Southern Ocean contribution to both steps in deglacial atmospheric CO2 rise

The transfer of vast amounts of carbon from a deep oceanic reservoir to the atmosphere is considered to be a dominant driver of the deglacial rise in atmospheric CO2. Paleoceanographic reconstructions reveal evidence for the existence of CO2-rich waters in the mid to deep Southern Ocean. These water masses ventilate to the atmosphere south of the Polar Front, releasing CO2 prior to the formation and subduction of intermediate-waters. Changes in the amount of CO2 in the sea water directly affect the oceanic carbon chemistry system. Here we present B/Ca ratios, a proxy for delta carbonate ion concentrations Δ[CO32−], and stable isotopes (δ13C) from benthic foraminifera from a sediment core bathed in Antarctic Intermediate Water (AAIW), offshore New Zealand in the Southwest Pacific. We find two transient intervals of rising [CO32−] and δ13C that that are consistent with the release of CO2 via the Southern Ocean. These intervals coincide with the two pulses in rising atmospheric CO2 at ~ 17.5–14.3 ka and 12.9–11.1 ka. Our results lend support for the release of sequestered CO2 from the deep ocean to surface and atmospheric reservoirs during the last deglaciation, although further work is required to pin down the detailed carbon transfer pathways.


1
, Δ 14 C (corrected 14 C activity) 2 , and atmospheric δ 13 C-values 3 . These patterns indicate that vast amounts of CO 2 were released from a reservoir (or reservoirs) with low 14 C and δ 13 C values during the last deglaciation. Evidence for the existence of such reservoirs has been found in all sectors of the Southern Ocean 4-8 , the North Pacific 9 , as well as northern permafrost soils 10,11 . However, evidence for the pathway of sequestered carbon from the deep-waters to the surface and ultimately the atmosphere is limited. Upwelling of Circumpolar Deep Waters (CDW) in the Antarctic and Subantarctic Zones of the Southern Ocean is the most likely pathway for the carbon-rich waters to debouch carbon to the atmosphere. These upwelled waters are subsequently subducted and exported northward as intermediate-and mode-waters (Fig. 1).
Any change in the sequestration and cycling of CO 2 in the ocean will affect the marine inorganic carbon system. In this respect, reconstructions of water mass carbon chemistry (carbonate ion concentrations; [CO 3 2− ]) can yield important insights into the effect of past changes in dissolved inorganic carbon (DIC), and the oceanic carbon reservoir.
Here we present reconstructions of [CO 3 2− ] and δ 13 C on benthic foraminifers (Cibicidoides wuellerstorfi and C. dispars) from an intermediate-water record recovered at 44.4°S offshore New Zealand (PS75/104-1; 835 m; Fig. 1). PS75/104-1 is bathed by Antarctic Intermediate Water (AAIW) 12 , and thus changes in the [CO 3 2− ] will reflect changes in DIC that could provide clues to air-sea gas exchange occurring upstream in the Southern Ocean prior to the subduction of these waters to intermediate depths. Our study presents evidence for deglacial changes in the Pacific DIC pool, highlighting a pathway of sequestered CO 2 from the ocean to the atmosphere during Heinrich Stadial 1 (HS1; ~ 18-14.5 thousand years before present (ka)) and the Younger Dryas (YD; ~ 12.9-11.6 ka).  71 . AAIW path (grey arrows) according to Bostock et al. 12 . Map created with GeoMapApp version 3.6.12 (https:// www. geoma papp. org).   (Fig. 2). Carbonate ion concentrations of core PS75/104-1 range from 31.6 µmol/kg up to 119.1 µmol/kg (Fig. 2). The most pronounced increase in [CO 3 2− ] (60.3-109.8 µmol/kg) occurred between 17.2 and 14.5 ka and is paralleled by an equally pronounced rise in δ 13 C benthic from ~ 1.25 to 1.6‰. Both records display a second, yet less pronounced, increase between 13.55 and 10.52 ka (Fig. 2). During the Holocene, both records diverge, with increasing δ 13 C and decreasing [CO 3 2− ]. Following the YD disturbance to the inorganic carbon system, carbonate compensation drives back the system back to its original state 14 , while δ 13 C continues to increase 15 .

Discussion
Changes in Southern Ocean deep-water ventilation are often used to explain the two deglacial pulses of rising atmospheric CO 2 5,7,8,16,17 (Fig. 3). However, while being an indicator for deep-water residence time, radiocarbon alone does not allow for the direct analysis of the past oceanic carbon pool and changes in DIC. Barring lateral water mass transport, the primary driver for changes in δ 13 C DIC and [CO 3 2− ] is the biological pump, the export of organic matter from the surface into deeper water masses, and its subsequent degradation is a primary driver for changes in δ 13 C DIC and [CO 3 2− ] 14,18 . With progressive export of carbon (CO 2 ), the biological carbon pump increases the DIC content of a given water mass, while decreasing its δ 13 C and [CO 3 2− ] values. Depending on the balance between CO 2 sequestration via the biological carbon pump, circulation and ventilation via the upwelling of deep water masses, the circumpolar Southern Ocean can alter between a carbon sink or source. 2− ] and δ 13 C during HS1 (Fig. 3), such as increased surface export, release of carbon-rich fluids from pockmarks, or the discharge of CO 2 from an oceanic carbon reservoir that built-up during the preceding glacial and fed into the formation area of SW Pacific AAIW. In this case the area, where surface waters are influenced by both the upwelling of deep waters south of the APF, and then air-sea exchange processes as it flows north due to Ekman transport and subducted as intermediate waters. During glacial times, a combination of multiple climatic factors enhanced the ability of the Southern Ocean to sequester CO 2 . Lower surface temperatures allowed for increased uptake of CO 2 19 , while higher fluxes of iron-rich dust 20 resulted in increased primary productivity as a result of iron fertilization of the nutrient-rich Subantarctic Sector of the Southern Ocean 21 . This led to enhanced export of carbon to the deep ocean via the biological pump in the Atlantic sector. However, there is no clear evidence of increased productivity in the Subantarctic Sector of the SW Pacific 22-25 , while it was decreased in the Antarctic Sector 23,24 . 230 Th fluxes of biogenic matter in our research area have shown no significant local change in export production off New Zealand since the LGM 25 . Thus, we assume that changes in productivity and surface export did not play a dominant role in driving SW-Pacific AAIW [CO 3 2− ] at our core site. Yet, as AAIW chemistry is also influenced by regional, zonal, and meridional processes, we acknowledge the further need for additional investigations to test the feasibility of our interpretation. The distinct anti-phased pattern of atmospheric δ 13 C (Fig. 3f) 3 and PS75/104-1 suggests that AAIW δ 13 C is not driven by the atmosphere via air-sea gas exchange. An additional factor that potentially influenced water mass carbon chemistry in the SW Pacific might have been the release of CO 2 from pockmarks that are documented in the research area 26 . During the late glacial and early deglacial, carbon-rich fluids were probably released from the seafloor off New Zealand and might have contributed to extremely low CDW and AAIW 14 C values 8,26 . The injection of CO 2 would also affect [CO 3 2− ] of their respective water masses (Fig. S1). However, as the records of PS75/104-1 14 C 8 does not point toward a deglacial influence of 14 C-dead CO 2 from pockmarks 26 (Fig. 4) and as the transient increase in [CO 3 2− ] is interpreted as a loss of CO 2 from a water mass 14,27 , we expect that the release of carbon-rich fluids did not play a role in the evolution of AAIW [CO 3 2− ] during the time interval covered by our study. The expansion of Antarctic sea ice toward the north 28 HS1 was marked by the most pronounced increase in atmospheric CO 2 1 and was paralleled by a similar increase in PS75/104-1 [CO 3 2− ] of ~ 58 µmmol/kg (Figs. 2 and 3). Transient increases in [CO 3 2− ], as recorded by PS75/104-1, reflect a loss of CO 2 and the subsequent return to the previous state of the marine carbon system via carbonate compensation 14,27 . The inverse relation of [CO 3 2− ] and water mass pCO 2 38 , imply that the observed [CO 3 2− ] rise is consistent with the HS1 release of CO 2 via the Southern Ocean (Fig. 5b). The release of sequestered CO 2 is furthermore supported by our record of benthic δ 13 C that closely parallels the increase in observed [CO 3 2− ] as well as the patterns observed in AAIW RR0503-79 17,39 (Figs. 3 and 6) and Bay of Plenty δ 18 O and δ 13 C gradients 40 . This trend is in good agreement with the loss of metabolic (high 12 C) CO 2 via air-sea gas exchange in the formation area of AAIW 40 (Fig. S1). However, other factors such as the thermodynamic effect, an overprint of the atmospheric signal 41 , changes in export production (EP), and changes in AAIW formation should be considered when interpreting our records. Air sea exchange under colder temperatures shifts the δ 13 C values toward higher values [42][43][44] . Thus, warming temperatures tend to shift the system in the direction of higher δ 13 C values 39 . Our HS1 and YD trends in δ 13 C and [CO 3 2− ] (Fig. S1) are more in line with the slope predicted for regenerated organic carbon 39 , and thus imply that carbon sequestration via the biological pump, and release via ventilation was the more dominant driver. Another process that was observed to have a pronounced influence, is the overprint of atmospheric δ 13 C values on surface and recently ventilated waters 41 . Several δ 13 C-records follow the atmospheric pattern, while minor differences can be attributed to different temperatures during air-sea exchange 41 . In contrast to these records, PS75/104-1 δ 13 C is antiphased to the atmospheric record 3 during HS1 and the YD (Fig. 3). During the Holocene, however, we observe a strong correlation between atmospheric and AAIW δ 13 C values that imply an overprint as proposed by Lynch-Stieglitz et al. 41 . Any pronounced decrease in www.nature.com/scientificreports/ EP might also affect and increase both δ 13 C and [CO 3 2− ]. In the South Atlantic, a reduction in EP coincided with a HS1 increase in [CO 3 2− ], pointing to decreased biological productivity as a contributing factor to rising CO 2 45 . A similar process could presumably have driven or contributed to both pulses observed during HS1 and the YD. The analysis of 230 Th normalized fluxes of biogenic opal, carbonate, and excess barium on a suite of sediment records from the SW-Pacific indicate no pronounced change in EP since the LGM 25 . Thus, while a contribution of changes in EP cannot be excluded, we assume that it would only have a subordinate effect on the patterns observed. Given that AAIW could integrate signals from broad Southern Ocean regions due to homogenization by Antarctic Circumpolar Current, we want to encourage further work. Another factor to consider is the formation of AAIW and its subduction toward our core location. If the formation of low [CO 3 2− ] AAIW is reduced, an increase in concentrations at our core site can be expected. The formation of AAIW is closely coupled to the applied wind stress 46 . During the glacial, stationary SW-Pacific ocean fronts 47 in combination with a northward  www.nature.com/scientificreports/ displacement of Southern Westerly Winds 24 , reduced the wind stress experienced in the formation area of SO AAIW. These processes in combination to changing salinity contrasts might have been coupled to the reduced the glacial subduction of SO AAIW (Ronge et al., 2015). Given these local effects, we expect only a subordinate role on the patterns recorded in PS75/104-1. The HS1 pulse of increasing atmospheric CO 2 was accompanied by the most dramatic drop of atmospheric δ 13 C as reconstructed Antarctic ice cores 3 (Fig. 3f). Thus, the release of CO 2 via the Southern Ocean, as implied by PS75/104-1 and Bay of Plenty records 39,40 , illustrates a likely mechanism that can account for the coevolution of atmospheric CO 2 and δ 13 C-values (Fig. 3) Throughout the ACR (14.5-12.9 ka) 48 , when atmospheric CO 2 plateaued 1 and its δ 13 C briefly returned to higher values 3 , intermediate-water [CO 3 2− ] and δ 13 C return to lower, glacial-like values (Fig. 3). This suggests that expanding sea ice during the ACR disrupted the communication of upwelling deep waters with the atmosphere, before being incorporated into AAIW. These findings contrast 14 C reconstructions from deep sea corals, in the South Tasman Sea bathed by AAIW 16 . The zonal asymmetry might be explained by regional differences in the movement of Southern Ocean fronts that are less constrained by sea floor topography in the Indo-Pacific sector 16 .
During the YD there is evidence in the ice cores for a second pulse of increasing CO 2 and decreasing δ 13 C, this is again accompanied by an increasing intermediate-water [CO 3 2− ] and δ 13 C in PS75/104-1 (Fig. 3). Thus, our AAIW record at PS75/104-1 points to a strong mechanistic link between Southern Ocean ventilation and atmospheric CO 2 that was active in the SW Pacific.
However, to understand the importance of this Southern Ocean pathway of CO 2 in the deglacial carbon system, we have to address several important questions: Which key regions contributed to the two-pulse, deglacial rise in atmospheric CO 2 7,34,39,49-52 ? What were the mechanisms and reservoirs that resulted in the release of CO 2 10,11,53,54 , and which role did these have on the patterns observed in intermediate waters off New Zealand (PS75/104-1)?
In combination with shifting Southern Westerly Winds 24 , the deglacial retreat of Antarctic sea ice 29 resulted in an intensification of upwelling of Circumpolar Deep Water throughout the Southern Ocean (Fig. 3e) 50 . Deepwater records from the Southern Ocean show significant perturbations during HS1 and the early deglacial. Radiocarbon values from the Atlantic, Indian, and Pacific Sectors of the Southern Ocean indicate an increase in deep-water ventilation that reflects the renewed contact of the deep and shallow overturning cells and thus exchange of the glacial carbon pool with surface waters and the atmosphere 5,7,8 . Simultaneously, a rise in CDW [CO 3 2− ] 39,52 and coral-derived δ 11 B 34 reflect the loss of CO 2 from this lower cell, coeval with increasing deepwater ventilation throughout HS1 (Fig. 3). Increasing pH in the lower overturning cell and decreasing pH in the upper cell indicate a release of CO 2 from the deep Drake Passage during HS1 34 . However, our understanding of the lower cell carbon chemistry in other sectors of the Southern Ocean is still poorly constrained. In this respect, our AAIW data from the Pacific Sector suggest a close link between sea ice and westerly winds 24,29 , upwelling intensity 50 , deep-water ventilation and carbonate chemistry 8,39,52 , and atmospheric CO 2 (Figs. 3 and 5). Intensified air-sea gas exchange triggered a transient increase in [CO 3 2− ] in the formation area of AAIW. The signal from this process was subsequently exported from the formation area of SW Pacific AAIW 12 toward the core location of PS75/104-1 (Figs. 1 and 5). During HS1, the abrupt reduction in Bay of Plenty intermediate-water Δδ 13 C (663-1165 m) suggests the loss of CO 2 from SW Pacific AAIW 40 . Radiocarbon reconstructions from a southeasterly bathymetric transect off New Zealand likewise identified the ventilation from upwelling deepwaters in the formation area of Southern Ocean AAIW 8 . HS1 upwelling and ventilation of carbon-rich deep waters did not affect the ( 14 C) ventilation of PS75/104-1 8 (Fig. 4), while [CO 3 2− ] in our record as well as nearby AAIW record RR0503-79 39 show a significant excursion to higher values. Collectively this argues for a loss of CO 2 in the formation area of AAIW.
In combination with other Southwest Pacific records that display a similar HS 1 pattern 8,15,17,39,52 our AAIW data highlight the importance the Southern Ocean's Pacific pathway had on the HS1 atmospheric CO 2 increase. The majority of AAIW reaching our core location is formed directly to the south as so called SO AAIW 12 . Nevertheless, given the similar pattern evident in the Bay of Plenty records 39 it could also be due to an upstream contribution from the SE Pacific (the primary formation region of AAIW) or the Indo Pacific via the Antarctic Circumpolar Current 7,16 .
Following HS1, the southern hemispheric ACR was marked by a reduction in Southern Ocean upwelling rates, however less pronounced than during the LGM (Fig. 3e) 50 . In PS75/104-1 [CO 3 2− ] and δ 13 C values rapidly decreased during the ACR (Fig. 2). The decrease in opal flux (suggested to be an indicator of upwelling of carbonrich deep-waters) 50 and an enhanced winter and spring sea ice cover 55 , reduced the carbon loss in the formation area of AAIW. Throughout the Austral summer and autumn, ACR sea-ice and biological feedbacks increased the sequestration of CO 2 in the high southern latitudes 55 . In combination, these factors provide a likely scenario as the mechanism for the ACR trends seen in our record and highlight the fact that AAIW off New Zealand (PS75/104-1) can trace upstream changes in the Antarctic Zone.
Following the ACR, during the YD AAIW [CO 3 2− ] (this study) and opal flux suggest a reinvigorated upwelling in the Antarctic zone of the Southern Ocean (Fig. 3e) 50 . This is supported by other records from the Southern Ocean. In the South Atlantic, two records from 4276 m 56 and 4981 m 57 point toward a progressive deepening in the erosion of the deep ocean carbon pool. Similar to the South Atlantic 57 , it is likely that the ventilating water masses came from below ~ 4300 m. Radiocarbon-based reconstructions of deep-water ventilation show that down to this depth, deep-water ventilation reached modern-like values at the end of HS1 8 . There is also evidence for a Southern Ocean contribution to the second pulse in atmospheric CO 2 from rapidly decreasing pH values of the lower cell in the Drake Passage (Fig. 3d) 34 , and the Southern Indian Ocean off the Kerguelen Archipelago where steepening isohalines and isopycnals decreased stratification and allowed for a resumption of deep-water ventilation 7 . Other records of deep-and intermediate water [CO 3 2− ] 39,52 , and radiocarbon 8 from Scientific Reports | (2021) 11:22117 | https://doi.org/10.1038/s41598-021-01657-w www.nature.com/scientificreports/ the SW Pacific, lack sufficient resolution across the YD. Our study provides more data points and significantly improved chronological constraints during this time interval 13 . Thus, while the main part of the atmospheric CO 2 -increase during the YD is thought to be due to the thawing permafrost soils on the northern hemisphere 10,11 , our data can now point to southern hemispheric contribution of CO 2 at this time via outgassing from the Southern Ocean (Fig. 5b). During the Holocene AAIW records RR0503-79 39 and PS75/104-1 begin to diverge (Fig. 3). The youngest values of both records agree with modern [CO 3 2− ] data 39,58 , likely the result of two different sources and pathways of AAIW at the core sites; Tasman AAIW at the Bay of Plenty site, and SO AAIW at our Chatham Rise site, as defined by Bostock et al. 12 .
During the last deglacial, the history of AAIW [CO 3 2− ] and δ 13 C in our record closely trace Southern Ocean upwelling rates from opal flux 50 , as well as atmospheric δ 13 C 3 and CO 2 1 (Fig. 3). Between ~ 18 and 11 ka, changes in the extent of Antarctic sea ice cover 59 and the meridional shift of the southern westerly winds 54 modulated the upwelling rates of CDW 50 . In combination with changes in the efficiency of the biological carbon pump 53 , the increased communication of CO 2 -rich deep-waters via Southern Ocean upwelling was the main driver of early deglacial atmospheric CO 2 . Both transient peaks, observed in PS75/104-1 during HS1 and the YD, are indicative of a loss in CO 2 14,27 in the formation area of AAIW. The very close relationship between AAIW [CO 3 2− ] and δ 13 C with atmospheric patterns (Figs. 3 and 5) highlights the Southern Oceans role on the deglacial climate. Deglacial deep-water records of Bay of Plenty [CO 3 2− ] 39 and δ 13 C 17 (Figs. 3c and 6), and Bounty Trough ΔΔ 14 C 8 (Fig. 4) indicate a progressive change in water mass properties, indicative of circulation induced shifts in water mass mixing and/or loss of respired CO 2 through ventilation 39 .
During time periods with cold northern hemispheric stadial conditions (HS1 or YD), the bipolar seesaw hypothesis 60 argues for a reduction in the efficiency of the AMOC. A diminished AMOC results in the reduced export of heat from the southern hemisphere to the northern hemisphere, ultimately triggering a decrease in Antarctic sea ice that contributed to Southern Ocean release of CO 2 . The warmer northern hemispheric Bølling-Allerød period again resulted in a strengthening of North Atlantic Deep Water formation and the AMOC 61 . This period roughly correlates to the ACR that saw an increase in Antarctic sea ice, a northward displacement of southern westerly winds and reduced upwelling of CDW 50 . As our data show, the ACR was also marked by a return to glacial-like [CO 3 2− ] and δ 13 C (Fig. 3) and thus reduced air-sea gas exchange in the Southern Ocean, the area of AAIW formation. The well constrained temporal evolution of AAIW [CO 3 2− ] and δ 13 C throughout the entire deglacial period (Fig. 3) provides important new insight into the key role, the Southern Ocean played in the two-step rise of atmospheric CO 2 .

Conclusions
Our investigation of foraminifer-based [CO 3 2− ] and δ 13 C records on an intermediate water core off New Zealand highlight the role SW-Pacific AAIW played during the deglacial rise in atmospheric CO 2 . In conclusion, we propose that: 1. Reconstructed [CO 3 2− ] and δ 13 C trends point to a potential release of respired CO 2 through ventilation in the upwelling region of circumpolar deep water. 2. Our findings agree with previous studies from the region that indicated that the mid-depth Pacific acted as a reservoir for CO 2 during the last glacial 8,39,40,52,62 . 3 2− ] and δ 13 C during HS1 are consistent with the release of CO 2 during this interval. While this interpretation is not unambiguous, it adds to a growing set of studies that indicate a similar process 8,39,40,50,52,63 . 4. In addition to northern hemisphere sources 10,11 , the YD rise in atmospheric CO 2 , might have experienced a contribution of released CO 2 from the South Pacific as well. 5. Throughout the Holocene, AAIW δ 13 C probably experienced an overprint from atmospheric values 41 . 6. C. dispars can be used for reconstructions, using our new calibration B/Ca = 2.27(Δ[CO 3 2− ]) + 152.37

Materials and methods
Sediments and sample treatment. We analyzed sediment core PS75/104-1 that was retrieved during expedition ANTXXVI/2 at S44° 46′ 9.012′′ E174° 31′ 31.8′′ in a water depth of 835 m (AAIW), using a BGR type piston corer. The core was split into an archive and a working half and subsequently sampled. All samples were frozen and freeze dried for 2-3 days.  13 . Some depths of PS75/104-1 might be affected by pronounced Zoophycos burrows that were mapped by densely spaced 14 C-samples and X-radiographies 13 . However, only four of our samples fall into these intervals (Fig. 2). Excluding these would not affect any of our interpretations. Hence, we are highly confident in the integrity of our records.

B/Ca measurements and [CO 3
2− ] calculation. B/Ca measurements were conducted on the 315-400 µm fraction of specimens of the epibenthic foraminifer species Cibicidoides wuellerstorfi and Cibicidoides dispars 64 , which showed no sign of alteration or secondary fillings. B/Ca analyses were conducted at the GEOMAR Helmholtz Center for Ocean Research in Kiel, using a Coherent GeoLasPro 193 nm Excimer laser ablation system, coupled to a Nu Instruments AttoM magnetic sector mass spectrometer. LA-ICP-MS is a well-established method for the analysis of foraminiferal calcite [65][66][67] . The analytical method we used for this study has been proven to be accurate and precise 68 (instrument details given in the supplementary information). For each sample 3-6 specimens were analyzed on four 90 µm spots in the three oldest chambers on the umbilical side. Before and after each set of five specimens, the NIST615 standard 69 was measured and used for calibration. Before beginning the analyses, each shell as well as the NIST615 standard were pre-ablated to prevent any surface contamination effects. Samples with ratios of Mn/Ca > 0.2 mmol/mol and Al/Ca > 0.4 mmol/mol were discarded from the dataset. For our calculations, we applied the calibration of Yu et al. 56  ]. While probably more pronounced in intermediate-waters, than deep-waters, changes in S, BWT, and P are within the uncertainty of the proxy's calibration 39,52 .