Contributions of biological and physical dynamics to deglacial CO_2 release from the polar Southern Ocean

Domain: earth_science
Published: 2025-03-18
Editorial outcome: minor_revision

Abstract
As a critical region regulating air-sea gas exchanges, the polar Southern Ocean has important implications for deglacial atmospheric CO2 rises. However, proxy data evidence is sparse to evaluate the respective roles of Southern Ocean biological and physical dynamics in affecting past air-sea CO2 exchanges due to longstanding challenges in obtaining carbonate material to reconstruct surface conditions in this region. Here, we circumvent these challenges by constraining polar Southern Ocean surface-water conditions based on preformed deep-water properties derived from paired carbonate ion-phosphate-oxygenation reconstructions during the last deglaciation. We show that polar Southern Ocean carbon losses coincided with increased deep-ocean preformed nutrient concentrations, highlighting reduced biological carbon utilization as a key process for deglacial CO2 outgassing. By comparing total carbon losses with those attributable to biological processes, we further show that enhanced physically-driven air-sea gas exchanges in the polar Southern Ocean strongly drove CO2 outgassing towards the end of the last deglaciation.

Introduction
Ice core measurements show rapid atmospheric CO2 increases during the Late Pleistocene glacial-interglacial transitions1–3. These deglacial atmospheric CO2 increases are thought to be linked to strengthened release of carbon previously stored in the deep ocean through the Polar Antarctic Zone (PAZ)4–6. Climate models suggest that changes in the PAZ enhanced the release of carbon previously sequestered via both biological and physical processes and thus raised atmospheric CO2 during deglaciations7–13. However, mechanisms controlling past atmospheric CO2 fluctuations remain elusive due to the lack of paleoclimate proxy records in the PAZ. This presents as an obstacle to confidently link PAZ processes to changes in deep-ocean carbon storage and atmospheric CO2. Given the critical role of the Southern Ocean in governing atmospheric CO2 on various timescales in the past and future4,14,15, it is imperative to fully evaluate the respective roles of PAZ biological and physical processes in affecting atmospheric CO2 using marine paleoclimate archives. Nevertheless, a longstanding challenge to investigate the PAZ is the difficulty to obtain sufficient carbonate materials for paleoclimate reconstructions.

Here, we circumvent this challenge by constraining PAZ conditions using two quasi-conservative properties of deep waters7,15,16: PO4* (refs. 16–18) and air-sea exchange signature of carbonate ion concentration ([CO32-]as)19–22. Following deep water formation, signatures of surface source waters can be preserved as “preformed” properties that are quasi-conservative in the ocean interior. PO4*, derived from concentrations of phosphate and oxygen ([PO43-] and [O2], respectively), has long been employed to infer the biological nutrient utilization efficiency and surface-water oxygen saturation in deep-water formation regions (Methods)16–18. PO4* is quasi-conservative because phosphate regeneration is stoichiometrically linked to oxygen consumption during biogenic matter decomposition in the ocean interior. For a given water mass, higher PO4* largely indicates less efficient nutrient utilization7,16,17,23. Recently, [CO32-]as has been developed to constrain CO2 exchanges between surface waters and the atmosphere19–22. [CO32-]as is quasi-conservative in the ocean interior and represents changes in seawater carbonate ion concentration ([CO32-]) due to the net air-sea CO2 exchange when source waters leave the surface ocean, after correcting for anthropogenic carbon influences19,20. In addition to air-sea CO2 exchanges in source-water regions, the distribution of deep-water [CO32-] is driven by changes in physical conditions in the ocean (salinity, temperature, and pressure), remineralization of sinking organic matter, and dissolution of carbonate. To reveal the effect of air-sea exchanges on [CO32-], calculation of [CO32-]as accounts for [CO32-] changes driven by (1) physical conditions, (2) redistributions of dissolved inorganic carbon (DIC) and alkalinity related to regeneration of biogenic organic matter and carbonate, which can be estimated by [PO43-] variations, and (3) when considering temporal evolutions, global ocean DIC and alkalinity changes due to imbalances between riverine inputs and marine carbonate burial (Methods). Higher deep-water [CO32-]as values reflect more net CO2 losses through air-sea gas exchanges in source regions of the deep water19,20.

Seawater PO4* and [CO32-]as in the deep ocean is largely determined by mixing between southern sourced waters (SSWs) formed in the Southern Ocean and northern sourced waters (NSWs) formed in the North Atlantic (Fig. 1). In the preindustrial ocean, SSWs have higher PO4* than NSWs, reflecting less efficient nutrient utilization in the Southern Ocean7,16,17. Preindustrial SSWs have higher [CO32-]as than NSWs because CO2 tends to be released in the PAZ due to upwelling of high-DIC deep waters and inefficient biological utilization of nutrients and carbon4,15,22. Consequently, SSWs are less efficient to sequester atmospheric CO2 than NSWs19. Given known water-mass mixing proportions, PO4* and [CO32-]as reconstructed in the deep ocean away from polar regions can be used to infer water-mass endmember compositions and thereby surface conditions in deep-water formation regions. By contrast, non conservative tracers such as [PO43-], [CO32-], and [O2] can vary substantially due to organic matter respiration within the ocean interior. The advantage of using PO4* and [CO32-]as is illustrated by preindustrial ocean hydrography. For example, deep waters in the Pacific show large meridional gradients in [PO43-], [CO32-], and [O2], largely due to biogenic matter respiration on their northward transport route (Fig. 1 and Supplementary Fig. 1). By contrast, PO4* and [CO32-]as in this ocean basin display nearly stable values, reflecting mixing of NSWs and SSWs at relatively fixed proportions (Fig. 1 and Supplementary Fig. 1).

It is worth noting that PO4* and [CO32-]as are related to different aspects of the marine carbon cycle. While biologically-driven air-sea CO2 exchange would be accompanied by PO4* changes, purely physically-driven CO2 outgassing is largely independent of PO4*. Seawater [CO32-]as changes are not tied to any specific mechanisms but the integrate net CO2 air-sea exchange effects from both biological and physical processes. Therefore, combinational use of PO4* and [CO32-]as may allow for distinguishing the respective roles of biological and physical processes in controlling atmospheric CO2 changes. In this study, we reconstruct deep-water PO4* and [CO32-]as during the last deglaciation (18.0–11.7 ka) using sediment core MD97-2106 (45.15°S, 146.28°E, 3310 m water depth) from the Southwest Pacific (Fig. 1). Based on previous work23, deep waters at site MD97-2106 are composed of about 30% NSWs and 70% SSWs in the preindustrial.

Results
Figure 2 shows reconstructed deep-water [PO43-], [O2], and [CO32-] during the last deglaciation and their late Holocene (LH; 4.2–0 ka) averages (Supplementary Data 1). From the Last Glacial Maximum (LGM; 18–22 ka) to the LH, our reconstructions show that deep-water [PO43-] declined by 0.10 ± 0.04 µmol/kg (aggregated 1 standard deviation of two compared periods), and [O2] increased by 75 ± 21 µmol/kg, which is similar to previous records30 (Supplementary Fig. 4). Combining [PO43-] and [O2], the calculated PO4* shows an increase by 0.3 ± 0.1 µmol/kg from the LGM to the LH (Fig. 2). During the same time interval, deep-water [CO32-] at our site increased by 16 ± 6 µmol/kg (Fig. 2). In the absence of CO2 exchanges, warmer temperature, lower salinity, and deeper water depth together would lower [CO32-] at our site by ~2 µmol/kg during the LH compared to the LGM (Fig. 3). Also, assuming no CO2 exchanges, DIC and alkalinity decreases caused by regeneration of biogenic organic matter and carbonate that are linked to the [PO43-] decrease of 0.10 ± 0.04 µmol/kg would raise [CO32-] by 5 ± 2 µmol/kg from the LGM to the LH (Fig. 3). Global ocean DIC and alkalinity declines associated with CaCO3 input/output imbalances would lower [CO32-] by 26 µmol/kg (Fig. 3). Because the magnitude of the global alkalinity change cannot yet be reconstructed by proxies, we here rely on model simulations to calculate its effect on [CO32-]as31. However, different choices of modelled deglacial glacial alkalinity change only marginally affect the structure of the deglacial [CO32-]as changes (Supplementary Fig. 5). The offset between the reconstructed [CO32-] change and the expected [CO32-] changes without air-sea CO2 exchange is attributed to air-sea CO2 exchange signals, that is, [CO32-]as. Based on our calculations, we derive an LGM-to-LH [CO32-]as increase of 39 ± 6 µmol/kg. Past changes in stoichiometry compositions of remineralized biogenic matter would introduce uncertainties in our calculated PO4* and [CO32-]as, but they minimally affect our conclusion as shown by sensitivity tests (Supplementary Figs. 5, 6).

To infer LGM SSW PO4* and [CO32-]as from reconstructions at site MD97-2106, SSW-NSW mixing ratios (represented by SSW% and NSW% hereafter; SSW% + NSW% = 100%) at our site and the LGM NSW endmember values need to be constrained. In the preindustrial ocean, SSW% at site MD97-2106 is about 70%, based on PO4* of 1.36 µmol/kg at our site with SSW and NSW PO4* endmember values being 1.64 ± 0.06 and 0.73 ± 0.06 µmol/kg, respectively (Fig. 4; Table 1). The preindustrial [CO32-]as of 6.7 µmol/kg at our site is consistent with 70% SSWs with [CO32-]as of 10 ± 3 µmol/kg and 30% NSWs with [CO32-]as of -17 ± 5 µmol/kg (Fig. 4, Table 1). During the LGM, SSW% in the deep South Pacific was higher than the LH as indicated by neodymium isotopes (εNd) records32,33, due to shoaling of the upper limb of the global overturning circulation15,34–36. We thus assume the LGM SSW% at our site to be 80%. LGM NSW PO4* has not been reconstructed based on proxies but is expected to be less variable than the LGM SSW PO4* (refs. 7, 8, 15, 37). We assume the LGM NSW PO4* to be 0.58 ± 0.06 µmol/kg, which is 80% of the modern level, following ref. 37. LGM NSW [CO32-]as is inferred to be -43 ± 5 µmol/kg from a reconstruction in the North Atlantic (Supplementary Fig. 7). Considering the above SSW% and NSW endmember values along with our reconstructions at site MD97-2106, the LGM PO4* and [CO32-]as of SSWs are estimated to be 1.18 ± 0.13 and -34 ± 10 µmol/kg, respectively (Table 1). Additional sensitivity tests show that uncertainties associated with SSW% and NSW endmembers only marginally affect our final SSW PO4* and [CO32-]as results during the LGM (Supplementary Fig. 8). We find that LGM SSW PO4* and [CO32-]as were lower than their LH levels (Fig. 4), suggesting enhanced carbon sequestration in the PAZ during the LGM.

The reduced LGM SSW PO4* compared to the LH could be driven by lower preformed [PO43-] and/or lower preformed [O2]. Lower LGM preformed [PO43-] would indicate declines in surface-water phosphate and associated DIC in the PAZ7,8 that were driven by biological processes. For lower LGM preformed [O2], insufficient O2 air-sea exchange imposed by physical barriers such as sea ice that would retain O2 deficiencies of upwelled-and-then-downwelled waters in the PAZ38. However, purely physically-driven component of the lowered PAZ preformed [O2] is deemed to be relatively small between LGM and LH based on model simulations38. Therefore, declines in PAZ preformed [O2] must largely originate from previous organic matter remineralization that consumed O2 and produced CO2 in tandem. Such a biologically-driven PAZ preformed [O2] decline must be accompanied by surplus CO2 that is retained in the newly formed deep waters, because of the longer air-sea exchange equilibration time of CO2 than O2 (refs. 39, 40). Therefore, lowered LGM SSW PO4*, regardless of relative contributions of preformed [PO43-] and preformed [O2], is stoichiometrically linked to an more DIC sequestered by biological processes in the PAZ. If completely driven by more efficient nutrient utilization, an LGM-LH change in SSW PO4* of 0.46 ± 0.14 µmol/kg would be associated with a DIC decrease of 58 ± 18 µmol/kg due to enhanced biological CO2 uptake in the PAZ during the LGM. Such a change in biologically-driven DIC uptake may be slightly overestimated due to a potential but minor contribution of physically-driven disequilibrium to the lowered LGM SSW PO4*. Our LGM PO4* reconstruction shows that more phosphate and DIC supplied from the subsurface was biologically utilized in the PAZ, despite lower glacial export productivities in the region6,41. The associated biological processes would draw down surface-water CO2 partial pressure, with an effect to strengthen carbon sequestration in the ocean by a maximum amount of 58 ± 18 µmol/kg.

Our SSW PO4* reconstruction provides new evidence for increased carbon uptake during the LGM caused by more efficient nutrient utilization in the surface Southern Ocean as revealed by published nitrogen isotope (δ15N) records42–45. Importantly, our PO4* reconstruction reflects an integrated influence of biological processes on the Southern Ocean biogeochemistry and thus can complement surface-ocean δ15N records, which are presumably biased towards spring-summer seasonality of phytoplankton blooms that are offset from wintertime deep water formation. Moreover, our PO4* reconstruction demonstrates that reduced preformed nutrient concentrations in the PAZ, driven by more complete utilization of nutrient and carbon supplied to the surface Southern Ocean in spite of lowered export productivity6,41,45, was propagated downstream to great water depths during the LGM. Our PO4* reconstructions thus lend strong support for an enhanced biological pump in the PAZ to lower atmospheric CO2 during the LGM5,7,8,15.

Lower LGM SSW [CO32-]as also more directly demonstrates that atmospheric CO2 was more efficiently sequestered via the PAZ. The sensitivity of [CO32-]as change to DIC change is previously shown to depend on [CO32-]as (ref. 19) and is estimated to be -0.4 on average over the range of [CO32-]as values reconstructed at our site19. Based on this sensitivity, the reconstructed LGM-LH change in SSW [CO32-]as (-44 ± 10 µmol/kg) corresponds to 110 ± 25 µmol/kg more CO2 sequestration in the PAZ during the LGM that includes CO2 sequestration caused by both biological and physical processes. This carbon sequestration during the LGM in the PAZ derived from [CO32-]as, although dependent on the global alkalinity derived from model simulations, is larger than the upper limit of the biologically-driven air-sea CO2 uptake of 58 ± 18 µmol/kg estimated from PO4*. The large difference between CO2 uptake changes based on [CO32-]as and PO4* suggests that in addition to biological processes, physical dynamics such as fluctuations associated with sea-ice extent10,11, Southern Ocean overturning circulation34,43, and southern hemisphere westerlies44,46,47 are an essential contributor to enhanced carbon sequestered in the PAZ during the LGM.

Biologically driven PAZ CO2 outgassing during Heinrich Stadial 1

High sedimentation rates and tightly constrained chronology in our core allow us to generate detailed deglacial reconstructions (Fig. 5; Methods). During Heinrich Stadial 1 (HS1; 18.0–14.6 ka), PO4* at our site remained stable during the first 2 ka and rapidly increased by 0.4 ± 0.1 µmol/kg towards the end of HS1 (Fig. 5d). [CO32-]as at our site displayed a similar pattern, showing minimal change during early HS1 and an increase of 19 ± 3 µmol/kg during late HS1 (Fig. 5c). During HS1, SSW% at our site likely increased due to substantially decreased NSW production as indicated by proxy data48 (Fig. 5b). Everything else being equal, an SSW% increase from an LGM level of 80% to the extreme but unlikely level of 100% would raise PO4* and [CO32-]as at our site by 0.11 and 5 µmol/kg, respectively. Any deglacial NSW PO4* increase likely occurred after HS1, making little contribution to the reconstructed PO4* change at our site during HS149. NSW [CO32-]as increased by 7 µmol/kg as shown by previous reconstructions20 (Supplementary Fig. 7), which could raise [CO32-]as by at most 2 µmol/kg at site MD97-2106 assuming no SSW% increase (a condition that maximizes the NSW-endmember influence at our site). Therefore, we conclude that PO4* and [CO32-]as increases observed at our site during HS1 can be largely attributed to changes in the SSW endmember values.

During early HS1, both PO4* and [CO32-]as remained invariant at our site, suggesting that changes in air-sea CO2 sequestration in the PAZ were too small to be resolved beyond our reconstruction uncertainties. The roughly stable SSW PO4* during early HS1 coincided with minor nutrient utilization changes shown by opal flux41 and δ15N records42–45 in the Open Antarctic Zone (OAZ) (Supplementary Fig. 9), the latter of which is well dated in deep-sea coral samples (Fig. 5f). Thus, our new and literature data suggest that PAZ biological processes changed little during early HS1. Nearly constant early-HS1 SSW [CO32-]as further suggests little changes in air-sea CO2 exchanges in the PAZ during this period, including those may be caused by physical processes. This seem to contradict with other proxy records, which indicate favourable physical conditions for CO2 outgassing during early HS1. During this period, Antarctic ice-core sea-salt sodium concentrations suggest Southern Ocean sea-ice extent possibly retreated46, and sediment core radiocarbon data indicate improved deep Southern Ocean ventilation at multiple locations including our site28,30,50. According to models simulations9,51, enhanced deep Southern Ocean radiocarbon ventilation might be partially driven by changing geometry of the global overturning circulation that is strongly modulated by the sea-ice extent34. Although retreating sea ice itself would promote air-sea gas exchange, the lack of strong CO2 outgassing resolvable by our [CO32-]as reconstruction might be related to minor warming in the PAZ and deep Southern Ocean during early HS152, which has been shown by model simulations to be crucial to modulate physically-driven CO2 outgassing9,13. Nevertheless, further investigations are needed to better understand how physically-driven air-sea exchanges in the PAZ affected deep-ocean radiocarbon and carbon storage during early HS1.

During late HS1, increased PO4* suggests more biologically unutilized nutrients relative to the supply and thus a weaker biological pump in the PAZ. This is corroborated by δ15N records in the OAZ42–45 (Fig. 5f and Supplementary Fig. 9), which show prominent decreases in nutrient utilization efficiency over the same period. Ignoring the minor contribution of purely physically driven preformed [O2] to PO4*, PO4* that was raised by 0.4 ± 0.1 µmol/kg at our site suggests 44 ± 10 µmol/kg of DIC became unutilized by biological processes during late HS1. Rising SSW [CO32-]as during late HS1 also suggests that CO2 released through air-sea exchange in the PAZ during this period (Fig. 5c). Based on the [CO32-]as increase of 19 ± 3 µmol/kg, 46 ± 11 µmol/kg DIC was released from deep waters at site MD97-2106 during late HS1. Comparable DIC changes estimated based on PO4* and [CO32-]as demonstrate that increased DIC related to biologically unutilized nutrients can largely account for CO2 outgassing in the PAZ during late HS1. Physical processes must have facilitated outgassing of these biology-related DIC during late HS1. For example, given the rising export productivity in the OAZ during this period41 (Fig. 5g), declines in deglacial nutrient utilization efficiency during HS1 must be driven by more nutrients supplied by enhanced deep water upwelling5,15,41. However, direct contribution of physical processes to promoting CO2 outgassing independent of biological processes appear to be small during late HS1.

Reduced PAZ CO2 outgassing during the Antarctic Cold Reversal

During the Antarctic Cold Reversal (ACR; 14.6–12.8 ka), PO4* and [CO32-]as declined at site MD97-2106 by 0.22 ± 0.08 and 11 ± 5 µmol/kg, respectively (Fig. 5). During this period, reinvigorated production of low-PO4* and low-[CO32-]as NSWs48 would reduce SSW% and thereby lower PO4* and [CO32-]as at our site (Figs. 1, 4 and Supplementary Fig. 1). Without changes in NSW and SSW endmember values, the reconstructed PO4* and [CO32-]as declines could be largely explained by 30% reduction in SSW% from 100% (highest possible HS1 level) to 70% (ACR level). However, this would require no NSWs at our site during HS1 that is inconsistent with deep Southern Ocean εNd records32,33. A large SSW% reduction would also raise [O2] during the ACR, opposite to the reconstructed trends at our site (Fig. 2d) and the Drake Passage53. Therefore, SSW% decline is inadequate to explain all the reconstructed PO4* and [CO32-]as changes at our site. A moderate SSW% decline by 15% from HS1 would lower PO4* and [CO32-]as, respectively, by 0.12 and 4 µmol/kg, which are too small to account for our reconstructed changes. Hence, we conclude that SSW PO4* and [CO32-]as endmember values likely decreased during the ACR.

Our inferred SSW PO4* reduction suggests a decrease in the preformed nutrient and DIC due to more efficient nutrient utilization in the PAZ during the ACR, in line with OAZ δ15N records41–45 and enhanced marine productivity deduced from marine-derived aerosols recorded by ice cores54. The biologically-driven air-sea CO2 sequestration is also registered by the reconstructed [CO32-]as decline (Fig. 5c). In addition to biological changes, the enhanced CO2 sequestration may be linked to physical processes including readvancing Antarctic sea ice46 that limited air-sea gas exchange and northwards shift of the Southern Hemisphere westerlies44,53. We suggest that both biological and physical changes contributed to enhanced air-sea CO2 sequestration during the ACR.

Strong physically driven PAZ CO2 outgassing during the Younger Dryas

During the Younger Dryas (YD; 12.8–11.7 ka), deep-water PO4* and [CO32-]as increased by 0.10 ± 0.11 µmol/kg and 16 ± 5 µmol/kg at our site, respectively (Fig. 5c, d). Proxy data suggest that SSW% in the deep Southern Ocean during the YD declined towards the Holocene level32,35. Any decline in SSW% tends to lower PO4*, opposite to the reconstructed increasing PO4* at our site. The NSW PO4* possibly increased by ~0.15 µmol/kg during the YD49, with an effect of raising PO4* by 0.04 µmol/kg, assuming Holocene SSW% level of 70% at our site (a condition that maximizes the impact of NSW endmember changes). The remaining PO4* increase of 0.06 ± 0.11 µmol/kg at our site must be due to an increase in SSW endmember PO4*. We estimate that SSW PO4* was increased by 0.08 ± 0.17 µmol/kg with SSW% level of 70% at our site. For [CO32-]as, an NSW endmember [CO32-]as increase of 15 ± 7 µmol/kg (Supplementary Fig. 7) could account for at most 5 ± 2 µmol/kg [CO32-]as increase at site MD97-2106, assuming an SSW% of 70%. The remaining 11 ± 6 µmol/kg increase in [CO32-]as at our site must be caused by an SSW endmember [CO32-]as increase of 16 ± 9 µmol/kg. Therefore, both PO4* and [CO32-]as increases at MD97-2106 are strongly driven by increases in their SSW endmember values during the YD.

Our inferred SSW PO4* increase during the YD agrees with OAZ δ15N records41–45, indicative of lowered nutrient utilization efficiency in the OAZ. The 0.08 ± 0.17 µmol/kg increase in SSW PO4* suggests 8 ± 18 µmol/kg increase in DIC, which could cause a maximum magnitude of 8 ± 18 µmol/kg CO2 outgassing, due to biological changes in the PAZ during the YD. Meanwhile, the 16 ± 9 µmol/kg increase SSW [CO32-]as indicate a total air-sea CO2 loss of 40 ± 23 µmol/kg. Thus, the PAZ CO2 outgassing during the YD must be mainly attributed to physical processes including substantial sea ice retreat46, Southern Ocean circulation changes43, and southward shift of Southern Hemisphere westerlies44,46,47.

Methods
Trace element and stable isotope analyses

We examined sediment samples from 0–17 cm and 54–225 cm at site MD97-2106. For each sample, the >212 μm fractions of sediments were examined for benthic foraminifera Cibicidoides wuellerstorfi (sensu stricto)56, Hoeglundina elegans, and Globobulimina affinis. C. wuellerstorfi can be continuously found throughout the section. H. elegans is abundant above 158 cm, but absent below 207 cm. G. affinis can be found below 96 cm, but is absent above this level, except in 0–17 cm of the core. For trace element analyses, 4–20 C. wuellerstorfi shells and 2–8 H. elegans shells were picked from each sample. When the number of C. wuellerstorfi shells was larger than 15 or the number of H. elegans was larger than 6, uncrushed shells were split into two subsamples and treated separately in further steps.

Foraminifera shells were first gently crushed between glass slides. C. wuellerstorfi shells were cleaned following the modified Mg-cleaning protocol57,58 and H. elegans samples were cleaned following the Cd-cleaning protocol including the reductive step59. The cleaned samples were dissolved in 0.1 M HNO3 and trace element-to-calcium ratios were measured on an ICP-MS at the Australian National University following the method of ref. 60. Analytical uncertainties (1 standard deviation) are >2% for B/Ca and Cd/Ca. Relative standard deviations of replicated C. wuellerstorfi B/Ca and H. elegans Cd/Ca are 1.5-2.6% (3 samples) and 1.0–3.1% (6 samples), respectively. These replicated subsamples are plotted separately in Supplementary Fig. 2.

For carbon and oxygen isotope analyses, 2–5 C. wuellerstorfi shells and 2–5 G. affinis shells were picked from each sample from the same sediment depths. The analyses were performed at the Institute of Earth Environment, Chinese Academy of Sciences. The precision is better than 0.08‰ for δ13C and δ18O analyses, based on repeated measurements of an in-house dolomite standard (TTB1).

Estimating seawater biogeochemistry

Modern hydrographic data near site MD97-2106 (temperature, salinity, DIC, alkalinity, silicate, phosphate, oxygen, and anthropogenic carbon) are retrieved from the GLODAPv2 gridded database61,62. The anthropogenic carbon is deducted from the measured DIC to obtain the preindustrial DIC levels. Carbonate chemistry parameters are calculated from hydrological data using CO2sys 2.163.

Down-core in-situ deep-water [CO32-] ([CO32-]in-situ) is reconstructed from B/Ca in benthic foraminifera C. wuellerstorfi. [CO32-]in-situ can be defined by calcite saturation state (Δ[CO32-]) and calcite saturation values ([CO32-]sat):

[CO32-]in-situ = Δ[CO32-] + [CO32-]sat (1)

Δ[CO32-] can be estimated from C. wuellerstorfi B/Ca using the calibration from Yu, et al.64:

B/Ca = (1.14 ± 0.04) * Δ[CO32-] + 176.6 ± 1.0 (2)

We adopt the modern [CO32-]sat for the deglacial reconstruction, because [CO32-]sat is insensitive to past temperature, salinity, and pressure changes which would only cause ~0.1 μmol/kg variations in [CO32-]sat at Site MD97-2106 on glacial-interglacial timescales. We use a Monte-Carlo approach to combine the uncertainties in both proxy data and the age model. To combine these uncertainties, 10,000 iterations of the record were firstly randomly generated considering uncertainties associated with the B/Ca-[CO32-] calibration, B/Ca measurements, and the age model. Secondly, the 10,000 iterations were smoothed by a LOESS function. The 1σ uncertainty is defined as half of the difference between the 84th and 16th percentile of the smoothed curves. The combined ±1σ uncertainty of [CO32-] is, on average, ±4.5 μmol/kg for the range of B/Ca observed in core MD97-2106.

Past deep-water [PO43-] is derived from seawater Cd concentration (Cdw) reconstructed from Cd/Ca in benthic foraminifera H. elegans. Cdw is estimated by the following equation65:

Cdw (nmol/kg) = [(Cd/Ca)foram / DCd] * 10 (3)

where DCd is the partition coefficient of Cd between foraminifera carbonate and seawater. For H. elegans, DCd is 1.0 and independent of water depth25,66. We apply 5% uncertainty to DCd when estimating the uncertainty of seawater [PO43-]. Cdw is used to estimate [PO43-] based on the relationship between Cdw and [PO43-] for sites from the GEOTRACES transect GIPY06 located at the South of Tasmania67:

Cdw = 1.23 * [PO43-] + 1.30 ([PO43-] > 1.5 μmol/kg, r2 = 0.76, n = 294, RMSE = 0.117) (4)

Using different Cdw-[PO43-] relations (e.g., a linear correlation for data from the entire GIPY06 transect and a quadratic relation based on the whole GEOTRACES dataset, 2021v268) would result in different absolute [PO43-] values, but downcore [PO43-] patterns remain similar (Supplementary Fig. 10). We choose the local calibration because it yields a late Holocene [PO43-] estimation of 2.23 μmol/kg that is consistent with the modern hydrographic data. Combining uncertainties of DCd, RMSE of the Cdw-[PO43-] relation, Cd/Ca measurements, and the age model using the Monte-Carlo approach similar to that used for [CO32-] reconstructions, the average uncertainty of our [PO43-] reconstructions is ±0.11 μmol/kg (±1σ). Due to the relatively long phosphorus residence time in the ocean (~20–100 ka)69, we assume little change in the global oceanic [PO43-] inventory during the last deglaciation.

Past bottom water [O2] is estimated from the δ13C gradient (Δδ13C) between epifaunal species C. wuellerstorfi and deep infaunal species G. affinis using the following equation27:

Δδ13C = 0.0077 * [O2] + 0.41 (5)

The total uncertainty of the [O2] reconstruction is ±17 μmol/kg (±1σ), which includes both calibration and δ13C analytical errors27.

Calculating quasi-conservative tracers

Deep-water PO4* is derived from in-situ [O2] and [PO43-] based on the following equation16,17:

PO4* = [PO43-] + [O2]/175 - 1.95 (6)

The average ±1σ uncertainty in PO4* is 0.15 μmol/kg, derived from errors in [PO43-], [O2], and chronology using an iterated LOESS smoothing method. By definition16,17, calculation of PO4* employs a R-O2:P value of 175. Using other R-O2:P values (e.g., 160 or 180) to calculate PO4* changes absolute PO4* values, but with little effect on its deglacial structure (Supplementary Fig. 6). By definition and due to its conservativeness in the ocean, deep-water PO4* incorporates preformed [PO43-] and preformed [O2], that is, [PO4*]pre, which are derived by combining different proxies, changes in these proxies underlying the preformed tracers are linked by the following equation:

ΔPO4* = Δ[PO43-]pre + Δ[O2]pre/175 (8)

Following ref. 38, Δ[O2]pre can be further divided into two components related to physical and biological changes:

Δ[O2]pre = Δ[O2]pre_phy + Δ[O2]pre_bio (9)

Δ[O2]pre_phy is determined by surface O2 solubility, mainly driven by temperature and air-sea gas exchange disequilibrium not related to biological nutrient utilization.

Calculation of [CO32-]as is based on the carbonate system sensitivities to changes in physical conditions, remineralization of biogenic organic matter and CaCO3, and air-sea CO2 gas exchanges, which are derived from hydrographic data in the preindustrial ocean19. The carbonate system sensitivities are shown to be similar in preindustrial and LGM conditions19. We calculate deglacial [CO32-]as changes at site MD97-2106 in four steps following refs. 19, 20. In step 1, downcore deep-water [CO32-] is normalized to a reference condition (T=3°C, S=35, P=2500 dbar), using [CO32-]-T-S-P sensitivities from ref. 19. At site MD97-2106, we estimate deglacial changes in salinity based on global sea level changes, temperature based on benthic foraminiferal δ18O, and water depth based on global sea level changes, following methods described in ref. 19. In step 2, seawater [CO32-] changes linked to within-ocean DIC-alkalinity redistributions are calculated by changing in-situ [PO43-] to the global average of 2.22 µmol/kg, following the sensitivity of [CO32-] change to [PO43-] constrained in ref. 19. [CO32-] changes are driven by DIC and alkalinity changes related to [PO43-] via Redfield Ratio (RC:P = 127) and rain ratio (the ratio between organic and inorganic carbon in remineralized biogenic matter, Corg:CaCO3 = 4). Using sensitivities of [CO32-] change to [PO43-] defined by or different RC:P and Corg:CaCO3 minimally affects deglacial patterns of [CO32-]as (Supplementary Fig. 5). In step 3, deep-water [CO32-] changes linked to the global ocean DIC and alkalinity inventory changes due to CaCO3 input/output imbalance over time is calculated following refs. 19, 20. Global DIC changes associated with this process are half of the global alkalinity changes, which are the offsets between in-situ alkalinity estimated from model simulations31 and a reference alkalinity of 2385 µmol/kg. Effects of global DIC-alkalinity changes on [CO32-] are derived using the sensitivity of [CO32-] to DIC constrained in ref. 19. CaCO3 input/output imbalance mainly contributes to long-term [CO32-]as changes. Choosing different modelled deglacial global alkalinity changes leads to varying magnitudes of deglacial [CO32-]as changes but minimally affects millennial-timescale structure of [CO32-]as, as shown by sensitivity tests employing different modelled global alkalinity changes (Supplementary Fig. 5). In step 4, [CO32-] changes calculated in Steps 2 and 3 are removed from the [CO32-] normalized to a reference condition in Step 1 to obtain the corrected [CO32-], whose variabilities are attributed to air-sea CO2 exchange effects. To derive [CO32-]as, a constant of 78 µmol/kg is deducted from the corrected [CO32-]. Deducting this constant brings the late Holocene [CO32-]as values close to 0 µmol/kg, which are similar to hydrographic results in the pre-industrial ocean (Figs. 1, 3).

The uncertainty of [CO32-]as is also derived using the Monte-Carlo approach. Each [CO32-]as was calculated 10,000 times following the abovementioned methods using [CO32-], [PO43-], temperature, salinity, pressure, and age randomly generated from their averages and standard deviations. The combined ±1σ uncertainty of [CO32-]as is derived from the 16th and 84th percentiles of the iterated LOESS-smoothed curves, and our calculation yields an average 1σ-uncertainty of 7.7 μmol/kg for [CO32-]as.

To reliably infer millennial-timescale changes in PO4* and [CO32-]as, which are derived by combining different proxies, changes in these proxies underlying the preformed tracers need to be synchronized on the same timescale. To ensure synchroneity of these proxies, only reconstructions of [PO43-], [CO32-], and [O2], derived from different foraminifera found at the same sediment horizons were combined to calculate PO4* and [CO32-]as. Although the benthic foraminifera that we utilized dwell at varying depths within sediments, they were homogenized by bioturbation within sediments after deposition. This is evident by similar deglacial structures of oxygen isotopes in epifaunal and deep infaunal species in our core (Supplementary Fig. 2). Meanwhile, bioturbation does not prevent millennial-timescale deglacial changes being resolved, due to relatively high sediment rates (15 cm/ka) at our site during the last deglaciation. Finally, from the perspective of tracer calculation, millennial-timescale changes in PO4* and [CO32-]as are mainly driven by [O2] and [CO32-] with a minor contribution from [PO43-] (Fig. 3). For example, the concomitant increases in PO4* and [CO32-]as during the early HS1 are independent of [PO43-] changes.

Calculation of SSW endmember values and sensitivity test

We follow ref. 23 to select hydrographic data (GLODAPv2 dataset)61,62 to derive NSW and SSW endmember values in the preindustrial ocean. The NSW properties are derived from the North Atlantic sites north of 50°N and deeper than 1500 m, with CFC-11 concentrations > 0.5 pmol/kg. The SSW properties are derived from the Southern Ocean sites south of 45°S, and deeper than 1500 m, with CFC-11 concentrations > 0.5 pmol/kg and neutral density higher than 28.3 kg/m3 (Supplementary Fig. 1). Following these selection criteria, SSWs defined here are mainly ventilated in the PAZ and reside at deep (~>1500 m) depths in the Pacific Basin. A minor component of SSWs at deep depths is ventilated via the Subantarctic Zone70, but its contribution during the last deglaciation is smaller than the late Holocene (LH) due to the shallower upper-cell overturning circulation34, and hence considered to be negligible here.

PO4* and [CO32-]as levels during the LH, reconstructed from the top part of core MD97-2106, are consistent with estimations based on preindustrial hydrographic data. As both PO4* and [CO32-]as are quasi-conservative, LH PO4* and [CO32-]as at site MD97-2106 are determined by linear mixing between NSW and SSW endmembers based on the following equations:

PO4* LH_MD97-2106 = NSW% × PO4* LH_NSW + SSW% × PO4* LH_SSW (10)

[CO32-]as LH_MD97-2106 = NSW% × [CO32-]as LH_NSW + SSW% × [CO32-]as LH_SSW (11)

NSW% + SSW% = 1 (12)

Similarly, for the LGM, we have:

PO4* LGM_MD97-2106 = gNSW% × PO4* LGM_NSW + gSSW% × PO4* LGM_SSW (13)

[CO32-]as LGM_MD97-2106 = gNSW% × [CO32-]as LGM_NSW + gSSW% × [CO32-]as LGM_SSW (14)

gNSW% + gSSW% = 1 (15)

Equations (13)-(15) can be used to derive PO4* LGM_SSW and [CO32-]as LGM_SSW from PO4* LGM_MD97-2106 and [CO32-]as LGM_MD97-2106, if NSW-SSW mixing ratio and NSW endmembers during the LGM are known. The LGM water mass mixing and NSW endmember values are either not firmly quantified or associated with some uncertainties, but their influences on the LGM SSW endmember estimations can be quantified by sensitivity tests.

Based on paleocirculation reconstructions, SSW% in the deep South Pacific was higher during the LGM than the LH32,33. We therefore use 80% for the LGM SSW%, compared to 70% for the LH. For PO4*, there currently lacks coupled [PO43-] and [O2] reconstructions to constrain PO4* LGM_NSW. [PO43-] reconstructions show that LGM NSW was slightly more nutrient depleted than the modern ocean19,20,66, while the only published LGM reconstruction suggests similar NSW [O2] during the LGM and the LH71. We thus assume PO4* LGM_NSW to be 80% of the LH level following ref. 37 for our preferred calculations (Table 1), and use PO4* LGM_NSW the same as the LH level for sensitivity tests. We adopt LGM [CO32-]as reconstructed at North Atlantic site NEAP4K (61.5°N, 24°W, 1.6 km) as the [CO32-]as LGM_NSW (-43 ± 5 µmol/kg)20. However, glacial deep waters at this site may contain a small fraction of SSWs during the LGM due to the shoaling of NSWs66. We thus use [CO32-]as LGM_NSW derived from 90% NSWs and 10% SSWs at site NEAP4K for sensitivity tests. [CO32-]as LGM_NSW can be solved from [CO32-]as at sites NEAP4K and MD97-2106, with assumed SSW-NSW mixing ratios at both sites. As shown in Supplementary Fig. 8, our calculated LGM SSW PO4* and [CO32-]as employing different SSW% at our site and different NSW endmember levels are well within the ±1σ uncertainties associated with tracer reconstructions, supporting the robustness of our estimates of the LGM SSW endmembers.

Peer review rounds
[{"round": 1, "reviewer_comments": {"reviewer_1": "Review of Dai and Yu, “Deglacial polar Southern Ocean carbon release driven by physical and biological dynamics,” submitted to Nature Communications, by Jesse Farmer\n\nDai and Yu present new deep ocean B/Ca-derived carbonate ion, Cd/Ca-derived phosphate concentration, and d13C gradient-derived oxygen reconstructions from a sediment core in the Southwest Pacific Ocean tracking (largely) the properties of Southern Sourced deep waters over the last deglaciation. The authors translate this proxy data into information on changes to preformed phosphate and air-sea carbonate ion change to track the influences of Southern Ocean biology and gas exchange on carbon dioxide partitioning between the ocean and atmosphere over the last deglaciation.\n\nThere is a lot to like in this manuscript. It is overall well written, the figures are excellent, and the combination of proxies applied is novel and potentially quite transformative. At the same time, I feel that the authors are trying to reach a conclusion on the role of sea ice cover in deglacial ocean-atmosphere CO2 exchange before they have fully “paved the road” of proxy understanding that would allow the reader to get on board with this interpretation and exclude other alternatives. For this reason, I think the results would actually be more influential as a longer contribution to a field-specific journal where the proxy interpretations, and, most crucially, the covariance between these new proxies (see below) can be more fully explored. It may even be beneficial to split this into two manuscripts, one longer-form outlining the combined proxy approach and its uncertainties, and the second a shorter-form application to the deglaciation with its climate system implications.\n\nMajor comments (not in order of importance):\n1. Age model development. The C. wuellerstorfi d18O (ED Fig. 1) does not lend confidence in the age model construction that is based on G. bulloides Mg/Ca correlation to EDC dD (Dai et al., 2020). This is of crucial importance for the sea ice interpretation of outgassing in early HS1 – if the timing of HS1 is not correctly assigned in this sediment core, then the relationship of PO4* to sea ice (ice core ss Na) is uncertain. Unfortunately, I don’t find the chronological establishment here sufficiently compelling/precise as to allow for the discussion on L166 and onward, which is the crux of the current paper. Perhaps the published radiocarbon evidence can be brought in here, or other SST proxies can be measured to refine this alignment?\n2. Sea ice, or not? The authors take a strong position on the role of sea ice in the abstract “…physical dynamics involving sea ice retreats (sic) dominated CO2 outgassing…” (L22) but are rightly more circumspect about this in the main text “However, relative roles of biological and physical processes to the PAZ outgassing remain ambiguous on both millennial and glacial-interglacial timescales…” (L157-158). More clarity on the relative contributions of sea ice and their timing would be beneficial.\n3. Tracer interdependence. PO4* and [CO32-]as are not independent tracers, as both depend on Cd/Ca-derived [PO43-] (L48-49, 66-69). That suggests the difference between the two is largely driven by carbonate ion and oxygen… but it is not at all clear what to make of the distinctions between the two tracers. On the one hand, it’s interesting that the two show slightly offset patterns in time… but on the other, it’s not surprising that the two look similar (Fig. 2f, g). I would appreciate more guidance from the authors on what to make of this interdependence.\n4. NSW and SSW endmembers + contributions over time. The authors note that the advantage of PO4* and [CO32-]as is that both tracers are quasi-conservative in the ocean interior (e.g., L50-60). However, the downside of this is that these proxies are set by the endmember values of NSW and SSW and their relative contributions to the core site. It is known from the authors’ previous work that these endmember values are not stable over time (e.g., Yu et al., 2019 for [CO32-]as in the N. Atlantic); indeed, the goal of this study is essentially to reconstruct the SSW endmember PO4* and [CO32-]as over the last deglaciation (e.g., L63-65). To this end, it is critical that the authors know the contributions of the relative endmembers. In my mind, their approach of assuming a percentage contribution of SSW and allowing it to vary (e.g., L290) leaves much to be desired, especially as proxy data (Nd isotopes) may help address this question for them.\n5. Presentation of results. The authors do not discuss their underlying proxy reconstructions, instead they move directly to the derived tracer products (PO4* and [CO32-]as). This is unfortunate – the most directly calculated results should be presented first, and only after that a discussion of the interpreted derived products. Such additional length would allow for further investigation of intriguing features in the proxy data – for instance the offsetting (in both depth + age space) of carbonate ion, oxygen, and phosphate concentration deglacial maxima/minima (Fig. 2c-e). Might this relate to the relative depth habitat of the foraminifer used? (C. wuellerstorfi being ~sediment/water interface, H. elegans perhaps being infaunal, G. affinis being deep infaunal)?\n6. Uncertainty on PO4* and [CO32-]as. There does not appear to be an effort to estimate errors on PO4* and [CO32-]as aside from those relating to the individual proxy reconstruction of carbonate ion, oxygen concentration, and phosphate concentration. At the very least, the errors of each required measurement will add in quadrature; in addition, further error will result from assumption of Redfield stoichiometry (see modern ocean uncertainties in these ratios from Anderson & Sarmiento, 1994). A more comprehensive handling of uncertainty in these new tracers will be needed than is presented in the main text, figures and methods.\n\nMinor methods comments.\nL236. How was anthropogenic carbon calculated and deducted? I would not expect much detectible DICant at this water depth…\nL262. Given all those contributions, the estimated uncertainty on phosphate concentration is only ±0.05 µmol/kg? Come on now, that is not realistic. For one point of comparison, the Cd/Ca measurement precision alone in Farmer et al. (2019) corresponded to a similar or larger phosphate concentration uncertainty. I have no doubt the authors are better analysts than myself, but the propagation of error should be handled appropriately.\nL273-275. The effect of global alkalinity change on the reconstructed [CO32-]as is quite significant (ED Fig. 1i), and also there is no reason to expect that a single core site would experience the global ocean alkalinity change (see, e.g., regional deviations in benthic d18O for whole-ocean d18O change, Lisiecki & Stern, 2016 and references therein). Can the authors please further elaborate on/justify this correction and illustrate its uncertainty?", "reviewer_2": "Key Results:\nIt is difficult to parse apart the drivers of air-sea CO2 exchange (and by extension atmospheric CO2 variability), particularly in the Polar Antarctic Zone where carbonate archives are generally not present. The authors of this paper attempt to circumvent this challenge by inferring changes in quasi-conservative properties of deep ocean waters that are sourced from the PAZ; these properties are set by a combination of endmember values and water-mass mixing and should be recorded by benthic foraminifera at the study site. By the authors’ logic, changes in these properties should reflect a combination of changes in circulation and changes in endmember composition (in this case largely set in the PAZ and delivered to the study site).\n\nThe authors have divided the discussion of their results by timescale. On glacial/interglacial timescales, they find that the deep pacific ocean was a net source of CO2 to the atmosphere, given the increase in [CO32-]as at the site. The concurrent increase in PO4* suggests that there was a combination of more preformed PO43- (and thus lower nutrient efficiency) and/or more preformed O2 as a result of a combination of biology and the removal of sea ice. The latter is corroborated by the decrease in ventilation age documented at the site in a previous publication by these authors.\n\nOn millennial timescales, the authors suggest that the deep South Pacific was a source of CO2 outgassing across HS1, and that this was matched by changes in preformed nutrients (and by extension nutrient utilization efficiency) and/or preformed O2 inferred from PO4*. Radiocarbon ventilation ages suggest that this was a period of better ventilation, perhaps as a result of sea-ice retreat and/or a shift southwards of the southern hemisphere westerlies during this period. The Antarctic Cold Reversal shows a decline in [CO32-]as and PO4 from the end of HS1, consistent with poorer ventilation of the waters reaching this region. Lastly, the Younger Dryas measurements of [CO32-]as and PO4* suggest that this region was again a source of CO2 outgassing as a result of biological and physical changes in the PAZ. These mechanisms are consistent with to those provided for the full glacial/interglacial change in these parameters.\n\nOn sub-millennial scales, the authors examine and attempt to delineate the timing of changes in [CO32-]as and PO4* at this site across HS1 using ~12 data points (the sedimentation rate of this core is 10-20 cm/kyr depending on chosen tie points from Dai, Yu and Rafter 2021). The authors suggest that the asynchrony between [CO32-]as and PO4* in early HS1 allows for the discrimination between the various influences on these parameters. They suggest that the lack of change in PO4* is consistent with an increased role for physical drivers (retreating sea ice and altered southern hemisphere westerlies) in early HS1. In contrast, according to the authors interpretation, the change in PO4* across late HS1 is indicative that biological changes likely spurred mid-late HS1 outgassing.\n\nValidity:\nThe authors make use of two quasi-conservative (and related) tracers: PO4* and [CO32-]as. PO4* is derived using inferred concentrations of PO43- and O2. The former is derived from the measured Cd/Ca ratio of the benthic foraminfera H. elegans using a local calibration for Cd/Ca to PO43-. The latter is derived from the gradient in carbon isotopes of co-occurring infaunal (G. affinis) and epifaunal (C. wuellerstorfi) species. Changes in PO4* thus reflect some combination of changes in nutrient utilization (via changes in preformed PO43-) and changes in oxygen related to either biology and/or physics.\n\nThe PO43- measurements are also used to calculate [CO32-]as, a parameter which reflects changes in CO32- (and by extension DIC and atmospheric CO2) with biological influences removed. When this parameter changes, it reflects changes in carbonate ion as a result of CO2 sequestration or outgassing as a result of air-sea gas exchange; increases in [CO32-]as reflect lower DICas and outgassing of atmospheric CO2. Because increased exchange of CO2 should also alter PO4*, differences in the behavior of [CO32-]as and PO4* downcore should provide constraints on the role of physical processes (relative to biological processes) in governing CO2 release to the atmosphere across the deglaciation.\n\nIn general the data appear to be of high-quality and their interpretation appears robust. However, there are some major flaws that I think need to be fixed if the paper is accepted for publication. In particular, the discussion of sub-millennial drivers of CO2 variability in HS1 relies on the statement that [CO32-]as increased throughout early HS1, in contrast to PO4* (line 166- 167). It would be extremely helpful to the reader to delineate the exact magnitude of the change in these variables over the early HS1 time period. The change in [CO32-]as between 18- 16.3 ka appears to be on the order of 5 μmol/kg at most, and perhaps less than that if the lowest data point on the edge of HS1 is excluded. The data for PO4* looks as though it is also slightly increasing over the same time period, with the exception of a slight decrease right at 16.3 kyr (Figures 2 and 3). To make things more confusing, the onset of HS1 on Extended Data Figure 1 appears at a different place in the [CO32-]as and PO4* curves than in Figures 2 and 3. Because the figure is plotted against depth it makes interpretation and comparison difficult, but if Extended Data Figure 1 is used in interpretation it appears that [CO32-]as actually decreased between 18-16.3 kyr.\n\nIn other places in the text the authors refer to the change in N. Atlantic [CO32-]as as small across all of HS1 (lines 117-118), but the change in early HS1 in the NEAP-4K record is almost double that of the MD97-2106 core. How much of the small change in [CO32-]as in early HS1 could be from a change in NSW/SSW proportion and/or end member composition? Again this is unclear, but if the change in [CO32-]as is on the order of 5 μmol/kg, even a small change induced by circulation/water column structure change may represent a significant proportion of this inferred increase.\n\nIt is not convincing to me from Figures 2/3 or from Extended Figure 1 that the change in [CO32-]as is significant either relative to the lack of change in PO4* or its starting value at the beginning of HS1. Given that this is one of the 3 major findings of the paper and is used evidence for physical changes initiating CO2 outgassing, this needs to be addressed prior to publication. The other evidence cited (radiocarbon ventilation, sea ice extent, and δ15N) are convincing in this respect, but have all been published elsewhere.\n\nSignificance:\nThe authors’ text is well-cited with regard to deglacial CO2 release from the deep Southern Ocean, including mechanisms and modelling. The literature they cite shows that, over the last 5-10 years, there have been numerous results connecting the deep Southern Ocean to the release of carbon from the deep across the deglaciation. There is, to my knowledge, already a consensus that the Pacific would have been a major player in this process both on glacial/interglacial and millennial timescales. Where this article therefore would be adding to the literature, then, would be if it could distinguish the role of physical and biological changes in driving this release for the PAZ. However, for the glacial/interglacial and millennial timescales of this article, the authors acknowledge that the co-variance of their tracers does not allow them to fully delineate between the dual roles of physics and biology. They do attempt to constrain the amount of carbon release implied by their data for glacial/interglacial transition, but this extrapolation is not fully explained nor is the role of the deep South Pacific (relative to, for example, the Atlantic) constrained by this extrapolation.\n\nThe authors suggest that the behavior of the two conservative tracers over HS1 allows them to discriminate between influence of physical and biological tracers at its onset. To me this would be a significant advance in understanding provided by this paper, as it would serve to confirm evidence from radiocarbon and sea-ice proxies. However, the authors must provide more information regarding the changes in these two tracers across this time period to better demonstrate their conclusions (regarding whether the change in [CO32-]as is significant relative to the lack of change in PO4*).\n\nThe significance of this paper would clearer with discussion in more detail of how the results can or cannot discern between mechanisms or how they could build upon the various modelling results that they cite. Khatiwala et al. (2019), for example, find that sea ice extent and ocean circulation are minor contributors to changes in carbon storage relative to the role of air-sea equilibrium in amplifying changes in ocean temperature and iron fertilization. Cliff et al. (2021) also suggests a minor role for circulation and the biological pump in driving deep ocean deoxygenation (and thus PO4*), but a large role for sea ice extent and iron fertilization. Such contradictions might be cleared up from paleoclimate-based inferences, but is largely not discussed here.\n\nOne major thing that is unclear to me is why the authors do not discuss existing PAZ δ15N data. The authors suggest that the lack of carbonate archives from the PAZ preclude analysis of nutrient utilization in the region, but diatom archives (plotted in Extended Figure 5) provide such a record. Likewise, deep sea corals provide a coherent and chronologically well-constrained record of δ15N. There is no discussion of these diatom records throughout the text, and only a single mention of deep sea coral-bound δ15N. Recent important work (Ai et al., 2020) provides important constraints on the role of Southern Hemisphere westerly winds and upwelling on nutrient conditions in the PAZ and should be discussed here in reference to the conclusions being drawn, especially with regard to the interpretation of δ15N records from the PAZ.\n\nAnother discussion point that should be addressed are the ‘Late Holocene’ values marked in Figure 2. It is not entirely clear from the methods section what age the “Late Holocene values reflect (Lines 214-215), but from inference (the disappearance of G. affinis until 0-17cm in the core) it seems likely these reflect values core top values. The values of [CO32- ]as and PO4* are both equal to their value at 8-10 kyr. However, we know from Studer et al. (2018) that there was a substantial (20 ppm) rise in CO2 across the late Holocene (between 10 kyr and 2 kyr) that was accompanied by a large change in nitrate supply, reflected in the δ15N of Southern Ocean PAZ and Polar Frontal Zone diatom-bound N records, as well as in Coral-bound δ15N records from the PAZ. Presumably this change in δ15N should also be reflected in the authors’ records of PO4*. The changes in overturning driven by migration of Southern Hemisphere Westerly Winds (hypothesized by Studer et. al 2018) should also perhaps be reflected as changes in [CO32- ]as. Please discuss why your core top values relative to 10 kyr do or do not contradict existing data regarding this period.\n\nPrevious records from this same core (from the same authors in 2022) show a strong role for altered efficiency of the solubility pump over the same time period (inferred from data presented in Fig. 2A). No discussion of this in relation to the findings of this paper are made, yet the inference from the 2022 paper is that Southern Ocean outgassing as a result of an altered solubility pump makes up a significant proportion of the observed change between 19-15 kyr – but the current paper would suggest much of the action is taking place in the PAZ instead of the subantarctic.\n\nData and methodology:\nThe data appeared to be of high quality, and technically sound. I have minor comments that can be addressed line by line in the text. The main major comment I have is for readability: the authors here assume the reader has a detailed understanding regarding the calculation and interpretation of [CO32-]as, and rely on the reader to have a fairly detailed prior understanding of how this is calculated in order to follow the discussion. A short discussion of the tracer and its relationship to atmospheric CO2 and DIC would be helpful.\n\nMethods:\nThere is evidence that morphotypes of C. wuellerstorfi (e.g. sensu lato vs. sensu stricto after Gottschalk et al. 2016 and references therein) can alter the values of δ13C and trace element ratios. Given their expertise in the field, the authors are likely aware of this – were selections made by morphotype within C. wuellerstorfi that are not mentioned in the text? If not, could this account for some of the variability in, for example, B/Ca and δ13C that would give rise to artificial changes in reconstructed [CO32-]as and PO4* values? δ13C in particular can vary by between 0.2-0.7‰, which would be significant for reconstructions of bottom water oxygen concentrations.\n\nIn lines 222-223, there is mention of splits of large samples into subsamples, which were presumably run in different batches and/or treated as independent samples. What is the standard deviation of these “true” replicates? Are they averaged and included in the figures (e.g. Extended Figure 1) or plotted independently? The 5th oldest point in Extended Figure 1g, for example, looks as if it may be 2 replicate samples, but one is significantly different than the other – but it is hard to tell.\n\nWere the benthic foraminiferal tests for stable isotope analysis (lines 230-233) treated in any way (e.g. crushed and sonicated)? Why not use splits of C. wuellerstorfi from the trace element measurements?\n\nδ18O is plotted in Extended Figure 1E but there is no reference to it in the text, and no discussion of the methodology (species?) for the δ18O measurements (line 233 is the only mention). It appears δ18O largely agree, but there are places with offsets of up to ~0.25‰ in co- occurring benthics. This difference is disguised by the scale of glacial/interglacial change in δ18O – is this a normal difference in δ18O between these two species? If so, perhaps this should be discussed, and if not, there should be a rationalization for how these particular points can be used for co-occurring measurements of δ13C gradients.\n\nA table or a more organized discussion of the results of sensitivity tests would also be helpful for the reader – at the moment the discussion feels haphazard. The results of the sensitivity tests are presented parenthetically (Lines 91-93 for example), which makes it hard to follow. The methodology behind the sensitivity tests is straightforward and described in the text (lines 290 onwards) but the results of these tests as they are discussed in the text are not as clear.\n\nAnalytical approach:\nThe data appeared to be of high quality, and technically sound. I have minor comments that can be addressed line by line in the text. The main major comment I have is for readability: the authors here assume the reader has a detailed understanding regarding the calculation and interpretation of [CO32-]as, and rely on the reader to have a fairly detailed prior understanding of how this is calculated in order to follow the discussion. A short discussion of the tracer and its relationship to atmospheric CO2 and DIC would be helpful.\n\nMethods:\nThere is evidence that morphotypes of C. wuellerstorfi (e.g. sensu lato vs. sensu stricto after Gottschalk et al. 2016 and references therein) can alter the values of δ13C and trace element ratios. Given their expertise in the field, the authors are likely aware of this – were selections made by morphotype within C. wuellerstorfi that are not mentioned in the text? If not, could this account for some of the variability in, for example, B/Ca and δ13C that would give rise to artificial changes in reconstructed [CO32-]as and PO4* values? δ13C in particular can vary by between 0.2-0.7‰, which would be significant for reconstructions of bottom water oxygen concentrations.\n\nIn lines 222-223, there is mention of splits of large samples into subsamples, which were presumably run in different batches and/or treated as independent samples. What is the standard deviation of these “true” replicates? Are they averaged and included in the figures (e.g. Extended Figure 1) or plotted independently? The 5th oldest point in Extended Figure 1g, for example, looks as if it may be 2 replicate samples, but one is significantly different than the other – but it is hard to tell.\n\nWere the benthic foraminiferal tests for stable isotope analysis (lines 230-233) treated in any way (e.g. crushed and sonicated)? Why not use splits of C. wuellerstorfi from the trace element measurements?\n\nδ18O is plotted in Extended Figure 1E but there is no reference to it in the text, and no discussion of the methodology (species?) for the δ18O measurements (line 233 is the only mention). It appears δ18O largely agree, but there are places with offsets of up to ~0.25‰ in co- occurring benthics. This difference is disguised by the scale of glacial/interglacial change in δ18O – is this a normal difference in δ18O between these two species? If so, perhaps this should be discussed, and if not, there should be a rationalization for how these particular points can be used for co-occurring measurements of δ13C gradients.\n\nA table or a more organized discussion of the results of sensitivity tests would also be helpful for the reader – at the moment the discussion feels haphazard. The results of the sensitivity tests are presented parenthetically (Lines 91-93 for example), which makes it hard to follow. The methodology behind the sensitivity tests is straightforward and described in the text (lines 290 onwards) but the results of these tests as they are discussed in the text are not as clear. The results of the sensitivity tests are presented parenthetically (Lines 91-93 for example), which makes it hard to follow. The methodology behind the sensitivity tests is straightforward and described in the text (lines 290 onwards) but the results of these tests as they are discussed in the text are not as clear. More discussion of the sensitivity tests relative to the findings is necessary - for example, in lines 89-93 the authors discuss the range of sensitivities in SSW PO4* implied by potential changes in NSW PO4* - but neglect to then relate this finding back to their results (potentially because the order of magnitude of the sensitivity test is on par with the entire change in PO4* in SSW). If the sensitivity tests are included in discussion of the results (which they correctly are), they should also be discussed in greater detail. It is also unclear why the authors do not conduct sensitivity tests for other transitions (e.g. YD) where they are suggesting there must have necessarily been changes in the balance of NSW/SSW at the site.\n\nA specific point that I found confusing in the sensitivity tests is that, with their help, the authors define the increased change at the site from the change in NSW endmember to be 7 μmol/kg and the SSW endmember contribution to be 30 μmol/kg. However, adding these up gives a change larger than the glacial interglacial change suggested in lines 73-74 (32 μmol/kg +/- 4). Apologies if I have misunderstood, but these numbers do not seem to add up correctly, even accounting for the sensitivity test results.\n\nThere is no discussion of whether the core used here is of high enough sedimentation rate to analyze the sub-millennial scale variability that makes up the latter part of the paper, with the exception of lines 163-165 which references chronological uncertainties in Extended Data Figure 1 that are not evident in the figure or discussed in detail anywhere I can find. If there are significant uncertainties then they need to be described. Regardless, Extended Data Figure 1 does not match Figure 2 or figure 3 with regard to the onset of HS1 in PO4* or [CO32-]as - as an example, there are 4 data points prior to the gray bar in Extended Data Fig. 1i, and only 2 or 3 in Fig. 2g. Likewise, the upward trend in the oldest part of the PO4* record in Fig. 2F contradicts the data points in Extended Fig. 1h. Further confusing the analysis of this record, Extended Fig. 1 lacks the 16.3 kyr dashed line, making it difficult to analyze how many points constitute the trend the authors are inferring exists for [CO32- ]as and not for PO4* between 18- 16.3 kyr.\n\nThe extrapolation of carbon release from the glacial-interglacial transition needs to be made clearer and discussed further, or removed. What is this amount relative, to say, the Atlantic or Indian – and if the amount released is comparable in size to the atmosphere then does this imply that the majority of the C release was accomplished through a reorganization of the deep Pacific alone? What was the role of the intermediate waters in this reorganization? Lines 110-113 are not clear to me and need further clarification.\n\nClarity and context:\nThere are several places where the writing is unclear and obscures the logic or argument. The authors’ use of the phrase “direct evidence” (lines 83-85) needs to be substantiated or made more clear - why is this “direct evidence” compared to the numerous other records of Southern Ocean deglacial carbon loss? These records also represent inferences, given that they are not from the PAZ and rely on inferred relationships between trace metal/calcium ratios, bottom water chemistry, pore water chemistry, etc, as well as sensitivity tests and assumptions regarding the characteristics of NSW, SSW, and their proportions. While the authors of existing δ15N records would not necessarily say their records record “direct” CO2 outgassing from the AZ, they might be inclined to argue that diatom records from the PAZ are a much more direct record of nutrient utilization efficiency than PO4* inferred here. Why is B-P radiocarbon from the site is not “direct” evidence (lines 103-106, for example)?\n\nText, minor comments:\nLine 37: Key to this paper is the idea that these properties are conserved, adding “along flow paths moving away from the PAZ” might be clearer.\nLines 46-48 are confusing: PO4* is noted to be largely governed by surface nutrient conditions here, but in the methods explanation has a solubility/physical component (lines 276-289). The statement in lines 48-49 about deconvolving these tracers makes it sound much more straightforward than the methods imply, which is shown by the fact that their deconvolution can not be achieved for glacial-interglacial or millennial timescales.\nLine 71: There are no calculation details for [CO32-]as.\nLines 75-76: Changes in SSW/NSW is vague wording: do you mean changes in the proportion of these or in the endmember composition?\nLines 76-79: This is confusing – if the proportion of NSW is increasing across the deglaciation, changes in NSW characteristics would have a larger impact as the proportion increased.\nLines 87: “Changes” could refer to proportion or endmember again.\nLines 107-110: “Data-based evidence” is a strong way to phrase this, given that most of the other studies cited (including those from this site) are also Data-based.\nLine 118: This refers to extended data fig. 4, I believe, not certainly not 5. Line 119: decreased fraction of NWS during HS1 needs a citation.\nLine 121-122: citation needed for this\nLine 136-138: Are your results adding to these conclusions listed? A reader might have arrived at the same conclusions (Driven by biological and physical changes) from the δ15N, radiocarbon, and sea salt records alone. The paper needs to be more specific about the constraints that this data is providing.\nLines 156-159: Again, as noted immediately above, it remains unclear what additional constraints are provided given the ambiguous conclusion. The authors would be well-served to make their specific conclusions stronger and/or clearer, with regard to how these add to the existing literature.\nLines 166-167: This does not appear to be true, at least not significantly. This needs to be addressed with the actual data, and concentrations/values need to be given.\nLines 169-170: 0.1 represents 1/3 of the Glacial-Interglacial transition, so it is not as small as the authors imply here.\nLines 186-187: Callout to figure 3 would be helpful here.\nLines 210-212: The finding that the biological pump plays a large role is not surprising, but could be discussed further relative to recent papers finding a large role for the solubility pump and other physical mechanisms, particularly for intermediate Southern Ocean waters.\n\nFigures, minor comments:\nCaption Fig. 1 has the words “Lower Circumpolar Deep Water” but no explanation of what this is in reference to with regard to the rest of the sentence (Lines 323-324).\nFig. 2: The late Holocene values for Figure 2 [CO32-]as are not at all close to the values in Figure 1D. While some of the absolute value change is due to calibration, the values are significantly higher than those given for the modern ocean. Why is this the case, if the PO4values are accurately reflected in the core tops? Why does Fig. 2F increase at the oldest point in the record without a data point on the plot that is older than ~18 kyr?\nFig. 3: It is not clear what the gradient in the color bars on the triangles in b,c,e, f and h are reflecting given the depiction of timeseries. Also, should Fig. 3D say B-P and not B-A age?\nExtended Fig. 1: This needs substantial improvements as it does not align with Figure 2 or 3, and does not have a marker for 16.3 kyr.\nExtended Fig. 2: Typo for a instead of (a) in line 567.\nExtended Fig. 3: Needs a better legend, given that there are actually 3 sets of data on the plot and not 2. This figure is not discussed in the main text. For example, hat does the substantial gradient (20 μmol/kg) between MD97-2106 and the S. Indian Core in reconstructed O2 imply for the deep South Pacific vs Indian Ocean during early HS1?\nExtended Fig. 4: Is there a reason that the Yu 2022 data for the South Atlantic is not also included here? I assume the reasoning is to show the endmembers for NSW and SSW, but the GGC90 record was used by Yu et al to infer significant release of carbon from the Atlantic sector of the Southern Ocean as well.\nExtended Fig. 5: It would be helpful to discuss Fig. 5C in the text. Also, putting the late Holocene values for the δ15N records and for PO4* would be helpful for any discussion of the late Holocene evolution of these tracers.\n\nReferences:\nThe paper is well-cited, though not many of the citations are used to further the discussion of the text (for example, δ15N references) despite being in the text. These citations may also be of interest for the discussion:\nRonge, Thomas A., et al. \"Southern Ocean contribution to both steps in deglacial atmospheric CO2 rise.\" Scientific reports 11.1 (2021): 22117.\nMoy, Andrew D., et al. \"Varied contribution of the Southern Ocean to deglacial atmospheric CO2 rise.\" Nature Geoscience 12.12 (2019): 1006-1011.\nFogwill, C. J., et al. \"Southern Ocean carbon sink enhanced by sea-ice feedbacks at the Antarctic Cold Reversal.\" Nature Geoscience 13.7 (2020): 489-497.", "reviewer_3": "Dai and Yu present a valuable dataset, interpreted in a cutting edge quantitative framework, to investigate glacial carbon storage and release. This is a really interesting contribution and following revisions would, I think, make a very suitable Nature Communications paper.\n\nI have provided detailed line by line comments, which include key conceptual questions as well as more minor points and typos. Note there are other typos and sections lacking clarity too - the paper would benefit from further proof reading and an attempt to more clearly describe processes to the reader alongside their quantification in different tracer signals. I also pick out some of the most critical issues below.\n\n*[CO32-]as - better describe what it can and can’t tell us*\nThe [CO32-]as tracer has been introduced and used in a series of recent papers from Yu’s group, yet is still relatively new and (as detailed below) needs further description to help the reader.\n\nMore critically, the interpretation I felt leaps in to this relatively derived quantity quite quickly, and would benefit from spelling out some of the signals in the raw data, and how these are brought together into [CO32-]as and other related properties. In particular, I was left wondering about the other components of ocean carbon storage (e.g. remineralised carbon) and their fingerprints on the carbonate chemistry and other reconstructions here. Can we learn something about remineralised carbon storage from these data or not? This is particularly important as the topic is currently in vogue (see e.g. the Volmer et al., 2023 paper in Paleo) and as this question naturally arises from several points in the text. For instance right from the abstract, line 19 highlights “carbon losses from the Deep South Pacific to the atmosphere”; this to many readers will I think give an image of release of a deep carbon reservoir which others have suggested has a large remineralised carbon component. But the main text description and discussion of the signals is very much focussed on preformed signals in the Southern Ocean surface. I can see ways to bring these different ends of the glacial carbon storage system together, but it would I think be very helpful for the authors to spell this out more explicitly, both by describing the tracer in more detail and its relationship with other ways of thinking about deep carbon storage and release (e.g. remineralised vs preformed carbon) in the introduction and the initial sections describing deglacial carbon loss from the deep ocean.\n\n*Which water are we really feeling?*\nThe authors claim that their record is indicative of processes in the PAZ. However this needs to be better substantiated for a couple of reasons. While I agree with the authors that the core will be dominantly influenced by SSW, not all SSW is necessarily sourced from the PAZ. For instance in the Gebbie and Huybers (2010) deconvolution of deep water sources, this core site would have about 50% water from the PAZ, with additional, subequal contributions from the PFZ/SAZ and NSW (the NSW component is indeed nicely picked up in the low PO4* blob in the section figure).\n\n[The proportion of SSW at this site and how this is tracked by PO4* is, in addition, strongly influenced by the role of entrainment in the SSW PO4* end member signature, discussed further below.] Related to the major comment above about preformed vs remineralised signatures, I’m also left wondering how I should think about timescales of change at this site compared to the atmosphere, given that it will have influences both from relatively recently ventilated water from the south alongside return flow from the Pacific. Should I think of this as being in water that has recently left the PAZ so quickly picks up changes sourced from PAZ-atmosphere interactions? Or does it also have a component of return flow from the Pacific, rich in remineralised carbon, having left its surface source regions perhaps several thousand years earlier? A mixture of these to me seems likely and I think some discussion of this, alongside the different carbon signatures being felt here as discussed above, would benefit the paper.\n\n*Structure of paper / structure of millennial vs HS1 changes*\nI read through the section on millennial changes wondering why the earlier and later parts of HS1, which clearly have very different signatures, were being lumped together. I then realised that they are discussed in more detail in the next section on Bio vs Physical carbon loss in HS1. However I think this current structure risks confusing readers. I’d suggest at least flagging in the millennial section that there are two distinct parts of HS1 that will be discussed further below.\n\nI also suggest using the general/orbital section to better discuss the tracers and the nature of carbon loss, addressing the major comments above.\n\n*Early HS1 [CO32-]as change?*\nThe text discusses there being a [CO32-]as change in both the early and late part of HS1, and from reading the text alone without looking at the figures one might assume these [CO32-]as changes are similar in size. However from looking at the figure, it’s unclear whether there really is a significant [CO32-]as change in early HS1, whereas the late HS1 change is substantial. Modify the text to address this directly - is there perhaps even something interesting one can say about the different sizes of these changes are their relationship/slope compared to other tracers that may further point to mechanisms?\n\n*Plotting error - is there an interesting centennial signal?*\nThe relative phasing of the data in Figure 2 and ED Figure 1 is not the same (and I realise that one is in age space, the other in depth - the issue lies in the phasing of proxy signals within the same core). In ED Figure 1 the notable spike to low [PO4] at 150 cm is simultaneous with a sharp jump up in [CO32-]as. Taken together these signatures are really interesting in terms of mechanism - and are also particularly interesting given some of the very rapid changes in deep sea coral data around this time (e.g. Li et al., 2020). However in Figure 2, the sharp drop in [PO4] at 16.3 ka has no accompanying feature in [CO32-]as - the jump up comes 500 years or so later.\n\nIf the jump up in [CO32-]as is associated with a drop in [PO4], then this does have implications for mechanisms and should be addressed. Specifically, at present the late HS1 rise in [CO32-]as is discussed as being associated with rising PO4, leading to the implication that this outgassing change comes about through reduced biological pump efficiency. However if this major jump up in [CO32-]as (which is about half the overall late HS1 change) is in detail associated with a drop in [PO4], then this surely challenges that interpretation, at least within this interval? Rather, it might fit better with the coral 14C younging that happen at this time, indicating a pulse of pronounced ventilation that outgasses CO2 and flushes well-ventilated, low nutrient waters through the water column.\n\nLine by line comments\n43-46 - [CO32-]as is the main focus of this paper and is still a relatively new tracer. I think the paper (and certainly its readers!) would benefit from some more description of this tracer here and also in the methods. This doesn’t need to be exhaustive (you can reference the previous work for more details), but it should be enough for the reader to get a feeling how this is calculated and what it indicates in more detail than currently provided. It should also discuss how DIC is calculated from this tracer, which appears out of the blue later in the paper, and would similarly benefit from introduction.\n73 - jumps straight in to very derived property/concept. My preference would be to step through the glacial state you have first and the different components to the carbon storage (e.g. both preformed and remineralised), then how they evolve over deglaciation.\n78 - unclear as written - an increase in Northern water, which is increasing its [CO32]as, sounds like it could contribute to the signal of increasing [CO32]as. Key missing logical step is I think the different [CO32]as signatures of NSW vs SSW: as NSW has lower absolute [CO32]as, this may be expected to dominate. I realise this is what the mass balance calculation in the sentence that follows is meant to spell out, but it’s not very clear for a non-expert. Add a short clause of descriptive explanation here and a fuller explanation to Figure ED4 caption, which is referenced here, but doesn’t have this level of explanation in the caption.\n81 - the NSW contribute to - unclear writing.\n82/83 - this seems like it would also benefit from being shown in a figure - perhaps an additional panel to Figure ED4. Also what does the PO4* suggest about water masses?\n85 - a key point (also mentioned above) that is not discussed in enough detail is whether we can explicitly link changes at this site to the PAZ, vs other regions of the Southern Ocean. The core is not in a region of pure PAZ-sourced AABW and likely also has a component of water ventilated along sloping isopycnals in the PFZ/SAZ (as well as the NSW which is discussed).\n86 - raised —> rose\n90 - being —> is\n92 - if of varying\n91-93 - again I don’t find this list of numbers particularly helpful in getting a sense for how big a deal this is - would much rather see in a figure or a table.\n94 - “the latter of which” is odd here - would seem like it would only be referring to O2, but think you mean PO4 and O2?\n110-111 - where do these numbers come from?! I understand how you calculate [CO32-]as, but not how you then go to (presumably) DIC to get total Pg carbon.\nAlso, critically, if you are going to extend the calculation to the whole deep Pacific, you really need to say something about the other proxy records in this basin (especially DCO32-, PO4, O2) - are they consistent with the signals here allowing you to extrapolate to the Pacific below 1500m?\n112 - is this really sizeable? Pacific below 1500 m is likely main deep ocean carbon store. If need to account for at least another 400 PG and maybe as much as 800 or 1200, then where does the rest of the carbon come from?\n115-129 - “throughout HS1” - actually the big change in [CO32-]as and PO4* is late HS1.\n130-139 - linked but conversely to the point above, the 14C change happens entirely during early HS1, and is associated with no change in [CO32-]as! Seems like this is trying to tell us something. Alternatively, if the authors don’t think the age model is robust to this, need to be up front about it. What do the benthic-planktic data look like?\n143 - Stewart et al. (2021) paper also nicely shows this and reaches similar conclusions.\n166 - only significant change seems late HS1. Certainly a very different magnitude / slope of change if there is any change at all, and can’t just brush over this.\n168-169 - this sentence is very hard to follow\n171 - given what you say about sea ice I think best to omit “likely stay unchanged” - if your sea ice and 14C interpretation is correct, then it likely increases.\n174 - if you are really in PAZ waters, then how influential is any likely change in temperature? Probably stays pretty cold!\n176-177 - this is really not clear to me and I think will be completely opaque to majority of readers. Rephrase to make more clear what you mean. I’m left unsure whether these are expected to drive a change in PO4* or not.\n180 - very unlikely that no change in physical O2 recharge if there is a change in 14C, as 14C takes much longer to equilibrate than O2.\n186 - this younging in 14C occurs notably earlier than coral 14C data: with the exception of the transient younging pulse at 16.1 ka, the main younging is all in late HS1. This is also true of most of the younging in the Skinner record. Why the discrepancy?\n191 - This is nicely supported by Gray et al. 2023 reconstruction.\n195 - “14C further declined” - this makes it sound like it’s comparable to the much larger earlier change. As with the points made above on the magnitude of [CO32-]as change in the different parts of HS1, you need to be up front about the different magnitudes of these signals.\n198-199 - hard to follow what is meant here. Even if more details are provided in the Yu et al. 2019 paper, the reader shouldn’t have to go to this to get a general sense for how this calculation works.\nAlso if this calculation is done for DIC in late HS1, what is the equivalent calculation for DIC in early HS1? I would think minimal as there’s to my eye very little [CO32-]as change, but the text discusses [CO32-]as change in early HS1, so maybe there’s something I’m missing? Having the equivalent calculation for both may be helpful.\n200-201 - this might not be minor if there was a major change in sea ice\n208-210 - interesting and important point but unclear as written\n229 - 1 sigma or 2 sigma?\n261-262 - rephrase - this isn’t considering *all* the uncertainties as it doesn’t take in the calibration uncertainty, which is more like ±0.2 umol/kg.\n268 - need to at least give a brief description\n269 - similar to point above, but even more importantly, you really need to give more information here (and ideally also in the main text). [CO32-]as is totally central to this paper and is still a relatively new tracer, being only previously used in a couple of other papers from this group. While I realise that thinking about [CO32-]as probably now seems very natural to Yu and co-workers, it is still new enough that the rest of the community will continue to need some explanation - especially when it’s as key to a paper as it is here. Indeed I think the paper would benefit from additional description in the main text too to help build intuition.\n271-275 - I agree in principle with the correction for global alkalinity change and it’s very helpful to see figure ED1 that shows the [CO32-]as signatures both with and without this correction. However this figure (ED1i) also makes it apparent that the vast majority of the glacial-interglacial change in [CO32-]as from the oldest to the youngest samples comes about from this correction. This really needs to be discussed explicitly in the main text section on orbital scale changes.\n286-287 - why does air sea gas exchange cause reduced nutrient utilisation?\n303 - this is highly dependent on what end member is used for the SSW - the AABW signature in the source region vs the signature that incorporates entrained CDW. Rae and Broecker (2018) show this changes the SSW:NSW fraction in the deep Indo-Pacific between 50:50 and 75:25 respectively. Some discussion of this is needed, as it also has implications for what you are reconstructing in terms of SSW end member. Similarly, the influence of entrainment on the PO4* signature of the Southern end member should also be mentioned - this has an important influence today, is notably different in different regions of the Southern Ocean, and may have also changed in the past. These are, I realise, not easy things to constrain, but it needs at least some discussion.\n323-324 - LCDW - not really a sentence.\nFigure 2\n328—329 - if crosses really indicate age ties then these look unconvincing. Are the crosses 14C control points and the triangles the ties?\n331 - is the PO4 data also really including calibration error? Based on the different choices possible in Figure ED2, I would expect calibration error to be larger?\nFigure ED1\n- Please show benthic-planktic 14C data here too", "reviewer_1_attached": "We thank all three reviewers for their very detailed and constructive comments. Below, we\nprovide our point-to-point responses to the reviewers’ comments.\nReviewer #1 (Remarks to the Author):\nReview of Dai and Yu, “Deglacial polar Southern Ocean carbon release driven by physical\nand biological dynamics,” submitted to Nature Communications, by Jesse Farmer\nDai and Yu present new deep ocean B/Ca-derived carbonate ion, Cd/Ca-derived phosphate\nconcentration, and d13C gradient-derived oxygen reconstructions from a sediment core in the\nSouthwest Pacific Ocean tracking (largely) the properties of Southern Sourced deep waters\nover the last deglaciation. The authors translate this proxy data into information on changes\nto preformed phosphate and air-sea carbonate ion change to track the influences of Southern\nOcean biology and gas exchange on carbon dioxide partitioning between the ocean and\natmosphere over the last deglaciation.\nThere is a lot to like in this manuscript. It is overall well written, the figures are excellent, and\nthe combination of proxies applied is novel and potentially quite transformative. At the same\ntime, I feel that the authors are trying to reach a conclusion on the role of sea ice cover in\ndeglacial ocean-atmosphere CO2 exchange before they have fully “paved the road” of proxy\nunderstanding that would allow the reader to get on board with this interpretation and exclude\nother alternatives. For this reason, I think the results would actually be more influential as a\nlonger contribution to a field-specific journal where the proxy interpretations, and, most\ncrucially, the covariance between these new proxies (see below) can be more fully explored.\nIt may even be beneficial to split this into two manuscripts, one longer-form outlining the\ncombined proxy approach and its uncertainties, and the second a shorter-form application to\nthe deglaciation with its climate system implications.\nWe thank the reviewer for the positive comments, and appreciation of the amount of work put\ninto this study. Following the reviewer’s comments, we have expanded discussions on original\nproxy records, contributions of original proxies to the derived tracers, and some alternative\nexplanations for our data that we dismissed. These will follow the reviewer’s comments.\nMajor comments (not in order of importance):\n1. Age model development.\nThe C. wuellerstorfi d18O (ED Fig. 1) does not lend confidence in the age model construction\nthat is based on G. bulloides Mg/Ca correlation to EDC dD (Dai et al., 2020). This is of crucial\nimportance for the sea ice interpretation of outgassing in early HS1 – if the timing of HS1 is\nnot correctly assigned in this sediment core, then the relationship of PO4* to sea ice (ice core\nss Na) is uncertain. Unfortunately, I don’t find the chronological establishment here sufficiently\ncompelling/precise as to allow for the discussion on L166 and onward, which is the crux of the\ncurrent paper. Perhaps the published radiocarbon evidence can be brought in here, or other\nSST proxies can be measured to refine this alignment?\nWe agree with the reviewer that a robust age model is the foundation to discuss changes\nduring HS1. Our age model enables us to discuss these changes within our records from the\nsame sediment core and between our records and other records. This point, which was not\nproperly emphasized in the previous version, is now made clear in the main text (Line 92). Our\nage model during HS1 is based on planktic Mg/Ca-EDC δD correlation in (Dai et al., 2021),\nas well as 4 radiocarbon dates during HS1 reported in (Dai et al., 2022)(2 in the early HS1\nand 2 in the late HS1). These age constraints allow us to resolve millennial timescale changes,\nincluding distinguishing the early and late HS1, despite uncertainties in absolute ages.\nTo demonstrate the robustness of our age model, we have plotted our data based on an\nalternative age model discarding the planktic Mg/Ca-based tie point. Similar patterns of the\nrecords to that shown in the main text, especially within HS1, prove that our age model is\nrobust in discussing these features. (Fig. S3).\nThe second point raised by the reviewer is the slightly delayed benthic δ18O decline during\nHS1. This appears to be a common pattern that has been observed in the region previously.\nIn Fig. X1, it is shown that our data is consistent with another well-dated site nearby, which\nfurther lends support to our age model.\nFig X1. benthic δ18O at site MD97-2106 compared to two other SW Pacific sites in (Sikes et\nal., 2016).\n2. Sea ice, or not? The authors take a strong position on the role of sea ice in the abstract\n“…physical dynamics involving sea ice retreats (sic) dominated CO2 outgassing…” (L22) but\nare rightly more circumspect about this in the main text “However, relative roles of biological\nand physical processes to the PAZ outgassing remain ambiguous on both millennial\nand glacial-interglacial timescales…” (L157-158). More clarity on the relative contributions of\nsea ice and their timing would be beneficial.\nWe clarify the contribution of the gas exchange disequilibrium imposed by sea ice in the\nabstract (Line 22) and the main text (e.g., Line 255). In short, our PO * and [CO 2-] changes\n4 3 as\nshow that sea-ice related gas exchange disequilibrium contributed to deglacial CO\n2\noutgassing during the YD and minimally contributed to HS1 CO outgassing.\n2\n3. Tracer interdependence. PO4* and [CO32-]as are not independent tracers, as both depend\non Cd/Ca-derived [PO43-] (L48-49, 66-69). That suggests the difference between the two is\nlargely driven by carbonate ion and oxygen… but it is not at all clear what to make of the\ndistinctions between the two tracers. On the one hand, it’s interesting that the two show slightly\noffset patterns in time… but on the other, it’s not surprising that the two look similar (Fig. 2f,\ng). I would appreciate more guidance from the authors on what to make of this\ninterdependence.\nThe reviewer’s concern is valid in that Cd/Ca-derived [PO 3-] contributed to both PO * and\n4 4\n[CO 2-] . In the revised version, we add a new figure (Fig. 3) to show contributions of\n3 as\nunderlying proxies to PO * and [CO 2-] . As shown in this figure, [PO 3-] is a minor contributor\n4 3 as 4\nto millennial-timescale variabilities in PO * and [CO 2-] .\n4 3 as\n4. NSW and SSW endmembers + contributions over time. The authors note that the advantage\nof PO4* and [CO32-]as is that both tracers are quasi-conservative in the ocean interior (e.g.,\nL50-60). However, the downside of this is that these proxies are set by the endmember values\nof NSW and SSW and their relative contributions to the core site. It is known from the\nauthors’ previous work that these endmember values are not stable over time (e.g., Yu et al., 2019 for\n[CO32-]as in the N. Atlantic); indeed, the goal of this study is essentially to reconstruct the\nSSW endmember PO4* and [CO32-]as over the last deglaciation (e.g., L63-65). To this end,\nit is critical that the authors know the contributions of the relative endmembers. In my mind,\ntheir approach of assuming a percentage contribution of SSW and allowing it to vary (e.g.,\nL290) leaves much to be desired, especially as proxy data (Nd isotopes) may help address\nthis question for them.\nIn the new version, we focus on discussing SSW endmember changes based on our best\nestimation of SSW% and NSW endmembers determined from literature data (including Nd\nisotope (εNd) data for SSW%, previous North Atlantic records for NSW endmembers).\nSensitivity tests (now moved to the supplementary materials, Fig. S8) support estimations\nbased on our assumptions. This is because our site was dominantly affected by SSWs over\nthe last deglaciation (with SSW% over 70%), any variability of the SSW% at our site during\nthe last deglaciation would be expected to be small.\nOur assumptions about SSW% at our site are based on εNd data in the region. However, we\ndo not directly use εNd data to infer SSW% changes in the SW Pacific. This is because εNd\nhas different endmembers (the North Atlantic and the North Pacific) from PO * and [CO 2-]\n4 3 as\n(the North Atlantic and the Southern Ocean).\n5. Presentation of results. The authors do not discuss their underlying proxy reconstructions,\ninstead they move directly to the derived tracer products (PO4* and [CO32-]as). This is\nunfortunate – the most directly calculated results should be presented first, and only after that\na discussion of the interpreted derived products. Such additional length would allow for further\ninvestigation of intriguing features in the proxy data – for instance the offsetting (in both depth\n+ age space) of carbonate ion, oxygen, and phosphate concentration deglacial\nmaxima/minima (Fig. 2c-e). Might this relate to the relative depth habitat of the foraminifer\nused? (C. wuellerstorfi being ~sediment/water interface, H. elegans perhaps being infaunal,\nG. affinis being deep infaunal)?\nIn this work, we aim to focusing on interpreting derived tracers, which have the advantage of\navoiding some alternative explanations related to local processes for the most directly\ncalculated proxies. However, the reviewer’s point that in the previous version we moved on to\ndiscussing PO * and [CO 2-] too quickly is valid.\n4 3 as\nTo address the reviewer’s concern, we now expand background of PO * and [CO 2-] in the\n4 3 as\nintroduction, and use distribution of deep-water [CO 2-], [PO 3-], [O ], PO * and [CO 2-] in the\n3 4 2 4 3 as\npreindustrial ocean to better illustrate the strength of PO * and [CO 2-] (Fig. 1, S1, Line 69).\n4 3 as\nIn this version, we also started our discussion from directly reconstructions (Line 98). We avoid\ndirectly interpreting their implications for carbon cycle changes but explain the contribution of\ntheir variabilities to the derived tracers aided by Fig. 3.\nWe discount the varying depth habitats of employed benthic foraminifera as the driver of the\nobserved changes in the supplementary material (Line 385). The possibility that some of the\nsubtle offsets between the proxies are caused by varying habitats of benthic foraminifera\nemployed for proxy reconstructions cannot be fully ruled out. These features are however not\nthe focus of our interpretation. We focus on millennial-timescale changes in our core (including\nthe different changes during the early and late HS1), which are not affected by this effect, due\nto bioturbation and relative high sedimentation rate. This is can be supported by similar δ18O\npatterns of the epifaunal species C. wuellerstorfi and the deep infaunal species G. affinis (Fig.\nS2e) (H. elegans being epifaunal to shallow-infaunal (within top 0.5 cm), thus between these\ntwo species in the habitat depth profile).\n6. Uncertainty on PO4* and [CO32-]as. There does not appear to be an effort to estimate\nerrors on PO4* and [CO32-]as aside from those relating to the individual proxy reconstruction\nof carbonate ion, oxygen concentration, and phosphate concentration. At the very least, the\nerrors of each required measurement will add in quadrature; in addition, further error will\nresult from assumption of Redfield stoichiometry (see modern ocean uncertainties in these\nratios from Anderson & Sarmiento, 1994). A more comprehensive handling of uncertainty in these\nnew tracers will be needed than is presented in the main text, figures and methods.\nThe method for propagating errors using an iterated LOESS smoothing method for individual\nproxies and the derived tracers are expanded (Lines 320, 339, 377). We also fixed the mistake\nassociated with the [PO 3-] estimation spotted by the reviewer (see reply to minor method\n4\ncomments) and updated related error envelopes in figures.\nConsistent with previous works on proxies (e.g., (Lynch-Stieglitz et al., 2024)) and modelling\n(e.g., (Khatiwala et al., 2019)), we adopt a constant set of Redfield ratios to derive tracers and\ninfer DIC changes. However, our conclusions are minorly influenced by variable Redfield ratios\nin the ocean, as shown by additional sensitivity tests in Figs. S6, S7. In these tests, we\ncalculated PO * using varying R (160, 175, 180), and calculated [CO 2-] using varying\n4 -O2:P 3 as\nsensitivities of [CO 2-] to [PO 3-] changes with varying R (127 or 140) and R (4 or\n3 as 4 C:P Corg:Ccarb\n8) defined in Yu et al. 2019. It is noted that, as shown in Yu et al 2019 (Shown here as Fig\nX2), the sensitivity of [CO 2-] to [PO 3-] is relatively insensitive to changes in R and\n3 as 4 C:P\nR at the low [CO 2-] end ([CO 2-] is similar to [CO 2-] without correcting for the\nCorg:Ccarb 3 norm 3 norm 3 as\nconstant offset of -78), where our data points are.\nFig. X2. Sensitivities of [CO 2-] changes to [PO 3-] with varying R and R .\n3 norm 4 C:P Corg:Ccarb\nMinor methods comments.\nL236. How was anthropogenic carbon calculated and deducted? I would not expect much\ndetectible DICant at this water depth…\nAnthropogenic DIC is estimated from (Lauvset et al., 2016). It contributes negligibly to DIC in\nthe deep SW Pacific indeed, but it does substantially contribute to the DIC in the N Atlantic.\nCorrecting anthropogenic DIC is critical to reveal linear SSW-NSW mixing trend of [CO 2-] in\n3 as\nthe ocean before the industrial era.\nL262. Given all those contributions, the estimated uncertainty on phosphate concentration is\nonly ±0.05 µmol/kg? Come on now, that is not realistic. For one point of comparison, the Cd/Ca\nmeasurement precision alone in Farmer et al. (2019) corresponded to a similar or larger\nphosphate concentration uncertainty. I have no doubt the authors are better analysts than\nmyself, but the propagation of error should be handled appropriately.\nThis was a mistake, which is now corrected (~0.11 µmol/kg, ~0.05 of the estimated [PO 3-],\ndetails about how the error of [PO 3-] is derived (Line 320).\nsignificant (ED Fig. 1i), and also there is no reason to expect that a single core site would\nexperience the global ocean alkalinity change (see, e.g., regional deviations in benthic d18O\nfor whole-ocean d18O change, Lisiecki & Stern, 2016 and references therein). Can the authors\nplease further elaborate on/justify this correction and illustrate its uncertainty?\nBoth data and models suggest that the average ocean alkalinity was substantially higher\nduring the LGM than the Holocene. The effect of global alkalinity change on [CO 2-] is\n3 as\nsignificant, and correcting for it helps to better quantify [CO 2-] in the past. Although regional\n3 as\nalkalinity changes would deviate from the global trend to certain degree, the magnitude of\nchange would be similar on the orbital timescale (also the case for δ18O).\nThe reviewer’s concern about the uncertainty of the global alkalinity change is valid, as it could\nbe the driving factor for some subtle feature in our record. For this reason, we no longer argue\nfor an early-HS1 [CO 2-] increase, which is dependent on the alkalinity correction and is not\n3 as\nstatistically significant compared to the updated [CO 2-] reconstruction uncertainty. In\n3 as\ncomparison, the discussed changes from the late HS1 onwards are independent of the\nalkalinity correction (Fig. 3). We also use three different model outputs to calculate [CO 2-]\n3 as\n(Fig. S7). Although the magnitudes of [CO 2-] changes vary, the millennial-timescale [CO 2-]\n3 as 3 as\nfeatures are consistent (Fig. S7)."}, "author_response": "We thank all three reviewers for their very detailed and constructive comments. Below, we\nprovide our point-to-point responses to the reviewers’ comments.\nReviewer #1 (Remarks to the Author):\nReview of Dai and Yu, “Deglacial polar Southern Ocean carbon release driven by physical\nand biological dynamics,” submitted to Nature Communications, by Jesse Farmer\nDai and Yu present new deep ocean B/Ca-derived carbonate ion, Cd/Ca-derived phosphate\nconcentration, and d13C gradient-derived oxygen reconstructions from a sediment core in the\nSouthwest Pacific Ocean tracking (largely) the properties of Southern Sourced deep waters\nover the last deglaciation. The authors translate this proxy data into information on changes\nto preformed phosphate and air-sea carbonate ion change to track the influences of Southern\nOcean biology and gas exchange on carbon dioxide partitioning between the ocean and\natmosphere over the last deglaciation.\nThere is a lot to like in this manuscript. It is overall well written, the figures are excellent, and\nthe combination of proxies applied is novel and potentially quite transformative. At the same\ntime, I feel that the authors are trying to reach a conclusion on the role of sea ice cover in\ndeglacial ocean-atmosphere CO2 exchange before they have fully “paved the road” of proxy\nunderstanding that would allow the reader to get on board with this interpretation and exclude\nother alternatives. For this reason, I think the results would actually be more influential as a\nlonger contribution to a field-specific journal where the proxy interpretations, and, most\ncrucially, the covariance between these new proxies (see below) can be more fully explored.\nIt may even be beneficial to split this into two manuscripts, one longer-form outlining the\ncombined proxy approach and its uncertainties, and the second a shorter-form application to\nthe deglaciation with its climate system implications.\nWe thank the reviewer for the positive comments, and appreciation of the amount of work put\ninto this study. Following the reviewer’s comments, we have expanded discussions on original\nproxy records, contributions of original proxies to the derived tracers, and some alternative\nexplanations for our data that we dismissed. These will follow the reviewer’s comments.\nMajor comments (not in order of importance):\n1. Age model development.\nThe C. wuellerstorfi d18O (ED Fig. 1) does not lend confidence in the age model construction\nthat is based on G. bulloides Mg/Ca correlation to EDC dD (Dai et al., 2020). This is of crucial\nimportance for the sea ice interpretation of outgassing in early HS1 – if the timing of HS1 is\nnot correctly assigned in this sediment core, then the relationship of PO4* to sea ice (ice core\nss Na) is uncertain. Unfortunately, I don’t find the chronological establishment here sufficiently\ncompelling/precise as to allow for the discussion on L166 and onward, which is the crux of the\ncurrent paper. Perhaps the published radiocarbon evidence can be brought in here, or other\nSST proxies can be measured to refine this alignment?\nWe agree with the reviewer that a robust age model is the foundation to discuss changes\nduring HS1. Our age model enables us to discuss these changes within our records from the\nsame sediment core and between our records and other records. This point, which was not\nproperly emphasized in the previous version, is now made clear in the main text (Line 92). Our\nage model during HS1 is based on planktic Mg/Ca-EDC δD correlation in (Dai et al., 2021),\nas well as 4 radiocarbon dates during HS1 reported in (Dai et al., 2022)(2 in the early HS1\nand 2 in the late HS1). These age constraints allow us to resolve millennial timescale changes,\nincluding distinguishing the early and late HS1, despite uncertainties in absolute ages.\nTo demonstrate the robustness of our age model, we have plotted our data based on an\nalternative age model discarding the planktic Mg/Ca-based tie point. Similar patterns of the\nrecords to that shown in the main text, especially within HS1, prove that our age model is\nrobust in discussing these features. (Fig. S3).\nThe second point raised by the reviewer is the slightly delayed benthic δ18O decline during\nHS1. This appears to be a common pattern that has been observed in the region previously.\nIn Fig. X1, it is shown that our data is consistent with another well-dated site nearby, which\nfurther lends support to our age model.\nFig X1. benthic δ18O at site MD97-2106 compared to two other SW Pacific sites in (Sikes et\nal., 2016).\n2. Sea ice, or not? The authors take a strong position on the role of sea ice in the abstract\n“…physical dynamics involving sea ice retreats (sic) dominated CO2 outgassing…” (L22) but\nare rightly more circumspect about this in the main text “However, relative roles of biological\nand physical processes to the PAZ outgassing remain ambiguous on both millennial\nand glacial-interglacial timescales…” (L157-158). More clarity on the relative contributions of\nsea ice and their timing would be beneficial.\nWe clarify the contribution of the gas exchange disequilibrium imposed by sea ice in the\nabstract (Line 22) and the main text (e.g., Line 255). In short, our PO * and [CO 2-] changes\n4 3 as\nshow that sea-ice related gas exchange disequilibrium contributed to deglacial CO\n2\noutgassing during the YD and minimally contributed to HS1 CO outgassing.\n2\n3. Tracer interdependence. PO4* and [CO32-]as are not independent tracers, as both depend\non Cd/Ca-derived [PO43-] (L48-49, 66-69). That suggests the difference between the two is\nlargely driven by carbonate ion and oxygen… but it is not at all clear what to make of the\ndistinctions between the two tracers. On the one hand, it’s interesting that the two show slightly\noffset patterns in time… but on the other, it’s not surprising that the two look similar (Fig. 2f,\ng). I would appreciate more guidance from the authors on what to make of this\ninterdependence.\nThe reviewer’s concern is valid in that Cd/Ca-derived [PO 3-] contributed to both PO * and\n4 4\n[CO 2-] . In the revised version, we add a new figure (Fig. 3) to show contributions of\n3 as\nunderlying proxies to PO * and [CO 2-] . As shown in this figure, [PO 3-] is a minor contributor\n4 3 as 4\nto millennial-timescale variabilities in PO * and [CO 2-] .\n4 3 as\n4. NSW and SSW endmembers + contributions over time. The authors note that the advantage\nof PO4* and [CO32-]as is that both tracers are quasi-conservative in the ocean interior (e.g.,\nL50-60). However, the downside of this is that these proxies are set by the endmember values\nof NSW and SSW and their relative contributions to the core site. It is known from the\nauthors’ previous work that these endmember values are not stable over time (e.g., Yu et al., 2019 for\n[CO32-]as in the N. Atlantic); indeed, the goal of this study is essentially to reconstruct the\nSSW endmember PO4* and [CO32-]as over the last deglaciation (e.g., L63-65). To this end,\nit is critical that the authors know the contributions of the relative endmembers. In my mind,\ntheir approach of assuming a percentage contribution of SSW and allowing it to vary (e.g.,\nL290) leaves much to be desired, especially as proxy data (Nd isotopes) may help address\nthis question for them.\nIn the new version, we focus on discussing SSW endmember changes based on our best\nestimation of SSW% and NSW endmembers determined from literature data (including Nd\nisotope (εNd) data for SSW%, previous North Atlantic records for NSW endmembers).\nSensitivity tests (now moved to the supplementary materials, Fig. S8) support estimations\nbased on our assumptions. This is because our site was dominantly affected by SSWs over\nthe last deglaciation (with SSW% over 70%), any variability of the SSW% at our site during\nthe last deglaciation would be expected to be small.\nOur assumptions about SSW% at our site are based on εNd data in the region. However, we\ndo not directly use εNd data to infer SSW% changes in the SW Pacific. This is because εNd\nhas different endmembers (the North Atlantic and the North Pacific) from PO * and [CO 2-]\n4 3 as\n(the North Atlantic and the Southern Ocean).\n5. Presentation of results. The authors do not discuss their underlying proxy reconstructions,\ninstead they move directly to the derived tracer products (PO4* and [CO32-]as). This is\nunfortunate – the most directly calculated results should be presented first, and only after that\na discussion of the interpreted derived products. Such additional length would allow for further\ninvestigation of intriguing features in the proxy data – for instance the offsetting (in both depth\n+ age space) of carbonate ion, oxygen, and phosphate concentration deglacial\nmaxima/minima (Fig. 2c-e). Might this relate to the relative depth habitat of the foraminifer\nused? (C. wuellerstorfi being ~sediment/water interface, H. elegans perhaps being infaunal,\nG. affinis being deep infaunal)?\nIn this work, we aim to focusing on interpreting derived tracers, which have the advantage of\navoiding some alternative explanations related to local processes for the most directly\ncalculated proxies. However, the reviewer’s point that in the previous version we moved on to\ndiscussing PO * and [CO 2-] too quickly is valid.\n4 3 as\nTo address the reviewer’s concern, we now expand background of PO * and [CO 2-] in the\n4 3 as\nintroduction, and use distribution of deep-water [CO 2-], [PO 3-], [O ], PO * and [CO 2-] in the\n3 4 2 4 3 as\npreindustrial ocean to better illustrate the strength of PO * and [CO 2-] (Fig. 1, S1, Line 69).\n4 3 as\nIn this version, we also started our discussion from directly reconstructions (Line 98). We avoid\ndirectly interpreting their implications for carbon cycle changes but explain the contribution of\ntheir variabilities to the derived tracers aided by Fig. 3.\nWe discount the varying depth habitats of employed benthic foraminifera as the driver of the\nobserved changes in the supplementary material (Line 385). The possibility that some of the\nsubtle offsets between the proxies are caused by varying habitats of benthic foraminifera\nemployed for proxy reconstructions cannot be fully ruled out. These features are however not\nthe focus of our interpretation. We focus on millennial-timescale changes in our core (including\nthe different changes during the early and late HS1), which are not affected by this effect, due\nto bioturbation and relative high sedimentation rate. This is can be supported by similar δ18O\npatterns of the epifaunal species C. wuellerstorfi and the deep infaunal species G. affinis (Fig.\nS2e) (H. elegans being epifaunal to shallow-infaunal (within top 0.5 cm), thus between these\ntwo species in the habitat depth profile).\n6. Uncertainty on PO4* and [CO32-]as. There does not appear to be an effort to estimate\nerrors on PO4* and [CO32-]as aside from those relating to the individual proxy reconstruction\nof carbonate ion, oxygen concentration, and phosphate concentration. At the very least, the\nerrors of each required measurement will add in quadrature; in addition, further error will\nresult from assumption of Redfield stoichiometry (see modern ocean uncertainties in these\nratios from Anderson & Sarmiento, 1994). A more comprehensive handling of uncertainty in these\nnew tracers will be needed than is presented in the main text, figures and methods.\nThe method for propagating errors using an iterated LOESS smoothing method for individual\nproxies and the derived tracers are expanded (Lines 320, 339, 377). We also fixed the mistake\nassociated with the [PO 3-] estimation spotted by the reviewer (see reply to minor method\n4\ncomments) and updated related error envelopes in figures.\nConsistent with previous works on proxies (e.g., (Lynch-Stieglitz et al., 2024)) and modelling\n(e.g., (Khatiwala et al., 2019)), we adopt a constant set of Redfield ratios to derive tracers and\ninfer DIC changes. However, our conclusions are minorly influenced by variable Redfield ratios\nin the ocean, as shown by additional sensitivity tests in Figs. S6, S7. In these tests, we\ncalculated PO * using varying R (160, 175, 180), and calculated [CO 2-] using varying\n4 -O2:P 3 as\nsensitivities of [CO 2-] to [PO 3-] changes with varying R (127 or 140) and R (4 or\n3 as 4 C:P Corg:Ccarb\n8) defined in Yu et al. 2019. It is noted that, as shown in Yu et al 2019 (Shown here as Fig\nX2), the sensitivity of [CO 2-] to [PO 3-] is relatively insensitive to changes in R and\n3 as 4 C:P\nR at the low [CO 2-] end ([CO 2-] is similar to [CO 2-] without correcting for the\nCorg:Ccarb 3 norm 3 norm 3 as\nconstant offset of -78), where our data points are.\nFig. X2. Sensitivities of [CO 2-] changes to [PO 3-] with varying R and R .\n3 norm 4 C:P Corg:Ccarb\nMinor methods comments.\nL236. How was anthropogenic carbon calculated and deducted? I would not expect much\ndetectible DICant at this water depth…\nAnthropogenic DIC is estimated from (Lauvset et al., 2016). It contributes negligibly to DIC in\nthe deep SW Pacific indeed, but it does substantially contribute to the DIC in the N Atlantic.\nCorrecting anthropogenic DIC is critical to reveal linear SSW-NSW mixing trend of [CO 2-] in\n3 as\nthe ocean before the industrial era.\nL262. Given all those contributions, the estimated uncertainty on phosphate concentration is\nonly ±0.05 µmol/kg? Come on now, that is not realistic. For one point of comparison, the Cd/Ca\nmeasurement precision alone in Farmer et al. (2019) corresponded to a similar or larger\nphosphate concentration uncertainty. I have no doubt the authors are better analysts than\nmyself, but the propagation of error should be handled appropriately.\nThis was a mistake, which is now corrected (~0.11 µmol/kg, ~0.05 of the estimated [PO 3-],\ndetails about how the error of [PO 3-] is derived (Line 320).\nsignificant (ED Fig. 1i), and also there is no reason to expect that a single core site would\nexperience the global ocean alkalinity change (see, e.g., regional deviations in benthic d18O\nfor whole-ocean d18O change, Lisiecki & Stern, 2016 and references therein). Can the authors\nplease further elaborate on/justify this correction and illustrate its uncertainty?\nBoth data and models suggest that the average ocean alkalinity was substantially higher\nduring the LGM than the Holocene. The effect of global alkalinity change on [CO 2-] is\n3 as\nsignificant, and correcting for it helps to better quantify [CO 2-] in the past. Although regional\n3 as\nalkalinity changes would deviate from the global trend to certain degree, the magnitude of\nchange would be similar on the orbital timescale (also the case for δ18O).\nThe reviewer’s concern about the uncertainty of the global alkalinity change is valid, as it could\nbe the driving factor for some subtle feature in our record. For this reason, we no longer argue\nfor an early-HS1 [CO 2-] increase, which is dependent on the alkalinity correction and is not\n3 as\nstatistically significant compared to the updated [CO 2-] reconstruction uncertainty. In\n3 as\ncomparison, the discussed changes from the late HS1 onwards are independent of the\nalkalinity correction (Fig. 3). We also use three different model outputs to calculate [CO 2-]\n3 as\n(Fig. S7). Although the magnitudes of [CO 2-] changes vary, the millennial-timescale [CO 2-]\n3 as 3 as\nfeatures are consistent (Fig. S7).\nReviewer #2 (Remarks to the Author):\nKey Results:\nIt is difficult to parse apart the drivers of air-sea CO2 exchange (and by extension atmospheric\nCO2 variability), particularly in the Polar Antarctic Zone where carbonate archives are\ngenerally not present. The authors of this paper attempt to circumvent this challenge by\ninferring changes in quasi-conservative properties of deep ocean waters that are sourced from\nthe PAZ; these properties are set by a combination of endmember values and water-mass\nmixing and should be recorded by benthic foraminifera at the study site. By the authors’ logic,\nchanges in these properties should reflect a combination of changes in circulation and\nchanges in endmember composition (in this case largely set in the PAZ and delivered to the\nstudy site).\nThe authors have divided the discussion of their results by timescale. On glacial/interglacial\ntimescales, they find that the deep pacific ocean was a net source of CO2 to the atmosphere,\ngiven the increase in [CO32-]as at the site. The concurrent increase in PO4* suggests that\nthere was a combination of more preformed PO43- (and thus lower nutrient efficiency) and/or\nmore preformed O2 as a result of a combination of biology and the removal of sea ice. The\nlatter is corroborated by the decrease in ventilation age documented at the site in a previous\npublication by these authors.\nOn millennial timescales, the authors suggest that the deep South Pacific was a source of\nCO2 outgassing across HS1, and that this was matched by changes in preformed nutrients\n(and by extension nutrient utilization efficiency) and/or preformed O2 inferred from PO4*.\nRadiocarbon ventilation ages suggest that this was a period of better ventilation, perhaps as\na result of sea-ice retreat and/or a shift southwards of the southern hemisphere westerlies\nduring this period. The Antarctic Cold Reversal shows a decline in [CO32-]as and PO4 from\nthe end of HS1, consistent with poorer ventilation of the waters reaching this region. Lastly,\nthe Younger Dryas measurements of [CO32-]as and PO4* suggest that this region was again\na source of CO2 outgassing as a result of biological and physical changes in the PAZ. These\nmechanisms are consistent with to those provided for the full glacial/interglacial change in\nthese parameters.\nOn sub-millennial scales, the authors examine and attempt to delineate the timing of changes\nin [CO32-]as and PO4* at this site across HS1 using ~12 data points (the sedimentation rate\nof this core is 10-20 cm/kyr depending on chosen tie points from Dai, Yu and Rafter 2021).\nThe authors suggest that the asynchrony between [CO32-]as and PO4* in early HS1 allows\nfor the discrimination between the various influences on these parameters. They suggest that\nthe lack of change in PO4* is consistent with an increased role for physical drivers (retreating\nsea ice and altered southern hemisphere westerlies) in early HS1. In contrast, according to\nthe authors interpretation, the change in PO4* across late HS1 is indicative that biological\nchanges likely spurred mid-late HS1 outgassing.\nWe thank the reviewer for the very comprehensive summary. We altered the interpretation of\nour data during HS1 based on the reviewers’ comments, which will be detailed following the\nrelated comments.\nValidity:\nThe authors make use of two quasi-conservative (and related) tracers: PO4* and [CO32-]as.\nPO4* is derived using inferred concentrations of PO43- and O2. The former is derived from\nthe measured Cd/Ca ratio of the benthic foraminfera H. elegans using a local calibration for\nCd/Ca to PO43-. The latter is derived from the gradient in carbon isotopes of co-occurring\ninfaunal (G. affinis) and epifaunal (C. wuellerstorfi) species. Changes in PO4* thus reflect\nsome combination of changes in nutrient utilization (via changes in preformed PO43-) and\nchanges in oxygen related to either biology and/or physics.\nThe PO43- measurements are also used to calculate [CO32-]as, a parameter which reflects\nchanges in CO32- (and by extension DIC and atmospheric CO2) with biological influences\nremoved. When this parameter changes, it reflects changes in carbonate ion as a result of\nCO2 sequestration or outgassing as a result of air-sea gas exchange; increases in [CO32-]as\nreflect lower DICas and outgassing of atmospheric CO2. Because increased exchange of CO2\nshould also alter PO4*, differences in the behavior of [CO32-]as and PO4* downcore should\nprovide constraints on the role of physical processes (relative to biological processes) in\ngoverning CO2 release to the atmosphere across the deglaciation.\nIn general the data appear to be of high-quality and their interpretation appears robust.\nHowever, there are some major flaws that I think need to be fixed if the paper is accepted for\npublication. In particular, the discussion of sub-millennial drivers of CO2 variability in HS1\nrelies on the statement that [CO32-]as increased throughout early HS1, in contrast to PO4*\n(line 166- 167). It would be extremely helpful to the reader to delineate the exact magnitude\nof the change in these variables over the early HS1 time period. The change in [CO32-]as\nbetween 18- 16.3 ka appears to be on the order of 5 μmol/kg at most, and perhaps less than\nthat if the lowest data point on the edge of HS1 is excluded. The data for PO4* looks as though\nit is also slightly increasing over the same time period, with the exception of a slight decrease\nright at 16.3 kyr (Figures 2 and 3).\nWe recognized that the early-HS1 [CO 2-] increase presented in the last version was not\n3 as\nstatistically significant compared to the updated reconstruction uncertainty (due to an error\nwith the uncertainty associated with [PO 3-], see responses to Reviewer 1’s comments for\n4\ndetails). We revised our discussion on HS1 changes focusing on the minimal changes in PO *\n4\nand [CO 2-] during the early HS1 and increases in these tracers during the late HS1.\n3 as\nTo make things more confusing, the onset of HS1 on Extended Data Figure 1 appears at a\ndifferent place in the [CO32-]as and PO4* curves than in Figures 2 and 3. Because the figure\nis plotted against depth it makes interpretation and comparison difficult, but if Extended Data\nFigure 1 is used in interpretation it appears that [CO32-]as actually decreased between 18-\n16.3 kyr.\nThere was a mistake in depth of [CO 2-] when plotting Figure ED1, which is corrected in Fig.\n3 as\nS2 of this revision.\nIn other places in the text the authors refer to the change in N. Atlantic [CO32-]as as small\nacross all of HS1 (lines 117-118), but the change in early HS1 in the NEAP-4K record is almost\ndouble that of the MD97-2106 core. How much of the small change in [CO32-]as in early HS1\ncould be from a change in NSW/SSW proportion and/or end member composition? Again this\nis unclear, but if the change in [CO32-]as is on the order of 5 μmol/kg, even a small change\ninduced by circulation/water column structure change may represent a significant proportion\nof this inferred increase.\nWe reorganized the discussion on the minor contribution of NSWs to [CO 2-] changes at our\n3 as\nsite during HS1. In short, even when overestimating the NSW% at our site during HS1 (20%)\nthe NSW [CO 2-] increase of ~7 µmol/kg can only make a small contribution to our site (~2\n3 as\nµmol/kg, Line 179).\nIt is not convincing to me from Figures 2/3 or from Extended Figure 1 that the change in\n[CO32-]as is significant either relative to the lack of change in PO4* or its starting value at the\nbeginning of HS1. Given that this is one of the 3 major findings of the paper and is used\nevidence for physical changes initiating CO2 outgassing, this needs to be addressed prior to\npublication. The other evidence cited (radiocarbon ventilation, sea ice extent, and δ15N) are\nconvincing in this respect, but have all been published elsewhere.\nWe agree with the reviewer that this [CO 2-] change during the early HS1 is not statistically\n3 as\nsignificant compared to the updated reconstruction uncertainty as mentioned above.\nSignificance:\nThe authors’ text is well-cited with regard to deglacial CO2 release from the deep Southern\nOcean, including mechanisms and modelling. The literature they cite shows that, over the last\n5-10 years, there have been numerous results connecting the deep Southern Ocean to the\nrelease of carbon from the deep across the deglaciation. There is, to my knowledge, already\na consensus that the Pacific would have been a major player in this process both on\nglacial/interglacial and millennial timescales. Where this article therefore would be adding to\nthe literature, then, would be if it could distinguish the role of physical and biological changes\nin driving this release for the PAZ. However, for the glacial/interglacial and millennial\ntimescales of this article, the authors acknowledge that the co-variance of their tracers does\nnot allow them to fully delineate between the dual roles of physics and biology. They do\nattempt to constrain the amount of carbon release implied by their data for glacial/interglacial\ntransition, but this extrapolation is not fully explained nor is the role of the deep South Pacific\n(relative to, for example, the Atlantic) constrained by this extrapolation.\nThe authors suggest that the behavior of the two conservative tracers over HS1 allows them\nto discriminate between influence of physical and biological tracers at its onset. To me this\nwould be a significant advance in understanding provided by this paper, as it would serve to\nconfirm evidence from radiocarbon and sea-ice proxies. However, the authors must provide\nmore information regarding the changes in these two tracers across this time period to better\ndemonstrate their conclusions (regarding whether the change in [CO32-]as is significant\nrelative to the lack of change in PO4*).\nWe made substantial changes to our manuscript to clarify that our conclusions on both orbital\nand millennial timescales are significant.\n Paired carbonate ion-phosphate-oxygenation reconstructions for deep-water preformed\nconditions\nOur work is a first attempt to use these proxies to quantify two different deep-water preformed\nproperties tracing different biogeochemical processes during the last deglaciation. Firstly, our\nreconstructions from the deep ocean are important in the context of the published surface\nrecords suggesting the influence of the nutrient utilization efficiency on PAZ CO outgassing.\ndemonstrated that the surface nutrient utilization changes previously reconstructed, which are\noutside main deep-water forming region and offset from the deep water forming season,\nindeed were associated with deep-ocean biogeochemistry changes. Secondly, our approach\nfocusing on reconstructing preformed properties can avoid some uncertainties associated with\nextrapolating local in-situ deep-water biogeochemistry changes to the global scale. (e.g., Line\n70).\n Contributions of biological and physical processes to deglacial CO outgassing\n4 3 as\n(including both biology- and physics- related DIC changes). On glacial-interglacial timescales,\nwe provide strong evidence that both biological and physical processes are contributors of\ndeglacial CO outgassing. More importantly, by quantifying the DIC changes associated with\n2\nother tracers, we reveal that the biological processes are a more dominant contributor during\nthe early deglaciation, but the physical processes are more important during the late\ndeglaciation.\nOur work is the first attempt to distinguish mechanisms of the deep Southern Ocean CO\n2\noutgassing using proxy data. it is difficult to use single proxy for in-situ properties (such as\n[O ], radiocarbon, and [CO 2-]) to discern contributions of biological and physical processes as\n2 3\ndid in previous works on deep Southern Ocean carbon cycle changes.\nThe significance of this paper would clearer with discussion in more detail of how the results\ncan or cannot discern between mechanisms or how they could build upon the various\nmodelling results that they cite. Khatiwala et al. (2019), for example, find that sea ice extent\nand ocean circulation are minor contributors to changes in carbon storage relative to the role\nof air-sea equilibrium in amplifying changes in ocean temperature and iron fertilization. Cliff et\nal. (2021) also suggests a minor role for circulation and the biological pump in driving deep\nocean deoxygenation (and thus PO4*), but a large role for sea ice extent and iron fertilization.\nSuch contradictions might be cleared up from paleoclimate-based inferences, but is largely\nnot discussed here.\nSome of the modelling results mentioned by the reviewer is added to our introduction (Line\n29). Our deep-water reconstructions can quantify CO outgassing from the PAZ and\n2\ndistinguish contributions from biological and physical processes. For biological processes, our\nestimation is consistent with PAZ nutrient utilization efficiency changes inferred from previous\nδ15N records. The key role of the PAZ as the key region modulating deep-ocean regenerated\nDIC changes can also be supported by moderate deglacial CO outgassing from the SAZ\n2\nattributable to iron fertilization (Dai et al., 2022). For the contribution of physical processes,\nwe cannot distinguish contributions of sea ice extent, position and strength of westerlies, and\nocean circulation, as these processes are not directly linked to the proxies we reconstructed.\nOne major thing that is unclear to me is why the authors do not discuss existing PAZ δ15N\ndata. The authors suggest that the lack of carbonate archives from the PAZ preclude analysis\nof nutrient utilization in the region, but diatom archives (plotted in Extended Figure 5) provide\nsuch a record. Likewise, deep sea corals provide a coherent and chronologically well-\nconstrained record of δ15N. There is no discussion of these diatom records throughout the\ntext, and only a single mention of deep sea coral-bound δ15N. Recent important work (Ai et\nal., 2020) provides important constraints on the role of Southern Hemisphere westerly winds\nand upwelling on nutrient conditions in the PAZ and should be discussed here in reference to\nthe conclusions being drawn, especially with regard to the interpretation of δ15N records from\nthe PAZ.\nIn this revision, δ15N data are discussed in more detail. We acknowledge that the idea of the\nPAZ control on air-sea CO partition through nutrient utilization efficiency is from the δ15N\n2\nrecords (Line 152). However, we also explain why our deep-water records, showing the deep-\nwater preformed changes is also significant. It is not yet known if the nutrient utilization\nchanges inferred from δ15N records led to changes in deep-water preformed properties and\nthus deep-water carbon storage. The reasons for this are that 1) the published δ15N records\nare in the Open Antarctic Zone, not in the main region forming deep waters (although part of\nsurface water in the OAZ contributes to the process), and 2) that the δ15N records changes in\nseasons offset from deep-water forming season (Line 155). Our deep-water PO *\n4\nreconstruction provides strong evidence from the deep ocean to complement the surface δ15N\nrecords for the PAZ nutrient utilization efficiency to affect deep-water chemistry.\nWe compare our reconstructions to records with comparable or better age models when\ndiscussing millennial-timescale changes. For this reason, we choose to discuss the coral-\nbound δ15N records in the main text to support our observations during HS1. The δ15N record\nat site MD12-3394 (Ai et al., 2020) indeed show a pattern nicely consistent with our PO * \n4\nrecord (Fig. S10). The reason that we did not discuss this record in detail is that its age model\nduring the last deglaciation, which was determined by two tie-points at 11.7 and 18.2 ka (Table\nS2 in Ai et al., 2020), does not permit discussing millennial-timescale changes, despite the\nhigh sediment rate at this site. Lack of carbonates for 14C dating during deglaciation and the\nLGM is common in high-latitude Southern Ocean sediment cores (e.g., at site MD12-3394),\nand it is one aspect of the challenge associated with PAZ reconstructions. We do now,\nhowever, refer to Ai et al (2020) when discussing westerlies’ influence on deglacial CO\n2\noutgassing (Line 235) following the Reviewer’s advice.\nAnother discussion point that should be addressed are the ‘Late Holocene’ values marked in\nFigure 2. It is not entirely clear from the methods section what age the “Late Holocene values\nreflect (Lines 214-215), but from inference (the disappearance of G. affinis until 0-17cm in the\ncore) it seems likely these reflect values core top values. The values of [CO32- ]as and PO4*\nare both equal to their value at 8-10 kyr. However, we know from Studer et al. (2018) that\nthere was a substantial (20 ppm) rise in CO2 across the late Holocene (between 10 kyr and 2\nkyr) that was accompanied by a large change in nitrate supply, reflected in the δ15N of\nSouthern Ocean PAZ and Polar Frontal Zone diatom-bound N records, as well as in Coral-\nbound δ15N records from the PAZ. Presumably this change in δ15N should also be reflected\nin the authors’ records of PO4*. The changes in overturning driven by migration of Southern\nHemisphere Westerly Winds (hypothesized by Studer et. al 2018) should also perhaps be\nreflected as changes in [CO32-]as. Please discuss why your core top values relative to 10 kyr\ndo or do not contradict existing data regarding this period.\nThe Late Holocene data are shown in the supplementary data table. The Late Holocene age\nof these samples is also mentioned in the supplementary material (Line 405).\nThe current study focuses on changes during the last deglaciation. The reviewer provides an\ninteresting hypothesis for the partitioning between the Southern Ocean and the atmosphere,\nwhich our methodology can be employed to test. However, to test this hypothesis, more proxy\ndata and age constraints during the Holocene are needed, which warrant a future study.\nPrevious records from this same core (from the same authors in 2022) show a strong role for\naltered efficiency of the solubility pump over the same time period (inferred from data\npresented in Fig. 2A). No discussion of this in relation to the findings of this paper are made,\nyet the inference from the 2022 paper is that Southern Ocean outgassing as a result of an\naltered solubility pump makes up a significant proportion of the observed change between 19-\n15 kyr – but the current paper would suggest much of the action is taking place in the PAZ\ninstead of the subantarctic.\nChanges in both PAZ and SAZ are both thought to be important to the air-sea CO exchange\n2\nin the Southern Ocean (e.g., (Jaccard et al., 2013)). The current study used materials from\nthe same sediment core as our work in 2022. The current study mainly uses deep-water\nreconstructions at site MD97-2016 to infer changes in the PAZ and show that the air-sea CO\n2\nexchange in the PAZ are important during HS1. Dai et al. (2022) by contrast used surface-\nwater reconstructions from the same site to track SAZ changes and suggested that solubility\npump is important for SAZ surface carbon cycle changes. Taken together these studies\nprovide insight about how different parts of the Southern Ocean contributed to changes in the\nair-sea CO exchange during the last deglaciation.\ndelayed carbon release from deep waters at our site during HS1 compared to the atmospheric\nCO rise implies important contribution of the SAZ to the early-HS1 atmospheric CO rise (Line\n2 2\n273).\nData and methodology:\nThe data appeared to be of high quality, and technically sound. I have minor comments that\ncan be addressed line by line in the text. The main major comment I have is for readability:\nthe authors here assume the reader has a detailed understanding regarding the calculation\nand interpretation of [CO32-]as, and rely on the reader to have a fairly detailed prior\nunderstanding of how this is calculated in order to follow the discussion. A short discussion of\nthe tracer and its relationship to atmospheric CO2 and DIC would be helpful.\nWe expand description of the two tracers in the introduction, especially the main components\nof these tracers, and what they are related to (Lines 42-58). We also describe how [CO 2-] is\n3 as\nrelated to DIC changes, in the context of its distribution in the modern ocean (Line 63).\nTo more clearly explain downcore PO * and [CO 2-] changes, we discuss in details\n4 3 as\ndistribution of these two tracers in the modern ocean, in comparison to their underlying\ncomponents (Line 70). We also describe how variabilities in deep-water [CO 2-], [PO 3-], and\n3 4\n[O ] contributed to downcore PO * and [CO 2-] changes (Line 90) before moving on to discuss\n2 4 3 as\ncarbon cycle implications of the derived tracers.\nMore details about [CO 2-] calculation are also given in the methods (Line 358).\n3 as\nMethods:\nThere is evidence that morphotypes of C. wuellerstorfi (e.g. sensu lato vs. sensu stricto after\nGottschalk et al. 2016 and references therein) can alter the values of δ13C and trace element\nratios. Given their expertise in the field, the authors are likely aware of this – were selections\nmade by morphotype within C. wuellerstorfi that are not mentioned in the text? If not, could\nthis account for some of the variability in, for example, B/Ca and δ13C that would give rise to\nartificial changes in reconstructed [CO32-]as and PO4* values? δ13C in particular can vary\nby between 0.2-0.7‰, which would be significant for reconstructions of bottom water oxygen\nconcentrations.\nBoth of our B/Ca and δ13C data are minimally affected by potential differences related to C.\nwuellerstorfi morphotypes. We identified C. wuellerstorfi specimen following description in\nHolbourn et al (2013), which is largely in line with C. wuellerstorfi (sensu stricto) described by\nGottschalk et al. (2016). The morphology of picked C. wuellerstorfi is consistent within the\ncore, ruling out any possibility for systematic offset in B/Ca and δ13C related to morphotype\nwithin the record between different periods. Any abrupt changes in morphotype of picked\nsamples within the records would also likely lead to outliers in the B/Ca and δ13C records,\nwhich is not seen in our relatively smooth records.\nIn lines 222-223, there is mention of splits of large samples into subsamples, which were\npresumably run in different batches and/or treated as independent samples. What is the\nstandard deviation of these “true” replicates? Are they averaged and included in the figures\n(e.g. Extended Figure 1) or plotted independently? The 5th oldest point in Extended Figure 1g,\nfor example, looks as if it may be 2 replicate samples, but one is significantly different than the\nother – but it is hard to tell.\nThe standard deviations of the replicated samples are added in the methods (Line 294),\nReplicated samples are plotted independently in Fig S2, but the averages of the subsamples\nare shown in Figs.2, 3, which are used to derive smoothed records. All data are shown in the\nsupplementary data tables.\nWere the benthic foraminiferal tests for stable isotope analysis (lines 230-233) treated in any\nway (e.g. crushed and sonicated)? Why not use splits of C. wuellerstorfi from the trace element\nmeasurements?\nThese samples are untreated, which is commonly done for benthic foraminifera (e.g., sample\nsubmission guidelines at MARUM, https://www.marum.de/en/Infrastructure/Guidelines-and-\nTerms-of-Use-Carbonates.html).\nAlthough using splits of the C. wuellerstorfi shells for both trace element and stable isotopes\nmay yield better records, using different C. wuellerstorfi shells for stable isotopes and trace\nelements is also common practice (e.g., (Allen et al., 2020)).\nδ18O is plotted in Extended Figure 1E but there is no reference to it in the text, and no\ndiscussion of the methodology (species?) for the δ18O measurements (line 233 is the only\nmention). It appears δ18O largely agree, but there are places with offsets of up to ~0.25‰ in\nco- occurring benthics. This difference is disguised by the scale of glacial/interglacial change\nin δ18O – is this a normal difference in δ18O between these two species? If so, perhaps this\nshould be discussed, and if not, there should be a rationalization for how these particular\npoints can be used for co-occurring measurements of δ13C gradients.\nThe methodology is updated in the Methods (Line 298). The oxygen isotope data are used to\nshown synchroneity of varying proxy records downcore which is now specified in the\nsupplementary materials (Line 389).\nThe deglacial trends of oxygen isotopes in C. wuellerstorfi and G. affnis are overall in good\nagreement as the reviewer mentioned. The δ18O offset between G. affnis (corrected by the\nspecies offset of 0.94‰) and C. wuellerstorfi is 0.00±0.21‰ (1 standard deviation).\nConsidering the deviation within the data set, it is not well-grounded to reject data with larger\noxygen isotope offsets, although the reviewer’s concern over these points is valid. Here, we\nshow an alternative [O ] record rejecting these few points in comparison with the one\n2\npresented in the main text. The general trends of the two [O ] records are the same.\nconcentration estimated from carbon isotope gradients between C. wuellerstorfi and G. affinis.\nThe dashed line is the smoothed curve without data points with large oxygen isotope offsets\nbetween C. wuellerstorfi and G. affinis (shown by crosses).\nA table or a more organized discussion of the results of sensitivity tests would also be helpful\nfor the reader – at the moment the discussion feels haphazard. The results of the sensitivity\ntests are presented parenthetically (Lines 91-93 for example), which makes it hard to follow.\nThe methodology behind the sensitivity tests is straightforward and described in the text (lines\n290 onwards) but the results of these tests as they are discussed in the text are not as clear.\nWe reorganized the way we present sensitivity tests. To be concise in the main text, we focus\non interpreting SSW changes calculated from preferred SSW% at our site and NSW\nendmembers. Sensitivity tests are used to show that SSW changes we calculated are\ninsensitive to uncertainties in SSW% at our site and NSW endmembers. Sensitivity tests are\nsummarized in the supplementary material and the results are shown in Fig. S8.\nAnalytical approach:\nThere are places where the analytical approach is not always clear, though largely the\nauthors have done a good job here. Several of the figures contain data that is not explained\nand/or referenced in the main text or supplement for the purpose of discussion. This includes\nFig. 2A and 2B, Extended Data Fig 1B and 1E, Extended Data Fig. 3, Extended Data Fig. 4b,\nand most notably Extended Data Fig. 5C.\nWe now refer to the figures where appropriate, all supplementary figures are cited in main text\nand Methods.\nMore and/or clearer discussion of the results of the sensitivity tests would also benefit the\nmanuscript. It is not always clear what the tests and results of the test are (Lines 89-93 are a\ngood example of this). More discussion of the sensitivity tests relative to the findings is\nnecessary - for example, in lines 89-93 the authors discuss the range of sensitivities in SSW\nPO4* implied by potential changes in NSW PO4* - but neglect to then relate this finding back\nto their results (potentially because the order of magnitude of the sensitivity test is on par with\nthe entire change in PO4* in SSW). If the sensitivity tests are included in discussion of the\nresults (which they correctly are), they should also be discussed in greater detail. It is also\nunclear why the authors do not conduct sensitivity tests for other transitions (e.g. YD) where\nthey are suggesting there must have necessarily been changes in the balance of NSW/SSW\nat the site.\nWe simplified discussion of the sensitivity tests in the main text (see responses to the previous\ncomments). The purpose of our sensitivity test is to show that our estimated SSW changes\nare robust.\nSensitivity tests are employed for LGM-Late Holocene changes, as the Late Holocene levels\nof PO * and [CO 2-] of SSWs and NSWs along with SSW% at our site are constrained by\n4 3 as\npreindustrial hydrographic data. For millennial-timescale changes, in most cases, NSW\nendmember values and SSW% at our site during neither of the two compared time periods is\nconstrained. For simplicity, we maximized contributions from NSW endmembers to PO * and\n4\n[CO 2-] , and show that changes in NSW endmember values and SSW% cannot fully account\n3 as\nfor the observed changes in PO * and [CO 2-] at our site. The fraction of PO * and [CO 2-]\n4 3 as 4 3 as\nchanges at our site unexplained by NSW endmembers and SSW% are attributed to SSW\nchanges (e.g., Line 247).\nA specific point that I found confusing in the sensitivity tests is that, with their help, the authors\ndefine the increased change at the site from the change in NSW endmember to be 7 μmol/kg\nand the SSW endmember contribution to be 30 μmol/kg. However, adding these up gives a\nchange larger than the glacial interglacial change suggested in lines 73-74 (32 μmol/kg\n+/- 4). Apologies if I have misunderstood, but these numbers do not seem to add up correctly,\neven accounting for the sensitivity test results.\nWe removed this part from our discussion.\nThere is no discussion of whether the core used here is of high enough sedimentation rate to\nanalyze the sub-millennial scale variability that makes up the latter part of the paper, with the\nexception of lines 163-165 which references chronological uncertainties in Extended Data\nFigure 1 that are not evident in the figure or discussed in detail anywhere I can find. If there\nare significant uncertainties then they need to be described. Regardless, Extended Data\nFigure 1 does not match Figure 2 or figure 3 with regard to the onset of HS1 in PO4* or\n[CO32-]as - as an example, there are 4 data points prior to the gray bar in Extended Data Fig.\n1i, and only 2 or 3 in Fig. 2g. Likewise, the upward trend in the oldest part of the PO4* record\nin Fig. 2F contradicts the data points in Extended Fig. 1h. Further confusing the analysis of\nthis record, Extended Fig. 1 lacks the 16.3 kyr dashed line, making it difficult to analyze how\nmany points constitute the trend the authors are inferring exists for [CO32- ]as and not for PO4*\nbetween 18- 16.3 kyr.\nWe provide more information about the sediment core, including sediment rates and age\nmodel (Lines 92, 392). All the raw data are also provided in supplementary data table.\nFig. S2 (old ED Fig. 1) is improved, correcting for a few errors in depths, also adding a dashed\nline inicating 16.3 ka.\nThe extrapolation of carbon release from the glacial-interglacial transition needs to be made\nclearer and discussed further, or removed. What is this amount relative, to say, the Atlantic or\nIndian – and if the amount released is comparable in size to the atmosphere then does this\nimply that the majority of the C release was accomplished through a reorganization of the\ndeep Pacific alone? What was the role of the intermediate waters in this reorganization?\nLines 110-113 are not clear to me and need further clarification.\nThis part is removed.\nClarity and context:\nThere are several places where the writing is unclear and obscures the logic or argument.\nThe authors’ use of the phrase “direct evidence” (lines 83-85) needs to be substantiated or\nmade more clear - why is this “direct evidence” compared to the numerous other records of\nSouthern Ocean deglacial carbon loss? These records also represent inferences, given that\nthey are not from the PAZ and rely on inferred relationships between trace metal/calcium ratios,\nbottom water chemistry, pore water chemistry, etc, as well as sensitivity tests and assumptions\nregarding the characteristics of NSW, SSW, and their proportions. While the authors of\nexisting δ15N records would not necessarily say their records record “direct” CO2 outgassing\nfrom the AZ, they might be inclined to argue that diatom records from the PAZ are a much\nmore direct record of nutrient utilization efficiency than PO4* inferred here. Why is B-P\nradiocarbon from the site is not “direct” evidence (lines 103-106, for example)?\nWe no longer use the phrase “direct evidence”, as we agree that it could be too strong as we\nused deep-water records to infer PAZ changes. Nevertheless, our approach to constrain\nchanges in deep-water preformed values do have its unique advantages to infer PAZ change.\nAll proxies, including δ15N, radiocarbon, deep-water in-situ properties we reconstructed\n([CO 2-], [PO 3-], and [O ]), as well as deep-water preformed tracers we derived (PO * and\n3 4 2 4\n[CO 2-] ), have to be interpreted with some assumptions. The strength of our approach is\n3 as\npartly that we use our new proxy data and published data based on other proxies to show a\ncomprehensive view of the PAZ’s role in contributing to the release of carbon from the deep\nocean to the atmosphere via the surface Southern Ocean. We test the robustness of our\nconclusion rigorously.\n1. We reconstruct deep-water biogeochemistry using established proxies ([CO 2-], [PO 3-],\n3 4\nand [O ]) that have successfully been implemented in different regions and during various time\n2\nperiods.\n2. The quasi-conservative nature of our derived preformed tracers is based on\nobservations in preindustrial ocean (Figs. 1, S1), and tested against variabilities in\nassumptions made on Redfield ratios and rain ratios (Figs S6, S7).\n3. Our site is predominantly affected by SSWs, and sensitivity tests show that despite\nuncertainties in SSW% and NSW endmember values, our conclusions on SSW endmember\nchanges are robust (Fig. S8).\nText, minor comments:\nLine 37: Key to this paper is the idea that these properties are conserved, adding “along flow\npaths moving away from the PAZ” might be clearer.\nWe revise the following sentence to stress that these properties are conservative: Following\ndeep water formation, signatures of surface source waters can be preserved as\n“preformed” properties that are quasi-conservative in the ocean interior (Line 40)\nLines 46-48 are confusing: PO4* is noted to be largely governed by surface nutrient conditions\nhere, but in the methods explanation has a solubility/physical component (lines 276-289). The\nstatement in lines 48-49 about deconvolving these tracers makes it sound much more\nstraightforward than the methods imply, which is shown by the fact that their\ndeconvolution can not be achieved for glacial-interglacial or millennial timescales.\nWe now mention that the LGM-Holocene change in physical component in PO * is small based\n4\non modelling study (Cliff et al., 2021) (Line 142). This component is often ignored in paleo\nreconstruction (e.g., deriving preformed [PO 3-] from AOU as in (Lynch-Stieglitz et al., 2024)).\nThis sentence is removed.\nLines 75-76: Changes in SSW/NSW is vague wording: do you mean changes in the proportion\nof these or in the endmember composition?\nWe added SSW% to make the sentence clear (Line 116).\nLines 76-79: This is confusing – if the proportion of NSW is increasing across the deglaciation,\nchanges in NSW characteristics would have a larger impact as the proportion increased.\nThis sentence is not correct and removed. We reorganized the argument about limited\ninfluences of NSWs to [CO 2-] changes observed at our site (Line 131).\n3 as\nLines 87: “Changes” could refer to proportion or endmember again.\nLines 107-110: “Data-based evidence” is a strong way to phrase this, given that most of the\nother studies cited (including those from this site) are also Data-based.\nThese sentences are removed. (Line).\nLine 118: This refers to extended data fig. 4, I believe, not certainly not 5.\nCorrected.\nLine 119: decreased fraction of NWS during HS1 needs a citation.\nReference added (Line 177).\nLine 121-122: citation needed for this\nWhy PO * at our site can be used reflect that of SSW is expanded with references in the\n4\ndiscussion (Line 179).\nLine 136-138: Are your results adding to these conclusions listed? A reader might have arrived\nat the same conclusions (Driven by biological and physical changes) from the δ15N,\nradiocarbon, and sea salt records alone. The paper needs to be more specific about the\nconstraints that this data is providing.\nLines 156-159: Again, as noted immediately above, it remains unclear what additional\nconstraints are provided given the ambiguous conclusion. The authors would be well-served\nto make their specific conclusions stronger and/or clearer, with regard to how these add to the\nexisting literature.\nWe now explain why our perspective from the deep-water preform properties is important for\nsimilar conclusions both on glacial-interglacial and millennial timescales (Lines). We focus on\nproviding evidence of relative importance of the biological and physical processes during\nvarying deglaciation stages (Line 22).\nLines 166-167: This does not appear to be true, at least not significantly. This needs to be\naddressed with the actual data, and concentrations/values need to be given.\nAs explained in responses to main points, we no longer discuss these features.\nLines 169-170: 0.1 represents 1/3 of the Glacial-Interglacial transition, so it is not as small as\nthe authors imply here.\nWe agree 0.1µmol/kg change of PO * is not small. This however cannot be discussed with\n4\nconfidence due to the reconstruction uncertainty of PO *.\nAdded.\nLines 210-212: The finding that the biological pump plays a large role is not surprising, but\ncould be discussed further relative to recent papers finding a large role for the solubility pump\nand other physical mechanisms, particularly for intermediate Southern Ocean waters.\nWe agree with the potential importance of physical mechanism of other water masses, which\nis briefly discussed later (Line 276). Here, we choose to focus on the PAZ ventilated deep\nwaters.\nFigures, minor comments:\nCaption Fig. 1 has the words “Lower Circumpolar Deep Water” but no explanation of what\nthis is in reference to with regard to the rest of the sentence (Lines 323-324).\nThis is removed from the figure and caption as LCDW is not used in our discussion.\nFig. 2: The late Holocene values for Figure 2 [CO32-]as are not at all close to the values in\nFigure 1D. While some of the absolute value change is due to calibration, the values are\nsignificantly higher than those given for the modern ocean. Why is this the case, if the\nPO4values are accurately reflected in the core tops?\n[CO 2-] in Fig 2 of the last did not correct for the constant of -78 µmol/kg. Both late-Holocene\n3 as\nPO * and [CO 2-] are close to the pre-industrial levels (Figs. 1, 2).\n4 3 as\nWhy does Fig. 2F increase at the oldest point in the record without a data point on the plot\nthat is older than ~18 kyr?\nThis was generated by smoothed curves of [PO 3-] and [O ], but was not explained. It is\n4 2\nremoved.\nFig. 3: It is not clear what the gradient in the color bars on the triangles in b,c,e, f and h are\nreflecting given the depiction of timeseries. Also, should Fig. 3D say B-P and not B-A age?\nColour bars with colour gradients are removed. The radiocarbon data shown is B-A age. A\nstands for atmosphere, and we change to Atm for clarity.\nExtended Fig. 1: This needs substantial improvements as it does not align with Figure 2 or 3,\nand does not have a marker for 16.3 kyr.\nThis figure (now Fig. S2) is improved. Error in depth is corrected, and a line indicating depth\ncorresponding to 16.3 ka is added.\nExtended Fig. 2: Typo for a instead of (a) in line 567.\nCorrected.\nExtended Fig. 3: Needs a better legend, given that there are actually 3 sets of data on the\nplot and not 2. This figure is not discussed in the main text. For example, hat does the\nsubstantial gradient (20 μmol/kg) between MD97-2106 and the S. Indian Core in reconstructed\nO2 imply for the deep South Pacific vs Indian Ocean during early HS1?\nThis figure generally show that these deglacial deep Southern Ocean [O ] records are similar,\n2\nand is now discussed (Fig. S5, Line 101).\nThe difference in [O ] between our site and the Indian Ocean site is close to the reconstruction\n2\nuncertainty preventing detailed interpretation.\nExtended Fig. 4: Is there a reason that the Yu 2022 data for the South Atlantic is not also\nincluded here? I assume the reasoning is to show the endmembers for NSW and SSW, but\nthe GGC90 record was used by Yu et al to infer significant release of carbon from the Atlantic\nsector of the Southern Ocean as well.\nGGC90 record is added (Fig. S8). GGC90 record mainly reflects intermediate depth carbon\nrelease, which complement with deep-water carbon release.\nExtended Fig. 5: It would be helpful to discuss Fig. 5C in the text. Also, putting the late\nHolocene values for the δ15N records and for PO4* would be helpful for any discussion of the\nlate Holocene evolution of these tracers.\nNot discussing δ15N records other than the coral bound record is because these records lack\nwell-defined age models allowing for discussing millennial scale changes, as well explained\nabove.\nLate Holocene δ15N changes are omitted in this figure as this work focus on deglacial changes.\nReferences:\nThe paper is well-cited, though not many of the citations are used to further the discussion of\nthe text (for example, δ15N references) despite being in the text. These citations may also be\nof interest for the discussion:\nRonge, Thomas A., et al. \"Southern Ocean contribution to both steps in deglacial atmospheric\nCO2 rise.\" Scientific reports 11.1 (2021): 22117.\nMoy, Andrew D., et al. \"Varied contribution of the Southern Ocean to deglacial atmospheric\nCO2 rise.\" Nature Geoscience 12.12 (2019): 1006-1011.\nFogwill, C. J., et al. \"Southern Ocean carbon sink enhanced by sea-ice feedbacks at the\nAntarctic Cold Reversal.\" Nature Geoscience 13.7 (2020): 489-497.\nIn this revision, we pay attention to specifying the types of cited proxies and referring to\nreferences where appropriate in the discussion. We also refer to some of the recommended\ncitations for discussion about changes during the Antarctic Cold Reversal (Line 235) and the\nYounger Dryas (Line 261).\nReviewer #3 (Remarks to the Author):\nDai and Yu present a valuable dataset, interpreted in a cutting edge quantitative framework,\nto investigate glacial carbon storage and release. This is a really interesting contribution and\nfollowing revisions would, I think, make a very suitable Nature Communications paper.\nI have provided detailed line by line comments, which include key conceptual questions as\nwell as more minor points and typos. Note there are other typos and sections lacking clarity\ntoo - the paper would benefit from further proof reading and an attempt to more clearly describe\nprocesses to the reader alongside their quantification in different tracer signals. I also pick out\nsome of the most critical issues below.\nWe thank the reviewer for the positive comments and apologize for typos in the last version.\n*[CO32-]as - better describe what it can and can’t tell us*\nThe [CO32-]as tracer has been introduced and used in a series of recent papers from Yu’s\ngroup, yet is still relatively new and (as detailed below) needs further description to help the\nreader.\n[CO 2-] is now better described in the introduction. We explain the underlying component of\n3 as\n[CO 2-] when the tracer is introduced, the quasi-conservative property it quantifies, its relation\n3 as\nto CO air-sea exchange, and its limitation in understanding mechanisms for carbon cycle\n2\nchanges (Line 77). A new figure (Fig. 3) is introduced to help readers understand how [CO 2-]\n3 as\nis calculated.\nMore critically, the interpretation I felt leaps in to this relatively derived quantity quite quickly,\nand would benefit from spelling out some of the signals in the raw data, and how these are\nbrought together into [CO32-]as and other related properties. In particular, I was left wondering\nabout the other components of ocean carbon storage (e.g. remineralised carbon) and their\nfingerprints on the carbonate chemistry and other reconstructions here. Can we learn\nsomething about remineralised carbon storage from these data or not? This is particularly\nimportant as the topic is currently in vogue (see e.g. the Volmer et al., 2023 paper in Paleo)\nand as this question naturally arises from several points in the text. For instance right from the\nabstract, line 19 highlights “carbon losses from the Deep South Pacific to the atmosphere”;\nthis to many readers will I think give an image of release of a deep carbon reservoir which\nothers have suggested has a large remineralised carbon component. But the main text\ndescription and discussion of the signals is very much focussed on preformed signals in the\nSouthern Ocean surface. I can see ways to bring these different ends of the glacial carbon\nstorage system together, but it would I think be very helpful for the authors to spell this out\nmore explicitly, both by describing the tracer in more detail and its relationship with other ways\nof thinking about deep carbon storage and release (e.g. remineralised vs preformed carbon)\nin the introduction and the initial sections describing deglacial carbon loss from the deep ocean.\n Discussing directly reconstructed proxies first then derived tracers\nSimilar points were raised by Reviewer 1. To address this point in this revision, provide more\ninformation about [CO 2-] and PO * in the introduction (Line 45), and discuss the derived\n3 as 4\ntracers and their underline components in the modern ocean to help readers understand\ndowncore [CO 2-] and PO * records. We also begin our discussion by describing how\n3 as 4\nchanges in deep-water [CO 2-], [PO 3-], and [O ] contributed to [CO 2-] and PO * changes\n3 4 2 3 as 4\nfrom the LGM to the late Holocene (Line 98).\n Remineralized (regenerated) carbon storage change\nWe aim to constrain the deep-water preformed changes, which are more homogenous for the\ndeep Indo-Pacific and deep Southern Ocean and can be tied the deep ocean changes to the\nPAZ processes.\nThe change in regenerated carbon/nutrient storage is the other side of the same token of the\npreformed changes. For example, deglacial increases in SSW PO * (largely preformed [PO 3-]),\n4 4\nwhen considering the minor changes in oceanic [PO 3-] inventory during the last deglaciation,\n4\nsuggests declined concentrations of regenerated nutrient and thus regenerated carbon. Our\ndeep-water reconstructions of [CO 2-], [PO 3-], and [O ] generally show deglacial loss of\n3 4 2\nregenerated carbon from this site. This is consistent with previous deep-water records.\nHowever, these observed changes related to deep-water regenerated carbon storage, at\nanywhere in the deep ocean can be alternatively accounted by other processes (e.g., Line 71).\nTherefore, observed loss in deep-water regenerated carbon (as at our site) does not mean\nthe lost carbon is necessarily ended up in the atmosphere. Moreover, in contrast to the\npreformed properties, regenerated carbon storage of deep waters is heterogenous in the deep\nocean depending on where they are on the route of deep-water circulation (shown by the O\n2\ntransect in Fig.1). These characters of regenerated carbon/nutrient, when compared with the\npreformed properties, show the advantage of preformed properties in inferring large scale\nocean carbon cycle changes.\n*Which water are we really feeling?*\nThe authors claim that their record is indicative of processes in the PAZ. However this needs\nto be better substantiated for a couple of reasons. While I agree with the authors that the core\nwill be dominantly influenced by SSW, not all SSW is necessarily sourced from the PAZ. For\ninstance in the Gebbie and Huybers (2010) deconvolution of deep water sources, this core\nsite would have about 50% water from the PAZ, with additional, subequal contributions from\nthe PFZ/SAZ and NSW (the NSW component is indeed nicely picked up in the low PO4* blob\nin the section figure).\n[The proportion of SSW at this site and how this is tracked by PO4* is, in addition, strongly\ninfluenced by the role of entrainment in the SSW PO4* end member signature, discussed\nfurther below.] Related to the major comment above about preformed vs remineralised\nsignatures, I’m also left wondering how I should think about timescales of change at this site\ncompared to the atmosphere, given that it will have influences both from relatively recently\nventilated water from the south alongside return flow from the Pacific. Should I think of this as\nbeing in water that has recently left the PAZ so quickly picks up changes sourced from PAZ-\natmosphere interactions? Or does it also have a component of return flow from the Pacific,\nrich in remineralised carbon, having left its surface source regions perhaps several thousand\nyears earlier? A mixture of these to me seems likely and I think some discussion of this,\nalongside the different carbon signatures being felt here as discussed above, would benefit\nthe paper.\n Contribution of the SAZ-ventilated SSWs to our site\nSSWs are indeed ventilated in both the PAZ and SAZ. The SAZ-ventilated SSWs are however\nmainly at intermediate depths and a relatively minor contribution to the deep Indo-Pacific\nOcean in the modern ocean. The minor impact is due to the even smaller influence of SAZ-\nventilated SSWs at deep depths during the LGM owing to the shallower upper cell circulation,\nwhich is explained in the supplementary material (Line 407). We do not discuss contribution\nof SAZ-ventilated SSWs to our reconstructions in the main text for clarity, and attribute inferred\nchanges in SSW [CO 2-] and PO * largely to PAZ processes.\n3 as 4\n SSW at our site\nNewly ventilated SSWs and recirculated SSWs from the North Pacific (PDW) have distinctive\ngeochemical signatures ([CO 2-], [O ], [PO 3-], and radiocarbon). As can be shown by relatively\n3 2 4\nhigh deep-water [O ] and well-ventilated radiocarbon, the returning PDWs are a minor\n2\ncontributor in the modern ocean. Due to the shoaled upper-cell overturning circulation, our site\nwas more strongly affected by the newly ventilated SSWs during the LGM, and likely also the\nlast deglaciation, as shown by εNd records.\nBy contrast, [CO 2-] and PO * of newly ventilated SSWs and PDW are similar in the\n3 as 4\npreindustrial ocean (as shown in Fig. 1, S1), which is an advantage of using the quasi-\nconservative tracers, when considering past changes in water mass mixing. For this reason,\ncontributions of PDWs to our site during the LGM, if different from Holocene, minimally affect\n[CO 2-] and PO * of deep waters, with the assumption that the LGM was close to a steady\n3 as 4\nstate similar to the Holocene.\nDuring the last deglaciation, [CO 2-] and PO * of recently ventilated SSWs and PDW could\n3 as 4\ndiffer due to changing preformed conditions in the Southern Ocean. Changes in [CO 2-] and\n3 as\nPO * of PDWs theoretically lag recently ventilated SSWs by the time taken for deep waters to\n4\ntransit from the Southern Ocean to the North Pacific. However, such a difference does not\nsignificantly affect millennial-timescale [CO 2-] and PO * changes at our site for two reasons.\n3 as 4\nFirstly, the deep-water transit time from the Southern Ocean to the North Pacific is relatively\nshort during the last deglaciation. It is estimated to be 550-850 years during the LGM, the\nlongest since the LGM (Rafter et al., 2022). This suggests that [CO 2-] and PO * of the PDW\n3 as 4\nwould be similar to the recently ventilated SSWs on millennial timescale, and the difference of\n[CO 2-] and PO * between the PDW and recently ventilated SSWs driven by millennial-\n3 as 4\ntimescale changes are small. Secondly, our site was more dominantly affected by recently\nventilated SSWs during the last deglaciation, as we explained above. Consequently, we do\nnot discuss the potential influence of changing PDW to our site.\n*Structure of paper / structure of millennial vs HS1 changes*\nI read through the section on millennial changes wondering why the earlier and later parts of\nHS1, which clearly have very different signatures, were being lumped together. I then realised\nthat they are discussed in more detail in the next section on Bio vs Physical carbon loss in\nHS1. However I think this current structure risks confusing readers. I’d suggest at least\nflagging in the millennial section that there are two distinct parts of HS1 that will be discussed\nfurther below.\nI also suggest using the general/orbital section to better discuss the tracers and the nature of\ncarbon loss, addressing the major comments above.\nWe have restructured the manuscript. After comparing LGM results with the Late Holocene,\nmillennial timescales changes are presented in chronological order.\nWe also expanded discussion on the changes on LGM-Holocene changes. Specifically, we\n describe how deep-water [CO 2-], [PO 3-], and [O ] changes contribute to [CO 2-] and\n3 4 2 3 as\nPO * changes (Line 102);\n4\n use DIC changes related to [CO 2-] and PO * to show the carbon loss was driven by both\n3 as 4\nbiological and physical processes (e.g., Lines 149, 161).\n*Early HS1 [CO32-]as change?*\nThe text discusses there being a [CO32-]as change in both the early and late part of HS1,\nand from reading the text alone without looking at the figures one might assume these\n[CO32-]as changes are similar in size. However from looking at the figure, it’s unclear whether\nthere really is a significant [CO32-]as change in early HS1, whereas the late HS1 change is\nsubstantial. Modify the text to address this directly - is there perhaps even something\ninteresting one can say about the different sizes of these changes are their relationship/slope\ncompared to other tracers that may further point to mechanisms?\nAs mentioned in responses to comments of Reviewers 1 and 2, the early HS1 [CO 2-] change\n3 as\nis not statistically significant and not discussed any more.\n*Plotting error - is there an interesting centennial signal?*\nThe relative phasing of the data in Figure 2 and ED Figure 1 is not the same (and I realise\nthat one is in age space, the other in depth - the issue lies in the phasing of proxy signals\nwithin the same core). In ED Figure 1 the notable spike to low [PO4] at 150 cm is simultaneous\nwith a sharp jump up in [CO32-]as. Taken together these signatures are really interesting in\nterms of mechanism - and are also particularly interesting given some of the very rapid\nchanges in deep sea coral data around this time (e.g. Li et al., 2020). However in Figure 2,\nthe sharp drop in [PO4] at 16.3 ka has no accompanying feature in [CO32-]as - the jump up\ncomes 500 years or so later.\nIf the jump up in [CO32-]as is associated with a drop in [PO4], then this does have implications\nfor mechanisms and should be addressed. Specifically, at present the late HS1 rise in\n[CO32-]as is discussed as being associated with rising PO4, leading to the implication that\nthis outgassing change comes about through reduced biological pump efficiency. However if\nthis major jump up in [CO32-]as (which is about half the overall late HS1 change) is in detail\nassociated with a drop in [PO4], then this surely challenges that interpretation, at least within\nthis interval? Rather, it might fit better with the coral 14C younging that happen at this time,\nindicating a pulse of pronounced ventilation that outgasses CO2 and flushes well-ventilated,\nlow nutrient waters through the water column.\nWe improved the discussed figure. The relation in Fig. ED1 in the previous version mentioned\nhere is indeed due to a plotting error with the depth of [CO 2-] . The [PO 3-] trough corresponds\n3 as 4\nto the relative low part of the [CO 2-] record at ~150 cm.\n3 as\nIt is also worth mentioning that we tend not to compare [CO 2-] or PO * with the underlying\n3 as 4\nproxies to avoid circular arguments.\nLine by line comments\n43-46 - [CO32-]as is the main focus of this paper and is still a relatively new tracer. I think the\npaper (and certainly its readers!) would benefit from some more description of this tracer here\nand also in the methods. This doesn’t need to be exhaustive (you can reference the previous\nwork for more details), but it should be enough for the reader to get a feeling how this is\ncalculated and what it indicates in more detail than currently provided. It should also discuss\nhow DIC is calculated from this tracer, which appears out of the blue later in the paper, and\nwould similarly benefit from introduction.\nWe agree with the reviewer on this point. As mentioned above, we substantially increase\nbackground of the [CO 2-] tracer in the introduction (Line 52) and details about its calculation\n3 as\nin the methods (Line 362).\n73 - jumps straight in to very derived property/concept. My preference would be to step through\nthe glacial state you have first and the different components to the carbon storage (e.g. both\npreformed and remineralised), then how they evolve over deglaciation.\nwe now start our discussion from describing how directly reconstructed tracers contributed to\nthe derived tracers. We tend not to interpret directly reconstructed tracers as they all contain\nboth preformed and remineralized component that are hard to tease out. We would like to\nfocus on the preformed properties derived from our derived tracers.\n78 - unclear as written - an increase in Northern water, which is increasing its [CO32]as,\nsounds like it could contribute to the signal of increasing [CO32]as. Key missing logical step\nis I think the different [CO32]as signatures of NSW vs SSW: as NSW has lower absolute\n[CO32]as, this may be expected to dominate. I realise this is what the mass balance\ncalculation in the sentence that follows is meant to spell out, but it’s not very clear for a non-\nexpert. Add a short clause of descriptive explanation here and a fuller explanation to Figure\nED4 caption, which is referenced here, but doesn’t have this level of explanation in the caption.\nThis sentence is removed. We simplify the discussion about the NSW endmember’s\ncontribution to [CO 2-] at our site (Line 128). We use Fig 4 to demonstrate the [CO 2-]\n3 as 3 as\ndifference between SSW and NSW.\n81 - the NSW contribute to - unclear writing.\nThis sentence is removed.\n82/83 - this seems like it would also benefit from being shown in a figure - perhaps an\nadditional panel to Figure ED4.\nThe glacial NSW and SSW endmembers are shown in the new Fig. 4.\nAlso what does the PO4* suggest about water masses?\nPO * at our site is also driven by SSW PO * change (Line 128).\n4 4\n85 - a key point (also mentioned above) that is not discussed in enough detail is whether we\ncan explicitly link changes at this site to the PAZ, vs other regions of the Southern Ocean. The\ncore is not in a region of pure PAZ-sourced AABW and likely also has a component of water\nventilated along sloping isopycnals in the PFZ/SAZ (as well as the NSW which is discussed).\nThis point is now mentioned in supplementary materials (Line 409) as mentioned in the\nresponse to a main point.\n86 - raised —> rose\nCorrected (Line 102).\n90 - being —> is\n92 - if of varying\nRemoved.\n91-93 - again I don’t find this list of numbers particularly helpful in getting a sense for how big\na deal this is - would much rather see in a figure or a table.\nAgreed. Sensitivity tests are now summarized in Fig S8.\n94 - “the latter of which” is odd here - would seem like it would only be referring to O2, but\nthink you mean PO4 and O2?\nRemoved.\n110-111 - where do these numbers come from?! I understand how you calculate [CO32-]as,\nbut not how you then go to (presumably) DIC to get total Pg carbon.\nMore details are provided how to use SSW [CO 2-] change to derive DIC changes (Line 163).\n3 as\nAlso, critically, if you are going to extend the calculation to the whole deep Pacific, you really\nneed to say something about the other proxy records in this basin (especially DCO32-, PO4,\nO2) - are they consistent with the signals here allowing you to extrapolate to the Pacific below\n1500m?\n112 - is this really sizeable? Pacific below 1500 m is likely main deep ocean carbon store. If\nneed to account for at least another 400 PG and maybe as much as 800 or 1200, then where\ndoes the rest of the carbon come from?\nWe discard the estimation of the ocean carbon inventory change. The reason is that we used\nmodelled global alkalinity changes to derive [CO 2-] , and the argument would be slightly\n3 as\ncircular if we use the derived [CO 2-] changes to infer global DIC change, which would be\n3 as\npartly dependent on the modelled global alkalinity change we employ.\n115-129 - “throughout HS1” - actually the big change in [CO32-]as and PO4* is late HS1.\nAgreed. We altered the way we describe the HS1 [CO 2-] changes, and discuss early- and\n3 as\nlate- HS1 changes separately (Line 186).\n130-139 - linked but conversely to the point above, the 14C change happens entirely during\nearly HS1, and is associated with no change in [CO32-]as! Seems like this is trying to tell us\nsomething. Alternatively, if the authors don’t think the age model is robust to this, need to be\nup front about it. What do the benthic-planktic data look like?\nOur age model resolves the early and late HS1 based on radiocarbon dates. We do not have\nenough planktic radiocarbon data to generate B-P age.\nDue to the large reconstruction uncertainty associated with [CO 2-] , there could be a small\n3 as\nphysically driven outgassing during this period, consistent with radiocarbon data. It is notable\nthat radiocarbon is always in disequilibrium with the atmosphere and can be more sensitive to\nimproved air-sea exchange than CO .\n2\n143 - Stewart et al. (2021) paper also nicely shows this and reaches similar conclusions.\nCited (Line 237).\n166 - only significant change seems late HS1. Certainly a very different magnitude / slope of\nchange if there is any change at all, and can’t just brush over this.\nAs explained above, we no longer argue for an early-HS1 [CO 2-] increase.\n3 as\n168-169 - this sentence is very hard to follow\n171 - given what you say about sea ice I think best to omit “likely stay unchanged” - if your\nsea ice and 14C interpretation is correct, then it likely increases.\n174 - if you are really in PAZ waters, then how influential is any likely change in temperature?\nProbably stays pretty cold!\n176-177 - this is really not clear to me and I think will be completely opaque to majority of\nreaders. Rephrase to make more clear what you mean. I’m left unsure whether these are\nexpected to drive a change in PO4* or not.\n180 - very unlikely that no change in physical O2 recharge if there is a change in 14C, as\n14C takes much longer to equilibrate than O2.\nDiscussion related to above comments is removed.\n186 - this younging in 14C occurs notably earlier than coral 14C data: with the exception of\nthe transient younging pulse at 16.1 ka, the main younging is all in late HS1. This is also true\nof most of the younging in the Skinner record. Why the discrepancy?\nThe reviewer’s observation is valid. However, this study does not focus on radiocarbon data,\nand thus we do not discuss this point in the main text, which could be related to the relative\nlow resolution of LCDW-level coral based radiocarbon during the period between 18 and 16.3\nka (Li et al., 2020).\n191 - This is nicely supported by Gray et al. 2023 reconstruction.\nCited (Line 261).\n195 - “14C further declined” - this makes it sound like it’s comparable to the much larger earlier\nchange. As with the points made above on the magnitude of [CO32-]as change in the different\nparts of HS1, you need to be up front about the different magnitudes of these signals.\nThis is removed.\n198-199 - hard to follow what is meant here. Even if more details are provided in the Yu et al.\n2019 paper, the reader shouldn’t have to go to this to get a general sense for how this\ncalculation works.\nDescription of this calculation is expanded (Line 361).\nAlso if this calculation is done for DIC in late HS1, what is the equivalent calculation for DIC\nin early HS1? I would think minimal as there’s to my eye very little [CO32-]as change, but the\ntext discusses [CO32-]as change in early HS1, so maybe there’s something I’m missing?\nHaving the equivalent calculation for both may be helpful.\nThe discussion on the early-HS1 [CO 2-] is removed.\n3 as\n200-201 - this might not be minor if there was a major change in sea ice\nWe specify that it is the purely physically driven preformed [O ] change is minor, as shown by\n2\n(Cliff et al., 2021) (Line 142).\n208-210 - interesting and important point but unclear as written\nThis point on relative contribution of biological and physical processes on millennial timescales\nis expanded when discussing changes during both HS1 and YD (Lines 211, 255).\n229 - 1 sigma or 2 sigma?\nSpecified (1 sigma, Line 297).\n261-262 - rephrase - this isn’t considering *all* the uncertainties as it doesn’t take in the\ncalibration uncertainty, which is more like ±0.2 umol/kg.\nWe specified all the source of the uncertainty (Line 323).\n268 - need to at least give a brief description\nCalculation of [CO 2-] is described in detail (Line 361).\n3 as\n269 - similar to point above, but even more importantly, you really need to give more\ninformation here (and ideally also in the main text). [CO32-]as is totally central to this paper\nand is still a relatively new tracer, being only previously used in a couple of other papers from\nthis group. While I realise that thinking about [CO32-]as probably now seems very natural to\nYu and co-workers, it is still new enough that the rest of the community will continue to need\nsome explanation - especially when it’s as key to a paper as it is here. Indeed I think the paper\nwould benefit from additional description in the main text too to help build intuition.\nFollowing the reviewer’s comments, we expanded description of the tracer both in the main\ntext (Line 52) and methods (Line 361).\n271-275 - I agree in principle with the correction for global alkalinity change and it’s very helpful\nto see figure ED1 that shows the [CO32-]as signatures both with and without this correction.\nHowever this figure (ED1i) also makes it apparent that the vast majority of the glacial-interglacial\nchange in [CO32-]as from the oldest to the youngest samples comes about from this correction.\nThis really needs to be discussed explicitly in the main text section on orbital scale changes.\nBecause of this concern, we discard discussion on the early HS1 [CO 2-] change.\n3 as\nthe global alkalinity correction is necessary to reveal the deglacial [CO 2-] change (although\n3 as\nadmittedly introducing uncertainty in our reconstruction).\nThe global alkalinity change is on the orbital timescale. We now show its contributions during\nthe last deglaciation in Fig. 3. Sensitivity tests show that the millennial timescale changes are\nindependent of the alkalinity correction (Fig. S7).\n286-287 - why does air sea gas exchange cause reduced nutrient utilisation?\nThis sentence is removed from the methods. However, we expand the description of the\nrelation between air-sea exchange and PO * in the main text (Line 75).\n4\n303 - this is highly dependent on what end member is used for the SSW - the AABW signature\nin the source region vs the signature that incorporates entrained CDW. Rae and Broecker\n(2018) show this changes the SSW:NSW fraction in the deep Indo-Pacific between 50:50 and\n75:25 respectively. Some discussion of this is needed, as it also has implications for what\nyou are reconstructing in terms of SSW end member. Similarly, the influence of entrainment on\nthe PO4* signature of the Southern end member should also be mentioned - this has an\nimportant influence today, is notably different in different regions of the Southern Ocean, and\nmay have also changed in the past. These are, I realise, not easy things to constrain, but it\nneeds at least some discussion.\nAs mentioned in responses to the reviewer’s main comments, we now discuss entrained CDW\nwith a substantial contribution from the deep North Pacific,\nWe agree that absolute values of estimated SSW changes indeed depend on how the SSW\nis defined, the influence of the SSW changes on the global carbon cycle changes is the same\n(a point made in Line ). This is because regardless of the definition of SSWs, the SSW% at\nour site is comparable to the deep Indo-Pacific. We defined SSW and NSW endmembers\nfollowing Rae and Broecker (2018) that gives the SSW:NSW of 75:25, which “best\ncharacterize(s) the ratio of deep Southern Ocean to North Atlantic water volume”. If we adopt\nthe Antarctic shelf-water definition of SSWs (also in (Rae & Broecker, 2018)), the SSW% at\nour site would be lower (~47% rather ~70%). Based on the LGM-Holocene changes at our\nsite, projected changes in SSW [CO 2-] and [PO 3-] would be higher. However, due to the\n3 as 4\nlower SSW% in the deep Indo-Pacific Ocean based on this definition (~50% rather than 75%),\nthe carbon cycle changes attributable to SSWs would be the same.\n323-324 - LCDW - not really a sentence.\nRemoved.\nFigure 2\n328—329 - if crosses really indicate age ties then these look unconvincing. Are the crosses\n14C control points and the triangles the ties?\nYes, this is clarified in the caption.\n331 - is the PO4 data also really including calibration error? Based on the different choices\npossible in Figure ED2, I would expect calibration error to be larger?\nAs in responses to reviewer 1, errors with [PO 3-] was not correct in the last version and\n4\ncorrected now. How [PO 3-] errors are calculated is also better explained.\n- Please show benthic-planktic 14C data here too\nBenthic radiocarbon data at this site is compared to the atmospheric value not the planktic\nradiocarbon at this site. For this reason, benthic radiocarbon data have to be shown in the age\nscale, thus it is not included in this figure.\nReferences\nAllen, K. A., Sikes, E. L., Anderson, R. F., & Rosenthal, Y. (2020). Rapid Loss of CO2\nFrom the South Pacific Ocean During the Last Glacial Termination.\nPaleoceanography and Paleoclimatology, 35(2). Unsp 2019pa003766\n10.1029/2019pa003766\nAnderson, R. F., Sachs, J. P., Fleisher, M. Q., Allen, K. A., Yu, J., Koutavas, A., &\nJaccard, S. L. (2019). Deep-Sea Oxygen Depletion and Ocean Carbon\nSequestration During the Last Ice Age. Global Biogeochemical Cycles.\n10.1029/2018gb006049\nCliff, E., Khatiwala, S., & Schmittner, A. (2021). Glacial deep ocean deoxygenation\ndriven by biologically mediated air–sea disequilibrium. Nature Geoscience,\n14(1), 43-50. 10.1038/s41561-020-00667-z\nDai, Y., Yu, J., & Rafter, P. (2021). Deglacial Ventilation Changes in the Deep\nSouthwest Pacific. Paleoceanography and Paleoclimatology, 36(2).\n10.1029/2020pa004172\nDai, Y., Yu, J., Ren, H., & Ji, X. (2022). Deglacial Subantarctic CO2 outgassing driven\nby a weakened solubility pump. Nature Communications, 13(1).\n10.1038/s41467-022-32895-9\nJaccard, S. L., C. T. Hayes, A. Martiń ez-Garciá , Hodell, D. A., Anderson, R. F., D. M.\nSigman, & Haug, G. H. (2013). Two Modes of Change in Southern Ocean\nProductivity Over the Past Million Years. Science, 339(1419-1423).\nKhatiwala, S., Schmittner, A., & Muglia, J. (2019). Air-sea disequilibrium enhances\nocean carbon storage during glacial periods. Sci Adv, 5(6), eaaw4981.\n10.1126/sciadv.aaw4981\nLauvset, S. K., Key, R. M., Olsen, A., van Heuven, S., Velo, A., Lin, X., et al. (2016). A\nnew global interior ocean mapped climatology: the 1° × 1° GLODAP version 2.\nEarth System Science Data, 8(2), 325-340. 10.5194/essd-8-325-2016\nLi, T., Robinson, L. F., Chen, T., Wang, X. T., Burke, A., Rae, J. W. B., et al. (2020).\nRapid shifts in circulation and biogeochemistry of the Southern Ocean during\ndeglacial carbon cycle events. Sci Adv, 6(42). 10.1126/sciadv.abb3807\nLynch-Stieglitz, J., Vollmer, T. D., Valley, S. G., Blackmon, E., Gu, S., & Marchitto, T.\nM. (2024). A diminished North Atlantic nutrient stream during Younger Dryas\nclimate reversal. Science, 384(6696), 693-696. 10.1126/science.adi5543\nRae, J. W. B., & Broecker, W. (2018). What fraction of the Pacific and Indian oceans'\ndeep water is formed in the Southern Ocean? Biogeosciences, 15(12), 3779-\n3794. 10.5194/bg-15-3779-2018\nRafter, P. A., Gray, W. R., Hines, S. K. V., Burke, A., Costa, K. M., Gottschalk, J., et al.\n(2022). Global reorganization of deep-sea circulation and carbon storage after\nthe last ice age. Sci Adv, 8(46), eabq5434. 10.1126/sciadv.abq5434\nSikes, E. L., Elmore, A. C., Allen, K. A., Cook, M. S., & Guilderson, T. P. (2016). Glacial\nwater mass structure and rapid δ18O and δ13C changes during the last glacial\ntermination in the Southwest Pacific. Earth and Planetary Science Letters, 456,\n87-97. 10.1016/j.epsl.2016.09.043\nReviewer #1 (Remarks to the Author):\nR1 review of Dai and Yu, “Deglacial polar Southern Ocean carbon release driven by\nbiological and physical dynamics,” submitted to Nature Communications, by Jesse\nFarmer\nThe authors have taken significant steps to revise their manuscript following detailed\nreviewer comments. I appreciated the time and care they put in the responses to\nreviewer and the clarifications added to the manuscript. There are still some\noutstanding issues that need addressing related to the clarity of the presentation; I’ve\nhighlighted these as major comments. Ultimately, although I may not fully agree with\nthe presented interpretations, I think the authors provide valuable additions to\nthe discussion of the ocean’s mechanistic role in deglacial atmospheric CO2 increases\nthat should be open to broader discussion. Thus, I support publication once the\nbelow comments are addressed.\nWe thank the reviewer for the positive comments. We’ve revised our interpretation\nfor the early HS1 data (Major comment 1). We’ve also improved writing for clarity,\nbased on the detailed and constructive comments.\nMajor comment 1. L198-203. I’m not following how it is plausible for sea-ice to be\nsufficiently reduced to allow for CO2 isotopic equilibrium (and hence 14C\nequilibrium), but not CO2 concentration equilibrium (which is ~an order of magnitude\nfaster). This is all the more confusing because the authors later call on a\nconcurrence of CO2 isotopic and concentration equilibrium during the Younger Dryas\n(L257-261). The authors should revisit the explanation; they can also allow that their\ninterpretation is not consistent with other paleodata and worth further investigation.\nWe’ve revised our interpretation of our [CO 2-] reconstruction and literature radiocarbon\n3 as\nrecords during early HS1. We now suggest that the large changes seen in radiocarbon\nrecords may be related to changes in the geometry of global overturning circulation, while\nthe roughly stable [CO 2-] at our site might be related to the lack of warming in the PAZ.\n3 as\nModel simulations show that warming is important to promote CO outgassing (Khatiwala et\n2\nal., 2019). Thus, little warming would result in negligible outgassing with limited change in\n[CO 2-] during early HS1. We fully appreciate that multiple processes can control air-sea CO\n3 as 2\nexchange and encourage further investigations to explain the contrasting changes in\ngeochemical tracers during early HS1 (Lines 231-239).\nMajor comment 2. L162-170. This paragraph needs some work. There should be a\nmore clear comparison between the [CO32-]_as-derived DIC sequestration estimate\nand the [PO4*]-derived DIC sequestration estimate. For instance, you could add in\non L166, “This DIC sequestration is larger than that estimated from [PO4*] of +58±18\nµmol/kg”. Then the sentence on L167-170 should be clarified: It appears you are\nascribing the DIC sequestration from [PO4*] directly to biological processes,\nwhereas the additional DIC sequestration from the [CO32-]_as indicates a further\ncontribution from physical processes, such as a sea ice barrier (and add references\nfor this mechanism).\nUltimately, the I found that the paragraph on L204-218 did the best job of outlining\nthe various quantitative datasets, comparing them, and reaching conclusions on\nrelevant mechanisms. This should be adopted to the other paragraphs.\nWe’ve improved writing for all related paragraphs following your useful advice. We\nclarify that DIC changes derived from PO * is ascribed to biological processes, and the\n4\nadditional DIC changes derived from [CO 2-] is ascribed to physical processes (Line 192).\n3 as\nWhen DIC changes are inferred, [CO 2-] and PO * changes that are used to make these\n3 as 4\ninferences are mentioned alongside (e.g., Lines 161,191).\nMajor comment 3. L116-136. As someone who loves thinking about endmember\ncompositions, even my head is spinning from all the values presented here. I’d\nstrongly recommend adding a table of modern NSW, modern SSW, LGM NSW and\nSSW endmember values, percentages estimated for the core site, and derived\nvalues for the core site to aid the reader through this text.\nWe’ve added Table 1 to show various values used for calculations and discussions\nfollowing the reviewer’s suggestion.\nMinor comments.\nTitle: “driven by biological and physical dynamics” is rather vague. Perhaps\n“Biological and physical contributions to deglacial polar Southern Ocean carbon\nrelease” or similar?\nWe changed the title to “Contributions of biological and physical dynamics to\ndeglacial CO2 release from the polar Southern Ocean”.\nL26-27. Suggest adding “during the Late Pleistocene glacial-interglacial transitions”\nL39-40. Suggest specifying “…quasi-conservative properties of deep waters set by\nthe PAZ surface”.\nL55. Remove unnecessary comma after “(3)”\nL58. Cite refs. 19 and 20 at the end of this sentence.\nChanges were made following these comments.\nL70-74 and Fig 1. This discussion has been greatly streamlined. One note, I’d\nhighly recommend making the color step the same unit in panels a, d (5 or 10 µmol/kg)\nand b, e (0.1 µmol/kg). This allows the reader to simply visualize how much PO4\nchanges relative to PO4*.\nWe appreciate the reviewer’s thought. After some trying, we found it is difficult to use the\nsame color steps for Fig. 1a and d, due to the smaller variability of deep ocean [CO 2-]\n3 as\n(PO *) than [CO 2-] ([PO 3-]). We thus have decided to use the full color map to illustrate the\n4 3 4\ndifferences between SSW and NSW [CO 2-] and PO *. Given the reviewer’s suggestion,\n3 as 4\nwe’ve added a short sentence in Fig. 1 caption to remind readers regarding different steps\nfor various panels. (Line 534)\nL80. Change to “may allow for distinguishing the…”\nChanged.\nL88-89. Strongly recommend briefly stating those proxies here in 1-2 sentences.\nThere is space in the main text and it prevents the reader from having to flip to the\nmethods three times.\nInformation added (Line 95).\nL101 and throughout. Note that the new supplementary figures appear out of order\nwith their presentation in the text; these will need to be reordered. For instance, Fig\nS5 should become Fig S2.\nAll supplementary figures are rearranged according to the cited order.\nL105. Add space between [CO32-] and \"at\"\nL108, L114 and elsewhere. Recommend changing to “organic matter” as \"matter\" is\ntypically indicative of a plural quantity.\nL111. Remove comma or swap “that” with “which”. The options are either “…change,\nwhich is” or “…change that is”\nL142. Suggest adding hyphen: “physically-driven”\nChanges were made following these comments.\nL144-145. The authors should cautiously note here that the atmosphere-ocean\nexchange times for O2 and CO2 are not equivalent (e.g., Broecker and Peng, 1974),\nbut that difference may be expected to be mediated by the millennial-scale\nintegration of their sediment records.\nReference: Broecker, W., and T. Peng (1974), Gas exchange rates between air and\nsea, Tellus, 26(1–2), 21–35.\nWe’ve clarified that the disequilibrium in O must be accompanied by CO\n2 2\ndisequilibrium due to the longer equilibration time of CO (Line 162).\nL156. Suggest changing to “which are presumably biased toward the spring-summer\nseasonality of phytoplankton blooms that are offset from wintertime deep water\nformation”. Although the OAZ-PAZ difference might also be important, it would take\nmore detail to explain why than is provided here, and so I think it best to not dwell on\nthis difference. (Alternatively, the authors could devote several sentences to\nexplaining why the OAZ would be expected to be distinct from the PAZ).\nL167. Change to “combines”\nL190. Change “using” to “in”\nL200. Change to “enhanced”\nChanges were made following the above comments.\nL204-218. As noted above, nice paragraph here – I find this argument convincing.\nThanks!\nL227. Delete the repetition of “at our site”\nChanged.\nL230-239. See Major Comment 2. As you did in L204-218, you should include the\nmagnitudes of DIC change inferred from both PO4* and CO32-_as to support the\nstatement of physical changes being necessary (L238-239).\nL257-261. See Major Comment 1. I find it really challenging to have radiocarbon\nventilation work with CO2 outgassing at one point, but without it at a previous point.\nPlease refer to our responses above (Major comment 1).\nL262. Suggest changing to “Conclusions”\nL265. Change to “partitioning”\nL267-268. Recommend changing to the following: “…were mediated by biological\nand physical processes both between the LGM and the Holocene interglacial, and on\nmillennial timescales during the deglaciation”\nL270. References should be added here.\nL274-275. Specifically, PAZ CO2 outgassing as diagnosed in this study\nL275. “was reduced”\nL278. Change to variability (it is also plural)\nL279-280. Specify your approach – namely, reconstructing CO2 exchange through\nPO4* and CO32-_as\nL281 (Methods throughout). The equations should all be set on separate lines, and\nnot in-text (see L312, 317, 336-337, 349, 355, 356, 358, vs. L415, 416, 419, 420...)\nChanges were made following these comments.\nL301-304. How was d18O and d13C precision quantified? What standards were\nused for normalization?\nWe added details for determining stable isotope precision in the Methods (Line 348).\nNormalization standards are commonly not reported.\nL310-346. It appears the [CO32-] reconstruction paragraph should be presented first,\nas the Monte Carlo approach is referenced in the [PO4] reconstruction paragraph.\nHowever, as the [PO4] paragraph currently comes first, this Monte Carlo approach\nhas not yet been introduced.\nL313-314. Add comma before “which”\nL351. Missing reference.\nL383-385. I believe these should be combined into one sentence.\nL506. Caption panel f. Specify “Deep-sea coral-bound”\nChanges were made following these comments.\nReviewer #2 (Remarks to the Author):\nGeneral remarks:\nThe authors have substantially reorganized the manuscript since the first\nsubmission, and the submission is much improved. The authors have clarified the\nconclusions of the paper in such a way as to make the significance clear, and there\nis considerably more discussion of the background necessary for the interpretation of\nthe PO4* and [CO32-]as data. For example, the utility of the tracer pair is made\nclear, particularly by Fig. S1. They have removed discussion of the sub millennial\nscale variability in early HS1 (18-16.3 kyr) from the text, which I think was not\njustified in the previous draft given the lack of statistical significance in the\nmeasurement, and instead discuss on the lack of change during this time.\nMore detail has been added clarifying the significance of their reconstructions in the\ncontext of existing records of nutrient utilization efficiency (e.g. δ15N records), which\nI felt was unclear or not present before. The authors results serve to confirm the\nconclusions of existing δ15N records from the OAZ and provide a complementary\nconstraint to existing records that may have been governed by different processes\n(although it appears that they are in fact consistent).\nThe discussion of the sensitivity tests was confusing before in the previous draft, and\nI think the text benefits from the authors' decision to summarize these in the\nsupporting material. I think that the added sensitivity tests details in Fig. S6 and S7\nalso substantially improve the paper.\nAll in all the authors should be congratulated on an improved manuscript that reads\nlogically and mostly clearly.\nWe thank the reviewer for the positive comments.\nSpecific Comments:\nI was perhaps unclear with regard to the core top values of PO4* and [CO32-]as in\nthe previous draft. My point was that, given the placement of the core top values in\nFig. 2F and Fig. 2G, it would seem that the values you are calling Late Holocene\n(4.2-0 kyr) identify no change between 8 kyr to Late Holocene in PO4* and CO32-\nas , which I found surprising. I take your general point in your response that this\nstudy may not be suitable to examine these features between 8kyr and Late\nHolocene. However, upon looking at the latest version of the data (I believe Fig. 3F is\ncomparable in the new paper to Fig. 2G in the first submission), it appears that there\nis now a substantial difference (~15 μmol/kg) between the Late Holocene and 8 kyr\nvalues in CO32-as. This wasn’t present in the previous version of the paper. What is\nchanging between the two versions of the paper that is driving this large change?\nThe change in the LH value is on the order of that observed across the YD or ACR in\nyour data, so is a substantial change to be seeing between two separate\nsubmissions.\nWe’ve updated the Late Holocene [CO 2-] in the 2nd version which is ~10 µmol/kg\n3 as\nhigher than that in the 1st version. The ~10 umol/kg difference is caused by two\nfactors below.\n(1) A ~7 µmol/kg difference is due to the global alkalinity effects during the\nHolocene (Fig. X1), which was not considered in the 1st version (see Fig. X1);\nand\n(2) A ~3 µmol/kg difference comes from different calculation method for the Late\nHolocene [CO 2-] averages: one includes all [CO 2-] reconstructions (n = 4)\n3 3\nwhile the other only includes [CO 2-] reconstructions (n = 2) with paired [PO 3-]\n3 4\nreconstructions (Supplement Data).\nFig. X1. Effects of the global alkalinity change on deep Southern Ocean [CO 2-]\n3\nassuming no air-sea CO exchange.\nfrom the LGM to the LH, which is ~75 μmol/kg – but then you say that that change is\nsimilar to previous records. The record you show in Fig. S5 has no Late Holocene\ndata whatsoever. The two records are impressively close in behavior and magnitude,\nbut Fig. S5 does not back up any statement about LGM to LH change being similar\nto previous records.\nIndeed, it would be ideal to have the LH [O ] reconstructions at site MD12-3396CQ,\n2\nbut such records do not yet exist. So, we now show the modern ocean [O ] to\n2\nindicate the LH [O ] levels at sites MD12-3396CQ in Fig. S4. As can be seen\nLGM and the LH.\nForgive my confusion, but I thought you might clarify lines 110-112 a bit further. It\nfeels like you are building towards an obvious conclusion for the reader about the\nsums of these multiple factors, but as I understand it, the measured increase in\n[CO32-] is 13 μmol/kg and the expected changes you list are an increase of 5\nμmol/kg (PO4 decrease), a decrease of 26 μmol/kg (global alkalinity and DIC\ndecline), and a decrease of 2 μmol/kg from T/S/Depth changes. I could easily be\nreading it wrong, but it feels like lines 110-112 imply that the numbers you give in the\npreceding lines should in some way sum or difference to 39 μmol/kg – but looking at\nthe difference between reconstructed [CO32-] and the three factors listed above\ndoesn’t give 39 μmol/kg (difference between an increase of 13 and a decrease of\n23). I say this because the way you listed the factors driving your calculation of\nCO32-as in lines 52-57 make it feel like these factors should give you the number\nback for your LGM to LH change that you give in the paragraph.\nThe confusion is largely due to different LH [CO 2-] averages as explained in the\n3\nresponse to your first specific comment. To make numbers consistent, we’ve revised\nthe amount of [CO 2-] change to 16 µmol/kg that is based on 2 data points with\n3\npaired [PO 3-] reconstructions.\n[CO32-]as during the ACR are ~0.25 and ~10 μmol/kg, respectively (Fig. 3C and 3F).\nIt seems like a 30% decline in SSW% (going from 100% to 70%) could then, in fact,\naccount for the reconstructed magnitudes almost entirely. Unless I am missing\nsomething with the calculation, I feel that this statement is not supported.\nWe’ve improved our argument here. It is theoretically possible to explain the\nobserved changes in PO * and [CO 2-] using SSW% changes as acknowledged in Lines\n4 3 as\n261-264. However, this was hardly the case, because SSW% was unlikely to 100% during HS1\nand a large SSW% decrease would contradict with deep-water [O ] changes during the ACR\n2\n(Line 265).\nLine 246: Why assuming little change in NSW%, if you mention that SSW% declined?\nI think I know what you mean but this could be clearer, as it seems like from line 242\nyou would assume some change in NSW%.\nWe’ve now clarified that we estimate the maximum influence on PO4* at MD97-2106\nfrom changes in NSW endmember values by using 70% for SSW% at our site (Line\n285). In other words, the influence from NSW endmember variations would be\ndecreased if using higher SSW%.\nLine 251-252 is not clearly written. Should this be \"driven by changes in SSW\nendmembers\"?\nRevised following your suggestion (Line 292).\nThere are small typos and grammatical errors throughout the text that I encourage\nthe authors to correct prior to another resubmission. For example, the first line of the\nabstract is, to my understanding, not grammatically correct – I think you cannot be\n\"deemed as a region regulating…\", and I think \"important implications\" would be\nreferring to changes in the polar Southern Ocean. I also was curious what the other\npart of \"in part\" is (line 13) – this seems like the main challenge for this region.\nAnother easy but obvious small error is that the references are not carefully\nformatted throughout -Line 35 for example. Also, they appear to have forgotten a\nreference listed in line 351 (ref. XX).\nWe have corrected typos and errors throughout the manuscript including those\nspecifically pointed out here.\nThis is a personal preference, but the use of δ in the figure labels for Fig. 3 is not\nintuitive to me, given its common usage as other notation for isotopes. I would prefer\npersonally if they were replaced with Δ if you are seeking to indicate change relative\nto some average (in this case LGM values). This would make it consistent with you\ncalculations lines in the Methods section (Lines 354-360, for example).\nWe chose to use δ rather than  to denote changes to avoid potential confusions.\nSpecifically, [CO 2-] is often used to present calcite saturation state, not temporal\n3\nchanges/anomalies.\nReviewer #3 (Remarks to the Author):\nThe authors have done a good job addressing my concerns and I found the paper\nsignificantly strengthened. I found the addition of the sensitivity tests in Figure 3\nparticularly helpful, along with the improved description of the [CO32-]as tracer and\nthe restructuring. It is worth noting though that despite the improved description (and\nhaving thought about this quite a bit), it took me some time to get my head back into\n[CO32-]as space, and this will likely be true of your readers. I suggest taking on a\ndescription similar to that in Gruber et al., 2009, as per the comment on line 53.\nI have a series of comments below that I think should be addressed prior to\npublication. These include minor typos and ways to improve the clarity and\nrobustness of the argument, but also a couple of more major comments, including an\nimportant point on the sensitivity of the [CO32-]DIC calculation under altered whole\nocean alkalinity.\nWe thank the reviewer for positive reviews. In the revision, we’ve provided more\ninformation about [CO 2-] and strengthened our arguments based on your comments.\n3 as\n12 – “is sparse” would be much better after “evidence”\n14 – “challenges in obtaining”\nChanges were made following these comments.\n21-22 – “biologically unutilized carbon as a key source for deglacial CO2 outgassing\nin the polar Southern Ocean” – not super clear. Suggesting leading this sentence\nwith what you actually see in the signals in HS1 (e.g. rising PO4* and [CO32-]as)\nand how this allows you to infer a shift from lower preformed nutrients and high\ncarbon uptake to higher preformed nutrients and lower carbon uptake.\n22-23 – “relieved gas exchange disequilibrium” is confusing as written. This\nsentence and the abstract in general would also benefit from more explicitly\ndiscussing the HS1 and ACR as well as the YD change.\nAlso note that, as discussed in more detail in comments on lines 164-170 below, the\nwhole ocean change in ALK and DIC is I think a significant contributor to the glacial-\ninterglacial change in carbon uptake calculated from [CO32-]as.\nWe’ve revised the latter part of Abstract following these comments. We’ve specified\nthat reduced biological nutrient utilization was a key process controlling the deglacial\nCO outgassing.\n2\nis deduced from comparing carbon changes estimated using PO * (reflecting biological\n4\nimpacts) and [CO 2-] (reflecting impacts from all processes) as described in Lines 18-24.\n3 as\n30 – I think some more nuance is required here, as in general biological processes\nare not thought to have much influence in the PAZ. While biology is always\nsomewhat tied in to carbon and nutrient balance, and may play a role indirectly too\nvia global increase in respired carbon storage at the LGM, the observation that\nexport is lower in PAZ at LGM and increases over deglacial goes in the wrong way,\nin isolation, for biology to be key in enhancing deglacial outgassing in the PAZ.\nThis sentence is rewritten to clarify that released carbon were previously\nsequestered via both biological and physical processes (Line 30).\n45 – “in deep-water formation regions”\n49 – “Recently, a tracer of the air-sea exchange”\nChanges were made following these comments.\n53 – both here and in figure 1 caption I think it would be very helpful to reference the\nGruber et al. (2009; GBC) DC_gas-exch property. This is conceptually equivalent to\n[CO32-]as and that paper explains it very clearly.\nMore generally, while I found the revised manuscript easier to follow, I would still\nencourage the authors to expand their introductory description of this tracer at some\npoint here or in the sections below, with a few lines on how it is set in the North\nAtlantic (which is a nice easy case) vs AAIW (another useful end member) vs SSW\nfrom the PAZ (which is less clear than NADW or AAIW yet is central to this paper).\nWe now refer to Gruber 2009 and 1996 when introducing [CO 2-] . (e.g., Line 51)\n3 as\nFollowing the reviewer’s suggestion, we explain driving factors for deep ocean [CO 2-]\n3\nvariation before describing how the air-sea exchange signature of [CO 2-] can be calculated\n3\n(Line 54). The reason why SSW [CO 2-] is high is also explained (Line 69).\n3 as\n64 – delete “on the other hand” as this sentence isn’t really in tension with the\nprevious one – they both have the same general sense.\nRemoved.\n78 – not strictly quite true – you discuss below quite extensively that both biological\nand physical processes might influence PO4* via gas exchange influence on O2.\nI agree this may not be a super sensitive effect, but it’s not nothing.\nAdded “largely” to be accurate (Line 84).\n93 – add “and uncertainties on derived properties are given at 1 sigma throughout”.\nAdded (Line 102).\n99 – 102 – briefly mention here or in text immediately prior to this results section\nabove that only perform the derived tracer calculations where you have paired data\nfrom the same sediment depth. This is important given the different temporal\nextents of the different datasets.\nAdded (Line 100).\n100 – unclear whether Late Holocene data is from forams or water column (line 483\nsays “late Holocene data” which seems like forams, but I can’t find these data\nanywhere – they could for instance easily go into the supplemental figure on core\ndepth). Forams would of course be preferable but I understand that this might be\nlimited by abundances. Water column is ok especially as the youngest data in the\ncore seem to align well with these values, however would still benefit from brief\njustification e.g. Have the youngest data (which from the figures appear to be ~8-10\nka) been aligned to water column values, such that the B/Ca data down core is\nanomaly from modern values (i.e. taking account of any calibration offset)? If so,\nthere needs to be justification about why it’s ok to do this with Early Holocene data\nand that we don’t need to worry about an EH peak in e.g. B/Ca.\nWe specify that these are core-top sediment samples in the caption of Fig. 2 (Line\n553).\n107 – add reference to figure 3 after “compared to the LGM”\nAdded.\n110-111 – note here that this model dependent – stating just that it influences by 26\numol/kg makes it sound as if we know this precisely. It might be 26 umol/kg, but it\ncould alternatively be 15 or 40, we really don’t know. Should also note that while\nthis is a big influence on the size of the glacial-interglacial change, it shouldn’t\nsubstantially influence the structure or timing of the deglacial changes – can also\nreference S7.\nWe’ve added a sentence citing the sensitivity test to show that our conclusion (both\nbiological and physical processes contributed to more carbon sequestration during\nthe LGM) is independent of the choices of model simulation results (Line 124).\n116 – Fig S8  S6\nSupplementary figures were reordered and references to them were checked.\n138-139 – might it be possible to give some feel for how much is physical vs\nbiological based on the O2 and PO4 signals themselves? i.e. O2 has a more striking\nchange than PO4 – might this point to an important role for physical processes?\nWe think in general it is difficult to tease out contributions of biological and physical\nprocesses by looking at in-situ properties, because of the large influence of organic\nmatter respiration on them.\n140 – I think jumping straight to DIC is a bit unclear. More straightforward would be\nclearer to say biologically unutilized phosphate and associated DIC.\nWe now adopt this suggestion throughout the manuscript.\n148 – “declined … ratios” – unclear, rephrase\nRephrased.\n150-153 – there’s some important nuance that I think should be added here (and\nbelow), in that biological export production in the PAZ is lower at the LGM compared\nto the late Holocene, so increased biological usage would need to be relative to\nsupply. Many readers not super familiar with the field would read this sentence as\nindicating greater biological productivity, rather than utilisation relative to supply.\nDiscussion here was reorganized to emphasize that it is the nutrient utilization that\ndrove changes in carbon sequestration, and the lowered PAZ export productivity is\nalso mentioned as a caveat (Line 182).\n154-155 – I don’t think “storage” is quite the right word here. Storage suggests\ncarbon at depth, whereas these preformed signals are I think better linked to\n“uptake” or “reduced outgassing” or “improved nutrient and carbon usage efficiency”.\nThis could then lead to net improvement in deep storage; I think this nuance should\nbe mentioned. This is particularly the case given the comparison to nitrogen isotopes\nthat follows – these explicitly track surface utilisation efficiency, rather than deep\nstorage.\nChanged to “uptake” (Line 175).\nWe also expand the discussion about how our PO * reconstruction supports more deep-\n4\nwater carbon storage by surface Southern Ocean carbon uptake (Line 181).\n162 – again slightly confusing sentence; at least needs an “improved” in front of\nefficiency\n163 – maybe add “more directly” between “also” and “demonstrates”\nRevised.\n164-167 – one key concern here is on the sensitivity of [CO32-]as to DIC change\n(which I think is assessed based on modern ocean data) and how this would change\nunder overall higher ocean ALK and DIC at the LGM?\nThere are two related issues:\nThe first is that authors are using higher ALK and DIC based on model results and\nwhile this will not alter the structure of the deglacial change (as discussed above and\nin Figure S7), it will impact the value calculated here, and I think this should be\nacknowledged.\nWe acknowledged that the estimated DIC change is partially dependent on the global\nalkalinity derived from model simulations, which has little effect on our conclusion (line\n193).\nThe second is that if the ocean has higher overall ALK and DIC, it seems the\nrelationship between [CO32-]as will likely be different. I have tried looking in to the\nreferences given here about how this sensitivity was established but couldn’t find an\neasy answer to this and I think it should be as least briefly discussed.\nSensitivities of carbonate system to physical conditions, phosphate, and DIC are\nsimilar as shown by comparing results using preindustrial and LGM data sets (Yu et\nal., 2019). That is, global ALK changes have negligible impacts on the [CO 2-] vs\n3 as\nDIC sensitivities. We’ve clarified this point in the supplementary material (Line 415).\n167 – “Different from PO4*” – there’s the need for a bit more nuance here, as you’ve\ndiscussed above how physical processes can also influence PO4* via O2\ndisequilibrium, so this sentence jars a bit compared to the text above. I think the\nresolution to this point is your argument that at least some of that O2 disequilibrium\nmay ultimately be derived from biological utilisation, and it would be helpful to\nreference this briefly again here.\nThe commented discussion was reorganized. We now say that PO * is\n4\ndominantly/largely driven by biological processes (Line 84). We’ve also clarified that\nDIC changes stoichiometrically associated with PO * are the upper limited of\n4\nbiologically-driven DIC uptake, should there be any effect of physical processes on\nPO * via changing the preformed [O ]. Therefore, the larger DIC change associated\n4 2\nwith [CO 2-] than PO *supports a physically-driven component in air-sea CO changes (Line\n3 as 4 2\n195).\nThe other important point here relates to the comment above, which is that a\nsubstantial portion of the [CO32-]as comes not from within ocean biological or\nphysical processes, but from whole ocean carbon cycle change, so doesn’t fit neatly\nwithin this biological vs physical partition. This should be explicitly discussed.\nIn terms of the procedures to calculate [CO 2-] , we must use [CO 2-], [PO 3-], TSP,\n3 as 3 4\nand global alkalinity changes. After the calculation, [CO 2-] reflects air-sea\n3 as\nexchange component CO changes, attributions of which can be partitioned into\n2\ninfluences from biological and physical processes.\nHere we use the LH-LGM deep-water [CO 2-] change (Fig. X2) to illustrate the\n3\nrelation between carbonate system changes caused by the whole ocean alkalinity-\nDIC changes related to CaCO input/output imbalance and by air-sea exchange\n3\neffects. If we consider the LH-to-LGM transition (that is backwards in time), reduced\nCaCO burial during the transition would raise alkalinity and DIC at 2:1 along the\n3\ntrend defined by LH-LGM’ (Fig. X2). [Note that from the LGM to the LH, increased\ndeep-sea CaCO preservation is observed (Cartapanis et al., 2018). Thus, CaCO\n3 3\nburial would be reduced from the LH to the LGM].\nWithout air-sea CO exchange, in-situ [CO 2-] would have been raised (LGM’). In\n2 3\ncontrast, the LGM [CO 2-] is observed to be lower than the LH level (LGM).\n2\nsequestration (LGM’-LGM). When calculating the magnitude of the DIC change\n(LGM’LGM), the global alkalinity change is required; in other words, the global ALK\nchange affect the length of the “LGM’LGM” vector. However, the DIC change\n(LGM’LGM) must be caused by air-sea exchange which can be classified into\nchanged due to physical and biological processes.\nFig. X2. Deep-water [CO 2-] changes in the ALK- DIC space.\n3\n168 – “combines”\nFixed.\n170 – again as above, a chunk of this is from whole ocean change and this should\nbe specified.\nSee our response above and Fig X2.\n182 – “contribution”\n194 – “changes in”\n202 – “only enhance CO2 outgassing in the PAZ negligibly”  “but had minimal\nimpact on CO2 outgassing”\nFixed.\n203-205 – confusing as written – seems to imply that warming, which lowers\nsolubility, limits outgassing! I think you are trying to say that it stays fairly cold, which\nmeans that there’s not much solubility-driven change\nThis sentence has been revised to “the lack of strong CO outgassing resolvable by our\n2\n[CO 2-] reconstruction might be related to minor surface warming in the PAZ and deep\n3 as\nSouthern Ocean during early HS1(Stewart et al., 2023), which has been shown by model\nsimulations to be crucial to modulate physically-driven CO outgassing(Khatiwala et al.,\n2\n2019; Menviel et al., 2018)” Line (234).\n206 – here and elsewhere “late HS1” is I think neater than “the late HS1” – but feel\nfree to ignore\nChanged throughout the manuscript following the advice.\n209 – as above, is this really minor if O2 change is much bigger than PO4 change?\n213 – again, I worry about sensitivity of this DIC value to higher ALK and DIC, as\nmentioned above\nThese were addressed in responses to previous comments.\n217 – the timing of this change in [CO32-] also fits carbon release during the deep\nconvection pulse suggested by Li et al. (2020, Science Advances)\nWe’ve cited Li et al. (2020) to suggest enhanced Southern Ocean convection as a\npotential physical process leading to CO outgassing (Line 300).\n2\n223-224 – again, thinking about O2 in isolation likely useful here – high NSW input\nwould be unlikely to drive low O2. The low O2 also nicely explains the gap in deep\nsea corals at this time (see Stewart et al., 2021, Paleo)\nWe’ve used the observation from Stewart et al. (2021) to support minimal increase in\nthe NSW contribution to the deep Southern Ocean during the ACR (Line 266).\n230 – why “possibly” given argument above – “likely” seems more likely!\n241 – remove “in the both”\n258 – 8 out of 40 would be better described as “minor” than “negligible”\n353 – ref xx!\n375 – unclear as written\nFixed.\n380 – give the rationale for this (I went digging through the previous papers and\nfound it but would be handy to mention briefly here)\nThe rationale has now been added (Line 448).\n484 – foram data or water?\nThis is now specified that the data are based on core-top reconstructions (Line 552)\nFig S2 – would be helpful to also see the Mg data\nMg/Ca-based SST was added to this figure.\nReferences\nCartapanis, O., Galbraith, E. D., Bianchi, D., & Jaccard, S. L. (2018). Carbon burial\nin deep-sea sediment and implications for oceanic inventories of carbon and\nalkalinity over the last glacial cycle. Climate of the Past, 14(11), 1819-1850.\n10.1029/2019cp-14-1819-2018\nKhatiwala, S., Schmittner, A., & Muglia, J. (2019). Air-sea disequilibrium enhances\nocean carbon storage during glacial periods. Sci Adv, 5(6), eaaw4981.\n10.1126/sciadv.aaw4981\nLi, T., Robinson, L. F., Chen, T., Wang, X. T., Burke, A., Rae, J. W. B., et al. (2020).\nRapid shifts in circulation and biogeochemistry of the Southern Ocean during\ndeglacial carbon cycle events. Sci Adv, 6(42). 10.1126/sciadv.abb3807\nMenviel, L., Spence, P., Yu, J., Chamberlain, M. A., Matear, R. J., Meissner, K. J., &\nEngland, M. H. (2018). Southern Hemisphere westerlies as a driver of the\nearly deglacial atmospheric CO2 rise. Nature Communications, 9(1).\n10.1038/s41467-018-04876-4\nStewart, J. A., Li, T., Spooner, P. T., Burke, A., Chen, T., Roberts, J., et al. (2021).\nProductivity and Dissolved Oxygen Controls on the Southern Ocean Deep‐\nSea Benthos During the Antarctic Cold Reversal. Paleoceanography and\nPaleoclimatology, 36(10). 10.1029/2021pa004288\nStewart, J. A., Robinson, L. F., Rae, J. W. B., Burke, A., Chen, T., Li, T., et al. (2023).\nArctic and Antarctic forcing of ocean interior warming during the last\ndeglaciation. Sci Rep, 13(1), 22410. 10.1038/s41598-023-49435-0\nYu, J., Menviel, L., Jin, Z. D., Thornalley, D. J. R., Foster, G. L., Rohling, E. J., et al.\n(2019). More efficient North Atlantic carbon pump during the Last Glacial\nMaximum. Nat Commun, 10(1), 2170. 10.1038/s41467-019-10028-z"}]

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