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Distribution patterns of nitrous oxide during the summer season in the Prydz Bay, eastern Southern Ocean
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Jiexia Zhang1, Damian L. Arévalo-Martínez2, 3, 4, 5, Liyang Zhan1, *, Liqi Chen1, Jun Zhao6, Wangwang Ye1, Man Wu1, Yuhong Li1, Jian Liu1
Acta Oceanologica Sinica | 2025, 44(4) : 90 - 100
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Acta Oceanologica Sinica | 2025, 44(4): 90-100
• Multiscale processes in polar oceans and their climate and ecological implications: Part Ⅱ •
Distribution patterns of nitrous oxide during the summer season in the Prydz Bay, eastern Southern Ocean
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Jiexia Zhang1, Damian L. Arévalo-Martínez2, 3, 4, 5, Liyang Zhan1, *, Liqi Chen1, Jun Zhao6, Wangwang Ye1, Man Wu1, Yuhong Li1, Jian Liu1
Affiliations
  • 1Key Laboratory of Global Change and Marine Atmospheric Chemistry, Third Institute of Oceanography, Ministry of Natural Resources, Xiamen 361005, China
  • 2RD2 Marine Biogeochemistry, GEOMAR Helmholtz Centre for Ocean Research Kiel, Kiel 24148, Germany
  • 3Institute of Geosciences, Kiel University, Kiel 24148, Germany
  • 4Microbiology Department, Radboud University Nijmegen, Nijmegen 6525XZ, The Netherlands
  • 5Leibniz Institute for Baltic Sea Research Warnemünde, Rostock 18119, Germany
  • 6Key Laboratory of Marine Ecosystem Dynamics, Second Institute of Oceanography, Ministry of Natural Resources, Hangzhou 310012, China
Published: 2025-04-25 doi: 10.1007/s13131-024-2445-1
Outline
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The global ocean is a major source of the climate-relevant atmospheric trace gas nitrous oxide (N2O). However, an accurate assessment of the global oceanic emissions of N2O is hampered by missing data on dissolved N2O from large regions such as the Southern Ocean. To address this deficit, N2O was measured in the Prydz Bay in February 2015 during the 31st Chinese National Antarctic Research Expedition. N2O concentrations (saturation) in the surface layer were generally low (undersaturation with respect to atmospheric equilibrium) and ranged from 13.3 nmol/L to 16.1 nmol/L (83%–102%) at the time of sampling. A comparison of our observations with archived data revealed that no discernible trend in N2O concentrations in the surface waters of Prydz Bay could be detected for the period between 2006 and 2015. Temperature and salinity changes driven by meltwater input were the predominant controls on N2O concentrations in surface waters. At depth, the distribution of N2O concentrations was dominated by production via nitrification in offshore deep waters and vertical convection in the shelf waters, where concentrations were lower and gradients were less steep. Our results suggest a rather unusual pattern of N2O distribution in the Prydz Bay (low N2O in shelf waters compared with the open ocean), providing important insights into the coastal dynamics of N2O in high-latitude polar regions.

nitrous oxide (N2O)  /  greenhouse gases  /  sea-to-air flux  /  Prydz Bay  /  Southern Ocean
Jiexia Zhang, Damian L. Arévalo-Martínez, Liyang Zhan, Liqi Chen, Jun Zhao, Wangwang Ye, Man Wu, Yuhong Li, Jian Liu. Distribution patterns of nitrous oxide during the summer season in the Prydz Bay, eastern Southern Ocean[J]. Acta Oceanologica Sinica, 2025 , 44 (4) : 90 -100 . DOI: 10.1007/s13131-024-2445-1
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1　Introduction
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Nitrous oxide (N2O) is an important atmospheric greenhouse gas because of its strong radiative effect, which is 200−300 times greater than that of carbon dioxide (Canadell et al., 2021), and its role as a major ozone-depleting substance in the stratosphere (Ravishankara et al., 2009). Since the Industrial Revolution, the volume fraction of N2O concentration in the atmosphere has increased by about 24%, from 270 ppbv in 1750 to about 336 ppbv in 2023 (Dutton and Hall, 2023).
Marine production mainly occurs via two microbial pathways, namely, nitrification and denitrification. Under oxic conditions, N2O results as a byproduct of the first step of bacterial and archaeal nitrification (i.e., the oxidation of ammonia to nitrite), and the N2O yield during nitrification increases with decreasing oxygen concentration (Goreau et al., 1980; Löscher et al., 2012). During denitrification, N2O results as an intermediate product through nitrite reduction under suboxic conditions. Under oxygen concentrations below approximately 5 μmol/L, N2O is reduced to nitrogen gas during complete denitrification (Zhang et al., 2020).
Oceans are generally considered a net source of atmospheric N2O. Based on in situ observations and model estimates, surface ocean waters constitute 3.4 Tg/a (2.5–4.3 Tg/a), in terms of N, of the total N2O source (about 17.0 Tg/a, in terms of N) (Tian et al., 2020). The eastern tropical Pacific, equatorial and coastal upwelling zones are thought to be the largest oceanic N2O emission regions (Yang et al., 2020), whereas open ocean waters are mostly near atmospheric equilibrium. High-latitude (polar) studies are scarce, and consequently, uncertainties regarding N2O source-sink dynamics and marine emissions from the open ocean (Yang et al., 2020) and coastal areas (Resplandy et al., 2024) remain largely uncertain.
The Southern Ocean (SO) occupies about 20% of the global ocean area, yet the extent of N2O exchange across sea-ice-air interfaces is poorly understood. To date, only a few studies have published results on N2O measurements from the SO (Rees et al., 1997; Farías et al., 2015; Zhan et al., 2015; Grefe et al., 2018). Based on the evaluation of high-resolution air N2O measurements at the atmospheric monitoring station Cape Grim, Tasmania (Australia), Nevison et al. (2005) estimated that approximately 0.9 Tg/a, in terms of N, is ventilated to the atmosphere as N2O in the SO (30°–90°S). Hence, the SO may account for approximately 21% of global oceanic N2O emissions (Tian et al., 2020), which suggests that the region might be a potential source for the atmosphere. However, the region estimated based on the extrapolation of the dataset (Weiss et al., 1992) covered latitudes between 40°S and 60°S, and field studies have shown slight undersaturation of N2O (90%–94%) in the Antarctic coastal zone (Rees et al., 1997; Zhan et al., 2015), indicating that the SO could also act as a N2O sink. The Antarctic coastal zone is covered by sea ice all year round except in summer. As shown by recent studies, sea ice can act as a reservoir of organic matter (Henley et al., 2023; Koch et al., 2023; Thomas et al., 2001), which might fuel N2O production during organic matter degradation. The produced N2O could then be released into the atmosphere at the beginning of sea ice melting. In addition, sea ice-influenced coastal waters might become sinks (Rees et al., 1997; Zhan et al., 2015) due to the dilution of meltwater (Rees et al., 2022). In recent decades, global climate change has caused rapid changes in the SO, including warming (Sallee, 2018), westerly wind strengthening (Thomson et al., 2002) and altered temporal variability in sea ice formation-melting cycles (Kurtz and Markus, 2012). Ocean warming might change the activity of bacteria or archaea and hence change the N2O production rate. The increase in ocean temperature also accelerates the melting of sea ice, which could promote productivity by facilitating iron inputs from the Antarctic ice sheet (St-Laurent et al., 2017) and increasing sunlight exposure (Castellani et al., 2022), potentially affecting surface productivity, which, upon sinking, would fuel N2O production. However, unlike in the open ocean, the presence of melting sea ice in the Antarctic coastal zone enhances stratification, which could result in a dilution effect that reduces surface N2O concentrations while increasing storage at greater depths (Randall et al., 2012; Zhan et al., 2015). At present, it is unclear how the interactions between these processes affect the SO N2O budget.
To determine whether these ongoing changes affect N2O distributions and sea-air emissions and how they affect the eastern SO, we conducted a multiyear survey of N2O in the Prydz Bay (the third largest embayment in Antarctica). Here, we present measurements of dissolved N2O from this area in February/March 2015. These measurements were complemented with archived records of N2O measurements from January 2006 (Zhan and Chen, 2009) and 2011 (Zhan et al., 2015) to decipher the variability in dissolved N2O in the Prydz Bay under a changing SO. The possible influencing factors were discussed in combination with the environmental parameters. Our findings offer new insights into N2O dynamics in high-latitude marine environments.
2　Material and methods
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2.1　 Study site
Seawater samples for the determination of N2O concentrations were collected from four sections (P5, P6, PA, and IS) in the Prydz Bay (64°00'–70°00'S, 72°00'–76°00'E; 3−10 February 2015) as part of the 31st Chinese National Antarctic Research Expedition (CHINARE) aboard the research vessel Xuelong (Fig. 1a). The transect extended from the front of the Amery Ice Shelf to the continental shelf (with depth < 1000 m) and to the offshore deep waters (Fig. 1a). The data from the 31st CHINARE were complemented by the data from the 22nd and 27th CHINARE cruises (Fig. 1b), which took place in January 2006 (Zhan and Chen, 2009) and January 2011 (Zhan et al., 2015) (Table 1). Notably, only surface measurements are available for the 22nd CHINARE.
2.2　 Analysis and data handling
2.2.1　 N2O concentrations
Water samples were collected from Niskin bottles that were mounted on a Conductivity-Temperature-Depth (CTD) rosette. Seawater was transferred from the Niskin bottles to 250 mL BOD bottles via a silicon tube. After overflowing 2–3 times the bottle volume, 0.1 mL of saturated mercuric chloride (HgCl2) was added to the bottles to inhibit microbial activity. The bottles were immediately sealed with greased ground-glass stoppers and fastened with a clip to avoid air contamination. All the seawater samples had no bubbles and were stored at 4℃. In the laboratory, bubble-free, triplicate subsamples were taken from 250 mL BOD bottles into 20 mL headspace vials and then crimped with a metal crimp cap with PTFE-coated silicone rubber septa. Approximately 10 mL of high-purity nitrogen (N2, 99.999%) was used to create a headspace by replacing the seawater with N2. The N2O contents of the samples were analyzed via a static headspace method coupled with gas chromatography (GC-2010, SHIMADZU, Japan) equipped with an electron capture detector (ECD) (Zhan et al., 2013). Aqueous standard solutions were made of N2O standard gas-equilibrated water, and a headspace was created by following the same protocol as for the seawater samples. Six standard gas mixtures (volume fraction 100 ppbv, 200 ppbv, 300 ppbv, 500 ppbv, 700 ppbv, and 900 ppbv N2O in N2) were used. The standard gas mixtures were purchased from the National Institute of Metrology (NIM), P. R. China, and were calibrated against standard gas mixtures from the NOAA Earth System Research Laboratory. The NIM standards were within ±0.5% (or better) of the NOAA reference standards. The overall mean relative precision of the dissolved N2O concentrations was ±2%. The calculation of the N2O concentrations is described by Zhan et al. (2013).
2.2.2　 N2O saturation and excess
The saturation (in %) and excess N2O concentration (ΔN2O) were calculated via the following equations:
$ \mathrm{{N}_2O}_\mathrm{saturation}=[{\mathrm{N }}_2{\mathrm{O}}] _{\mathrm{measured}} /[{\mathrm{N}} _2{\mathrm{O}}] _{\rm{eq}} \times 100{\text{%}}, $
$ \Delta[\mathrm{N}_2{\mathrm{O}}]=[\mathrm{N_2O]}_{\mathrm{measured}}-[\mathrm{N_2O]}_{\mathrm{eq}}, $
where [N2O]eq is the equilibrium concentration of N2O at the in situ seawater temperature and salinity at the time of sampling, computed with the solubility equation by Weiss and Price (1980) and using a mean monthly atmospheric N2O dry mole fraction of 326.81 nmol/mol measured at mole fraction the Syowa Station (Antarctica; https://www.esrl.noaa.gov/gmd/hats/combined/N2O.html).
2.2.3　 Data comparability
We calibrated N2O concentrations during different cruises by comparing the values in deep waters. We assumed that the N2O concentrations in deep waters (>3000 m) of Prydz Bay were stable within the relatively short period considered here (Bange and Andreae, 1999; Toyoda et al., 2019). The N2O concentrations in deep ocean waters were compared to correct the measurement error between different cruises. As a result, we found that the N2O concentrations in deep waters were slightly higher during the 31st CHINARE than during the 27th CHINARE. Consequently, a correction factor of 0.93 was applied. The correction factor resulted from the ratio of the mean N2O concentrations in deep waters between the 27th CHINARE and 31st CHINARE. All N2O concentrations from the 31st CHINARE were multiplied by this factor. The N2O concentration data from the 22nd CHINARE were not corrected because only surface water data were available. The analytical precision of this dataset was ±5%.
2.2.4　 Other parameters
Nutrients and dissolved oxygen (O2) were obtained from the Second Institute of Oceanography, Ministry of Natural Resources. The nutrients were measured via Ultraviolet-Visible spectrophotometry, with a precision of ±0.1 μmol/L for $ {\mathrm{NO}}_{3}^- $ according to the WOCE protocol (Sun et al., 2012). The O2 concentrations were determined via the automatic Winkler titration method, with a precision of ±1 μmol/L. Apparent oxygen utilization (AOU) was calculated via the following equation:
$ \mathrm{AOU} ={[{\mathrm{O}}} _2 ] _{ \mathrm{eq}} -{[{\mathrm{O}}} _2] _{\mathrm{meas}} ,$
where [O2]eq is the equilibrium concentration calculated with the in situ seawater temperature and salinity via the O2 solubility equation from Weiss (1970) and where [O2]meas is the measured in situ O2 concentration.
The temperature and salinity data were obtained from a calibrated SBE-911 conductivity-temperature-depth (CTD) system (Sea-Bird Electronics Inc., America). The CTD temperature and salinity precision were 0.001℃ and 0.002, respectively.
2.2.5　 Sea-air flux
Sea-to-air flux densities (Fsea-air in µmol/(m2·d) of N2O) were computed via surface layer (3–5 m depth) data by applying Eq. (4):
$ F_{\mathrm{sea}\text{-}\mathrm{air}}=k_{ }\times\Delta[\mathrm{N}_2\mathrm{O}], $
where the gas transfer velocity (k, in cm/h) was computed according to Wanninkhof (2014):
$k_w=0.251(u _{10} )^2( S c_{{{\mathrm{N}}_2{\mathrm{O}}}} /660)^{-0.5} ,$
where u10 represents the wind speed at a height of 10 m above sea level, which was obtained from the ship’s anemometer. $S c_\mathrm{{N}_2 O} $ is the Schmidt number for N2O computed via the empirical equation from Wanninkhof (2014):
$\begin{split}{S c} _{{{\rm{N}}_2{\rm{O}}} } =\ &2\;356.2-166.38\;{t} +6.395\;2\;t ^2-\\& 0.134\;22\;{t} ^3 +0.001\;150\;6\;{t} ^4,\end{split}$
where t is the temperature of seawater.
2.2.6　 N2O mass balance calculation in the surface mixed layer
The depth of the surface mixed layer (SML) was defined based on a potential density difference of 0.125 kg/m3 and a temperature difference of 0.5℃ within the surface water (<2 m) (de Boyer Montégut et al., 2004; Ye et al., 2023). N2O fluxes through the pycnocline into the SML due to turbulent diffusivity (Fdeep) were estimated as follows:
$ {F}_{{\mathrm{deep}}}={K}_{z}\frac{\partial \left[{{\mathrm{N}}}_{2}{\mathrm{O}}\right]}{\partial {{z}}} ,$
where Kz represents the diffusion coefficient at depth z (in m) and 10–4 m2/s was used as the value in the Southern Ocean (Mashayek et al., 2017).$ \dfrac{\partial \left[{{\mathrm{N}}}_{2}{\mathrm{O}}\right]}{\partial z} $ is the vertical N2O gradient. Considering that the surface water in the Prydz Bay is a closed cyclonic and weak gyre, the effect of upwelling due to Ekman offshore water transport was assumed to be of minor importance (Charpentier et al., 2010). We used a simple mass balance approach between the upward N2O fluxes through the pycnocline and the outward fluxes through the sea-air interface (Fig. 2). The balance is expressed as follows:
$ -{F}_{\mathrm{sea}{\text{-}}\mathrm{air}}+{F}_{\mathrm{deep}}+{P}_{i}+{F}_{x}=0, $
where $ {P}_{i} $ [μmol/(m2·d)] represents the integrated N2O production in the SML (Charpentier et al., 2010). $ {F}_{x} $ is the other N2O source/sink generated by physical processes such as meltwater intruding the SML. A positive value of $ {F}_{x} $ corresponds to the input of N2O from meltwater, a negative value of $ {F}_{\mathrm{s}\mathrm{e}\mathrm{a}{\text{-}}\mathrm{a}\mathrm{i}\mathrm{r}} $ corresponds to ocean uptake of atmospheric N2O, and a positive value of $ {F}_{\mathrm{d}\mathrm{e}\mathrm{e}\mathrm{p}} $ corresponds to the input of N2O from the subsurface to the SML.
3　Results
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3.1　 Oceanographic setting
Prydz Bay is the third largest bay of the SO and has a unique topographic setting. The shelf oceanic circulation is characterized by a closed cyclonic gyre centered in a deep channel, known as the Amery Depression (Smith et al., 1984; Vaz and Lennon, 1996) (Fig. 1b). A relatively narrow and cold coastal current (Antarctic Coastal Current, ACoC) runs from the West Ice Shelf and enters the southeastern part of the bay. When the ACoC reaches the shallow region and near the Amery Ice Shelf, two components affect the flow, resulting in the formation of the typical gyre circulation in the area: one component of the current toward the west and the other eastward back to the inner region of the shelf (Smith et al., 1984; Sabu et al., 2021). The flow north of the continental shelf is associated with the Antarctic Circumpolar Current (ACC), which is driven eastward by the prevailing westerlies (Fig. 1b; Smith et al., 1984).
The continental shelf and offshore deep-water regions in the Prydz Bay are separated by a frontal structure, the Antarctic slope front (ASF), which extends from the shelf break (about 66.8°S) and ascends northward to subsurface layers (Williams et al., 2010). The major water masses identified in the Prydz Bay during austral summer are Antarctic surface water (AASW), continental shelf water (SW), Circumpolar Deep Water (CDW), and Antarctic Bottom Water (AABW). AASW has variable characteristics in terms of temperature and salinity (T > –1.80℃, S < 34.45) (Wong et al., 1985), with a neutral density (γn) < 28.00 kg/m3 (Shadwick et al., 2014), and is located in the upper layer from about 20 m within the Prydz Bay to approximately 80 m in the oceanic domain (Wong et al., 1985). Below the AASW lies a temperature minimum layer (−1.9℃ < T < –1.5℃, 34.2 < S < 34.5) (Wong et al., 1985), which is a layer of remnant “memory” of surface water formed in winter and called winter water (WW). To distinguish the WW in the continental shelf and offshore oceanic domain separated by the ASF (Smith et al., 1984; Middleton and Humphries, 1989), the continental shelf water, which is called SW, occupies most of the water column due to deep overturning derived from cool air temperature, strong winds and salt rejection during the formation of sea ice in winter. In the oceanic domain, it is termed WW and reaches a depth that gradually decreases northward from 300 m at about 67°S to disappear at about 64°S (Anilkumar et al., 2006). Below WW is CDW, which is warm, saline (0℃ < T < 2℃, 34.67 < S < 34.75) (Wong et al., 1985), and depleted of oxygen with a density of 28.00 kg/m3 < γn < 28.27 kg/m3 (Shadwick et al., 2014). The southward flowing warming CDW mixes with ambient water masses in the shelf region and forms modified Circumpolar Deep Water (mCDW) (Lin et al., 2016). Below the CDW is the AABW, which has a temperature of about 0℃ and a salinity of about 34.67 (Fig. 2).
3.2　 N2O in surface waters
The surface (3−5 m water depth) concentrations of N2O in the Prydz Bay (during the 31st CHINARE) were generally low, ranging from 13.3 nmol/L to 16.1 nmol/L (mean ± std: 14.7 nmol/L ± 0.7 nmol/L). Spatially, the lowest N2O concentrations were found in front of the Amery Ice Shelf (14.0 nmol/L ± 0.5 nmol/L), whereas the highest values were observed at about 67°–66°S (Fig. 3), where these values coincided with relatively low temperatures and salinities (Table 1). Most waters at the time of sampling were either near atmospheric equilibrium or slightly undersaturated (83%–102%; mean ± std: 94% ± 4%) (Table 1). Surface N2O undersaturation in the Prydz Bay has been reported to be associated with physical environment variabilities, such as the input of ice-melt water and high solubility due to low temperatures (Zhan and Chen, 2009; Zhan et al., 2015). Compared with cruises, we found that the N2O concentration was slightly higher in 2011 than in 2015 and 2006. However, the classical onshore-offshore gradient observed in most oceanic regions (Yang et al., 2020) was not attained, and no discernible temporal trend in N2O concentrations for the period 2006−2015 could be distinguished (Fig. 3).
3.3　 Cross-shelf variability in N2O in the Prydz Bay
The section along 73°E from 64°S to 68°S from the 31st CHINARE was taken as a representative profile section, which covers the continental shelf and slopes to the adjacent deep ocean. The vertical profiles of N2O were characteristically different in the Prydz Bay between the continental shelf and offshore deep ocean (Fig. 4a), which was consistent with the findings of a previous study (Zhan et al., 2015).
In the central Amery depression, the N2O concentrations below the surface water were relatively homogeneous, ranging from 14.3 nmol/L to 17.5 nmol/L (mean ± SD: 16.5 nmol/L ± 0.5 nmol/L) throughout the water column (Fig. 4a). In contrast, N2O distributions north of the ASF (i.e., in open ocean waters north of 67°S) clearly varied with increasing water depth (Fig. 4a). Between the ASF and about 65°S, N2O concentrations below the surface water slowly increased and then reached a weak N2O maximum value (about 18.8 nmol/L) between 100 m to 500 m. However, at the northernmost station P05-1 (65°S), a steep N2O concentration gradient and N2O maximum (up to 23.4 nmol/L) were observed at a depth of about 200 m below the N2O maximum layer, and N2O concentrations gradually decreased to about 18.2 nmol/L at about 2000 m. Below 2000 m, N2O concentrations were relatively constant, with a mean concentration of about 17.5 nmol/L.
4　Discussion
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4.1　 Driving factors of N2O in surface waters
The variability of N2O in surface waters reflects the balance between biological production/consumption and physically driven changes in solubility and/or water mass sources. Surface N2O production by biological processes is thought to be strongly limited by light (Horrigan et al., 1981) and dissolved oxygen inhibition (Averill and Tiedje, 1982). However, recent studies have demonstrated that significant rates of nitrification can be measured in the euphotic zone (Yool et al., 2007; Beman et al., 2012), suggesting that N2O production in the surface layer might be nonnegligible (Zamora and Oschlies, 2014). However, N2O production through nitrification generally requires low O2 concentrations (about 4–10 μmol/L), high ammonium (${\mathrm{NH}}_4^+ $) (Zamora and Oschlies, 2014; Mdutyana et al., 2020) and abundant amoA genes (marker genes for nitrification by archaea) in surface water. In our study, the high O2 concentrations (mean ± std: 376 μmol/L ± 42 μmol/L) measured seem to indicate inadequate conditions for nitrifiers. This is in line with the observation that rates of nitrification are generally extremely low in southern regions of the SO (Mdutyana et al., 2020, 2022). Moreover, amoA genes have been found to be essentially absent in the surface water of the SO (Kalanetra et al., 2009). Thus, the contribution of surface N2O by nitrification in the Prydz Bay seems negligible.
To explore the role of physical processes as controlling factors of N2O distribution in the Prydz Bay, we carried out multiple regression analyses of N2O concentration vs. temperature and salinity (Fig. 5). R² indicates how well the regression line approximates the real data points. Our regression analysis revealed that N2O was significantly negatively correlated with temperature during all three cruises (Pearson, p < 0.005), indicating that temperature affected mainly the changes in N2O concentrations in surface waters. However, the effect of salinity on surface water is also important because the level of salinity is determined by the volume of meltwater in the sea ice region, which directly decreases the N2O concentration in surface water (Randall et al., 2012).
We used a simple box model to calculate the N2O mass balance in the SML [Eq. (8)]. The results revealed that the cross-pycnocline fluxes from the subsurface to the SML (Fdeep) (mean ± std: 0.04 m/s ± 0.03 m/s in 2015, Table 2) were almost zero, indicating that the upward transport of N2O from subsurface water was of minor importance; this results from strong stratification of the upper water column, reducing mixing between the surface and N2O-enriched subsurface waters. In this study, we had no data to calculate the contribution of integrated N2O production in the SML; however, we assumed that the value of Pi is close to zero because N2O production in the SML could be minor, as discussed above. In terms of the N2O mass balance, Fsea-air was affected mainly by melt water. Fsea-air was mostly negative [range: –3.2 μmol/(m2·d) to 1.4 μmol/(m2·d); mean ± SD: –0.6 μmol/(m2·d) ± 1.0 μmol/(m2·d)] during the austral summer of 2015. Spatially, the lowest sea-air flux densities were observed in front of the Amery Ice Shelf. Similarly, Fx was clearly lower in front of the Amery Ice Shelf (mean ± std: –2.0 m/s ± 1.1 m/s in 2015) than in the northern region (mean ± std: –0.2 m/s ± 0.6 m/s in 2015), suggesting a greater input of meltwater, which caused dilution of N2O in front of the Amery Ice Shelf (Randall et al., 2012; Zhan et al., 2015). These results were in line with the values of the 27th CHINARE. Thus, overall, Prydz Bay represented a weak sink of atmospheric N2O at the time of sampling (austral summers in 2006 and 2015). During the summer, sea ice melting plays an important role in the N2O cycle in the surface water of Prydz Bay as strong stratification, caused by sea ice melting water, hinders the vertical exchange between the subsurface and surface waters, which combined with low nitrification rates results in low N2O production in surface waters.
4.2　 Driving factors of N2O in subsurface waters
The different distributions of N2O in the Prydz Bay indicated that it was influenced by different driving factors. In general, three processes influence N2O concentrations in the subsurface and deep layers of the ocean: in situ production, consumption and physical transport. The prevailing high O2 concentrations (158−424 µmol/L) in the Prydz Bay exclude N2O production or consumption via denitrification but favor nitrification as the main production process of N2O in subsurface/deep waters.
The ASF and depression separate Prydz Bay into two relatively independent regions, the continental shelf and offshore deep ocean (Vaz and Lennon, 1996), with only limited exchange between them (Smith et al., 1984). Therefore, the effect of the exchange of N2O between the continental shelf and offshore deep ocean can be considered negligible. We assumed that vertical convection and in situ production were possibly responsible for the observed N2O concentration and distribution pattern on the continental shelf of Prydz Bay.
The continental shelf is a relatively high net primary productivity region (Wang et al., 2021). Considering that the shelf water column environment was well ventilated dissolved oxygen (338 μmol/L ± 29 μmol/L) (Fig. 4a), a denitrification source seems unlikely. This is further supported by the positive N* values (understood here as deviations in the ratio of dissolved nitrogen to phosphorus from Redfield stoichiometry, i.e., N* = [${\mathrm{NO}}_3^- $] + [${\mathrm{NO}}_2^- $] + [${\mathrm{NH}}_4^+ $] – 16 × [${\mathrm{PO}}_4^{3-} $] + 2.9) (Deutsch et al., 2001) and found throughout the study area (mean ± std: 2.5 μmol/L ± 4.8 μmol/L), indicating that canonical denitrification was unlikely. Linear regression between Δ[N2O] and AOU did not reveal the expected relationship for a system where nitrification is dominant (Fig. 6a), which may result from strong vertical convection caused by increased density through the cooling of air temperature and salt rejection during sea ice formation in fall and winter, which can reach a very deep depth throughout the bay (Smith et al., 1984). We assume that the biological production of N2O was low because low temperature inhibits enzyme activity (Sabba et al., 2015) and had a low rate of nitrification (Mdutyana et al., 2020, 2022).
Unlike those in most coastal regions, the N2O concentrations and Δ[N2O] values in the offshore deep ocean were higher than those in the continental shelf of Prydz Bay. A positive relationship between AOU and N2O was observed in the offshore deep ocean (Fig. 6a). However, the profiles of N2O clearly differed between 65°–67°S and north of 65°S. The N2O profiles north of 65°S showed clear maxima between 150 m and 200 m, which coincided with the dissolved oxygen minimum (Fig. 4b), indicating that they were associated with the core of the upper Circumpolar Deep Water (UCDW). A similar pattern was not observed between 65° and 67°S, where the N2O concentration was relatively low and presented N2O subsurface maxima at about 150–300 m and then a very small decrease until 2000 m. These profiles agree with those of a previous study published by Zhan et al. (2015). We assumed that hydrographic processes are likely to result in these two types of profiles. Between 65°S and 67°S, subsurface water was primarily occupied by lower Circumpolar Deep Water (LCDW) (200–2000 m), which originated from North Atlantic Deep Water. In contrast, the water masses north of 65°S were occupied by UCDW (100–300 m) and LCDW (300–2000 m). UCDW originates primarily from the eastern South Pacific Ocean and western Indian Ocean.
We assumed that N2O in the eastern South Pacific Ocean/western Indian Ocean and North Atlantic Deep Water was taken into the UCDW and LCDW, respectively, and that an accumulation of N2O occurred during transport over time (Bange and Andreae, 1999). N2O at the starting point of the South Pacific Ocean/western Indian Ocean tends to be higher than that in North Atlantic Deep Water (Bange and Andreae, 1999), resulting in higher N2O concentrations in the UCDW than in the LCDW. A positive linear correlation between Δ[N2O] and AOU was observed in the offshore deep ocean (Fig. 6a), indicating that the observed N2O accumulation over time was caused mainly by nitrification. There were no discernible differences in N2O concentrations between 27th CHINARE (2011) and 31st CHINARE (2016) in the offshore deep ocean because the subsurface water in the SO is influenced mainly by large-scale water masses in the global circulation system, and differences cannot be observed in such a short time. However, under the influence of global climate warming and increasing sea water melting, the global circulation system might still have an important impact on the physical control of N2O cycling in the SO.
5　Conclusions
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The dissolved N2O concentrations in the surface waters of Prydz Bay were generally low and featured undersaturated with the overlying atmosphere, such that this area was a weak sink of atmospheric N2O at the time of sampling. A comparison with N2O concentrations from the same region during different cruises revealed that N2O concentrations and saturations in the surface waters were variable, but no temporal trend for the period 2006−2015 could be identified. Temperature and salinity changes played a dominant role in setting the N2O concentrations for surface ocean waters in the Prydz Bay. Excess N2O was observed in the offshore deep ocean below the surface layer, mostly due to N2O accumulation from nitrification during transportation. No signs of significant N2O production were detected in the shelf waters between 2011 and 2015.
Unlike most coastal regions, our results suggest a different pattern of N2O distribution in the Prydz Bay (low N2O on the continental shelf vs. high N2O in offshore deep waters), a finding that might serve future model studies assessing N2O emissions from the SO. To detect the interannual variability and long-term trends in dissolved N2O concentrations in the Prydz Bay (which are in turn induced by ongoing environmental changes in the SO, such as warming and increasing atmospheric N2O mole fractions), we suggest establishing time series measurements regularly as part of the proposed global N2O ocean observation network.
Funding
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  • The National Natural Science Foundation of China under contract No. 41906193; the Scientific Research Foundation of Third Institute of Oceanography, Ministry of Natural Resources, under contract No. 2019033; the Natural Science Foundation of Fujian Province under contract No. 2019J05147; the Federal Ministry of Education and Research of Germany under contract No. FKZ 03F03F0783A; the National Polar Special Program under contract Nos IRASCC 01-01-02 and IRASCC 02-02.
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Year 2025 volume 44 Issue 4
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doi: 10.1007/s13131-024-2445-1
  • Receive Date：2024-10-14
  • Online Date：2026-11-05
  • Published：2025-04-25
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  • Received：2024-10-14
  • Accepted：2025-01-21
Funding
The National Natural Science Foundation of China under contract No. 41906193; the Scientific Research Foundation of Third Institute of Oceanography, Ministry of Natural Resources, under contract No. 2019033; the Natural Science Foundation of Fujian Province under contract No. 2019J05147; the Federal Ministry of Education and Research of Germany under contract No. FKZ 03F03F0783A; the National Polar Special Program under contract Nos IRASCC 01-01-02 and IRASCC 02-02.
Affiliations
    1Key Laboratory of Global Change and Marine Atmospheric Chemistry, Third Institute of Oceanography, Ministry of Natural Resources, Xiamen 361005, China
    2RD2 Marine Biogeochemistry, GEOMAR Helmholtz Centre for Ocean Research Kiel, Kiel 24148, Germany
    3Institute of Geosciences, Kiel University, Kiel 24148, Germany
    4Microbiology Department, Radboud University Nijmegen, Nijmegen 6525XZ, The Netherlands
    5Leibniz Institute for Baltic Sea Research Warnemünde, Rostock 18119, Germany
    6Key Laboratory of Marine Ecosystem Dynamics, Second Institute of Oceanography, Ministry of Natural Resources, Hangzhou 310012, China

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表12种不同金属材料的力学参数
科
Family		 属数
Number of
genus		 种数
Number of
species		 占总种数比例
Percentage of
total species (%)		 属
Genus		 种数
Number of
species		 占总种数比例
Percentage of total
species (%)
鹅膏菌科Amanitaceae 2 11 5.26 鹅膏菌属 Amanita 10 4.78
小菇科 Mycenaceae 2 12 5.74 丝盖伞属 Inocybe 5 2.39
多孔菌科 Polyporaceae 8 14 6.70 蜡蘑属 Laccaria 5 2.39
红菇科 Russulaceae 3 23 11.00 小皮伞属 Marasmius 6 2.87
小菇属 Mycena 11 5.26
光柄菇属 Pluteus 5 2.39
红菇属 Russula 17 8.13
栓菌属 Trametes 5 2.39
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