Methane production , consumption and its carbon isotope ratios in the Southern Ocean during the austral summer

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Introduction
Considerable attention has recently been focused on biogenic trace gases in ecosystems because they include a significant amount of greenhouse gases.Methane (CH 4 ) is an effective greenhouse gas that is approximately 20 times more effective in radiative forcing than CO 2 on a per mole basis and contributes significantly to global warming (Schneider, 1989).The atmospheric concentration of CH 4 is currently increasing faster than that of other biogenic greenhouse gases (Pearman and Fraser, 1988;Bouwman, 1990).Each year, 50 million tons of CH 4 are added to the atmosphere, resulting in a relative annual increase of more than 1% (Bouwman, 1990).The primary sources of CH 4 include wetlands, paddy fields, intestinal fermentation, and pyrolysis of organic matter, i.e., biomass burning (Cicerone and Oremland, 1988).However, methane produced in marine environments also contributes to atmospheric greenhouse gas concentrations and organic carbon cycle pathways.Furthermore, the isotopic signature of CH 4 is recognized as providing constraints on relative source strength and information on reaction dynamics concerning relevant formation, destruction, and biological pathways.Introduction

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Full The Southern Ocean, one of the most productive water bodies in the world, has some characteristics unique among the world's oceans, i.e., the near-circular symmetry of most of its physical and chemical characteristics; the immense expanse of cold surface water, which is globally by far the largest; and the most pronounced seasonal variation in sea surface temperature of any comparable ocean basin (Lutjeharms, 1990).The temporal and spatial distribution of these physical and chemical variables should also be reflected in the distribution patterns of biological processes.Here we describe the results of isotopic studies of dissolved CH 4 in the Southern Ocean at 140 • E to elucidate the origins of CH 4 in the ocean and to estimate its CH 4 contribution to the atmosphere.

Chemical analysis
Salinity and temperature were recorded using a CTD sampler (SBE 911plus).Dissolved oxygen was measured with a dissolved oxygen auto-titrator based on the Scripps Institute of Oceanography design.

Dissolved CH 4 and carbon isotope analysis
In the laboratory, water samples were transferred to a stripping chamber using helium gas replacement in the sample vials.After H 2 O and CO 2 were removed in a magnesium perchlorate [Mg(ClO 4 ) 2 ] and Ascarite (sodium hydroxide-coated silica) column, CH 4 was collected in a U-shaped tube filled with activated charcoal (AC1) at liquid nitrogen temperature.At the end of the quantitative extraction of CH 4 , the U-tube was heated to room temperature (25 • C) by removing a liquid nitrogen trap.The desorbed gas from the AC1 was separated in a molecular sieve column.The purified CH 4 fraction was collected again in an activated charcoal (AC2) tube at liquid nitrogen temperature.The AC2 was then heated to room temperature by removal of the liquid nitrogen.The desorbed CH 4 was cryofocused before being introduced into the system GC and then separated in a PoraPlot Q column (25 m) at 30 • C. The GC was connected to a combustion furnace made of quartz (0.6 mm ×15 cm) filled with CuO at 960 • C. Methane was then combusted into CO 2 and H 2 O, and the produced H 2 O was removed by diffusion through a Nafion membrane tube.He-carrier gas was then partially separated using a helium separator, and CO 2 was introduced into the isotope ratio mass spectrometer (Finnigan, MAT 252).The carbon isotopic composition results were expressed as δ 13 C values, defined as: where R SAMPLE and R STANDARD are the isotope ratios ( 13 C/ 12 C) for samples and standards, respectively.Carbon isotope ratios of dissolved CH 4 were expressed in ‰ de- to within ± 0.3‰.The atmospheric equilibrium concentration of CH 4 was calculated from in situ water temperatures and the atmospheric mixing ratio (1.7 ppmv; Quay et al., 1991;Holmes et al., 2000) following the methods of Yamamoto et al. (1976).

Results and discussion
Methane concentrations and isotope ratio profiles in the Southern Ocean water column during the austral summer (Fig. 2) revealed that methane distribution and production were related to physical and biological properties in the area was observed in the subsurface (at a depth of approximately 125 m) and decreased to a constant value below 1000 m at station 1.The δ 13 C value at this station was generally enriched below 200 m, in contrast to the CH 4 concentration.However, small enrichments in δ 13 C were also observed in the CH 4 maximum layer.While CH 4 at station 5 exhibited no subsurface maximum, the CH 4 concentration gradually decreased with increased δ 13 C from the surface and reached a constant value below 500 m.At station 8, the CH 4 concentration and δ 13 C value remained close to atmospheric equilibrium from the surface to the deep region.
As CH 4 is produced and/or oxidized by bacteria and also lost to the atmosphere through gas exchange, the shape of a vertical profile must be determined by the balance between in situ biological production and consumption and physical processes of diffusion, advection, and gas exchange.Thus, dissolved CH 4 and its isotope ratio could be categorized into three parts: surface, subsurface (including the CH 4 maximum), and deeper regions.
At the surface where gas exchange occurs, the CH 4 concentration is governed by physical factors, mainly temperature and wind, with gas saturation values up to about Introduction

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Full 5% above or below saturation (Bieri et al., 1966(Bieri et al., , 1968)).The surface equilibrium of CH 4 in the Southern Ocean could be caused by high wind speeds, given that temperatures in the area are quite stable during the austral summer.The δ 13 C values of surface dissolved CH 4 at all stations were enriched by 0.2-1.2‰relative to the atmospheric value (−47 to −47.4‰; Holmes et al., 2000;Stevens, 1993), possibly because of the kinetic fractionation factors during CH 4 invasion and evasion.Fuex (1980) found that the equilibrium fractionation (α eq ) for methane was about 1.00033, a value similar to our result and those observed in other areas (e.g., Holmes et al., 2000;Tsunogai et al., 2000;Valentine et al., 2001).The δ 13 C of dissolved CH 4 , however, was enriched by more than 1‰, a value greater than can be attributed to analytical error at station 1 where wind speeds were high and were expected to result in a high CH 4 exchange rate with surface water.This result suggests that the exchange of CH 4 between ocean and atmosphere and its isotopic fractionation should be treated kinetically and considered temporally and spatially variable, as well as related to high in situ δ 13 C(CH 4 ), even at the surface.
Given that the microbial production of CH 4 cannot occur in an oxic environment (Wolfe, 1971), it is thought that CH 4 forms mostly within the reducing interiors of particles, as supported by incubation experiments (Karl and Tilbrook, 1994;Marty et al., 1997;Owens et al., 1991).In the Southern Ocean waters surrounding Antarctica, phytoplankton blooms fueled by nutrient-rich waters lead to the growth of vast swarms of krill (Euphausia superba) each summer.Consequently, settling particles, such as fecal pellets, may be produced during the life cycle of the krill, either directly or indirectly.Chlorophyll-a is the most important phytopigment and can be used as a semiquantitative measure of phytoplankton abundance.Hirawake et al. (2003) reported that chlorophyll-a concentrations show clear seasonal variation with a complicated spatial distribution; concentrations were very low (< 0.3 mg m −3 ) in October, but rapidly increased to 1-5 mg m −3 through December near the sea-ice edge and in the open ocean.Furthermore, these particles do not consist only of inorganic or dead organic matter; they also host microbial communities, including bacteria, various protozoa, and Introduction

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Major substrates for methanogenic bacteria include acetate and CO 2 /H 2 (Cicerone and Oremland, 1988;Whiticar et al., 1986).In the former substrate, CH 4 is derived from the methyl group of acetate, and this process depends on the production of acetate at anaerobic sites during decomposition of organic matter.The carbon isotope fractionation factor for reduction of CO 2 to CH 4 has been reported for a pure culture system, and the high fractionation factor of 1.045 was reported for wetlands rich in labile organic matter (Games et al., 1978).The general trend that governs the variation in δ 13 C of bacterial methane has been examined through paddy soil incubation experiments.The δ 13 C value of CH 4 from acetate was close to that of methyl carbon in acetate (−43 to −30‰), while that from CO 2 /H 2 reduction was −70 to −60‰ (Sugimoto and Wada, 1993).The δ 13 C(CH 4 ) value at the ∆CH 4 maximum, i.e., about −43.4‰ at station 1, suggested that the formation of CH 4 in this subsurface was produced via acetate fermentation.Blair and Carter (1992) reported the δ 13 C value of acetate and estimated the δ 13 C of CH 4 derived from acetate to be −43‰ in anoxic marine sediment, which agrees well with our results.However, the large δ 13 C variation in settling organic matter in the Antarctic Ocean (−26.7 to −20.5‰;Wada et al., 1987) may affect the δ 13 C of CH 4 produced via acetate fermentation.The oxidation of the methyl position of acetate to CO 2 may also cause the isotopic fractionation of the produced CH 4 (Whiticar et al., 1986).However, the ∆CH 4 maximum peak was only observed at station 1, suggesting that the magnitude of CH 4 production varied with abundance, distribution, and migration of phytoplankton and zooplankton in the area.
In the deeper subsurface waters, the ∆CH 4 value became negative in association with 13 C enrichment, suggesting the oxidation of dissolved CH 4 with less or without new CH 4 production.Thus, the residual, heavier δ 13 C CH 4 should be found based on the kinetic isotope effect during oxidation.Several researchers have demonstrated this process with aerobic culture systems and have found that the magnitude of carbon Introduction

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Full isotope fractionation varies from 5 to 31‰ (e.g., Barker and Fritz, 1981;Zyakun et al., 1984).Such water column CH 4 oxidation has also been observed in other areas, such as at the ALOHA station and in the North Atlantic (Holmes et al., 2000), suggesting that the oxidative consumption could occur normally in oxic water column environments.
Other parameters that could indicate CH 4 oxidation include the preferential loss of CH 4 relative to higher molecular weight hydrocarbons (Whiticar, 1999) and the systematic shift in H isotope ratios (Coleman et al., 1981;Whiticar, 1999).Furthermore, in deep water, at approximately 2500 m, the horizontal and vertical movements of cold water from inshore induced by the geopotential anomaly of the strong steering of the current by the ridge system around Antarctica (see Lutjeharms, 1990;Orsi et al., 1995) should affect the δ 13 C level of CH 4 .The ∆CH 4 values did not decrease below 500 m and 1000 m at stations 5 and 1, respectively, probably because of the lack of CH 4 consumption in deeper water (Scranton and Brewer, 1978).Also, the isotope ratio varies with depth because of the mixing of different water masses with different histories of CH 4 input (Holmes et al., 2000).It should be noted that at station 8, located in a shallow water column close to shore, the strong steering of the summer current may promote vertical homogeneity of the CH 4 concentrations and isotopic compositions.
The degassing of supersaturated surface seawater could release CH 4 and sources of isotopically heavy CH 4 into the earth's atmosphere by simple diffusion and thereby contribute to the net flux of CH 4 across the sea-air interface.It is premature to quantitatively calculate the influence of the Southern Ocean on the global CH 4 budget from the present data alone; however, our results do provide some insight into the contribution of the Southern Ocean.The estimated amount of CH 4 that will eventually reach the atmosphere after transport through the unsaturated zone can be determined from the empirical relationship between wind speed and gas transfer rate with measured values for wind speed, temperature, and surface concentrations of CH 4 , according to the following equation (Liss and Slater, 1974): Tables Figures

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Full where K L is the liquid-phase gas-transfer coefficient, C S is the gas concentration in the surface water; and C E is the gas concentration of the solution during equilibrium with the overlying gas phase.The liquid-phase gas-transfer coefficient, K L , used in calculating the flux was obtained from Liss and Merlivat (1986): and Wanninkhof (1992): where V is wind velocity (m s −1 ), and Sc is the Schmidt number.The Schmidt number for CH 4 in seawater is expressed as a decreasing function of temperature (Wanninkhof, 1992): The calculated net sea-air flux of CH 4 in the Southern Ocean is an average of 0.32 µmol m −2 d −1 (−0.09 to 0.74 µmol m −2 d −1 ), using the Liss and Merlivat (1986) and Wanninkhof (1992)  with sea-air exchange because of the long turnover time (an average of 65 yr; Jones, 1991;Ward et al., 1987).The CH 4 flux in this area was much higher than that reported for Drake Passage and the Weddell Sea during the austral summer and autumn, respectively (Heeschen et al., 2004;Tilbrook and Karl, 1994); both locations showed negative flux values (Table 1).However, positive flux values have been found in coastal areas with CH 4 accumulation (Tilbrook and Karl, 1994).This could suggest that sea-air flux of CH 4 varies greatly in time and space depending on physical factors as well as the magnitude of CH 4 diffusion from the CH 4 maximum layer.If this flux value could be applied to the entire Southern Ocean, which has an area of about 37 × 10 12 m 2 (Peng, 1984), for the duration of the austral summer, the total CH 4 flux to the atmosphere from this region would be 0.19 Gg d −1 .The value of δ 13 C for the CH 4 flux to the atmosphere was calculated using the following equation from Quay et al. (1993) and Hoefs (1987): where R atm is the 13 C/ 12 C ratio of atmospheric CH 4 , R ml is the 13 C/ 12 C ratio of CH 4 in the mixed layer, α k is 0.9992 (Knox et al., 1992;Happell et al., 1995), α sol is 1.00033 (Fuex, 1980), pCH 4 in the methane partial pressure, δ 13 C CH 4[atm] is −47.0 to −47.4‰ (Quay et al., 1991;Holmes et al., 2000), and R PDB , the isotopic ratio of the PDB standard, is 0.011237 (Craig, 1957).We determined that the average δ 13 C value supplied to the atmosphere in the CH 4 flux from this area was −42.7‰ (Table 1).This contribution (approx.0.19 Gg d −1 ) is a significant part of the global oceanic CH 4 flux of 5-50 Tg yr −1 (2.4-24 µmol m −2 d −1 ; Prather, 1995).Thus, the Southern Ocean could be contributing 0.14-1.39% of the total oceanic flux of CH 4 .However, if we consider other results from the circum-Antarctic region (Heeschen et al., 2004;Tilbrook and Karl, 1994), the total net sea-air CH 4 flux from the Southern Ocean would be about 0.15 Tg yr −1 , equivalent to 0.30-3.0% of the total oceanic CH 4 contribution.Moreover,

BGD Introduction
Full Discussion Paper | Discussion Paper | Discussion Paper | were collected during the 43rd Japanese Antarctic Research Expedition (JARE-43), 2002 Marine Science Cruise on the R/V Tangaroa from 6 February to 7 March 2002.The purpose of the expedition was to study the biogeochemical cycles and biological processes of the Southern Ocean as they relate to global environmental issues.The production and consumption of dissolved CH 4 in Southern Ocean ecosystems were investigated by collecting seawater samples at stations 1 (open water), 5 (marginal ice zone), and 8 (previous ice zone; Fig. 1).Water samples were collected at the indicated depths using a CTD water sampler (SBE 32 24x10-L Carousel Water Sampler).For CH 4 analyses, water samples were put in 125-ml glass vials and sterilized with mercury chloride (1 ml of saturated HgCl 2 solution per vial).The vials were then sealed with a butyl-rubber septum and an aluminum cap, taking care to avoid bubble formation, and stored at 4 • C in the laboratory until analysis.Discussion Paper | Discussion Paper | Discussion Paper | viations from PeeDee belemnite (PDB) carbonate.The δ 13 C values were reproducible Discussion Paper | Discussion Paper | Discussion Paper | Screen / Esc Printer-friendly Version Interactive Discussion Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | transfer coefficients.The ∆CH 4 maximum observed at the base of the mixed layer would result in a diffusive flux of CH 4 into the surface layer.We also obtained δ 13 C of about −41.8‰ for the source of CH 4 formation from the isotopic correlation with the inverse concentration between the surface and the CH 4 maximum and estimated the vertical net flux to be about 0.26 µmol m −2 d −1 from the one-dimensional vertical mixing coefficient (Li et al., 1984) at station 1.This input accounted for about 80% of the estimated net flux of CH 4 in the Southern Ocean to the atmosphere and suggested that CH 4 was produced by in situ biological processes, diffused through the mixed layer, and then released by sea-air exchange.The consumption loss rate of 0.01 µmol m −2 d −1 , calculated from an average CH 4 concentration in the upper 70 m and an average CH 4 turnover time, was negligible compared to the losses associated Discussion Paper | Discussion Paper | Discussion Paper | Figure1
. CH 4 concentrations and δ 13 C values of the dissolved CH 4 in surface water were almost in equilibrium with the atmosphere, with mean excess CH 4 (∆CH 4 ) and δ 13 C at 0.06 nM (or 102% saturation) and −46.3‰, respectively.The ∆CH 4 maximum (2.94 nM or about 200% saturation)