Introduction
The ocean surface is a major source of submicron aerosols in both number and
mass concentrations (e.g., Spracklen et al., 2008). These aerosols play an
important controlling role in the atmospheric radiative budget because they
determine the number of cloud condensation nuclei (CCN) and ice nuclei (IN),
particularly over the remote ocean. In general, organic matter (OM) is
concentrated in the sea surface microlayer relative to the bulk seawater. OM
is further concentrated in aerosols during the bubble-bursting process,
which produces primary submicron sea spray aerosol (SSA) that is enriched in
OM (O'Dowd and de Leeuw, 2007). It has been recognized that marine
microorganisms play a large role in marine aerosol formation and its
composition. Marine-derived submicron organic aerosol (OA) can affect marine
aerosol optical depth (AOD) as well as CCN and IN concentrations.
Nevertheless there is still uncertainty in the chemical signatures of SSA,
leading to uncertainty in determining their climate impact.
The oceanic surface chlorophyll a (Chl a) concentration has been used as a
proxy for marine phytoplankton biomass. However, linear source functions
based on Chl a underpredict organic carbon (OC) enrichments for nascent SSA
produced from oligotrophic waters and overpredict OC enrichments for nascent
SSA produced from highly productive waters (Long et al., 2011). Recent field
and laboratory studies suggest that organic enrichment of SSA might be more
closely related to the concentration of oceanic dissolved organic carbon
(DOC), rather than to the concentration of Chl a in surface seawater (Prather
et al., 2013; Quinn et al., 2014).
To assess the impact of marine biological activity on ambient aerosols and
the subsequent formation of clouds, it is important to differentiate
marine-derived natural aerosols from anthropogenic aerosols over the oceanic
regions. For this purpose, methods need to be established to discriminate
between ocean- and land-derived aerosols found in marine atmospheres. A
method using the isotopic composition of aerosol carbon has been used
successfully to determine the contributions of marine and terrestrial
sources to aerosol found in the remote marine atmosphere (Chesselet et al.,
1981; Cachier et al., 1986; Kawamura et al., 2004; Miyazaki et al., 2011;
Ceburnis et al., 2011). The marine-derived OC (δ13C
∼ -22 to -18 ‰) is enriched in
13C relative to terrestrial C3 vegetation (which uses the Calvin–Benson
cycle as a metabolic pathway for carbon fixation in photosynthesis) and
fossil fuel OC (δ13C ∼ -30 to
-23 ‰ for both) (e.g., Fry and Sherr, 1984).
Ocean-derived submicron particles contain a large fraction of water-soluble
OC (WSOC), which can significantly alter the hygroscopic property of
aerosols (Prather et al., 2013) and act as CCN and IN. Only a few studies
have used the 13C of WSOC for source apportionment (Fisseha et al.,
2009; Kirillova et al., 2010; Miyazaki et al., 2012). The WSOC-specific
13C analysis in combination with organic molecular markers provides
robust tools for the source apportionment of WSOC in marine aerosols.
To better characterize submicron OAs in the marine boundary layer (MBL) and
differentiate them from those with terrestrial sources, we investigated the
stable carbon isotopic signature of WSOC in submicron aerosols collected
over the eastern equatorial Pacific. Primary productivity in that oceanic
region accounts for ∼ 23 % of the total productivity of the
entire Pacific Ocean (Pennington et al., 2006), and the potential of
enhanced OM in the surface microlayer is present. Deng et al. (2014) found
that long-chain organic molecules and humic-like substances (HULIS) were
prevalent in the marine aerosol sampled in a latitudinal cruise over the
eastern Pacific. Here we investigated possible sources of submicron WSOC
over the oceanic region by the analysis of WSOC-specific 13C combined
with several organic molecular markers, such as monosaccharides (glucose and
fructose) and low-molecular-weight (LMW) fatty acids (FAs).
Experimental
Submicron aerosol samplings
(Left) Cruise track of the National Oceanic and
Atmospheric Administration (NOAA) ship Ka'imimoana in the eastern equatorial Pacific
between 25 January and 1 March 2012 in three categorized oceanic areas (see
text). (Right) Typical 5-day back trajectories (red lines) that started
along the cruise track, together with monthly averaged concentrations of
chlorophyll a (Chl a) for February 2012. A vertical cross section of the back
trajectories is shown in the right bottom panel.
Aerosol samplings were conducted in the MBL on board
the National Oceanic and Atmospheric Administration (NOAA) R/V Ka'imimoana during the
Tropical Ocean tRoposphere Exchange of Reactive halogens and Oxygenated VOCs
(TORERO) field experiment (Coburn et al., 2014; Volkamer et al., 2015).
Figure 1 presents the cruise track over the eastern equatorial Pacific
between 25 January and 1 March 2012. The cruise originated in Honolulu,
Hawaii, and headed to Puntarenas, Costa Rica (KA-12-01), from
157 to 83∘ W longitude and
21∘ N to 8∘ S latitude.
A cascade impactor (CI; Series 230, Tisch Environmental, Cleves, OH, USA)
attached to a high-volume air sampler (HVAS; Model 120SL, Kimoto Electric,
Osaka, Japan) was used to collect submicron particles (Miyazaki et al.,
2012). The sampler was located on the upper deck of the ship. Aerosol
samplings were made using quartz fiber filters (25 × 20 cm) set on
the bottom stage of the impactor at a sampling flow rate of ∼ 1100 L min-1. The sampling time for each
sample was approximately 24 h. The samples were collected on precombusted (450 ∘C for 3 h) quartz filters, and the average total volume of each sample was 1318.7 m3.
The collected samples were stored individually in glass jars with a
Teflon-lined screw top cap at -20 ∘C prior to analysis.
The aerosol data taken during 1–28 February 2012 were used in this study.
To obtain the average size distributions of the WSOC mass, size-segregated
aerosol samplings were also performed with an Andersen-type
CI running in parallel to the HVAS. The size-segregated aerosol samples
were collected on precombusted quartz filters (8 cm ID) every
∼ 3 days. The sampling was made according to the 50 %
equivalent aerodynamic cutoff diameters, with nine stages between 0.39 and
10.0 µm (Miyazaki et al., 2010). Ambient air was drawn at a flow rate
of 120 L min-1 per sample without temperature and humidity control.
Other data were also obtained using an in situ ozone (O3) monitor,
sonic anemometer and Chl a fluorometer along the cruise track.
Aerosol chemical analysis
To determine the WSOC concentration, a portion of each filter sample (19.63 cm2) was extracted with ultrapure water (> 18 MΩ cm-1)
using an ultrasonic bath. The ultrapure water was generated by a
Sartorius Stedim Biotech arium pro ultrapure water system (Model 611:
Sartorius AG, Goettingen, Germany). The extracts were then filtrated with a
disc filter (Millex-GV, 0.22 µm, Millipore, Billerica, MA, USA),
followed by the injection of the DOC in the extracts into a total organic
carbon analyzer (Model TOC-Vcsh, Shimadzu, Kyoto, Japan) (Miyazaki et al.,
2011). The WSOC value of a field blank corresponded to less than
∼ 13 % of the WSOC concentration of the ambient samples. All
WSOC data presented here were corrected against field blanks.
For the determination of δ13CWSOC, a filter (14.13 cm2) for each sample was acidified to pH 2 with hydrochloric acid (HCl)
to remove inorganic carbon prior to extraction. The decarbonated filter
samples were then dried under a nitrogen stream for approximately 2 h. WSOC
was extracted from the filters in 20 mL of the ultrapure water using the
method as described above for measuring the WSOC concentration. The
extracted samples were concentrated by rotary evaporation, and 40 µL of
each sample was transferred to be absorbed onto 10 mg of pre-combusted
Chromosorb in a pre-cleaned tin cup. The δ13CWSOC was then
measured using an elemental analyzer (EA) (NA 1500, Carlo Erba, Milan,
Italy) interfaced to an isotope ratio mass spectrometer (IRMS) (Finnigan MAT
Delta Plus, Thermo Finnigan, San Jose, CA, USA). The δ13C data
are reported relative to an established reference of carbon Vienna Pee Dee
Belemnite (VPDB). The nitrogen isotope ratio (δ15N) of
water-soluble total nitrogen (WSTN) (δ15NWSTN) in aerosols
was also measured with basically the same procedure as δ13CWSOC, but without any acidification using HCl. In addition,
the concentrations of total carbon (TC) and the δ13C of TC
(δ13CTC) (i.e., without water extraction) were also
measured with the EA–IRMS for the same aerosol samples (Miyazaki et al.,
2010). Further details of the analytical method used for isotopic analysis
are given by Miyazaki et al. (2012).
For the determination of inorganic ions, another filter cut was extracted
with ultrapure water. The total extract was filtrated through a membrane
disc filter, and major anions including methanesulfonic acid (MSA) and
cations were determined using an ion chromatograph (Model 761 compact IC;
Metrohm, Herisau, Switzerland) (Miyazaki et al., 2011). The MSA value of
field blanks corresponded to less than ∼ 12 % of the
concentrations of the ambient samples, whereas the blank values of Na+,
Cl-, and Mg2+ were less than 1 % of the ambient concentrations.
For the analysis of possible tracers of marine DOC,
a filter portion was extracted with dichloromethane/methanol. The
–COOH and –OH functional groups in the extracts were reacted with
N,O-bis-(trimethylsilyl)trifluoroacetamide (BSTFA) to derive trimethylsilyl
(TMS) esters and TMS ethers, respectively. The TMS derivatives were then
analyzed for α-glucose, β-glucose, α-fructose, β-fructose, and a homologous series of straight-chain fatty acids
(C12–C19 saturated acids) using a capillary gas chromatograph
(GC7890, Agilent, Santa Clara, CA, USA) coupled to a mass spectrometer (5973
MSD, Agilent, Santa Clara, CA, USA) (Fu et al., 2011; Miyazaki et al.,
2012). The values of a field blank were less than ∼ 24 % of
the concentration of these molecular compounds in the ambient samples.
Trajectory analysis
To investigate air mass histories along the TORERO cruise track,
back trajectories were computed with the Real-time Air Quality Modeling
System (RAQMS) (Pierce et al., 2007), which calculated chemical and
meteorological forecasts. RAQMS has a horizontal resolution of
1∘ × 1∘, with 35 hybrid eta theta
vertical levels. Meteorological forecasts are initialized with operational
analyses from the National Centers for Environmental Prediction (NCEP)
Global Data Assimilation System (GDAS). The RAQMS calculations in
conjunction with reverse domain filling (RDF) techniques (Sutton et al.,
1994) are based on an analysis of back trajectories initialized along the
cruise track. A three-dimensional 7-day back trajectory was calculated using
the Langley Trajectory Model (LTM) (Pierce and Fairlie, 1993) and
initialized at model hybrid levels along the TORERO cruise tracks.
Results and discussion
Characteristics of sea salt particles
Time series of (a) the concentrations of Cl-,
Na+, and Mg2+; (b) local wind speeds measured on the ship; and (c) the
Cl- / Na+ molar ratios. The local wind speed data were merged into
the time interval of each filter sampling, and the average values with the SD
over each sampling duration are shown. R1, R2, and R3 denote oceanic regions
1, 2, and 3, respectively, which are defined in the text and shown in Fig. 1.
Figure 2 presents a time series of concentrations of Na+, Cl-, and
Mg2+ as tracers of sea spray, together with the daily-averaged local
wind speed measured on the ship at each aerosol sampling location. In
general, temporal variations of Na+, Cl-, and Mg2+ in the
submicron particles were correlated with variation in local wind speeds with
r2 of 0.41–0.60 (n=21). This is consistent with the wind-driven
production of primary marine aerosol particles, whereas the moderate linear
correlation can be explained by a power law relationship between sea spray
mass and local wind speed (e.g., Ovadnevaite et al., 2012).
In this study, the aerosol sampling regions were classified into three
categories according to the differences in oceanic areas and patterns of
backward trajectories (Fig. 1). Region 1 (R1), sampled during the period of
1–7 February 2012, corresponded to open oceans at 5–15∘ N and
112–133∘ W. Most of Region 2 (R2) covered the oceanic area in the
Southern Hemisphere (8∘ S–2∘ N and 93–110∘ W),
where the sampling was conducted during 9–19 February. Region 3 (R3) was
close to the coastal region at 0–8∘ N and 84–90∘ W,
where observations were made during 20–28 February. R1 and R2 are characterized
by very low anthropogenic impact on marine ecosystem (Halpern et al., 2008)
and represent some of the most pristine ocean environments at tropical
latitudes with the low Chl a concentrations (Fig. 1). According to the
back trajectories, the air masses sampled in R1 and R3 (i.e., in the
Northern Hemisphere) had been transported over the ocean for at least 48 h
prior to aerosol sampling on the ship. The trajectories further indicated
that those air masses were not significantly influenced by the land surface
for at least 5 days. The air masses sampled in R2 had been transported
over the ocean in the Southern Hemisphere for at least 5 days without any
significant influence from the land surface or pollution. The relative
influence of ocean surface and land on the observed aerosols will be
discussed in Sect. 3.3.
In R3, enhanced marine biological activity at the sea surface was observed,
with an average Chl a concentration of 0.15 ± 0.04 mg m-3 (Fig. 1). Because this value is substantially larger than the average
concentration in R1 and R2, R3 is characterized as
a high-Chl a region. The
enhancement of the Chl a concentration in R3 (up to 0.33 mg m-3) could
be attributed to surface mixing in the Pacific Eastern Boundary Upwelling
System (EBUS) (Rossi et al., 2009) and the coastal region (Pennington et
al., 2006). R1 was characterized by high concentrations of sea salt
particles with the average molar ratio of chloride to sodium
(Cl- / Na+) close to unity. This is not necessarily expected in
submicron aerosols in the tropical oceanic regions, because rapid
acidification of sea salt particles occurs on the timescale of seconds
(e.g., Pszenny et al., 2004; Keene et al., 2009). The fact that a depletion
of Cl- is apparently less pronounced in R1 indicates that the
concentrations of gas species including organic acids (e.g., Laskin et al.,
2012) responsible for the Cl- loss were substantially low in R1. The
current results suggest that the submicron particles collected in R1 were
more similar to nascent sea spray aerosols compared to those in R2 and R3.
The Cl- / Na+ ratio tended to decrease from R1 (av 1.06 ±
0.23) to R2 (av 0.60 ± 0.24) and R3 (av 0.32 ± 0.25). In
fact, the Cl- / Na+ ratio tended to decrease with increasing sulfate
concentration (not shown in the figure), whereas this trend is not apparent
for nitrate. This result suggests the Cl- depletion by acid
substitution in seawater-derived NaCl and indicates production of more
chemically aged particles in R3 relative to R1 and R2.
Size distribution and time series of WSOC and related parameters
Typical mass size distributions of water-soluble organic
carbon (WSOC) in R1, R2, and R3.
Figure 3 shows the typical mass size distributions of WSOC for each regional
category during the TORERO cruise. In general, WSOC displayed a bimodal size
distribution, with peaks in the submicron and supermicron particle-size
ranges. Bimodal size distributions of WSOC in marine aerosols were also
observed in the western North Pacific (Miyazaki et al., 2010), whereas both
unimodal and bimodal size distributions of water-soluble organic species
were also reported in particles collected at a coastal site facing the
eastern North Pacific (Maudlin et al., 2015). The bimodal size observed in
this study can be attributed to the difference in the formation processes of
WSOC between the two size ranges. The two distinct size modes include (i) direct co-emissions associated with sea salt particles in both size ranges,
(ii) aqueous-phase products in the submicron size range, and (iii) partitioning to the surface of coarse particles (i.e., sea salt) and/or
heterogeneous reactions in the supermicron size range (Mochida et al.,
2002). Although it is difficult to provide a clear explanation by this dataset
alone, the observed WSOC size distributions might be explained by some
combination of these possible origins and processes. Here we focus on the
submicron size of WSOC relevant to its isotope ratios and several chemical
tracers.
Time series of (a) WSOC, WSOC / total carbon (TC); (b) the
mass ratio of WSOC / Na+; (c) δ13CTC and δ13CWSOC; (d) O3 mixing ratios; and (e) in situ Chl a
concentrations in surface seawater, along with the ship's position
(latitude). The data for the WSOC / Na+ ratio during 19–21 February are
not shown because the Na+ concentrations were extremely low (< 0.03 µg m-3; Fig. 2a), which substantially increased the ratio.
Figure 4a shows a time series of the mass concentration of WSOC and its
ratio to TC in the submicron particles. In R1 and R2, the WSOC
concentrations ranged between 50 and 160 ngC m-3, with averages of 130 ± 27 and 85 ± 24 ngC m-3 for the two regions, respectively.
The WSOC / TC ratios ranged between 20 and 50 %, with averages of 36 ± 10 % (R1) and 31 ± 10 % (R2). Although black carbon (BC)
concentration in the TC fraction was not measured in this study, most of the
TC can be attributed to OC under the assumption of extremely low concentrations
of BC (< 2 ng m-3) previously observed over this oceanic region
(Shank et al., 2012). The lower WSOC / TC ratios in R1 and R2 can be explained
by fresh primary marine aerosols enriched in water-insoluble organic carbon
(Facchini et al., 2008).
Average (± standard deviation) and median concentrations and
ratios in the different oceanic areas during the Tropical Ocean tRoposphere
Exchange of Reactive halogens and Oxygenated VOCs (TORERO) cruise
observation. Values in parentheses show those of the 67th percentile. The
precision of each measurement including the blank subtraction is also shown.
R1 (1–7 Feb 2012)
R2 (9–19 Feb 2012)
R3 (20–28 Feb 2012)
Precision
Na+ (µg m-3)
2.78 ± 0.73 2.51 (3.46/2.49)
1.25 ± 0.47 1.19 (1.20/0.95)
2.00 ± 0.60 1.78 (1.83/1.77)
4 %
Cl- / Na+ molar ratio
1.06 ± 0.23 1.15 (1.16/0.95)
0.60 ± 0.24 0.58 (0.60/0.54)
0.32 ± 0.25 0.22 (0.25/0.19)
6 %
WSOC / Na+ (ngC ng-1)
0.05 ± 0.01 0.05 (0.06/0.04)
0.08 ± 0.05 0.07 (0.08/0.04)
0.31 ± 0.13 0.35 (0.37/0.32)
16 %
WSOC (ngC m-3)
130 ± 27 113 (159/109)
85 ± 24 83 (84/70)
515 ± 268 588 (603/572)
15 %
TC (ngC m-3)
380 ± 132 337 (377/324)
322 ± 160 259 (323/238)
978 ± 345 973 (1106/840)
9 %
WSOC / TC (%)
36 ± 10 32 (43/29)
31 ± 10 30 (30/24)
62 ± 19 62 (66/58)
17 %
δ13CWSOC (‰)
-19.1±0.7 -19.3 (-18.4/-19.5)
-19.6±2.2 -18.4 (-18.0/-19.0)
-18.8±1.2 -18.7 (-18.5/-19.0)
0.7 ‰
δ13CTC (‰)
-22.0±2.0 -21.3 (-21.0/-23.9)
-22.1±1.4 -22.5 (-22.0/-22.7)
-20.7±0.7 -21.0 (-20.9/-21.2)
0.5 ‰
δ15NWSTN (‰)
10.1 ± 1.3 10.0 (10.9/9.9)
11.6 ± 1.3 11.4 (11.5/11.3)
12.8 ± 5.4 15.4 (16.5/14.3)
0.8 ‰
Glucose (ng m-3)
0.11 ± 0.04 0.11 (0.12/0.10)
0.05 ± 0.08 0.03 (0.03/0.02)
1.55 ± 0.66 1.52 (1.89/1.15)
13 %
Fructose (ng m-3)
0.02 ± 0.01 0.03 (0.03/0.02)
0.01 ± 0.01 0.01 (0.01/0.00)
0.48 ± 0.30 0.34 (0.39/0.28)
13 %
Fatty acids (C12-C19) (ng m-3)
1.38 ± 0.47 0.97 (1.40/0.69)
3.60 ± 2.20 1.49 (1.58/1.37)
5.82 ± 3.02 3.32 (3.77/2.86)
28 %
MSA (ng m-3)
92 ± 13 99 (102/82)
141 ± 21 144 (146/131)
123 ± 20 118 (121/114)
18 %
Chl a (mg m-3)
0.112 ± 0.016 0.107 (0.129/0.100)
0.106 ± 0.015 0.101 (0.107/0.097)
0.147 ± 0.037 0.125 (0.127/0.123)
NA
In contrast, both the WSOC concentrations (515 ± 268 ngC m-3) and
the WSOC / TC ratios (62 ± 19 %) were substantially higher in R3 than
those in R1 and R2. Further, the correlation of WSOC with Na+ was
strongest in R3 (r2=0.40). In the submicron particles, the average
WSOC / Na+ ratio was 0.15 ± 0.14 (Fig. 4b), which is within (though
near the lower end of) the OC / Na+ ratio range (0.1–2.0) previously
reported for submicron marine primary OA (Russell et al., 2010; Frossard et
al., 2014). This result is consistent with our understanding that the
submicron SSA is enriched in OC relative to seawater (O'Dowd et al., 2004:
Keene et al., 2007). The enrichment of water-soluble organics in the
submicron particles is particularly significant for R3 (Fig. 4b), where the
average WSOC / Na+ ratio (0.31 ± 0.13) was substantially higher
than that in R1 (0.05 ± 0.01) and R2 (0.08 ± 0.05) (Table 1).
The enrichment of organics can be attributed to the phytoplankton blooms
identified by the increased concentrations of Chl a in seawater (up to 0.33 mg m-3) in R3 (Fig. 4e), together with the spatial distributions
measured by the satellite (Fig. 1). Previous studies have shown a linkage
between organics and high Chl a concentrations on timescales of months
(O'Dowd et al., 2004; Sciare et al., 2009). However, Quinn et al. (2014)
found no well-defined relationship between instantaneous Chl a in seawater
and organic-mass enrichment in sea spray, suggesting no significant
variability in the OC content of freshly emitted sea spray aerosol, despite
significant variability in seawater Chl a levels. This point will be
discussed in Sect. 3.3. The higher WSOC / Na+ ratio in R3 can be also
interpreted as an indicator of secondary contributions of photochemical
products of primary OA and/or marine biogenic organic gas species to the
observed aerosols during the aging, as indicated by the enhanced levels of
O3 (up to 25 ppbv) (Fig. 4d) and the decreased Cl- / Na+ ratio
(Fig. 2).
Isotopic characterization of aerosol WSOC and TC
Average values of the WSOC concentration, the WSOC / TC
ratio, δ13CTC, and δ13CWSOC in each
oceanic area.
As shown in Fig. 4c, the δ13CWSOC ranged from -23.0 to
-15.7 ‰, with an average of -19.8 ± 2.0 ‰ during the cruise. The δ13CWSOC
values were systematically higher than the δ13CTC ranging
from -25.5 to -19.7 ‰, with an average of -21.8 ± 1.4 ‰ (Fig. 5). On average, WSOC was enriched in
13C by ∼ 2.0 ‰ relative to TC,
indicating that 13C-enriched submicron carbonaceous aerosol is
preferentially water soluble. Regardless of the oceanic area, the average
δ13CWSOC values of -19.6 to -18.8 ‰
(Table 1) were within the typical range of δ13C in the DOC pool
of seawater (-22 to -18 ‰; Fontugne and Duplessy,
1981). This range is influenced by factors such as local ocean temperatures
and phytoplankton species, whereas changes in δ13C resulting
from trophic transfers are minimal (e.g., Guo et al., 2003). In the eastern
North Pacific and in tropical oceans, the δ13C of DOC typically
ranges from -22 to -20 ‰ in surface seawater (Bauer
and Druffel, 1998). In contrast, relatively few studies have measured the
δ13C signature in aerosol WSOC, which ranges from -25.5 to
-23 ‰ at rural and urban sites, and is generally
attributable to terrestrial and fossil sources (Kirillova et al., 2010;
Wozniak et al., 2012). The δ13CWSOC measured in this study
indicate an enrichment of sea-surface-derived DOC in submicron WSOC aerosols
throughout the study region, and the 13C-enriched WSOC over TC cannot
be explained by influences of land surface. It should be noted that this
enrichment of 13C in WSOC could be partly due to isotopic fractionation
throughout the partitioning of semi-volatile organics between the gas and
particle phases (Fisseha et al., 2009). In equilibrium, partitioning between
the gas and particle phases leads to larger 13C of particle-phase
organic compounds than the corresponding gas-phase compounds (Gensch et al.,
2014). However, even if this effect (±2.0 ‰) is
taken into account, the δ13CWSOC values were still within
the range of δ13C of DOC.
A combination of both carbon and nitrogen isotopic signatures can provide
better information on the sources of dissolved organic matter (DOM) in
marine aerosols than carbon isotopes alone. Figure 6 shows the ranges of the
nitrogen isotope ratio of the water-soluble total nitrogen (δ15NWSTN) and δ13CWSOC in the submicron aerosols
for each oceanic region. The δ15NWSTN ranged between 3.5
and 16.7 ‰, with an average of 11.8 ± 3.1 ‰. The wide range of δ15NWSTN
values was partly due to the fact that WSTN contains inorganic nitrogen,
such as NO3- and NH4+, in addition to water-soluble
organic nitrogen (ON). In general, the observed values were similar to the
δ15N values in surface seawater (i.e., 2 m depth). Benner et
al. (1997) reported a dataset of δ15N values for marine
high-molecular-weight DOM samples obtained in the Gulf of Mexico and the
Pacific and Atlantic oceans, which ranged from 6.6 to
10.2 ‰. The δ15NWSTN in aerosol also
provide evidence of a significant contribution of DOC to the observed
submicron aerosols.
Figure 7 shows the percentage exposure (percent of time over 7 days) of the
sampled air mass to ocean and land surfaces along the cruise track as
functions of altitude and time. The calculated air parcels were initialized
at each sampling location along the cruise track. This Lagrangian trajectory
analysis showed very low exposure (< 20 %) of air parcels at the
sampling points on the ship to boundary layers over land, consistent with
the results of the isotopic analysis, and suggested that the majority of
submicron WSOC originated from the sea surface during the study period. It
is noted that the observed aerosols in R3 had been transported by low-level
air flow from the Atlantic, as indicated by the back trajectories (Fig. 1). In fact, the trajectories had passed over the Isthmus of Panama at
higher altitudes, followed by descent to the sampling point in R3 as seen in
Fig. 1, indicating less influence from the land surface. This is
consistent with the results from the isotopic analysis of WSOC, which
suggest that the influence of the land surface on the observed WSOC was
insignificant.
The ranges of the δ15NWSTN and δ13CWSOC in the submicron aerosols obtained in each oceanic area.
The rectangle in the panel indicates the typical ranges of δ15N
and δ13C for dissolved organic matter (DOM) in the eastern
equatorial Pacific (see text).
The percentage exposure of the 7-day air mass to the (a) maritime and (b) continental planetary boundary layer as functions of
altitude and time. The calculation was made along the TORERO cruise track
with the Real-time Air Quality Modeling System (RAQMS).
In R3, the elevated levels of WSOC along the cruise track were not always
accompanied by the increase of Chl a on a daily timescale. Specifically, the
Chl a concentrations displayed an insignificant increase on 22 and 27–28 February, whereas the WSOC concentrations increased, ranging from 300 to 900 ngC m-3
during the same periods (Fig. 4a and e). Deng et al. (2014)
also observed the lack of correlation between organics and Chl a over the
eastern Pacific. They attributed it to the small variation in Chl a and the
fact that aerosol composition is only sensitive to major changes in Chl a.
Rinaldi et al. (2013) observed time lag between Chl a and OM enrichment in
aerosol, suggesting that biological processes in oceanic surface waters and
their timescales should be considered when modeling the production of
primary marine OA. Quinn et al. (2014) assessed the relationship between the
OC content of seawater and freshly emitted SSA in the presence and absence
of phytoplankton blooms in the North Atlantic and the coastal waters of
California. They concluded that there is a large reservoir of OC in surface
seawater that results in the enrichment of OC in SSA. They also reported
that the oceanic source of OC in the region is uncoupled from, and
overwhelms any influence of, local biological activity as measured by Chl a over large ocean regions. O'Dowd et al. (2015) showed that a correlation
between OM in sea salt particles and Chl a increased as the timescale
increased from daily to monthly intervals and suggested that OM production
is closely linked with the decay phase of the bloom and is driven by
nanoscale biological processes that release large quantities of transferable
OM in surface seawater. The results of our study support those of previous
studies in showing that linear source functions based on Chl a might not
properly predict OC enrichments for SSA on the timescale considered here.
Monosaccharides, fatty acids, and MSA as marine biogenic tracers
We also used several organic molecular markers to further investigate the
contribution of DOC to the submicron organics in concert with the isotope
tracers described previously. The analysis of sea surface waters for
organics has revealed a significant carbohydrate concentration, including
glucose (Aluwihare et al., 1997), whereas primary saccharides (e.g.,
glucose) in aerosol have been suggested as possible tracers for surface soil
dust and/or biomass burning (Simoneit et al., 2004). Electron
ionization–mass spectrometry (EI-MS) measurements of marine aerosol in the
western Pacific revealed substantial contributions from carbohydrates such
as glucose and levoglucosan, and the former is partially attributed to
organics from the ocean surface (Crahan et al., 2004). Low-molecular-weight
LMW FAs have multiple sources associated with marine
microbial activity, vascular plants, and microbes (Mochida et al., 2002;
Kawamura et al., 2003). Burrows et al. (2014) introduced a framework for
parameterizing the fractionation of marine OM into SSA and partitioned
marine OM into different classes, including a polysaccharide-like mixture
associated with semilabile DOC, a lipid-like mixture associated with labile
DOC, and others. In this study, we investigated the possible contributions
of types of DOC to submicron organics using the molecular markers of DOC.
Time series of the concentrations of (a) glucose and
fructose with their ratios to Na+, (b) fatty acids
(C12–C19) with their ratios to Na+, and (c) methanesulfonic
acid (MSA) and the Cl- / Na+ molar ratios in the submicron aerosol
samples collected during the cruise.
Figure 8a–b show a time series of concentrations of glucose, fructose, and
LMW FAs (C12–C19) in the submicron particles collected. The
concentrations of both glucose and fructose were elevated in R3, with
average values of 1.6 ± 0.7 and 0.5 ± 0.3 ng m-3, respectively. The observed concentrations of glucose and LMW FAs
were similar to those observed in total suspended particulate matter (TSP)
over coastal areas in California and western Mexico (1.0–1.4 and 1.0–6.0 ng m-3 for glucose and LMW FAs, respectively) (Fu et al., 2011). The
temporal trends of these saccharides were similar to that of WSOC, with
r2 of 0.82 (n=21). The mass ratio of (glucose +
fructose) / Na+ was substantially higher in R3 compared to the other
regions, indicating an enrichment of these monosaccharides in submicron sea
salt particles over oceanic areas with high biological activity. In
contrast, the correlation between LMW FAs and WSOC was less significant
(r2=0.31). The combined results of the organic molecular tracers
and δ13CWSOC indicate a substantial contribution of
saccharide-related DOC associated with sea spray to submicron WSOC. The
results also suggest that the monosaccharides detected here might be
suitable indicators for the ocean-derived submicron WSOC over the study
region.
Russell et al. (2010) used reference Fourier transform infrared (FTIR)
spectra of 11 different saccharides, including glucose, and found that a
majority of organic component in ambient marine submicron aerosol consisted
of organic hydroxyl groups characteristic of saccharides. Frossard et al. (2014) observed a significant amount of monosaccharides and disaccharides in
model-generated primary marine aerosols from bubbled seawater, whereas the
organic mass hydroxyl group in seawater was mostly characterized by
polysaccharides. They attributed this finding to the larger saccharides
preferentially remaining in the seawater during the primary OA production.
Miyazaki et al. (2014) found lactic and glycolic acids, which are LMW
hydroxyacids that can be produced as the major metabolic end products of
carbohydrate fermentation, in marine aerosols obtained from biologically
active oceanic regions of the western North Pacific. The results of our
study on glucose and fructose in the submicron WSOC were consistent with the
chemical signatures of marine OA reported in those studies. Moreover, our
result is in line with a modeling study by Burrows et al. (2014), who
simulated that, in regions such as the Southeast Pacific, semilabile DOM
contributes significantly to estimated aerosol organic mass as saccharides
and proteins. They also reported an anticorrelation between Chl a and OM
fraction in their model. Contributions of proteins to the submicron WSOC in
the same samples are discussed in Chen et al. (2016) using excitation–emission matrices.
MSA in aerosol is also considered a marker of marine biogenic origin,
because it is a major oxidation product of dimethyl sulphide (DMS). The mass
concentration of MSA increased in R2, with an average of 141 ng m-3,
which was greater than in R3 (123 ng m-3) (Fig. 8c). The increase in
the background level of MSA did not necessarily accompany the increase in
the background level of WSOC in R2 (Fig. 4a). In fact, a globally coupled
ocean–atmosphere model calculation showed a “hot spot” of mean sea surface
DMS in the upwelling zones of the eastern equatorial Pacific from December
to May whose oceanic area corresponded to R2 (Kloster et al., 2006).
The observed MSA is considered to be either produced by gas-phase MSA
directly scavenged by aerosols or rapidly produced in aqueous phase from
scavenged dimethylsulfoxide (DMSO) and methanesulfinic acid (MSIA) (Zhu et
al., 2006), particularly under conditions with high relative humidity
typical of the MBL. Assuming a typical average OH concentration of 1 × 106 cm-3, a lifetime of DMS is roughly estimated to be
< ∼ 1 day with respect to oxidation by OH in the MBL
(Davis et al., 1999; Kloster et al., 2006). The rapid oxidation of the
intermediate reaction products of DMS to produce MSA (on a timescale of
hours) and the typical residence time of submicron aerosols in the MBL
(∼ 5–7 days) indicate that the measured MSA likely reflects
the larger emission of DMS around the sampling locations of R2.
Additionally, a large missing source of MSA photolytically enhanced during the
daytime has been as suggested by Zhang et al. (2014) and would be
consistent with the lowest solar zenith angle in R2 (Coburn et al., 2014).
The MSA concentrations generally increased with the decreasing
Cl- / Na+ ratios (Fig. 8c), which is consistent with the chemical
aging of the observed aerosols. Another possible reaction of DMS with
species other than OH to produce MSA includes BrO+DMS (Saiz-Lopez et al.,
2004). However, BrO in the MBL over this region was extremely low (generally
below 0.5 pptv) during the same observational period (Volkamer et al.,
2015), indicating that the reaction of BrO+DMS is likely insignificant
source for MSA. Gaston et al. (2010) suggested a possible catalytic role of
vanadium in MSA formation. The observed increases in the MSA concentrations
were most evident in R2, in which the impacts of anthropogenic sources
appeared to be very low. Therefore, the effects of such a catalytic reaction
on the increases in MSA concentrations in R2 are likely insignificant. The
lack of correlation between MSA and WSOC implies that the presence of DMS in
seawater and its subsequent oxidation to MSA were not necessarily linked to
the formation of submicron WSOC over this oceanic region. This confirms the
difficulties of connecting Chl a with DMS concentrations in seawater over
subtropical and tropical ocean as previously suggested (Bell et al., 2010).
Contribution of marine OC sources to the WSOC aerosol
To estimate the relative contribution from marine and terrestrial OC sources
to the observed WSOC, an isotopic mass balance equation assuming a
two-end-member isotopic mixing was used (e.g., Turekian et al., 2003;
Miyazaki et al., 2010). We applied 13C values ranging from -22 to
-18 ‰ for marine OC and those ranging from -27 to
-26 ‰ for terrestrial OC (e.g., Kirillova et al., 2010)
typically found in the Northern Hemisphere. The effect of isotopic
fractionation by heterotrophic degradation on OM is considerably small
(∼ 1 ‰ for 13C; Shaffer et al., 1999).
Our calculation indicates that, on average, marine sources contribute
∼ 90 ± 25 % of the aerosol carbon. As discussed
previously, the higher WSOC / Na+ ratio in R3 indicates some contribution
of a secondary, ocean-derived source to WSOC, although it is difficult to
quantify their contributions to the WSOC mass.
The results of our study contradict those of Shank et al. (2012), who
suggested that there was little to no marine source of submicron OA to the
atmosphere in a similar oceanic region (corresponding to R1 and R2 in the
current study) over the eastern South Pacific. Shank et al. (2012) reported the
average concentrations of non-refractory organics in submicron aerosols to
be as low as 70 ng m-3 with a maximum of 170 ng m-3 at most
measured with an Aerodyne high-resolution time-of-flight mass spectrometer
(HR-ToF-AMS). Assuming that most of the TC in this study can be attributed
to OC in R1 and R2 and given OC-to-OM conversion factors of 1.8 for
water-soluble OM and 1.4 for water-insoluble OM reported for the marine OA
(Facchini et al., 2008), the average OA concentration in R1 and R2 is
estimated to be ∼ 490–580 ng m-3 (cf. Table 1). These
values are substantially larger than those reported by Shank et al. (2012).
One possible explanation for the contradiction between our study and Shank
et al. (2012) is that the studies were conducted in different seasons with
different meteorological conditions and microbial activity at the sea
surface. Another possible explanation is that the AMS could not detect a
significant fraction of refractory material (e.g., HULIS) found in primary
marine OA over the study region (Deng et al., 2014). Our analysis of δ13CWSOC and organic molecular markers indicated that DOC in
surface seawater contributed substantially to the submicron WSOC levels
regardless of the oceanic area of the study region. It is noted that the
contribution of anthropogenic sources cannot be negligible although this is
indicated by the isotopic analysis. Nevertheless, the present study
demonstrates that DOC is closely correlated with the submicron WSOC aerosol
concentration and implies that it may characterize background OA in the MBL
over the study region.