Journal of Oceanology and Limnology   2022, Vol. 40 issue(2): 530-550     PDF       
http://dx.doi.org/10.1007/s00343-021-1010-4
Institute of Oceanology, Chinese Academy of Sciences
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Article Information

YE Wangwang, DU Guanxiang, ZHANG Honghai, ZHANG Guiling
Methane in the Yellow Sea and East China Sea: dynamics, distribution, and production
Journal of Oceanology and Limnology, 40(2): 530-550
http://dx.doi.org/10.1007/s00343-021-1010-4

Article History

Received Jan. 7, 2021
accepted in principle Feb. 26, 2021
accepted for publication Apr. 2, 2021
Methane in the Yellow Sea and East China Sea: dynamics, distribution, and production
Wangwang YE1,2,3#, Guanxiang DU1,3#, Honghai ZHANG1,3, Guiling ZHANG1,3     
1 Frontiers Science Center for Deep Ocean Multispheres and Earth System, and Key Laboratory of Marine Chemistry Theory and Technology, Ministry of Education, Ocean University of China, Qingdao 266100, China;
2 Key Laboratory of Global Change and Marine-Atmospheric Chemistry, Third Institute of Oceanography, Ministry of Natural Resources, Xiamen 361005, China;
3 Laboratory for Marine Ecology and Environmental Science, Pilot National Laboratory for Marine Science and Technology(Qingdao), Qingdao 266237, China
Abstract: The Yellow Sea (YS) and East China Sea (ECS) are important marginal seas of the western Pacific. Understanding the dynamics of methane (CH4) in the YS and ECS are essential to evaluate the role of coastal seas in global warming. We measured dissolved CH4 at various depths in the water column of the YS and ECS during a cruise from March to April 2017. The concentrations of CH4 varied greatly in different water masses, suggesting that the hydrographic conditions can substantially affect the CH4 distribution. The CH4 budget in the shelf of the ECS, which was estimated with a box model, suggests CH4 consumption in the water column was the major sink (>95%), followed by a loss with a total of 2.2% CH4 released to the atmosphere. Overall a local CH4 production of 0.28 nmol/(L·d) was needed to maintain the CH4 excess. Results from laboratory incubations showed an increase in CH4 (1.5 times higher than the value of the control) after the addition of dimethylsulfoniopropionate (DMSP). Field incubations result in a CH4 production rate of 1.2 nmol/(L·d) under a N-stressed conditions (N: P<1), indicates that the DMSPdependent CH4 production prefer to occur in the oligotrophic seawaters, where nitrogen is depleted. This study demonstrates that the marginal seas of China is a hotspot for CH4 dynamics, and the cycling of methylated sulfur compounds (such as DMSP) may contribute importantly to locally formed CH4. This may have further implication to carbon and sulfur biogeochemical cycles in the western Pacific.
Keywords: methane    East China Sea    Yellow Sea    aerobic production    dimethylsulfide (DMS)    dimethylsulfoniopropionate (DMSP)    
1 INTRODUCTION

Methane (CH4) is an important greenhouse gas that influences the global climate (IPCC, 2013). Emissions of individual sources have been poorly quantified, even though the total global emissions of CH4 into the atmosphere have been reasonably constrained (Saunois et al., 2020). Oceans cover 71% of the planet's surface and are considered a minor source of atmospheric CH4 (Saunois et al., 2020). The CH4 concentrations in the surface water of most of the world's oceans are 5%-75% supersaturated with respect to atmospheric CH4 (Karl et al., 2008; Damm et al., 2009), implying a local in-situ methane source. In particular, studies showed that near-shore environment contributes the largest but most uncertain diffusive flux of the oceanic CH4, but comprise only 3% of ocean surface areas (Weber et al., 2019). This observation can be partially attributed to the high CH4 production rates and fast ventilation of the shallow water column (Borges et al., 2018; Schmale et al., 2018), or CH4 escape from localized gas seeps via bubbles (Sakai et al., 1990; Judd, 2004; Shakhova et al., 2014; Di et al., 2020).

The Yellow Sea (YS) and East China Sea (ECS) are marginal seas in the northwestern Pacific Ocean, with average water depths of 44 m and 350 m, respectively. The distribution of dissolved CH4 in the YS and ECS exhibits great temporal and spatial variability, which has been well documented in previous studies (Zhang et al., 2004, 2008a, b; Yang et al., 2010; Ye et al., 2016; Sun et al., 2018). Although several studies have reported that the surface and subsurface layers is supersaturated with CH4 with respect to the atmosphere (Zhang et al., 2004; Ye et al., 2016; Sun et al., 2018), the mechanisms are still unclear.

Physical processes can contribute to the CH4 accumulation at the ocean surface, including lateral transport from the CH4-rich sources (e.g. estuaries; Borges et al., 2018) or vertical diffusion from subsea CH4-hotspots (e.g. gas hydrates, cold seeps, and pockmarks) (Judd, 2004; Shakhova et al., 2014; Di et al., 2020). In the ECS, CO2-rich fluids with other residual gas (including CH4) were observed to emerge from the sea floor in the hydrothermal field (Sakai et al., 1990), suggesting the gas hydrates can serve as potential source to the aqueous CH4 (Luong et al., 2019). Biotic factors contributing to the excess CH4 was generally attributed to local production by microorganisms in anoxic microniches (de Angelis and Lee, 1994; Karl and Tilbrook, 1994; Stawiarski et al., 2019; Wäge et al., 2020). A recent mechanism that was proposed for maintaining surface CH4 excess is that CH4 is released as a byproduct via the decomposition of methylated compounds, such as methylphosphonate (MPn) and dimethylsulfoniopropionate (DMSP), when heterotrophic bacteria compete for nutrients (Repeta et al., 2016; Teikari et al., 2018). Studies for MPn identified the gene content and expression and thus confirmed the possibility of a Carbon-Phosphorus pathway for CH4 formation at the biological level (Metcalf et al., 2012; Martínez et al., 2013; Sosa et al., 2017, 2020). Alternatively, DMSP is mainly produced in the cells of phytoplankton and can be transferred to the ambient water by grazing or sloppy feeding, so the yield and metabolism ratio of DMSP might be important for CH4 formation. Damm et al. (2008) reported CH4 production in the Arctic sea that was associated with the metabolism of DMSP and its degradation products. However, evidence for the formation of CH4 from DMSP remains ambiguous and is mainly based on occasional correlations between CH4 and DMSP concentrations (Damm et al., 2008; Florez-Leiva et al., 2013; Zindler et al., 2013; Damm et al., 2015; Zhai et al., 2019). In addition, the cleavage of the Carbon-Sulfur bond pathway (DMSP-to-CH 4) for CH4 production has been rarely reported in other marine regimes, such as the marginal seas of the Pacific.

In this study, we examine the excess CH4 in the upper water column of the YS and ECS and combine hydrographic conditions to discuss the factors that may influence the CH4 dynamics in oxygenated waters. Then, we develop a preliminary CH4 budget using a box-model to quantify CH4 sources and sinks in the shelf of the ECS and evaluate the contribution from local CH4 production to maintenance of the water column's CH4 inventory. Finally, in combination with the field observation, we conduct laboratory incubation experiments alongside field observations to provide evidence that DMSP can serve as a potential precursor of CH4, probably via demethylation in the YS and ECS. Thus, the understanding of CH4 dynamics in the YS and ECS was improved by providing new data, conducting incubation experiments, interpreting these data and comparing with previous results (Zhang et al., 2004; Ye et al., 2016; Sun et al., 2018).

2 MATERIAL AND METHOD 2.1 Research area

The YS is a semi-enclosed continental shelf sea and is connected to the ECS and the Bohai Sea. The boundary between the ECS and YS spans from the northern tip of the mouth of the Changjiang (Yangtze) River to Cheju Island (Fig. 1a). In winter and early spring, strong northerly winds that accompany a surge of cold, dry continental air tend to vertically homogenize most of the YS (Su, 1998). The coastal current flows southward near the Shandong Peninsula in the north and along the 50-m isobath in the south (Fig. 1a). Water exchange between the YS and ECS mainly occurs near this boundary. Thus, three dominant water masses exist in the YS in spring, namely, the coastal water (CW) in the west, YS Central Water (YSCW) in the central trough, and YS-ECS Mixed Water (YEMW) near the boundary (Li, 1989; Su, 1998; Guan and Fang, 2006).

Fig.1 Schematic diagram of the circulation pattern in spring a. water circulation modified from Guan and Fang (2006) and Yang et al. (2018); b. sampling locations in the Yellow Sea and East China Sea from March to April 2017. The six representative stations discussed in this study are shown in (a). The red-dashed line marks the boundary of the YS and ECS. The base map is derived from the Ministry of Natural Resources of China (No. GS(2020)4388). CW: Coastal Water; YEMW: YS-ECS Mixed Water; YSCW: YS Central Water; TWCW: Taiwan Warm Current Water; ESW: ECS Shelf Water; KW: Kuroshio Water.

In spring, the western YS is mainly influenced by the Coastal Water, which is fresh (salinity<32.0) and nutrient poor (nitrate<1.0 μmol/L) (Jin et al., 2013). The coastal water contains two components in the YS: Shandong coastal water forms around the Shandong Peninsula and flows southward along the 40-50-m isobaths, while the Subei CW lies in the southern, coastal regions, flowing southeast to the ECS and merging with remnants of Changjiang diluted water (Hickox et al., 2000). The YSCW is one of the main water masses that exist in the central YS throughout the year. This water displays great seasonal differences in salinity (31.0-32.0) and temperature (5.0-25.0 ℃) because of seasonal stratification (Su, 1998). The YEMW is present in the southeastern YS, which consists of a branch of warm and saline water (from the ECS) that is mixed with YSCW. Thus, the YEMW exhibits high temperature (> 9.0 ℃) and salinity (>32.5) in spring.

In the ECS, the Changjiang River introduces freshwater from the west and flows southward in spring (Guan and Fang, 2006). From the east, the Kuroshio introduces water of high temperature and salinity, flowing northeast along the edge of the ECS and eventually turning east to the Korea Strait and Japan Sea (Fig. 1a) (Tang, 1997; Li et al., 2012; Qi et al., 2014). Water exchange between the ECS and Kuroshio occurs across the shelf break through frontal processes at the surface and upwelling in the subsurface (Yang et al., 2018; Zhou et al., 2018). Importantly, this exchange introduces an intrusion of shelf water to the Kuroshio as turbidity tongue, increasing the exchanges in materials (such as nutrients and rare earth element) (Matsuno et al., 2009; Luong et al., 2018; Wang et al., 2019) and may affect the CH4 concentrations at the Kuroshio subsurface (Luong et al., 2019). From the south, the Taiwan current flows to the north throughout the year between the 50- and 100-m isobaths (Su, 1998). In this study, we classified four dominant water masses in the ECS, namely, the coastal water (CW), ECS Shelf Water (ESW), Taiwan Warm Current Water (TWCW), and Kuroshio Water (KW) (Qi et al., 2014). The Kuroshio Water is usually considered to consist of four water masses, namely, the Kuroshio Surface Water (KSW), Kuroshio Sub-Surface Water (KSSW), Kuroshio Intermediate Water (KIW), and Kuroshio Deep Water (KDW) (Su, 1998; Zhou et al., 2018). The coastal water in the YS and ECS can be divided into three components: the Shandong, the Subei, and the Zhe-Min CW (Fig. 1a). For simplicity, we refer to these three water masses as coastal water in this study.

The Zhe-Min CW (Fig. 1a) forms from the mixing of seawater with runoff from the Changjiang River, Minjiang River, and other rivers. From spring to summer, the coastal water expands eastward and retreats northward, becoming warmer and fresher (salinity<32.0; Fig. 2b) (Qi et al., 2014). In contrast, the Kuroshio water on the outer shelf of the ECS showed relatively high temperature (> 23.0 ℃) and salinity (> 34.2). The TWCW originates from the Taiwan Strait, bringing warm (> 17.0 ℃) and saline water (34.0-34.5) to the ECS (Yang et al., 2018; Zhou et al., 2018). This water runs northward along the 60-m isobath parallel to the coastal line and encounters the Subei CW in the eastern region off the Changjiang River (Fig. 1a). The water mass in the middle continental shelf consisted of ECS water, TWCW, and a mix of KW and CW. These water masses have similar temperature-salinity (T-S) characteristics (Qi et al., 2014), so we collectively call them ESW.

Fig.2 T-S diagrams and CH4 concentrations (nmol/L) with isopycnic contours of the density (dashed lines, kg/m3) a. Yellow Sea; b. East China Sea. The classification of the water masses was based on Tang et al. (1997), Su (1998), Qi et al. (2014), Yang et al. (2018), and Zhou et al. (2018). CW: Coastal Water; YEMW: YS-ECS Mixed Water; YSCW: YS Central Water; TWCW: Taiwan Warm Current Water; ESW: ECS Shelf Water; KSW: Kuroshio Surface Water; KSSW: Kuroshio Sub-Surface Water; KIW: Kuroshio Intermediate Water; KDW: Kuroshio Deep Water.
2.2 Field sampling

In early spring of 2017, a cruise was conducted in the YS and ECS on the R/V Dong Fang Hong 2 from 27 March to 11 April. A total of 57 stations along 10 lines were investigated during the cruise to obtain the overview distribution of biogenic gases in the study area (Fig. 1b). Surface seawater for laboratory incubation experiments was collected from the coastal areas of the YS from November 2017 to March 2018. Seawaters for field incubation experiment were collected from the western North Pacific in a cruise during October 2018.

Water samples were collected at various depths by using 10-L Niskin bottles that were mounted on a rosette. At each station, we used Sea-Bird conductivity, temperature, and depth (CTD) probes (SBE 911 plus), which were equipped with ancillary sensors, to measure the salinity, temperature, and dissolved oxygen (DO) as functions of depth. Replicate samples for CH 4 analysis were obtained from the Niskin bottles in 100-mL glass vials by an overflow of approximately 1.5-2 times the bottle volume without bubbles. Afterwards, saturated mercuric chloride (HgCl2) was added into the vials to inhibit microbial activity. Then, the vials were immediately sealed with poly tetra fluoroethylene (PTFE)/silicone septum and aluminum caps to avoid air contamination. Water samples for determining the dimethylsulfide (DMS) and DMSP concentrations were collected by using acid-cleaned polycarbonate bottles through silicone tubing that was attached to the Niskin bottles.

2.3 Box-model

The CH4 budgets for the ECS shelf were evaluated by using a box-model, which was based on the concept of the conservation of water and salt masses (Chen and Wang, 1999; Zhang et al., 2007, 2018; Günthel et al., 2019). Briefly, we assume that the CH4 inventory was temporarily constant in a specific water mass. Thus, the CH4 concentration in these different water masses was the mean value of the stations that belonged to these distinct water masses (classified by the salinity and temperature, Fig. 2). Hence, the CH4 fluxes (Fi, mol/s) during water transport could be obtained simply by the water flux (Qi, m3/s) times the CH4 reservoir (Ci, nmol/L) in a specific water mass: ΣQi×Ci=Fi. Then the CH4 mass balance in the ECS shelf can be expressed as: ΣFi+Rair+Rsed+Rcon+Rpro=0, where Fi represents the water-transported CH4 input (+) and outflow (-) over the shelf (Table 1), R (mol/s) indicates the rates of CH4 accumulation (+)/ elimination (-), including sea-to-air exchange (Rair), sedimentary release (Rsed), microbial consumption (Rcon), and production (Rpro). The Rair and Rsed were computed by the interface-fluxes (μmol/(m2·d)) times the shelf area (~5.1×105 km2) as indicated by Ye et al. (2019). The Rcon was calculated by the shelf water volume (~2.7×1014 m3) multiplies the CH4 oxidation rate (OR, nmol/(L·d)). The OR was calculated using a first-order equation (OR=k×CESW) as reported by Rogener et al.(2018, 2019), where k is the turnover rate constant and CESW is the mean CH4 concentration in the ESW (Table 1). The net production rate (Rpro) was the balance of other sources and sinks.

Table 1 CH4 concentration in water mass and the CH4 rate used in box-model in the ECS
2.4 Incubation experiment

Four separate incubation experiments were conducted to study the links between DMSP/DMS degradation and CH4 production. The laboratory experiments were conducted at the Ocean University of China (Qingdao) and the field incubation experiment was conducted onboard. The seawater for these experiments was placed into an acid-cleaned polycarbonate carboy before the start of the experiment. For the first experiment, a final concentration of 100-μmol/L DMSP was directly added into the carboy. After mixing homogeneously, the seawater in the carboy was transferred into 160-mL glass bottles as the DMSP-amended group. Subsamples were dispensed into the glass bottles prior to the addition of DMSP as the control treatment. This experiment was conducted in November 2017 and repeated in January 2018. The second incubation experiment was set to identify the effect of nutrient-stressed conditions (in particular, N deficiency) on DMSP-dependent CH4 production. Specifically, the carboy was amended with phosphate (P) to a final concentration of 10.0-μmol/L. Subsamples were dispensed into the glass bottles as the control treatment. Afterwards, DMSP was added to the carboy to a final concentration of 1.0 mmol/L, and then the water was placed into another group of glass bottles as the (P+DMSP) treatment (N: P<1, N-depleted). Finally, additional subsamples were collected after the carboy was sequentially amended with 50.0-μmol/L and 160.0-μmol/L nitrate, which were the (P+DMSP+N) treatments (5<N: P<16, N-stressed or N-rich). The third incubation experiment was implemented to test the direct effect of spiked DMS on CH4 production. Similarly, subsamples were sequentially transferred from a carboy that was amended with different nutrients, including glucose (100.0-μmol/L C), nitrate (16.0-μmol/L N), phosphate (1.0-μmol/L P), and DMS (100.0 mmol/L). Then, the bottles were assigned to the control (without addition), the (C+N+P) treatment, and the (C+N+P+DMS) treatment. The fourth incubation was conducted in the western North Pacific to test the DMSP-dependent CH4 production in the oligotrophic areas (station E24, 130°E/5°N, see detail in Ye et al. (2020)). The DMSP (final concentration of 10.0 μmol/L) was amended directly into the surface waters without any other additions, which was the DMSP treatment. Natural seawaters incubated under the same conditions then assigned to the control. All the bottles were filled with 150-mL seawater and 10-mL headspace volume (except for the fourth experiment with 50-mL seawater and 10-mL headspace volume), sealed with PTFE/silicone septum, crimped with aluminum caps, and incubated in a 12-h: 12-h light-dark cycle incubator or in the onboard-surface-water-circulation-system that was maintained at nearly in-situ temperatures (~25 ℃). At each time point per day, triplicate bottles per treatment were randomly sacrificed to obtain the CH4 and DO concentrations. The DO was measured by using the Winkler titration method (Bryan et al., 1976).

2.5 Chemical analyses

CH4 was determined by the headspace method for the incubation samples (Bange et al., 2010) or the gas-stripping method for the field samples (Zhang et al., 2004). The CH4 in the headspace was measured by a Shimadzu GC-14B gas chromatograph that was packed with Porapak Q (80-100 mesh) and fitted with a flame ionization detector (FID). The CH4 was quantified by calibrating peak areas to the FID's response to a three-point calibration by using known-volume injections of CH4 standards (CH4: N2 mixtures of 2.0×10-6, 4.0×10-6, and 50.0×10-6, China Institute of Metrology). For the field samples, seawater was introduced into a stripping chamber and purged with high-purity nitrogen. After bubbling, the dissolved CH4 was passed through a desiccant tube to remove water vapor. The CH4 was concentrated onto a stainless steel trap that was filled with 80-100 mesh Porapak Q and then was released into a Shimadzu GC-14B gas chromatograph for separation and quantification (Zhang et al., 2004).

A purge-and-trap system was used to analyze the DMS concentration, as previously described by Yang et al. (2008). Dissolved DMSP (DMSP d) was determined by using the small-volume drip filtration method by Li et al. (2016). Briefly, a known volume of seawater was gently filtered by a Whatman GF/F glass fiber filter (0.7 μm). The DMSP in the filtrate was then converted completely to DMS by adding sodium hydroxide (5-mol/L NaOH), and the generated DMS was analyzed by the same technique. The total DMSP (DMSP t) concentrations were directly analyzed from unfiltered alkaline subsamples. The concentrations of particulate DMSP (DMSPp) were calculated by subtracting DMSPd from the DMSPt value (Li et al., 2016).

The seawater for chlorophyll a (chl a) analysis was first filtered through a 0.45-μm pore-size filter (Whatman GF/F). Then, the chl a was extracted by soaking in 90% acetone and measured by using a fluorescence spectrophotometer (Turner Designs, USA) (Zhai et al., 2019).

2.6 Statistical analysis

Principal component analysis (PCA) is a multivariate statistical analysis technique, in which a group of correlated variables is transformed into a new group of variables that are uncorrelated or orthogonal to each other (Jackson, 1991). The datasets of the field cruise were analyzed by using PCA and the correlation matrix. Prior to the statistical analysis, the variables were standardized to eliminate the influence of differences in the data magnitude and measurement scales (Webster, 2009). A VARIMAX rotation was performed to aid with the interpretation of the results. Factors with eigenvalues of 1.0 or greater were selected and considered significant. Factor loading (FL) is defined as the correlation coefficient between the principal component score and each variable (Yamamoto et al., 2014). Additionally, univariate analysis of variance (ANOVA) was used to determine whether the CH4 concentrations and production rates varied between treatments with or without the addition of DMS/ DMSP. The homogeneity of the variance was checked by using Levene's test prior to the ANOVA. Differences were considered statistically significant when P was less than 0.05.

3 RESULT AND DISCUSSION 3.1 CH4 in different water masses in the YS and ECS

The main water masses that existed in the YS during March 2017 were classified in the T-S plot to constrain the CH4 biogeochemical processes (Fig. 2a). The CH4 concentrations varied greatly within the different water masses in the shallow YS, ranging from near atmospheric equilibrium (3.3 nmol/L, 107%) in the YSCW to a mean value of 4.8±1.1 nmol/L in the coastal water and a peak value of 6.9 nmol/L (241%) in the YEMW (Table 1). This distribution resulted in a north-south concentration gradient on the shelf because the concentrations tended to increase towards the south of the YS. Indeed, the highest CH4 concentration (14.8 nmol/L) was observed in the southern YS at station D1 (Fig. 1b). This high value could be attributed to the CH4-rich Changjiang diluted water with low salinity (25.8) and density (< 23.5 kg/ m3). Consequently, the CH4 concentration at station D2 (mean: 6.1±1.5 nmol/L) was the result of mixing with diluted water and Coastal Water, corresponding to relatively high salinity (32.3) and density (24.4 kg/ m3). The surface water of the Changjiang River was reported to contain high CH4 (e.g., > 100.0 nmol/L) (Zhang et al., 2008a; Ye et al., 2016; Sun et al., 2018). Thus, mixing between diluted water and the coastal water near the boundary of the YS could have partially contributed to the north-south CH4 gradient. In contrast, lower CH4 concentrations (mean: 3.9±0.2 nmol/L) were observed in the northern central regime, as influenced by YSCW (such as stations A3 and A4; Fig. 1b). Station D7 was located in the western area off Jeju Island. Although this area belongs to the YS in this study, the dominant water mass at this station was ECS Water according to the T-S characteristics (temperature: 12.2-12.8 ℃, salinity: 33.8-34.1, density: 25.5-25.8 kg/m3; Fig. 2a & b), which implies the exchange of YS and ECS water near the boundary.

The CH4 varied greatly in the different water masses across the shelf of the ECS (Fig. 2b), ranging from undersaturation (2.2 nmol/L, 66%) to significant oversaturation (15.7 nmol/L, 566%) with respect to the atmosphere. In the shallow coastal areas (depth<50 m) and middle continental shelf (depths of 50- 100 m), the CH4 concentrations generally decreased with increasing longitude, ranging from an average of 10.9±3.8 nmol/L in the CW to 4.8±0.5 nmol/L in the ESW and 4.8±0.4 nmol/L in the TWCW (Table 1). The rapid CH4 decrease in the ESW compared to the CW was likely the result of microbial CH4 oxidation in the water column and emission into the atmosphere (Zhang et al., 2008b; Ye et al., 2016; Sun et al., 2018). At stations with water depths greater than 100 m, the CH4 concentrations were typically less than or near equilibrium to the atmosphere. These areas were essentially influenced by the Kuroshio (Fig. 1b), and the low measured CH4 concentrations in the KIW (mean: 3.9±0.4 nmol/L) and KDW (3.2±0.7 nmol/L) indicated that CH4 had been consumed in the water column. An exception occurred at water depths of 200-300 m, where CH4 peaks (4.8-8.2 nmol/L) were observed. No other source (such as lateral transport or vertical diffusion) was present in the Kuroshio-influenced area (Zhang et al., 2004), so we speculate that these maxima were correlated with in-situ CH4 production, which will be discussed in the following sections. Thus, the CH4 concentrations were believed to be associated with different water masses (e.g., the CH4 in the coastal water was more than two times higher than the CH4 in the Kuroshio). However, the vertical CH4 distribution showed great spatial variations, and the water masses could only generally explain the CH4 sources in the YS and ECS.

3.2 Vertical distribution of dissolved CH4 in the YS and ECS

The depth profiles of hydrographic features and CH4 concentrations varied greatly in the entire study area, and they could be classified into different sub-groups according to the dominant water masses. We found that stations in the same group had similar features in the vertical profiles. Thus, the depth profiles at six representative stations were chosen to show the spatial variability of CCW, YSCW, and YEMW, while stations T2, E4, and F6 in the ECS were dominated by TWCW, ESW, and KW (Fig. 1b).

Station C1, which had low salinity (31.1) and temperature (9.4 ℃), is located in the western YS, where the coastal water is the dominant water mass. The average water depth was quite shallow (16 m), and the physical parameters were well mixed from the surface to the bottom in early spring (Fig. 3a). Because of this strong vertical mixing, the dissolved CH 4 barely varied in the water column, ranging from 5.2 nmol/L at the surface to 5.0 nmol/L at the bottom. The water is oversaturated in CH 4 with respect to the atmospheric equilibrium, implies that CH4 can be significantly released from the water column to the atmosphere. Station B6 is located in the central trough of the YS, where YSCW was the dominant water mass during the investigated period (Fig. 3b). The overall CH4 concentration at station B6 (mean value of 4.4±0.2 nmol/L) was lower than that at station C1 (mean value of 5.1±0.1 nmol/L); as indicated previously, the CH4 had lower concentrations in the YSCW than in the CW (Fig. 2a). The input of fresh water (i.e., Sheyang River) near the coast can explain the CH4 difference between the coastal area (CW) and the central YS (YSCW). However, the CH 4 concentrations at both stations were 1.5-1.8 times higher than the atmospheric equilibrium and were also higher than the reported values from previous studies for the same season (Zhang et al., 2004). The excess CH4 in the well-mixed water could have been released into the atmosphere, with a mean rate of 1.3 μmol/(m2·d) (Zhang et al., 2004). Thus, additional CH4 sources should have existed to sustain a stable value of CH4 (such as 4.4 nmol/L at station B6) in the water column. Sediment release did not appear to be an important source for the water column's CH4 in the YS and barely contributed to the atmospheric CH4 because of the low organic-carbon concentrations (~0.3%) in the sediment and the presence of abundant DO (> 9.0 mg/L) in the bottom water (Zhao et al., 2018). Moreover, only 5.4% of the sedimentary organic carbon was degraded in the YS sediments, which was three times lower than that in the ECS (Zhao et al., 2018). The low remineralization rates of organic carbon restrict the upward flux of CH 4 in sediments (Valentine, 2002). Consequently, limited CH 4 could diffuse through the sediment-water interface in the YS because of anaerobic/aerobic CH4 oxidation (Jørgensen et al., 2001; Treude et al., 2005). In addition, CH 4-rich CW could not be transported to the central YS (such as station B6), so the extra CH4 source should have been attributed to in-situ CH4 production and our incubation experiments suggested that DMSP may have been a potential substrate for CH4 production in the study area (see details in Section 3.4).

Fig.3 Vertical profiles for the temperature (℃), salinity, DO (mg/L), chl a (μg/L), and CH4 (nmol/L) at six representative stations in the YS (a-c) and ECS (d-f) a. station C1; b. station B6; c. station D6; d. station T2; e. station E4; f. station F6.

At station D6 which is the boundary of ECS and YS (Fig. 3c), the surface layers were influenced by the YEMW, but the bottom layer was likely a component of the ECS water. CH4 concentrations increased from 5.5 nmol/L at the surface to 6.5 nmol/L at a depth of 30 m and then dropped to 4.3 nmol/L at the bottom. Thus, the high CH4 concentrations (> 5.0 nmol/L) in the upper 30 m were the result of mixing, while the relatively low values (< 5.0 nmol/L) below this layer were characteristic of ECS water, further indicating the influence of the water masses on the CH4 distribution.

Significant correlations between excess CH4 (ΔCH4, Cin-situ-Cequilibrium) and salinity (P<0.05, Fig. 4a) were found in the water columns both in the YS and ECS, suggesting that the mixing with fresh water has important effect on CH4 distribution in the nearshore regions. In the ECS, however, more complex hydrography influences CH4 distribution. For example, at station T2 near the coast (Fig. 3d), the water layer of the surface 30 m was influenced by the coastal water and while water below 30 m could be attributed to the TWCW (Li et al., 2012). Consequently, the chl-a concentration dropped to<0.5 μg/L below 30 m, and the CH4 concentration was higher in the surface layer (5.6 nmol/L) than in the bottom layer (4.4 nmol/L). Hence, we carefully chose stations that could represent the properties of different water masses to calculate the CH4 inventories in these waters, and the mean value of CH4 in the TWCW was 56% of the mean value in the CW (Table 1). However, the CH4 concentrations in the water masses in the ECS exhibited seasonal variations. For example, our observed CH4 in the TWCW in early spring was 21% lower than what was reported for summer (Zhang et al., 2008b).

Fig.4 Correlations between ΔCH4 (nmol/L) and other parameters a. salinity vs. ΔCH4. Blue circles and red triangles represent data constrained in the YS and the ECS at depth below 20 m. Black line showed the best fit of the total data; b. DMS (blue triangles)/DMSP (red diamonds) vs. ΔCH4. Date were constrained from those offshore stations (water depth > 800 m) within the water column of 0-300 m.

Station E4, which is located in the middle of the continental shelf, had a well-mixed water column below a depth of 15 m (Fig. 3e). The CH4 only decreased by 0.4 nmol/L within the top 50 m and then reached a maximum of 5.1 nmol/L in the bottom water, suggesting a possible CH4 source from sediments. Previous studies showed that the sedimentary release of CH4 in the ECS is an important source of dissolved CH4 in the water column, and the release rates from the sediment to the water column varied 0.6-2.3 μmol/(m2·d) in season (Zhang et al., 2008a; Sun et al., 2018). Thus, the sediment-water CH4 flux was ~10% of the sea-air CH4 flux (Sun et al., 2018), indicating that sediment was a potential source of the dissolved CH4 in the ECS. Our results showed that the CH4 in the ESW was comparable to that in the TWCW (Table 1), matching the summer observations by Zhang et al. (2008b).

Station F6 is located at the edge of the Okinawa Trough, where the mainstream of Kuroshio flows northeastward and intrudes the ECS along the continental slope (Figs. 1a & 3f). Thus, the main features of the Kuroshio can be captured from this station. Specifically, the high temperature (22.1 ℃) and salinity (34.8) in the surface layer indicate the existence of KSW. The maximum and minimum salinity at depths of ~150 and ~500 m showed the characteristics of KSSW and KIW, respectively. In addition, the low temperature (< 6.0 ℃) and high density (> 27.0 kg/m3) at depths below 700 m showed the features of KDW. Consequently, the CH4 increased in the subsurface layer (50-300 m) and reached a maximum of 4.9 nmol/L. This value subsequently decreased between depths of 300 and 800 m and fluctuated around the atmospheric equilibrium in deep water (800-1 000 m). This CH4 profile showed the characteristics of the Kuroshio, with relatively rich CH4 occurring in the KSSW (4.6±1.1 nmol/L) and KIW (3.9±0.4 nmol/L) and poor CH4 occurring in the KSW (3.2±0.5 nmol/L) and KDW (3.2±0.7 nmol/L). The CH4 in the KSSW was comparable to that in the ESW, probably resulting from an intrusion of shelf water as a turbidity tongue (Luong et al., 2018, 2019). In contrast, CH4 in the KSW and KDW were 33% and 71% lower than that in the ESW and CW, respectively (Table 1). Compared to previous studies, the observed CH4 in the KSW in this study was 1.1-1.4 times higher than that in the spring and summer (Zhang et al., 2004, 2008b), but 20% lower than that in the autumn (Sun et al., 2018). The temperature may be a driver of CH4 variations in the KSW during different seasons (Sun et al., 2018). Meanwhile, the upward diffusion of CH4 in the subsurface layer may be another factor that determines the magnitude of surface CH4. Subsurface CH4 maxima at depths between 200 and 300 m were observed at the deep stations (> 800 m, stations P7, FJ7, E7, and F6). Consequently, CH4 gradients existed within the top 200-300 m, with the intensity ranging from 0.006 nmol/(L·m) to 0.026 nmol/(L·m). The presence of these gradients implies that excess CH4 formed locally in the subsurface water, subsequently diffused to the surface, and eventually entered the atmosphere.

3.3 CH4 budget in the shelf of the ECS

The ECS shelf is likely a transition zone between the coast and open ocean. The exchanging of water masses, heat, nutrients, and other materials occur on this shelf, which has far-reaching effects on the northern Pacific (Jiang et al., 2018). Understanding the CH4 dynamics in the shelf water, therefore, is crucial for investigating the CH4 biogeochemical processes in the ECS. In early spring, the Changjiang flows southward and becomes the coastal water by mixing with shelf water, whereas the Kuroshio flows northeastward along the edge of the ECS and intrudes the shelf mainly through the KSW and KSSW (more than 80%) (Su, 1998; Zhou et al., 2018). When examining the water budgets of the ECS, the discharges of KSW and KSSW inflows into the shelf were two orders of magnitude higher than that of the coastal water (Table 1). However, the TWCW comprised ~64% of the total water inflow in spring, and the ECS received ~0.01×106 m3/s influx from the YS. Consequently, the outflows through the Tsushima Strait and Japan Sea were estimated as 2.83×106 m3/s. We estimated the CH4 that was carried by the coastal water to be only 0.11 mol/s, which was 15-19 times lower than what was transported via the KSW and KSSW but comparable with the input through the YS (0.05 mol/s), suggesting that the influx of freshwater was a minor source (0.01%) of shelf CH4 in early spring because of the low water discharge (Table 1; Fig. 5). The gas hydrates located at the Okinawa Trough may play an important role in contribution of high CH4 to the Kuroshio and the CH4 concentrations venting from the hydrothermal field were recorded at millimole level (Sakai et al., 1990; Luong et al., 2019), raising the concerns of gas hydrates in contribution to the shelf CH4. However, the fluids were diluted quickly and diffused northwest carried by the Kuroshio, resulting in a relatively low CH4 inventory (2.2-5.3 nmol/L) (Luong et al., 2019) that matches with our observation at the KSSW (4.6±1.1 nmol/L, Table 1). This means the gas hydrates located at the Okinawa Trough (water depth > 800 m) would have small impact on the CH4 distribution in the ECS shelf (< 200 m) due to diluting and consuming. The TWCW brought the most CH4 to the shelf in spring (8.64 mol/s). Thus, the TWCW could be considered an important source of the shelf's CH4 inventory. The total out-flux was 8% higher than the total input, implying that the ECS shelf was a net CH4 source to the open ocean. The differences between the total inflow and outflow also suggested the existence of other CH4 sources in the ECS (e.g., in-situ production, groundwater, and sedimentary release). However, these CH4 fluxes may exhibit seasonal variations. For example, Zhang et al. (2008b) reported that the net CH4 out-flux from the shelf in summer was more than five times higher than our results in early spring. The difference in freshwater discharge, which was more than three times higher in summer than in winter (Zhang et al., 2007), was the main reason for these seasonal CH4 changes. In addition, the net CH4 outflow in this study was 56% of what was previously reported in the same season (Zhang et al., 2004). The estimation of the net out-flux in this previous study did not consider TWCW, which is believed to be an important CH4 source according to our study. Thus, our data showed a reasonable CH4 budget in the shelf.

Fig.5 Preliminary CH4 budget in the shelf of the ECS The CH4 fluxes in box-model calculations are provided in Table 1.

The sedimentary release of CH4 was proven a potential source of dissolved CH4 in the ECS shelf (Fig. 3e) (Zhang et al., 2008a; Sun et al., 2018). The quantity of CH4 that was emitted from sediments showed great seasonal variations, presumably lower in winter and higher in summer, especially under hypoxia conditions off the Changjiang River estuary (Zhang et al., 2008a; Ye et al., 2016; Sun et al., 2018). The CH4 fluxes from sediment to the water in the ECS were documented in previous studies, such as 0.8 μmol/(m2·d) in April/May 2002 (Zhang et al., 2008a) and 1.7-2.2 μmol/(m2·d) in March 2011 (Sun et al., 2018). We used the average of these studies to represent the mean sediment-water CH4 flux in the ECS in spring since we did not measure the sediment-to-water flux in our study. Thus, the sediment-water CH4 flux was estimated as 8.19 mol/s based on the shelf area. This result indicates that the sediment in the ECS was a comparable source to the TWCW for the shelf CH4. From literatures, we found benthic CH4 fluxes varied spatially. Fox example, in the Baltic Sea, the fluxes ranged from 100 μmol/(m2·d) in coastal area to 26 000 μmol/(m2·d) in the inner eutrophic estuary (Sawicka and Brüchert, 2017). Another study reported the CH4 flux at the sediment-water interface was estimated 20.9-25.1 μmol/(m2·d) in the Godavari and Krishna estuaries (Rao and Sarma, 2016). Our results were comparable with that reported in the west coast of India (4.71 μmol/(m2·d)) (Araujo et al., 2018) and in the continental shelf of the Gulf of Cádiz (0.9-24 μmol/(m2·d)) (Ferrón et al., 2009). With respect to the CH4 emissions to the atmosphere, we used the mean value of 3.3 μmol/ (m2·d) that was calculated by Du et al. (unpublished data) for this cruise to represent the sea-air CH4 flux in the ECS in spring. Thus, the sea-air CH4 flux (19.41 mol/s) was one of the major CH4 sinks in the ECS shelf and was more than ~1.5 times higher than what was transported by water masses, suggesting that the ECS is a net source of atmospheric CH4. The CH4 oxidation rate (OR) can be measured by radiotracer techniques using tritiated CH4 according to the method described by Bussmann et al. (2017). Here, we used the turnover rate constant (k) in the North Sea, with a mean value of 0.058/d, reported by Osudar et al. (2015), to calculate the CH4 oxidation rate in the ECS shelf in our study because the environmental conditions in the North Sea were similar to that in the ECS. Thus, the CH4 oxidation rate was computed as 0.278 nmol/(L·d) when considering the mean CH4 concentration of 4.8 nmol/L in the ECS shelf (Table 1). Meanwhile, a comparable CH4 oxidation rate (0.265 nmol/(L·d)) was measured by a 13C-tracking method deployed in this survey. The consistency of CH4 oxidation rate from different methods indicates that the result we used in the CH4 budget is reasonable. Consequently, the Rcon in the shelf water was calculated as 868.75 mol/s by multiplying the shelf water volume by the CH4 oxidation rate. Taken together, a CH4 production rate of 881.01 mol/s was calculated after balance of other sources and sinks in spring (Fig. 5). Considerable CH4 should be produced in the water column to sustain the CH4 deficit. According to the shelf water volume, our mass balance indicated that a local CH4 production rate of 0.28 nmol/(L·d) was required to maintain the CH4 loss. This value is in the range of the reported CH4 production rates from Mexico Bay (-0.04- 2.00 nmol/(L·d)) (Bange et al., 1994) but much lower than what was reported for oligotrophic lakes (~50.00 nmol/(L·d)) (Grossart et al., 2011). Hence, our results showed that local production/consumption is the major source/sink (> 95%) of CH4 in the ECS shelf, following by a total of 2.2% CH4 diffused to the atmosphere. It should be noted that the estimation contains great uncertainties due to insufficient field data (e.g. gas fluxes at seep sites) and the source flux was presumably underestimated without consideration of the contribution of the gas hydrates. It is hard to quantify because rare data was published to constrain the size, area, and gas spreading velocity in various seeps. This means the local CH4 production rate was overestimated to some extent but would not affect the fact that the ECS was a hotspot for CH4 dynamics (production, consumption, and emission).

3.4 CH4 formation via DMSP/DMS degradation

Field observations suggested the presence of excess CH4 in the subsurface layers and the box model indicated the significant potential for CH4 production in the study area. As DMSP has been recently suggested as the possible substrate for CH4 production (Damm et al., 2015; Stawiarski et al., 2019), we measured the dissolved and particulate fractions (DMSP d and DMSPp) (data not shown) in the field survey. No correlations between DMSP/DMS and CH4 were found when all the stations in the ECS were considered. However, a negative correlation between DMS and ΔCH4 (R2=0.59, P<0.01) and positive correlation between DMSPd and ΔCH4 (R2=0.38, P<0.05) were found at offshore stations (water depth > 800 m) within the water column of 0-300 m (Fig. 4b). Further, to weaken the influence of multiple factors, the parameters at deep-water stations (> 800 m, stations P7, FJ7, E7, and F6) were analyzed by using PCA to determine potential correlations between the Carbon-Sulfur bonded compounds and CH4 distribution (Fig. 6). Weak correlations between CH4 and DMSP/DMS were found in the YS (Supplementary Tables S1-S2). In contrast, two factors were extracted in the ECS, explaining a cumulative variance in the data of 69% (Supplementary Tables S3-S4). Thus, the first factor was associated with the features of dissolved gases because the temperature (FL=0.94) and dissolved gas content (DO: FL=0.65, DMS: FL=0.89, CH4: FL=-0.54) were strongly loaded (Fig. 6b). The second factor, however, was significantly loaded by DMSPd (FL=-0.89), indicating that this factor was related to the cycling of methylated sulfur compounds. CH4 (FL=-0.74) was strongly loaded in the second factor, suggesting correlations between CH4 and DMSP d. Thus, we speculate that DMSPd and its degradation product (DMS) may serve as CH4 precursors in the study area, particularly in oligotrophic areas.

Fig.6 Loading plots that correspond to the first two factors following the VARIMAX rotation of the principal components a. in the YS; b. in the ECS. Dataset from the ECS is derived from deep stations (> 800 m).

To assess this hypothesis, incubation experiments were conducted in the laboratory as well as in the field. In our first experiment, spiking treatments with DMSPd significantly increased the CH4 concentration compared to the control groups (Fig. 7a, ANOVA, P<0.01 and Fig. 7c, ANOVA, P<0.05). The oxygen content in the glass bottles was constantly above 5.0 mg/L, illustrating the well-oxygenated conditions during the incubation period (Fig. 7b & d). Thus, these results showed that the amendment with DMSP induced CH4 production, and DMSP can act as a precursor for CH4 formation in oxic water. Furthermore, different N: P ratios were set in the second experiment to test the effect of N deficiency on DMSP-dependent CH 4 production, as indicated by Damm et al. (2009). In particular, the addition of DMSP increased the CH4 to ~15.0 nmol/L in less than a week, 1.5 times higher than the value of the control (Fig. 8a). The DO in the control was almost constant during the entire incubation, whereas the DO in the DMSP treatments rapidly depleted over the first three days and gradually decreased to approximately 5.0 mg/L after seven days of incubation, suggesting strong respiration (Fig. 8b). The effect of N deficiency on the microbial utilization of DMSP was ambiguous in our experiments, as reflected in the insignificant difference between the rates under N-rich treatment (5<N: P<16) and N-starved treatment (N: P<1) (Fig. 8a, ANOVA, P=0.65; Supplementary Fig.S1, ANOVA, P=0.89), suggesting that N-stressed conditions may not stimulate the transformation of DMSP to CH4 in the coastal waters. In the third experiment (seawater incubation with direct addition of DMS, see Section 2.4) CH4 production was observed in both the positive control (amended with C+N+P) and DMS-added treatment within the first three days, but no CH4 accumulations were observed during the last seven days in all the treatments (Fig. 8c). Meanwhile, the water in the incubation bottles never became hypoxic or anoxic during this experiment (Fig. 8d). The insignificant difference between the concentrations in the (C+N+P) treatment and DMS-added treatment suggested that the addition of DMS did not induce CH4 production (ANOVA, P=0.98).

Fig.7 CH4 production and DO consumption during DMSP-amended incubation experiments a & b. in November 2017; c & d. in January 2018. The error bars represent one standard deviation for the triplicate bottles.
Fig.8 CH4 and DO concentrations in surface waters incubated over 16:28:44 a & b. amended with DMSP; c & d. amended with DMS. The linear regression slopes of the CH4 concentration over time (R2=0.33, 0.89, 0.97, and 0.86 for control, N-depleted, N-stressed, and N-rich groups, respectively, P<0.01) represent the potential CH4 production rates (0.06, 0.89, 0.87, and 0.85 nmol/ (L·d) for control, N-depleted, N-stressed, and N-rich groups, respectively) in Fig. 8a. The error bars represent one standard deviation for the triplicate bottles.

CH4 significantly accumulated with the addition of DMSP in the first seven days compared to the control (Figs. 7 & 8a), indicating that the catabolism of DMSP can contribute to the dissolved CH4 in the presence of oxygen. The evidence for the formation of CH4 from DMSP/DMS remains circumstantial and is mainly based on correlations between the CH4 and DMSP/ DMS concentrations (Borges et al., 2018). Two mechanisms of DMSP degradation exist in marine systems. First, the cleavage of DMSP results in DMS and acrylate, which is called dissimilation (Kiene et al., 1986; Visscher et al., 1994). Second, DMSP is demethylated to 3-mercaptopropionate (MPA), with 3-methiolpropionate (MMPA) as a possible intermediate (Kiene and Taylor, 1988); MMPA can also degrade with demethylation, yielding methanethiol (MeSH) (Fig. 9) (van der Maarel and Hansen, 1997). Thus, CH 4 is known to be produced through three possible pathways: 1) Methanogenic archaea can decompose DMS to CH4 via demethylation if the DMS concentration is high enough that sulfate-reducing bacteria do not compete with methanogens (Kiene and Taylor, 1988); 2) Microbes can utilize MeSH (demethylation) to produce CH4 under anoxic environments (Tallant and Krzycki, 1997); 3) Acrylate can decompose into fatty acids (such as acetate and propionate, Fig. 9), which are the precursors for methanogenesis, so CH can form indirectly from acrylate degradation. These possible mechanisms for the transformation of DMSP to CH4 are thought to occur in the absence of oxygen. However, reducing conditions can form in micro-niches or inside bacterial cells (Jørgensen, 1977) despite the presence of oxygen in the ambient water. A theoretical model was recently built to explain the maintenance of anaerobic conditions for CH4 production inside bacterial cells in association with the transformation of DMSP (Damm et al., 2015). In addition, such organisms would metabolize MeSH in cells because they must rid themselves of MeSH before it accumulates to toxic levels (Kiene et al., 1986). Thus, CH4 may form inside cells and then be released into the oxygenated water via cell lysis (Reisch et al., 2011; Damm et al., 2015).

Fig.9 Schematic diagram for the conversion of DMSP and the connection with the production of CH4 Modified from van der Maarel and Hansen (1997).

One concern is that the transformation efficiency from DMSP to CH4 in our study was low; the addition of DMSP was 100-1 000 μmol/L and the CH4 concentration only increased by 2.0-5.0 nmol/L (Fig. 8a). We suppose that this result was associated with the mechanisms of DMSP degradation. Previous studies showed that the dissimilation pathway comprises 30% of DMSP metabolism (Kiene and Service, 1991; Kiene et al., 2000; Reisch et al., 2011; Damm et al., 2015) and only 6%-17% in surface seawater (Archer et al., 2002), which means that DMS is a minor product of DMSP metabolism. However, the first and second pathways of CH4 formation were directly connected to the degradation of DMS (Fig. 9). If limited DMS was produced from DMSP, then the yielding of CH4 would also have been limited. Moreover, over 80% of DMS in the marine system is known to be bacterially and photochemically oxidized to dimethylsulfoxide (DMSO) (Davis et al., 1998; Nowak et al., 2001), which further reduces the conversion efficiency from DMSP to CH4. In combination with the results of the second incubation, where CH4 did not obviously accumulate when the treatment was directly amended with DMS (Fig. 8c, ANOVA, P=0.98), we speculate that CH4 somehow has difficulty forming directly from DMS, at least in environments such as the YS and ECS. The exact mechanism is unknown; one possibility is that CH4-producing organisms prefer substrates other than DMS under environments where abundant carbon and sulfur sources are present (Lin et al., 2000, 2002). This observation is consistent with the negative correlations between DMS and ΔCH4 in the ECS (Fig. 4b), which means that the conversion of DMSP to CH4 was restricted if the DMSP preferred to dissimilate to DMS. Thus, DMS accumulated in the ambient seawater but CH4 production was limited because of the low efficiency of the DMS-to-CH4 process. On the other hand, DMS production would be limited if the DMSP mainly degraded via demethylation (to form MPA and acrylate), whereas CH4 would accumulate via the second and third pathways (Fig. 9). Consequently, in either case, we could expect to observe positive correlations between the accumulated DMSP and ΔCH4 and negative correlations between DMS and ΔCH4, as we did in the field observations (Fig. 4b). Another independent study that was conducted in the ECS also showed a positive correlation between DMSP and CH4 (Zhai et al., 2019). Thus, the CH4 in the study area may prefer to be produced via the demethylation of MeSH and/or degradation of acrylate rather than the decomposition of DMS.

Another concern is that DMSPd can be low in seawater (e.g., ~5.5 nmol/L in the YS and ECS in this study), and its contribution to the dissolved CH4 concentration can be negligible. However, studies showed that the intracellular concentration of DMSP could reach millimolar levels (Reisch et al., 2011). Thus, the potential importance of DMSP for CH4 production should be brought to the forefront if sulfur compounds transform in cells, as speculated above. According to the incubation experiments in the coastal waters, the DMSP-to-CH4 production rates were estimated as an average of 0.17 nmol/(L·d) (after subtracting the rate of 0.48 nmol/(L·d) in the control) with the addition of 100-μmol/L DMSP (Fig. 7a & c) and ~0.80 nmol/(L·d) with the addition of 1 000-μmol/L DMSP (Fig. 8a). Our results were lower than the recently reported production rate of ~1.8 nmol/(L·d) in the central ECS with the addition of 5-10-μmol/L DMSP (Zhai et al., 2019). Damm et al. (2009) suggested that a low N: P ratio in oligotrophic areas may enhance the microbial utilization of DMSP as a C source, with CH4 released as a by-product. However, our experiments conducted in coastal waters may weaken the effect of DMSP on CH4 concentrations, where abundant C sources (e.g., dissolved organic carbon: 3.0 mg/L) were present for microbial assimilation (Liu et al., 2015). Thus, DMSP may be out-competed by those substrates in coastal areas because the Gibbs free energy for DMSP-dependent CH4 production was low (-35.7 kJ/mol) (Damm et al., 2009). Consequently, even N deficiency could not trigger DMSP-dependent CH4 production in CW (Fig. 8a) because the low N: P ratio probably could not effectively stimulate the microbial utilization of DMSP in the presence of other competitive energy sources. In contrast, the experiment conducted in the oligotrophic surface seawaters of the western North Pacific (N: P<1) showed that directly addition of DMSP (10 μmol/L) induced significantly increase in CH4 (Fig. 10, ANOVA, P<0.01) and the CH4 production rate (1.2 nmol/(L·d)) is comparable to that reported in the central ECS (Zhai et al., 2019). This suggests that the DMSP-dependent CH4 production prefer to occur in the oligotrophic seawaters, where nitrogen is depleted.

Fig.10 Incubation experiment for seawaters directly amended with DMSP in the oligotrophic Western North Pacific a. CH4 production; b. oxygen consumption. The final concentration of DMSP was 10 μmol/L. The error bars represent one standard deviation for the triplicate bottles.

In addition, studies showed that the DMSPd concentration in the ECS can exceed 40.0 nmol/L in summer (Yang et al., 2011) and exceed 250.0 nmol/L during blooms (Matrai and Keller, 1993), which increases the possibility of a DMSP-dependent pathway for CH4 production in the water column. Moreover, additional CH4 sources (such as anaerobic methanogenesis in micro-niches and the degradation of Carbon-Phosphorus/Carbon-Nitrogen bonded compounds) alongside DMSP pathways should exist in the oxygenated water column to maintain the loss of CH 4 into the atmosphere. Recently, CH4 production from MPn was attributed to microbial activity in P-starved oceanic areas (Karl et al., 2008; Metcalf et al., 2012; Repeta et al., 2016). Our recent experiments in the YS also indicated the possibility of a Carbon-Phosphorus cleavage pathway for the enrichment of CH4 in oxic water (Ye et al., 2020). Hence, the CH4 dynamics in the YS and ECS are complicated, and CH4 production in well-oxygenated water may be attributed to several mechanisms. According to our field observations and incubation experiments, the Carbon-Sulfur cleavage pathway for CH4 production in well-oxygenated water is a potential mechanism for the production of excess CH4 in the western North Pacific. Although the exact molecular mechanism is unclear and no direct evidence exists of carbon transfer from DMSP to CH4, the implications for the biogeochemical cycling of carbon and sulfur in the marine system is essential to future work, especially when considering the effect of CH4 and DMS on global warming and atmospheric chemistry.

4 CONCLUSION

The CH4 concentrations in the ECS and YS showed great variations among different water masses, which profoundly affected the CH4 distribution. In spring, the vertical profiles of CH4 were homogenous in the YS but varied greatly in the ECS because of the latter's more complicated hydrological conditions. The surface and subsurface waters are oversaturated in CH4 indicated that the YS and ECS are the net sources of atmospheric CH4 in spring. A box model was used to evaluate the CH4 dynamics in the ECS. The results suggested that the coastal water plays an important role in distributing CH4 among the near-shore areas, while it was a minor source (0.01%) for CH4 in the continental shelf because of the low water discharge. Over 95% CH4 was considered to be produced and consumed in the shelf water, subsequently with 2.2% CH4 diffused to the atmosphere, suggesting ECS is a hotspot for CH4 dynamics. Further incubation experiments were carried out to figure out the source of this hotspot. Results show that DMSP can be a potential CH4 precursor in oxygenated coastal water and DMSP-dependent CH4 production has potential to be enhanced in oligotrophic areas.

This study demonstrates that the in-situ production could be a major source of dissolved CH4 in the marginal seas of China and its correlation with cleaving of Carbon-Sulfur bonded compounds (such as DMSP and MeSH) may have further implications for carbon and sulfur biogeochemical processes in the western Pacific. Further work is required to confirm the intrinsic connections between methylated sulfur compounds and CH4, presumably by molecular-biology techniques.

5 DATA AVAILABILITY STATEMENT

All data included in this study are available upon request by contact with the corresponding author.

6 ACKNOWLEDGMENT

The authors wish to thank the crews of the R/V Dong Fang Hong 2 and colleagues from the Laboratory of Marine Biogeochemistry at the Ocean University of China for their assistance with the collection of field samples. We also thank Haibing DING from the Ocean University of China for providing CH4 oxidation data.

7 CONFLICT OF INTEREST

The authors declare that they have no conflict of interest.

Electronic supplementary material

Supplementary material (Supplementary Tables S1-S4 and Fig.S1) is available in the online version of this article at https://doi.org/10.1007/s00343-021-1010-4.

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