A winter cruise was undertaken on board the R/V SA Agulhas II in July 2017 between Cape Town, South Africa, and the marginal ice zone (MIZ; encountered at 61.7∘ S; de Jong et al., 2018), returning to South Africa along the meridional World Ocean Circulation Experiment (WOCE) I06 transect (30∘ E) (Fig. 1). Sampling was conducted on two legs – between 37 and 62∘ S on the southward leg (Leg 1) and between 59 and 41∘ S on the northward return leg along the WOCE I06 line (Leg 2). During Leg 1, only surface samples were collected, while on Leg 2, the deployment of conductivity–temperature–depth (CTD) hydrocasts allowed for depth-profile sampling.

The positions of the major hydrographic fronts during both legs of the cruise are shown in Fig. 1a. At the hydrocast stations (Leg 2), the mixed-layer depth (MLD) averaged 143 m in the OAZ, 146 m in the PFZ, 205 m in the SAZ, and 113 m in the STZ, which is within the reported climatological range for the western Indian sector of the Southern Ocean in winter (Sallée et al., 2010). Underway ambient NO2- concentrations (Leg 1) ranged from 74 to 232 nM (transect average of 168±48 nM, median of 177 nM) and generally increased with latitude, albeit with a high degree of variability (Figs. 1a, S2). The ambient NO2- concentrations at the hydrocast stations were fairly constant throughout the mixed layer (ranging from 55±35 to 159±73 nM), decreasing rapidly to values below detection by 150–200 m (Fig. 1b). Mixed-layer NO2- showed no clear latitudinal trend mainly because of the anomalously low concentrations measured at St 09 (54∘ S; mixed-layer average of 64±30 nM, compared to 144±56 nM for the seven other hydrocast stations). The NO3- concentrations were also near-homogenous throughout the mixed layer, decreasing from an average of 28.4±0.2 µM at the southernmost station (St 08; 59∘ S) to 3.7±1.1 µM at the northernmost station (41∘ S) and increasing below the mixed layer as expected (Fig. 1c).

The rates of NO3- uptake (ρNO3-) were low and relatively homogenous over the upper 75 m at each station (Fig. S3a). Average euphotic zone ρNO3- increased northwards, from 2.9±1.1 nM d-1 at St 08 in the AZ to 12±2.0 nM d-1 at St 11 in the SAZ, with a transect average of 6.2±3.4 nM d-1. The euphotic zone PON concentrations also increased northwards, from 0.24±0.02 µM at St 08 to 0.47±0.08 µM at St 11 (Fig. S3b). Integrated over the mixed layer, corrNO2ox- accounted for an average of 122 % of ρNO3- (range of 63 % at St 09 to 237 % at St 08; Table S2), consistent with previous observations from the wintertime Southern Ocean (Mdutyana et al., 2020).

Measurements of NO2- oxidation rates are limited in the Southern Ocean, with only two studies that have directly measured this pathway in open ocean waters (Bianchi et al., 1997; Mdutyana et al., 2020). For NO2- oxidation kinetics, there are no data at all for the Southern Ocean. This scarcity of measurements is unsurprising given that in situ NO2- oxidation kinetics studies are generally limited; indeed, to our knowledge, there are only two studies from the coastal ocean (Olson, 1981a; Zhang et al., 2020) and two from the eastern tropical North Pacific oxygen deficient zone (ETNP ODZ; with these experiments conducted across a range of ambient oxygen concentrations; Sun et al., 2017, 2021). By contrast, there exist numerous estimates of NO2- oxidation kinetic parameters determined using cultured marine NOB (e.g., Jacob et al., 2017; Kits et al., 2017; Nowka et al., 2015; Sorokin et al., 2012; Zhang et al., 2020). In general, culture experiments suggest far higher kinetic constants compared to the limited in situ observations from the ocean, particularly for Km (i.e., culture-based Km estimates of 9–544 µM; Blackburne et al., 2007; Nowka et al., 2015; Ushiki et al., 2017).

The high Km values derived for cultured NOB suggest that the affinity of these organisms for NO2- is low. However, this is not what is observed in the environment, which indicates that the most abundant marine NOB are not represented in the culture collection. For the Southern Ocean, we report high substrate affinities of NOB, with Km values ranging from 134 to 403 nM, which is within the range documented for oxygenated coastal and open ocean waters (27–506 nM; Olson, 1981a; Zhang et al., 2020) (Table 2). In the low- to zero-oxygen waters of the ETNP ODZ, similarly low Km values have been reported (254±161 nM; Sun et al., 2017), although values >5 µM have also been observed (Sun et al., 2021), with these latter estimates associated with ambient NO2- concentrations >1 µM. We explore the relationship between ambient NO2- concentration and Km in detail in Sect. 4.2 below. Here, our focus is on the Km values derived under conditions of low ambient NO2- (i.e., < 250 nM) given that (some of) the environmental factors affecting NO2- oxidation at high ambient NO2- concentrations appear to be unique. For example, oxygen has been shown to decrease the rate of NO2- oxidation in the ODZs (Sun et al., 2017, 2021) where novel clades of NOB have been detected (Sun et al., 2021). Additionally, NO2- concentrations in the oxygenated open ocean seldom exceed 250 nM (Zakem et al., 2018), in contrast to the ODZs (Bristow et al., 2016; Füssel et al., 2012).

Across our Southern Ocean transect, Vmax ranged from 5 to 14 nM d-1, which is relatively low compared to estimates from other regions (Table 2), although such a comparison may not be particularly informative as our rates (and typically those of others) are not normalized for NOB abundance. Our Vmax estimates are also low compared to a previous study of mixed-layer nitrification in the winter Southern Ocean (Mdutyana et al., 2020). This difference may be partly due to the fact that the kinetics experiments were conducted using surface (∼7 m) seawater (and thus, the surface NOB community that had been exposed to surface conditions, including elevated light), yet the highest rates of NO2- oxidation typically occur near the base of the mixed layer, including in the Southern Ocean (Fig. 3b–e; Mdutyana et al., 2020; Peng et al., 2018; Sun et al., 2017). The opposite pattern has also been observed, however (although not in the Southern Ocean), with deeper samples yielding a lower Vmax than samples collected in shallow waters (Sun et al., 2017; Zhang et al., 2020).

We report maximum NO2- oxidation rates that generally increase towards the south and with decreasing sea surface temperature (SST; recognizing that these parameters co-vary), although St 01 in the STZ and St 06 and St 07 in the MIZ deviate from this trend (Fig. 4a and b; R2=0.019, p=0.77 and R2=0.12, p=0.45, respectively, when all the stations are considered and R2=0.92, p=0.041 and R2=0.94, p=0.029, respectively, when Sts 01, 06, and 07 are excluded). It is possible that changes in the NOB community (composition and/or abundance) across the transect explain some of the observed variability. Nonetheless, taking latitude as a qualitative proxy for light, it is perhaps unsurprising that the maximum NO2- oxidation rates increase southwards given that NOB are known to be at least partially light inhibited (Olson, 1981b; Peng et al., 2018; Ward, 2005). This explanation does not hold for the stations in the MIZ, however, at which Vmax decreases sharply despite these waters receiving the least light (less than 5 h of weak sunlight versus ∼7 h at 55∘ S to ∼9 h at 37∘ S). The temperature at the MIZ stations was <0∘ C, which raises the possibility of a temperature effect on Vmax. Indeed, we previously observed a strong decline in the Vmax associated with NH4+ oxidation at SSTs <0∘ C in the Southern Ocean, while at SSTs ranging from 0.6 to 16∘ C, Vmax was near invariant (Mdutyana et al., 2022a).

Marine nitrification has been reported to be largely unaffected by temperature variations (Baer et al., 2014; Bianchi et al., 1997; Horak et al., 2013), although NH4+ and NO2- oxidation may respond differently to similar changes in temperature. For example, marine NOB incubated at temperatures ranging from 10 to 35∘ C responded far more slowly to an increase in temperature than co-incubated AOA, resulting in an accumulation of NO2- in the incubation bottles (Schaefer and Hollibaugh, 2017). By contrast, we previously observed no robust relationship between temperature and the maximum NH4+ oxidation rate in the Southern Ocean (Mdutyana et al., 2022a), a finding that is consistent with studies of NH4+ oxidation in the Arctic and temperate coastal ocean (Baer et al., 2014; Horak et al., 2013). Far less work has been done to assess the response of NOB to temperature changes. In the absence of experiments specifically designed to test the response of Southern Ocean NOB to temperature, it is difficult to disentangle the effect(s) on NO2- oxidation of temperature versus light (and possibly other parameters that co-vary with latitude, such as NO2- and/or micronutrient availability).

Plotting Vmax as a function of the ambient substrate concentration ([NO2-]amb) reveals a strong positive relationship for all but the MIZ stations (Fig. 4c; R2=0.73, p=0.065 if the MIZ stations are excluded). In particular, the STZ station (St 01), which appeared anomalous in the plots of Vmax versus latitude and SST, is consistent with the other non-MIZ stations when evaluated in Vmax versus [NO2-]amb space. The positive relationship of Vmax to [NO2-]amb could be taken as evidence that NO2- availability strongly controls the maximum achievable rate of NO2- oxidation. However, Vmax varies four-fold across the transect, while [NO2-]amb only changes by a factor of 2, and [NO2-]amb is also correlated with latitude (R2=0.51, p<0.001 for all surface [NO2-]amb data; Fig. S2). Additionally, previous wintertime Southern Ocean NO2- oxidation rates (albeit not Vmax) showed no relationship with ambient NO2- concentration (Bianchi et al., 1997; Mdutyana et al., 2020). The extent to which Vmax is directly controlled by [NO2-]amb is thus unclear, and it is likely that NOB community composition, light availability, and temperature also play a role, with SST perhaps becoming more important at very low temperatures (i.e., in the MIZ).

Our estimates of Km reveal that NOB in the wintertime Southern Ocean have a high affinity for NO2- that appears to decrease (i.e., the Km rises) at higher latitudes (i.e., lower light) and lower temperatures, with St 01 in the STZ again emerging as an exception (Fig. 4d and e; R2=0.86, p=0.008 and R2=0.86, p=0.008, respectively). Plotting our Km values as a function of [NO2-]amb reveals a strong positive relationship (Fig. 4f; R2=0.83, p=0.004; black data points), implying that NO2- availability rather than temperature or light exerts the dominant control on Km. This trend further suggests that NOB are well-adapted to the environment (or Southern Ocean region) in which they are found. Southern Ocean mixed-layer NO2- concentrations are almost never <150 nM, regardless of the season (Fripiat et al., 2019; Mdutyana et al., 2020; Zakem et al., 2018), yet the relationship of Km to [NO2-]amb also holds at far lower NO2- concentrations. The colored data points in Fig. 4f show Km versus [NO2-]amb for four additional regions where a Michaelis–Menten relationship of NO2- oxidation rate to NO2- concentration was observed and where [NO2-]amb was <250 nM (two coastal ocean sites, the South China Sea (SCS; Zhang et al., 2020) and Southern California Bight (SCB; Olson, 1981a); one oligotrophic ocean site, the subtropical southeast Atlantic (SSA; Sarah Fawcett et al., unpublished data); and two stations from the ETNP ODZ, where oxygen concentrations ranged from 0 to 16.8 µM (Sun et al., 2017)). The robust positive relationship of Km to [NO2-]amb that emerges when these previous results are combined with our Southern Ocean data (R2=0.68, p<0.001) strongly implicates [NO2-]amb as the dominant control on the Km of NO2- oxidation in the ocean, particularly at low [NO2-]amb (i.e., <250 nM).

The production of NO2- from NH4+ oxidation has recently been hypothesized to be vulnerable to iron limitation (Mdutyana et al., 2022a) since AOB rely on iron-rich cytochrome c proteins (Arp et al., 2002; Walker et al., 2010), and some AOA appear to have a low affinity for inorganic iron (Shafiee et al., 2019). NOB also contain iron-rich enzymes, such as nitrite oxidoreductase, which is responsible for converting NO2- to NO3- (Meincke et al., 1992; Spieck et al., 1998). While we have no iron data with which to compare our kinetic parameters, dissolved iron concentrations ([DFe]) were measured throughout the euphotic zone at the depth-profile stations (St 08 to St 11; Mdutyana et al., 2022a). The revised NO2- oxidation rates at these stations are weakly positively correlated with [DFe] (R2=0.35, p=0.016; Fig. 5), indicating a potential role for iron in controlling NO2- oxidation. Combined with the evidence that iron may also constrain marine NH4+ oxidation (Shafiee et al., 2019), this observation implies that mixed-layer nitrification in the Southern Ocean may be iron limited. Since phytoplankton consumption of regenerated NO3- yields no net removal of atmospheric CO2 in a mass balance sense (Dugdale and Goering 1967; Yool et al., 2007), an iron-related control on mixed-layer nitrification would help to limit the extent to which this process can weaken the Southern Ocean's biological pump and would lead to enhanced competition between phytoplankton and nitrifiers for iron.

While still limited, there is growing evidence that marine AOA have a very high affinity for NH4+ (more correctly, ammonia (NH3), the substrate for NH4+ oxidation; Horak et al., 2013; Martens-Habbena et al., 2009; Mdutyana et al., 2022a; Newell et al., 2013; Peng et al., 2016). Marine NOB also appear able to access low concentrations of substrate, based on the few in situ studies conducted to date, including this one (Fig. 4f; Olson, 1981a; Sun et al., 2017; Zhang et al., 2020). This high substrate affinity is perhaps unsurprising given that NO2- concentrations are generally near-zero throughout the oxygenated ocean, rising modestly to values typically <500 nM at the PNM in (sub-)tropical waters (Lomas and Lipschultz, 2006; Zakem et al., 2018) and <400 nM over the mixed layer in (sub-)polar regions (Zakem et al., 2018). The average surface NO2- concentration measured during Leg 1 of our cruise was 168±48 nM (Fig. 1a), and the average mixed-layer concentration for Leg 2 was 137±57 nM (Figs. 1b and 3a). Similar concentrations have been observed previously across the Southern Ocean, including in other seasons (Cavagna et al., 2015; Fripiat et al., 2019; Mdutyana et al., 2020). Thus, while NO2- oxidation in Southern Ocean surface waters is characterized by a low Km, the affinity of NOB for NO2- is apparently not high enough to completely remove the available NO2-.

The persistence of elevated NO2- concentrations in the mixed layer at high latitudes has previously been attributed to the inability of iron- and/or light-limited phytoplankton to fully consume NO2- transported to the surface with NO3- during deep mixing events (Zakem et al., 2018). However, subsurface NO2- concentrations in the Southern Ocean are typically below detection (Figs. 1b and 3a; Olsen et al., 2016), so it is unclear how deep mixing could supply measurable NO2- to the euphotic zone. We thus discount subsurface mixing as a primary explanation for the elevated Southern Ocean mixed-layer NO2- concentrations observed during our study and in other seasons (e.g., Fripiat et al., 2019).

A second possible source of elevated mixed-layer NO2- is efflux following partial NO3- reduction to NO2- by phytoplankton (Lomas and Lipschultz, 2006), which has been extensively documented in laboratory and field studies (see Collos, 1998, for a review). The release of NO2- by phytoplankton is hypothesized to result from light limitation of intracellular NO2- reduction (Kiefer et al., 1976; Vaccaro and Ryther, 1960), short-term increases in irradiance to which phytoplankton cannot adapt (Lomas and Lipschultz, 2006), iron limitation of NO3- assimilation (Milligan and Harrison, 2000), and/or release of phytoplankton from NO3- limitation following a period of starvation (Sciandra and Amara, 1994). While some of these mechanisms may be ongoing in the Southern Ocean, they all require the initial uptake of NO3- by phytoplankton. This process occurs in the winter mixed layer at rates that are too low to support NO2- efflux to the extent that it would allow NO2- to accumulate to concentrations of 100–400 nM (Fig. S3; Mdutyana et al., 2020; Philibert et al., 2015) while simultaneously being removed by NO2- oxidation. Additionally, we observe a reasonable correlation between the NH4+ oxidation rates and the ambient NO2- concentration (R2=0.46, p<0.001; Fig. S4), which implies that NO2- derives mainly from NH4+ oxidation rather than phytoplankton efflux.

A third potential explanation for elevated mixed-layer NO2- is a decoupling of NH4+ and NO2- oxidation, which appears to be widespread in the environment (e.g., Beman et al., 2013; Ward and Zafiriou, 1988). In the oxygenated ocean, NH4+ oxidation has been considered the rate-limiting step in the nitrification pathway because NO2- seldom accumulates in the mixed layer (Kendall, 1998; Kowalchuk and Stephen, 2001; Vajrala et al., 2013; Walker et al., 2010). However, rate measurements from numerous ocean regions show contrasting results, with NO2- oxidation sometimes outpacing NH4+ oxidation (Bristow et al., 2015; Dore and Karl, 1996; Horrigan et al., 1990; Peng et al., 2018), while in other cases, NH4+ oxidation is dominant (Clark et al., 2008; Kalvelage et al., 2013; Ward and Kilpatrick, 1991). The limited data available from previous Southern Ocean investigations show no clear trend (Bianchi et al., 1997; Mdutyana et al., 2020). In the present study, mixed-layer corrNO2ox- rates are two- to seven-times lower than the coincidentally measured corrNH4ox+ (Figs. 3 and 6). Additionally, the maximum rates of NO2- oxidation (Vmax) that we measure in this study for the surface NOB community (∼5 to 13 nM d-1; Fig. 2) are on average half those determined at the same stations for NH4+ oxidation (14 to 23 nM d-1; Mdutyana et al., 2022a). At the time of our sampling, therefore, NO2- oxidation was rate-limiting for nitrification, which likely accounts for at least some of the NO2- accumulated in the Southern Ocean's winter mixed layer.

If a decoupling of NH4+ and NO2- oxidation is predominantly responsible for NO2- accumulation, an obvious question is why these rates are not balanced. Environmental factors like temperature and light may play a role (Ward, 2008), as may iron limitation and the different ecophysiologies of NH4+ and NO2- oxidizers. AOA have been shown to adapt more rapidly than NOB to a change in temperature (Schaefer and Hollibaugh, 2017); however, seasonal SST changes within the various zones of the Southern Ocean are fairly small and the aforementioned study showing the differential thermal response of AOA and NOB was conducted at higher temperatures than those experienced in much of the Southern Ocean. With regards to light, there is evidence from culture and field studies that NOB are more photosensitive than AOA and AOB (Bock, 1965; Olson, 1981b; Qin et al., 2014). Our data are consistent with this notion insofar as the Vmax associated with NO2- oxidation in surface waters rises with increasing latitude (and thus decreasing light; Fig. 4a), while the Vmax derived for NH4+ oxidation remains largely unchanged across >30∘ of latitude (Mdutyana et al., 2022a). However, the ambient NO2- concentration in Southern Ocean surface waters rises near linearly with latitude (Fig. S2a), while the NH4+ concentration resembles a step function, increasing from ∼100 nM north of the SAF to ∼700 nM south of the SAF, over a distance of roughly 1∘ of latitude (Fig. S2b). The differing trends in Vmax may thus have more to do with substrate availability than photoinhibition.

Mixing, particularly deep winter overturning, might also contribute to a decoupling of NH4+ and NO2- oxidation. In coastal waters, deep mixing has been shown to dilute the nitrifier community, with AOO subsequently observed to recover more rapidly than NOB. This differential rate of recovery has been hypothesized to result in a period of low rates of NO2- oxidation during which the co-occurring NH4+ oxidation rates remain elevated, ultimately causing NO2- to accumulate in the surface layer (Haas et al., 2021). While a similar effect may play a role in NO2- accumulation in the open Southern Ocean, it is unlikely that the entire NO2- reservoir can be attributed to this process. The rates of NH4+ oxidation are only slightly higher than the NO2- oxidation rates in the winter mixed layer (Fig. 3), and the mixed-layer NH4+ concentrations are elevated (Fig. 3f). These observations imply that NH4+ oxidizers are limited by something other than the NH4+ substrate. This limitation prevents AOO from catalyzing higher rates of NO2- production (and thus NO2- accumulation).

Nitrite oxidoreductase (NXR), the enzyme possessed by NOB that is responsible for aerobic NO2- oxidation to NO3-, is an iron–sulfur molybdoprotein (Lücker et al., 2010; Meincke et al., 1992; Sundermeyer-Klinger et al., 1984). As such, NO2- oxidation has a significant iron requirement (Bayer et al., 2021; Saito et al., 2020), intimated by the relationship we observe between corrNO2ox- and DFe (Fig. 5). Additionally, NO2- accumulation at the PNM in the California Current has been hypothesized to be caused by the iron limitation of NOB (Santoro et al., 2013). AOB also require iron, in particular for the oxidation of hydroxylamine, which is catalyzed by the heme-rich hydroxylamine oxidoreductase complex (Arp et al., 2002; Walker et al., 2010). By contrast, AOA, the dominant marine NH4+ oxidizers, rely mainly on copper-containing proteins to mediate NH4+ oxidation (Amin et al., 2013; Santoro et al., 2015; Walker et al., 2010). In the iron-limited Southern Ocean, it is thus possible that iron scarcity more strongly limits NO2- than NH4+ oxidation. However, recent culture and proteomic work suggests that some AOA may actually have a high iron requirement (Carini et al., 2018; Qin et al., 2018; Santoro et al., 2015; Shafiee et al., 2019), and we have previously hypothesized an iron-related control on NH4+ oxidation in the Southern Ocean (Mdutyana et al., 2022a). Deeper investigation is thus required to characterize the role of iron in controlling the relative rates of NH4+ and NO2- oxidation, as well as the implications for the complete nitrification pathway.

A further consideration is differences in the ecology of AOA and NOB. Marine NOB are an order of magnitude less abundant than AOA (e.g., Beman et al., 2013; Damashek et al., 2019; Füssel et al., 2012; Kitzinger et al., 2020; Pachiadaki et al., 2017) and roughly 3-times larger (Könneke et al., 2005; Martens-Habbena et al., 2009; Pachiadaki et al., 2017; Watson and Waterbury, 1971). While marine NOB appear to have a high affinity for ambient NO2-, the in situ Km values derived to date are not as low as those reported for NH4+ oxidation (Horak et al., 2013; Mdutyana et al., 2022a; Peng et al., 2016; Xu et al., 2019; Zhang et al., 20202), which is perhaps to be expected given the larger size of NOB versus AOA. Resource limitation theory posits that nitrifiers (NOB and AOA) require a subsistence concentration of substrate (R∗) to maintain their population and that those with the lowest R* will outcompete all other organisms limited by the same resource, provided that their Vmax is higher than their loss rate due to grazing and/or viral lysis (Zakem et al., 2018). Because NOB are larger than AOA, they will have a higher R∗ even before grazing pressure is factored in. Their larger size also means that NOB are more likely to be grazed than AOA, which will further increase their R∗, as will the fact that their maximum growth rates are low and thus vulnerable to being outpaced by their loss rate. Taken together, these factors will increase R∗, potentially resulting in the accumulation of NO2- in the water column, and may help to explain why the Km for NO2- oxidation, in the Southern Ocean and elsewhere, is considerably higher than the Km derived for NH4+ oxidation. Additionally, the fact that NOB will be preferentially grazed over AOA may contribute to NO2- oxidation being rate-limiting for nitrification.

That NO2- oxidation was rate-limiting at the time of our sampling does not necessarily explain the accumulation of NO2- in the Southern Ocean mixed layer year-round. Neither NH4+ nor NO2- oxidation occurs at elevated rates in summer or autumn (Bianchi et al., 1997; Mdutyana et al., 2020), yet the elevated NO2- concentrations persist during these seasons (Cavagna et al., 2015; Fripiat et al., 2019; Mdutyana et al., 2020). To fit a Michaelis–Menten function to our experimental data required amending the classical equation (Eq. 2) to allow for a positive x intercept (i.e., a non-zero S value at which V was still zero, the C parameter in Eq. 3) (Archontoulis and Miguez, 2014). Additionally, at most stations, the NO2- oxidation rates did not increase substantially following the initial two or three substrate amendments (i.e., in Fig. 2, the slope of the relationship between V and S is less steep for the initial two to three values of S than at higher S values). Practically, our findings suggest that Southern Ocean NOB require a minimum (i.e., “threshold”) NO2- concentration below which NO2- becomes severely limiting. Coupled with weak NO2- drawdown by iron- and/or light-limited phytoplankton during their incomplete consumption of the NO3- + NO2- pool, a threshold substrate requirement of NOB can explain the year-round persistence of non-zero mixed-layer NO2- since it implies that there is no mechanism by which NO2- can be completely exhausted.

The existence of a NO2- concentration threshold may indicate limitation of the membrane-bound NXR enzyme, either by NO2- or by another essential nutrient. Recently, using NXR concentrations, estimates of NXR-specific activity, and direct measurements of in situ NO2- oxidation rates, Saito et al. (2020) deduced that Nitrospina NXR is undersaturated with NO2- in the tropical Pacific possibly due to iron limitation. The authors suggest that under iron-scarce conditions, it becomes increasingly difficult for NOB to synthesize NXR and thus to oxidize NO2-. A similar dynamic may be at play in the Southern Ocean, with limited synthesis of NXR at low iron concentrations resulting in a decrease in the efficiency of the NO2- oxidation pathway that manifests most strongly when the ambient NO2- concentration is also low. This inefficiency could be alleviated at higher NO2- concentrations since NOB (even with a paucity of NXR) are less likely to experience diffusion limitation with respect to NO2- when there is more of this substrate available (Pasciak and Gavis, 1974). Regardless of its mechanistic basis, limitation of NOB NXR would help to explain the perennially high concentrations of NO2- in the Southern Ocean mixed layer. Moreover, environmental factors unique to the Southern Ocean, such as limited iron availability, may be instrumental in setting the NO2- threshold and associated elevated mixed-layer NO2- concentrations.

Our observations raise the question of why a similar NO2- concentration threshold has not been reported for other ocean regions, particularly those characterized by similar conditions to the Southern Ocean. This may partly be due to the very limited number of NO2- oxidation kinetics experiments that have been conducted in the open ocean and/or to the fact that a classic Michaelis–Menten function is usually imposed upon kinetics data, with V assumed to increase as soon as S>0. Additionally, depending on the maximum substrate concentration added during kinetics experiments (i.e., the maximum concentration on the x axis of the V versus S plot), it can be difficult to discern a possible threshold NO2- concentration by simply examining the plots. Inspection of published Michaelis–Menten curves does reveal the possibility of a non-zero C value in some cases, including in the ETNP ODZ (Sun et al., 2021) and associated with the PNM in the South China Sea (Zhang et al., 2020). However, there are also published curves that clearly intercept the origin in V versus S space (Olson, 1981a; Sun et al., 2017), underscoring the need for further investigation of the conditions that lead to a threshold NO2- concentration requirement of NOB.