Larger CO2 source at the equatorial Pacific during the last deglaciation

While biogeochemical and physical processes in the Southern Ocean are thought to be central to atmospheric CO2 rise during the last deglaciation, the role of the equatorial Pacific, where the largest CO2 source exists at present, remains largely unconstrained. Here we present seawater pH and pCO2 variations from fossil Porites corals in the mid equatorial Pacific offshore Tahiti based on a newly calibrated boron isotope paleo-pH proxy. Our new data, together with recalibrated existing data, indicate that a significant pCO2 increase (pH decrease), accompanied by anomalously large marine 14C reservoir ages, occurred following not only the Younger Dryas, but also Heinrich Stadial 1. These findings indicate an expanded zone of equatorial upwelling and resultant CO2 emission, which may be derived from higher subsurface dissolved inorganic carbon concentration.

U nderstanding the past condition of the surface ocean carbonate system and air-sea CO 2 exchange is crucial to projecting future changes in the carbon cycle under ongoing anthropogenic global warming and ocean acidification. Atmospheric CO 2 concentration increased by as much as 80 matm during the last deglaciation, with ,50 matm released during Heinrich Stadial 1 (HS1, from 17.5 to 14.6 kyr), followed by an additional ,30 matm during the Younger Dryas (YD, from 12.9 to 11.7 kyr) (ref. 1). While the Southern Ocean is generally considered to be central to the deglacial CO 2 rise 2-9 , the contribution from other oceanic regions remains relatively uninvestigated [10][11][12][13] . Information on the partial pressure of CO 2 (pCO 2 ) is needed to directly constrain past air-sea CO 2 exchange, and this can be reconstructed from boron isotopes (d 11 B), a marine carbonate pH proxy 14,15 . Regions where surface seawater CO 2 is out of equilibrium with the atmosphere are ideal for such studies, and the equatorial Pacific is particularly well suited because it represents the largest global CO 2 source at present (e.g. ref. 16). Corals broadly distributed in tropical to subtropical areas constitute excellent high-resolution geochemical archives for paleo-CO 2 studies because they may be precisely radiogenically dated (U-series) 3,17 , unlike foraminifers that are affected by uncertainty in radiocarbon reservoir age (R).
Two previous studies have attempted to constrain equatorial Pacific CO 2 changes. Palmer and Pearson 10 showed increased CO 2 emission during the last deglaciation in the western equatorial Pacific (WEP) from d 11 B measurements on the planktonic foraminifer (Globigerinoides sacculifer) in a sediment core recovered from offshore Papua New Guinea (ERDC-92, Fig. 1a). Further east, Douville et al. 11 performed d 11 B on fossil corals from the Marquesas (9.5uS 139.4uW, Fig. 1a) and also demonstrated increased CO 2 release at the end of the YD. However Douville et al. 11 did not observe a significant CO 2 release during HS1, complicating interpretation of the equatorial contribution to deglacial atmospheric CO 2 rise. Integrated Ocean Drilling Program Expedition 310 (IODP Exp. 310) 18 drilled the outer reef slope at Tahiti (17.6uS 149.5uW, Fig. 1a) recovering fossil corals from an open ocean environment spanning HS1, which enable us to assess the issue.
Results d 11 B-pH calibration. This study establishes a new empirical d 11 B-pH calibration utilizing, for the first time, anthropogenic ocean acidification. Empirical calibration is needed to overcome the observed offsets from a theoretical d 11 B-pH curve in culture experiments for zooxanthellate corals 14,15 . (pH is reported using the total hydrogen scale, hereafter pH for simplicity). There are two primary approaches to overcoming the reported offsets. One is an empirical approach 14 that assumes constant offsets in measured and theoretical d 11 B (''offset a;'' see Methods), while the other is an observational approach 15 that considers potential pH modification by calcifiers.
The potential for pH-modification is of great concern for d 11 Bbased reconstruction of pH due to the implications for atmospheric pCO 2 calculation. Such a phenomenon is consistent with indirect pH measurements of internal calcification fluid using pH sensitive dye that suggests a higher pH than ambient seawater, creating better conditions for calcification 19 . A 'DpH' concept in the d 11 B-pH calibration that reflects pH differences in seawater (pH SW ) and internal calcification fluid has been proposed 15 , however usage of this proposed relationship to calibrate d 11 B of modern Porites spp. from Tahiti and Marquesas resulted in unrealistically high values (e.g. ,8.34 in AD 1991), well above reported estimates (e.g. refs. 20,21). Therefore the present study employs the empirical equation 14 here (see Methods).
Ocean acidification was estimated from a combination of in situ fCO 2 values in the surface ocean, atmospheric CO 2 concentration directly measured at the Mauna Loa observatory in Hawaii since AD 1960 (ref. 22), and CO 2 concentrations within bubbles trapped in an Antarctic ice core 23 (Figs. 2, S1-S4; Supplementary Methods). These data were then fit to the previously reported d 11 B measurements of Porites spp. (refs. 11,24), which are for the years AD 1991, 1950, and 1700 ( Fig. 2, see Methods and Supplementary Methods for details of the d 11 B-pH calibration and pH estimation since the Industrial Revolution).
pH and pCO 2 reconstruction. Using our revised calibration, we reconstructed pH from our new d 11 B measurements on Tahitian corals, as well as from previously reported data 11 from both the Marquesas and Tahiti, and the overall result is consistent with the WEP foraminifer d 11 B variations 10 (Fig 3a and b). The oldest coral sample, dated to 20.7 ka BP during the last glacial maximum (LGM), exhibits a relatively high pH (8.26). From 15.5 to 9.0 ka BP, pH is generally constant within uncertainty (8.15-8.22) and consistent  11 . The low pH following HS1 had been previously undetected at this location. Calculation of pCO 2 (see Methods) reveals deglacial values significantly above those of the atmosphere (Figs. 3c and 4a). Conversely, DpCO 2 during last glacial and the early Holocene was nearly zero, suggesting air-sea CO 2 equilibrium.
Results from a different portion of the same 14.99 ka BP coral sample deviate by as much as 1.4%, which corresponds to 0.11 in pH and 100 matm in pCO 2 (310-M0024A-11R-1W_77-90 and 310-M0024A-11R-1W_60-75, Table S1). Considering the average ,4 year temporal resolution of each sample, these excursions occurred abruptly and persisted for several years, which differs from modern observations that show no clear interannual or decadal variability (Supplementary Methods, Fig. S1). This enhanced variability, which is also observed in Sr/Ca derived SST results from another Porites colony recovered from IODP Exp. 310 (ref. 25), may relate to Tahiti's location at the rim of equatorial upwelling cell (Fig. 1a). Taken together, pCO 2 (pH) records indicate that the equatorial Pacific became a larger CO 2 source during the last deglaciation with excursions at the end of HS1 and the YD.  Table S2). Larger and more variable values of R are evident in Tahiti during HS1 and the YD, and enhanced R variability is also seen in the Marquesas (Figs. 4 and S5). Paterne et al. 26 sub-sampled different parts in the same fossil coral skeleton and analyzed both 14 C and U/Th. They observed no difference in U/Th dates, but a much larger difference in 14 C.  (Table S1). Estimated pH from the AD 1700 Marquesas coral was scaled by 0.04 to correct for offset from Tahitian coral values. Error bars are 2s.  Possibilities of either a diagenetic alteration or a change in R were suggested. The latter is more probable because a large variation in R is also suggested from Vanuatu coral at 11.7-12.4 ka (,400 years; during the YD, ref. 27).
Though reported R diff (difference between calculated R and modern R; see Methods) data around the upwelling zone during the LGM are sparse, calculations with the new Lake Suigetsu datasets 28 suggests no substantial change in R (see Methods; Figs. 4 and S5). This implies that the CO 2 exchange rate in the surface equatorial Pacific during the last glacial was almost the same as present, which supports the above-mentioned observation that DpCO 2 is essentially equivalent to zero and indicates that anomalous R values are limited to the last deglaciation.
Discussion pCO 2 variability in subtropical oligotrophic water can be explained by mixing of water masses that exhibit distinctly different dissolved inorganic carbon (DIC) concentrations. A southward migration of the intertropical convergence zone (ITCZ) that partly controls thermocline depth is hypothesized during Heinrich Events including HS1 and the YD (e.g. ref. 29). At present, the ITCZ does not seem to affect surface pCO 2 variability (Fig. 1a), and if it is displaced southward, the locus of equatorial upwelling remains at the equator due to the influence of inter-hemispheric asymmetry of Coriolis force (e.g. refs. 30,31). Enhanced upwelling (shallower thermocline, La Niña-like conditions) or increased subsurface DIC concentration are more likely to drive pCO 2 variability based on sedimentary evidence from the equatorial Pacific for higher nutrient content, e.g., enhanced biogenic opal export production and lower stable carbon isotopes (d 13 C) (TT013-PC72, ODP Site 1240 and TR163-19) 32-35 (Fig. 1a). Semi-conservative radiogenic neodymium isotopes (eNd) from sediment cores at the eastern equatorial Pacific (EEP) (ODP Site 1240) and off Baja California (MV99-MC99-GC31/PC08) indicate stronger subsurface water transport from the south 33,36 (Fig. 1). Covariation of geochemical properties between the Southern Ocean and the equatorial Pacific suggest a subsurface connection during the last deglaciation (e.g. refs. 32-38). Thus, pCO 2 variability may be explained by an increase in DIC in the upwelled, subsurface water masses as opposed to physical processes alone.
Water mass subduction along the subantarctic front, mainly off Chile 39 , forms Subantarctic Mode Water (SAMW) and Antarctic Intermediate Water (AAIW) that upwells at the equatorial Pacific via the Equatorial Undercurrent (EUC) (Fig. 1). SAMW and AAIW are characterized by higher/lower concentrations of oxygen/silicic acid (Fig. 1b,c). It is suggested that the abyssal DIC reservoir around the Southern Ocean increased during the last glacial period [6][7][8]40 , which would have contributed to lower atmospheric pCO 2 . Carbon dioxide was released to the surface through deep ocean ventilation during HS1 and the YD (refs. 2-9), however export production was insufficient to fully compensate the increased carbon flux 41 . This is consistent with residual radiocarbon content (D 14 C) of intermediate water at the EEP (V21-30) 42 and off Baja California (MV99-MC99-GC31/PC08) 43 that indicates anomalously older water was incorporated into SAMW/AAIW (Fig. 1), as well as with depleted d 13 C of surface and lower thermocline dwelling foraminifers from sediment cores at both equatorial (TT013-PC72, ODP Site 1240 and TR163-19) and South Pacific sites [32][33][34][35]37,38 . Moreover, enhanced export production of biogenic opal suggest more silicic acid was transported via the EUC to thermocline water at the equatorial Pacific (V19-30 and TT013-PC72) without being consumed completely within the Southern Ocean 4,35,41 . Stronger Ekman transport in association with sea ice retreat and a poleward shift of southern westeries is suggested to be a driver 4,5 .
A similarity between R and pCO 2 variability during the last deglaciation supports an interpretation that older DIC was incorporated to subtropical surface water through mixing with SAMW/AAIW, though, contrary evidence comes from the current formation sites off Chile 44 and New Zealand 6 . However, a key sediment record off Chile was recently reevaluated, and the new interpretations indicate stronger upwelling and subsequent larger R in surface water in the Southern Ocean 9 , which agrees well with our interpretation. Yet, further work is still needed to fully understand both the physical and biogeochemical dynamics in the Southern Ocean and the equatorial Pacific 2 .
Positive DpCO 2 indicates CO 2 flux from the ocean to the atmosphere. Previous studies 10,11 indicated that the equatorial Pacific contributed to deglacial CO 2 rise, however the timing of anomalously higher pCO 2 events recorded in radiogenically dated fossil corals do not systematically correspond to those of atmospheric CO 2 rise recovered from Antarctic ice core on the GICC05 timescale 1 (Figs. 3 and 4). Moreover our new calibration reveals a modest CO 2 emission continued through the Bølling/Allerød/Antarctic Cold Reversal when no atmospheric CO 2 increase is observed (Figs. 3 and 4). Though we demonstrate that the equatorial Pacific became a larger CO 2 source during the last deglaciation, it is too early to conclude its exact contribution to atmospheric CO 2 rise. The Southern Ocean is suggested to be central in CO 2 degassing 4-9 and the contribution of the terrestrial biosphere should be further evaluated 45 . More evidence spanning the YD and the early part of HS1, in particular the sharp rise in atmospheric CO 2 and the sudden drop of d 13 C of CO 2 (refs. 1,46), as well as more spatial coverage is needed.
Methods d 11 B analyses. The d 11 B values of fossil Porites spp. were measured following the protocol of Ishikawa and Nagaishi 47 . Fossil corals were screened for diagenetic alteration with X-ray diffraction and geochemical analyses, as well as visual using a Scanning Electron Microscope 48 . Bulk sampling was conducted along the growth axis, and time resolution of each sample is several years (1-8 years) depending on growth rate of each coral 48 . Typically 6 mg of carbonate was used for d 11 B measurement. After removals of organic matter using 30% H 2 O 2 for ,12 hours, boron was purified by cation and anion exchange using AG 50 W X12 and 1-X4 resin (Bio-Rad, USA) and then d 11 B were measured using the positive polarity thermal ionization mass spectrometer (P-TIMS; Thermo Finnigan TRITON) installed at Kochi Core Center, Japan Agency for Marine-Earth Science and Technology. All reported d 11 B values are the mean of duplicate analyses (Table S1). Repeated analysis of the JCp-1, carbonate standard provided by Geological Survey of Japan yielded 24.21 6 0.18% (2s, n 5 18), which is the finest precision to date 47 . Differences between the duplicates are 0.08% on average with the largest one of 0.28% (Table S1), which is within the measurement uncertainty of JCp-1. We conservatively report 60.18% as the analytical uncertainty of our d 11 B measurements. d 11 B-pH calibration and pCO 2 calculation. First, the analytical procedure-specific isotopic offset 49  R compilation. Published 14 C (radiocarbon years) and U/Th ages of fossil coral samples obtained during IODP Exp. 310 were compiled in order to calculate residual radiocarbon activities (D 14 C) and R. We verified via IODP sample ID and core photographs 18 that the exact same samples were selected (Table S2). In some cases different portion of the skeleton of the same coral was dated. Given that lifetimes of coral are generally less than several decades, temporal gaps derived from subsampling are negligible in calculations of D 14 C and R. We did not use 14 C ages from either microbialite (carbonate created by bacteria) or encrusting coralline algae from equivalent down-core depths due to a possibility of post-depositional growth (for details, see ref. 53). Calculation was done according to equations (3) and (4) where 14 C-age is an original radiocarbon data before a local R correction 54,55 .  58 ), thus it differs from DR that conventionally represents local 14 C reservoir age. We estimate that the accumulated uncertainty in the R diff calculation are the sum of errors in 14 C dating, U/ Th dating, and modern R ( Fig. S5b; Table S2).