addition, the [B] and δ 11 B of ancient seawater are poorly constrained [ Lemarchand et al. , 2000 ; Raitzsch and Hönisch , 2013 ; Simon et al. , 2006 ], but the combination of these proxies can be used to reconstruct the timing, duration, and relative magnitude of ocean acidification during the PETM, which can be calculated independently of δ 11 B seawater by assuming pre‐PETM pH.

Current observational constraints on the changes in the carbonate chemistry of the surface ocean across the Paleocene‐Eocene Thermal Maximum (PETM) are largely limited to the abundance and/or morphology of marine calcifiers preserved in sediments [ Bown and Pearson , 2009 ; Gibbs et al. , 2006 , 2010 ; Raffi et al. , 2005 ], which should be sensitive to rapid ocean acidification. Planktic foraminifers show transient population changes and ecosystem disruption over the PETM, including the temporary disappearance of species abundant before the event, and the evolution of “excursion taxa” [ Kelly et al. , 1996 ; Raffi and De Bernardi , 2008 ; Raffi et al. , 2009 ]. In addition, changes in shell morphology may reflect rapid acidification during the CIE followed by elevated saturation during the “overshoot” phase of the recovery [ Kelly et al. , 1996 , 2010 ]. Calcareous nannoplankton experienced significant yet transient changes in species abundance, though without the major extinction that might be expected during rapid acidification [ Bown and Pearson , 2009 ; Gibbs et al. , 2006 ]. Calcareous plankton likely also responded to coeval changes in temperature, salinity, and nutrient concentrations [ Bown and Pearson , 2009 ; Gibbs et al. , 2010 ]. Inferring surface ocean saturation from pelagic fossil assemblages is further complicated by preservational biases, such as CaCO 3 dissolution at the seafloor [ Gibbs et al. , 2010 ]. Therefore, biotic records provide ambiguous evidence of acidification, and we turn to geochemical proxies for independent evidence for ocean acidification.

The Paleocene‐Eocene Thermal Maximum (~56 million years, Ma) is marked by a 3–4‰ decrease in the carbon isotopic composition (δ 13 C) of both organic and inorganic carbon in marine and terrestrial records [ Kennett and Stott , 1991 ; Koch et al. , 1992 ]. Concurrent with this carbon isotope excursion (CIE), marine sediments show a rapid decrease in calcium carbonate (CaCO 3 ) content [ Colosimo et al. , 2006 ; Thomas and Shackleton , 1996 ; Zachos et al. , 2005 ] and a transient global warming of 4–8°C as indicated by various paleothermometers [ Dunkley‐Jones et al. , 2013 ; McInerney and Wing , 2011 ; Zachos et al. , 2003 ]. This combined evidence points to a massive release of 13 C‐depleted carbon into the ocean‐atmosphere system [ Dickens et al. , 1997 ; Pagani et al. , 2006 ]; thus, this interval presents an opportunity to examine the response of ocean chemistry to a geologically rapid increase in the atmospheric concentration of carbon dioxide (CO 2 ) [ Hönisch et al. , 2012 ]. Ocean carbon cycle modeling [ Panchuk et al. , 2008 ; Zeebe et al. , 2009 ], using sediment CaCO 3 content to constrain the changes in the calcite compensation depth (CCD), estimates the mass of carbon released to between 3000 and 9000 Pg C, with marine methane clathrate, terrestrial and/or marine organic carbon, or some combination thereof as the probable source of carbon [ Dickens et al. , 1997 ; Pagani et al. , 2006 ]. In all cases, a significant decrease in surface seawater pH is predicted, the extent of which scales with the rate and magnitude of carbon release [ Hönisch et al. , 2012 ; Ridgwell and Schmidt , 2010 ].

Age models were constructed by correlating the fine‐fraction δ 13 C record of Sites 1209 and 1263 with the bulk δ 13 C record for ODP Site 690, which has an orbitally tuned age model [ Röhl et al. , 2007 ] (Figure S3). For Site 865, the benthic foraminiferal δ 13 C record was correlated with that at Site 690 [ Zachos et al. , 2001 ] to generate an age model. Our age models lead to a slightly shorter duration for the CIE than other estimates [ Farley and Eltgroth , 2003 ; Murphy et al. , 2010 ]. The resulting age model for Site 1209, similar to that for other pelagic PETM sites, implies greatly reduced sedimentation rates immediately below and within the core of the CIE, resulting from chemical erosion of uppermost Paleocene CaCO 3 sediments and decreased production and/or preservation of CaCO 3 during the CIE, followed by enhanced sedimentation rates during and after the recovery interval [ Farley and Eltgroth , 2003 ], representing a carbonate saturation overshoot phase [ Kelly et al. , 2010 ; Zachos et al. , 2005 ].

Applying simple assumptions about seawater chemistry (described below and in the supporting information ), we estimate not the absolute pH values, but the pH excursion (ΔpH), from δ 11 B values across the P‐E boundary. As a second constraint, we compare measured B/Ca records with model‐derived estimates of B/Ca based on the observed sensitivity of modern foraminifera. These two techniques complement each other in that the more complete understanding of the δ 11 B proxy allows a more quantitative interpretation, while the smaller sample size requirement of B/Ca analyses allows generation of a much higher resolution record, so that timing and structure can be resolved more precisely.

We measured δ 11 B using standard negative thermal ionization mass spectrometry methods [ Hemming and Hanson , 1994 ] and complemented the δ 11 B data at Site 1209 by higher‐resolution records of B/Ca and Mg/Ca measured by inductively coupled plasma–mass spectrometry, and δ 13 C and δ 18 O measured by standard dual‐inlet techniques (see supporting information for detailed description of analytical methods). In addition to M. velascoensis , trace elements and stable isotopes were analyzed in shells of A. soldadoensis and Subbotina .

During Ocean Drilling Program (ODP) Leg 198, three holes were drilled at Site 1209 (Shatsky Rise, N. Pacific 32°39.1081′N, 158°30.3564′E) at a water depth of 2387 m [ Bralower et al. , 2002 ], equivalent to a paleodepth during the PETM of ~ 1900 m [ Takeda and Kaiho , 2007 ]. The Paleocene‐Eocene boundary interval, recovered at 109 m composite depth, is composed primarily of a carbonate‐rich nannofossil ooze. Investigations of this interval documented a prominent carbon isotope excursion (CIE), warming, carbonate dissolution, and the benthic foraminiferal extinction horizon [ Colosimo et al. , 2006 ; Takeda and Kaiho , 2007 ; Zachos et al. , 2005 ]. Site 1209 was chosen for this study because of its location within a subtropical gyre, away from regions of upwelling. Moreover, its high %CaCO 3 and foraminiferal abundance facilitated collection of a high‐resolution B/Ca record and large samples for measuring boron isotopes. Sediment samples, collected at 1–3 cm resolution across a 2 m interval spanning the CIE, were washed and sieved, and specimens of the mixed‐layer‐dwelling planktic species Morozovella velascoensis and Acarinina soldadoensis were picked from the 250–300 and 300–425 µm size fraction. On the basis of shell size‐δ 13 C relations, these species likely harbored photosynthetic algal symbionts and were thus restricted to the photic zone of the surface ocean [ D'Hondt et al. , 1994 ]. Additionally, specimens of the smooth‐walled, thermocline‐dwelling genus Subbotina were picked from the 250–300 µm size fraction. Isotopic depth ranking suggests that this taxon was nonsymbiotic and occupied the thermocline [ Berggren and Norris , 1997 ]. Boron isotope analyses at Site 1209 were restricted to M. velascoensis and complemented by low‐resolution δ 11 B analyses of the same taxon from Sites 1263 (Walvis Ridge, Southeast Atlantic, 28°31.98′S, 02°46.77′E, 2717 m depth; paleodepth ~1500 m; [ Zachos et al. , 2004 ]) and 865 (Allison Guyot, Equatorial Pacific, 18°26.41′N, 179°22.24′W, 1518 m depth; paleodepth ~ 1400 m; [ Bralower et al. , 1995 ]) to evaluate whether the Site 1209 record is representative of a global signal or compromised by local or preservational effects. At each of these sites, previous work has identified the Paleocene‐Eocene (P‐E) boundary and accompanying CIE [ Colosimo et al. , 2006 ; Kelly et al. , 1996 ; Zachos et al. , 2005 ], which is used to correlate between sites.

B/Ca data generally parallel the Site 1209 boron isotope excursion and recovery. The upper Paleocene baseline M. velascoensis and A. soldadoensis B/Ca values of ~70 µmol/mol fall toward the low end of those of modern symbiont‐bearing foraminifers [ Allen et al. , 2012 ], consistent with lower than modern seawater pH and [B] [ Lemarchand et al. , 2000 ]. Values in both species show a rapid drop from an average of 70 to 45 µmol/mol across the onset of the CIE (Figure 1 ), remain low for 15 cm (at least ~70 kyr) above the boundary, then recover gradually, in step with δ 11 B to pre‐CIE levels. B/Ca data for the thermocline‐dwelling Subbotina follow the same trend as the data for mixed‐layer‐dwelling species but are offset toward lower values over the entire record, likely reflecting the lower environmental pH of their habitat. Subbotina record a smaller B/Ca excursion (from 45 to 30 µmol/mol) than mixed‐layer dwellers, consistent with model predictions that the magnitude of acidification decreases with depth in the water column [ Ridgwell and Schmidt , 2010 ].

Site 1209 δ 11 B values decrease from an average of ~15.5‰ below the CIE to an average of ~14.7‰ within the body of the CIE (Figure 1 ). This decrease is rapid, occurring within 5 cm, and values remain low for at least 15 cm (at least 70 kyr) before recovering more gradually to preevent levels, more or less in parallel with the δ 13 C recovery. δ 11 B values at Sites 1263 and 865 also show a significant decrease within the CIE (Figure 4 ), supporting that the Site 1209 record reflects global environmental change. However, the excursion is covered by only one or two data points in each of these complementary records, and we therefore restrict calculation of the magnitude of acidification to Site 1209 data.

Data from Core 198‐1209B‐22H plotted against distance in the core in centimeters relative to the CIE onset, at 134 cm in Section 1209B‐22H‐1W. Shown are a core photograph, foraminiferal δB, B/Ca, Mg/Ca, and δC []. The grey shaded area indicates the baseline conditions before the onset of the PETM; brown shading indicates the body of the CIE; yellow shading represents the CIE recovery interval; and the unshaded interval is considered postevent. Error bars on Mg/Ca and B/Ca are 2 standard deviations of repeat measurements of an in‐house carbonate standard: 7% on B/Ca and 4% on Mg/Ca. Error bars on δB are 2 SE of repeat sample analyses (> 3), or 2 SE of repeat analyses of an in‐house vaterite standard given the same, whichever is larger.

The δ 13 C and Mg/Ca values of the mixed‐layer‐dwelling foraminifera accurately replicate previously published records for Site 1209 [ Zachos et al. , 2003 ] and extend them by 200 kyr into the early Eocene (Figure 1 ). The 50% rise in Mg/Ca (~2 mmol/mol) at the onset of the CIE is consistent with a ~5°C increase in sea surface temperature (SST) [ Zachos et al. , 2003 ], whereas the comparably smaller decrease in oxygen isotopes presented in Zachos et al. [ 2003 ] is consistent with a rise in local δ 18 O sw and 1.5 ppt increase in salinity ( S ) [ Zachos et al. , 2003 ]. These excursion SST and S estimates are used to constrain the boric acid dissociation constant and boron isotopic fractionation factor required for pH calculations from δ 11 B across the CIE. Before and after the event, a baseline of 30°C and S = 37 are assumed, similar to other estimates of late Paleocene low‐latitude SST and S [ Kozdon et al. , 2011 ].

4 Discussion

4.1 Evidence for Acidification The similarity and timing of the δ11B and B/Ca anomalies suggest that both proxies are documenting the surface acidification caused by the geologically rapid (within 1 to 20 kyr) release of thousands of Pg C [e.g., Zeebe and Zachos, 2012]. However, given the drastic environmental changes documented during the PETM (including warming, salinity changes, and dissolution), we must consider whether there may be factors other than pH that might influence the B proxies. In culture experiments documenting the sensitivity of planktic foraminiferal δ11B and B/Ca to pH, no effect was observed with temperature and for B/Ca only a minor effect with salinity (~2.6 µmol/mol/psu) [Allen et al., 2012]. Moreover, culture studies would predict an increase in B/Ca at the PETM onset in response to the local salinity increase [Zachos et al., 2003], the opposite of what is observed. As for preservation artifacts, Yu et al. [2007] found no correlation between B/Ca in planktic foraminifers and bottom water carbonate saturation state (Ω) in core top samples, concluding that seafloor dissolution likely does not alter B/Ca. Coadic et al. [2013], however, found a correlation between depth and planktic B/Ca along a depth transect, suggesting that dissolution may be an issue. The observed dissolution effect of 10–15 µmol/mol between 2500 and 5000 m water depth, however, is small compared to the much larger signal in our records, and the paleodepth for Site 1209 was only ~ 1900 m[Takeda and Kaiho, 2007]. Similarly, δ11B records have been observed to be affected by dissolution [Hönisch and Hemming, 2004], and the change is in the same direction as that observed in our PETM records. Importantly, these dissolution artifacts are associated with a concomitant decrease in Mg/Ca [Coadic et al., 2013; Hönisch and Hemming, 2004], whereas our record shows an increase in Mg/Ca, consistent with the global average temperature increase [Dunkley‐Jones et al., 2013]. Furthermore, the most extreme phase of dissolution at Site 1209, as reflected in %CaCO 3 and shell fragmentation [Colosimo et al., 2006], occurred over ~5 cm during the onset of the event, and dissolution recovered below the δ11B and B/Ca recoveries. Because δ11B and B/Ca remain low after recovery of preservation, we do not consider dissolution a likely explanation for the observed trends. Furthermore, the preservation of species‐specific offsets in B/Ca suggests that recrystallization/secondary calcification has not altered the primary signal of the shells. Detailed scanning electron microscope (SEM) photographs of CIE interval planktic foraminifera [Colosimo et al., 2006] provide evidence for some diagenetic carbonate overgrowth and pore infilling, but an overgrowth effect cannot significantly bias our records, given that most of the B in pelagic sediments resides in calcite, pore water contributes comparably little B, and the B/Ca of inorganic calcite is similar to that of biogenic calcite [Sanyal et al., 2000]. A B content similar to that of foraminiferal calcite is most likely in overgrowths, but even if overgrowths contained no B, it would require anomalous carbonate overgrowths of ~30–40% (restricted to only the CIE interval) to explain our B/Ca records, an amount not supported by SEM investigations at Site 1209 [Colosimo et al., 2006]. And, if addition of B‐free carbonate overgrowth were to account for the B/Ca decrease, then it could not affect the decrease documented in δ11B. Perhaps the strongest evidence against diagenetic alteration of the δ11B records is the close agreement of resulting pH recorded at three different sites with varied lithologies and diagenetic regimes (see section 4.3). Additionally, photosynthetic activity of symbiotic algae elevates the pH of the mixed‐layer dwellers' calcifying microenvironment [Hönisch et al., 2003; Jorgensen et al., 1985; Rink et al., 1998], but severe warming could have led to the loss of symbionts and thus lowered both δ11B and B/Ca. Symbiont loss has been inferred from a reduced δ13C‐shell size relationship during the Middle Eocene Climatic Optimum [Edgar et al., 2013], and culturing studies have documented an effect of symbiont activity on δ11B [Hönisch et al., 2003]. However, the asymbiotic, thermocline‐dwelling Subbotina also record a B/Ca excursion, which means that potential loss of symbionts cannot explain the entire B/Ca signal. Additionally, we have measured δ11B in two size fractions (250–300 µm and 300–425 µm) below and during the CIE, as well as within the recovery interval, and there is little or no difference between these size fractions (Table S2). Since symbiont activity affects δ11B‐size relationships [Hönisch and Hemming, 2004], we conclude that the effects of symbiotic activity or its possible disappearance during the CIE on our δ11B record is minimal. The mixed‐layer species might also have migrated to deeper, cooler waters in response to warming, though this seems unlikely because the +5°C temperature increase recorded by Mg/Ca in the same shells is comparable to the global average [Dunkley‐Jones et al., 2013; Sluijs et al., 2006; Zachos et al., 2003; Zachos et al., 2006]. We conclude that symbiont loss and/or depth migration cannot explain the entire trends in our records but may have amplified the signal in B/Ca and δ11B in Morozovella and Acarinina, so that the estimated pH changes should be viewed as maxima if these records are considered complete. However, it is also possible that a short phase of highly acidified conditions at the onset of the CIE is not represented in our B records due to dissolution and chemical erosion. Compared to other pelagic PETM sections, Site 1209 features minimal dissolution, and just a millimeter‐scale clay‐rich seam [Colosimo et al., 2006] representing the CIE onset. The time represented by this seam is unknown, but local recovery of carbonate deposition from this initial pulse of emissions would be relatively fast (<10 kyr), so most of the main portion of CIE is represented as constrained by the correlation between δ13C records (Figure S3). Finally, changes in the Ca budget related to the global carbonate dissolution pulse [Komar and Zeebe, 2011] would produce changes in B/Ca and Mg/Ca much smaller than the analytical uncertainty in the techniques employed in this study. Moreover, the B/Ca signals are much larger than the analytical uncertainty, in contrast to those in studies of much smaller pH variations of the late Cenozoic [Allen and Hönisch, 2012]. In sum, the potential contributions of complicating factors are minor compared to the magnitude of the anomalies in both the δ11B and B/Ca records. We therefore interpret the decreases in both proxies at the onset of the PETM as ocean acidification signals.

4.2 Quantification of ΔpH From δ11B To constrain the magnitude of the acidification documented by the B proxies, we calculate a range of ΔpH values from the M. velascoensis δ11B record across the CIE onset (pre‐PETM values versus CIE values). Computation of absolute pH from foraminiferal δ11B requires knowledge of δ11B seawater , which is unknown for the Paleogene, and of species‐specific vital effects, which are difficult to calibrate in these now extinct species. However, δ11B seawater can be assumed constant over the ~300 kyr time window considered here, because the oceanic residence time of B is >10 Myr [Lemarchand et al., 2000]. Furthermore, estimates of ΔpH independent of those parameters can be made by assuming a pre‐PETM baseline pH (see supporting information for full calculations). Using the magnitude of the δ11B‐excursion, we calculate ΔpH for a range of assumed initial pH values between 7.5 and 8.2 (Figure 2). Due to reduced sensitivity in the fractionation of B isotopes at lower pH [Hemming and Hanson, 1992], the resulting ΔpH is larger at lower assumed initial pH and smaller at high (near‐modern) assumed initial pH. Within the range of estimates of pre‐PETM pH given by model simulations [Panchuk et al., 2008; Ridgwell and Schmidt, 2010; Zeebe et al., 2009], the 0.8‰ decrease in δ11B at the onset of the event is consistent with a ΔpH of −0.27 units, assuming a pre‐PETM pH of 7.80 (total scale, consistent with 750 ppm atmospheric CO 2 ) [Panchuk et al., 2008], to −0.34 units, assuming a pre‐PETM pH of 7.67 (total scale, consistent with 1000 ppm CO 2 )(Figure 3). Figure 2 Open in figure viewer PowerPoint Calculated ΔpH as a function of assumed initial (pre‐PETM) pH. Two sets of calculations are shown: (1) (black line and grey shading) δ11B data from the Paleocene (n = 7) and the CIE interval (n = 6) are treated as populations, and calculations are performed on the variance‐weighted mean of those populations with associated uncertainty propagated through the ΔpH calculation (pre‐CIE δ11B = 15.51‰ ± 0.12, CIE δ11B = 14.71‰ ± 0.11). (2) (red) Only the two δ11B data points spanning the P‐E boundary are considered (15.35 ± 0.30 and 14.76 ± 0.40 for the latest Paleocene and earliest Eocene, respectively), and error is analytical uncertainty of δ11B analysis propagated through the ΔpH calculation. Vertical lines represent the preevent pH of two model simulations of the PETM. Figure 3 Open in figure viewer PowerPoint Zeebe et al. [ 2009 Panchuk et al. [ 2008 Zeebe et al.'s [ 2009 Calculated pH versus time after the onset of the PETM. Two scenarios are plotted, one assuming an initial pH of 7.67 (after], red) and one assuming initial pH of 7.80 (after], blue). Error bars reflect 2 SE of individual boron isotope measurements as displayed in Figure 1 . For comparison, the grey line shows the evolution of the surface pacific pH in.'s [] simulation. Propagated uncertainties on ΔpH are calculated for two cases (Figure 2): (1) treating the Paleocene and CIE interval δ11B as coherent, normally distributed populations and using a variance‐weighted mean with associated propagated uncertainties, and (2) considering only the two δ11B data points which span the P‐E boundary and using analytical uncertainty as a source of error (see supporting information for description of statistics and error propagation). Using the weighted mean allows smaller uncertainty (± 0.08 units in the pH = 7.8 pre‐PETM case and ± 0.11 units in the pH = 7.67 pre‐PETM case) than the second technique (± 0.13 units in the pH = 7.8 pre‐PETM case and ± 0.18 units in the pH = 7.67 pre‐PETM case). The first technique, however, rests on the assumption that the two populations of δ11B are normally distributed around a coherent mean, which may not be the case, especially within the CIE when pH likely varied over time (Figure 5c). The second technique avoids this assumption, but the two boundary‐spanning data points are separated by 16 kyr in our age model, so that this comparison does not necessarily capture the full magnitude of the ΔpH across the boundary. Applying the same vital effect offset and δ11B seawater values at Sites 1209, 865, and 1263, pre‐PETM pH estimates from δ11B are within ±0.05 pH units (Figure 4), suggesting that diagenesis is not a major concern, because it would have affected the three sites differently due to different burial depths, carbonate content, bottom water, and pore water chemistry. The P‐E interval at Site 865 (paleodepth ~1400 m) is a winnowed foraminiferal sand (nearly 100% CaCO 3 ), bearing planktic foraminifers with significant overgrowth [Kelly et al., 1996; Kozdon et al., 2013]. Site 1209 (paleodepth ~1900 m) foraminifers are lightly overgrown [Colosimo et al., 2006], while Site 1263 (paleodepth ~1500 m) features clay‐rich sediments (~85% CaCO 3 ) with excellent foraminiferal preservation [Kelly et al., 2010] but a larger clay layer indicating more intense dissolution [Zachos et al., 2005]. Depending on the initial pH, our Site 1209 δ11B data allow for a wide range of ΔpH values (i.e., larger if a lower initial pH is assumed and smaller if a higher initial pH is assumed; Figure 2), though an initial pH far outside the range in Figure 2, and assuming realistic values for other carbonate chemistry parameters, would be irreconcilable with independent estimates for Paleocene pCO 2 [Beerling and Royer, 2011]. Figure 4 Open in figure viewer PowerPoint 11B values and two scenarios for pH (total scale) from M. velascoensis at all three sites plotted against time. Age models for Sites 1263 and 1209 were produced by correlating the bulk/fine‐fraction δ13C excursion to the δ13C excursion at Site 690 in the age model of [Röhl et al., 2007 13C record [Zachos et al., 2001 11B are 2 SE of repeat analyses (n > 3) or 2 SE of repeat analyses of an in‐house vaterite standard given the same n, whichever is larger. The pH in the middle and lower panels were calculated by assuming the initial pH for the Site 1209 record of 7.8 and 7.67 (total scale), respectively, and then applying the same δ11B‐pH relationship to all sites. Error bars reflect the uncertainty reported for δ11B analyses. The lower error bar on one point in the CIE from Site 865 is incalculable because the lower error limit on the δ11B of that point is below the minimum δ11B in the δ11B‐pH relationship. The δB values and two scenarios for pH (total scale) fromat all three sites plotted against time. Age models for Sites 1263 and 1209 were produced by correlating the bulk/fine‐fraction δC excursion to the δC excursion at Site 690 in the age model of []; for Site 865 the benthic foraminiferal δC record [] was correlated to that at Site 690, using the same age model. Error bars on δB are 2 SE of repeat analyses (> 3) or 2 SE of repeat analyses of an in‐house vaterite standard given the same, whichever is larger. The pH in the middle and lower panels were calculated by assuming the initial pH for the Site 1209 record of 7.8 and 7.67 (total scale), respectively, and then applying the same δB‐pH relationship to all sites. Error bars reflect the uncertainty reported for δB analyses. The lower error bar on one point in the CIE from Site 865 is incalculable because the lower error limit on the δB of that point is below the minimum δB in the δB‐pH relationship.

4.3 Significance of the B/Ca Excursion The B/Ca data allow us to independently evaluate the magnitude of acidification in surface waters at the onset of the PETM. Based on theoretical and laboratory studies, a pH decrease as estimated from the anomaly in δ11B should also lower the B/Ca of foraminifera [Allen et al., 2012], as indeed recorded at Site 1209. In order to estimate changes in pH from B/Ca, however, we must apply calibrations for modern foraminifer species, which, in contrast to δ11B versus pH relationships, vary in sensitivity between species. We assume that extinct species shared a similar range of sensitivities as living species, but the approach is complicated by the dual response of planktic B/Ca to varying [B(OH) 4 −] and DIC, where carbonate and borate ion appear to compete for lattice sites during calcification[Allen et al., 2012]. Although the concentrations of aqueous borate and carbonate ions are coupled during a rapid ocean acidification event and fall with decreasing pH [Zeebe et al., 2009], the carbon pulse at the onset of the PETM led to an increase in DIC, which may have amplified the B/Ca decrease through decreasing [B(OH) 4 −]/DIC (Figure 5c). In the later phase of the event (>10 kyr after the CIE onset), terrestrial weathering increased the flux of alkalinity and DIC to the ocean which decoupled pH and carbonate ion concentration [Hönisch et al., 2012], leading to slower recovery of [B(OH) 4 −]/DIC (and thus B/Ca) than surface ocean carbonate saturation (Figure 5c). As such, a direct, quantitative estimate of carbonate chemistry from raw B/Ca data carries uncertainty at present. Nevertheless, the direction, magnitude, and timing of the B/Ca excursions are qualitatively consistent with rapid and sustained reductions in thermocline and surface ocean pH, lasting as long as the peak carbon isotope excursion (>70 kyr). Figure 5 Open in figure viewer PowerPoint Röhl et al., 2007 Murphy et al., 2010 Allen et al., 2011 2012 Zeebe et al., 2009 13C, Ω calcite , and B(OH) 4 −/DIC (where 0 represents initial Paleocene values and 1 represents PETM minima) demonstrating the different timescales of recovery between model parameters. Horizontal grey line represents preevent conditions for all parameters. Model‐data comparison of the PETM B/Ca anomaly. Symbols with error bars are B/Ca measured in mixed‐layer‐dwelling foraminifers from Core 198‐1209B‐22H, plotted against a timescale derived from correlating the fine‐fraction CIE at that site to the orbitally tuned age model of ODP Site 690 [] (see supporting information Figure S3). Note that this age model produces a conservative estimate of the CIE duration (~100 kyr) compared to other age models which suggest a duration up to ~200 kyr [], so time after the CIE onset of the B/Ca data should be taken as a minimum. Colored lines are model‐derived predictions generated by applying the “absolute” and “relative” cultured B/Ca calibrations of three modern foraminifer species [] to the mixed‐layer water carbonate chemistry conditions calculated for the low‐latitude surface Pacific by the best fit PETM simulation of the LOSCAR carbon cycle model [] (see supporting information for discussion of timescale and calibrations). (a) Measured B/Ca ratios (µmol/mol) compared with model‐derived curves using the absolute calibration. (b) The B/Ca anomaly as percent change relative to the Paleocene baseline along with model‐derived curves using the relative calibration. (c) Normalized excursions in pH, surface Pacific δC, Ω, and B(OH)/DIC (where 0 represents initial Paleocene values and 1 represents PETM minima) demonstrating the different timescales of recovery between model parameters. Horizontal grey line represents preevent conditions for all parameters.