New constraints on biological production and mixing processes in the South China Sea from triple isotope composition of dissolved oxygen

. The South China Sea (SCS) is the world’s largest marginal sea, playing an important role in the regional biogeochemical cycling of carbon and oxygen. However, its overall metabolic balance, primary production rates and links to East Asian Monsoon forcing remain poorly con-strained. Here, we report seasonal variations in triple oxygen isotope composition ( 17 (cid:49) ) of dissolved O 2 , a tracer for biological O 2 , gross primary production (GP; inferred from δ 17 O and δ 18 O values) and net community production (NP; evaluated from oxygen–argon ratios) from the SouthEast Asian Time-series Study (SEATS) in the SCS. Our results suggest rather stable mixed-layer mean GP rates of ∼ 1500 ± 350 mg C m − 2 d − 1 and mean NP of ∼ − 13 ± 20 mg C m − 2 d − 1 during the summer southwest monsoon season. These values indicate, within uncertainties and


Introduction
The South China Sea (SCS) is the largest marginal sea of the world and significantly influences the regional biogeochemistry and climate (Wong et al., 2007a). The SCS also contributes to global circulation. A pathway through the SCS connects the tropical Pacific with the Indian Ocean, impacting the Indonesian Throughflow, a current which plays a pivotal role in the coupled ocean and climate system (Qu et al., 2005). It has been suggested that marginal seas may act as a significant global atmospheric carbon dioxide (CO 2 ) sink, primarily due to CO 2 absorption by continental shelf waters (Tsunogai et al., 1999;Liu et al., 2000;Yool and Fasham, 2001;Chen et al., 2003;Thomas et al., 2004). However, the heterogenous nature together with latitudinal differences between ocean margins makes joint extrapolations to a global scale highly uncertain. Most observations have revealed that seas at mid-latitude shelves, which experience strong spring blooms and clear seasonal patterns, function particularly well as CO 2 sinks (e.g. the North Sea, Frankignoulle and Borges, 2001; the Gulf of Biscay, Thomas et al., 2004; the Celtic Sea, Seguro et al., 2019; the East China Sea, Tsunogai et al., 1999;Wang et al., 2000; or the Mid-Atlantic Bight, De-Grandpre et al., 2002). Conversely, tropical and subtropical shelves and marginal seas are most likely CO 2 sources (Cai and Dai, 2004), and a similar scenario may also be anticipated for the SCS. While several studies reaffirmed that the SCS indeed acts as a source of CO 2 to the atmosphere in the spring, summer and autumn (Rehder and Suess, 2001;Zhai et al., 2005), Tseng et al. (2005) reported on the uptake of CO 2 during winter from the SCS. The observed CO 2 invasion, driven by an unusual seasonal pattern in the oligotrophic open northern part of the SCS with elevated chlorophyll concentrations (0.30-0.35 mg m −3 ) and primary production (300 mg C m −2 d −1 ), was apparently large enough to compensate for the CO 2 evasion during the rest of the year, resulting in only minor net annual sea-atmosphere CO 2 fluxes (0.24 g C m −2 yr −1 ; Tseng et al., 2007). Although the primary production and phytoplankton biomass are low for most of the year in the SCS, it appears that a clear winter maximum can be found regularly -a distinct seasonal pattern from other low-latitude water bodies. Clearly, the role of the SCS (and marginal seas in general) is complex, and their seasonal carbon cycling demands further study.
Owing to its geographical position between the Tibetan Plateau and the western Pacific warm pool, the SCS is continuously exposed to the East Asian Monsoon, which plays a fundamental role in its oceanography and biogeochemistry (see Wong et al., 2007a, for an overview). From June to September, the weaker southwest summer monsoon (SWM) drives the anti-cyclonic circulation gyre, while the strong northeast winter monsoon (NEM) propels a basin-wide cyclonic circulation gyre between November and April (Fig. 1). This intense seasonal reversal drives the short-and longterm physical, chemical and biological processes that control the distribution of phytoplankton communities (Ning et al., 2004). As a result, intermediate chlorophyll a (Chl a) concentrations are typically associated with the SWM, while the highest phytoplankton biomass is expected during the NEM due to increased diapycnal nutrient supply from the thermocline. Winter mean Chl a concentrations often peak at about 0.50 mg m −3 in the subsurface chlorophyll maximum and at 0.20 mg m −3 at the surface (Liu et al., 2002). Upwelling produced by the convergence of currents in the cyclonic gyre near Luzon Strait, where the Kuroshio intrudes, can at times enhance the mean Chl a concentration (to about 0.65 mg m −3 ) and primary production in winter to about 8 times the summer values (Chen et al., 2006). Conversely, lowest Chl a values have been observed during intermonsoon seasons (Liu et al., 2002;Wong et al., 2007b;Li et al., 2017). Superimposed on the main seasonal monsoondriven pattern, episodic events such as typhoons may temporarily elevate primary production due to wind-enhanced vertical mixing, bringing nutrients from the nutricline to the mixed layer and stimulating production. For instance, Lin et al. (2003) reported a bloom patch with average surface Chl a concentrations of 3.20 ± 4.40 mg m −3 during the passing of tropical cyclone Kai-Tak in July 2000. In addition to typhoons, the SCS has been suggested to be sensitive to various types of short-term physical forcings including tides, internal waves, eddies or topography-flow interactions (Lin et al., 2010;Tai et al., 2020;Lai et al., 2021). Generally, these tend to enhance vertical mixing, supplying nutrient-rich waters to the mixing layer, which enhance phytoplankton production and Chl a concentration (to about 0.30-0.40 mg m −3 ) in the oligotrophic waters of the SCS. Interannual variability in the SCS is primarily driven by the ENSO (El Niño-Southern Oscillation), which modulates the strength of the monsoon forcing, which in turn affects the regional marine biogeochemistry. During warm (cold) El Niño (La Niña) episodes in the Pacific, the monsoon tends to have a late (early) onset, and the monsoon intensity is generally weaker (stronger; Zhou and Chan, 2007). Consequently, weakened wind mixing and strengthened water column stratification results in anomalously low Chl a concentrations in the northern SCS. For example, the 1997-1998 El Niño event was one of the most powerful ENSO events in recorded history and caused concentration of surface Chl a to drop from 0.20 to 0.10 mg m −3 in the northern SCS and the mean winter production to be reduced by about 40 % (Shang et al., 2005;Tseng et al., 2009).
Accurate quantification of phytoplankton production rates, a fundamental property of the ocean system, remains a challenge primarily due to methodological biases. This has been a subject of increasing debate over the past years, resulting in augmented efforts to compare and resolve production rates through different methods (e.g. Juranek and Quay, 2013;Regaudie-de-Gioux et al., 2014). In the SCS and at the SouthEast Asian Time-series Study (SEATS;Wong et al., 2007a), our understanding of primary production is predominantly limited to opportunistic assessments using the 14 C-assimilation method (Liu et al., 2002) or satellitebased SeaWiFS observations (Liu et al., 2002;Lin et al., 2003). While the traditional 14 C approach (Steeman-Nielsen, 1952) is limited due to its in vitro nature, which cannot reflect the time-averaged mixed-layer phytoplankton production (Marra, 2002), the latter relies on calibrations against field measurements that are spatially and temporally scarce (Carr et al., 2006). Although not exempt of uncertainties (Juranek and Quay, 2013), as these are inherent to any productivity determination, the triple oxygen isotope composition ( 17 ) technique (Luz et al., 1999;Luz and Barkan, 2000) combined with O 2 /Ar measurements has proved to be a powerful tool to provide a new perspective on evaluating primary production (e.g. Sarma et al., 2005;Reuer et al., 2007;Stanley et al., 2010;Hamme et al., 2012;Castro-Morales et al., 2013;Jurikova et al., 2016). The key advantage of this technique is that 17 allows for distinguishing photosynthetic O 2 input from other sources directly in situ, while the co-variation of δ 17 O and δ 18 O, the "dual-delta approach" (Prokopenko et al., 2011;Kaiser, 2011), enables an estimation of the integrated gross productivity in the mixed layer.
In order to evaluate the photosynthetic O 2 production and its contribution to the local carbon balance, as well as improve our understanding of seasonal variabilities in primary production in the SCS, we performed triple isotopic analyses and determined the O 2 /Ar ratio of dissolved O 2 samples from five vertical profiles during the occupation of SEATS in October 2013, August 2014 and April 2015. We combined the 17 and the O 2 /Ar tracer data to study the seasonal trends in photosynthetic vs. atmospheric O 2 input in the upper water column (∼ 200 m), which we relate to the main monsoon seasons. Gross and net primary production rates are also estimated and discussed. Finally, owing to the limited contribution from photosynthesis and air-sea gas exchange to the 17 signal in a parcel of deep water (200 to 3500 m), the potential for the application of the tracer for assessing mixing processes is discussed.
The accuracy of dissolved oxygen concentration measurements from the CTD was verified and calibrated against in vitro measurements. Briefly, water samples were siphoned into triplicate 60 mL bottles (Wheaton), and the Winkler titration method by Pai et al. (1993) was adopted for in vitro dissolved O 2 determination with a precision of 0.2 % r.s.d. (relative standard deviation; full scale). Concentrations of Chl a were measured by a fluorometer (Chelsea Aqua-Tracka III) attached to the CTD to monitor vertical profiles of fluorescence and calibrated by in situ Chl a measurements with a Turner Designs fluorometer (10-AU-005) after extraction with 90 % acetone using a non-acidification method (Gong et al., 1996). The precision of Chl a measurements using a Turner fluorometer is usually better than 8 % for any chlorophyll values exceeding 0.50 mg m −3 (Strickland and Parsons, 1972), with any uncertainties in the estimation of Chl a linked to the presence of Chl b being below 5 % (Lorenzen, 1981). We compared two mixed-layer depth definitions: (1) temperature-based definition defined by 1 • C ( T ) threshold from reference temperature value at 10 m depth, and (2) dissolved-O 2 -based definition defined by 1 % ( O 2 ) threshold from reference O 2 concentration at 10 m depth. Selected mixed-layer depths were further verified by careful visual inspection of vertical temperature, density and dissolved oxygen profiles. Although the preferred dissolved-O 2 -based mixed-layer definition considers a relative difference of 0.5 % O 2 concentration to a reference value at 10 m (Castro-Morales and Kaiser, 2012), we did not find this definition suitable for our study location, as the oscillations in O 2 concentration in the mixed layer alone were of the order of 0.5 %. Instead, we opted for the 1 % definition, which also closely agreed with the visual inspection of the profiles. The limit of the photic zone was defined as the depth where the photosynthetically active radiation (PAR) was 1 % of the surface value. We used Ocean Data View (ODV; Schlitzer, 2020) for profile visualization.
Triple oxygen isotope analyses were carried out at Academia Sinica, Taiwan. The triple oxygen isotope composition or 17 O-excess (Luz et al., 1999;Luz and Barkan, 2000) is defined as where the isotopic compositions δ 17 O and δ 18 O represent the deviation of the abundance ratio of an isotopic and normal species in a sample relative to that of a standard:  Luz and Barkan (2005) the factor λ is taken to be 0.518. As suggested by Luz and Barkan (2011), we note that a slope of λ = 0.516 might present a more appropriate choice. However, in order to enable a direct comparison to other studies, we preferred the earlier value, which has been largely applied in studies on marine production to date. The use of a slope λ = 0.516 would result in only a minor increase in 17 (about 1 %-8 % of the reported values), which for most of our samples remains within analytical uncertainties. Our laboratory protocols for dissolved oxygen sample preparation and analysis are detailed in Jurikova et al. (2016). Note that O 2 -Ar data are not available for the profile from October 2013 due to the setting of the gas chromatograph condition at a dry-ice-acetone slush temperature for complete separation of O 2 . In summary, dissolved gases were extracted from water following Emerson et al. (1995) and Luz et al. (2002). δ 17 O and δ 18 O values of O 2 in the purified oxygen-argon mixture (or pure oxygen for October 2013 samples) were determined by dual-inlet mass spectrometry (Thermo Scientific Finnigan MAT 253 isotope ratio mass spectrometer). Similarly as in Jurikova et al. (2016), an Ar correction was performed to correct for the distribution of gases between the headspace and water in the sampling flasks and normalized to air. A size correction for the total amount of gas in the sample was not required at our current mass spectrometer setting and hence not applied. This was tested by the measurement of air samples and a working reference at varying gas concentrations, ranging approximately 20 to 60 µmol of O 2 -Ar mixture. The reduction of sample size by a factor of ∼ 2 did not affect the measured ratios, with the standard deviation between the air samples (n = 8) being ±0.36 ‰, ±0.070 ‰, ±1 per meg (1000 per meg = 1 ‰), and ±9 ‰ for δ 17 O, δ 18 O, 17 and δ(O 2 /Ar), respectively (see the Supplement). Potential fractionation of the working reference gas in the bellows during analysis (also known as the "bleed effect"; see also Mahata et al., 2012) was monitored by noting the change in oxygen isotope ratios throughout the analysis (in the range of 100 % to 20 % of the bellows capacity). Based on 1 year of analyses (more than 20 batches measured, with one batch consisting typically of 12 measurements against the same working reference aliquot in the bellows), we found the bleed effect to be small and corrected. The bleed effect on δ 18 O was ∼ 0.15 ‰-0.20 ‰ every 10 sample analyses (∼ 20 h of mass spectrometer time) and < 10 per meg for 17 . To minimize any potential bleed effects, only a limited number of sample analyses were performed for each reference gas refill. To monitor the bleed effect, two to three atmospheric air samples were analysed alongside dissolved O 2 samples in each batch (against the same working reference aliquot), with no significant fractionation observed. Overall, our actual and long-term precision (1σ , standard deviation) established from routine measurements (n = 36) of atmospheric air O 2 was 0.017 ‰, 0.030 ‰ and 6 per meg for δ 17 O, δ 18 O and 17 , respectively, and our O 2 scale (Jurikova et al., 2016;Liang and Mahata, 2015;Liang et al., 2017) is in agreement with that of Luz and Barkan (2011). The O 2 /Ar ratio was obtained by peak jumping following Barkan and Luz (2003), and it is expressed as δ(O 2 /Ar) (‰) = [(32/40) sample /(32/40) standard − 1] × 10 3 . The longterm precision (1σ ) of routine measurements of atmospheric air was better than 5 ‰. Results from the equilibrated water samples (n = 3, 1σ ) gave a mean 17 eq = 11 ± 3 per meg, δ 17 O eq = 0.323 ± 0.020 ‰, δ 18 O eq = 0.603 ± 0.037 ‰ and δ(O 2 /Ar) eq = −102 ± 4 ‰ (Jurikova et al., 2016).

Primary production calculations
To quantify gross production rates from δ 17 O and δ 18 O values, we followed the standard dual-delta approach following Prokopenko et al. (2011) and Kaiser (2011), where the gross oxygen production (GOP) may be calculated as follows: where δ * O is the measured value in a sample, δ * O eq is the relative abundance between air O 2 and dissolved O 2 in equilibrium with the atmosphere (i.e. the results from the equilibrated water samples, Sect. 2.1; Jurikova et al., 2016), and δ * O p represents the photosynthetic O 2 (Luz and Barkan, 2011). C o is the O 2 concentration at saturation using solubility coefficients from Benson and Krause (1984), and K is the piston velocity, which is the coefficient for gas exchange. K was calculated using the quadratic relationship appropriate for wind speeds between 3 and 15 m s −1 (K = 0.24 × U 2 10 × (Sc/660) −0.5 ) and normalized to Schmidt number 660 (Sc 660 ) based on mixed-layer temperatures (Wanninkhof et al., 2009). We compared two different approaches for deriving K. The resulting K values are available in Table 1. First, we used a simple approach where K was derived from mean NCEP (National Centers for Environmental Prediction) wind speeds ( Fig. 2) and averaged over the O 2 residence time in the mixed layer preceding sampling (K avg ; 12, 5.6 and 4.4 d for October 2013, August 2014 and April 2015, respectively), based on the average mixed-layer depth and the gas transfer coefficient. Second, we used the approach by Reuer et al. (2007), where K was calculated using a weighting technique, which considers variable wind speeds and accounts for the fraction of mixed layer ventilated each day (K wgh ). For this, K was derived from satellite wind speed measurements of hourly resolution using the ERA5 dataset (ECMWF, European Centre for Medium-Range Weather Forecasts, https://www.ecmwf.int/; the Supplement). Following Reuer et al. (2007), the fraction of the mixed layer ventilated on the collection date (f 1 ) was determined from the mixed-layer depth (Z MLD ) and gas transfer velocity on the collection day (K 1 ) as f 1 = K 1 × 1 d/Z MLD , and was assigned a weight ω 1 = 1. The fraction of the mixed layer ventilated prior to the sample collection day (day 2) was similarly calculated as f 2 = K 2 ×1 d/Z MLD but was assigned a reduced weight according to the fraction of the mixed layer ventilated on day 1 (ω 2 = ω 1 × (1 − f 1 )). Since the SEATS station was occupied for only a limited time during each cruise, we used the Z MLD of the collection date for all calculations. Considering the rather regular interannual pattern and minimal daily variations in the mixed-layer depth, the used mean value should be a suitable representation for the different sampling months (see also Sect. 3.1). The weight on the tth day prior to the sampling is described by the general term ω t = ω 1−1 × (1 − f 1 ). A weighted gas transfer velocity for each day was then calculated as K t ω t , and the weighted gas exchange rate for the mixed layer was calculated as where the term "(1 − ω 30 )" accounts for the residual unventilated portion of the mixed layer. We utilized 30 d for each sampling date (October 2013, August 2014 and April 2015) as the residual fraction on the 30th day was already minimal. The two approaches for deriving K resulted in different production rates, with the K avg either underestimating or overestimating production when compared to K wgh , by ∼ 38 % in October 2013, ∼ 47 % in August 2014 and ∼ 21 % in April 2015 (this applies to both gross and net production rates since the choice of K affects both proportionally). We therefore used K wgh for calculating production rates at SEATS. Mixed-layer O 2 production time (O 2 concentration in the mixed layer divided by O 2 gross production rate) was determined to evaluate the rate at which O 2 was produced biologically against the physical O 2 residence time. The O 2 production time was estimated from the measured O 2 concentrations and the GOP and was generally lower than the O 2 residence time (0.3 d for October 2013, 1.3 and 1.4 d for 5 and 6 August 2014, respectively, and 1.0 and 3.6 d for 24 and 25 April 2015, respectively).
To assess the net oxygen production rates (NOP), we used the O 2 /Ar measurements consistent with the biological O 2 supersaturation concept for net photosynthetic production. Because the physical properties of O 2 and Ar are similar and because Ar has no biological sources or sinks, measurements of Ar concentration in water may be used to remove physical contributions to O 2 supersaturation. The biological oxygen supersaturation (O 2 /Ar) (given in %) is defined as the relative deviation of the O 2 /Ar in a sample to the O 2 /Ar at equilibrium with the atmosphere (e.g. Craig and Hayward, 1987;Emerson et al., 1995;Kaiser et al., 2005) and may be calculated as follows: NOP can be calculated from (O 2 /Ar) values following Luz et al. (2002): The estimation of both GOP and NOP via the presented approaches relies on the assumption that the mixed layer is at a steady state and that there is no significant entrainment or upwelling of low-O 2 subsurface water into the mixed layer nor lateral advection from adjacent waters. Production rates were converted from O 2 to C units following a commonly applied approach (e.g. Hendricks et al., 2004;Juranek et al., 2012). To scale GOP to gross C production, we accounted for the fraction of O 2 linked to Mehler reaction and photorespiration following Laws et al. (2000) by applying a photosynthetic quotient (PQ) of 1.2. For NOP conversion, we used a PQ of 1.4 for new production (Laws, 1991). Hereinafter, we refer to the scaled production rates in C units as GP and NP.

Oceanographic setting
Vertical distributions of physical parameters, chlorophyll and dissolved O 2 composition were measured from profiles collected at SEATS during October 2013, August 2014 and April 2015. Generally, during the sampling campaigns the temperature-based mixed layer (see Sect. 2.1 for definition) variations were only minor and varied depending on the month. Highest surface temperatures of 29 • C were recorded during the summer in August 2014. In October 2013, the surface temperature was 28 • C. Lowest surface temperature of 27 • C was observed in April 2015. Temperature-based mixed-layer depth limit was deepest at 49 m on the 16 October 2013. In August 2014, the mixed layer was relatively shallow but changed from 25 m on the 5 August 2014 to 34 m on the 6 August 2014. The mixed-layer depth limit was 28 and 26 m on the 24 and 25 April 2015, respectively ( Table 1) (Table 1). The mixed-layer depths determined by the different criteria were in close agreement within 2 m, except for the one on the 5 August 2014 when the difference between the definitions was 5 m. As the O 2 -based definition was more conservative and is directly related to the species of interest of this study, we used the O 2 -based mixed-layer depths for estimating integrated mixed-layer production from the isotopic composition of dissolved O 2 .
The observed mixed-layer depths and interannual pattern fit well within the trend expected for SEATS, which appears to stay rather regular between years (Wong et al., 2007a;Tai et al., 2017). A shallow mixed layer (about 20 m deep in sum- Table 1. Mixed-layer dissolved oxygen composition and estimated seasonal primary production rates at the SouthEast Asian Time-series Study (station 55, SEATS) in the South China Sea (SCS) for October 2013, August 2014 and April 2015. NP and GP were calculated using the O 2 -based mixed-layer depth and weighted gas exchange rate (K wgh ). The GP and NP uncertainties are based on (O 2 /Ar), δ 17 O and δ 18 O variations between dissolved O 2 samples that were collected from a vertical profile through the mixed layer (1σ of the mean, n = 3 for each sampling day). mer and up to 100 m in winter) and a persistent stratification throughout the year are characteristic features of SEATS. The average maximum mixed-layer depth at SEATS is ∼ 80 m, occurring in December and January. Throughout spring, the mixed-layer depth steadily decreases, reaching a minimum of ∼ 25 m in May. Afterwards, the mixed layer increases again gradually, reaching ∼ 35 m in June, and remains approximately constant through to October, after which it increases sharply to reach its maximum winter values (Tai et al., 2017). The chlorophyll fluorescence was generally low and restricted to the thermocline in the upper 50-100 m (Fig. 3), as expected for the oligotrophic northern SCS (see Sect. 1), with the absolute magnitude of the subsurface maximum peak varying between seasons. Interestingly, Chl a was highest with 0.60 mg m −3 in October 2013 (Fig. 3a). In August 2014 the concentration remained at 0.20-0.30 mg m −3 without pronounced variations or diurnal trends (Fig. 3b). In April 2015, we observed a minor increase in the subsurface chlorophyll maximum (but mostly restricted to the dawn hours) of up to 0.50 mg m −3 , which gradually declined throughout the day and was lowest at night with approximately 0.20 mg m −3 (Fig. 3c).
The dissolved O 2 saturation in the upper 400 m on the different sampling days is shown in Fig. 3d. In October 2013, the mixed layer was saturated between 100 % and 102 %, and below in the thermocline the O 2 saturation decreased. In August 2014, the O 2 was saturated throughout the mixed layer (100 %) on the first collection day and was between 97 %-98 % on the second day. In April 2015, the mixed-layer O 2 saturation hovered between 97 %-98 % on the first day and 102 %-103 % on the second day. Below the limit of the mixed layer, O 2 saturation increased by a few percent on both days in August 2014 as well as on 24 April 2015. A more prominent supersaturated O 2 peak reaching 110 % below the mixed layer was observed on 25 April 2015.
Based on our sampling months and the observed physical parameters, the profile from October 2013 appears to reflect the transition from summer to winter conditions. The lack of basin-wide prevailing monsoon forcing is also evi-dent from surface wind maps (Fig. 2a), indicating that this collection date might largely represent an inter-monsoon period. The shallow mixed layer in August 2014 and southwest wind direction point towards typical summer monsoon conditions (Fig. 2b). Conversely, the mixed-layer characteristics in April 2015 are suggestive of spring conditions, although the surface wind maps still indicate the presence of the northeast winter monsoon in the area (Fig. 2c). We therefore conclude that the April 2015 collection days likely reflect late northeast (winter) monsoon season during spring.

Dissolved O 2 composition: 17 and (O 2 /Ar)
The triple isotope composition of dissolved O 2 profiles from SEATS is shown in Fig. 4. We observed broad variations in both the 17 and the (O 2 /Ar) between the different months as well as the days when the samples were collected, overall ranging between 22 and 229 per meg and −72 % to 2.2 %, respectively. Overall, the upper-ocean 17 profiles outlined a common trend, with lower 17 in the mixed layer, a peak in the values below and a gradual decrease towards 200 m depth. However, the average mixed-layer values (Table 1) and the depth of the 17 maximum peak, as well as its magnitude, varied considerably between months and collection days. Highest mixed-layer 17 values averaging 90 ± 28 per meg were observed in October 2013. On 5 and 6 August 2014 and 24 April 2015 the O 2 compositions were comparable, yielding mean mixed-layer 17 values of 59 ± 9, 54 ± 18 and 52 ± 11 per meg and (O 2 /Ar) values of −0.3 ± 0.5 %, −0.2 ± 0.2 % and −0.5 ± 3.5 %, respectively. Lowest mixedlayer 17 values of 26 ± 5 per meg with positive (O 2 /Ar) of 1.8 ± 0.3 % were observed on 25 April 2015.
Largest variations in 17 and (O 2 /Ar) were observed within the thermocline. In October 2013, the 17 gradually increased with depth to a maximum of 182 per meg at 80 m and then decreased (Fig. 4a). On the 5 August 2014, the highest 17 values reached 218 per meg at 100 m (Fig. 4b). The depth trend on the 6 August fairly resembled the one from the 5 August, but the 17 values between the 2 sampling

Seasonal trends in photosynthetic vs. atmospheric O 2 input in the upper water column
The combined approach of 17 and (O 2 /Ar) composition of dissolved O 2 has been shown to be a valuable tracer for distinguishing biologically mediated O 2 from that supplied by atmospheric air input to the euphotic zone (Luz et al., 1999;Luz and Barkan, 2000). This is because atmospheric O 2 has a unique isotopic signature generated by stratospheric photochemical reactions involving O 3, O 2 and CO 2 which fractionate its isotopes in a mass-independent way (see Lämmerzahl et al., 2002), while photosynthesis fractionates O 2 isotopes in a mass-dependant way. By definition, the atmospheric 17 atm = 0, although the air-water equilibrium 17 eq deviates slightly from the atmospheric value (    sured at 30 m depth, seem particularly unusual, as any instantaneous increase in photosynthetic 17 signal in the mixed layer is expected to be limited due to continuous exchange with atmospheric O 2 and thus averaged against the background signal. It is also unlikely that these samples could have been affected by contamination, as any leak during the sampling or preparation would result in a decreased 17 value due to influence from atmospheric O 2 . A likely explanation for the observed high 17 would be the rather short O 2 production time (< 1 d) against the relatively long residence time of O 2 in the mixed layer (12 d; see Sect. 2.2), suggesting a sustained accumulation of biologically produced O 2 . The timing of the high 17 values in the mixed layer (Fig. 4a) also coincided with the overall highest observed fluorescence in this study (Fig. 3a). The Chl a maximum was situated below the mixed layer in the thermocline where we also recorded a 17 peak. The elevated mixed-layer 17 values, therefore, could also reflect a transient increase in biological O 2 due to upward flux of photosynthetically produced dissolved O 2 from the Chl a maximum horizon. In such a case the 17 and the primary production would also integrate O 2 from below the mixed layer, and this may complicate the application of the steady-state model. Assuming that these values really reflect the integrated O 2 production, it may suggest rather high photosynthetic activity for an inter-monsoon period, probably enhanced by the onset of cooler temperatures after the summer.
The average mixed-layer 17 and (O 2 /Ar) values from 5 and 6 August 2014 as well as from 24 April 2015 were, within the variations, indistinguishable from each other (about 55 ± 13 per meg and −0.3 ± 0.4 %; Table 1). The observed 17 values reflect the presence of photosynthetic O 2 , which based on the (O 2 /Ar) values was in near-equilibrium concentrations to potentially slightly undersaturated (especially towards the bottom of the mixed layer if individual measurements are considered). Previous studies have largely disregarded negative mixed-layer (O 2 /Ar) values and attributed them to vertical mixing (Reuer et al., 2007;Stanley et al., 2010). During our sampling period, we did not observe significant upwelling of water (internal wave) that could be indicative of diapycnal mixing and contribution from deeper water to the mixed layer, although data spanning the entire residence time of O 2 in the mixed layer would be needed for a more conclusive statement. Recently, Qin et al. (2021) reported (O 2 /Ar) from the northern slope of the SCS collected by continuous measurement of gases in the surface waters (5 m) during approximately 2 weeks in October 2014 and June 2015. While most of their values yielded positive (O 2 /Ar), the authors also reported negative values which they linked to upwelling caused by a localized cold eddy. Hence, it can be expected that vertical mixing would result in negative (O 2 /Ar) values at 5 m in the SCS which, however, was not observed in our data (our (O 2 /Ar) values at 5 m ranged between 0 % to 1.6 %). Although some degree of exchange across the mixed-layer boundary cannot be excluded, we postulate that, overall, our data reflect a minimal photosynthetic activity characteristic of the summer months at SEATS. This may be attributed to the strong seasonal thermal stratification and nutrient depletion (Wong et al., 2017b), as supported by the measured low fluorescence in August 2014 (Fig. 3b), in agreement with past observations (Liu et al., 2002).
Distinct 17 and (O 2 /Ar) values were observed on 25 April 2015 (∼ 26 per meg and ∼ 1.8 %; Table 1). Although the average mixed-layer signal from 24 April 2015 (52 ± 11 per meg and −0.5 ± 3.5 %) is similar to the values from August 2014, if the deepest mixed-layer sample from 20 m depth were removed from the calculation (despite still being situated within the mixed layer), then the newly obtained mixed-layer (O 2 /Ar) for 24 April 2015 would agree well with the values from 25 April 2015 (1.5 ± 0.2 %; with the 17 remaining unaffected 46 ± 7 per meg). This points towards variable conditions near the mixed-layer boundary, potentially influenced by O 2 contribution from below waters where the (O 2 /Ar) was negative. The lower, closer to equilibrium, 17 values on the second day of sampling indicate increased air-sea gas exchange rates, which drive oxygen production as evidenced by the elevated (O 2 /Ar) in the mixed layer (Fig. 4c). This is further supported by the intensified fluorescence (Fig. 3c) and supersaturated dissolved oxygen waters (Fig. 3d) below the mixed layer. The distribution and concentration of the deep chlorophyll maximum corresponds to characteristic monsoon-forced trends (Liu et al., 2002) and demonstrates the vitality of the thermoclinedwelling phytoplankton and the important role of NEM winds on determining the metabolic balance of the system. The overall lower 17 and (O 2 /Ar) values in the upper water column observed in April 2015 when compared to August 2014 (Fig. 4) may illustrate the extent of wind-induced vertical mixing, which could be sufficient to reach the upper limit of the nutricline (e.g. Ning et al., 2004;Tseng et al., 2005) and supply nutrients to the phytoplankton communities. Similar (O 2 /Ar) values were also reported by Qin et al. (2021) for the slope region of the SCS and are typically attributed to nutrients, especially nitrogen, becoming available to phytoplankton. Alternatively, between February and April monsoon winds tend to carry minerals and iron-rich dust particles from the deserts in Central Asia to the northern SCS and SEATS (Lin et al., 2007;Duce et al., 1991;Fung et al., 2000), which could fuel the enhanced biological production. These profiles thus serve as a good example of the local ecosystem interactions and underscore the close dependence of the phytoplankton communities on the NEM forcing.

Primary production rates in South China Sea
Primary production, the synthesis of organic compounds from carbon-containing species, is of critical importance to biogeochemical cycling of carbon and oxygen that sustains the marine ecosystem. In a steady-state system, we may distinguish gross production (GP) and net production (NP), where the former represents the total C fixed by primary producers and the latter the C available to the heterotrophic community. The NP thus amounts to the difference between the GP and community respiration. NP is positive when GP exceeds respiration, and the ecosystem exports or stores organic C, while negative values result when respiration exceeds GP, and the ecosystem respires more organic C than it was able to produce. These terms are of fundamental interest to ocean studies. However, their quantification is not straightforward, and thus far, only limited information is available globally and especially from SEATS and the SCS.
Our production rates are summarized in Table 1, derived from δ 17 O and δ 18 O values of dissolved oxygen using a steady-state mixed-layer oxygen budget model that allows for determining integrated productivity in the mixed layer over the residence time of O 2 (as detailed in Sect. 2.2). We note that these estimates, however, do not account for complex physical processes (vertical mixing, lateral advection) and non-steady-state effects on the mass balance. Furthermore, as discussed in Sect. 2.2, the choice of parametrization method and approach for calculating gas exchange rates introduces uncertainties and merits a careful consideration. The definition of the mixed-layer depth is also highly relevant, although as shown both the temperature-based and dissolved-O 2 -based definition resulted in similar depths in this study. We found that largest uncertainties on the derived production rates (especially on the 16 October 2013 and on the 6 August 2014; Table 1) resulted from variations in δ 17 O, δ 18 O and (O 2 /Ar) between samples collected from a vertical profile through the mixed layer. These variations are reported as the standard deviation of the mean composition of dissolved O 2 in the mixed layer, and they are considered in the calculation of our NP and GP rates.
We found comparable production rates on the two consecutive sampling days in August 2014, with mean GP ∼ 1500 mg C m −2 d −1 , and low NP rates averaging −13 mg C m −2 d −1 . Although these rates appear to indicate that the system metabolism was net heterotrophic, within the uncertainties (based on the (O 2 /Ar), δ 17 O and δ 18 O variations in the mixed layer) it was in net balance. Similar values and overall net (or close to net) balance likely prevails in the SCS during the SWM season, as with the exception of sporadic typhoon events, the environmental conditions can be expected to remain rather stable and the water column strongly stratified. The production was also within the errors comparable on 24 April 2015, yielding GP of ∼ 2000 mg C m −2 and NP of ∼ −40 mg C m −2 d −1 . On the second sampling day, 25 April, the GP was considerably lower (∼ 600 mg C m −2 d −1 ), and the NP was higher (∼ 140 mg C m −2 d −1 ). This points towards a more dynamic system, likely influenced by the NEM conditions. It appears that the NEM forcing could shift the metabolism of the system from being close to net balance to net autotrophy, due to cooler temperatures and wind-induced mixing of the water column, providing nutrients from the subsurface waters into the mixed layer. Potentially, minor contribution to the signal could also come through the input of photosynthetically produced O 2 from below the mixed layer. Highest GP estimates of 6600 mg C m −2 d −1 were obtained in October 2013, which is rather surprising since during inter-monsoon periods phytoplankton production is expected to be limited. The origin of the high GP rates in October 2013 appears to be related to the deeper mixed layer on this sampling day. Elevated 17 values (122 per meg) were measured at 30 m, with some contribution to the signal potentially coming from the photic layer below. These estimates should, therefore, be treated with cau-tion as diapycnal mixing across the base of the mixed layer, and/or heterogenous distribution of phytoplankton in the water column and potential in situ production at depth cannot be ruled out, in which case the steady state no longer applies.
It is to be stressed that our estimates represent the mixedlayer production rates rather than total water column rates. During our sampling campaigns at SEATS, the euphotic zone was persistently deeper (∼ 30-70 m deeper) than the mixed layer (Table 1, Fig. 4). This may lead to underestimation of the true mixed-layer NP values due to mixing or entrainment of low-O 2 waters into the mixed layer and conversely overestimation of the true mixed-layer GP values due to mixing or entrainment of high-17 waters into the mixed layer, since the share of the production that takes place within the euphotic zone below the mixed layer cannot be accounted for by the present model. Nonetheless, it is likely that if respiration exceeds gross production in the mixed layer (and hence the NP is negative), the overall NP in the euphotic zone will also be negative, since deeper regions tend to have higher respiration rates. Thus, production estimates from paired 17 and (O 2 /Ar) profiles are still useful for indicating trends in ecosystem metabolism even in instances when the depths of the mixed and euphotic layers differ. Our findings indicate that over the year respiration might be close to the GP, with phenomena such as the NEM forcing playing an important role in fuelling productivity. Hence, production during winters with cooler temperatures and windy days may play a decisive role in determining the amount of organic C fixed. For instance, an oligotrophic system that is net heterotrophic most of the time could maintain overall net autotrophy with episodic production which occurs 10 % of the time and is 3 times the background rate (Karl et al., 2003). Weakening of the East Asian Monsoon by anthropogenically induced global warming (e.g. Hsu and Chen, 2002;Xu et al., 2006) is, however, likely to limit vertical transport and nutrient supply to the phytoplankton. It is to be seen to what extent this will affect the primary production and overall C balance at SEATS and in the SCS.
Previous primary production estimates from the SCS based on 14 C observations and modelling fall within the range of approximately 120-170 g C m −2 yr −1 (Ning et al., 2004;Chen, 2005;Liu et al., 2002). Ning et al. (2004) studied seasonal patterns in phytoplankton and productivity, and based on 14 C bottle incubation experiments, they estimated the summer and winter daily production within the euphotic zone to be 390 ± 342 and 546 ± 409 mg C m −2 d −1 , respectively. Although the comparison between incubation-based 14 C productivity estimates and those derived from triple oxygen isotopes is not straightforward due to the different natures of the methods and temporal and spatial scales over which they integrate, it is important to attempt to place the different estimates in context of each other. To convert 14 C production estimates from bottle incubations to GP, we used a scaling factor of 1.74 to correct for dark respiration and excretion of DO 14 C (dissolved organic carbon; Hendricks et al., 2004). This gives mean summer-to-winter GP rates of approximately 680-950 mg C m −2 d −1 for the values reported by Ning et al. (2004). Omitting our high GP rates from October 2013, which are likely biased by mixed-layer entrainment, this is lower than our mean rates of about 1410 ± 590 mg C m −2 d −1 . Similarly, previous attempts to compare the 14 C incubation-based and triple oxygen isotope-based estimates have generally found that the latter estimates tend to overestimate the primary production by a factor of about 1-3 (Juranek and Quay, 2005;Quay et al., 2010;Juranek and Quay, 2013;Jurikova et al., 2016). Whether this is due to a high bias from the triple oxygen approach or a low bias from the bottle incubation method remains a matter of on-going debate, which only further measurements of marine productivity using both incubationbased and incubation-independent methods and comparison studies can resolve. We postulate that some of the high bias we found could be due to vertical mixing with deeper waters (which leads to higher 17 ). To a larger extent, the observed differences probably result from the low bias of the bottle incubation approach due to containment effects and/or missed infrequent episodes or patches of high productivity (Quay et al., 2010;Juranek and Quay, 2013), which are known to occur in the SCS. Much of our understanding and discussion of productivity in the SCS, however, remains limited by the availability of data, which at present likely only offers a snapshot of the full picture. Crucially, further measurements of primary productivity at an increased spatial and temporal resolution will provide a more comprehensive picture of production at SEATS and within the wider SCS, reduce the uncertainties involved and inform the discussion of methodological biases.

Comparison to other tropical time series
Initially launched in 1998, and becoming part of the Joint Global Ocean Flux Study 1 year later (JGOFS;Shiah et al., 1999), the SEATS station has often been compared with the time series off Hawaii (the Hawaii Ocean Time-series, HOT), which together with the time series off Bermuda (the Bermuda Atlantic Time-series Study, BATS) were two key components of the former JGOFS programme. In contrast to the typical features of tropical waters, characterized by minimal seasonal variations, and comparing to the low-latitude HOT station (Karl and Lukas, 1996), the SEATS station located at an even lower latitude is characterized by a distinct phytoplankton biomass and primary production pattern (Tseng et al., 2005;Wong et al., 2007a). This distinct pattern is largely governed by the East Asian Monsoon, which brings seasonal changes that affect the oceanography and biogeochemistry of the SCS (Chao et al., 1996a;Liu et al., 2002).
Our seasonal depth profiles from SEATS share some similarities with the tropical oligotrophic HOT station (Juranek and Quay, 2005), albeit with different 17 magnitudes. Notably, the 17 depth distribution pattern at SEATS was com-parable to that at HOT, with a broad summer 17 peak (above 200 per meg at 80 and 100 m depth) from August 2014 comparable to that at HOT during the same month (with values above 140 per meg at 120 and 150 m depth), as well as a high peak in October 2013 (above 180 per meg at 80 m) rather similar to that at HOT during October (above 140 per meg at 100 m). In February, the 17 values were overall much lower at HOT, reaching the highest values in the deep (above 90 per meg between 150 and 200 m). Possibly, such trends could also be expected for SEATS; in fact our data from April 2015 appear to bear a close resemblance to this pattern, although a comparison of the same months would be preferable. Our maximum 17 values within the euphotic zone at SEATS (218 per meg at 100 m observed on 5 August 2014, Fig. 4b) are, however, to our knowledge much higher than any previously documented upper-ocean values at HOT or elsewhere, which typically do not exceed ∼ 160 per meg.
Obviating the estimates from October 2013, our GP and NP rates are comparable to those from HOT derived using the same approach, where seasonal variation in GP and NP were in the range of 800-1470 mg C m −2 d −1 and −120 and 180 mg C m −2 d −1 (Juranek and Quay, 2005). Generally, we would, however, expect higher rates at SEATS, where both seasonal monsoon forcing and/or episodic typhoon events induce sufficient vertical mixing to bring nutrients to the mixed layer and stimulate primary production. Assuming that our rates present an underestimation of the productivity due to the relatively very shallow mixed layer, these differences could reconcile. Our observed variations in NP/GP (∼ −0.01 in August 2014 and between −0.02 and 0.23 in April 2015) also compare well with the seasonal trends reported from HOT (between −0.13 and 0.13 during summer and winter; Juranek and Quay, 2005) with tendency towards heterotrophy in the summer and autotrophy during winter months. Very low NP/GP ratios were also observed from other lowlatitude regions such as the equatorial Pacific (Hendricks et al., 2005;Stanley et al., 2010). This supports the general parallels in ecosystem metabolism in the oligotrophic regions, but also given the broader variations in NP/GP ratios at SEATS, it emphasizes the importance of the monsoon forcing on introducing dynamics to this system.

New insights into the 17 in deep water
The 17 has been traditionally applied for evaluating primary production in the upper ocean, and thus far, only little is known on the 17 composition of the deep ocean. Due to the conservative behaviour of O 2 in a parcel of deep water where it may no longer be influenced by air-sea gas exchange or photosynthesis, the 17 could also present a valuable tracer for deep water mixing processes since any variations in 17 should principally result from mixing of waters with different 17 values. While respiration alone does not affect the tracer, the 17 may, however, behave non-conservatively and be altered by the combined effects of respiration and mixing. As shown by Nicholson et al. (2014), if two hypothetical parcels of water with very different δ 17 O and δ 18 O values (but same 17 values) mix (one with the starting composition of surface water and one that underwent a Rayleigh fractionation until 5 % of oxygen remained), the resulting 17 values can become negative. Subsurface (∼ 100-300 m) measurements in the equatorial Pacific indeed reported few negative values (Hendricks et al., 2005). Measurements from deeper profiles (700 m) were carried out in the Gulf of Elat and showed that the 17 values below the thermocline varied considerably with seasons, a likely result of vertical as well as horizontal mixing (Wurgaft et al., 2013). In order to evaluate the behaviour of 17 in the deep water of the SCS and its potential utility as a mixing tracer in an oceanographically very distinct system, we measured the 17 composition of deep O 2 profiles (down to 3500 m depth) from SEATS.
An overview of the oceanography of SCS is available in Wong et al. (2007a). The subsurface water masses in the SCS consist of three main water masses: (1) the Tropical Water situated at around 150 m, originating from near the international dateline at 20-30 • N in the North Pacific (Suga et al., 2003); (2) the North Pacific Intermediate Water centred around 500 m with a source in the subpolar regions in the North Pacific (You, 2003); and (3) the Deep Water below 2200 m. The Deep Water in the SCS basin is formed by Pacific water masses, which in the western Philippine Sea overflow the sill that separates it from the SCS. The characteristics of the deep water in the SCS are rather uniform and similar to those in the western Philippine Sea. They are maintained by a mass balance between the inflowing deep water from the Philippine Sea, upwelling and mixing with the shallower North Pacific Intermediate Water, and outflow at an intermediate depth through Luzon Strait (Gong et al., 1992;Chao et al., 1996b;Hu et al., 2000).
Our data showed overall elevated 17 values (> 140 per meg) below 200 m depth for both August 2014 and April 2015 profiles. Largest variations were found at 200 m with a decrease of 116 per meg from August 2014 to April 2015 (Fig. 5a), coinciding with changes in the temperature-salinity characteristics of the Tropical Water mass (Fig. 5b). The synergic decrease in 17 and increase in temperature-salinity at this depth illustrate the increased winter inflow of water to the SCS from Kuroshio through the Luzon Strait or possibly also partially from the East China Sea through the Taiwan Strait (Fig. 1). This highlights the importance of NEM winds in driving the seasonal circulation inducing vertical mixing, which extends down to 400 m and leads to a full exchange of water masses during a winter (Fig. 5a). Historic records also support intrusions of North Pacific water masses to the SCS all-year round with greatest strength in winter (Qu et al., 2000). Below, the deeper water remained relatively homogenous, and it did not appear to be influenced by seasonal changes, marking the limit of the extent of monsoon-driven circulation influence on the mixing. Surprisingly, variations in 17 (around ∼ 20 per meg) were found beneath the thermocline base; however, considering the low O 2 content at these depths, even a very minor change in O 2 concentration may result in a large effect on the 17 signal. Although further observations from the SCS are needed for a more comprehensive picture, our first results ad-vocate for the 17 as a valuable tracer of mixing processes, which brings new insights into some of the key aspects of our understanding of the circulation in the SCS.

Conclusions
In summary, in this study we provided the first insights into the 17 composition of dissolved O 2 at the SEATS station and within the South China Sea (SCS). We used the combined 17 and (O 2 /Ar) approach to monitor the seasonal changes in atmospheric vs. photosynthetic O 2 input to the upper part of the ocean and to calculate the integrated mixedlayer GP and NP rates. Our results showed that the net biological production at SEATS was negligible during most sampling days and close to net zero but increased when the system was more dynamic during a spring influenced by the northeast monsoon forcing. Mixed-layer variations and potential O 2 contribution from below the mixed layer, however, contribute uncertainties to these estimates and should be taken into consideration in further work applying this tracer, in particular within the SCS. Furthermore, we found the 17 of the deep water to be a promising tracer for physical mixing process (in addition to biological processes). This permitted us to evaluate the extent of the basin-wide monsoon-driven circulation in the water column and at depth, as well as to revisit the deep mixing processes. Although further work is required before the deep 17 may be confidently applied as a tracer of water mass mixing, our study showed that it could bring new perspectives on the renewal rate of the deep water, at least within the SCS. Thus, further deep 17 measurements within the region but also globally would be desirable. Similarly, further 17 and (O 2 /Ar) measurements within the SCS at an increased spatio-temporal resolution would be beneficial for gaining a more comprehensive picture of primary productivity dynamics at SEATS and its link and responses to the East Asian Monsoon as well as other episodic or interannual phenomena.
Data availability. The data used in this study are available in the Supplement including geochemical data (Tables S1 and S2), CTD data (Table S3) and wind speed data (Table S4).
Author contributions. HJ and MCL designed the study and performed the analyses. OA and FKS contributed analytics tools and data. HJ prepared the article with contributions from all co-authors.
Competing interests. The contact author has declared that neither they nor their co-authors have any competing interests.
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