Samples were analyzed by Inductively Coupled Plasma Mass Spectrometry (ICP-MS, Agilent 7900), after offline preconcentration by Solid Phase Extraction (SPE) on a seaFAST SC-4 DX module (Cloete et al., 2019). Briefly, 10 mL of seawater was taken up by the seaFAST (Elemental Scientific Inc.) module, buffered with an ammonium acetate buffer to a pH of 6.0 ± 0.2, and loaded onto a high affinity metal chelating resin column (Nobias PA1). Here the metal ions including Cd were bound to the resin and separated from the seawater matrix elements, e.g., Na, Mg, and Cl which passed through the column. The metal ions were subsequently eluted (in a 2% HNO₃ solution; Ultrapur^®, Merck) from the resin column in low volumes (250 μL) resulting in a preconcentration factor of 40. On the ICP-MS, samples were introduced using a low self-aspirating perfluoroalkoxy (PFA) nebulizer with a flow rate of 0.2 mL min^–1. Cd was measured using the instruments Octopole Reaction System (ORS) in helium (He) collision mode to reduce plasma and matrix based interferences, e.g., molybdenum oxide (MoO⁺) interference on Cd, although the molybdenum concentration was extensively reduced by the seaFAST matrix removal system. In order to quantify any remaining MoO⁺ interference on Cd, a Cd-free Mo standard at typical seawater concentrations (10–100 nmol kg^–1) was analyzed. Instrument parameters were optimized to reduce the interference to a maximum of 1 ppt Cd (∼9 pmol kg^–1). The Cd-free Mo standard was included throughout the analysis to monitor MoO⁺ production and correct the Cd results. Online internal standard addition for drift correction was not possible using the self-aspirating nebulizer. Instead, instrument drift was monitored by running a multi-element calibration standard (MES; supplied by Inorganic Ventures) every six samples. Where drift exceeded 5% relative to the starting concentration of the MES for a specific element, a drift correction was applied using the formula: 2^∗Conc_{MES_start}/(Conc_{MES_a} + Conc_{MES_b})^∗Conc_Sample, where a and b are the MES before and after each set of six samples.

The accuracy of the method was verified by the analysis of numerous community reference seawater samples for dCd (Table 1). For the GEOTRACES SAFe D2, GSC and GSP reference seawater, our dCd concentrations were within analytical uncertainty estimates of consensus values confirming the method’s accuracy over a wide concentration range. Our mean dCd value from the analysis of the NASS-7 certified reference material (CRM) was within analytical uncertainty of the certified dCd value. To monitor ICP-MS precision, the Winter Indian Southern Ocean Seawater (WISOS) internal control seawater was placed in the analysis sequence and results compared against each other and the calibrated mean (Table 1). The WISOS internal control was collected in bulk (20 L) from surface water at 55°S; 28°E during the 2017 Winter Cruise after filtration through a 0.2 μm filter into an acid cleaned LDPE container. Sub-samples (1 L) of the bulk seawater were then acidified (pH = 1.7) using hydrochloric acid (30% HCl, Ultrapur^®, Merck). The calibrated dCd mean (888 ± 21 pmol kg^–1) concentration was established by replicate analysis (n = 10) of the WISOS internal control seawater in conjunction with the SAFe and NASS-7 reference materials. Results from the analysis of the in-run WISOS control sample were 899 ± 27 pmol kg^–1 (n > 30) confirming the method’s precision over multiple analytical sessions.

The method blank (seaFAST and ICP-MS) was determined by preconcentrating and analyzing a solution of 2% HNO₃ (Ultrapur^®, Merck) in ultrapure deionized water (the same composition as the eluent used in sample pre-concentration). The method blank was subjected to the same pre-concentration procedure as the seawater samples. The mean method blank was 1.22 ± 0.30 pmol kg^–1 (n = 5). The limit of detection (LOD) of the instrument, calculated as 3 times the standard deviation of the preconcentrated blanks, was 0.89 pmol kg^–1 (n = 5).

Samples for dCd were analyzed in duplicate. The final values represent the mean of duplicate measurements. In the two cases where the percentage relative standard deviation (%RSD) between the duplicate measurements was >10%, one of the values was deemed a suspected outlier and not used further. This was determined by curve fitting the data points based on the values above (shallower depth) and below (deeper depth) the suspect value as well as by comparison with other parameters (salinity, temperature, and nutrients) measured from the same GO-FLO bottle.

The transect intersected several frontal systems which, in turn, separated the study area into physically and biogeochemically distinct regions (Figure 1A). Frontal positions were determined previously (Weir et al., 2020) following temperature, salinity and oxygen criteria outlined in the literature (Orsi et al., 1995; Belkin and Gordon, 1996; Pollard et al., 2002). The Subtropical Front (STF) located at 42.4°S, defined the boundary between the Subtropical Zone (STZ) to the north and the Subantarctic Zone (SAZ) to the south. The Subantarctic Front (SAF) was identified at 46.2°S and separated the SAZ in the north from the Polar Frontal Zone (PFZ) to the south which was bounded in the south by the Antarctic Polar front (APF) at 49.3°S. Further south, the Antarctic Zone (AAZ) extended until the Southern Boundary (SBdy) at 58.5°S, incorporating the southern Antarctic Circumpolar Current Front (sACCF) at 56.5°S. The Marginal Ice Zone (MIZ), defined by 30% ice cover, was encountered at 61.7°S (de Jong et al., 2018). The location of the sample stations meant that, from north to south, station IO08 was found in the STZ, stations IO07 and IO06 in the SAZ, station IO05 in the PFZ and stations IO04, IO02, and IO01 were within the AAZ.

Water masses (Figure 1B) were identified using potential temperature, salinity and oxygen criteria outlined in the literature (Orsi et al., 1995; Anilkumar et al., 2006). The high temperature and salinity signature of surface waters at the northernmost station reflected Subtropical Surface Water (STSW; >12°C, S > 35.3). Moving southward, the decrease in surface water temperature marked the transition from STSW to Subantarctic Surface Water (SASW; 7–12°C) and Antarctic Surface Water (AASW; −1–2.5°C). Below the surface waters in the STZ, a layer of South Indian Central Water (SICW) was sampled. This water mass was underlain by the northwards subducting Subantarctic Mode Water (SAMW; 12–15°C; 26.2–26.6 kg m^–3) which was sampled between ∼300–500 m at 41°S. Similarly, Antarctic Intermediate Water (AAIW; 2.5–5.0°C, 26.7–27.4 kg m^–3) was found subducting northwards below SAMW. In contrast to the intermediate waters, sampled deep waters consisted predominantly of southward upwelling Upper Circumpolar Deep Water (UCDW) and the deeper lying North Atlantic Deep Water (NADW) and Lower Circumpolar Deepwater (LCDW). The UCDW (27.2–27.8 kg m^–3) was further characterized by oxygen minima (155–180 μmol kg^–1) and PO₄ maxima (2.1–2.6 μmol kg^–1). North of the APF, NADW (27.0–27.9 kg m^–3) was identified by high salinity (S > 34.7) and deep PO₄ minima (1.9–2.3 μmol kg^–1) while LCDW (27.8–27.9 kg m^–3), sampled below UCDW and NADW, was characterized by lower temperatures (∼1°C). Finally, Antarctic Bottom Waters (AABW; 27.8–27.9 kg m^–3), distinguished by characteristic local temperature minimums (–0.5–1.0°C) were sampled at depth in the AAZ (IO02 and IO04).

The SML at each station was defined as the depth interval between the sea surface and the Mixed Layer Depth (MLD) which was calculated as the depth (between 10 and 400 m) at which the Brunt Väisälä frequency (BVF) squared reaches a maximum while the euphotic zone depth was approximated as the depth at which 1% of surface PAR was measured (Weir et al., 2020). The MLD was shallowest at 122 m in the STZ, deepest in the SAZ at 190–220 m and intermediate in the PFZ (134 m) and AAZ (125–161 m). The euphotic zone depth was always shallower than the MLD and generally shoaled northwards from >56 m in the AAZ to 18 m in the STZ.

Depth profiles of dCd concentrations along the transect are shown in Figure 2A and Supplementary Figure 1. At each station, dCd profiles displayed typical nutrient-type behavior consistent with low surface concentrations and increased concentrations at depth. Across the transect, dCd was fairly uniform in the SML and was characterized by a strong north-south gradient which was associated with the frontal positions. Within the SML, dCd concentrations were 10 ± 2 pmol kg^–1 in the STZ and increased rapidly across the APF to 749 ± 16 pmol kg^–1 in the southern AAZ. The difference between surface and deep water concentrations was most apparent at the northern stations and decreased southwards. For example, the mean dCd in STSW was 33 ± 21 pmol kg^–1, ∼20 times lower than the concentration of 688 pmol kg^–1 (n = 1) in UCDW. At the southernmost station, this factor was ∼1.4 between AASW (713 ± 30 pmol kg^–1) and UCDW (1018 ± 65 pmol kg^–1). Mid-depth maxima in dCd varied in concentration and with depth and coincided with local maxima in PO₄ and minima in O₂. The mid-depth dCd maximum ranged between 700 pmol kg^–1 at 1,250 m depth in the STZ and 1,150 pmol kg^–1 at 200 m depth in the AAZ. Below the mid-depth peak, concentrations were typically lower in NADW (598 ± 47 pmol kg^–1) and showed increases in the LCDW (800 ± 120 pmol kg^–1) and AABW (856 ± 88 pmol kg^–1).

Major nutrient concentrations have been reported previously (Weir et al., 2020). Here we briefly describe the distribution of PO₄ (Figure 2B and Supplementary Figure 2) because of the extensive reference made to the dCd-PO₄ relationship. There was a strong southward increase in SML PO₄ concentrations, from 0.28 ± 0.08 μmol kg^–1 in STSW to 1.85 ± 0.09 μmol kg^–1 in AASW. Below the SML, PO₄ increased to maximum concentrations in UCDW at each station. The intermediate depth PO₄ maximum was most evident around the APF where PO₄ reached 2.43 ± 0.16 μmol kg^–1. The underlying NADW (2.10 ± 0.16 μmol kg^–1) and LCDW (2.14 ± 0.14 μmol kg^–1) were characterized by local PO₄ minima before slight increases in PO₄ in AABW (2.20 ± 0.02 μmol kg^–1).

The depth profiles of pCd (Figure 2C and Supplementary Figure 1) generally showed the opposite vertical distribution in comparison to dCd. Highest pCd concentrations were typically found within the surface water masses at each station, while strong negative gradients through subsurface water masses resulted in pCd concentrations having uniformly low (<1 pmol kg^–1) concentrations below 1,000 m at all stations. UCDW, LCDW, and AABW had mean pCd concentrations of 1.34 ± 1.08, 0.35 ± 0.21, and 0.39 ± 0.12 pmol kg^–1, respectively. In the SML, mean pCd was highest near the APF (St. 48°S; 18.83 ± 3.76 pmol kg^–1), and generally decreased with distance north and south of this maximal zone. To the north, mean pCd in the SML of the STZ was 6.18 ± 3.49 pmol kg^–1 and 9.44 ± 2.90 pmol kg^–1 in the SAZ. To the south, mean pCd in the SML of the AAZ was 14.72 ± 4.04 pmol kg^–1.

The depth profiles of P (Figure 2D and Supplementary Figure 2) resembled those of pCd whereby surface waters were characterized by high P concentrations which decreased rapidly in subsurface water masses to extremely low values at depth. Like pCd, highest SML P concentrations (14.29 ± 1.78 nmol kg^–1) were measured near the APF. North of the APF, concentrations in the SML ranged between 9.77 ± 3.26 and 12.84 ± 4.49 nmol kg^–1 in the SAZ and STZ, respectively. South of the APF, concentrations in the SML were between 8.54 ± 1.19 and 12.98 ± 0.53 nmol kg^–1. At greater depths, mean P concentrations in UCDW, LCDW, and AABW were 2.47 ± 1.40, 1.93 ± 0.87, and 1.48 ± 0.25 nmol kg^–1, respectively.

The data presented here contributes the first measurements of dCd and pCd from the 30°E longitude in the Indian sector of the Southern Ocean and one of only a handful of winter trace metal datasets produced globally, thus closing key spatial and seasonal gaps in the global database. For comparative purposes, a global compilation of dCd and pCd measurements (Figures 3A–D) has been created from all available data included in the GEOTRACES Intermediate Data Product (IDP) 2017 (Schlitzer et al., 2018; see figure caption for full list of references). For dCd, our Southern Ocean data fell within the range of concentrations found in other ocean basins. Globally, the majority of vertical ocean profiles of dCd showed surface minima, from 0.04 pmol kg^–1 in the Mediterranean Sea (Gerringa et al., unpublished) to 36 pmol kg^–1 in the Atlantic Southern Ocean (Baars et al., 2014). Below the surface, most vertical ocean profiles displayed intermediate to deep dCd maxima, from 368 pmol kg^–1 in the Mediterranean Sea (Gerringa et al., unpublished) to 1,246 pmol kg^–1 in the North Pacific Ocean (Minami et al., 2015). For the full depth profiles, our Southern Ocean dCd data had a range of 9–1,150 pmol kg^–1 which compared most similarly to that of the Atlantic Southern Ocean (36–970 pmol kg^–1; Baars et al., 2014) and Indian Ocean (10–1,100 pmol kg^–1; Thi Dieu Vu and Sohrin, 2013). For pCd, our Southern Ocean data fell within the range of concentrations found in other ocean basins and most closely resembled those observed in the South Atlantic, particularly in the upper 1,000 m. Highest pCd concentrations were found in surface waters of all datasets while extremely low values, between 0.02 pmol kg^–1 (South Atlantic; Henderson et al., unpublished) and 0.06 pmol kg^–1 (this study) characterized deep ocean distributions. For the upper 1,000 m, mean pCd concentrations were between 1.13 and 5.04 pmol kg^–1 in the North Atlantic (Hayes et al., 2018) and Equatorial Pacific (Lee et al., 2018), respectively, while the mean value for our Southern Ocean data was 5.92 pmol kg^–1.

Considering only surface (∼25 m) dCd measurements allowed for some interesting observations to be made regarding inter-basin and seasonal differences within the Southern Ocean (Figure 4). At the lower latitudes (north of ∼45°S), surface dCd was low across all datasets (Figure 4A) and our winter concentrations (9–34 pmol kg^–1) were most similar to those measured in the Pacific sector of the Southern Ocean (17–23 pmol kg^–1). At the higher latitudes (south of ∼45°S), surface dCd increased rapidly across all datasets. While the sharpest dCd increase was observed in the winter Indian sector of the Southern Ocean, dCd:PO₄ spot ratios calculated for all datasets were more comparable (Figure 4B), particularly for the Pacific sector (winter and summer) and Indian sector (winter) of the Southern Ocean, suggesting dCd increases were proportional to increases in PO₄. Between 55°S and 60°S there was a decoupling between winter (only Indian sector dataset considered as the Pacific sector measurements stop at 52°S) and summer dCd:PO₄ whereby winter ratios were considerably higher. These seasonal differences are likely caused by two primary mechanisms. First, intense summer productivity at the higher latitudes is driven by diatoms with high intracellular pCd:P (Baars et al., 2014) causing reduced surface dCd and dCd:PO₄ compared to winter values. Second, deep winter mixing provides access to subsurface water masses with enriched dCd:PO₄ signatures (from diatom remineralisation and regional upwelling of UCDW with high dCd:PO₄; see section “Physical Resupply of dCd Through Upwelling and Deep Winter Mixing”) thereby preferentially enriching surface waters with additional dCd. While these comparisons do present interesting insights into the seasonal cycling of dCd, interpretations should be made with caution considering that the data are not direct seasonal crossover comparisons.

Seawater samples were collected in early to mid-July and therefore represent early winter conditions. Owing to increased surface mixing and decreased light exposure (Dong et al., 2008; Swart et al., 2015), the winter months are generally considered biologically dormant in comparison to the spring and summer months. However, estimates of relative productivity (inferred from measurements of Chl-a; Supplementary Figure 3), were comparable to spring measurements and were at the lower end of summer measurements (Weir et al., 2020). For example, north of the APF, our early winter (July) Chl-a concentrations averaged over the SML (0.31–0.54 μg kg^–1) were similar to Chl-a concentrations measured to the east during spring (0.34 μg kg^–1; Becquevort et al., 2000) and summer (0.39–1.23 μg kg^–1; Rembauville et al., 2016). In the AAZ, our SML Chl-a (0.20–0.30 μg kg^–1) was within the range of previous spring concentrations (0.10–0.39 μg kg^–1; Becquevort et al., 2000) and below previous summer concentrations (0.49–0.98 μg kg^–1; Rembauville et al., 2016). This suggests that the winter euphotic layer was not biologically dormant (Weir et al., 2020) and that biological uptake by phytoplankton could potentially play an important role in controlling dCd and pCd distributions during winter.

The distribution of Cd and P were closely correlated across the transect, particularly in the upper water column (above the depth of the O₂ minimum) where a strong linear relationship between dCd vs. PO₄ [dCd (pmol kg^–1) = 431 ± 44 PO₄ (μmol kg^–1) – 198; r² = 0.89; n = 93] and pCd vs. P [pCd (pmol kg^–1) = 0.65 ± 0.11 P (nmol kg^–1) + 0.40; r² = 0.86; n = 85] was observed. Assuming that all P is of biogenic origin, and considering the close coupling of pCd and P, it is reasonable to assume that pCd is majorly biogenic in composition. Low dCd and PO₄ concentrations near the ocean surface and enrichments in subsurface waters are widely accepted to result from biological uptake by phytoplankton in the euphotic zone and subsequent remineralisation at depth (Abouchami et al., 2014; Baars et al., 2014; Sieber et al., 2019; Janssen et al., 2020). Biological uptake and assimilation of dCd and PO₄ drives the production of organic pCd and P as observed by the high surface pCd and P concentrations at each station (Figure 2 and Supplementary Figures 1, 2). Furthermore, vertical profiles of pCd co-varied with those of Chl-a in the upper 150 m and pCd maxima were either at or very close to the depth of the Chl-a maxima (Supplementary Figure 3). In the subsurface, decreasing pCd and P concentrations correspond to generally increasing dCd and PO₄ concentrations partly reflecting the aerobic remineralisation of sinking phytoplankton cells by the microbial community. The extremely low (<1 pmol kg^–1) concentrations of pCd, and relatively high concentrations of dCd (>500 pmol kg^–1), below 1,000 m suggests that variations in Cd distributions at depth were primarily controlled by the advection of water masses (with preformed Cd signatures) into the sampling region and were therefore not related to local particle transformations.

The availability of Cd and the relative availability of several essential trace metals (e.g., Fe, Zn, and Mn) has been shown to impact Cd uptake and therefore directly influence phytoplankton pCd:P stoichiometries (Price and Morel, 1990; Lee and Morel, 1995; Sunda and Huntsman, 2000; Cullen and Sherrell, 2005; Middag et al., 2018). In an effort to better understand the link between Cd uptake and the availabilities of Cd, Fe, Zn and Mn, we compare surface (∼25 m) pCd:P with dCd and dissolved ratios of Cd to Fe (dCd:dFe), Cd to Zn (dCd:dZn) and Cd to Mn (dCd:dMn; Figure 5; see Supplementary Tables 1, 2 for validation of dFe, dZn, and dMn). Ratios of pCd:P were low in the STZ and SAZ (0.37–0.52 mmol mol^–1), increased in the PFZ (0.86 mmol mol^–1) and were highest in the AAZ (0.91–1.07 mmol mol^–1). In the STZ, dCd was 9 pmol kg^–1 and increased roughly 80-fold to 768 pmol kg^–1 in the AAZ. All dCd:metal ratios were lowest in the STZ and increased at various rates through the SAZ and PFZ. In the PFZ, dCd:dZn was highest (0.38 mol mol^–1) and decreased in the AAZ (0.18–0.23 mol mol^–1) while dCd:dFe (13.78–15.36 mol mol^–1) and dCd:dMn (1.62–2.29 mol mol^–1) were highest in the AAZ.

The uptake of Cd by phytoplankton is directly related to the concentration of free (bioavailable) dCd and inversely related the concentration of free dFe, dZn, and dMn. Since Mn forms only weak organic complexes, the bioavailable Mn concentration is assumed to be approximated by dMn (Sunda and Huntsman, 2000). For Cd and Zn, organic complexation may potentially impact biological uptake in the STZ and SAZ where dCd and dZn are inherently low, however, in the PFZ and AAZ, bioavailable dCd and dZn are high owing to increased concentrations which saturate metal binding ligands (Baars and Croot, 2011; Baars et al., 2014). For Fe, organic complexation dominates Fe speciation in the Southern Ocean (Croot et al., 2004) and, coupled with the low surface dFe (0.03–0.12 nmol kg^–1) measured throughout the transect, suggests that dFe was potentially at or near limiting concentrations. It follows that in regions with higher dCd and dCd:metal ratios, phytoplankton may assimilate more Cd relative to P providing at least one potential driver behind the elevated pCd:P in the PFZ and AAZ. While our data suggest dCd availability is the primary factor determining Cd uptake, as suggested previously (Middag et al., 2018), assessing the relative importance of the various dCd:metal ratios in influencing pCd:P is more complex. Lowest dFe was measured in the AAZ (0.03–0.05 nmol kg^–1) potentially limiting phytoplankton growth in this region resulting in increased cellular pCd:P via growth biodilution or increased uptake of Cd (relative to P) through non-specific divalent metal transporters, a mechanism which is intensified under low Fe (Lane et al., 2009). In the PFZ, pCd:P was relatively high (comparable to the lower end of pCd:P in the AAZ) yet dCd:dFe was more than 4 fold lower than in the AAZ. Based on dCd:dFe alone, we would expect pCd:P in the PFZ to be much lower in comparison to the AAZ. However, dCd:dZn in the PFZ was almost two-fold higher than the AAZ and dCd:dMn was high (although not as high as values calculated for the AAZ) possibly indicating increased competition success for dCd over dZn (and possibly dMn) or passive uptake of Cd to prevent toxicity (Horner et al., 2013) thereby increasing pCd:P ratios. In the northern AAZ, pCd:P was higher, and dCd lower, at 50°S compared to 56°S, which contradicts the primary role for dCd availability in governing pCd:P. At both stations, dCd:dFe was equally high, however, dFe at 50°S (0.03 nmol kg^–1; the lowest surface value measured along the transect) was lower than at 56°S (0.05 nmol kg^–1) potentially limiting phytoplankton growth (and increasing pCd:P) at 50°S to a greater degree than 56°S. Further south, dCd, pCd:P, and dCd:dFe decreased between 56°S and 58°S and dCd:dZn and dCd:dMn increased suggesting that, of the Cd:metal antagonisms considered as secondary controls (after the primary control by dCd), the dCd:dFe antagonism had the largest effect on pCd:P variations in this region.

The literature regarding plankton metal stoichiometries in natural phytoplankton assemblages is scarce (Twining and Baines, 2013). Nevertheless, phytoplankton cellular Cd quotas are reported to range by a factor of 100 (0.01–1 mmol mol^–1 P) compared to a factor of 20 for other bioactive metals such as Fe, Mn, Zn, and Cu (Ho et al., 2003; Twining and Baines, 2013). We compare our surface water (∼25 m) pCd:P ratios with plankton derived Cd:P ratios to assess whether or not the surface water particulate signature is a good representation of the observed phytoplankton community based on observed phytoplankton community data from this transect (Weir et al., 2020, Supplementary Figure 4). We acknowledge the limitations of such a comparison which include differences in oceanographic setting (e.g., open ocean vs. coastal upwelling) and inherent environmental conditions (e.g., oligotrophic vs. nutrient replete conditions, differences in resident phytoplankton community assemblage and grazing pressures by local zooplankton) which may influence Cd uptake systematics. Regardless, we find that pCd:P do provide a reasonable estimation of the resident phytoplankton community. To illustrate, a flagellate-dominated assemblage from the North Atlantic had a pCd:P ratio of 0.51 ± 0.09 mmol mol^–1 (Kuss and Kremling, 1999), almost exactly matching our pCd:P ratio from the flagellate-dominated SAZ stations (0.52 mmol mol^–1 for both SAZ stations). However, these ratios were higher than the STZ (0.37 mmol mol^–1) station and lower than the PFZ station (0.86 mmol mol^–1) despite significant contributions from flagellates to the phytoplankton community at both stations. There was, however, a much larger contribution of diatoms relative to flagellates at the PFZ station compared to the STZ providing a potential driver of variations in pCd:P whereby diatom growth rates have been shown to directly influence pCd:P ratios while flagellates show no functional response in pCd:P (Finkel et al., 2007). In the PFZ, changes in diatom growth rate, potentially due to the availability of other trace metals [see section “Cd Uptake Influenced by the Availability of Dissolved Cd and Other Metals (Fe, Zn, and Mn)”], would therefore have a greater bearing on the total phytoplankton community pCd:P ratio compared to the STZ. In the AAZ, the high pCd:P ratios (0.91–1.07 mmol mol^–1) compared most similarly to a ratio from a diatom-dominated assemblage from the high latitude Southern Ocean (1.29 ± 0.07 mmol mol^–1; Cullen et al., 2003) in agreement with our phytoplankton community assemblage data. The higher pCd:P ratios in the Cullen study could be the result of lower summer dFe and dMn concentrations (0.03 and 0.08 mol kg^–1, respectively) which may decrease diatom growth rates (and therefore increase pCd:P) to a greater extent compared to winter.

In order to investigate remineralisation rates of pCd and P at each station, we plot pCd and P concentrations individually (as a percentage of their respective maximum concentrations at each station) with depth as well as corresponding pCd:P ratios with depth (Figures 6A–D). We show one representative profile for each Southern Ocean Zone (see Supplementary Figure 5 for all other stations and Supplementary Figure 6 for full depth profiles). Maximum pCd and P concentrations generally occurred at the same depth and were most commonly located between the euphotic depth and the base of the mixed layer (with an exception in the PFZ where maxima were co-located above the euphotic depth). We observe different trends in the vertical pCd:P profiles across the transect. In the AAZ (Figure 6A), pCd:P increased to a subsurface maximum (1.2 mmol mol^–1) at 100 m, consistent with a preferential remineralisation of P over this depth range. Below the pCd:P maximum, ratios decreased rapidly and the relative concentrations of pCd and P suggest a preferential remineralisation of Cd at depth. These observations are consistent with previous findings in the AAZ of the Southern Ocean, and in high-nutrient, low-chlorophyll areas throughout the global ocean, where it was hypothesized that a labile P fraction (associated with ATP and DNA in the cell) is preferentially remineralised over pCd near the surface resulting in an increase in pCd:P (Bourne et al., 2018). Below the pCd:P maximum, pCd is preferentially remineralised over a second, more labile, P fraction (associated with phospholipid membranes) causing a decrease in pCd:P through the mesopelagic zone (Bourne et al., 2018).

A major difference between vertical profiles of pCd:P in the AAZ and those to the north was the rate at which pCd:P decreased with depth below the subsurface maximum. For example, pCd:P in the AAZ decreased by a factor of 7.5 from a subsurface maximum of 1.2 mmol mol^–1 to 0.16 mmol mol^–1 at 1,000 m. This factor was 3.6 in the PFZ (Figure 6B), 1.7 in the SAZ (Figure 6C) and 0.9 in the STZ (Figure 6D; there was no distinct pCd:P maximum here although a local subsurface maximum at 100 m was evident). Since all vertical pCd:P profiles ultimately converge on relatively consistent deep water (>1,500 m) ratios, stations with higher ratios in surface waters (i.e., PFZ and AAZ) must decrease more rapidly, relative to depth, compared to stations in the SAZ and STZ. Vertical and latitudinal differences in pCd:P profile shapes may simply be related to intracellular Cd concentrations, which vary as a function of the availability of Cd and other metals [see section “Cd Uptake Influenced by the Availability of Dissolved Cd and Other Metals (Fe, Zn, and Mn)”]. In addition, variations in pCd:P profiles are potentially driven by the proportion of diatoms in the resident phytoplankton community. In the AAZ (diatom dominated), preferential remineralisation of pCd below the pCd:P maximum will occur to a greater extent because of the elevated cellular pCd:P in diatoms. In the PFZ (diatom and flagellate dominated), SAZ (flagellate dominated), and STZ (flagellate dominated), the increasing contributions of flagellates, likely with lower cellular pCd:P compared to diatoms, provides a potential mechanism behind the slower rates of pCd remineralisation below the pCd:P maximum. Data pertaining to taxonomic differences in pCd:P are scarce and have not been equivocally demonstrated in the field. For other metals, however, notably for Fe, Mn, and Zn, metal:P ratios of diatoms are higher than flagellates from the Southern Ocean (Twining and Baines, 2013) providing at least some support for a taxonomic effect on pCd:P. In addition to the near-surface pCd:P maximum, there is a secondary, deeper, maximum visible at the northern stations. This maximum deepens from 200 m at 48°S to 400 m at 45°S and 800 m at 41°S (Figures 6B–D) and appears to be spatially associated with isopycnals corresponding to AAIW (Figure 7A). This feature may represent the advection of particles with relatively high pCd:P from the upper water column of the PFZ along the northward subducting flow path of AAIW. However, not all profiles fit the described vertical behavior of pCd:P. For example, in the SAZ (45°S), pCd:P decreased in the upper 150 m, a direct contrast to the other stations. Here, and likely other stations where deviations from general trends are evident, pCd:P may include signatures of dead organic matter, bacteria and possibly small zooplankton which may affect pCd:P ratios. Ultimately, the near-surface pCd:P maxima were consistent with increasing dCd:PO₄ ratios with depth at each station and in agreement with previous suggestions of pCd and P remineralisation length scales impacting ratios of dCd to PO₄ in surface and subsurface waters (Quay et al., 2015; Bourne et al., 2018).