2.1 Sampling and background weather during cruises

Total suspended particulate (TSP) samples were collected using a high-volume sampler (TE-5170D; Tisch Environmental, Inc.) with Whatman^®41 cellulose filters during two research cruises (Fig. 1) aboard the R/V Dongfanghong II. The first cruise (Fig. 1a) spanned from 17 March to 22 April 2014 (44 samples were collected in total; detailed sampling information can be found in Luo et al., 2016), and the second cruise (Fig. 1b) lasted from 30 March to 3 May 2015 (38 samples were collected in total; detailed sampling information, including the date, time period and location for each sample are listed in Table S1 in the Supplement). To avoid self-contamination from the research vessel, the TSP sampler was installed on the top of the tower at the ship head, and aerosols were sampled only during travel. More information about self-contamination from the ship exhaust can be found in Luo et al. (2016). Both cruises were undertaken during the East Asian monsoon transition period. The 2-month average (March and April) wind streamlines at 1000 hPa over the NWPO show that the wind speed ranged from 2 to 6 m s⁻¹ in 2014 (Fig. 1a) and from 1 to 3 m s⁻¹ in 2015 (Fig. 1b). In general, the wind was stronger in 2014 than in 2015 over the open ocean during the sampling periods.

Figure 1Regional wind streamlines (in m s⁻¹) at 1000 hPa during the Asian winter monsoon period (a, March and April in 2014 and b, March and April in 2015) based on an NCEP dataset. Cruise tracks are also shown (orange, pink and black indicate sea fog, dust and background aerosol, respectively). The aerosol number and collection range are shown in orange, pink and black for sea-fog modified, dust and background aerosol, respectively. The blue open circles in (b) indicate the locations of surface seawater sample (at a depth of 5 m) during the 2015 cruise. Hereafter, the Bgd. is the abbreviation of background aerosol in all the figures and tables.

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Meteorological data, including wind speed, direction, relative humidity (RH) and ambient temperature, are shown in Fig. S2 for the 2015 cruise (data for the 2014 cruise are reported in Luo et al., 2016). In the ECSs, both cruises encountered sea fog, which inevitably influenced aerosol sampling and aerosol chemistry. Because of the analogous weather conditions experienced during the two cruises, we used the techniques of Luo et al. (2016) to classify the 2015 marine aerosol samples into three types (namely, sea-fog-modified aerosol (orange triangles) collected in the ECSs, dust aerosol (pink circles) and background aerosol (black squares) sampled in the NWPO; Fig. 1) based on the meteorological conditions (Fig. S2), concentrations of aluminium (data not shown) and the lidar browse images from NASA (Fig. S3). Compared with the ECSs, which were strongly influenced by anthropogenic emissions, the NWPO (open ocean) was relatively clear. Hereafter, we define “background aerosol” as aerosol collected in the NWPO without influence from dust and sea fog during the investigation period.

To examine the relationships between the isotopic compositions of WSON in marine aerosol and DON in SSW, we collected SSW at a depth of 5 m (sampling locations are shown in Fig. 1b as open blue circles) using Niskin bottles during the 2015 cruise. The SSW samples were filtered using a 0.22 µm MILLEX^®-GP filter and kept frozen at −20 ^∘C in 50 mL 450 ^∘C pre-combusted brown glass tubes until analysis.

2.2 Chemical analyses

2.3 Data analysis

The concentrations of WSON in marine aerosol, which could not be measured directly (as mentioned previously), were calculated using the following equation:

where [WSON], [WSTN], [NO${}_{\mathrm{3}}^{-}$] and [NH${}_{\mathrm{4}}^{+}$] are the molar concentrations (nmol N m⁻³) of the given water-soluble nitrogen species in marine aerosol. The calculated WSON was subject to relatively large and variable uncertainties by error propagation since [NO${}_{\mathrm{3}}^{-}$] and [NH${}_{\mathrm{4}}^{+}$] were generally much higher than WSON concentrations in most observations. Such error propagation was a common and unavoidable problem (Mace and Duce, 2002; Cornell et al., 2003; Cape et al., 2011; Lesworth et al., 2010; Zamora et al., 2011). In previous studies, data points with high relative uncertainties were excluded (> 100 %, Lesworth et al., 2010; Zamora et al., 2011) and the negative values were taken as zero (Mace and Duce, 2002; Cornell et al., 2003; Violaki et al., 2015). Following previous studies, we excluded both the negative data and those with high relative uncertainties to reduce the uncertainty of mean WSON.

Reduced nitrogen (RN, i.e. NH${}_{\mathrm{4}}^{+}+$ WSON) and the δ¹⁵N-RN in the aerosol were calculated via mass balance:

where [WSTN] and [NO${}_{\mathrm{3}}^{-}$] are the molar concentrations (nmol N m⁻³) of the given water-soluble nitrogen species in the marine aerosol. The average propagated standard error for RN was 9 % for both 2014 and 2015, and the propagated error for the calculation of δ¹⁵N-RN was ±0.6 ‰.

Similar to aerosol WSON, the SSW DON concentration and δ¹⁵N-DON were calculated using the following equations:

where [TDN] and [NO${}_{\mathrm{3}}^{-}$] are the molar concentrations (µmol N L⁻¹) of the given species in SSW. The standard error propagated through the DON calculation was 5.3 %. Since the average [NH${}_{\mathrm{4}}^{+}$] in SSW at the selected sites during the 2015 cruise (12 sites and 23 samples) was 0.05 µM, which is much less than DON at µM level, [NH${}_{\mathrm{4}}^{+}$] was neglected in Eqs. (4) and (5). Unfortunately, most of the NO${}_{\mathrm{3}}^{-}$ concentrations in the SSW samples were < 0.5 µmol L⁻¹, which is too low for the measurement of δ¹⁵N-NO${}_{\mathrm{3}}^{-}$. We attempted to evaluate the interference from nitrate in the δ¹⁵N-DON calculations. For all the SSW samples on average, NO${}_{\mathrm{3}}^{-}$ comprised 5.7 % of the total NO${}_{\mathrm{3}}^{-}$ plus DON; the δ¹⁵N-NO${}_{\mathrm{3}}^{-}$ in SSW ranged from +8.2 to +16.4 ‰ in our measurements, and the bias of the calculated δ¹⁵N-DON varied from +0.5 to +0.9 ‰ .

The dry deposition nitrogen fluxes were calculated as follows:

where C_i is the concentration of a given water-soluble nitrogen species in the aerosol, and V_i is the given dry deposition velocity of the given nitrogen species. In addition to particle size, V_i was controlled by the meteorological conditions (wind speed and relative humidity; Duce et al., 1991; Hoppel et al., 2002) and underlying surface (smooth or rough; Piskunov, 2009). The V_i simulated by the model varied from 0.01 to 10 cm s⁻¹ under wind speeds from 5 to 30 m s⁻¹ and particle size from 0.1 to 100 µm (Hoppel et al., 2002). In our observations, wind speed ranged from 0.1 to 18 m s⁻¹, with relative humidity ranging from 40 to 100 % (for 2014 data, see Luo et al. (2016) and for 2015 data see Fig. S2), which was variable, thus prohibiting an accurate deposition estimate. Moreover, we did not obtain information about the size distribution. Previous studies showed that most of the NO${}_{\mathrm{3}}^{-}$ is distributed in supermicron size (ranging from 1 to 10 µm) in marine aerosol, with a small fraction in submicron size (ranging from 0.1 to 1 µm). However, NH${}_{\mathrm{4}}^{+}$ is mainly distributed in submicron size and only partly in supermicron size (Nakamura et al., 2005; Baker et al., 2010; Jung et al., 2013), except in coastal areas with mixed pollution and marine aerosols, where NH${}_{\mathrm{4}}^{+}$ was present in the coarse mode and NO${}_{\mathrm{3}}^{-}$ in the fine mode (Yeatman et al., 2001). In our case, the weather conditions, such as fog and dust, further affected the size distribution of aerosol N_r (Mori et al., 2003; Yao and Zhang, 2012; Hsu et al., 2014). Therefore, in this study, the deposition velocities of water-soluble nitrogen species were set to 2 cm s⁻¹ for NO${}_{\mathrm{3}}^{-}$, 0.1 cm s⁻¹ for NH${}_{\mathrm{4}}^{+}$ and 1.0 cm s⁻¹ for WSON, which have also been widely used to estimate the marine aerosol N_r dry deposition (Nakamura et al., 2005; Jung et al., 2013; Luo et al., 2016). However, bearing in mind, that any use of fixed deposition velocities to calculate the depositional flux of aerosol N_r may cause under- or over-estimation.

Figure 2Concentrations of aerosol WSTN (a, b), NH${}_{\mathrm{4}}^{+}$ (c, d), NO${}_{\mathrm{3}}^{-}$ (e, f) and WSON (g, h) with longitude for the 2014 and 2015 cruises. The orange open triangles denote sea-fog-modified aerosol in the ECSs, the pink circles denote dust aerosol, and the black open squares denote background aerosol in the NWPO. The error bars in (g) and (h) were the uncertainties of WSON caused by error propagation during the calculation.

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3.1 Spatial and temporal variations of water-soluble nitrogen species in the aerosol

Overall, significant logarithmic decreases occurred from the shore seaward for all water-soluble nitrogen species and WSTN in both 2014 and 2015 (Fig. 2). The seaward gradient was caused primarily by continental emissions influenced by sea fog (Luo et al., 2016); thus, concentrations were high in the ECSs (orange triangles in Fig. 2) and low offshore in the NWPO background aerosol (black squares in Fig. 2). Dust aerosol (pink circles in Fig. 2) appeared sporadically in the NWPO and generally featured higher NH${}_{\mathrm{4}}^{+}$ and NO${}_{\mathrm{3}}^{-}$ (but not WSON) values (Table 1 and Fig. 2).

The measured WSTN concentrations in TSP varied from 21 to 2411 nmol m⁻³ (Table 1 and Fig. 2a and b), lower than those in PM₁₀ sampled during spring in Xi'an, China (which ranged from 786 to 3000 nmol m⁻³; Wang et al., 2013), but higher than those in TSP sampled in Sapporo, Japan (which ranged from 20.9 to 108.6 nmol m⁻³; Pavuluri et al., 2015); Okinawa Island (which ranged from 5 to 216 nmol m⁻³; Kunwar and Kawamura, 2014); and the North Pacific (which ranged from 1.4 to 64.3 nmol m⁻³ in May–July; Hoque et al., 2015). This wide range of aerosol WSTN content illustrates the influence of the distance between sampling locations and emission sources (Matsumoto et al., 2014), seasonality (Kunwar and Kawamura, 2014) and meteorological conditions, such as sea fog (Luo et al., 2016).

Table 1Concentration ranges and means for WSTN, NH${}_{\mathrm{4}}^{+}$, NO${}_{\mathrm{3}}^{-}$, WSON and RN in aerosols.

^a Arithmetic mean. ^b Volume-weighted mean.

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The concentrations of marine aerosol WSTN in the ECSs ranged from 444 to 2411 nmol m⁻³ in 2014 (with a volume-weighted mean of 1136 nmol m⁻³) and from 92.9 to 1195 nmol m⁻³ in 2015 (with a volume-weighted mean of 287 nmol m⁻³), which were clearly higher than those in dust aerosol (with volume-weighted means of 242 nmol m⁻³ in 2014 and 154 nmol m⁻³ in 2015) and background aerosol (with volume-weighted means of 85.6 nmol m⁻³ in 2014 and 42.3 nmol m⁻³ in 2015) collected in the NWPO (Table 1). The air mass backward trajectories (see Fig. S5 for 2015 and Luo et al. (2016) for 2014) reveal that the high aerosol water-soluble nitrogen species in the ECSs arose from anthropogenic Nr emissions from eastern China (Gu et al., 2012). In addition, frequent formation of sea fog in the ECSs in spring (Zhang et al., 2009) may also enrich the amount of water-soluble nitrogen in sea-fog-modified aerosol via chemical processing (Luo et al., 2016). The much higher concentrations of water-soluble nitrogen species in the ECSs marine aerosol (compared to that in the NWPO aerosol) indicate that continental and/or anthropogenic N_r strongly affected the marine aerosol. The amounts of sea-salt ions (such as Na⁺) in the ECSs aerosols sampled in both 2014 (123 ± 98 nmol m⁻³; Luo et al., 2016) and 2015 (151 ± 164 nmol m⁻³; Luo et al., unpublished data) were higher than those in land aerosol sampled during spring (23 ± 7.8 nmol m⁻³ in Beijing; Zhang et al., 2013), which implies that those aerosols sampled in the ECSs were also significantly influenced by sea salt. Thus, we define the aerosol collected by ship over the ECSs as marine aerosol.

Figure 3Box plots for spring concentrations of aerosol NH${}_{\mathrm{4}}^{+}$ (a), NO${}_{\mathrm{3}}^{-}$ (b) and WSON (c) in the ECSs (the orange boxes denote sea-fog-modified aerosol) and NWPO (the pink boxes denote dust aerosol and the black boxes denote background aerosol), and summer aerosol (the blue box denotes NH${}_{\mathrm{4}}^{+}$ and NO${}_{\mathrm{3}}^{-}$ from Miyazaki et al. (2011) and Jung et al. (2013) and the WSON from Miyazaki et al., 2011). The large boxes represented the interquartile range from the 25th to 75th percentile, the line inside the box indicates the median value, and the whiskers extend upward to the 90th and downward to the 10th percentiles. Significant differences at the p < 0.05 level between different years are marked with coloured uppercase letters.

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In the NWPO, higher WSTN values were observed in dust aerosol than in background aerosol in both 2014 and 2015; the dust aerosol WSTN consisted predominantly of NH${}_{\mathrm{4}}^{+}$ and NO${}_{\mathrm{3}}^{-}$ rather than WSON (Table 1, pink circles in Fig. 2), which implies that dust can carry more NH${}_{\mathrm{4}}^{+}$ and NO${}_{\mathrm{3}}^{-}$ during long-range transport from East Asia to the NWPO during the Asian winter monsoon in spring. The air mass back trajectories of those dust aerosols arose mainly from high-N_r regions, as evidenced by dust plumes captured by lidar browse images from NASA (Fig. S3 for 2015 and Luo et al., 2016 for 2014).

Figure 4Aerosol δ¹⁵N-NO${}_{\mathrm{3}}^{-}$ (a, b) and δ¹⁵N-WSTN (c, d) in the ECSs (the orange open triangles denote for sea-fog-modified aerosol), and in the NWPO (the pink open circles denote dust aerosol and the black open squares denote background aerosol) with longitude during the 2014 and 2015 cruise, respectively.

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In our observations, the concentrations of NH${}_{\mathrm{4}}^{+}$ and NO${}_{\mathrm{3}}^{-}$ were higher in all aerosols in 2014 than they were in 2015 (Fig. 3a and b; Table 1). The difference between the two years was caused by a stronger Asian winter monsoon in 2014. Additionally, the cruise in 2014 (17 March to 22 April) occurred during a period of intensive fossil fuel combustion for heat supply in northern China; in contrast, the 2015 cruise started on 30 March and finished on 3 May, during a decrease in heating demand. The influence of heating on aerosol emissions can be seen in the atmospheric aerosol optical depth over heat-generating areas in China; for example, Xiao et al. (2015) reported that the aerosol optical depth was 5 times higher during heat generation periods than during non-generation periods. The consistent variations between heat supply in northern China and higher NO${}_{\mathrm{3}}^{-}$ in marine aerosol sampled in 2014 than 2015 underscore the influence of anthropogenic NO_x emissions on marine aerosol NO${}_{\mathrm{3}}^{-}$. Our spring observations in 2014 and 2015 showed average concentrations of NH${}_{\mathrm{4}}^{+}$ and NO${}_{\mathrm{3}}^{-}$ in background aerosol (black boxes in Fig. 3a and b; Table 1) higher than the average concentrations (3.5 ± 3.3 nmol m⁻³ for NH${}_{\mathrm{4}}^{+}$ and 2.1 ± 1.5 nmol m⁻³ for NO${}_{\mathrm{3}}^{-})$ reported in the western Pacific Ocean in summer (blue boxes in Fig. 3a and b, data from Miyazaki et al., 2011 and Jung et al., 2013), which suggests the far-reaching influence of anthropogenic emissions during the monsoon transition. Moreover, concentrations of NH${}_{\mathrm{4}}^{+}$ and NO${}_{\mathrm{3}}^{-}$ were higher in 2014 due to the stronger Asian winter monsoon, which further supports the idea that the monsoon exerts an important role in annual and seasonal variations in marine aerosol N_r via atmospheric long-range transport.

Table 2Ranges and means for stable nitrogen isotopes of WSTN, NO${}_{\mathrm{3}}^{-}$ and RN in aerosols.

^a Arithmetic mean. ^b Volume-weighted mean.

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Unlike NH${}_{\mathrm{4}}^{+}$ and NO${}_{\mathrm{3}}^{-}$, WSON concentrations in the background aerosol sampled in the NWPO in 2014 (average = 12.2 ± 6.3 nmol m⁻³) were similar to those in 2015 (average = 13.3 ± 8.5 nmol m⁻³; black boxes in Fig. 3c). In the open ocean, apart from terrestrial and anthropogenic WSON long-range transport (Mace et al., 2003; Lesworth et al., 2010), the ocean itself is the most likely source of marine aerosol WSON. For instance, in situ observations in the subtropical North Atlantic found that aerosol WSON had strong positive relationships with surface ocean primary productivity and wind speed (Altieri et al., 2016). Another study in the South Atlantic Ocean showed that the WSON in marine aerosol associated with high SSW chlorophyll a was 9 times higher than that associated with low SSW chlorophyll a (Violaki et al., 2015). In our observations, the WSON in the background aerosol (black bar in Fig. 3c) was significantly higher in spring than in summer (blue bar in Fig. 3c), which is consistent with the higher SSW chlorophyll a concentration over the NWPO in spring relative to that in summer (Fig. S6). However, the sources of marine aerosol WSON are a complex mixture composed of primary marine organic N and secondary N-containing organic aerosol. Biogenic organic material in SSW can be injected into the atmosphere to form an ice cloud via bubble bursting at the atmosphere–ocean interface (Wilson et al., 2015); this is probably the primary WSON aerosol source. Volatile organic compounds emitted from the surface ocean can react with NO_x and NH_x in the atmosphere to form secondary N-containing organic aerosol (Fischer et al., 2014; Liu et al., 2015).

Figure 5Scatter plots of δ¹⁵N-WSTN against the NH${}_{\mathrm{4}}^{+}$ ∕ WSTN ratio in (a) aerosol sampled in the ECSs, (b) dust aerosol and (c) background aerosol collected in the NWPO. Scatter plots of aerosol δ¹⁵N-WSTN against the WSON ∕ WSTN ratio in the (d) ECSs, (e) dust aerosol and (f) background aerosol in the NWPO. The solid and open symbols indicate aerosol sampled in 2014 and 2015, respectively.

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3.2 Isotopic composition of nitrogen species

Aerosol δ¹⁵N-NO${}_{\mathrm{3}}^{-}$ values over the ECSs and NWPO in 2014 and 2015 ranged from −9.2 to +10.2 ‰ (Table 2 and Fig. 4a and b). All the observed δ¹⁵N-NO${}_{\mathrm{3}}^{-}$ values fell within the ranges previously reported for atmospheric δ¹⁵N-NO${}_{\mathrm{3}}^{-}$ over land (Elliott et al., 2009; Fang et al., 2011; Felix and Elliott, 2014) and in the marine boundary layer (Hastings et al., 2003; Morin et al., 2009; Altieri et al., 2013; Gobel et al., 2013; Savarino et al., 2013). The mass-weighted mean aerosol δ¹⁵N-NO${}_{\mathrm{3}}^{-}$ values in 2014 (+1.6 ‰ in the ECSs, −1.1 ‰ for dust aerosol and −2.6 ‰ for background aerosol sampled in the NWPO) were similar to those in 2015 (+1.9 ‰ in the ECSs, −3.8 ‰ for dust aerosol and −1.1 ‰ for background aerosol sampled in the NWPO; Table 2) for all the aerosols, suggesting that the aerosol NO${}_{\mathrm{3}}^{-}$ arose from similar origins and atmospheric chemical pathways in both 2014 and 2015.

The δ¹⁵N-WSTN values for all the aerosols ranged from −10.7 to +5.6 ‰ (Table 2, Fig. 4c and d), which is consistent with the δ¹⁵N-WSTN ranges reported in precipitation, namely −4.2 to +12.3 ‰ in the Baltic Sea (Rolff et al., 2008); −8 to +8 ‰ in Bermuda (Knapp et al., 2010); −4.9 to +3.2 ‰ in a forest in southern China (Koba et al., 2012); and −12.1 to +2.9 ‰ in Cheju, Korea (Lee et al., 2012). In contrast, our results were lower than the δ¹⁵N-WSTN in TSP sampled in Sapporo, Japan (+12.2 to +39.1 ‰ ; Pavuluri et al., 2015) and the Sapporo Forest (+9.0 to +26.0 ‰; Miyazaki et al., 2014). These authors attributed higher isotopic values to biogenic sources, nitrogenous aerosol ageing and fossil fuel combustion. However, WSTN is, in fact, composed of various nitrogen species, and the relative proportions of NH${}_{\mathrm{4}}^{+}$, NO${}_{\mathrm{3}}^{-}$ and WSON to WSTN, coupled with their isotopic compositions (i.e. δ¹⁵N-NH${}_{\mathrm{4}}^{+}$, δ¹⁵N-NO${}_{\mathrm{3}}^{-}$ and δ¹⁵N-WSON), jointly mediate variations in aerosol δ¹⁵N-WSTN (where δ¹⁵N-WSTN ⋅ [WSTN] =δ¹⁵N-NO${}_{\mathrm{3}}^{-}$ ⋅ [NO${}_{\mathrm{3}}^{-}$] +δ¹⁵N-NH${}_{\mathrm{4}}^{+}$ ⋅ [NH${}_{\mathrm{4}}^{+}$] +δ¹⁵N-WSON ⋅ [WSON]). Taking our study as an example, the inconsistent trends in both the positive relationships between δ¹⁵N-WSTN and δ¹⁵N-NO${}_{\mathrm{3}}^{-}$ in the ECSs aerosols and NWPO background aerosols (Fig. S7a and c) and the negative relationship between δ¹⁵N-WSTN and the NO${}_{\mathrm{3}}^{-}$ concentration (Fig. S7d and f) imply that the δ¹⁵N of other species (NH${}_{\mathrm{4}}^{+}$ and WSON) in WSTN affected the δ¹⁵N-WSTN.

Figure 6(a) δ¹⁵N-DON (open circles) and δ¹⁵N-NO${}_{\mathrm{3}}^{-}$ (black squares), and (b) concentrations of DON (open circles) and NO${}_{\mathrm{3}}^{-}$ (black squares) in SSW with longitude during the 2015 cruise.

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The δ¹⁵N-WSTN values in 2014 (−10.7 to +1.0 ‰) were lower than those in 2015 (−5.6 to +5.6 ‰; Table 2), whereas the NH${}_{\mathrm{4}}^{+}$ ∕ WSTN ratios were higher in 2014 than in 2015 (Table 1) for all aerosol types. The negative linear relationships between NH${}_{\mathrm{4}}^{+}$ ∕ WSTN and δ¹⁵N-WSTN for all aerosol types (Fig. 5a–c) may be attributed to higher proportions of NH${}_{\mathrm{4}}^{+}$ in WSTN and negative δ¹⁵N-NH${}_{\mathrm{4}}^{+}$ values, which originated from the anthropogenic and marine emissions (Yeatman et al., 2001; Jickells, 2003; Altieri et al., 2014; Koba et al., 2012; Xiao et al., 2012; Liu et al., 2014). In fact, low δ¹⁵N-NH${}_{\mathrm{4}}^{+}$ values have been reported in precipitation in many places, such as Beijing (−33.0 to +14.0 ‰ with an arithmetic mean of −10.8 ‰; Liu et al., 2014); Guiyang City in south-western China (−38.0 to +5.0 ‰ with an average of −15.9 ‰; Xiao et al., 2012); Gwangju, Korea (−15.9 to +2.9 ‰ with volume-weighted means of −6.0 ‰ in 2007 and −6.8 ‰ in 2008; Lee et al., 2012); and a forest in southern China (−18.0 to +0.0 ‰ with a concentration-weighted mean of −7.7 ‰; Koba et al., 2012). Low atmospheric δ¹⁵N-NH${}_{\mathrm{4}}^{+}$ has also been associated with marine air masses (e.g. −8 to −5 ‰, Jickells et al. 2003; −9 ± 8 ‰, Yeatman et al., 2001; and −4.1 ± 2.6 ‰, Altieri et al., 2014). Together, this low atmospheric average δ¹⁵N-NH${}_{\mathrm{4}}^{+}$ (−15.9 to −4.1 ‰) supports our findings of higher NH${}_{\mathrm{4}}^{+}$ ∕ WSTN and lower δ¹⁵N-WSTN in aerosol.

There were positive linear relationships between the WSON ∕ WSTN ratio and δ¹⁵N-WSTN for all the aerosols types (Fig. 5d–f). This implies that the aerosol δ¹⁵N-WSON may be positive. The WSON in the marine aerosol either originated from terrestrial long-range transport or the DON from SSW and N-containing secondary organic marine aerosol as discussed in Sect. 3.1. Terrestrial aerosol δ¹⁵N-WSON was reported in a wide range (−15.0 to +14.7 ‰) with mean values from −3.7 to +5.0 ‰, and δ¹⁵N-WSON in marine aerosol with more positive δ¹⁵N (Fig. S1). In addition, our own observations for δ¹⁵N-DON in SSW showed positive δ¹⁵N in the ECSs (varying from +5.1 to +12.9 ‰, with an average of +7.9 ± 2.3 ‰) and NWPO (ranging from +1.9 to +11.6 ‰, with an average of +5.7 ± 2.0 ‰; Fig. 6a). At the same time, DON concentration in our observations ranged from 4.4 to 11.8 µmol L⁻¹ (Fig. 6b), which is within the DON concentration range reported in global SSW (Li et al., 2009; Van Engeland et al., 2010; Knapp et al., 2011; Letscher et al., 2013; Lønborg et al., 2015). This high DON concentration in SSW may be ejected into the atmosphere during bubble-bursting (Wilson et al., 2015).

To better clarify the sources of aerosol WSON, we removed the aerosol NO${}_{\mathrm{3}}^{-}$ and its δ¹⁵N effect from the WSTN and its δ¹⁵N-WSTN, respectively, by mass balance. The remaining NH${}_{\mathrm{4}}^{+}$ and WSON was defined as reduced nitrogen (RN $={\mathrm{NH}}_{\mathrm{4}}^{+}+$ WSON). The δ¹⁵N-RN ranged from −11.8 to +7.4 ‰ for all aerosol types in both 2014 and 2015 (Table 2), which is consistent with the reported δ¹⁵N-RN in precipitation (−12.6 to +7.8 ‰) collected in Bermuda, in the Atlantic Ocean (Knapp et al., 2010). The negative relationships between δ¹⁵N-RN and NH${}_{\mathrm{4}}^{+}$ concentration (Fig. S8a, S8b and S8c), as well as between δ¹⁵N-RN and the ratios of NH${}_{\mathrm{4}}^{+}$ ∕ RN (Fig. S8d–f) for all aerosol types, further support the low values of δ¹⁵N-NH${}_{\mathrm{4}}^{+}$ in marine aerosol. Three important endmembers, compiled from atmospheric δ¹⁵N-NH${}_{\mathrm{4}}^{+}$ (−15.9 to −4.1 ‰, green bars in Fig. 7) and continental δ¹⁵N-WSON (−3.7 to +5.0 ‰, grey bars) versus the δ¹⁵N-DON observed in SSW in our cruise (+7.9 ± 2.3 ‰ in the ECSs and +5.7 ± 2.0 ‰ in the NWPO, red dots with error bars), were added to Figure 7 to facilitate discussion. Note that nearly all the aerosol δ¹⁵N-RN values in 2014 were lower than those in 2015, which may be attributed to higher NH${}_{\mathrm{4}}^{+}$ concentrations in 2014 than in 2015 (Table 1). Moreover, higher values of δ¹⁵N-RN can be seen with higher WSON ∕ RN ratios for all aerosol types in Fig. 7. The high values of δ¹⁵N in continental WSON and marine DON in SSW may cause these positive relationships. Thus, although higher WSON ∕ RN values accompany higher δ¹⁵N-RN, we may still not conclude a significant DON in SSW contribution to aerosol WSON. In the open ocean, to some extent, background aerosol WSON was more likely influenced by DON in surface seawater, judging by nitrogen isotopic information.

Note that some data points collected in 2015 for the open-ocean case in Fig. 7b and c fell outside the mixing field, deviating toward higher δ¹⁵N-RN values; these high δ¹⁵N-RN values may be attributable to δ¹⁵N fractionation and ¹⁵N enrichment in the WSON during processes such as secondary N-containing organic aerosol formation by the reaction of NH_x or NO_x with organic aerosol (Fischer et al., 2014; Liu et al., 2015), complex atmospheric chemical reactions (i.e. the photolysis of organic nitrogen into ammonium; Paulot et al., 2015), the aerosol WSON ageing process and in-cloud scavenging (Altieri et al., 2016). More studies are needed to explore nitrogen transformation processes, especially those focusing on secondary N-containing organic aerosol in the atmosphere from an isotopic perspective.

Figure 7Scatter plots of aerosol δ¹⁵N-RN against the WSON ∕ RN ratio in the (a) ECSs, (b) dust aerosol in the NWPO and (c) background aerosol in the NWPO. The green bar indicates the sources of anthropogenic, terrestrial and oceanic δ¹⁵N-NH${}_{\mathrm{4}}^{+}$, the grey bar indicates the sources of terrestrial and anthropogenic δ¹⁵N-WSON, and the red bar indicates the δ¹⁵N-DON in SSW.

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Table 3Dry deposition fluxes of water-soluble nitrogen species.

^a Arithmetic mean. ^b Volume-weighted mean.

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Figure 8Dry deposition of aerosol NO${}_{\mathrm{3}}^{-}$-N (a, b), NH${}_{\mathrm{4}}^{+}$-N (c, d) and WSON-N (e, f) in the ECSs (orange open triangles) and in the NWPO (pink open circles for dust aerosol and black open squares for background aerosol) along longitude during the 2014 and 2015 cruise, respectively.

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3.3 N_r dry deposition and its biogeochemical role

The dry deposition of aerosol NH${}_{\mathrm{4}}^{+}$, NO${}_{\mathrm{3}}^{-}$ and WSON is summarized in Table 3. The calculated depositional fluxes of water-soluble nitrogen species in the ECSs were significantly higher than those in the NWPO (Fig. 8). The averaged dry depositional fluxes of NH${}_{\mathrm{4}}^{+}$ and NO${}_{\mathrm{3}}^{-}$ in 2014 were 2 to 5 times higher than those in 2015 for all aerosol types (Table 3). The dry depositional fluxes of NH${}_{\mathrm{4}}^{+}$ and NO${}_{\mathrm{3}}^{-}$ in dust aerosol were clearly higher than those in background aerosol in the NWPO (Table 3, Fig. 8a–d). Comparisons of these dry fluxes with other similar studies and estimations of the contribution of atmospheric N_r deposition to primary production in the NWPO are discussed in Luo et al. (2016), specifically for 2014; here, we focus on the influence of atmospheric N_r deposition on the nitrogen cycle in the ocean.

The influences of atmospheric N_r deposition on the marine nitrogen cycle are obvious over long timescales. For example, by analysing concentrations of NO${}_{\mathrm{3}}^{-}$ and phosphorus in seawater over the NWPO from 1980 to 2010, Kim et al. (2011) reported that the higher N ∕ P ratio in the upper ocean (in contrast to the deep ocean) in the NWPO was caused primarily by the accumulation of atmospheric anthropogenic N_r deposition. Another recent study found higher atmospheric anthropogenic N_r deposition to be associated with lower δ¹⁵N in surface sediment over the NWPO (Kim et al., 2017), the authors also posited that atmospheric anthropogenic N_r deposition can reach as far down as the deep ocean through biological action and lower the δ¹⁵N in surface sediment. The atmospheric δ¹⁵N values for water-soluble nitrogen species in our observations (Table 2) are lower than the δ¹⁵N-NO${}_{\mathrm{3}}^{-}$ in deep ocean water (+5.6 ‰; unpublished data from Kao); thus, it is possible that atmospheric δ¹⁵N-N_r can lower δ¹⁵N-NO${}_{\mathrm{3}}^{-}$ in the thermocline, as mentioned in previous studies (Knapp et al., 2010; Yang et al., 2014). However, it is hard to quantify the contribution of atmospheric δ¹⁵N-N_r to δ¹⁵N-NO${}_{\mathrm{3}}^{-}$ in the thermocline from the perspective of ¹⁵N for the following reasons: first, there are large spatial and temporal uncertainties in the dry and wet depositional fluxes of atmospheric N_r. For example, the dry depositional fluxes of NH${}_{\mathrm{4}}^{+}$ and NO${}_{\mathrm{3}}^{-}$ (23.1–43.3 µmol N m⁻² d⁻¹; Table 3) in our observations are significantly higher than those in summer (4.9 µmol N m⁻² d⁻¹; Jung et al., 2013). Moreover, according to a previous study, wet N_r deposition is 2–3 times higher than dry deposition (Jung et al., 2013), and N_r wet deposition in spring over the NWPO is unknown. Second, the δ¹⁵N of N fixation (−2 to 0 ‰; summarized by Knapp et al., 2010) is similar to the atmospheric δ¹⁵N-N_r (Table 2), but the fluxes of N fixation in the global ocean also vary considerably both spatially and temporally (Mulholland and Bernhardt, 2005; Needoba et al., 2007; Karl et al., 1997). Third, ¹⁵N fractionation occurs in the complicated marine nitrogen cycle (Knapp et al., 2005, 2010, 2011, 2012), which hampers the use of ¹⁵N in estimating the influence of atmospheric deposition N_r on marine δ¹⁵N under our limited current understanding.