# Varied contribution of the Southern Ocean to deglacial atmospheric CO2 rise

Andrew Moy, Martin Palmer, William Howard, Jelle Bijma, Matthew Cooper, Eva Calvo, Carles Pelejero, Michael Gagan, Thomas Chalk

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Andrew Moy, Martin Palmer, William Howard, Jelle Bijma, Matthew Cooper, et al. Varied contribution of the Southern Ocean to deglacial atmospheric CO2 rise. Nature Geoscience, 2019, 12 (12), pp.1006-1011. 10.1038/s41561-019-0473-9. hal-04702501

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# Varied contribution of the Southern Ocean to deglacial atmospheric CO₂ rise

Andrew D. Moy<fnref n="1" />,²*, Martin R. Palmer³, William R. Howard⁴, Jelle Bijma<fnref n="1" />,⁵, Matthew J. Cooper<fnref n="1" />,⁳, Eva Calvo<fnref n="1" />,⁶, Carles Pelejero⁶,⁷, Michael K. Gagan⁸,⁹ and Thomas B. Chalk³

Glacial-interglacial changes in atmospheric CO₂ are generally attributed to changes in seawater carbon chemistry in response to large-scale shifts in the ocean's biogeochemistry and general circulation. The Southern Ocean currently takes up more CO₂ than any other and it is likely to have played a crucial role in regulating past atmospheric CO₂. However, the physical, biological and chemical variables that control ocean-atmosphere CO₂ exchange during glacial-interglacial cycles are not completely understood. Here we use boron isotopes and carbon isotopes in planktonic foraminifera and an alkenone-based proxy of temperature to reconstruct seawater pH and CO₂ partial pressure in sub-Antarctic surface waters south of Tasmania over the past 25,000 years, and investigate the mechanisms that regulate seawater CO₂. The new record shows that surface waters in this region were a sink for atmospheric CO₂ during the Last Glacial Maximum. Our reconstruction suggests changes in the strength of the biological pump and the release of deep-ocean CO₂ to surface waters contributed to the last deglacial rise in atmospheric CO₂. These findings demonstrate that variations in upwelling intensity and the distribution of Southern Ocean water masses in this sector played a key role in regulating atmospheric CO₂ during the last glacial-interglacial cycle.

Atmospheric CO₂ concentrations range from minima of ~180ppmv during glacial intervals to maxima of ~280–300ppmv during interglacial intervals<fnref n="1" />. The ocean is the largest active carbon reservoir; hence, changes in ocean-atmosphere interactions, ocean circulation and marine biogeochemistry probably played major roles in modulating glacial-interglacial (GIG) atmospheric CO₂ variability²⁻⁵. Despite this understanding of the ocean's overall role in modulating atmospheric CO₂, information on the temporal and spatial evolution of ocean source-sink behaviour remains scarce. Determining the factors that control ocean-atmosphere CO₂ dynamics requires an understanding of the physical and biological processes that affect the state of CO₂ in the ocean, and is of key importance to validate the different mechanisms used in geochemical models to explain GIG variations in atmospheric CO₂ (for example, Sarmiento and Toggweiler⁶, Sigman and Boyle⁷ and Toggweiler et al.⁸). Changes in these processes in the Southern Ocean are hypothesized to have been crucial⁸⁻¹⁷. However, proxy histories of past GIG changes in surface-ocean CO₂ partial pressure (pCO₂) are limited to one record for the Atlantic sector of the Southern Ocean¹³ that only captures part of the GIG change in atmospheric CO₂. Our study focuses on the Indo-Pacific sector of the Southern Ocean south of Australia and we present a Southern Ocean surface-ocean carbonate chemistry reconstruction that allows the quantification of the ocean-atmosphere CO₂ exchange that accompanies the full GIG rise in atmospheric CO₂.

We analysed δ¹⁸O, δ¹³C, δ¹¹B and shell weight for the planktonic foraminifer *Globigerina bulloides*, δ¹⁸O and δ¹³C in benthic foraminifera and alkenone unsaturation ratios in sediment core MD972106 (Fig. 1 and Methods) and used these proxies to reconstruct surface water pH and CO₂ partial pressure (pCO₂) over the past 25kyr in the Southern Ocean region south of Tasmania.

Long-term climatological studies¹⁸ show that the core site is positioned at the southern edge of the Sub-Tropical Front (STF), which marks the northern boundary of the Southern Ocean and separates the Sub-Tropical Zone and Sub-Antarctic Zone (SAZ) surface waters (Fig. 1 and Supplementary Fig. 1). As a result, the core site is located just north of the high nutrient, low chlorophyll area of the Southern Ocean¹⁹ and the seafloor there is bathed by Lower Circumpolar Deep Water. Sea-surface temperature (SST) reconstructions for the Late Quaternary indicate that the STF migrated north of its present position during glacial periods and was at its most northerly position at the Last Glacial Maximum (LGM)²⁰. Therefore, we infer that the surface waters over MD972106 during the LGM had SAZ properties.

## The GIG proxy record

Down-core variations in *G. bulloides* δ¹⁸O exhibit a GIG amplitude of ~2‰ (Fig. 2a). The GIG change in global ocean δ¹⁸O is estimated to have been ~1.1‰ (ref. ²¹), hence the ~2‰ shift in the *G. bulloides* δ¹⁸O record includes a SST (and carbonate ion effect²²) component, which we account for here using alkenone-derived SSTs. The *G. bulloides* average δ¹³C values for the Holocene and LGM (~18–22 thousand years ago (ka)) are similar, but the GIG transition displays δ¹³C minima at ~12ka and during the Holocene at ~7.4ka (Fig. 2b). *G. bulloides* shell weights are heavier during glacial times (Fig. 2c). There is also a clear difference in *G. bulloides* δ¹¹B values, with Holocene δ¹¹B about 2‰ lower than that in the LGM (Fig. 2d). Alkenone-based SST estimates show a GIG change of ~4.1°C with a minimum of 8.2°C at ~21.4ka, followed by postglacial warming to 13.9°C at ~11ka and cooling to ~13°C over the past ~6kyr (Fig. 2e). The most recent samples (~2.6 to 2ka) show deviations in δ¹⁸O, shell weight and δ¹¹B compared to earlier Holocene values.

<footnote n="1">¹Australian Antarctic Division, Kingston, Tasmania, Australia. ²Antarctic Climate and Ecosystems Cooperative Research Centre, University of Tasmania, Hobart, Tasmania, Australia. ³Ocean and Earth Science, University of Southampton, Southampton, UK. ⁴Climate Change Institute, The Australian National University, Canberra, Australian Capital Territory, Australia. ⁵Marine Biogeosciences, Alfred-Wegener-Institut Helmholtz-Zentrum für Polar- und Meeresforschung, Bremerhaven, Germany. ⁶ICM-CSIC, Institut de Ciències del Mar, Barcelona, Spain. ⁷ICREA, Institució Catalana de Recerca i Estudis Avançats, Barcelona, Spain. ⁸School of Earth and Environmental Sciences, The University of Queensland, Brisbane, Queensland, Australia. ⁹Research School of Earth Sciences, The Australian National University, Canberra, Australian Capital Territory, Australia. *e-mail: Andrew.Moy@aad.gov.au</footnote>
<figure><img src="image_2.png" /><figcaption>**Fig. 1** | Location of sediment core MD972106 and modern positions of the Southern Ocean water masses and fronts. MD972106 (blue dot) is located at 45°09′S, 146°17′E (water depth 3,310 m). SAF, Sub-Antarctic Front; PF, Polar Front; SB, southern boundary of the Antarctic Circumpolar Current (ACC); NZ, New Zealand. Arrows indicate the ACC direction. The dashed red line shows the WOCE SR3 repeat hydrographic transect between Tasmania and Antarctica along ~140°E. Figure adapted from Trull et al.⁵¹, John Wiley and Sons.</figcaption></figure>

Downcore G. *bulloides* δ¹⁸O, δ¹¹B and shell weights co-vary with each other (Fig. 2) and with atmospheric CO₂ over the past 25 kyr (Fig. 3b). This co-variation is interpreted here to arise from the GIG evolution of the surface water temperature and chemistry²²⁻²⁷. As MD972106 was recovered at a depth above and close to the modern calcite saturation horizon (Supplementary Fig. 2), it is important to consider the possible effects of postdepositional calcite dissolution on shell weights²⁶,²⁸, particularly when considering GIG lysocline changes versus primary calcification driven by surface water carbonate ion concentration ([CO₃²⁻]). Holocene core-top G. *bulloides* shell weights indicate that the foraminiferal lysocline at the South Tasman Rise is at ~3,600 m water depth²⁹. Carbonate preservation indicators, percent calcium carbonate and percent whole foraminifera (Supplementary Fig. 3) suggest that seafloor calcite dissolution did not play a significant role in changing shell weights over the past 25 kyr at this site.

## Reconstructing seawater pH and $pCO_2$

Past variations in sea-surface water pH at the MD972106 core site were calculated from the δ¹¹B–pH relationship for G. *bulloides*³⁰ and from the alkenone SST record (Methods). The estimated pH shows a decline from LGM values of ~8.38 to Holocene values of ~8.16 (Fig. 3a). By comparison, the modern anthropogenic-influenced surface water pH at the MD972106 site was 8.11 in 2001 CE (Supplementary Note 1). Variations in surface water $pCO_2$ (Fig. 3b) over the past 25 kyr can be estimated using the SST alkenone record, the δ¹¹B surface water pH record and estimates of either alkalinity or dissolved inorganic carbon (DIC) (Methods). There is a close agreement between $pCO_2$ calculated using estimates of pH combined with either estimates of alkalinity (Method 1) or DIC (Method 2) (Fig. 3b). Methods 1 and 2 both use pH estimated from G. *bulloides* δ¹¹B, and hence any differences result from the assumptions that underlie

<figure><img src="image_3.png" /><figcaption>**Fig. 2** | δ¹⁸O, δ¹³C, shell weight and δ¹¹B for the planktonic foraminifera, G. bulloides and SST estimates from alkenones in sediment core MD972106.

**a–d**, Measured G. *bulloides* δ¹⁸O (with y axis reversed) (**a**), δ¹³C (**b**), shell weight (**c**) and δ¹¹B (**d**). **e**, SST estimates from alkenone unsaturation ratios ($U_{37}^{K}$). The age model was constructed using calibrated ¹⁴C ages (black triangles, Methods). Error bars represent 1σ uncertainties for δ¹⁸O, δ¹³C and shell weight, and 2σ uncertainties for δ¹¹B based on replicate measurements (Methods). Error bar on the SSTs represent 2σ analytical uncertainties for replicate measurements on a homogeneous laboratory standard (Methods).</figcaption></figure>

estimates of the alkalinity and DIC parameters. These calculations show that surface water $p_{CO_2}$ rose from ~165 ppmv at the LGM to ~280 ppmv in the early Holocene (Fig. 3b). The interval ~6.3–4.2 ka is marked by an elevated $p_{CO_2}$ of ~315–360 ppmv that decreased to ~270 ppmv at ~2 ka (Fig. 3b). The 2001 CE surface water $p_{CO_2}$ at this site was ~330 ppmv, broadly consistent with the known anthropogenic rise in atmospheric CO₂ (Supplementary Note 1).

The difference between the estimated $p_{CO_2}$ of the surface water and the atmosphere ($\Delta p_{CO_2}$) can be calculated by using ice core CO₂ data from Antarctica¹ (Fig. 3b). Estimated $\Delta p_{CO_2}$ (Figs. 3c and 4e) indicates that surface waters near MD972106 were a net sink for atmospheric CO₂ during the LGM ($\Delta p_{CO_2} = -25$ ppmv) and during the GIG transition until ~12 ka, when surface waters became a net source of CO₂ during the early Holocene ($\Delta p_{CO_2} = 15$ ppmv). Between ~6.3 and 4.2 ka the surface waters were a strong net source for CO₂ ($\Delta p_{CO_2} \approx 50–90$ ppmv), before returning to a weak net sink ($\Delta p_{CO_2} \approx -6$ ppmv) at ~2 ka (Figs. 3c and 4e). In addition, past
surface seawater [CO₃²⁻] calculated from estimates of pH, SST and alkalinity (and DIC) show [CO₃²⁻] falling from ~265 µmol kg⁻¹ in the LGM to ~190 µmol kg⁻¹ for most of the Holocene (Fig. 5a). The anthropogenically perturbed surface water [CO₃²⁻] at this site is ~183 µmol kg⁻¹, whereas the pre-industrial [CO₃²⁻] is estimated to be ~212 µmol kg⁻¹ (Supplementary Note 2), which is within the uncertainty of the most recent Holocene [CO₃²⁻] (Fig. 5a).

The shell weight–[CO₃²⁻] relationship²⁵ was also applied to the MD972106 data to reconstruct [CO₃²⁻] for surface waters over the past 25 kyr (Fig. 5b). The close agreement between the estimated surface water [CO₃²⁻] derived from these two independent proxies (G. *bulloides* δ¹¹B values and shell weight) reinforces our findings that the δ¹¹B of G. *bulloides* yields reliable reconstructions of seawater pH.

## Southern Ocean contribution to atmospheric CO₂ change

Our reconstruction is consistent with the hypothesis that the GIG modulation of the Southern Ocean carbon balance probably arose from changes in SST, biogeochemistry and ocean circulation, specifically: (1) colder LGM seawater increased the CO₂ solubility, (2) a stronger ‘biological pump’ during the LGM reduced the DIC in surface waters and (3) surface water pH and [CO₃²⁻] were higher during the LGM.

Two hypotheses that involve the Southern Ocean have been offered to explain GIG cycles in atmospheric CO₂. The first invokes an increased strength of the biological (or ‘soft-tissue’) pump in the high-latitude Southern Ocean during the LGM⁶,⁹ in response to an increase in nutrient utilization, possibly through iron fertilization³¹, in sub-Antarctic surface waters. This is supported by studies that indicate export production in the area between the STF and Polar Front were higher during the LGM than during the Holocene¹²,³²⁻³⁴. The second invokes a decreased exchange between surface waters and the deep-ocean south of the Polar Front³⁵⁻³⁷. The increased stratification reduces the upwelling and the exchange of nutrient and CO₂-rich deep waters with the surface ocean. This is supported by increased nutrient utilization, as shown by δ¹⁵N records³⁵, reduced ventilation of the deep Southern Ocean³⁸ and subsequent increase in the ocean CO₂ storage⁵,¹⁷ during the LGM. No single mechanism³⁹ can explain the full glacial draw-down in atmospheric CO₂, although a combination of mechanisms has been proposed²,¹²,¹⁴.

## Asymmetry of deglacial change in the Southern Ocean

Our full deglacial SAZ pH and $p_{CO_2}$ reconstruction can be compared to a reconstruction of late deglacial pH and $p_{CO_2}$ in the Atlantic sector of the SAZ of the Southern Ocean¹³ to provide further insights into the sub-Antarctic role in CO₂ change. That reconstruction of surface ocean $p_{CO_2}$ for the Atlantic sector of the SAZ, which spans ~16–2 ka and also utilizes δ¹¹B and δ¹³C of G. *bulloides*¹³, was undertaken at a similar latitude to MD972106. That study, although it did not span the full deglacial amplitude of CO₂ change, suggested surface waters in that sector were in approximate equilibrium with the atmosphere at ~16 ka and became a strong net source of CO₂ to the atmosphere (Δ$p_{CO_2}$ of ~50 ppmv (Fig. 4e)) by ~15 ka, before declining intermittently to reach an approximate atmospheric equilibrium at ~4 ka. This pattern was interpreted to reflect an enhanced upwelling of deep waters in the Atlantic sector of the SAZ during the deglaciation.

At our site south of Tasmania, surface waters were a net sink for atmospheric CO₂ (Δ$p_{CO_2}$ = -25 ppmv) during the LGM (Figs. 3c and 4e). The northward migration of the STF at the LGM²⁰ resulted in an expansion of the Southern Ocean water masses, and coincided with an increased delivery of bioavailable dust-borne iron to the ocean³¹; both of these processes also lead to the expansion of the zone of high biological production in the Southern Ocean. At the same time, the stratification of the deep Southern Ocean during the LGM reduced the contribution of the upwelling of dissolved

<figure><img src="image_4.png" /></figure>

**Fig. 3** | Reconstructed surface water pH and $p_{CO_2}$ at the MD972106 site and atmospheric CO₂ over the past 25,000 years. **a**, Surface water pH (with y axis reversed). **b**, $p_{CO_2}$ calculated from pH, alkenone SST and alkalinity (Method 1, solid blue circles) and from pH, alkenone SST and DIC (Method 2, open blue circles), and atmospheric CO₂ record (Antarctic ice core CO₂ composite¹, green diamonds). **c**, Δ$p_{CO_2}$ is the difference between the reconstructed surface water $p_{CO_2}$ and atmospheric $p_{CO_2}$ (ref. ¹) (Method 1, solid black squares; Method 2, open black squares). Error bars in **a**, **b** and **c** represent the 95% confidence intervals around the Monte Carlo mean, where 2σ uncertainties on the individual input variables and other carbonate system parameters are propagated via the Monte Carlo simulation (n = 10,000) in the program R (Methods).

CO₂ to surface waters. At the site of MD972106, the temperature and [CO₃²⁻] adjustments on G. *bulloides* δ¹³C (δ¹³C<sub>adj</sub>) (Methods gives details) and the δ¹³C gradient between planktonic and benthic foraminifera (Δδ¹³C gradient) were largest during the LGM (Fig. 4a,c), which indicates high primary productivity⁴⁰. Similarly, alkenone concentrations (Fig. 4d) were highest during the LGM and provide qualitative information on the past productivity of coccolithophorid algae⁴¹. Together, the alkenone concentrations (Fig. 4d) and Δ$p_{CO_2}$ (Fig. 4e) show that CO₂ degassing was suppressed when there was maximum biological production in this region. Our data suggest that the biological pump was stronger during the LGM (that is, the DIC gradient between the surface and deep ocean increased), possibly through iron fertilization³¹,⁴², and steadily decreased over the last deglaciation. Changes in the strength of the biological pump could also be related to changes in circulation⁴³, via a more isolated deep ocean, in particular south of the Polar Front³⁵.

The δ¹³C of G. *bulloides* in the Atlantic sector of the SAZ decreased by ~1.5‰ from the LGM to the onset of maximum CO₂ degassing at ~11 ka (not shown)¹³, and the MD972106 site shows a smaller δ¹³C decrease of ~0.3‰ (Fig. 2b) over the same period. The inferred increase in upwelling in the Atlantic sector is not reflected in the reconstructed Δ$p_{CO_2}$ record at MD972106 (Figs. 3c and 4e).
<figure><img src="image_5.png" /><figcaption>**Fig. 4** | Planktonic and benthic foraminiferal $\delta^{13}$C, $\Delta\delta^{13}$C gradient, alkenone concentrations and $\Delta p_{CO_2}$ over the past 25,000 years.

a, $\delta^{13}$C$_{adj}$ for G. bulloides. b, Benthic $\delta^{13}$C for *Cibicidoides* spp. c, $\Delta\delta^{13}$C gradient ($\delta^{13}$C planktonic$_{adj}$ – $\delta^{13}$C benthic foraminifera); the horizontal dashed line shows the average Holocene $\Delta\delta^{13}$C gradient. d, Alkenone concentrations. e, $\Delta p_{CO_2}$ is the difference between surface water $p_{CO_2}$ and atmospheric $p_{CO_2}$ (ref. 1) as in Fig. 3c. Also shown in e is the calculated $\Delta p_{CO_2}$ for the Atlantic sector SAZ (ref. 13) (sediment core PS2498-1 at 44° 09' S, 14° 13' 48" W, 3,783 m water depth; the grey open diamonds and error bars are the 95% uncertainty bounds). In a, $\delta^{13}$C values for G. bulloides are temperature and [$CO_3^{2-}$] adjusted (Methods). The error bars in a and b represent $1\sigma$ uncertainties based on replicate measurements, the error bar in d represent $2\sigma$ uncertainties for replicate extractions and measurements on a homogeneous laboratory standard and the error bars in e represent 95% confidence intervals around the Monte Carlo mean, based on $2\sigma$ uncertainties on the individual input variables and on other carbonate system parameters propagated via the Monte Carlo simulation (n=10,000) in the program R (Methods).</figcaption></figure>

Instead, the surface waters at MD972106 during the last deglaciation remained a net sink for CO₂ (average $\Delta p_{CO_2} = -25$ ppmv (Figs. 3c and 4e)). The steadily decreasing $\delta^{13}$C$_{adj}$ (Fig. 4a) and $\Delta\delta^{13}$C gradient (Fig. 4c) during the deglaciation is consistent with a progressive increase in the contribution of upwelling and advection of ‘old’ CO₂-rich (and ¹³C-depleted) Circumpolar Deep Water (CDW) to sub-Antarctic surface waters. Nevertheless, surface waters remained a net sink for atmospheric CO₂ during this period, which suggests that productivity (draw-down of CO₂) probably kept pace with any

<figure><img src="image_6.png" /><figcaption>**Fig. 5** | Reconstructed surface water carbonate ion concentration.

a, [$CO_3^{2-}$] calculated from $\delta^{11}$B-pH estimates, alkenone SST and alkalinity (Method 1, solid green circles) and from $\delta^{11}$B-pH estimates, alkenone SST and DIC (Method 2, open green circles). b, [$CO_3^{2-}$] calculated from shell weight data using the shell weight-[$CO_3^{2-}$] relationship²⁵. Error bars in a represent 95% confidence intervals around the Monte Carlo mean, with the $2\sigma$ uncertainties on the individual input variables and other carbonate system parameters propagated via the Monte Carlo simulation (n=10,000) in the program R (Methods).</figcaption></figure>

augmented nutrient supply and compensated for any increased CO₂ from upwelling in the MD972106 region. Alkenone concentrations (a proxy for coccolithophorid algae primary production (Fig. 4d)) show that enhanced biological productivity was maintained during this interval. There is compelling evidence for a build-up of carbon in the deep Southern Ocean towards the LGM¹⁰,¹⁵,¹⁷, and subsequent release and ventilation³⁸ of this deep carbon store during the last deglaciation, which is consistent with the rise in atmospheric CO₂ (refs. ¹⁵,¹⁷).

The interval between ~6.3 and 4.2 ka, when surface waters at MD972106 were a net source of CO₂ ($\Delta p_{CO_2}$ of ~50–90 ppmv (Figs. 3c and 4e)), coincides with a period of strengthened Southern Hemisphere westerly winds (SWWs) inferred from an analysis of terrestrial records⁴⁴,⁴⁵. Changes in the latitude and intensity of the SWWs are potential drivers of changes in atmospheric CO₂, as suggested by modelling⁸ and opal burial rates⁴⁶, and are in agreement with past changes in equatorial and Antarctic seawater carbon isotope variations⁴⁷. Intensification of the SWWs increases the contribution of upwelled CO₂-rich CDW (with a low $\delta^{13}$C) to surface waters within the Antarctic Zone. The $\delta^{13}$C$_{adj}$ of G. bulloides at MD972106 reaches a minimum at ~7.4 ka (Fig. 4a) and the $\Delta\delta^{13}$C gradient falls below the Holocene average gradient (Fig. 4c). Our data are thus consistent with the hypothesis that the net CO₂ source during this period resulted from an increased upwelling of CO₂-rich CDW within the Antarctic Zone that advected northward to SAZ surface waters by the strengthened SWWs⁴⁴,⁴⁵.

The deviations in the most recent G. bulloides samples ($\delta^{18}$O, shell weight and $\delta^{11}$B (Fig. 2)) stand in contrast to earlier Holocene values, and reflect changing surface water CO₂ dynamics. Surface waters at ~2 ka were a weak CO₂ sink ($\Delta p_{CO_2} = -6$ ppmv) that probably resulted from a changing balance between a stronger ‘biological pump’ and a corresponding decrease in the contribution and/or intensity of upwelling and advection of CO₂-rich CDW. The increasing $\delta^{13}$C$_{adj}$ and $\Delta\delta^{13}$C gradient (Fig. 4a,c) and the corresponding increase in alkenone concentrations (Fig. 4d) from ~4.1 to 2 ka are indicative of increased primary productivity, which corresponds to a decrease in the SWWs strength from ~4 to 1 ka (refs. ⁴⁴,⁴⁵).
Our findings demonstrate the interplay and changing balance between removal of CO₂ by biological production in surface waters and CO₂ outgassing via circulation changes during the last GIG rise in atmospheric CO₂. The modern Southern Ocean shows considerable zonal variability in annual and seasonal air-sea CO₂ flux⁴⁸, so such spatial heterogeneity in ocean-atmosphere CO₂ exchange in the past is perhaps not surprising. This study provides evidence that the CO₂ dynamics of Holocene surface waters in the Southern Ocean are dependent on competing factors such as the ocean's biological pump and changes in the location and intensity of the SWWs. The evolving relationship between atmospheric dynamics and carbon exchange is critically important to understand the role of Southern Hemisphere winds in modulating the uptake of anthropogenic CO₂ (refs. ⁴⁹,⁵⁰).

## Online content

Any methods, additional references, Nature Research reporting summaries, source data, statements of code and data availability and associated accession codes are available at https://doi.org/10.1038/s41561-019-0473-9.

Received: 21 December 2018; Accepted: 13 September 2019; Published online: 21 October 2019

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## Acknowledgements

This work was supported by the Australian Antarctic Division (AAS 4061) and the Australian Government's Cooperative Research Centres Programme through the Antarctic Climate and Ecosystems Cooperative Research Centre (ACE CRC). The boron isotope analyses were supported by the European Union 5th Framework Programme project 6C (Project ID: EVK2-CT-2002-00135 6C). E.C. and C.P. acknowledge Graham Logan and Geoscience Australia for providing analytical facilities for alkenone analyses. We thank the French Polar Institute and the crew of the RV *Marion Dufresne* for their efforts in recovering sediment core MD972106.

## Author contributions

A.D.M. and W.R.H. designed the study with input from M.R.P. and J.B. The manuscript was written by A.D.M., W.R.H., M.R.P. and J.B. with contributions from M.J.C., E.C., C.P., M.K.G. and T.B.C. Analysis and interpretation of the measurements were completed by A.D.M., W.R.H., M.R.P., J.B., M.J.C., E.C., C.P., M.K.G. and T.B.C. All authors contributed to improving the final manuscript.

## Competing interests

The authors declare no competing interests.

## Additional information

**Supplementary information** is available for this paper at https://doi.org/10.1038/s41561-019-0473-9.

**Correspondence and requests for materials** should be addressed to A.D.M.

**Peer review information** Primary Handling Editor(s): James Super, Rebecca Neely.

**Reprints and permissions information** is available at www.nature.com/reprints.

**Publisher's note** Springer Nature remains neutral with regard to jurisdictional claims in published maps and institutional affiliations.

© Crown 2019
## Methods

**Materials and methods.** Sediment core MD972106 (45°09' S, 146°17' E, water depth 3,310 m) is a 32 m long piston core collected using the CALYPSO Kullenberg corer on board the RV Marion Dufresne in 1997 (in this article, we present data from the top 2.4 m of the core) and was recovered from the northern flank of the South Tasman Rise. The MD972106 site is located at the southern edge of the STF (Fig. 1) and the modern bottom water that covers the South Tasman Rise is CDW, a mixture of high δ¹³C (low nutrient) North Atlantic Deep Water (or northern source deep water) and recirculated low δ¹³C (high nutrient) deep water from the Indian and Pacific Oceans (Supplementary Fig. 1).

MD972106 was sampled every 5 cm on board RV Marion Dufresne. Raw sediment samples were oven dried at 60 °C, and then disaggregated in distilled water at the ACE CRC, University of Tasmania. Water-saturated sediment samples were wet sieved through a >150 µm mesh. The dried >150 µm fraction was sieved to isolate the 300–355 µm fraction from which 50 or more planktonic foraminifera (G. bulloides) were picked. We selected G. bulloides because this species shows a seasonal range in δ¹⁸O values in this region of the Southern Ocean that is consistent with calcification in near-surface waters⁵², and comparisons between flux-weighted sediment trap and surface sediment δ¹⁸O values indicate the sedimentary record retains the seasonal isotopic imprint for G. bulloides⁵².

Each batch of foraminifera was ultrasonically cleaned in methanol and oven dried at 60 °C. Any broken tests were discarded. The cleaned samples were used to determine the average shell weights and isotopic compositions (δ¹⁸O, δ¹³C and δ¹¹B), with the shells undergoing further cleaning prior to the determination of their δ¹¹B values. Average shell weight, δ¹⁸O and δ¹³C measurements were conducted at 5 cm intervals. Determinations of δ¹¹B were made at a 20 cm resolution.

**Age model.** Age control for MD972106 is based on accelerator mass spectrometry radiocarbon (¹⁴C) dating of the planktonic foraminifer *Globorotalia inflata*. Five published accelerator mass spectrometry ¹⁴C dates for MD972106 (ref. ⁵³) were recalibrated to calendar ages using the Calib 7.1 program²⁴ with the Marine13 dataset⁵⁵. Ages were corrected for the local ¹⁴C reservoir correction (ΔR) for surface waters in this region of the Southern Ocean using a constant ΔR regional mean value (ΔR = 107 ± 13 yr) calculated for the Chatham Islands from the marine reservoir correction database⁵⁶. The ΔR (107 ± 13 yr) incorporates the measured ¹⁴C age of 560 ± 40 yr BP obtained from gastropods⁵⁷. Using a different and older ΔR for the ¹⁴C dates at 16.1 and 23.5 ka slightly changes the chronology, but this does not affect the conclusions of our study for the LGM and early deglacial periods. For example, using ΔR = 900 yr for the ¹⁴C dates at 16.1 and 23.5 ka changes the chronology by ~0.8–0.9 kyr, and using a ΔR = 900 yr for the ¹⁴C date at 16.1 ka and a ΔR = 1,500 yr for ¹⁴C date at 23.5 ka changes the chronology by ~0.8 to 1.4 kyr. The average sedimentation rate at MD972106 was ~9.3 cm kyr⁻¹ over the past 25 kyr.

**Planktonic foraminifera shell weights.** G. bulloides shell weights were determined by picking 50 or more whole individual shells from the 300–355 µm size fraction. The cleaned whole shells were counted and weighed on a microbalance (precision = 0.1 µg). The average shell weights were calculated by dividing the measured weight by the total number of whole foraminifera. Replicate shell-weight measurements were determined by selecting a separate aliquot of 50 or more shells from the same sample. The mean difference of the replicate shell weights was ±1.2 µg (1σ, n = 55).

**Stable isotope analyses.** Oxygen (δ¹⁸O) and carbon (δ¹³C) isotope ratios were determined for the planktonic foraminifera G. bulloides (300–355 µm), and the benthic foraminiferal taxa Cibicidoides spp. (>150 µm size fraction). Measurements of δ¹⁸O and δ¹³C were conducted on an automated individual-carbonate reaction Kiel Device coupled to a Finnigan MAT-251 isotope ratio mass spectrometer at the Research School of Earth Sciences, Australian National University. Calcite samples that weighed ~200 µg were reacted with 103% phosphoric acid at 90 °C to liberate sufficient CO₂ for isotopic analysis. The δ¹⁸O and δ¹³C values are reported as per mil (‰) deviations relative to the Vienna Pee Dee Belemnite (VPDB) standard. The results are normalized on the VPDB scale such that the NBS-19 calcite standard yields δ¹⁸O VPDB (−2.20‰) and δ¹³C VPDB (+1.95‰) and NBS-18 yields δ¹⁸O VPDB (−23.0‰) and δ¹⁴C VPDB (−5.0‰). The 2σ uncertainties for replicate in-run measurements of NBS-19 (n = 197) were ±0.05‰ and ±0.02‰ for δ¹⁸O and δ¹³C, respectively. The average s.d. for repeat measurements of two planktonic samples were ±0.17‰ for δ¹⁸O and ±0.09‰ for δ¹³C, and for benthic samples it was ±0.05‰ for δ¹⁸O and ±0.12‰ for δ¹³C.

Boron isotope (δ¹¹B) values were determined for the planktonic foraminifera G. bulloides (300–355 µm) at Southampton University using previously described methods⁵⁸⁻⁶⁰. Oxidative cleaning procedures followed those in Barker et al.⁶¹. We analysed ten G. bulloides (150–250 µg of calcite) for each in-run measurement and this provided ~2–3 ng of boron. The average 2σ in-run measurement precision was <0.2‰. The average agreement of repeat measurements on separate sample aliquots was ±0.4‰ (n = 7) and the average external 2σ precision was, at most, ±0.4‰. Foraminiferal cleaning procedures⁶¹ were verified for all the samples by Al/Ca ratios <100 µmol mol⁻¹, which provides a sufficient screen against clay contamination⁶².

**Palaeo-SSTs and alkenone concentrations.** Alkenones were analysed at Geoscience Australia following the methods of Calvo et al.⁶³ and, with them, SST estimates were obtained from the alkenone unsaturation ratio (U%K/37 index). Alkenone concentrations were quantified by using n-hexatriacontane as an internal standard. The s.d. for alkenone concentrations was 50 ng g⁻¹ (2σ, replicate extractions and measurements of a homogeneous laboratory sediment standard). Alkenone-derived SSTs were reconstructed using the U%K/37-SST relationship of Müller et al.⁶⁴, where U%K/37 = 0.033 × SST + 0.044 for annually averaged SSTs. U%K/37 measurements were completed at a 10 cm sample resolution. The s.d. for replicate measurements of a homogeneous laboratory standard was, at the most ±0.6 °C (2σ). The standard error of SST estimates from the U%K/37–SST calibration⁶⁴ was ±1.5 °C. Long-term climate data from the World Ocean Atlas 2013 (0.25° grid) gives a modern annual SST of 12.3 °C at the MD972106 locality, which is close to the most recent core-top sediment U%K/37-derived SST of 13.0 °C.

**Reconstruction of surface water $pCO_2$.** Seawater $pCO_2$ can be calculated from temperature, salinity and pressure, and knowledge of two out of the six main carbonate system parameters (free aqueous carbon dioxide ($CO_2(aq)$), bicarbonate ion ($HCO_3^-$), carbonate ion ($CO_3^{2-}$), hydrogen ion ($H^+$), DIC and total alkalinity). Accordingly, past changes in seawater $pCO_2$ can be calculated using estimates of seawater pH and either DIC or total alkalinity. Here we used an alkenone-based proxy for seawater temperature, planktonic foraminiferal δ¹¹B to estimate pH, planktonic foraminiferal δ¹³C to estimate DIC and salinity-derived seawater alkalinity to constrain the major parameters of the seawater carbonate system. Past variations in sea surface water pH at the site of MD972106 were calculated from a recently published δ¹¹B–pH relationship for G. bulloides³⁰ and the alkenone SST record. The δ¹¹B–pH relationship for G. bulloides³⁰ expands on the previously published δ¹¹B–pH relationship for the same species¹³. The δ¹¹B–pH relationship³⁰ is similar to and within error of a previously published δ¹¹B–pH relationship¹³. Even though the calibration uncertainty for Martínez-Boti et al.¹³ is considerably lower than that of Raitzsch et al.³⁰, we used the extended δ¹¹B–pH relationship data set for G. bulloides³⁰.

Variations in surface water $pCO_2$ (Fig. 3b) over the past 25 kyr were estimated using the SST alkenone record, the δ¹¹B surface water pH record and estimates of either alkalinity or DIC. One set of estimates was derived from δ¹¹B-based pH and seawater alkalinity, assuming that past variations in alkalinity were proportional to salinity, which can be estimated from past sea level⁵⁸,⁶⁵ (Method 1) (Fig. 3b). An alternative set of estimates was derived from δ¹¹B-based pH and foraminiferal δ¹³C-based DIC (Method 2) (Fig. 3b). Both approaches use a Monte Carlo method to calculate $pCO_2$ from pH, DIC, salinity and temperature. This method generates quasi-normally distributed errors around central values with 10,000 repetitions, and sea water carbonate system calculations using the R-based⁶⁶ seacarb package⁶⁷. 2σ uncertainties on the individual input variables and other carbonate system parameters were included to the end (δ¹¹B (±0.4‰), U%K/37-derived SST (±0.6 °C), salinity (±3 (practical salinity units), alkalinity (±200 µmol kg⁻¹) and DIC (±76 µmol kg⁻¹). A disequilibrium uncertainty of ±20 ppm (2σ) is also included. We note the importance of propagating all the uncertainties, but we also note that pH is by far the dominant variable on $pCO_2$ and that it is well constrained by δ¹¹B alone⁶⁸. Thus, we are very confident that the change in $pCO_2$ forcing over this time is much more robust than the $pCO_2$ change in terms of absolute uncertainty⁶⁸. The average $pCO_2$ difference between calculations using Methods 1 and 2 is ~7 ppm (Fig. 3b).

Surface water $pCO_2$ and hence ocean circulation and productivity exert an important control on atmospheric $CO_2$ levels. However, the difference between the two ($\Delta pCO_2$) is also affected by terrestrial components of the carbon cycle such as (re)growth of the biomass or the dynamics of (coastal) permafrost. Hence, more important than the absolute magnitude of $\Delta pCO_2$ is the change in this value, because this provides direct information as to whether a particular area of the ocean was more or less important in driving the direction of atmospheric change. Thus, although there are intervals in our record in which the analytical and/or age uncertainties preclude us from being certain whether our study area acted as source or sink of $CO_2$ to the atmosphere, we can be confident that the change in surface water $\Delta pCO_2$ levels in this area of the Southern Ocean made an important contribution to the last deglacial atmospheric $CO_2$ rise.

We used the total pH scale, equilibrium constants for the dissociation of carbonic acid from Dickson⁶⁹ and the boron isotope equilibrium constant from Klochko⁷⁰. Modern silicate and phosphate values were applied to the calculations, but they had very little effect on the final results⁷¹. The proxies recorded in G. bulloides and in coccolithophorid alkenones predominantly reflected austral spring conditions in the region, because this is the seasonal maximum of production for both taxa⁵²,⁷². Thus, the estimates of $pCO_2$, probably reflect austral spring conditions.

The calculated $pCO_2$ from boron isotope-based pH estimates and DIC (Method 2) utilizes an estimate of surface water DIC over time and requires a DIC estimate for the most recent sediment sample. The WOCE SR3 repeat hydrographic transect between 44° and 46°S (refs. ⁷³⁻⁷⁵) indicates a modern surface water DIC of ~2,061 µequiv. kg⁻¹ at this site, where G. bulloides is known to calcify²². Estimates of DIC are then varied over time as a function of G. bulloides δ¹³C, with the difference between modern surface and deep water DIC derived via the difference between planktonic and benthic foraminiferal δ¹³C (that is
ΔDIC<sub>SW-DW</sub> is proportional to Δδ<sup>13</sup>C<sub>plankton-benthic</sub>). The biological drivers of surface water δ<sup>13</sup>C are mainly modulated by organic carbon production; carbonate precipitation does not significantly fractionate surface water DIC. Therefore, any changes in the CaCO<sub>3</sub>:C<sub>organic</sub> rain ratio would have a minor effect on DIC. Studies of the mechanisms that modulate DIC in the surface waters in this region (for example ref. ⁷⁶) and of the Southern Ocean more broadly (for example, Sarmiento et al.⁷⁷) suggest that the CaCO<sub>3</sub>:C<sub>organic</sub> rain ratio is very low, with only a minor effect on surface DIC and surface δ<sup>13</sup>C<sub>DIC</sub>. Any changes to the rain ratio also affect the surface water pH, which we estimated through boron isotopes. Modern surface and deep water DIC are from the WOCE SR3 repeat hydrographic transect between 44° and 46° S (refs. ⁷³⁻⁷⁵), δ<sup>13</sup>C<sub>SW</sub> from King and Howard⁷⁸ and δ<sup>13</sup>C<sub>DW</sub> from Moy et al.⁵³.

**G. bulloides** δ¹³C<sub>adj</sub>. Temperature and carbonate ion concentration ([CO₃²⁻] have an additional, often neglected, effect on planktonic foraminiferal δ¹³C (refs. ²²,²⁴,⁷⁸). For Method 2, temperature and [CO₃²⁻] adjustments on the *G. bulloides* δ¹³C used to estimate DIC do not significantly change the *pCO₂* estimates. The average difference in *pCO₂* estimates calculated from (1) boron isotope-based pH estimates and DIC with the δ¹³C for *G. bulloides* not temperature and [CO₃²⁻] adjusted and (2) boron isotope-based pH estimates and DIC with the δ¹³C for *G. bulloides* temperature and [CO₃²⁻] adjusted is ~2 ppm (Supplementary Fig. 4).

*G. bulloides* δ¹³C values were adjusted for temperature using the δ¹³C disequilibrium-temperature relation 0.11‰ °C⁻¹ (ref. ²⁴). The δ¹³C disequilibrium-temperature²⁴ relation we applied here is similar to the relationship estimated from a Southern Ocean sediment trap field study²⁸. Temperature adjustments were made using the alkenone-based SST estimates in MD972106. *G. bulloides* δ¹³C values were also adjusted for the so-called carbonate ion ([CO₃²⁻]) effect using the relationship between δ¹³C disequilibrium and [CO₃²⁻] (−0.014‰ [CO₃²⁻]⁻¹) (ref. ²²). Estimates of [CO₃²⁻] are from surface water *pCO₂* calculations that use boron isotope-based pH estimates and estimates of seawater alkalinity (Method 1). The 'close agreement' between the [CO₃²⁻] outputs from Methods 1 and 2 (Fig. 5a) yields a similar adjustment to *G. bulloides* δ¹³C in terms of [CO₃²⁻] adjustments.

**Data availability**

All data are archived at the Australian Antarctic Data Centre (https://data.aad.gov.au/) and are publicly accessible at https://doi.org/10.26179/5d5df822acfc6.

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