Throughout the water column, the observed glacial Δ14C-range in our cores exceeds the modern and Holocene values by a factor of ∼5 and indicates strong age differences and therefore enhanced stratification of the intermediate and deep glacial South Pacific. Along its pathway in the global thermohaline circulation PDW is fed into circumpolar waters and constitutes today’s oldest water mass. The 14C-depleted PDW presently extends to 2,000–2,500 m north of the Chatham Rise close to New Zealand25 (Fig. 1b). From our transect, we are able to show that during the LGM, significantly 14C-depleted and aged water masses occupied depths between ∼2,000 and ∼4,300 m in the Southern Westerly (SW) Pacific (Fig. 2). Our data locate the core of the 14C-depleted water mass at the NZM in a of ∼2,500 m (PS75/100-4; modern depth of Upper Circumpolar Deep Water; Fig. 4), yielding a maximum deep water to atmosphere offset in radiocarbon activities (ΔΔ14C) of approximately −1,000‰. Analysing ΔΔ14C corrects for any impacts of changes in 14C production26 as well as for variable ocean–atmosphere exchange rates. A corresponding apparent ventilation age, based on benthic minus reservoir-corrected planktic 14C ages would equate to ∼8,000 years (Supplementary Fig. 3d). A similar glacial ΔΔ14C depletion of about −870‰ was reported from sediment core U938, which was recovered at the NZM in a water depth of 2,700 m (ref. 5) (Figs 1a and 4). We hypothesize that these extremely low ΔΔ14C-values might be the result of the admixture of 14C-dead hydrothermal CO 2 into a water mass with an initial high ventilation age, which was estimated to at least 2,700 years6. The upper and lower boundary of this old water mass are marked by higher ΔΔ14C-values of −550‰ to −600‰ indicating a highly stratified water column (Fig. 4b). Similar ΔΔ14C-values were reported at the NZM north of Chatham Rise at 2,314 m (ref. 6). This confines the most radiocarbon-depleted waters to a depth below ∼2,300 m. The observed trend of 14C between 830 and 4,300 m parallels the highest glacial nutrient concentrations off New Zealand, between 2,000 and 3,000 m (ref. 27) (Fig. 5), likewise indicative for the presence of aged, nutrient rich (low δ13C) and radiocarbon-depleted waters. Yet, the δ13C reconstructions might yield a certain bias, as endobenthic (Uvigerina) instead epibenthic (Cibicidoides) foraminifera were used27.

Figure 4: Comparison of oceanic and atmospheric proxy records. ΔΔ14C-values (ocean–atmosphere) of (a) the intermediate Pacific region; PS75/104-1 and SO213-84-1 (this study); MD97–2120 (ref. 15) (green line: Bounty Trough); MV99-MC19/GC31/PC08 (ref. 12) (black line: Northeast Pacific); (b) CDW ΔΔ14C; PS75 and SO213 records (this study); MD97–2121 (ref. 6) (light blue line: north of Chatham Rise); U9385 (red square: Bounty Trough); Coral dredges8 (red line: Drake Passage); MD07-3076 (ref. 4 (brown line: South Atlantic); and Modern SW-Pacific UCDW ΔΔ14C (ref. 67) (pink triangle). (c) Atmospheric CO 2 concentrations1 (red line); WAIS δ18O record69 (orange line) and atmospheric Δ14C-values20 (black line). UCDW, Upper Circumpolar Deep Water. Full size image

Figure 5: Vertical distribution of carbon isotopes off New Zealand throughout the LGM. Orange line—ΔΔ14C (this study). Green line δ13C Uvigerina 27. Full size image

We traced the 14C-depleted glacial carbon reservoir off New Zealand to the central South Pacific (EPR) 4,000 km east of the NZM. At this location, ΔΔ14C-values are as negative as −900‰ at 3,600 m (PS75/059-2; Figs 1a and 4). Hence, we are confident that this water mass was not only restricted to the NZM but seems to have occupied large parts of the South Pacific. Further off, in the Drake Passage and the South Atlantic, glacial water masses have been identified in CDWs with ΔΔ14C-values as low as −330‰ (ref. 8) and −540‰ (ref. 6), respectively (Fig. 4). In their timing and amplitudes, these records are similar to our Pacific radiocarbon signature characterizing the upper and lower boundary of the old carbon pool (Fig. 4b). Our intermediate-water record (SO213-84-1; modern depth of AAIW) shows the highest glacial ΔΔ14C-values of our transect (approximately −90‰). We suggest that the 14C-depleted deep waters represent the very old return flow from the North Pacific (PDW), similar to the modern circulation pattern (Fig. 1b). The distribution of radiocarbon in our reconstruction might indicate a floating carbon pool instead of a stagnant bottom layer. Several records from the North Pacific might corroborate this assumption. In the Gulf of Alaska28 (MD02-2489) and off Kamchatka29 (MD01-2416), as well, the glacial mid-depth water mass (Kamchatka) shows a considerable higher benthic to planktic 14C offset than the deeper water mass off Alaska. Additional data from the northwest Pacific suggest the lowest glacial 14C values in a water depth of ∼2,300 m with better ventilated waters above and below21. In the Atlantic Ocean as well, Ferrari et al.30 and Burke et al.31 observed a mid-depth (floating) radiocarbon anomaly. A floating carbon pool might furthermore explain why no sign of old carbon was found in the deep equatorial Pacific below 4,000 m (ref. 32). Therefore, this record might have ‘missed’ the old, mid-depth carbon pool above.

Any explanation for the pronounced glacial radiocarbon-depletion of the deep SO has to involve a limited ocean–atmosphere exchange due to strengthened ocean stratification under glacial boundary conditions3 (Fig. 6a). Northward-expanded Antarctic sea ice and SW Winds33 contributed to reduced air–sea gas exchange and upwelling of deep-waters30. Surface freshening by melting sea-ice in the source regions of intermediate waters34 and enhanced formation of highly saline Antarctic Bottom Water35 may have set a density structure that led to reduced mixing and the encasement of old PDW (Fig. 6a). In addition, the shoaling of North Atlantic Deep Water30 might have reduced the contribution of freshly ventilated waters into South Pacific CDW below ∼2,000 m. These interacting key processes may have significantly contributed to the low radiocarbon values and are consistent with an enhanced glacial storage of carbon in the deep ocean.

Figure 6: Schematic representation of South Pacific overturning circulation. (a) Glacial pattern: northernmost extent of sea ice and SWW. Increased AABW-salinity by brine rejection favours stratification. Increased dust input promotes primary production and drawdown of CO 2 . (b) Deglacial pattern: upwelling induced by southward shift of Antarctic sea ice and SWW. The erosion of the deep-water carbon pool releases 14C-depleted CO 2 towards the atmosphere. Following air–sea gas exchange, the outgassing signal is incorporated into newly formed AAIW (light blue shading). Blue shading: poorly ventilated old and CO 2 -rich waters; Darkest shading 2,500–3,600 m: water level influenced by hydrothermal CO 2 . Green arrows: intermediate water; orange arrows: deep-water; light-blue areas: sea ice; SWW: Southern Westerly Winds; coloured circles: sediment cores (colour coding according to Fig. 2); black circle: SO213-79-2—no glacial data; and circular arrows: diffusional and diapycnal mixing. Full size image

Old water masses of 5,000–8,000 years are expected to be strongly oxygen depleted36. According to Sarnthein et al.7, water masses with Δ14C values lower than −350‰ would be completely anoxic. However, pronounced anoxia have not been documented in the deep South Pacific between 2,500 and 3,600 m. Therefore, an admixture of 14C-dead carbon via submarine tectonic activity along MOR18 into a an old water mass in the deep South Pacific might have contributed to the extremely low radiocarbon values of the water mass at ∼2,500–3,600 m. During the LGM, sea floor eruption rates along tectonically active plate boundaries may have intensified due to the lower glacial sea level17,18,19. This process might have released significant amounts of 14C-dead CO 2 into the water column. Using a simple 1-box model (Methods), we tested our hypothesis and calculated if the injection of hydrothermal CO 2 into the deep Pacific has the potential to amplify the ΔΔ14C minimum throughout the LGM (Fig. 3). To overcome the influence of the variable atmospheric 14C-levels26,37, we compared our simulated Δ14C to our reconstructed ΔΔ14C values (deep ocean-to-atmosphere offset). The probably time-delayed response of submarine volcanism to changes in sea level complicate our flux calculations19. A crucial prerequisite for our hypothesis of the admixture of hydrothermal CO 2 is the presence of an already aged water mass with high nutrient concentrations and low Δ14C levels (Fig. 7). According to the record of MD97–2121 (ref. 6) (2,314 m), the glacial ventilation age off New Zealand is at least 2,700 years. However, radiocarbon values might have been even lower as the MD97–2121 record lacks any data points between ∼25 kyr and ∼18 cal. ka (Fig. 4b). Our ΔΔ14C record of SO213-82-1 (2,066 m; Fig. 4b) is −600‰ at ∼20kyr, ∼100‰ lower than the minimum observed in MD97–2121 ∼25 cal. ka, potentially indicating even higher turnover times. When we combine the radioactive decay, caused by an estimated water mass age of 2,700 years for the time of 35–18 cal. ka, with submarine 14C-free volcanic CO 2 influx, our model calculates a decrease in Δ14C for the corresponding water mass by additional −500‰ to −600‰ (Fig. 3a). This hydrothermal CO 2 outgassing (potentially accompanied by carbonate compensation and/or CO 2 sequestration) would lead to a maximum atmosphere-to-deep ocean offset of −800‰ to −1,000‰ ΔΔ14C (Supplementary Fig. 4e–g), comparable to the maximum depletion observed in PS75/059-2 and PS75/100-4. Today, in a water depth between ∼2,500 and ∼3,500 m, pronounced volcanic outgassing occurs along the southern EPR38,39. The resulting hydrothermal plume spreads towards the west and can be traced by the 3He-signal in the broader western Pacific (Fig. 8)40 and off northern New Zealand, right in the water depth under debate of ∼2,500 m (ref. 41). Therefore, we argue that increased glacial outgassing of 14C-dead volcanic CO 2 into a stratified ocean has the potential to significantly lower the Δ14C-content of an old (at least 2,700 years) water mass. The prominent Chatham Rise (Fig. 1a) might have acted as a physical barrier, blocking MD97–2121 (ref. 6) (∼2,300 m) from the volcanic plume. As MD97–2121 lacks data for most of the last glacial (∼18–25 cal. ka), we cannot fully exclude that this core might have been affected by volcanic CO 2 to some extent. Nevertheless, the 14C-data of MD97–2121 are already significantly higher at ∼18 cal. ka compared to the values of PS75/100-4. Therefore, we argue that the influence (if any) of hydrothermal activity must have been lower to the north of the Chatham Rise and/or at 2,300 m water depth. Despite the in detail unknown processes accompanying such a hydrothermal carbon flux, its admixture might add additional carbon to the ocean–atmosphere–biosphere system (0.08–0.16 PgC per year; Fig. 3b). However, the net carbon injection depends in detail on the strength of the additional processes carbonate compensation and CO 2 sequestration and might also be zero. Once the glacial processes, favouring stratification, are reversed, any net injected carbon might eventually be released to the atmosphere along with the carbon already stored within the deep ocean (Fig. 5b). A further quantification of the net carbon injection and its contribution to atmospheric CO 2 is not yet possible, since future investigations with process-based models are necessary. Furthermore, as the distribution of MOR is inhomogeneous in the world ocean, it is difficult to compare our local results to global CO 2 flux estimates24. In Fig. 3b, we illustrate that the water mass affected by hydrothermal CO 2 might span an area of between 3 and 17% of the global glacial ocean. While the results for the minimum area, covered by our sediment cores, are below the maximum estimate of present day global estimate of hydrothermal outgassing, the results for an area representative for most of the South Pacific are a factor of 2–4 times higher (Fig. 3b). This suggests that if the admixture of hydrothermal CO 2 is the process that can explain the minimum ΔΔ14C values recorded in the mid-depth South Pacific (PS75/100-4; PS75/059-2; U938 (ref. 5)) the global CO 2 -fluxes from MOR throughout the LGM might have been much larger than today. However, if the initial water mass was older than the 2,700 years assumed in our model, the resulting fluxes might also have been smaller than stated here. Although mantle-CO 2 is depleted in both, Δ14C and also in δ13C (−5±3‰ (refs 42 and 43)), its influence might be stronger on radiocarbon. As 14C is by far less common in the ocean than 13C, it can be diluted (lowered) more easily than its non-radiogenic counterpart.

Figure 7: Processes linking the glacial release of hydrothermal CO 2 and water mass Δ14C. The drop in global sea level triggers increased volcanic activity at MORs. The plume of 14C-dead hydrothermal CO 2 is mixed into an aged water mass, in which the combined effects of surface reservoir age and deep-ocean turnover time, of ∼2,700 years, led already to a background ocean-to-atmosphere offset in ΔΔ14C of −300‰ to −400‰. This admixture of hydrothermal CO 2 further lowers the water masses ΔΔ14C by another −500‰ to −600‰, yielding a total ΔΔ14C of about −1,000‰ (purple layer). Modified after Hand70. Reprinted with permission from AAAS. Full size image

Figure 8: Dispersal of hydrothermal 3He in the Southwest Pacific. The modern hydrothermal 3He plume, emanating at the southern EPR in a water depth of ∼2,500 m (refs 38 and 39), can be traced throughout the southern Pacific towards New Zealand41. 3He distribution along (a) WOCE line P17 (ref. 67) (∼135° W) and (b) WOCE line P16 (ref. 67) (150° W). (c) Locations of 3He sections P16 and P17. (d) Hypothesized dispersal of the glacial hydrothermal plume emanating from the EPR. Core locations and depths indicated by red bars and yellow squares. Panels generated using ODV 4.7.2 (ref. 68). WOCE, World Ocean Circulation Experiment. Full size image

We assume that the previously outlined glacial/interglacial changes in the SO climate system (position of sea ice and westerlies; changes in water mass densities; changes in upwelling and circulation) are the major factors influencing the spatio-temporal evolution of the oceanic carbon pool. However, the extreme minima in ΔΔ14C between 2,500 and 3,600 m water depth in the South Pacific are most plausibly explained by the hypothesized admixture of hydrothermal CO 2 into an old and already 14C-depleted water mass. Admittedly, this process would complicate the use of 14C as a ventilation proxy for the water masses affected by hydrothermal CO 2 . However, as the presence of an already existing 14C-depleted glacial water mass is a crucial prerequisite for our model, our theory does not interfere with the concept of stratification, decreased ventilation and the presence of a glacial oceanic carbon pool.

In the Pacific, other processes affect the radiocarbon inventory of water masses as well. Stott and Timmermann44 already suggested the release of 14C-depleted carbon from gas clathrates. As the stability of such clathrates is located in shallow waters ∼400 m (ref. 44), this process is not applicable for our mid-depth anomaly below ∼2,500 m. Therefore, we reject the possibility of any large influence of 14C-depleted CO 2 and CH 4 clathrates.

At the end of the LGM and during the transition into the Holocene (∼20–11.5 cal. ka), converging ΔΔ14C-values argue for a progressive destratification (Fig. 4). During this interval, the deep-water-to-atmosphere offset in radiocarbon between ∼2,000 and ∼4,300 m decreases significantly (Fig. 4b). Although the resolution of our deep-water sediment cores is rather low, their ΔΔ14C-values increase within error synchronous to the rise in atmospheric CO 2 rise (Fig. 4).

During Termination 1, when the most radiocarbon-depleted deep waters rejuvenate, no pronounced depletion in AAIW ΔΔ14C is recorded (Fig. 4a). However, the intermediate-water ΔΔ14C-values remain low from ∼18–15 cal. ka. The decrease in the deep water-to-atmosphere ΔΔ14C-offset and the abrupt drop in δ13C of atmospheric CO 2 (ref. 45) suggests increased air–sea gas exchange and the oceanic release of upwelled old CO 2 (Fig. 5b). The intermediate-waters from the NZM (this study) and the Chile Margin14 significantly deviate from the (sub)tropical East Pacific, which shows two prominent drops in ΔΔ14C at the intermediate-water level during Termination 1 (refs 12, 13) (Fig. 4a). Therefore, it seems unlikely that southern sourced AAIW represents the source for the deglacial radiocarbon signals in the (sub)tropical East Pacific.

As we mentioned before, because of the close proximity, similar reservoir ages for all sediment cores are a requirement for our interpretations. However, the Holocene reservoir ages of two of our cores differ by ∼1,000 years (Supplementary Fig. 5). Although this offset does not change the overall story, it prevents us from discussing the data of this time interval in more detail.