te
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us
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equ
81%
ide
y of
ary
ate constant of 3.8×10−3 h−1. Based on these results and a conceptual model, we
2O) are
ic lifeti
f enha
he stra
impacts O (Nevison and Holland, 1997). hampers shelf-basin exchange by restricting the wave energy and
Marine Chemistry 121 (2010) 80–86
Contents lists available at ScienceDirect
Marine Ch
j ourna l homepage: www.e lse3
Constraining the marine contribution to tropospheric CH4 and N2O
suffers from data scarcity in key areas and inherent uncertainties in
scaling sea-to-air fluxes. Current best estimates are that marine waters
provide b5% of tropospheric CH4 and 29–55% of tropospheric N2O; the
highest contributions are from estuaries and shallow coastal shelves
(Bange et al., 1994, 1996; Bates et al., 1996; Seitzinger and Kroeze,
1998; Upstill-Goddard et al., 2000). Methanogenesis, an anaerobic
microbial processmediatedby archaea (MadiganandMartino, 2006), is
thought to dominate marine CH4 emissions, although cold seepage of
thermogenic CH4 and the dissolution of CH4 hydrates may also be
thereby vertical diffusivity (Carmack and Chapman, 2003). Changes to
the internal dynamics and sea–air fluxes of CH4 and N2O are therefore
expected with decreasing sea ice cover. Additional CH4 release might
result from temperature destabilisation of gas hydrates on the shallow
continental shelves (Westbrook et al., 2009), and organic matter stored
in terrestrial and marine permafrost might be mobilised onto the
shelves leading to further CH4 andN2O release throughmethanogenesis
and denitrification.
As far as we are aware there is only one previously published N2O
dataset for the Arctic Ocean (Hirota et al., 2009), and existing CH4significant (Judd, 2004; Kvenvolden and Rog
product of aerobic nitrification and an in
anaerobic denitrification (Hattori, 1983; Kapl
⁎ Corresponding author. Plymouth Marine Laboratory
E-mail address: vak@pml.ac.uk (V. Kitidis).
0304-4203/$ – see front matter. Crown Copyright © 20
doi:10.1016/j.marchem.2010.03.006tosphericwater cycle and
l radical (OH) and ozone
tratospheric NOx which
(Comiso, 2002; Comiso et al., 2003; Gregory et al., 2002; Holland et al.,
2006; Stroeve et al., 2007b; Teng et al., 2006; Wadhams, 1995). Sea ice
presents a barrier to turbulent diffusion across the sea–air interface andin the photochemistry of tropospheric hydroxy
(O3) (Crutzen, 1991), and N2O generates sNitrous oxide
Greenhouse effect
Sea ice
1. Introduction
Methane (CH4) andnitrousoxide (N
gases with relatively long atmospher
active, together accounting for ∼20% o
(Forster et al., 2007). CH4 is involved in tCrown Copyright © 2010 Published by Elsevier B.V. All rights reserved.
important biogenic trace
mes. They are infrared-
nced greenhouse forcing
The Arctic Ocean is particularly sensitive to global warming through
polar amplification and is thus recognised as one region where the
resulting effects on ecosystems will be among the most dramatic
(Denman et al., 2007; Holland and Bitz, 2003). Arctic sea ice thickness
and extent are in decline due to increasing sea-surface temperaturesMethane suggest that future sea-ice retreatmay decrease the residence times ofmethane and nitrous oxide in the surface
Arctic Ocean and thus enhance the sea–air flux of these climatically active gases.Arctic zone
Polar waters in surface seawater with a rMethane and nitrous oxide in surface wa
Arctic Ocean
Vassilis Kitidis a,b,⁎, Robert C. Upstill-Goddard a, Leif
a Ocean Research Group, School of Marine Science and Technology, Newcastle University,
b Plymouth Marine Laboratory, United Kingdom
c Department of Chemistry, University of Gothenburg, Sweden
a b s t r a c ta r t i c l e i n f o
Article history:
Received 21 March 2009
Received in revised form 17 March 2010
Accepted 17 March 2010
Available online 24 March 2010
Keywords:
Dissolved methane and nitro
Passage between the North
with respect to atmospheric
range of 58–528% and 82–1
low methane and nitrous ox
was also found in the vicinit
water column and sedimenters, 2005). N2O is a by-
termediate product of
an, 1983).
, United Kingdom.
10 Published by Elsevier B.V. All rigr along the North-West Passage,
Anderson c
ed Kingdom
oxide in seawater were measured along a 6700 km transect of the North-West
ntic Ocean and Beaufort Sea in the Arctic Ocean. Over- and under-saturation
ilibrium were observed for both gases. Methane and nitrous oxide were in the
saturation, respectively. Under-saturation was attributed to melt-water with
, while over-saturation was found under multi-year sea-ice. Elevated methane
the marginal ice zones and the Mackenzie River plume. Our data support both
sources of methane and nitrous oxide. We found first-order methane oxidation
emistry
v ie r.com/ locate /marchemobservations are restricted to localised studies that all showed strong
CH4 super-saturation relative to atmospheric air: theWhite Sea (Barents
Sea) (Savvichev et al., 2004), the Beaufort Sea (Kvenvolden et al., 1993;
Macdonald, 1976), the East Siberian Sea (Shakhova and Semiletov,
2007; Shakhova et al., 2005), and coastalwaters off Spitsbergen (Damm
et al., 2007, 2008).
In this paper we present data for CH4 and N2O collected during an
approximately 7000 km transect from the south of Iceland, through
hts reserved.

Baffin Bay, the North-West Passage and the Beaufort Sea during July
2005. Due to the nature of the cruise being a transit leg for subsequent
shorebasedfieldwork, samplingwasnecessarily limited toopportunistic
surface underway (7 m) measurements. Nevertheless, we were able to
also carry out a single 918 mhydrocast in Baffin Bay. Ourmeasurements
represent the first ever basin-wide survey of CH4 and to the best of our
knowledge only the second N2O dataset from the Arctic Ocean.
2. Methods
Samplingwas carried out onboard IBOden during Beringia 2005 (4–
28 July, 2005), a research initiative organized by the Swedish Polar
Research Secretariat (http://www.polar.se/beringia2005/eng/index.
html). The cruise track is shown in Fig. 1. Underway sampling for CH4
and N2O was in 1 L volumetric flasks with the exclusion of all bubbles,
via a branch from the ship's pumped thermosalinograph supply (Sea-
Bird, SBE 45; inlet 7 m). In order to remove ice-blockages, the supply's
inlet was fitted with a back-flush systemwhich was operated from the
main lab following an audible alarm. Samples were only collected after
the supply of seawater was restored for at least 5min. Samples for
dissolved silicate were similarly collected in polypropylene containers
gases (UpstillGoddard et al., 1996). For the incubation samples the
procedure was modified to allow in-syringe equilibration (Upstill-
Goddard et al., 1990).Method calibration used gravimetrically prepared
primary standards (BOCgases Ltd.: 0.72 and 5.47 ppmv for CH4 and 0.71
and 0.41 ppmv for N2O). Each sample run was followed by a
corresponding calibration standard and air sample collected from a
gas line outside the bridge. Percent CH4 and N2O saturations relative to
atmospheric air were calculated from the in situ concentrations and
solubility data for CH4 and N2O (Weiss and Price, 1980). Some of our
observationswere below the temperature range of data onwhich these
solubility functions were based (N−1 °C). Although this is not entirely
satisfactory, there are currently no formulations for the required tem-
perature range to our knowledge. Averagemixing ratios of atmospheric
CH4 was 1.86 ppmv (range: 1.81–2.00 ppmv) and N2O was 320 ppbv
(range: 279–358 ppbv).
Daily sea-ice cover was obtained from the National Snow and Ice
Data Centre (NSIDC; http://www.nsidc.org) Nimbus-7 SMMR and
DMSP SSM/I 25 km resolution datasets (Knowles et al., 2002).
3. Results
rs fr
G), M
81V. Kitidis et al. / Marine Chemistry 121 (2010) 80–86(NUNC cryotubes, cat. no. 379146) and stored at−20 °C until analysis
by standard colorimetry (Hansen and Koroleff, 1999). In addition CH4
andN2Owere sampled on a 918 mNiskin bottle hydrocast in Baffin Bay
(73°44.78 N, 77°3.72 W) and at a surface melt-pond during an ice
station (73°24.73 N, 96°18.13W).Water depth at the hydrocast station
was approximately 950 m.
Rates of microbial CH4 and N2O production/consumption were
investigated by incubating surface seawater from under sea ice, and
melt-pondwater fromanadjacent ice station. Incubations of headspace
free duplicate samples were at 4 °C under fluorescent lighting, in gas-
tight, 100 mL glass syringes (Samco), fitted with gas-tight, 3-way
polycarbonate stopcocks. The melt-pond samples were collected
directly from the pond edge in glass syringes while the seawater
samples were collected with a 10 L Niskin bottle. On deck, the Niskin
bottle was connected to a polycarbonate vessel fitted with a 0.45 μm
filter at the outflow, thereby concentrating the plankton community in
the vessel 15-fold by gravity filtration. Glass syringes containing the
concentrated plankton community were then filled directly from the
Niskin and incubated as above.
All water samples were analysed for CH4 and N2O by single phase
equilibration (UHP N2) gas chromatography with flame ionisation and
electron capture detection for CH4 and N2O respectively; analytical
precision for selected samples (n=3) was better than ±2% for both
Fig. 1. Sampling locations (dots) along the cruise track of RV Oden with 1000 km marke
Bay, Lancaster Sound (LS), Peel Sound (PS), McClintock Channel (MC), Amundsen Gulf (A
the Ocean Data View software. The locations of the Baffin Bay depth profile and ice stationAn overview of the distribution of CH4 and N2O grouped into six
geographic regions is given in Table 1. These regionswere: a) the North
Atlantic, b) Labrador Sea—Baffin Bay South, c) Baffin Bay Marginal Ice
Zone, d) Baffin Bay North—Lancaster Sound, e) Canadian Archipelago
and f) Amundsen Gulf—Beaufort Sea. The classification of the entire-
dataset into these regions was based on geographic criteria as well as
ice-cover characteristics. For example, BaffinBaywasbroken down into
three regions based on ice cover (see below and Fig. 2). Surface salinity
showed anoverallwestwarddecrease along the transect (Fig. 2A). High
variability to the west of Baffin Bay and into the Beaufort Sea reflects
dilution by freshwater. A broadly inverse relationship between salinity
and silicate (Fig. 2A and B) identifies this variability to result mostly
from river discharge rather than localised icemelting; note for example
the exceptionally low salinity and high silicate (N6.8 μmol L-1) in the
vicinity of the Mackenzie River Plume (MRP) 6200–6500 km along the
transect (Figs. 1 and 2).
Dissolved CH4 at this site exceeded atmospheric equilibrium (4.6±
1.0 nmol L−1; 144±30% saturation) and was higher than in the
adjacent Beaufort Sea, consistent with the general observation of high
CH4 super-saturation in river outflow (Upstill-Goddard et al., 2000).
Nevertheless, over the cruise track in general CH4 did not appear to
correlate closelywith either salinity or silicate; evidently river discharge
is not the major control on surface CH4 concentrations throughout this
om the first sample. The east-west cruise track crossed the North Atlantic Ocean, Baffin
ackenzie River plume (MR; in square) and Beaufort Sea. The figure was compiled usingare marked by a white circle and square respectively.

te, C
Sou
ch r
82 V. Kitidis et al. / Marine Chemistry 121 (2010) 80–86region. Rather, the variability in dissolved CH4 along the cruise track
(1.7–19.9 nmol L−1; 58–528% saturation) appears to reflect sea-ice
cover. Highest CH4 was observed under multi-year sea-ice in the
Canadian Archipelago; (N4 m thick) in Peel Sound and the McClintock
Channel at 4462 km (19.9 nmol L−1, 528% saturation), while lowest CH4
was in the Amundsen Gulf at 5814 km: 1.7 nmol L−1, 58% saturation
(Figs. 1 and 2C). Sharp increases in CH4 concentrationswere observed in
themarginal ice zones (MIZs) thatmark the transition fromopenwater-
to-ice cover (e.g. Baffin Bay MIZ at 2578 km: 9.2 nmol L−1, 297%
saturation, Figs. 1 and 2C). By contrast, for the most part surface N2O
was relatively mildly supersaturated but with some under-saturation.
Pronounced maxima under multi-year sea ice (33.5 nmol L−1; 181%
saturation) and in the Baffin Bay MIZ (24.9 nmol L−1, 172% saturation)
coincided with those for CH4 (Fig. 2C,D). However at other MIZ's with
high CH4, N2O maxima were generally absent. The lowest N2O
concentration was in the North Atlantic at 517 km (10.7 nmol L−1, 91%
saturation) whereas the lowest N2O saturation was in the McClintock
Channel at 4842 km: 14.7 nmol L−1, 82% saturation (Fig. 2D); the lower
seawater temperature leading to higher solubility.
The depth profile from Baffin Bay (Fig. 3) revealed strong near
surface gradients in CH4 and N2O, with maxima within the upper
150 m. The highest concentrations of both were in a sample from13 m
(CH4: 8.8 nmol L−1, 266% saturation; N2O: 28.1 nmol L−1, 179%
saturation); however the vertical sampling resolution precludes
identifying the location and magnitudes of these maxima more
precisely. The near surface distributions are consistent with slight
under-saturation of CH4 and N2O in surface melt-water and their
production within the upper mixed layer via methanogenesis and
Table 1
Average and standard deviation values for each parameter (Temperature, Salinity, Silica
South (Lab. S-BBS), Baffin Bay Marginal Ice Zone (BBMIZ), Baffin Bay North and Lancaster
Beauf. S). The range in cumulative distance along the track shown in Fig. 1 is given for ea
atmospheric equilibrium.
Region Distance
(×103 km)
Temp.
(°C)
Salinity
N. Atl. 0–1.1 10.0 35.0
±0.9 ±0.1
Lab.S-BBS 1.1–2.5 3.8 33.2
±1.7 ±0.7
BBMIZ 2.5–3.3 −0.3 30.7
±1.3 ±1.1
BBN-LS 3.3–3.9 3.1 30.8
±0.8 ±0.8
Can-Arch 3.9–5.6 −1.0 27.3
±0.7 ±2.8
AG-Beauf. S 5.6–6.7 3.9 23.3
±2.4 ±5.8nitrification. We previously observed significant N2O under-saturation
in the Bellingshausen Sea following seasonal sea ice retreat (Rees et al.,
1997). A sink for N2O is apparent at around 180 m but below this N2O
gradually increases, with depth becoming supersaturated below500 m
and thus indicative of net N2O production in deeper water. By contrast
CH4 consistently decreased below 13 m although it remained signifi-
cantly supersaturated down to 918 m (Fig. 3).
The incubations of under-ice surface seawater indicated CH4
consumption whereas the melt-pond results were inconclusive (Fig. 4).
By contrast N2O did not change significantly in either incubation.
Consumption of CH4 in the under-ice surface seawater apparently
followed first-order kinetics; we estimate a corresponding rate constant,
corrected for 15-fold pre-concentration of the microbial community, of
3.8×10−3 h−1.
4. Discussion
The variability in dissolved CH4 and N2O along the ∼6750 km cruise
track must reflect biological and physical source-sink imbalances andtheir interactions. The observed CH4 super-saturations across large areas
are consistent with findings from less extensive prior surveys in the East
Siberian and Beaufort Seas (Kvenvolden et al., 1993; Shakhova and
Semiletov, 2007). Relatively high CH4 in the vicinity of MIZs, may reflect
high biological activity resulting in the accumulation of particulate
organic matter which can act as a substrate for methanogenesis.
Methanogenesis during zooplankton grazing has been known for some
time (DeAngelis and Lee, 1994) and elevated CH4 close to Spitzbergen
was attributed to post phytoplankton bloom methanogenesis (Damm
et al., 2008). Evidence for a shallowwater columnCH4 source consistent
with methanogenesis was seen in the depth profile from Baffin Bay
(Fig. 3). An additional or alternative origin for the CH4 may involve its
release from the seabed. However, the decreasing concentration of CH4
with increasing depth from 13 m to within 30 m off the seafloor
suggests that benthic sources are of minor importance at this location.
Nevertheless, it is also possible that a CH4 plume originating from
nearby sediments was advected into surface waters at this location
leading to CH4 accumulation under sea-ice. Subsequent sea-ice retreat
may have led to loss of CH4 to the atmosphere in the uppermost surface
layer shortly before sampling. High CH4 concentrations in the vicinity of
the Beaufort Sea MIZ were previously linked to sedimentary sources
(Kvenvolden et al., 1993; Macdonald, 1976) and high dissolved CH4 in
the East Siberian Sea was ascribed to methanogenesis in sub-seafloor
permafrost (Shakhova and Semiletov, 2007; Shakhova et al., 2005). In
the region of the Beaufort Sea MIZ sub-seafloor shallow gas and/or
hydrate deposits are another potential CH4 source (Judd and Hovland,
2007).
In contrast to the generally strong CH super-saturation observed in
H4 and N2O) measured by region: North Atlantic (N. Atl.), Labrador Sea and Baffin Bay
nd (BBN-LS), Canadian Archipelago (Can. Arch) and Amundsen Gulf-Beaufort Sea (AG-
egion. CH4 and N2O data are given in concentration and as percent saturation relative to
SiO42−
(μmol L−1)
CH4
(nmol (L−1); % sat
N2O
(nmol (L−1); % sat
nd 4.2; 165 11.5; 1.3
±1.2; 48 ±1.1; 7
1.2 4.2; 139 15.4; 109
±0.3 ±1.3; 45 ±1.6; 9
3.2 5.9; 170 19.8; 117
±1.2 ±1.6; 51 ±3.1; 24
1.5 5.3; 164 17.5; 116
±0.7 ±2.1; 65 ±1.6; 12
2.6 8.2; 226 21.0; 117
±0.9 ±3.5; 94 ±2.8; 15
3.7 4.9; 151 17.0; 115
±1.9 ±2.8; 76 ±1.1; 74
all MIZs, high N2O only occurred in the Baffin Bay MIZ. This most likely
reflects a nitrification source in near surface waters which is further
supported by the near surface N2O maximum in the depth profile from
Baffin Bay (Fig. 3). Nevertheless, the mineralisation of organic matter
derived from sea-ice algae has been shown to result in substantial
denitrification activity in anaerobic micro-niches in the sea-ice matrix
(Rysgaard and Glud, 2004). We can therefore not exclude the
possibility that denitrification also contributed to surface water N2O
super-saturation here. Sediment denitrification likely contributed to
deeper water N2O, again as reflected in the Baffin Bay depth profile
(Fig. 3). Denitrification activity has been identified in sediments of the
Chukchi Sea (Devol et al., 1997) and in sediments off Greenland
(Rysgaard et al., 2004). Furthermore, depth profiles and the N stable
isotope signature of N2O in the Bering and Chukchi Seas to the west of
our sampling transect, suggest that sediment denitrification is a
substantial source of N2O in thesewaters (Hirota et al., 2009).Whether
or not the apparent decoupling of surface N2O and CH4 concentrations
in other MIZs reflects insufficient accumulation of NH4+ for nitrification
for example, must remain speculative at present and overall the

83V. Kitidis et al. / Marine Chemistry 121 (2010) 80–86relationship between N2O and other Arctic Ocean N cycle components
requires more detailed study.
For CH4, microbial oxidation is a major sink that moderates water-
to-air CH4 emissions in sub-Arctic marine waters, the North Atlantic
and Labrador Sea (Niemann et al., 2005, 2006; Rehder et al., 1999;
Treude et al., 2003). Our incubation experiments with seawater
collected from beneath the ice (Fig. 4) gave a first-order rate constant
for CH4 oxidation (−3.8×10−3 h−1)whichwashigher than inprevious
experiments in the coastal Beaufort Sea: average −0.4×10−4 h−1;
range: 0.0 to −4.2×10−4 h−1 (Lorenson and Kvenvolden, 1995).
Rehder et al. (1999) estimatedmuch lower rate constants equivalent to
−2.3×10−6 h−1 for the North Atlantic and Labrador Sea. The
discrepancy between these results may be due to the use of different
methodologies: a) direct CH4 consumption measured here versus
measured CO2 and biomass production from isotopically labelled CH4
Fig. 2. The distribution of sea-surface temperature (SST) and Salinity (A.); dissolved
Silicate (B.), CH4 saturation with respect to atmospheric equilibrium and sea-ice cover
(C.) andN2O saturationwith respect to atmospheric equilibrium and sea-ice cover (D.) in
underway surface samples against cumulative distance along the sampling track shown
in Fig. 1. The solid line in panels (C.) and (D.) represents 100% saturation with respect
to atmospheric equilibrium. The Mackenzie River plume (MRP) and marginal ice zone
in Baffin Bay (MIZ) showed elevated super-saturation with respect to atmospheric
equilibrium for both CH4 and N2O.uptake (Lorenson and Kvenvolden, 1995) or apparent CH4 utilisation
relative to a conservative tracer, CCl3F (Rehder et al., 1999); b)
timescales concerned (hours/days here vs. decades—e.g. Rehder et al.,
1999) or c) sample manipulation. Lorenson and Kvenvolden (1995)
removed “(via shaking) most of the CH4 in” their samples prior to
replacing this with labelled CH4 and incubating the collected seawater.
Vigorous shaking may have led to mechanical damage of the
methanotrophic community and hence explain lower methanotrophy
rates than those observedhere. On the other hand, pre-concentration of
the planktonic community here may have enhanced methanotrophy
disproportionately to the seawater volume concentration factor.
Nevertheless, the CH4 oxidation rate constant observed here is lower
than the respective rate constant we previously found in temperate
marine environments, up to −2.0×10−2 h−1 (Kitidis et al., 2007); as
expected given the lower seawater temperature in Arctic waters. We
did not detect CH4 oxidation in the melt-pond sample (Fig. 4). This is
intriguing because melt ponds are often highly productive environ-
ments with visible colouration from algal pigments (Thomas and
Dieckmann, 2002). However,wedid not observe such colouration here,
andwenote that dissolved CH4 (andN2O)was close to the atmospheric
equilibriumat this site. These observations lead us to speculate that this
particular melt-pond had zero or very low productivity, and that
consequently the supply of CH4 to the potential oxidation pathwaywas
negligible.
The role of sea ice in the distribution and sea–air exchange of CH4
and N2O is clearly important. It was previously suggested that the
accumulation of CH4 under-saturated surface seawater due to sea-ice
melting can result in a significant drawdown of atmospheric CH4
(Heeschen et al., 2005), and previous work in the Beaufort Sea showed
higher surface seawater CH4 concentrations under sea ice than in open
water (Kvenvolden et al., 1993). Similarly, N2O accumulation under
sea-ice and its subsequent release to the atmosphere during sea-ice
melting in the Bellingshausen Sea was proposed by Rees et al (1997)
and also for the Sea of Okhotsk (Lammers et al., 1995). Such a
mechanism is entirely consistent with the under-saturation we
observed for both CH4 and N2O and especially the depth profile from
Baffin Bay which showed a surface salinity minimum and under-
saturation of CH4 and N2O at 7 m depth (Fig. 3). While ice meltingmay
represent a significant sink for CH4 andN2O in openwater, sea-ice itself
may reduce turbulent diffusivity in the water column (Carmack and
Chapman, 2003) and inhibit sea–air gas exchange. Sea-ice may
therefore increase the concentrations and residence times of CH4 and
N2O in the under-icewater column. Similarly, sea-icewas found to play
a critical role in preventing sea–air gas exchange of CO2 during our
cruise (Fransson et al., 2009). Consequently we observed the highest
CH4 andN2O concentrationsundermulti-year sea-ice inPeel Sound and
in the McClintock Channel (Figs. 1 and 2). One important consequence
of the enhanced residence time for CH4 under sea-ice is that potentially
more of it is available for utilisation by the microbial community,
consistent with the results of our incubation experiments in under-ice
seawater (Fig. 4).
To illustrate the potential significance of sea-ice to the degree of CH4
partitioning between the microbial oxidation and sea-to-air exchange
sinks, we consider a hypothetical surface mixed layer of 10 m depth
with a temperature of−0.3 °C, a salinity of 28 and a CH4 concentration
of 7.4 nmol L−1. With the exception of the mixed layer depth, these
values were averages measured during the Beringia 2005 cruise when
sea-ice was present and are therefore representative of ambient
conditions. Fig. 3 shows that the estimated mixed layer depth of 10 m
is realistic, if not a slight over-estimate. Our calculations demonstrate
that if themixed layer depth is indeed shallower than 10 m, this would
decrease the fraction of CH4 oxidised and thereby strengthen our
conclusion (see below). The 10 m mixed layer depth is therefore a
conservative estimate. We consider only the sea–air exchange and
microbial oxidation; loss terms; for the purposes of this exercise any
mixing of CH4 across the base of the mixed layer is neglected and we

Fig. 3. The distribution of CH4 and N2O saturationwith respect to atmospheric equilibrium (A.), salinity and temperature (C.) in the Baffin Bay depth profile. Surface detail is shown in
panels (B.) and (D.), respectively.
84 V. Kitidis et al. / Marine Chemistry 121 (2010) 80–86also ignore the possibility of diffusional CH4 transfer through the ice.
We specify a first-order microbial oxidation rate (FMO) with a rate
constant of−3.8×10−3 h−1 and parameterise a sea–air gas exchange
rate (FSA; μmol m−2 h−1) according to:FSA = kw × Δ CH4½  ð1Þ
Fig. 4. CH4 and N2O concentrations during incubation experiments with surface seawater (A
show the respective concentrations in equilibrium with the atmosphere.where kw is the gas transfer velocity and Δ[CH4] is the sea–air CH4
concentration difference. kw was calculated according to a wind-
driven turbulent diffusion model (Nightingale et al., 2000):
kw = 0:333 × u + 0:222 × u
2 ×
Sc

0:5
ð2Þ
660
and C) and melt-pond water collected during an ice station (B and D). The dashed lines

methane supply from the Sea-of-Okhotsk induced by the seasonal ice cover. Global
Biogeochemical Cycles 9 (3), 351–358.
85V. Kitidis et al. / Marine Chemistry 121 (2010) 80–86where u is the wind speed at 10 m above sea level and Sc is the
Schmidt number for CH4 (2076 at−0.3 °C). In our model u was 6.2±
3.1 m s−1 in the absence of sea-ice (the measured average value from
the ship's meteorological package situated at 35 m above sea level and
corrected for height) while in the presence of sea-ice u was 0 ms−1.
The sea–air gas exchange term in this model was thus equal to zero in
the presence of sea-ice. Our model shows that in the absence of sea-
ice, sea–air exchange was the dominant sink for CH4 removing 54% of
the CH4 in excess of atmospheric equilibrium over 72 h while CH4
oxidation accounted for 37% (9% excess CH4 remained in the water
after 72 h). In the presence of sea-ice the corresponding values for
sea–air exchange and CH4 oxidation were 0% and 46% respectively. In
the absence of sea–air loss (with sea-ice), CH4 concentration in the
hypothetical mixed layer was higher than in the ice-free model and
therefore CH4 oxidation increased since this is a first-order process
with respect to CH4 concentration.
This simplified treatment does not consider potential loss terms
such as sea–air gas exchange through sea-ice brine channels or a
biological CH4 sink in sea-ice. Such loss terms have been shown for CO2
(Gosink et al., 1976; Semiletov et al., 2004), but remain unexplored in
the case of CH4. The model clearly demonstrates the potential of a
changingCH4 residence timeunder sea ice in influencing the fate of CH4
innear surface seawater. Given that progressive seasonal sea-ice retreat
is an established recent phenomenon in the Arctic (Holland et al., 2006;
Stroeve et al., 2007a)we should expect accelerated sea ice retreat year-
on-year to result in increasing CH4 emissions to the atmosphere in
future. CH4 will thereby be oxidised in the atmosphere and contribute
to the greenhouse effect rather than undergomicrobial oxidation in the
water column. We may thus be witnessing another example of a
positive climate feedback loop, in which ice melt through warming
stimulates further warming through greenhouse gas emissions. Given
that the radiative forcing potential of CH4 is ∼20 times that of its
microbial oxidation product CO2, the amplification effect is significant.
Furthermore, although the super-saturations we observed for N2O
were comparatively modest, its high radiative forcing potential, ∼300
times that of CO2, amplifies its warming contribution ten-fold relative
to CH4. Nevertheless, a positive climate feedback loop betweenmelting
sea-ice and increased CH4 emission to the atmosphere remains
speculative due to the simplified nature of our conceptual model. We
thus recommend that futurework in the Arctic should involve efforts to
establish the detailed spatio-temporal surface distributions of CH4 and
N2O and their controlling process rates, so that the feedback effect
identified here may be more fully evaluated.
5. Conclusions
Elevated CH4 and N2O concentrations with respect to atmospheric
equilibrium were found along the ∼6750 km cruise track through the
Canadian Archipelago, indicating substantial sources for both gases in
these waters and from the Arctic Ocean to the atmosphere. Sea-ice
cover, marginal ice zones and river plumes appear to play a dominant
role indetermining the variability in the concentrations of CH4 andN2O.
Importantly, sea-ice melting may be an important, transient sink for
atmospheric CH4 and N2O in agreement with previous work. Microbial
oxidation is an important sink for CH4 in the Arctic Ocean. Sea-ice cover
may moderate sea–air exchange and thereby increase the residence
time of CH4 in surface waters, potentially allowing a larger fraction of
CH4 to be oxidised by the microbial community.
Acknowledgements
The authors would like to thank Dr. R. Torres for help with sea-ice
data, Dr. A. Fransson and Prof. K. Abrahamsson for their help and
support during the cruise, the Swedish Polar Research Secretariat as
well as the captain and crewof IBOden.Wewould also like to thank two
anonymous reviewers for their constructive remarks which haveLorenson, T.D., Kvenvolden, K.A., 1995. Methane in coastal seawater, sea ice and bottom
sediments, Beaufort Sea, Alaska. US Geological Survey, Menlo Park, CA.
Macdonald, R.W., 1976. Distribution of low-molecular-weight hydrocarbons in theimproved this manuscript. This work was supported by the UK Natural
Environment Research Council (grant number: NE/C000706/1).
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