A tremendous amount of work has been devoted to developing foraminiferal proxies for temperature and pH, using global calibrations derived from core-top samples (for example, the Mg/Ca seawater temperature proxy41). Low I/Ca ratios of planktonic foraminifera unambiguously reveal the presence of low-oxygen waters, but a global calibration approach cannot establish planktonic foraminifera I/Ca as a linearly quantitative proxy for the continuum of dissolved O 2 concentration. Due to the stepwise nature of redox reactions42, quantitative IO 3 − reduction does not occur before the dissolved oxygen is depleted to a certain threshold value, triggering nitrate reduction43. IO 3 − concentrations at water depths matching planktonic foraminiferal habitats are often not in equilibrium with the in situ O 2 concentrations, and O 2 contents which are sufficiently low to initiate major IO 3 − reduction may be detrimental to many species44. Instead, the I/Ca (recording the in situ IO 3 − concentration) is determined by the depth habitat of the foraminifera and the upper ocean IO 3 − mixing gradient. This mixing gradient is largely controlled by the surface water IO 3 − concentration and the depth of the IO 3 − reduction zone28. Nonetheless, a planktonic foraminifera proxy that semi-quantitatively approximates dissolved O 2 concentrations, indicative of the presence of an OMZ, can still be highly valuable for the paleoceanography community.

Before interpreting the down-core record from site TC493/PS2547, we identify the characteristic I/Ca signals for modern OMZs. IO 3 − depth profiles in the open ocean generally fall into two types (Fig. 2d): (1) the OMZ-type, with low surface water values and near-zero subsurface values in the OMZ; and (2) the normal open ocean type (for example, in a well-oxygenated water column), with relatively high surface water values and even higher subsurface values. A threshold O 2 concentration will cause complete IO 3 − reduction in the subsurface, and there may be a surface water IO 3 − threshold concentration below which complete IO 3 − reduction is likely to happen in the water column. Combined with modern water column IO 3 − and O 2 data, the I/Ca values measured on modern and late Holocene planktonic foraminifera consistently indicate that I/Ca <2.5 μmol mol−1 is equivalent to a surface water IO 3 − concentration of <0.25 μmol l−1, thus providing a marker for the presence of oxygen-depleted water with a subsurface O 2 concentration <20–70 μmol kg−1 (Fig. 2a–c).

Modern surface water IO 3 − concentrations are influenced by productivity and the presence of a subsurface OMZ25,28. To visualize this relationship, we compiled surface water IO 3 − concentrations from the literature and plotted them against the minimum O 2 concentrations in the subsurface water (Fig. 2b). The IO 3 − concentration broadly increases with the minimum O 2 concentration when the surface water IO 3 − concentration is >0.25 μmol l−1 (Fig. 2b). This correlation is likely a reflection of surface productivity versus subsurface respiration, because lower productivity leads to lower iodine uptake in surface water and less oxygen consumption by subsurface organic matter decomposition. In areas with a strong OMZ and near-zero O 2 values, the surface water IO 3 − concentrations are below 0.25 μmol l−1 (Fig. 2b). A partition coefficient K d (K d =[I/Ca]/[IO 3 −] with units of [μmol mol−1]/[μmol l−1]) of ∼10 was reported from abiological calcite synthesis experiments17,20. Using this K d value, an IO 3 − concentration <∼0.25 μmol l−1 results in I/Ca values <∼2.5 μmol mol−1 in calcite. This estimate is consistent with modern I/Ca at OMZ Sites 658, 849 and 1242, as well as the last interglacial I/Ca value at Site 720 (Fig. 2a). Therefore, a surface water I/Ca value <2.5 μmol mol−1 indicates that a pronounced subsurface O 2 minimum exerted the dominant control on the upper ocean IO 3 − profile. This I/Ca threshold value does not seem to depend on foraminiferal species (Fig. 2a).

The O 2 threshold for maintaining an OMZ-type IO 3 − profile is useful for the paleoceanographic application of the planktonic I/Ca proxy. At O 2 concentrations <20 μmol kg−1, microbial processes become dominant29, and IO 3 − likely would be completely reduced to I− since the reaction is biologically mediated45 (for example, ODP Sites 1242, 720 and 849 in Fig. 2a,c). ODP Site 658 is located at the northern edge of a shallow pocket of distinctively low-oxygen water with mean O 2 concentrations of ∼70 μmol kg−1 in the upper 200 m (ref. 46), which may be sufficiently low to generate an OMZ-type iodate profile. Three species of planktonic foraminifera analysed at ODP Site 1242 show exceptionally low I/Ca ratios around 0.5 μmol mol−1, corresponding to an IO 3 − concentration of ∼0.05 μmol l−1. Such a low IO 3 − concentration is comparable to that reported for a location where an extreme hypoxic event occurred47. Moreover, this low IO 3 − concentration implies that IO 3 − reduction should occur shallower than at Site 849 and at two sites with classic OMZ-type IO 3 − profiles (Eastern Equatorial Pacific28 and Arabian Sea48; Fig. 2c). A comparison of the O 2 profiles of these sites reveals that the O 2 threshold needs to be >50 μmol kg−1 to achieve a shallower IO 3 − reduction at Site 1242. Therefore, we suggest that I/Ca values lower than ∼2.5 μmol mol−1 indicate O 2 minima <20–70 μmol l−1. This O 2 range cannot be further narrowed down with the available information, and we refer to this range as the O 2 threshold for an OMZ-type IO 3 − profile. However, the threshold behaviour of IO 3 − reduction (relative to O 2 ) in subsurface waters does not necessarily lead to step changes in down-core records of planktonic I/Ca. This is because planktonic foraminifera typically record the IO 3 − mixing gradient in the top part of water column, above the O 2 -depleted zone where rapid step changes in IO 3 − concentrations occur. Low planktonic I/Ca values may be driven by shoaling of O 2 -depleted water, and/or by increasing productivity, both of which could change gradually over time.

The available data from modern and late Holocene planktonic foraminifera suggest that the I/Ca ratio acts as a robust (paleo-) proxy for determining the signature of O 2 -depletion in the upper ocean (Fig. 2). At site TC493/PS2547, I/Ca was high (5–7 μmol mol−1) during the Holocene and interglacial MIS 5 when compared with the lowest values (<2 μmol mol−1) characterizing peak glacial periods MIS 2 and 6 (Fig. 3). Changes in salinity, temperature and foraminiferal habitat, most likely, are not the main drivers for this record (Supplementary Discussion). The glacial I/Ca values of N. pachyderma (s) are best explained by the presence of a water mass with a dissolved O 2 content <70 μmol kg−1 close to, i.e., above or near, this site (Figs 2 and 3). We reiterate that the low I/Ca does not necessarily imply O 2 -depleted seawater within the foraminiferal habitat.

At present, CDW wells up to a water depth of approximately 250–300 m in the Amundsen Sea31 and has O 2 concentrations notably lower than the top 200 m of the water column (Fig. 2c). Although the interpretation of absolute values of planktonic δ13C is far from straightforward in the seasonal ice zone (for example, disequilibrium from seawater49), it is reasonable to assume that CDW had a strong influence on the local water column during glacial periods, as its upwelling along the continental margin was probably responsible for the opening of the glacial polynyas. The CDW upwelling at site TC493/PS2547 today partly originates from Pacific Deep Water (PDW) moving southward from the equator, with a low-oxygen and high nutrient signature (Fig. 1)50,51. δ30Si data from fossil diatoms and sponges indicate higher silicic acid concentrations in the Pacific sector of the Southern Ocean during the LGM, which further imply that either the southward transport of PDW was more efficient or PDW was less ventilated than today52. So glacial CDW was likely more O 2 depleted than during interglacials, and upwelling of this water contributed to the glacial I/Ca signal at site TC493/PS2547.

The oxidation of I− to IO 3 − is thought to take from a few months up to 40 years53. Long-distance transport of well-oxygenated deep water with low IO 3 − concentrations (<0.25 μmol l−1) has not been documented in the modern ocean, but this scenario should be tested with further work on I− oxidation kinetics. Today our site is bathed in CDW transported from a Pacific OMZ and the interglacial I/Ca values at site TC493/PS2547 do not show any remnant signal of the OMZ from the Pacific Ocean. On the basis of the knowledge about iodine speciation change in modern ocean, we interpret the observed glacial I/Ca values as a local signal, in principle, indicating the presence of a water mass with low O 2 and low IO 3 − vertically or horizontally close to the planktonic foraminiferal habitat.

In the setting of site TC493/PS2547 a coherent conceptual model for N. pachyderma (s) recording the presence/absence of O 2 -depletion needs to integrate changes in productivity, sea–ice extent and the opening/closing of polynyas on time scales of glacial to seasonal cycles (Fig. 4). Although the polynyas complicate the interpretation of the proxy data, their presence arguably provides the only window for sufficient accumulation of planktonic microfossils to record upper ocean conditions during glacial periods at such high latitudes.

Figure 4: Conceptual illustration of paleo-environmental changes. Upper ocean IO 3 − and O 2 profiles were influenced by circulation, productivity and polynyas over glacial cycles. (a) Well-oxygenated interglacial condition; (b) Relatively oxygen-depleted glacial conditions with expanded sea–ice cover; (c) Episodic polynya opening during glacials. Full size image

The modern O 2 profile at site TC493/PS2547 is defined by equilibration with the atmosphere at 0–250 m, and CDW influence below 250 m, as shown by the distinctively low O 2 concentrations (Fig. 2c). With O 2 above the threshold for complete IO 3 − reduction in the entire water column, the IO 3 − profile at site TC493/PS2547 should be similar to those at other high latitude locations, for example, site PS71/179–1 at 69°31′ S and 0°3′ W in the Weddell Sea54 (Fig. 2d). An interglacial scenario of relatively high seasonal productivity, high O 2 and surface water IO 3 − (>0.3 μmol l−1) concentrations (Fig. 4a), is well described for the modern Atlantic sector of Southern Ocean54.

Relative to the interglacial periods, the Southern Ocean experienced expanded sea–ice cover during glacial periods, and was less ventilated9,36. A more dynamic seasonal sea–ice cycle during ice ages would have increased water column stratification. Increased winter sea–ice formation (spatially and volumetrically) may have generated waters dense enough to sink ultimately to the bottom of the ocean55. On the other hand, melting of thicker sea ice during glacial-time summers in the seasonal sea–ice zone would have strengthened the halocline (not considering the influence of polynyas). So, the glacial seasonal stratification was likely stronger than today. These factors overall should have lowered the glacial O 2 concentrations in the Southern Ocean (Fig. 4b). At site TC493/PS2547, glacial I/Ca demonstrate that the IO 3 − profile was OMZ-like with complete IO 3 − reduction near the foraminiferal habitat (Fig. 2d). However, the dynamics of polynyas must be considered when interpreting the location of the low O 2 water mass, and the means by which the signal was recorded by N. pachyderma (s).

Without a polynya above site TC493/PS2547, glacial phytoplankton productivity under perennial sea–ice cover would have been relatively low due to the scarcity of light34, and planktonic foraminifera depending on algae could not flourish. The water column would have been relatively poorly ventilated and strongly stratified during these times, creating the ideal environment for developing low O 2 conditions and an OMZ-type IO 3 − profile (Fig. 4b). The episodic opening of a polynya re-established primary production (mainly by diatoms) and thus a planktonic foraminiferal habitat, vertical mixing and oxygenation in, at least, the uppermost part of the water column (Fig. 4c). While overall glacial-time production was reduced30,34, the planktonic foraminifera preserved in the glacial sediments probably recorded transient I/Ca changes in the water column associated with polynya-induced peaks in glacial productivity. Modern open ocean productivity pulses do not lower IO 3 − concentrations to <0.25 μmol l−1 in oxygenated water (Supplementary Discussion)54, thus the glacial I/Ca signal is most likely driven by changes in O 2 and not productivity.

The likely short-lived nature of glacial polynyas makes it difficult to envisage that very brief plankton blooms alone could produce a utilization-driven O 2 depletion in a cold, well-oxygenated Southern Ocean. For the same reason, it is difficult to imagine that the vertical mixing cells restricted by the size of the polynya could rapidly oxygenate voluminous nearby waters outside of the polynya, if most of the sea–ice covered areas were O 2 -depleted. The more likely scenario is that the O 2 concentrations in the deep and abyssal Southern Ocean were generally lower during glacial periods than during interglacial periods. Upwelling of a more O 2 -depleted CDW in the generally stratified upper ocean was mainly responsible for the IO 3 − reduction at site TC493/PS2547 (Fig. 4b), while the episodic opening of polynyas created habitable conditions for planktonic foraminifera to record the deoxygenation in the upper ocean (Fig. 4c). We suggest that the I/Ca proxy should be used as a local proxy, in principle. However, it is probably a reasonable speculation that this record (Fig. 3) shows oxygenation changes integrated over a regional volume of water (e.g. CDW).

The timing of glacial deoxygenation and deglacial reoxygenation at site PS2547 shows potential linkages to global climate changes (Fig. 3). The appearance of OMZ-type I/Ca values (<∼2.5 μmol mol−1) during past glacial periods coincided with the lowering in atmospheric pCO 2 level below the long-term mean value56. Identical timing was reported for a strongly stratified Antarctic Zone coincident with pCO 2 decrease under the same threshold value (225 p.p.m.) in the Atlantic sector of the Southern Ocean57. Stronger stratification may be the common driving force for the productivity change (ODP Site 1094) and oxygenation change (PS2547/TC493) in the Antarctic zone. Furthermore, during the last interglacial period, the recovery of N. pachyderma (s) I/Ca values is offset from the δ18O trend, with peak I/Ca occurring about 10 kyr after the peak δ18O (Fig. 3), an observation worthy of future investigation.