Characteristics of methanesulfonic acid, non-sea-salt sulfate and organic carbon aerosols over the Amundsen Sea, Antarctica

To investigate the characteristics of particulate methanesulfonic acid (MSA(p)), non-sea-salt sulfate (nss SO2− 4 ) and organic carbon (OC) aerosols, aerosol and seawater samples were collected over the Southern Ocean (43–70 S) and the Amundsen Sea (70–75 S) during the ANA06B cruise conducted in the austral summer of 2016 aboard the Korean icebreaker IBR/V Araon. Over the Southern Ocean, the atmospheric MSA(p) concentration was low (0.10± 0.002 μg m−3), whereas its concentration increased sharply up to 0.57 μg m−3 in the Amundsen Sea where Phaeocystis antarctica (P. antarctica), a producer of dimethylsulfide (DMS), was the dominant phytoplankton species. Unlike MSA(p), the mean nss SO2− 4 concentration in the Amundsen Sea was comparable to that in the Southern Ocean. Water-soluble organic carbon (WSOC) concentrations over the Southern Ocean and the Amundsen Sea varied from 0.048 to 0.16 and 0.070 to 0.18 μgC m−3, with averages of 0.087±0.038 and 0.097±0.038 μgC m−3, respectively. For water-insoluble organic carbon (WIOC), its mean concentrations over the Southern Ocean and the Amundsen Sea were 0.25±0.13 and 0.26±0.10 μgC m−3, varying from 0.083 to 0.49 and 0.12 to 0.38 μgC m−3, respectively. WIOC was the dominant organic carbon species in both the Southern Ocean and the Amundsen Sea, accounting for 73 %– 75 % of the total aerosol organic carbon. WSOC/Na and WIOC/Na ratios in the fine-mode aerosol particles were higher, especially in the Amundsen Sea where biological productivity was much higher than the Southern Ocean. The fluorescence properties of water-soluble organic aerosols investigated using a fluorescence excitation–emission matrix coupled with parallel factor analysis (EEM–PARAFAC) revealed that protein-like components were dominant in our marine aerosol samples, representing 69 %–91 % of the total intensity. Protein-like components also showed a significant positive relationship with the relative biomass of diatoms; however, they were negatively correlated with the relative biomass of P. antarctica. These results suggest that the protein-like component is most likely produced as a result of biological processes of diatoms in the Amundsen Sea.

concentration in the marine atmosphere, and that ocean-derived OC is a significant component of the aerosol-cloud-climate feedback system involving marine biota.
Although the importance of marine biogenic source contribution to CCN concentration has motivated numerous studies, quantitative measurements of the size-dependent composition of marine aerosols over high southern latitudes, especially the Antarctic remain sparse, due to its inaccessibility. Because of the difficulty in conducting a field observation, the sources and 5 evolution of aerosols over the Antarctic are still a subject of many open questions (Giordano et al., 2017). It is, therefore, necessary to fill the data gap in the knowledge of marine aerosols in the Antarctic to improve understanding of the effect of ocean ecosystem on the marine aerosol-cloud-climate system.
Polynyas, recurring areas of seasonally open water surrounded by sea ice in high latitude regions, often exhibit high primary productivity (e.g., Arrigo et al., 2012;Yager et al., 2012;Hahm et al., 2014) because they are the 10 first polar marine systems to be exposed to the increasing springtime solar radiation (Arrigo and van Dijken, 2003;Criscitiello et al., 2013). Coastal polynyas surrounding Antarctica exhibit massive phytoplankton blooms during the austral summer (December-February), with most peaking in January. The productive polynyas are located in the Ross and Amundsen Seas while annual production in the polynyas of East Antarctica is generally low (Arrigo and van Dijken, 2003). The Amundsen Sea, located in West Antarctica, hosts two coastal polynyas: the Amundsen Sea Polynya (ASP) and Pine Island Bay Polynya 15 (PIBP) (Arrigo and van Dijken, 2003). The ASP is the most productive polynya (per unit area) among 37 identified coastal polynya systems in the Antarctic (Arrigo and van Dijken, 2003) due to the combined effects of the enhanced light condition (Park et al., 2017), the abundance of macronutrients, supply of iron (Fe) from melting sea ice and/or glaciers and continental shelf sediments resulted from the intrusion of relatively warm, salty, and nutrient-rich Circumpolar Deep Water (CDW) Dutrieux et al., 2014;Sherrell et al., 2015). Consequently, the Amundsen Sea is an ideal place to monitor a direct 20 link between biological production and local emissions of sulfur compounds and organics, not only because of its remoteness from anthropogenic activity but also because it is an area of exceptionally high seasonal primary production. However, little is known about the distributions of atmospheric sulfur and organic species in the Amundsen Sea.
To understand the influence of marine biological activities on atmospheric marine aerosols in the Amundsen Sea, we have investigated characteristics of MSA(p), nss-SO4 2-, water-soluble organic carbon (WSOC), and water-insoluble organic 25 carbon (WIOC) in marine aerosols collected over the Amundsen Sea, Antarctica. Besides, we have carried out a hydrographic survey to examine the link between biological production and atmospheric sulfur and organic species in the Amundsen Sea.
The objectives of this study are to (1) investigate the distributions of MSA(p) and nss-SO4 2over the Amundsen Sea, (2) examine the factors influencing atmospheric MSA(p) concentration in the Amundsen Sea, (3) investigate the distributions of atmospheric WSOC and WIOC over the Amundsen Sea, and (4) estimate the source of atmospheric water-soluble organic aerosols using a 30 fluorescence technique.
Aerosol and seawater samples were collected during the ANA06B cruise conducted over the Southern Ocean and the Amundsen Sea, Antarctica aboard the Korean icebreaker IBR/V Araon (Fig. 1). The cruise started from Christchurch, New Zealand on 6 January 2016, sailed over the Amundsen Sea for 23 days (14 January−5 February), and returned to Christchurch, New Zealand on 24 February 2016. In this study, the Southern Ocean and the Amundsen Sea are defined as the regions between 5 43°S and 70°S, and between 70°S and 75°S, respectively. Although the cruise track covered the Southern Ocean (43°S−70°S) and the Amundsen Sea (70°S−75°S), a significant portion of our cruise was allocated to the ASP and near-ice shelf surveys (< 2 km from ice shelves) adjacent to the Dotson, the Getz, and the Pine Island ice shelves.

Aerosol collection
Two high-volume aerosol samplers (HV-1000R, Sibata Scientific Technology Ltd.) were installed on the front of the 10 upper deck (20 m a.s.l.) and used to collect marine aerosols on pre-combusted (at 550°C for 6 hours) quartz filters (25 × 20 cm,Sibata Scientific Technology Ltd.). Particle size selectors for PM2.5 and PM10 were installed to each aerosol sampler to collect fine (D < 2.5 µm) and coarse modes (2.5 µm < D < 10 µm) aerosols on the filters, respectively. A windsector controller was used to avoid contamination from the ship's exhaust during the cruise (Jung et al., 2013(Jung et al., , 2014. The windsector controller was set to collect marine aerosol samples only when the relative wind directions were within plus or minus 15 to be < 5%. The concentrations of nss-SO4 2− were calculated as total SO4 2− minus Na + concentration times 0.2516, the SO4 2− /Na + mass ratio in bulk seawater (Millero and Sohn, 1992).

Water-soluble organic carbon
Other subsamples of the quartz filters were ultrasonically extracted using the same method for ionic species measurements. The filtrates were analyzed by a total organic carbon (TOC) analyzer (Model TOC-L, Shimadzu Inc.) for 5 determination of total dissolved organic carbon (TDOC), which was defined as water-soluble organic carbon (WSOC) in this study. In the analytical system, inorganic carbon was removed by acidifying the samples to pH 2 by 2 M HCl and sparging for 8 min before analysis of the total carbon content. Carbon dioxide (CO2) derived from the conversion of TDOC by high temperature (680°C) catalytic oxidation was measured by a nondispersive infrared (NDIR) detector to quantify TDOC (Miyazaki et al., 2011). Milli-Q water and consensus reference material (CRM, 42-45 µM C for DOC, deep Florida Strait 10 water obtained from University of Miami) were measured every sixth analysis to check the accuracy of the measurements. The procedural mean blank for WSOC was 180 µgC L −1 , which represented 14% of mean WSOC concentration in aerosols. The detection limit calculated as three times of the standard deviation of the procedural blanks was 21 µgC L −1 . The relative standard deviations of WSOC analysis for reproducibility test (at least three measurements per sample) was less than 3%.

Organic and elemental carbon 15
Concentrations of OC and elemental carbon (EC) were measured using the thermal optical transmission (TOT) method on a Sunset lab OC/EC analyzer (Model-4, USA) (Birch and Cary, 1996). The analytical procedures for OC and EC measurements are described in detail elsewhere (e.g., Miyazaki et al., 2011;Niu et al., 2012Niu et al., , 2013. In brief, a filter punch of 1.54 cm 2 was placed in the oven and heated in a completely pure helium environment up to 850°C to convert all OC into CO2. For EC measurement, the oven was cooled to 550°C and then heated until the oven temperature stepping up back to 850°C 20 under a 2% oxygen-containing helium environment. All CO2 derived from the conversion of both OC and EC was measured by an NDIR detector. The equivalent OC concentration from filed blank accounted for ~15% of the average OC concentrations of the actual samples. Based on field blank uncertainties, detection limits for OC and EC were 0.1 µgC cm −2 and 0.02 µgC cm −2 , respectively. In this study, WIOC was defined as the difference between OC and WSOC (i.e., WIOC = OC − WSOC) (Miyazaki et al., 2011). Using the propagating errors of each parameter, the uncertainty of WIOC concentration was estimated 25 to be 5.8%.

Optical measurements and excitation-emission matrix coupled with parallel factor analysis
Absorption spectra of atmospheric WSOC were obtained from 240 to 800 nm on a Shimadzu 1800 ultraviolet-visible (UV-Vis) spectrophotometer (Shimadzu Inc.). Three-dimensional fluorescence excitation-emission matrices (EEMs) were scanned using a Hitachi F-7000 luminescence spectrometer (Hitachi Inc.) at the excitation/emission (Ex/Em) wavelengths of spectra were used for inner filter correction for the EEMs according to McKnight et al. (2001). Further details on the EEM measurements and the procedure of post-acquisition corrections are available in previous studies (Chen et al., 2010(Chen et al., , 2017(Chen et al., , 2018. The procedure for Raman Unit (RU) normalization can be found elsewhere (Lawaetz and Stedmon, 2009). Parallel factor (PARAFAC) modeling was performed using MATLAB 7.0.4 with the DOMFluor toolbox. All corrected EEMs of aerosol samples were used for modeling. The number of fluorescent components was determined based on split-half validation. 5 The biological index (BIX), an index of recent biological and autochthonous contribution, was calculated according to Huguet et al. (2009) Seawater samples were collected from 4 to 5 layers in the upper 100 m at 46 stations in the Amundsen Sea using a conductivity-temperature-depth (CTD) and a rosette system holding 24-10L Niskin bottles (SeaBird Electronics, SBE 911 plus) ( Fig. 1). Seawater sample for chlorophyll-a (Chl-a) analysis was filtered through a GF/F filter (47 mm, Whatman), which was then extracted with 90% acetone for 24 hours. Chl-a was measured onboard using a fluorometer (Trilogy, Turner Designs, USA) (Lee et al., 2016). 15

Satellite observation
Monthly composite of sea surface Chl-a concentration (mg m -3 ) was obtained from Moderate Resolution Imaging Spectroradiometer (MODIS) Aqua data available from the Goddard Space Flight Center, NASA. The spatial resolution of the data was approximately 9 km per pixel (http://oceandata.sci.gsfc.nasa.gov).

Dissolved organic carbon 20
Seawater sample was drawn from the Niskin bottle by gravity filtration through an inline pre-combusted (at 550°C for 6 hours) Whatman GF/F filter held in an acid-cleaned (0.1 M HCl) polycarbonate 47 mm filter holder (PP-47, ADVANTEC).
The filter holder was attached directly to the Niskin bottle spigot. The filtrate was collected in an acid-cleaned glass bottle and then distributed into two pre-combusted 20 ml glass ampoules with a sterilized serological pipette. Each ampoule was sealed with a torch, quick-frozen, and preserved at -24°C until the analysis in our land laboratory. Analysis of dissolved organic 25 carbon (DOC) was performed by high-temperature combustion using a Shimadzu TOC-L analyzer. Milli-Q water (blank) and consensus reference material (CRM, 42-45 µM C for DOC, deep Florida Strait water obtained from University of Miami) were measured every sixth analysis to check the accuracy of the measurements. Analytical errors based on the standard deviations for replicated measurements (at least three measurements per sample) were within 5% for DOC.

Particulate organic carbon
For determination of particulate organic carbon (POC), seawater sample was drawn from the Niskin bottle into an amber polyethylene bottle. Known volumes (500 ml-1 L) of seawater were filtered onto pre-combusted Whatman GF/F filters (25 mm) under gentle vacuum at < 0.1 MPa. The filter samples were stored at -80°C until the analysis in our land laboratory.
Before POC analysis, the filter samples were freeze-dried and then exposed to HCl fumes for 24 hours in a desiccator to 5 remove inorganic carbon from the samples. Measurement of POC was carried out with a CHN elemental analyzer (vario MACRO cube, Elementar, Germany). Acetanilide was used as a standard. The precision of these measurements was ± 4%.

Phytoplankton taxonomic composition
Phytoplankton taxonomic composition was assessed using High Performance Liquid Chromatography (HPLC) analysis of accessory photosynthetic pigment concentrations. For the analysis of photosynthetic pigments, 1−2 L of seawater samples 10 were filtered onto 47 mm Whatman GF/F filters and stored at −80°C until the analysis in our land laboratory. The filters were extracted with 3 ml of 100% acetone, ultrasonicated for 30 seconds and maintained under 4°C in dark for 15 hours. Debris was removed by filtering through 0.45 µm Teflon syringe filters. Before injection, the extracts were diluted with distilled water (1 ml of extract plus 0.3 ml of distilled water) to avoid the peak distortion of the first eluting pigments. Pigments were assessed by HPLC analysis following the method of Zapata et al. (2000). Before analysis, the instrument (Agilent series 1200 15 chromatographic system, Germany) was calibrated using standard pigments (DHI, Denmark). A C8 column (250 mm × 4.6 mm, 5 µm particle size, Agilent XDB-C8, USA) was used for pigment separation. HPLC pigment data was processed using CHEMTAX (CHEMical TAXonomy), a matrix factorization program to calculate the absolute Chl-a biomass of major algal groups. Twelve pigments were chosen for CHEMTAX analysis, and seven pigment algal groups were defined according to Wright et al. (2010), including Phaeocystis antarctica (P. antarctica), diatoms, and cryptophytes. 20

Factors influencing atmospheric MSA(p) concentration over the Amundsen Sea
Various factors appear to influence atmospheric MSA(p) concentration in the Amundsen Sea. According to Arrigo and 15 van Dijken (2003), 5 years (1997-2002 averaged Chl-a level in the ASP during the month of January (6.98 ± 3.32 mg m -3 ) is more than 2 times higher than that in the Ross Sea Polynya (2.67 ± 0.82 mg m -3 ). Indeed, satellite ocean color images (http://oceancolor.sci.gsfc.nasa.gov) exhibited persistently high Chl-a levels in the ASP during the sampling period ( Fig. 3), implying a strong influence of biogenic source on atmospheric MSA(p) in the Amundsen Sea.
In addition to the high productivity, phytoplankton taxonomic assemblages could be a significant factor influencing 20 atmospheric MSA(p) abundance. In general, phytoplankton taxonomic abundances in the Amundsen Sea are dominated by haptophyte P. antarctica (e.g., Lee et al., 2016;Yager et al., 2016), which produces large amounts of DMS in Antarctic waters (Liss et al., 1994;Schoemann et al., 2005). During the cruise, P. antarctica was the dominant phytoplankton species in the upper 50 m, accounting for 42 ± 19% of phytoplankton biomass (Chl-a), with lesser abundances of diatoms (39 ± 17%) found throughout the polynya and sea ice zone (see Supplement, Fig. S1). In addition, extremely high concentrations (> 150 nM) and 25 fluxes (85 ± 119 µmol m -2 d -1 ) of DMS have been observed in the region where phytoplankton assemblages were dominated by P. antarctica during the cruise (Kim et al., 2017), which are consistent with previous results observed in the Amundsen Sea (Tortell et al., 2012).
We investigated the relationships between atmospheric MSA(p) concentration and the variables mentioned above, to examine factors influencing atmospheric MSA(p) concentration in the Amundsen Sea. Atmospheric MSA(p) concentration 30 showed no relationship with either in situ sea surface Chl-a concentration (r = 0.029, p > 0.05) (Fig. S2) or the relative biomass of P. antarctica (r = 0.30, p > 0.05). This suggests that Chl-a concentration and phytoplankton taxonomic composition are not direct factors determining atmospheric MSA(p) concentration over the Amundsen Sea.
To investigate the relationship between atmospheric MSA(p) and DMS flux, we used the DMS flux data reported by Kim et al. (2017), who calculated sea-air DMS fluxes using sea surface DMS concentrations and shipboard wind speed data monitored during our cruise. Details of the measurement of sea surface DMS concentration and the sea-air DMS flux calculation are given by Kim et al. (2017). The DMS flux (reported by Kim et al., 2017) averaged for the duration of each aerosol sampling showed a somewhat similar variation trend to that of atmospheric MSA(p) concentration (Fig. 4a), but no 5 correlation was found between atmospheric MSA(p) and DMS flux (r = 0.18, p > 0.05, Fig. 4b). DMS fluxes typically rely on gas transfer velocity (k), which is frequently parameterized as a function of wind speed (Wanninkhof, 2014). Measurement and parameterization of the gas transfer velocity are more challenging and subject to greater uncertainty, particularly at high wind speeds (Smith et al., 2018). Wanninkhof et al. (2014) reported that there is a considerable uncertainty in k, especially, under the strong wind condition. About 20% uncertainty was estimated at a global mean wind speed (7.3 m s -1 ). When we 10 applied four different k values (the units of k are in cm hr -1 ) calculated from the equations suggested by Wanninkhof, 1992, Wanninkhof and McGillis, 1999, Nightingale et al. 2000, and Wanninkhof, 2014 water, and where both low temperatures and high winds are typical (McGillis et al., 2000). In addition to the uncertainty in DMS flux, the insignificant relationship between atmospheric MSA(p) and DMS flux could result from various complexities in the rate of oxidation of DMS to form atmospheric MSA(p) and long-range transport of atmospheric MSA(p) from biogenically active region, given the lifetime of DMS is approximately 1-2 days (Kloster et al., 2006;Read et al., 2008). Although we found no significant relationship between atmospheric MSA(p) concentration, in situ sea surface Chl-a concentration, the 20 relative biomass of P. antarctica and the local sea-air DMS flux, the higher atmospheric MSA(p) concentrations observed over the Amundsen Sea compared to those over the Southern Ocean and in coastal Antarctic regions most likely resulted from complex linkage between these factors.

Atmospheric nss-SO4 2over the Southern Ocean and the Amundsen Sea
Concentration of nss-SO4 2in bulk (fine + coarse) aerosols during the cruise ranged from 0.30-0.87 µg m -3 , with ~79% 25 (median values for all data) of nss-SO4 2being present on fine mode aerosols (Fig. 2b). Mean concentration of nss-SO4 2during the cruise was 0.61 ± 0.17 µg m -3 . Over the Southern Ocean (43°S−70°S, samples A1-A2 and A11-A14), the nss-SO4 2concentration ranged from 0.30-0.87 µg m -3 (mean: 0.60 ± 0.23 µg m -3 ), whereas its concentration over the Amundsen Sea respectively. In addition, the mean nss-SO4 2concentration in the Amundsen Sea was also a factor of 1.6-4.4 higher than those observed at Palmer Station (0.24 ± 0.16 µg m -3 , December 1990-March 1991, Savoie et al., 1993), Halley Station (0.14 ± 0.017 µg m -3 , monthly mean in January from 1991-1993, Legrand and Pasteur, 1998), Neumayer Station (0.38 ± 0.13 µg m -3 , monthly mean in January from 1983-1995, Minikin et al., 1998, and Dumont D'Urville Station (0.34 ± 0.039 µg m -3 , monthly mean in January from -1995, Minikin et al., 1998 during the austral summer. Unlike MSA(p), the mean nss-SO4 2concentration in the Amundsen Sea was comparable to that in the Southern Ocean, 5 although the variation trend of nss-SO4 2in the Amundsen Sea was similar to that of MSA(p), suggesting that nss-SO4 2was affected by marine source and large-scale transport (Korhonen et al., 2008). Indeed, nss-SO4 2showed a strong correlation (r = 0.98, p < 0.01) with MSA(p) in the Amundsen Sea, whereas no relationship was found between them in the Southern Ocean (r = 0.51, p > 0.05) (Fig. S3), suggesting that the local emission of DMS is a significant source of nss-SO4 2in the Amundsen Sea. It is worth mentioning that nss-SO4 2can be formed from the homogeneous nucleation of new particles involving H2SO4 10 or from the condensation of gas-phase DMS oxidation products onto existing particles (e.g., Covert et al., 1992;Quinn and Bates, 2011). Recently, Sanchez et al. (2018) identified two types of SO4 2particles using an Event-Trigger Aerosol-Mass-Spectrometer (ET-AMS), and reported that 63% of SO4 2was derived from newly formed particles in the free troposphere and 38% of SO4 2was formed from the condensation of DMS products onto existing particles in the clean marine conditions, revealing the importance of phytoplankton-produced DMS emission for CCN in the Atlantic. In this study, it is hard to 15 distinguish nss-SO4 2derived from new particle formation from nss-SO4 2formed by the condensation of DMS products onto existing particles because of the limitations related to the method to collect data. However, the significant correlation of nss-  et al. (2018) reported that the lack of correlation between SO4 2particle and atmospheric DMS (or its oxidation products) could result from the competition for DMS and its oxidation products with the competing sinks of condensation onto existing particles and vertical transport to the free troposphere. The lack of correlation between nss-SO4 2and MSA(p) in the Southern 25 Ocean, therefore, could have resulted from the input of nss-SO4 2from the free troposphere. Although our data set is not sufficiently complete to allow a meaningful analysis of this likely explanation, the result for nss-SO4 2concentrations from this study would be valuable for filling the data gap, especially for the Amundsen Sea during the austral summer, and be helpful for validation of modeling of sulfur-containing aerosols.
During the cruise, the MSA(p)/nss-SO4 2ratio in bulk aerosols varied from 0.12 to 0.70 (mean: 0.35 ± 0.17), with lower ratios in marine aerosols collected over the Southern Ocean (range: 0.12-0.51, mean: 0.26 ± 0.14) and higher values over the Amundsen Sea (range: 0.20-0.70, mean: 0.44 ± 0.16), showing a similar variation trend (r = 0.92, p < 0.01) to that of MSA(p) 5 (Fig. 5). This result suggests that atmospheric MSA(p) plays a key role in the variation in MSA(p)/nss-SO4 2ratio over the Southern Ocean and the Amundsen Sea during the austral summer since atmospheric nss-SO4 2concentrations in the Southern Ocean were quite comparable to the values in the Amundsen Sea.

Atmospheric WSOC and WIOC over the Southern Ocean and the Amundsen Sea
Concentration of WSOC in bulk (fine + coarse) aerosols during the cruise ranged from 0.048-0.18 µgC m -3 , with an 10 average of 0.092 ± 0.037 µgC m -3 (Fig. 6a) (Fig. 6b). We expected much higher WSOC and WIOC concentrations in the 15 Amundsen Sea than the Southern Ocean because of extremely high Chl-a concentration in the Amundsen Sea (Fig. 3). However, no significant differences of mean WSOC and WIOC concentrations were found between the Southern Ocean and the Amundsen Sea, suggesting that Chl-a concentration is not a direct factor controlling atmospheric OC concentration in our study area (Quinn et al., 2014), although a significant correlation between atmospheric OC and Chl-a concentrations was observed in the Austral Ocean (Amsterdam Island, 37°48′S, 77°34′E, Sciare et al., 2009). 20 Both WSOC and WIOC mainly existed in fine mode particles, and the percentages of WSOC and WIOC present in fine aerosol particles were ~93% and ~74%, respectively (median value for all data). During the cruise, WIOC was the dominant OC species in both the Southern Ocean and the Amundsen Sea, accounting for 75% and 73% of total aerosol organic carbon, respectively (Fig. 6c). These results were consistent with the previous studies. North Atlantic. Moreover, Facchini et al. (2008a) reported that OC observed through a bubble bursting experiment and a field measurement (at Mace Head) was mainly water-insoluble, accounting for 77 ± 5% of the primary marine aerosol fraction in the submicron size range, and that WIOC consisted of colloids and aggregates exuded by phytoplankton. Given that atmospheric WIOC is mechanically produced through bubble bursting processes from hydrophobic organic matter that accumulates in the ocean surface (Facchini et al., 2008a;Gantt et al., 2011;Miyazaki et al., 2011), the dominance of WIOC 30 suggests that water-insoluble organic matter exuded by phytoplankton is more accumulated in sea surface water and emitted into the marine atmosphere via bubble bursting and breaking waves by local wind.
Our mean values of WSOC and WIOC in the Amundsen Sea were comparable to the results by Sciare et al. (2009) who reported that the mean concentrations of WSOC and WIOC observed at Amsterdam Island (37°48′S, 77°34′E) during the austral summer (January) were 0.083 ± 0.028 µgC m -3 and 0.19 ± 0.062 µgC m -3 , respectively. It is worth noting that atmospheric WSOC and WIOC show seasonal variations, with maximum values during austral summer (January) and minimum concentrations during winter (Sciare et al., 2009). These variations are attributable to pronounced seasonal variations 5 in biogenic marine productivity. Given that our study was carried out during the austral summer, the concentrations of WIOC and WSOC from this study would be considered as maximum values in the Amundsen Sea. Although our results were supported by previous studies as mentioned above, the spatial variability in WSOC and WIOC concentrations has been observed over various oceanic regions by previous studies. For instance, Fu et al. (2015) reported that atmospheric OC species concentrations observed at Alert, in the Canadian High Arctic were 0.186 µgC m -3 (range: 0.041-0.30 µgC m -3 ) for WSOC 10 and 0.068 µgC m -3 (range: 0.022-0.12 µgC m -3 ) for WIOC. The results by Fu et al. (2015) were 1.9 times higher and 3.8 times lower than our mean concentrations of WSOC and WIOC in the Amundsen Sea, respectively. In addition, Miyazaki et al. (2011) reported that mean concentrations of WSOC and WIOC observed over the western North Pacific (40°N-44°N) were 0.65 ± 0.27 µgC m -3 and 0.56 ± 0.19 µgC m -3 , which were a factor of 6.7 and 2.2 higher than those in the Amundsen Sea, respectively.
These differences in WSOC and WIOC concentrations among the oceanic regions presumably reflect regional differences in 15 factors influencing atmospheric WSOC and WIOC concentrations, such as source strength for volatile organic compounds emitted from biogenic sources (BVOCs), atmospheric oxidative capacity (e.g., OH, NO3 and ozone), meteorological condition (e.g., wind speed), DOC and POC concentrations in sea surface water, atmospheric transport and removal (e.g., Facchini et al., 1999;Kanakidou et al., 2005;Sun et al., 2006).

WIOC/Na + and WSOC/Na + ratios and relationships of WIOC and WSOC with Na + over the Southern Ocean and 20 the Amundsen Sea
Breaking waves on the ocean surface generate air bubbles that scavenge organic matter from seawater. When injected into the atmosphere, these bubbles burst, yielding sea spray aerosols enriched in organic matter, relative to seawater (Quinn et al., 2014). Sea spray aerosols have been defined as the hydrated droplets encapsulating dissolved sea-salt and entrained organic matter . Previous studies have revealed that organic matter is enriched in sea spray aerosol particles 25 produced by bubble bursting processes in the fine and ultrafine aerosol size fractions, suggesting that sea spray aerosol particles have an important role in transferring organic matter from the sea surface to the atmosphere (Facchini et al., 2008a;O'Dowd et al., 2008). During the cruise, ~76% of Na + , a tracer of sea spray, was associated with the coarse mode particle (Fig. 7a).
Moreover, statistically significant relationships were found between mean wind speed and Na + concentrations in fine (r = 0.54, p < 0.05) and coarse (r = 0.69, p < 0.01) mode aerosols, reflecting that Na + was formed from bubble bursting by local wind. 30 Although Na + was predominantly associated with the coarse mode particle, WSOC/Na + and WIOC/Na + in the fine mode aerosol particles were higher than those in the coarse mode aerosol particles, especially in the Amundsen Sea where biological productivity was much higher than the Southern Ocean (Fig. 7b). In the Southern Ocean, the WSOC/Na + ratio in the fine mode particles varied from 0.045-0.40, whereas the WSOC/Na + ratio in the Amundsen Sea ranged from 0.17-0.89. The average WSOC/Na + ratio in the Amundsen Sea (0.40 ± 0.24) was substantially higher than that in the Southern Ocean (0.16 ± 0.12).
For the WIOC/Na + ratio in the fine mode particles, similar results were observed. The WIOC/Na + ratios in the Southern Ocean and the Amundsen Sea varied from 0.038-0.97 and 0.26-2.4, with averages of 0.35 ± 0.31 and 0.91 ± 0.73, respectively; however, WIOC/Na + ratio in the fine mode aerosol particles was much higher than WSOC/Na + , suggesting that bubble bursting 5 at the ocean surface is a major source of atmospheric WIOC and that WIOC is more accumulated in the sea surface water.
These results show that the higher marine biological activities in the Amundsen Sea can be a significant factor leading to the higher OC/Na + ratios, indicating a linkage between OC emission and biological activities (Fig. 3).
Because sea spray aerosols are emitted as a result of wind-driven bubble bursting, correlations of organic aerosol mass concentrations with sea spray aerosols (i.e., Na + ), whose concentration are related to local wind speeds, have been used to 10 attribute their origin to marine sources because submicron Na + is known to form primary aerosols from evaporating seawater droplets (Russell et al., 2010). In this study, we investigated relationships of WIOC and WSOC concentrations with Na + concentration in the fine modes since both WIOC and WSOC were primarily associated with the fine mode particles (Figs. 6a and 6b). The submicron WIOC showed no statistically significant relationships with submicron Na + over the Southern Ocean and the Amundsen Sea (Fig. 8a), although WIOC/Na + ratio in the fine mode aerosol particles was much higher (Fig. 7b). 15 Similarly, Boreddy et al. (2018) found no correlation between sea-salt and WIOC in the western North Pacific. The increase in sea-salt particle flux under higher wind speed conditions shifts the sea spray aerosol size distribution towards larger sizes and accelerates their dry deposition and gravitational settling from the atmosphere (de Leeuw et al., 2011). Thus, theses insignificant relationships between WIOC and Na + in the fine modes could result from the differences in local wind speeds and local biological activities, because wind speed, a key factor determining sea spray aerosols, controls the local flux rather 20 than local concentration of marine particles (Monahan and O'Muircheartaigh, 1986). Our results are further supported by the study of Ceburnis et al. (2008), who found WIOC and sea-salt exhibited upward fluxes observed through gradient flux measurements, suggesting a primary production mechanism for WIOC. Although we found no significant relationships between WIOC and Na + in the fine modes, the high WIOC/Na + ratio in the fine mode aerosol particles (Fig. 7b) indicates that WIOC was predominantly of primary origin . However, the WIOC production by secondary processes 25 cannot be completely excluded either (Ceburnis et al., 2016), but we have no evidence of that.
Unlike the relationship between WIOC and Na + , submicron WSOC showed a strong correlation (r = 0.94, p < 0.01) with submicron Na + in the Amundsen Sea (Fig. 8b). In addition, we also found a significant correlation (r = 0.93, p < 0.01) between WSOC and MSA(p) concentrations in the Amundsen Sea (Fig. 8c). However, in the Southern Ocean, WSOC showed no significant relationship with submicron Na + or MSA(p). MSA(p) is produced by atmospheric oxidation of DMS, which is 30 released as a gas phase from marine biological activities and thus can be used as an indicator of secondary aerosols of marine biogenic origin (Miyazaki et al., 2011). Consequently, the strong correlations between WSOC, Na + and MSA(p) in the Amundsen Sea implies that the Amundsen Sea that has the most productive polynya in the Antarctic is a strong source region of BVOCs, and that WSOC was formed by the condensation of BVOCs released from sea surface onto preexisting submicron sea spray aerosols through gas-to-particle conversion due to a higher surface-to-volume ratio of submicron aerosols (Romakkaniemi et al., 2011). On the other hand, the poor correlations between WSOC, Na + , and MSA(p) in the Southern Ocean implies the differences in local source strength of BVOCs and that the presence of DMS in seawater and its subsequent oxidation to MSA(p) were not necessarily linked to the formation of submicron WSOC over the Southern Ocean (Miyazaki et al., 2016). 5

Fluorescence properties of water-soluble organic aerosols over the Southern Ocean and the Amundsen Sea
Fluorescence excitation-emission matrix coupled with parallel factor analysis (EEM-PARAFAC) has been widely used to investigate the sources and optical properties of dissolved organic matter in terrestrial and oceanic systems (e.g., Coble, 1996Coble, , 2007Stedmon et al., 2003;Yamashita et al., 2011;Retelletti Brogi et al., 2018). Moreover, recent field studies demonstrated that EEM-PARAFAC provides useful information for characterizing atmospheric OC in aerosols and rainwater 10 (e.g., Fu et al., 2015;Miyazaki et al., 2018;Yang et al., 2019).
As described in section 3.6, our results strongly suggested that the submicron WSOC observed in the Amundsen Sea was formed by the condensation of BVOCs onto preexisting submicron sea spray aerosols by showing the strong correlations with Na + and MSA(p). To further elucidate the sources of water-soluble organic aerosols, we investigated the fluorescence properties of submicron aerosols using EEM-PARAFAC. Fluorophores in water-soluble organic aerosols were divided into 15 three primary types on the basis of their peak position (Fig. 9). The spectral features of component 1 (C1, Ex/Em: 300/344 nm) was similar to the component previously identified in in coastal and oceanic waters as well as the polar ocean and was thought to be phytoplankton-derived (or ice algae-derived) protein-like component (Stedmon et al., 2007, Retelletti Brogi et al., 2018. Component 2 (C2, Ex/Em: 276/326 nm) was assigned as a tryptophan-like fluorophore, which has been considered to be a labile component produced as a result of biological production in marine environments (Coble et al., 1998;Yamashita 20 and Tanoue 2003). Component 3 (C3, Ex/Em: <260/458 nm) spectra were characterized as representing terrestrial humic-like fluorophores (Coble et al., 1996;Yamashita and Tanoue, 2003;Chen et al., 2010;Chen et al., 2018). During the cruise, the C1 fluorescence intensity was much higher and variable, varying between 0.0133 and 0.139 RU (Fig. 10a). In comparison, the fluorescence intensity of C2 (range: 0.0195-0.0518 RU) and C3 (range: 0.00857-0.0351 RU) was much less variable. Among the three components, the protein-like C1 was the dominant fluorescence component in our marine aerosol samples, accounting 25 for 31-73% of the total intensity, and the relative contributions of tryptophan-like C2 and terrestrial humic-like C3 were found to represent 17-50% and 8-31%, respectively (Fig. 10b). In our marine aerosol samples, protein-like components (i.e., C1 and C2) represented 69-91% of the total intensity. Despite the extremely high Chl-a concentration in the Amundsen Sea (Fig. 3), we found no significant difference of the average values of protein-like C1 and tryptophan-like C2 fluorescence intensity between the Amundsen Sea and the Southern 30 Ocean. However, relatively much higher values of C1 and C2 fluorescence intensity were observed when the ship approached the Amundsen Sea (i.e., samples A3 and A4), passing through the sea ice zone (Figs. 1b and 10a). The C1 and C2 fluorescence intensity values sharply decreased and remained relatively low in the Amundsen Sea, and then gradually increased from the the Amundsen Sea, whereas diatoms formed a major group in the marginal sea ice zone (Lee et al., 2016).
Ice algae, commonly found in polar sea ice and surrounding waters, are largely dominated by diatoms (Roberts et al., 2007), which are an important contributor to aerosols by emission of aerosol-forming volatile (e.g., alkyl-amines) and nonvolatile (e.g., mycosporine-like amino acids) organic nitrogen in the Antarctic sea ice region (Dall'Osto et al., 2017). Previous 5 studies (e.g., Facchini et al., 2008b;Miyazaki et al., 2011) provided the evidence that volatile emissions of alkyl-amine from marine algae can represent an important source of marine secondary organic aerosol. Moreover, Dall'Osto et al. (2017) observed that the fluorescence signal for protein-like component was positively correlated to organic nitrogen originated from the melted Antarctic sea ice floes, indicating that protein-like component was associated with organic nitrogen derived from the microbiota of sea ice and sea ice-influenced ocean. Although we have no aerosol water-soluble organic nitrogen dataset, 10 our results provide additional evidence that marine algae can influence fluorescent property of marine aerosols. Interestingly, we found that fluorescence intensity of C1 showed a significant positive relationship with the relative biomass of diatoms (r = 0.89, p < 0.01); however, it was negatively correlated with the relative biomass of P. antarctica (r = -0.79, p < 0.05) (Figs. 10c and 10d). Given the dominance of diatoms in the marginal sea ice zone during the cruise and the significant positive relationship between fluorescence intensity of protein-like C1 and the relative biomass of diatoms, it is plausible, therefore, 15 that biological processes of diatoms are an important factor in controlling the abundance of protein-like component in watersoluble organic aerosols over the Southern Ocean and the Amundsen Sea, although further studies are necessary to clarify this point.
The biological index (BIX) has been used to estimate the contribution of autochthonous biological activity (Fu et al., 2015;Miyazaki et al., 2018). An increase in BIX is related to an increase in the contribution of microbially derived organics. 20 High BIX values (> 1) have been shown to correspond to a predominantly biological or microbial origin of dissolved organic matter and to the presence of organic matter freshly released into water, whereas low values (< 0.6) indicate little biological material (Huguet et al., 2009). In this study, the BIX values ranged from 1.17-3.61, with an average of 2.23 ± 0.807 (Fig. 10b).
The high BIX values also supported that the fluorescence properties of WSOC were influenced by marine biological activities.

Conclusions 25
Characteristics of atmospheric sulfur (i.e., MSA(p) and nss-SO4 2-) and OC (i.e., WSOC and WIOC) species in marine aerosols, and the environmental factors influencing their distributions were investigated over the Southern Ocean and the Amundsen Sea during the austral summer. In the Amundsen Sea, atmospheric MSA(p) concentration drastically increased (up to 0.57 µg m -3 ), suggesting significant influences of marine biological activities on atmospheric MSA(p). The higher MSA(p) concentration was attributed to exceptionally high seasonal primary production during the austral summer, the dominance of WIOC was the dominant OC species in both the Southern Ocean and the Amundsen Sea, accounting for 75% and 73%, respectively. Despite extremely high Chl-a concentration in the Amundsen Sea, no significant differences of mean WSOC and WIOC concentrations were found between the Southern Ocean and the Amundsen Sea. However, the higher WSOC/Na + and WIOC/Na + ratios were observed in the submicron aerosol particles, especially in the Amundsen Sea where biological productivity was much higher than the Southern Ocean. 5 It is worth noting that the simultaneous measurements of chemical species in marine aerosols as well as chemical and biological properties of seawater in the Amundsen Sea allowed a better understanding of the effect of ocean ecosystem on OC species. Moreover, the fluorescence properties of water-soluble organic aerosols revealed that protein-like components are most likely produced as a result of biological processes of diatoms.
West Antarctica is one of the fastest-warming regions globally (Bromwich et al., 2013). Ice shelves and glaciers in the 10 Amundsen Sea have been shrinking at a remarkable rate (Rignot et al., 2008). Moreover, sea ice coverage is decreasing fast in the western Antarctic (Stammerjohn et al., 2012). Because ocean buoyancy, stratification, and trace metal distribution are affected by these changes, the regional oceanography, phytoplankton community structure and biogeochemical cycles of sulfur and carbon in the Amundsen Sea are likely affected as well (Yager et al., 2012). Further studies, therefore, are required to understand more clearly biogeochemical cycles of sulfur and carbon between the ocean and the marine atmosphere and should 15 focus on long-term monitoring of atmospheric sulfur and OC species in the Amundsen Sea.

Data availability
The data used in this study is available on request to the correspondence author Jinyoung Jung (jinyoungjung@kopri.re.kr).

Author contributions
JJ designed the research, carried out the experiments, processed the data, and wrote the paper. SBH, MC and LJ analyzed the aerosol samples. YL and EJY provided the marine biological data. JOC and JP helped in obtaining satellite products. JH, KP, DH and EJY contributed the scientific discussion and paper correction. TWK and SHL organized the field campaign.