Epipelagic nitrous oxide production offsets carbon sequestration by the biological pump

The removal of carbon dioxide from the atmosphere by the marine biological pump is a key regulator of Earth’s climate; however, the ocean also serves as a large source of nitrous oxide, a potent greenhouse gas and ozone-depleting substance. Although biological carbon sequestration and nitrous oxide production have been individually studied in the ocean, their combined impacts on net greenhouse forcing remain uncertain. Here we show that the magnitude of nitrous oxide production in the epipelagic zone of the subtropical ocean covaries with remineralization processes and thus acts antagonistically to weaken the radiative benefit of carbon removal by the marine biological pump. Carbon and nitrogen isotope tracer incubation experiments and nitrogen isotope natural abundance data indicate enhanced biological activity promotes nitrogen recycling, leading to substantial nitrous oxide production via both oxidative and reductive pathways. These shallow-water nitrous oxide sources account for nearly half of the air–sea flux and counteract 6–27% (median 9%) of the greenhouse warming mitigation achieved by carbon export via the biological pump. Substantial nitrous oxide production in the epipelagic zone of the subtropical ocean partially offsets carbon sequestration by the marine biological pump, according to observations from the South China Sea and subtropical northwest Pacific.

The removal of carbon dioxide from the atmosphere by the marine biological pump is a key regulator of Earth's climate; however, the ocean also serves as a large source of nitrous oxide, a potent greenhouse gas and ozone-depleting substance. Although biological carbon sequestration and nitrous oxide production have been individually studied in the ocean, their combined impacts on net greenhouse forcing remain uncertain. Here we show that the magnitude of nitrous oxide production in the epipelagic zone of the subtropical ocean covaries with remineralization processes and thus acts antagonistically to weaken the radiative benefit of carbon removal by the marine biological pump. Carbon and nitrogen isotope tracer incubation experiments and nitrogen isotope natural abundance data indicate enhanced biological activity promotes nitrogen recycling, leading to substantial nitrous oxide production via both oxidative and reductive pathways. These shallow-water nitrous oxide sources account for nearly half of the air-sea flux and counteract 6-27% (median 9%) of the greenhouse warming mitigation achieved by carbon export via the biological pump.
The ocean plays a crucial role in the global climate system through modulating atmospheric greenhouse gases by absorbing nearly 30% of anthropogenic carbon dioxide (CO 2 ) and releasing 20% of total nitrous oxide (N 2 O) emissions to the atmosphere 1,2 . The marine biological pump (defined as the biologically driven processes that transfer carbon from the surface ocean to the ocean's interior) is the dominant mechanism driving long-term CO 2 exchange across the air-sea interface and plays a critical role in regulating atmospheric CO 2 and climate [3][4][5] . The efficiency of the marine biological pump, defined by the ratio of carbon export to net primary production at a specific reference depth (for example, the base of the euphotic zone), is often estimated to be ~10% in the global ocean 5 , and even lower in the subtropical oceans 6,7 . In the subtropical oceans, a large fraction of newly produced organic material undergoes remineralization in the upper 200 m (the epipelagic zone), resulting in rapid carbon and nitrogen transformations between organic and inorganic forms. Numerous studies on the ocean's biological pump have focused on the magnitude and controls of carbon removal; however, the potential counter effect of N 2 O emissions in offsetting the radiative effect of CO 2 removal by the biological pump has been largely overlooked.
In the marine nitrogen cycle, N 2 O is mainly produced as a by-product of nitrification and as an intermediate during denitrification, both of which are largely controlled by organic matter supply and remineralization 8 . In the oxygenated ocean, nitrification is considered the dominant source of N 2 O (refs. [8][9][10]. However, recent studies on marine N 2 O suggest the sources are more complex, including both aerobic nitrification and anaerobic nitrate (NO 3 − ) and nitrite (NO 2 − ) reduction in anaerobic micro-niches associated with marine aggregates 11 or zooplankton guts 12 . Ammonia-oxidizing archaea, the dominant ammonia oxidizer in the open oceans 13 , utilize a hybrid N 2 O production pathway that is distinct from that of their bacterial counterparts, in which intermediates sourced from ammonium (NH 4 + ) and NO 2 − co-contribute to N 2 O formation 14 . The physiological and Article https://doi.org/10.1038/s41561-022-01090-2 productivity, where regeneration drives recycling (including N 2 O production), while export determines organic carbon removal to depth. Intensive organic matter remineralization not only reduces CO 2 sequestration efficiency, but also contributes to N 2 O production via nitrogen recycling. The potential counter effect of N 2 O emission in offsetting the radiative effect of CO 2 sequestration via export has been investigated in a few geoengineering and nitrogen deposition experiments [22][23][24] ; however, the source of N 2 O, and links between biological CO 2 sequestration and N 2 O production, have not been quantified. We hypothesized that in the vast subtropical oligotrophic oceans, where export production is inefficient [3][4][5][6][7] , the counter effect of N 2 O emission on carbon removal by the biological pump could be substantial. We measured rates of N 2 O production and carbon export in the epipelagic ocean extending from the South China Sea (SCS) into the North Pacific Subtropical Gyre (NPSG) during seven cruises conducted over eight years (Supplementary Fig. 1 and Supplementary Tables 1 and 2). We demonstrate that a large fraction of N 2 O in the surface ocean is locally produced in the epipelagic waters via both oxidative and reductive pathways. Furthermore, the magnitude of this shallow N 2 O source appears to covary spatially with biological productivity, offsetting greenhouse warming mitigation achieved by carbon export. enzymatic details of this pathway remain mostly unresolved 15,16 . Moreover, sources of N 2 O production to the epipelagic ocean are less well studied than in the mesopelagic waters, despite the fact that remineralization of nitrogen is more intensive in the epipelagic ocean. This lack of investigation stems from the recognition that ammonia oxidation is inhibited by light 17 and the abundance of ammonia-oxidizing archaea is relatively low in the upper ocean 13 . Hence, the pathways, controls, relative contributions from various N 2 O sources to the N 2 O pool and linkages between primary productivity and N 2 O production in the epipelagic ocean remain unclear, hampering our ability to constrain the role of the ocean in the atmospheric N 2 O budget.
Nitrogen is a primary limiting nutrient to phytoplankton growth over much of the low-latitude oceans 18,19 . Supply of exogenous sources of nitrogen to the euphotic zone can control net productivity in these waters, with new nitrogen introduced via NO 3 − supply from subsurface, N 2 fixation and nitrogen deposition 20 . In addition, a large fraction of primary productivity is controlled via nitrogen supplied from remineralization within the epipelagic zone, supporting regenerated production 20,21 . Under the assumption of steady state, new production should be quantitatively related to export of material out of the epipelagic zone into the interior waters. Thus, regenerated and export production are two competing sides of biological  Incorporating N 2 O production that accompanies organic matter remineralization is necessary to quantify net climatic consequences associated with the marine biological pump.

Low export efficiency in the oligotrophic subtropical ocean
The epipelagic waters at all study sites exhibited characteristics typical for the thermally stratified SCS and the NPSG, with high oxygen and low nutrient concentrations throughout the shallow mixed layer (Extended Data Figs. 1 and 2 and Supplementary Table 3). NH 4 + concentrations were persistently low (mean ± s.d., 24.4 ± 23.6 nmol l −1 in the SCS and 21.0 ± 24.3 nmol l −1 in the NPSG) with occasional peaks of 50-200 nmol l −1 below the mixed layer. By contrast, NO 2 − concentrations consistently showed a primary nitrite maximum (PNM) slightly below the deep chlorophyll maximum (DCM). A prominent maximum in the NO 3 − to Si concentration ratio was also frequently observed in the vicinity of the PNM layer. Particulate organic carbon (POC) and particulate nitrogen (PN) concentrations generally decreased with depth, and concentrations were greater in the shelf region than the open ocean (Extended Data Fig. 3).   Rates of primary production in the nutrient-depleted waters were consistently low (<1 μmol C l −1 d −1 ) and decreased with depth at both the SCS and the NPSG stations (Extended Data Fig. 4). Depth-integrated (0-125 m) primary production ranged from 10.4 ± 0.4 to 44.1 ± 3.0 mmol C m −2 d −1 , similar to long-term observations at the NPSG Station ALOHA and the BATS station in the Atlantic 6,7 . Carbon export at 200 m in the SCS and NPSG stations ranged from 0.3 ± 0.1 to 2.0 ± 0.4 mmol C m −2 d −1 , with low export ratios (averaging 3.9 ± 2.3%; Table 1 and Supplementary Discussion 1). Such results are typical of the subtropical oceans, where most primary production undergoes rapid remineralization, fuelling intensive and rapid nitrogen recycling in the epipelagic zone 7,25 .  Table 4), in agreement with previous observations in the subtropical oceans [26][27][28] .

A large shallow source of N 2 O in the epipelagic zone
Our high-resolution vertical profile sampling of the epipelagic waters provided insights into the vertical variations in N 2 O concentrations. N 2 O concentrations and the resulting air-saturation states generally increased with depth. Distinctive N 2 O concentration peaks, which deviate from simple vertical mixing, were observed at four stations ( Fig. 1 and Extended Data Fig. 5). At these sites, peaks in N 2 O concentrations occurred within narrow depth intervals of 10-20 m and could be easily missed by coarser vertical sampling resolution.
The location of these N 2 O peaks varied in depth, temperature, salinity and density (Extended Data Fig. 6), but consistently overlapped with the PNM and NO 3 − /Si maximum layers (Fig. 1). The proximity of an N 2 O peak to the PNM suggests a spatial coupling between N 2 O accumulation and intensive nitrogen recycling at or around the PNM layer, where high ammonia oxidation rates occur 29 , leading to enhanced N 2 O production and accumulation of NO 2 − (refs. 30,31 ). The NO 3 − /Si maximum provides additional evidence of intensive remineralization of organic nitrogen and subsequent nitrification at this depth 32,33 , reinforcing the contribution of local sources to the N 2 O accumulation.
High-resolution vertical profiles of N 2 O stable isotopes (δ 15 N-N 2 O and δ 18 O-N 2 O) at additional stations provided further evidence for near-surface in-situ production of N 2 O in the subtropical ocean (Fig. 2). The dual isotopes of N 2 O covaried with depth, decreasing from near equilibrium with the atmosphere in the near-surface waters (average δ 15 N = 7.4 ± 0.4‰; δ 18 O = 45.3 ± 0.6‰), to minima in both isotope ratios in the vicinity of the NO 3 − /Si maximum and the PNM (minima were 0.4 ± 0.2‰ to 4.9 ± 0.3‰ and 0.5 ± 0.4‰ to 4.0 ± 0.4‰ lower than the values in surface water for δ 15 N and δ 18 O, respectively). Below the PNM, the δ 15 N and δ 18 O values of N 2 O increased with depth. These results indicate that the prominent dual isotope minima around the PNM probably do not derive from vertical mixing of surface and deep waters. Lateral advection of water with low isotopic signatures is also unlikely because the isotopic minimum layers occurred within waters of varying density and salinity (Extended Data Fig. 7). Hence, the most likely cause for the local δ 15 N and δ 18 O minima is in-situ N 2 O production, as has previously been observed at Station ALOHA in the NPSG 10,34 . The widespread dual isotope minima in our study area (12 out of 15 stations) further suggest that shallow N 2 O production is ubiquitous in the subtropical ocean. The shallowness of the feature in the SCS is noteworthy (60-120 m in the epipelagic zone versus ~300 m in the upper mesopelagic zone at ALOHA 10,34 ). This shallow feature points to a potentially sensitive  climatic consequence of nitrogen recycling along ocean margins due to the shorter distance from the isotope minimum layer to the air-sea interface and more vigorous physical dynamics in the epipelagic ocean.
The relative contribution of in-situ N 2 O production from the isotope minimum layer to air-sea flux can be constrained using an isotope mass balance model under the assumption that N 2 O in the isotopic minima layer derives from a mixture of N 2 O diffusing from the concentration maximum layer and locally produced N 2 O (ref. 10 ). Applied to our stations, this two-component model reveals that shallow N 2 O production contributes 41.6 ± 21.0% by using δ 15 N mass balance and 31.3 ± 11.0% by using δ 18 O mass balance (Supplementary Discussion 2  and Supplementary Table 5), implying the shallow N 2 O source is a substantial contributor to air-sea flux in the oligotrophic oceans.

Multiple biological N 2 O sources in the oxygenated water
The underlying mechanisms that cause the N 2 O isotope minimum are not fully resolved. Nitrification was previously considered a primary N 2 O source to the well-oxygenated open ocean [8][9][10] . However, later observations of δ 18 O-N 2 O and isotope labelling incubations suggest part of the N 2 O in the isotope minima may be produced through nitrifier denitrification or denitrification in particle-associated microenvironments 34 Table 6), which is lower than the reported δ 15 N-NO 3 − (4.8 ± 0.3‰) and δ 15 N-PN (4.2 ± 1.0‰) at the base of the euphotic zone in the SCS 36 . These results indicate NO 2 − could be an important precursor to N 2 O, contributing to the dual isotope minima observed in the PNM layer. Nevertheless, because NO 2 − can be incorporated into N 2 O via denitrification, nitrifier denitrification or the hybrid pathway, the relative contribution of these potential sources cannot be determined using natural abundance data alone.
We conducted a set of 15 N isotope tracer incubations aimed at identifying sources and quantifying their relative contributions to N 2 O production. Results from these experiments show that multiple precursors contribute to N 2 O production in the epipelagic ocean (Fig. 3). Notably, N 2 O production was sometimes detected in the upper mixed layer even though ammonia oxidation rates were below detection limits (Extended Data Fig. 8). Gross N 2 O production increased with depth to a maximum in the vicinity of the N 2 O isotopic minima layer, providing additional evidence for active in-situ N 2 O production. Both NH 4 + and   15 NO 3 − labelling incubations suggests production of N 2 O via nitrifier denitrification and/or denitrification in micro-anoxic niches in the oxygenated ocean 11,12 . As nitrification is widely used as a key component for model parameterization to estimate N 2 O production in the oxygenated ocean 8,9,41 , our results strongly support the contribution of multiple precursors and pathways of N 2 O in the epipelagic ocean that need to be considered in biogeochemical models aiming to estimate marine N 2 O sources and air-sea flux.

N 2 O production offsets CO 2 removal by the biological pump
The export of organic matter to the ocean's interior through the marine biological pump is a primary control on the oceanic CO 2 sink on long-term timescales [3][4][5] . However, the magnitude of the marine biological pump depends on complex interactions, including those that alter the vertical length-scale of organic matter remineralization, altering the timescales over which carbon is sequestered 42 . Rapid (days to weeks) remineralization of organic matter and concomitant nitrogen recycling in the epipelagic zone sustains a large fraction of biological productivity throughout the subtropical oceans [43][44][45] . Our results highlight that this process also promotes N 2 O production. We observed that rates of N 2 O production and ammonia oxidation were significantly positively Article https://doi.org/10.1038/s41561-022-01090-2 correlated (Fig. 4a), highlighting the covariance between N 2 O production and nitrogen regeneration in the epipelagic zone. Moreover, both the integrated ammonia oxidation rates and N 2 O production rates were also significantly correlated with the POC and PN inventories (Fig. 4b,c and Extended Data Fig. 9), implying the strength of nitrogen recycling and N 2 O production scales with the availability of organic nitrogen. Particles are known hotspots of microbial metabolism and can provide a source of organic and inorganic substrates to the surrounding seawater. The microenvironment formed in the particle also favours various nitrogen transformation pathways contributing to N 2 O production 11 . In our study, larger POC and PN stocks, presumably sustained by efficient epipelagic recycling, seem to promote N 2 O production. Carbon and nitrogen cycling are intimately coupled to each other because both elements are required by all organisms. Although the regulation of carbon cycling by nitrogen supply has been extensively studied, we provide new perspectives on potential climatic impacts associated with nitrogen recycling and N 2 O production in the epipelagic ocean. Such shallow N 2 O production is of particular importance in the vast subtropical oligotrophic oceans, where export efficiencies are low and nutrient recycling is rapid. For example, the average export ratio in our study was 3.9 ± 2.3% (Table 1), suggesting >95% of primary production was remineralized in the epipelagic zone. Our observations indicate that nitrogen recycling promotes production of N 2 O via multiple pathways in these remineralization-intensive systems. In comparing the potential offset in radiative warming due to N 2 O production relative to carbon export, we assumed a 100-year time horizon of global warming potential (GWP 100 ) for both processes (where 1 mol N 2 O would be equivalent to 300 mol CO 2 in radiative energy) 1 . We estimate the integrated N 2 O production rate associated with nitrogen recycling would be equivalent to offsetting 5.6 ± 0.6% to 27.2 ± 6.7% (median value 8.8%) of the greenhouse gas mitigation capacity supported by carbon export measured at the SCS and the NPSG stations ( Fig. 5 and Table 1). However, there are uncertainties associated with this estimate, variation in time and length-scales of particle remineralization and water mass ventilation would alter these radiative warming offset estimates. Therefore, our estimates would probably fall at the lower end of the potential offset attributable to N 2 O production in this region. For example, ventilation times between 200 and 300 m in this region average 32 ± 5 years, with that age increasing to 50 years at 400 m and 100 years at 500 m 46 . Assuming the vertical attenuation of sinking particulate matter follows a power-law function 47 , we estimate ~57 ± 5% of the measured exported carbon would be remineralized above 500 m and could exchange with the atmosphere in <100 years, leading to less CO 2 sequestration and a higher N 2 O offset value (Supplementary Discussion 4). This offset of the effectiveness of the CO 2 sink, attributable to a largely overlooked epipelagic N 2 O source, requires re-examination of the warming mitigation capacity of the marine biological pump.
Ongoing warming of the ocean and atmosphere may lead to a decline in export efficiency and decreased length-scale of remineralization due to intensified upper-ocean stratification and shifting of phytoplankton communities towards smaller cells 48 . Together with increased temperatures, these dynamics may enhance organic matter recycling in the epipelagic ocean 49,50 , with concomitant impacts of N 2 O production 41 . Our study suggests enhancement of surface N 2 O production, through intensified organic matter remineralization, could further exacerbate warming of the climatic system through decreased export and greater N 2 O production. Our results were derived from a limited number of stations at one time and cannot be directly extrapolated to the large spatial-temporal variation in both carbon export and nitrogen regeneration in the ocean 3,6,7 . Nevertheless, our findings show that active N 2 O production, driven by intense organic matter recycling in the epipelagic ocean, can offset a considerable fraction of the benefit of radiative forcing achieved by CO 2 sequestration via the marine biological pump. Future work should investigate and compare the N 2 O/CO 2 offset between systems with different export efficiencies. A better integrated assessment should take N 2 O generation into account for understanding the climatic impact of the marine biological pump in order to devise the best greenhouse gas mitigation strategy.

Online content
Any methods, additional references, Nature Portfolio reporting summaries, source data, extended data, supplementary information, acknowledgements, peer review information; details of author contributions and competing interests; and statements of data and code availability are available at https://doi.org/10.1038/s41561-022-01090-2. During the operation of biological pump processes, only a small fraction of newly produced organic matter (in our study, only around 5% of primary production) was exported to depth. This export (EP) drives slow recycling of C and N, during which N 2 O and CO 2 are produced and accumulate at longer timescales (the residence time (τ): centuries to millennia) before exchanging with the atmosphere. In contrast, most of the newly produced organic matter is rapidly remineralized (τ: days to weeks) in the epipelagic ocean to drive intense recycling of C and N, during which N 2 O is produced through both oxidative and reductive pathways and emitted to the atmosphere. This more rapid recycling in the epipelagic waters can offset a substantial part (6-27%, median 9%) of the decreased radiative forcing (GWP 100 ) due to biological CO 2 removal to depth in our study.  Fig. 1 and Supplementary Table 1). Temperature, salinity, depth and fluorescence concentrations were measured using a Seabird SBE 911 CTD sensor package equipped with fluorometer. Photosynthetically active radiation (PAR) was measured using PAR sensors (LI-COR Biosciences, LI-193 on RV Dongfanghong II and Biospherical QCP-2300L-HP on RV Tan Kah Kee). Discrete seawater samples were collected using 24 12-litre Niskin bottles mounted to the conductivity, temperature and depth (CTD) rosette. The base of the mixed layer was defined as the depth where a difference of 0.8 °C relative to the surface value was observed 51 . The nitracline was derived as the mid-point (average) of the steepest nitrate concentration gradient with depth 51 . The depth with 0.1% surface PAR was defined as the base of euphotic zone 6 .
On Samples for chemical, biological and rate measurements were collected from the same casts. Triplicate 150 ml high-density polyethylene Nalgene bottles were used for nutrient collection; 120 ml glass serum bottles were used to collect samples for subsequent N 2 O concentration measurements in 2012 and 2013; and triplicate 250 ml glass serum bottles (Wheaton) were used for subsequent N 2 O concentration and isotope measurements from 2014 to 2018. Ammonia oxidation and N 2 O production rate incubations were conducted in 120 ml glass serum bottles. Seawater for subsequent analyses of POC and PN were collected into 4 l polycarbonate Nalgene bottles. All bottles and equipment were acid washed and rinsed with in-situ seawater at least three times prior to sample collection. During sample collection, glass sample bottles were overfilled two to three times before sealing with 20 mm butyl stopper and aluminium crimp seals (Wheaton). Samples were preserved by adding 0.1 ml to 0.2 ml saturated HgCl 2 and were stored at 4 °C. For POC and PN samples, 4-8 l of seawater was gently (<200 mm Hg, 26.6 kPa) filtered through a pre-combusted (450 °C for 4 h) Whatman GF/F filter (25 mm diameter). After filtration, the filters were folded and wrapped in pre-combusted (450 °C for 4 h) aluminium foil and stored at −80 °C.
A comprehensive set of incubations was carried out on-board to determine rates of nitrification and N 2 O production using 15 Table 2). All incubations were conducted in the dark at near in-situ temperatures (±2 °C). On the 2015 cruise, 0.2 ml of tracer was injected into each bottle to obtain final concentrations of 15 15 NO 2 − and 15 NO 3 − labelling incubations, respectively. 15 NH 4 + labelling incubations were conducted for deriving both N 2 O production and nitrification rates. Immediately after the 15 NH 4 + tracer injection, 10 ml of sample was pushed out by pure N 2 and then filtered through a 0.2 μm syringe filter to represent the initial condition (t 0 ) for the nitrification incubations. The remaining water was preserved with 0.1 ml saturated HgCl 2 . The remaining bottles were incubated in the dark at near in-situ temperature. At each timepoint, 10 ml of water was sampled and then filtered for subsequent nitrification rate measurements, and the remaining water was preserved using HgCl 2 for subsequent determinations of N 2 O. The filtrate was stored at −20 °C for subsequent analyses. For 15 NO 2 − and 15 NO 3 − incubations, the same procedures were used for on-board incubation, except that the incubation was terminated by adding 0.1 ml HgCl 2 without replacing by N 2 . Primary production rate was also measured in selected stations on 2015, 2018 and 2019 cruises using H 13 CO 3 − tracer (99 atom% 13C, Cambridge Isotope Laboratories), and the final concentration of H 13 CO 3 − was 100 μmol l −1 , accounting for ~5% of the substrate pool. On-deck incubation (duplicates) was performed in 4 l polycarbonate Nalgene bottles for 24 h. Light conditions of the incubators were manipulated by neutral density filter. Seawater was gently (<200 mm Hg, 26.6 kPa) filtered through a pre-combusted (450 °C for 4 h) Whatman GF/F filter (25 mm diameter) and stored at −80 °C.

Nutrient, POC and PN measurements
NH 4 + concentrations were measured on-board the research vessels using a fluorometric method with detection limit of 1.2 nmol l −1 and precision of ±3.5% 52 . Nutrient concentrations below the nitracline were measured using a four-channel Continuous Flow Technicon AA3 Auto-Analyzer. The detection limits for NO x (NO 3 − + NO 2 − ) and Si(OH) 4 were 0.03 μmol l −1 and 0.05 μmol l −1 , respectively, with precision better than 1% and 2.8%, respectively 53 . NO 2 − and NO 3 − concentrations above the nitracline were determined using the standard colorimetric method coupled with a Flow Injection Analysis-Liquid Waveguide Capillary Cell system (World Precision Instruments) 54 ; the detection limit was 5 nmol l −1 and precision was better than 3.1%. For POC and PN concentration measurement, the filters were freeze dried and then acidified with 1 ml of 1 N HCl solution to remove carbonates. All filters were dried at 60 °C for 48 h. The decarbonated samples were then analysed for POC and PN using an EA-IRMS (Thermo Finnigan Flash EA 2000 interfaced to a Delta V PLUS isotopic ratio mass spectrometer) system. The precision for both PN and POC concentration is <1% (ref. 55 ).

N 2 O concentration measurement
During 2012 to 2013, N 2 O concentrations were measured using a purge and trap system coupled with a gas chromatograph (Hewlett-Packard model 6890 equipped with a micro-electron capture detector). Calibration of N 2 O concentrations was determined from peak areas with standard gases of 1.0-5.0 ppmv N 2 O/N 2 (Research Institute of China National Standard Materials), which were run at six-sample intervals. The precision of this method was estimated to be better than ±5% (ref. 56 ). Beginning in 2014, N 2 O concentrations were also derived from ion peak area (m/z = 44) during isotope analysis using the gas chromatography-isotope ratio mass spectrometry (GC-IRMS) system (see below). The two methods yielded comparable results; thus, N 2 O concentrations are shown as the mean value from these independent methods.

Th measurement
The thorium-deficit method was used to estimate export production. Total 234 Th samples were processed using a manganese oxide co-precipitation technique 57 . Briefly, total 234 Th in the seawater was co-precipitated with MnO 2 particles and the resulting particles were collected on a 25 mm, 1.0 μm quartz micro-filter (QMA). Suspended particles in the seawater were also analysed for 234 Th; for these samples, ~8 l water was filtered onto a QMA filter. All total and particulate 234 Th samples were beta counted on a gas flow proportional low-level Risø beta-counter for 16 h until total counts >2,500. A second counting was carried out after >150 days for background correction. The recovery for 234 Th was monitored by 230 Th spike addition in the seawater and Article https://doi.org/10.1038/s41561-022-01090-2 quantified by an alpha-counter with addition of a 228 Th internal standard 58 . The recoveries of 234 Th were better than 90%. 238 U (dpm l −1 ) was calculated from the linear relationship of 238 U with salinity 59 .

Isotopic analyses of NO x − and N 2 O
δ 15 N of NO x − samples for nitrification rate were determined using the denitrifier method 60,61 . Briefly, NO x − was quantitatively converted to N 2 O using the bacterial strain Pseudomonas aureofaciens. The evolved N 2 O was then introduced to the GC-IRMS (Delta V PLUS isotopic ratio mass spectrometer) through an online N 2 O cryogenic extraction and purification system. δ 15 N of NO x − values were calibrated against nitrate isotope standards USGS 34, IAEA N3 and USGS 32, which were run before, after and at ten-sample intervals. Accuracy was better than ±0.2‰ according to analyses of these standards at an injection level of 20 nmol N. For samples with NO x − concentrations lower than 0.5 μmol l −1 , 1 ml of 5 μmol l −1 of in-house NO 3 − standard was added as carrier to 9 ml of sample, and the isotopic composition of the sample was then calculated from the measured composition of the mixture and the known in-house standard via mass conservation.
Concentrations and isotopes of N 2 O were measured using a modified GC-IRMS with large volume purge and trap system 62 . Briefly, two needles were used for sample transfer and He pressurization, and the sample was transferred into a sparging flask (Pyrex) using ultra-high-purity He (>99.999%) and purged with He. For a 250 ml bottle, the sample was purged for 60 min at a flow rate of 50 ml min −1 , and for a 120 ml bottle, the purge time was 30 min. The extracted gases were passed through an ethanol trap with dry ice and a chemical trap filled with magnesium perchlorate and Ascarite to remove H 2 O and CO 2 . N 2 O was trapped by liquid nitrogen twice for purification and concentration and then injected into the GC-IRMS with He as carrier gas. N 2 O concentrations were determined by ion peak area (m/z = 44), and calibration of N 2 O concentration was calculated from ion peak areas (m/z = 44) with standard gases of 199.6 and 501.0 ppmv N 2 O/He, which were run at ten-sample intervals. The serum bottle was weighed before and after transfer to calculate the amount of water transferred. The precision of this method was estimated to be better than ±3% (ref. 62

Surface N 2 O saturation and air-sea flux
Surface N 2 O saturation was calculated using equation (1): where R (%) is the saturation of surface N 2 O; C obs represents N 2 O concentration at 5 m depth; C eq is the expected equilibrium concentration, which is computed using Henry's law 63 ; and the solubility depends on temperature and salinity 64 (2) and (3): where F (μmol m −2 d −1 ) is air-sea flux of N 2 O; k (cm h −1 ) is the gas transfer velocity depending on wind and water temperature; u is daily mean wind speed at 10 m above sea surface during the cruise, as measured by the on-board meteorological station; and Sc is the Schmidt number calculated from temperature 64 .

Estimation of the fraction of N 2 O source derived from the isotope minimum layer
A two-endmember mixing model of isotopically enriched N 2 O mixing upward from the N 2 O concentration maximum layer and isotopically depleted N 2 O produced at the isotope minima layer was used to calculate the fraction of N 2 O contributed by shallow in-situ production using equation (4) 10 : where δ total is the lowest measured isotopic value of N 2 O at the isotope minimum. δ deep is the isotopic signature of N 2 O mixing upward from deep layers; here, we use the N 2 O concentration maximum layer as an endmember; the measured δ 15 N was 9.52 ± 0.28‰ and the δ 18 O was 52.25 ± 0.74‰ in our study sites (Extended Data Fig. 10). δ shallow is the isotopic value of the in-situ source in the isotope minimum layer, which is unknown. f is the fraction of N 2 O contributed from the shallow source to the isotope minimum layer, with the remainder equal to that diffusing upward from the [N 2 O] maximum. The lower limit of f could be constrained by assuming δ shallow was represented by the lowest value in an existing database from the North Pacific, and the δ 15 N and δ 18 O was 1.8‰ and 24.5‰, respectively 65 .

Nitrification and N 2 O production rate calculation
Rates of nitrification were determined based on the accumulation of 15 N in the product pool relative to the initial 15 N signature of that pool. Rates were computed based on equation (5): where R NR is the bulk nitrification rate for all substrates following 15 NH 4 + enrichment (nmol N l −1 d −1 ). C NO x − is the product concentration at the beginning of incubation (nmol N l −1 ), f 15 is the atom% 15 N of the NH 4 + pool at the beginning of incubation (the fraction of 15 N-NH 4 + in the gross NH 4 + pool after tracer enrichment), and n t and n 0 are the atom% 15 N of the product pool (NO 2 − + NO 3 − ) at the ending and beginning of incubation (%), respectively. t is the duration of incubation (h). This equation quantifies the transformation rate including the concentration due to the tracer addition (that is, ambient substrate + tracer) and thus represents a potential reaction rate.
Rates of N 2 O production from a particular labelled substrate (for example, 15 (6) and (7): Nature Geoscience ) can be then derived using equations (8) and (9): (9) where R15 N−N 2 O (pmol N l −1 d −1 ) is the measured production of 15  ). The gross N 2 O production rate was derived from the sum of NH 4 + sourced, NO 2 − sourced and NO 3 − sourced N 2 O. Therefore, the rate of gross N 2 O production was calculated using equation (10): (10) where R gross is the total N 2 O production rate during our incubation (pmol N l −1 d −1 ). The errors of the NH 4 + sourced, NO 2 − sourced and NO 3 − sourced N 2 O rate are based on the increase of N 2 O of our incubation in the 2015 cruise (duplicates), 2018 cruise (triplicates) and 2019 cruise (triplicates), and propagation of the errors during the calculation using the equations listed above.

Detection limits of rate measurements
For nitrification rate measurements, the detection limits depend on the concentration of the product pool and the fraction of 15 N in the substrate pool during the incubation 66,67 . As mentioned, the precision of δ 15 N-NO x − was better than ±0.2‰, and we here use three times the standard deviation as a reliable enrichment of 15 N in the product pool. Therefore, we calculated a detection limit of 0.04-0.16 nmol N l −1 d −1 for nitrification. Similarly, for N 2 O production rate, the precision of δ 15 N-N 2 O and δ 18 O-N 2 O was better than ±0.3‰ and ±0.4‰, respectively, and we here use three times the standard deviation as a reliable enrichment of 45 N 2 O and 46 N 2 O in the product pool. Therefore, we calculated a detection limit of 0.1-0.3 pmol N l −1 d −1 for 45 N 2 O production rate and 0.2-0.6 pmol N l −1 d −1 for 46 N 2 O production rate in 15

Data availability
All data needed to evaluate the conclusions in the paper are deposited in the Zenodo database and can be accessed through https://doi. org/10.5281/zenodo.6867932.