Records of the δ¹³C of atmospheric CH₄ over the last 2 centuries as recorded in Antarctic snow and ice

1.1. Previous Studies of Recent Atmospheric Methane Trends

[3] Ice cores provide the primary means of reconstructing the composition of past atmospheres. High-resolution atmospheric records have been reconstructed for the last 1000 years from the Law Dome ice cores [Etheridge et al., 1998]. Methane concentration records from Law Dome have documented atmospheric loadings that have increased in an exponential fashion starting from ∼670 ppb in 1000 A.D. to ∼1600 ppb in 1970 A.D. The major cause of this increase is thought to be the result of anthropogenic emissions (rice, cattle, biomass burning, and fossil fuels) with a smaller contribution from a net reduction in the primary atmospheric sink (tropospheric OH).

[4] To date, only one set of ice core analyses have been made to estimate the δ¹³C of paleoatmospheric CH₄. Craig et al. [1988] made δ¹³CH₄ measurements on large ice samples from the Crete ice core (central Greenland) and concluded that atmospheric δ¹³CH₄ increased by ∼2‰ over the past 2 centuries. Unfortunately, Craig et al. [1988] failed to account for a number of important factors (gas age–ice age difference, gravitational and thermal fractionation, and the diffusional impact of increasing [CH₄] on the δ¹³CH₄ of firn air CH₄) making their estimates of the net change in atmospheric δ¹³CH₄ uncertain though the sense of the change is probably correct.

[5] Analyses of air from the interstitial spaces around snow near the surface of polar ice sheets (firn air) provides the means of splicing ice core records with direct measurements of the troposphere which began in the 1950s. Firn air δ¹³CH₄ measurements indicate a positive recent trend of 0.03–0.10‰ per year for the later part of the twentieth century [Bräunlich et al., 2001; Francey et al., 1999]. By way of comparison, the current peak-to-peak amplitude of the seasonal δ¹³CH₄ variations range from 0.1 to 0.4‰ in both hemispheres while the pole-to-pole difference is ∼0.5‰ with a tendency toward larger gradients during the boreal winter [Lowe et al., 1999; Platt et al., 2004; Quay et al., 1999; Rice et al., 2001]. Since loss of CH₄ is controlled primarily by reaction with OH in the troposphere, the potential influence of isotope effects associated with loss processes in the stratosphere has been omitted in a number of budget studies [e.g., Craig et al., 1988; Francey et al., 1999]. Gupta et al. [1996] pointed out, however, that the increase in the stratospheric Cl (with concomitant CH₄ removal) should have raised tropospheric δ¹³CH₄ by as much as 0.7‰ over the past century. Estimates suggest that stratospheric Cl loading may have increased by ∼200% owing to the introduction of chlorine containing freons [Flocke et al., 1999]. The magnitude of the Cl impact on tropospheric δ¹³CH₄ has been recently reassessed indicating the net impact is probably between 0.2 and 0.5‰ [McCarthy et al., 2003; Wang et al., 2002].

[6] Recently, measurements of firn air from Dome C and Dronning Maud Land have yielded new δ¹³CH₄ data that extend the atmospheric record back to ∼1950 A.D. Results indicate the δ¹³CH₄ in the far Southern Hemisphere has increased by ∼1.7‰ since 1950 A.D. The trend toward higher δ¹³CH₄ was thought to be the result of increased CH₄ emissions from biomass burning and fossil fuel consumption [Bräunlich et al., 2001].

[7] The present contribution extends the atmospheric record to the early 1900s with analyses of firn air from the South Pole that were drawn in January 1995 and 2001 [Battle et al., 1996]. Additionally, samples from a shallow ice core from the Siple Dome area are used to establish the δ¹³CH₄ value during the nineteenth century. Moreover, the two separate firn air samplings provide an opportunity to establish the evolution of the firn air profiles over a 6-year period from the same locale. In the sections that follow, we initially discuss how firn air from South Pole has been sampled and used to reconstruct atmospheric records over the past century. We then discuss the methods we used to measure δ¹³CH₄ and present results from two separate South Pole sampling expeditions measured in two labs. In order to establish a preanthropogenic δ¹³CH₄ value, a new technique was developed to extract and analyze the δ¹³CH₄ of air trapped in ice from a shallow core at Siple Dome. Finally, we discuss the significance of the results in terms of estimating historical CH₄ emissions from anthropogenically mediated biomass burning.

1.2. Retrieving Interstitial Air From Polar Snow

[8] The composition of air within the interstitial spaces in polar snowfields is dictated by the historical atmospheric record and the rate at which air mixes in a vertical sense [Schwander et al., 1993]. At the South Pole, the firn/ice transition is located ∼123 m below the surface (mbs) [Battle et al., 1996]. Below 123 mbs, air is occluded in bubbles and can no longer communicate with the overlying atmosphere. Firn air at the South Pole is reasonably well mixed in the upper 2–5 m by wind and atmospheric pressure changes. Below 5 m, the air mixes solely by diffusion down to ∼116 m, corresponding to the top of the “lock-in” zone [Battle et al., 1996]. At this depth, the open porosity decreases to a level where the tortuous nature of the firn no longer allows the air within the lock-in zone to mix with the overlying firn air. The air becomes effectively sealed in the firn although very few bubbles are apparent. Below 116 m, the age of the air increases with depth at a rate that is dictated by the snow accumulation rate (∼8 cm of ice equivalent/year).

[9] Over time, atmospheric compositional changes are propagated down into the firn by diffusion. At the South Pole, the average age of CH₄ molecules just above the bubble close-off region date to ∼1900 A.D. Moving up the firn column, the age of the air decreases rapidly between 123 and 116 m and then more slowly from 116 m to the surface. The composition of the air at any depth in the firn is therefore dependent on the atmospheric history and the diffusion processes controlling the movement of each constituent down into the firn [Battle et al., 1996].

[10] Two separate expeditions were mounted to retrieve firn air from the geographic South Pole. The first expedition occurred in January of 1995 where two holes (30 m apart) were drilled incrementally and sampled at 30 different depths using standard techniques to fill replicate flasks [Battle et al., 1996; Schwander et al., 1993]. In January 2001, a similar expedition was mounted using effectively the same equipment. In 2001, two additional holes were drilled incrementally and sampled at a total of 31 depths. A subset of flasks from both experiments was measured at NOAA/CMDL for trace gas species. Another suite of flasks was measured at Princeton for the elemental and isotopic composition of O₂, N₂, and Ar. The results from these analyses provide the basis for the firn air modeling that we discuss in a later section.

2.1. LGGE Analyses

[11] The δ¹³C measurements at LGGE were made with a Finnigan MAT 252 mass spectrometer coupled to a Finnigan GC/combustion interface in continuous flow mode (CF-IRMS). CH₄ was initially isolated from other air constituents using a preconcentration device. Air samples are initially expanded into a pre-evacuated 150-mL sample loop that is subsequently flushed with pure He carrier gas for 20 min at 70 mL/min. The He carrier stream was passed through a HaySep D column (80/100 mesh, 20 cm length, 0.32cm OD) held at −130°C to trap CH₄ and CO₂ with the remainder of the air/He directed to vent. A six port Valco valve is then switched before the HaySep D column is instantaneously warmed to 50°C to release the trapped CH₄ sample and other condensable gases from the HaySep D column. The CH₄ and CO₂ were then cryofocused onto a 3-m capillary column (fused silica Poraplot Q, 0.32 mm ID) at −130°C. Finally, the focusing column was warmed to 50°C to release the CH₄ and CO₂ into a 30-m Poraplot Q column where the CH₄ was separated from CO₂ and N₂O. The effluent from the Poraplot column was then fed into an oxidation furnace (Ni, Pt, and Cu at 980°C) where CH₄ was converted to CO₂. The sample stream then passes through a reduction oven to remove N₂O and O₂ from the effluent stream. After a drying stage (Nafion membrane), the sample was transferred to the mass spectrometer using an open split design, where about one third of the effluent from the chromatographic column was admitted via a capillary into the ion source. The mass to charge ratio (m/e) for the 44-, 45-, and 46-isotopologue peaks were integrated before the ¹³C/¹²C ratio was then calculated using the ¹⁷O correction following Santrock et al. [1985]. The δ¹³CH₄ values were reported on the VPDB scale using a pure CO₂ working standard with a δ¹³CH₄ value of −45.98 ± 0.02‰. System checks are performed daily using a compressed air standard (CSIRO 1636) that was filled at Cape Grim on March 1995 in collaboration with the Commonwealth Scientific and Industrial Research Organisation (CSIRO), who subsequently determined the δ¹³CH₄ value as −47.12 ± 0.03‰. Throughout the 3 years over which measurements have been made at LGGE for this study, the average δ¹³CH₄ on this tank was −47.07 ± 0.18‰ (N = 201).

2.2. PSU Analyses

[12] At PSU, we have developed a CF-IRMS technique to measure the δ¹³CH₄ in air samples that is similar to previously developed techniques [Miller et al., 2002; Rice et al., 2001]. The analytical system consists of a substantially modified PreCon device (Finnigan MAT) that interfaces with a MAT 252 mass spectrometer (Figure 1).

[13] The system was designed to measure δ¹³C of CH₄ as well as the δ¹⁵N and δ¹⁸O of N₂O from a single air sample. The details of the N₂O analyses have been previously discussed [Sowers et al., 2002]. Initially, samples are loaded onto the PreCon between two 1/4″ ultratorr adapters. The connections are flushed with He for 3 min (15 cc/min) before the glass valves on the 100-cc sample flask are opened allowing the air to be flushed (30 cc/min) into a chemical trap containing ascarite and MgClO₄ for CO₂ and H₂O removal. The air sample is then passed through a 1/32″ stainless steel (SS) loop in liquid nitrogen (LN₂) where the N₂O is trapped. The air +CH₄ stream is then routed through trap 2 at −130°C (pentane slush) where the CH₄ is trapped on a 0.32 cm × 10 cm SS column packed with 80/100 HaySep D material. The sample container is flushed for 1400 s to insure quantitative trapping of CH₄ in trap 2. Valco valves 1 and 2 are then switched so the CH₄ is transferred from trap 2 to trap 3, which contains a 1-m section of the 0.32-mm poraplot column immersed in the −130°C pentane bath. After 500 s, trap 3 is raised out of the pentane bath mobilizing the CH₄ that is then separated from other trace impurities in the remainder of the poraplot column before entering the oxidation furnace (Pt, Cu, and Ni wires at 1050°C) for quantitative conversion to CO₂. The effluent from the furnace passes through a 15-cm Nafion drying system cooled to −60°C for efficient water removal [Leckrone and Hayes, 1998]. The sample gas stream then enters the open split leading to the ion source of the mass spectrometer. The CO₂ peak is monitored with three separate faraday cups that record each isotope of CO₂ (m/e = 44,45 and 46). The area of each isotope peak was integrated before ratios of the peak areas were compared to CO₂ standard peaks that were introduced before and after the sample for reference.

[14] The ¹³C/¹²C of the CH₄ is referenced to a working standard of liquefied CO₂ that was previously calibrated to VPDB using NIST CO₂ reference standards (8563 and 8564). The δ¹³CO₂ of the working standard was −33.26 ± 0.07‰ (1σ, N = 16) based on periodic analyses of the working standard against a primary lecture bottle of gaseous CO₂. In all cases, the measured 45/44 ratios were corrected for ¹⁷O following Santrock et al. [1985].

2.3. Ice Core Methodology for δ¹³C of CH₄

[15] The analytical methodology for extracting and analyzing trapped gases in ice can be broken into two distinct parts: the liberation of the trapped gases from the ice core sample and the δ¹³CH₄ analysis. The extraction of the gases from the 1–1.5 kg ice core samples was accomplished using a “wet” extraction technique. The outside of the ice samples were initially shaved to remove the outermost ∼5 mm of ice. The shaved samples were then inserted into a 7.6 cm × 35 cm stainless cylinder that was sealed with a copper gasket prior to evacuation. After 1 hour of evacuation, the cylinder was isolated and inserted into a warm water bath to melt the ice samples and liberate the trapped air into the headspace above the meltwater. The cylinder was then transferred to a large SS Dewar where the meltwater is refrozen using liquid nitrogen. After 30 min, the headspace was flushed with UHP He (60 cc/min) through a water trap (−110°C) and a HaySep D trap at −130°C where the CH₄ was quantitatively removed from the air. After 1 hour, the CH₄ trap was isolated and removed from the system. The trap was then connected to the PreCon and processed in an identical fashion as the air samples described above.

3.1. PSU Procedure Verification

[16] Each stage of the analytical procedure was checked to quantify the degree to which the measured δ¹³CH₄ reflects the δ¹³C of CH₄ in the original sample. To establish the integrity of the preconcentration/mass spectrometry, we analyzed 50–100 cc (STP) aliquots of two compressed air standards that had been previously calibrated by Stan Tyler (UCI) as part of their atmospheric δ¹³CH₄ program [Rice et al., 2001; Tyler et al., 1999]. The assigned δ¹³CH₄ values for the two standards (CH₄ Std 1 and 2 in Table 1) are −19.6 ± 0.03‰ and −47.16 ± 0.04‰, respectively. Aliquots of these compressed air standards are routinely expanded into numerous 50–100 cc flasks at the beginning of each day to insure the complete analytical protocol provides accurate and precise δ¹³CH₄ values. If the results on the daily standards are more than 0.2‰ away from the assigned value, system checks and additional standards are run until the results fall within the 0.2‰ tolerance. Over the 3-year period since we began measuring δ¹³CH₄, we have measured 51 aliquots of CH₄ Std 1 and 116 aliquots of CH₄ Std 2 yielding average values of −19.6 ± 0.3‰ and −46.86 ± 0.2‰, respectively (Table 1). The difference between the average values and the assigned values is less than 0.2‰, suggesting that our analytical procedures provide accurate δ¹³CH₄ data with an external precision of 0.2‰.

[17] Aliquots of CH₄ Std 2 air standards were processed through the entire ice core procedure using degassed ice from Siple Dome. We performed eight simulated trapped gas transfers between October 2002 and November 2003 using 100–150 cc (STP) aliquots of CH₄ Std 2 with an assigned δ¹³CH₄ value of −47.16‰ (Table 1). The resulting average δ¹³CH₄ was −47.0 ± 0.3‰. The results are indistinguishable from the assigned value, indicating that the complete procedure did not measurably alter the isotopic composition of CH₄. The standard deviation about the mean for the simulated transfers is ∼0.1‰ higher than that based on replicate analyses of the standards expanded directly into glass flasks. We attribute the added uncertainty to variable amounts of water that get transferred to the PreCon with the fossil air sample. Small changes in the water background in the source accompanying the sample CO₂ influence the measured 45/44 ratio through protonation of CO₂ to HCO₂⁺ (m/e = 45) in the ion source [Leckrone and Hayes, 1998].

3.2. South Pole Firn Air

[18] The NOAA/CMDL [CH₄] data show a gradually decreasing trend with depth to 114 m with a much steeper rate of decrease below 114 m to the final bubble close off at 122 m where the measured CH₄ was 910 ppb (Figure 2). The 114-m horizon represents the top of the lock-in zone below which exchange with the overlying firn is substantially restricted. The δ¹³CH₄ data show a gradually decreasing trend between the surface and 110 m. Between 110 m and 122 m, the δ¹³CH₄ values exhibit an interesting oscillation that has not previously been observed in firn air studies because of their younger firn air age. The minimum values approach −51.7‰ at 117 m with increasing δ¹³CH₄ values below 117 m. This oscillation is generated by the faster diffusion coefficient of ¹²CH₄ in air compared to ¹³CH₄, superimposed on a present-day atmospheric δ¹³CH₄ value which is comparatively heavier than the pre-industrial one.

[19] There are three firn air δ¹³CH₄ data sets from the South Pole plotted in Figure 2. Two of the three δ¹³CH₄ data sets were generated at LGGE, one on each of the two expeditions (1995 and 2001). At each depth, multiple flasks were collected to provide enough firn air for all analyses. All data plotted in Figure 2 are average results from all flasks that were measured at each depth by each lab (generally duplicate flasks with duplicate or triplicate analyses on each flask in each laboratory). The δ¹³CH₄ measurements were also made at PSU on the same 2001 flasks as an intercalibration exercise. These data, along with the corresponding LGGE data for each flask, are tabulated in Table 2. All data have been corrected for gravitational fractionation using the δ¹⁵N of N₂ measured on each flask at Princeton (M. Bender, personal communication, 2002). Eleven flasks were measured at PSU before shipping 20 flasks to LGGE for analyses. Five of those flasks were remeasured at PSU after the LGGE analyses to check for alterations associated with the trans-Atlantic shipments and the LGGE analyses. The average difference between the PSU analyses performed before and after the LGGE analyses was −0.05‰, indicating no measurable change in the δ¹³CH₄ values.

[20] Comparison of the δ¹³CH₄ results from each flask measured at LGGE and PSU showed the average difference between the 19 flasks was 0.32 ± 0.28‰. This value is comparable to the external precision associated with replicate analyses on a given flask (∼0.1‰). We interpret this close agreement as indicating the analytical procedures and working standards at the two laboratories are consistent within our external precision.

[21] In order to compute the net change in the δ¹³CH₄ between the two firn air expeditions (2001–1995 = Δδ¹³CH₄), we interpolated the 1995 δ¹³CH₄ profile at depths for which samples were taken during the 2001 expedition. We then used the average LGGE data at each depth for the two samplings to calculate the difference (Δδ¹³CH₄, Figure 2). With the exception of the 53 mbs sample, the upper 100 m showed the 2001 values to be ∼0.3‰ higher than the 1995 values. The δ¹³CH₄ values below 100 mbs are effectively identical given our sample resolution. This was expected as the air in the deeper portions of the firn has not communicated with the overlying atmosphere during the 6 years that separate the two expeditions. We interpret this close agreement below 100 mbs as evidence that the sampling and analytical techniques employed during the two expeditions are indistinguishable from one another.

3.3. Siple Dome Ice

[22] We have analyzed 10 individual samples of ice from a shallow Siple Dome ice core that was drilled during the 1996/1997 field season (81°24.18′S, 148°18.14′W, elevation = 621 masl, mean annual temp = −25.4°C, accumulation ∼11 cm ice/yr [Taylor et al., 2004]). The depths ranged from 62.4 m to 70.6 m, with corresponding gas ages between 1837 and 1907 A.D. The average δ¹³CH₄ was −48.8 ± 0.37‰ (1σ). This value has been corrected for gravity with three measurements of the σ¹⁵N of N₂ (δ¹⁵N₂ = 0.22 ± 0.02‰) (1σ) using a previously described technique [Sowers et al., 1989; Sowers and Jubenville, 2000]. One sample from 66.5 mbs (gas age = 1873 A.D.) provided a δ¹³CH₄ value of −49.8‰. While we cannot attribute the anomalous result to any identified artifact associated with the extraction, we consider this value questionable given that it falls more than 1‰ below the average of the other nine samples (−48.7 ± 0.1‰). We consider the average of the nine samples (−48.7 ± 0.1‰) to be the best estimate of the average δ¹³CH₄ for the later part of the nineteenth century.

5.1. Evolution of δ¹³CH₄ Between 1995 and 2001 at the South Pole

[29] One advantage of performing multiple firn air experiments at the same locale is the fact that seasonal variations in surface air are smoothed out in the firn by diffusion. Resampling firn air thereby provides a unique means of deciphering the compositional changes without having to continually sample and measure air from a site. In this case, we sampled the interstitial firn air at South Pole in January of 1995 and again in January of 2001, providing a 6-year integrated window.

[30] To determine the net change in δ¹³CH₄, we utilized the 1995 and 2001 δ¹³CH₄ data generated at LGGE (Figure 2b). To calculate the δ¹³CH₄ difference, we interpolated the 1995 profile at depths sampled in 2001. The resulting Δδ¹³CH₄ profile is plotted in Figure 2c. Discounting the anomalous 2001 sample at 53 mbs, the average Δδ¹³CH₄ between 20 and 103 m is 0.56 ± 0.14‰ (N = 6). This value should represent the true δ¹³CH₄ change over the 6-year period as it was determined by measurements made in the same laboratory using identical equipment and procedures. However, we have documented a measurable offset between the CSIRO air standard cylinder measurements that were made during the two analytical windows. For the 1995 suite, 43 measurements of the CSIRO standard were made with a mean δ¹³CH₄ value of −47.22 ± 0.08‰. This value is 0.21‰ lower than the corresponding δ¹³CH₄ value obtained from the same tank in 2001 (−47.01 ± 0.012‰, N = 25). We attribute the different results to slightly higher water background levels associated with the 2001 samples [Leckrone and Hayes, 1998], and a drift in the δ¹³C value of the liquid CO₂ tank that served as the working reference. We have thus applied the 0.21‰ correction to all 2001 LGGE δ¹³CH₄ values (Table 2). Applying this difference to the measured Δδ¹³CH₄ values, we calculate that the net change in atmospheric δ¹³CH₄ over the 6-year period was 0.35 ± 0.2‰. This value translates to +0.06 ± 0.03‰/yr for our best estimate of the average annual δ¹³CH₄ change recorded in the firn at South Pole. This value is slightly higher than that determined by Bräunlich et al. [2001] and Francey et al. [1999] of +0.04 ± 0.01‰/yr for the last 2 decades of the twentieth century. The increasing δ¹³CH₄ trend is thought to be the combined result of an increase in the proportion of Δδ¹³CH₄ emissions with elevated ¹³C/¹²C values and/or simply the added time needed for atmospheric δ¹³CH₄ to adjust to a change in sources/sinks [Tans, 1997].

5.2. Preliminary Tests With an Atmospheric Isotope Model

[31] The twentieth century δ¹³CH₄ record from the South Pole firn air provides additional constraints on the cause of the 150% mixing ratio increase. Given the plethora of sources, sinks, and their characteristic isotope fingerprints, it is not possible to derive a single set of historical emission records that will satisfy the δ¹³CH₄ data. To address this problem, we utilized an eight-box atmospheric CH₄ model (BOSCAGE-8) that was constructed for performing long-term sensitivity tests of various emission scenarios [Marik, 1998]. Briefly, the BOSCAGE-8 model consists of eight boxes: six boxes cover the troposphere in 30° meridional bands and two for each hemisphere of the stratosphere with the boundary between the troposphere and stratosphere set at 200 hPa. Model dynamics were tuned with SF₆ following Levin and Hesshaimer [1996]. The model is based on seasonal and latitudinal variations in the emission and sink terms as well as the tabulated characteristic isotopic values adopted from the inverse modeling of Hein et al. [1997]. Starting with an initial inventory of sources and their distribution (Table 3), the model calculates the mixing and isotopic ratios inside each box taking into account the exchange times and the methane lifetime (about 8 years) estimated from the sinks. The free parameters of the model are the strengths of each source and sink and the characteristic isotope value of each source group (animals, rice, bogs, swamps, landfills, biomass burning, coal, oil/gas). In order to reproduce the industrial increase of CH₄ mixing ratio, Marik [1998] keeps the natural sources constant and sets the anthropogenic sources (animals, rice, landfills, coal, oil/gas, and part of the total biomass burning) proportional to human population, as suggested by Khalil and Rasmussen [1985]. The BOSCAGE-8 model does not have a chemical module to predict changes in [OH] that result from increased CH₄ and CO loadings. To account for the decreasing [OH] levels, we adopt a 25% linear decrease between 1885 A.D. and the present that was suggested by Marik [1998] in line with results from Khalil and Rasmussen [1985]. Model runs with these settings compared extremely well with direct measurements of [CH₄], δ¹³CH₄, and δD_CH4 during the 1990s from three sites: Alert (82°N), Izana (28°N), and Neumayer Station (70°S) [Marik, 1998].

5.3. Sensitivity Tests With the BOSCAGE-8 Model

[32] There are effectively two bacterial processes that liberate methane. In terrestrial ecosystems, acetate fermentation is the dominant pathway whereby methanogenic bacteria produce methane with characteristic δ¹³CH₄ values between −55 and −65‰ [Francey et al., 1999; Quay et al., 1999, 1991; Whiticar, 2000, 1993]. In deep sea marine sediments, CO₂ reduction is the primary pathway yielding δ¹³CH₄ values between −55 and −70‰ [Whiticar et al., 1986]. The large δ¹³CH₄ overlap between these two metabolic pathways makes it difficult to ascribe a given change in the atmospheric δ¹³CH₄ to any particular source. However, CH₄ associated with biomass burning has a distinct δ¹³CH₄ value (−27 ± 3‰ [Quay et al., 1991]) relative to the dominant bacterial sources. Thus, while we cannot rule out changes in the characteristic δ¹³CH₄ value for any of the anthropogenically mediated sources that would account for the increasing δ¹³CH₄ values we recorded over the past 150 years, we suggest that the more likely explanation for the increasing δ¹³CH₄ values is increased CH₄ emissions from biomass burning. To assess the magnitude of the expected increase in biomass burning that would account for the δ¹³CH₄ increase, we performed three sensitivity tests involving changes in the percentage of total biomass burning CH₄ that is associated with anthropogenic activities.

[33] The optimum scenario from Marik [1998] identified 30% of the total biomass burning CH₄ as being anthropogenic in origin. In our first scenario, we assumed that 70% of the 53.2 Tg/yr of biomass burning methane in 1985 was related to anthropogenic activities. Two other runs were made in which we assigned the anthropogenic contribution as 20% and 30% of total. In all runs, the natural emissions were held constant for the duration of the model run. To maintain the atmospheric loading for all three runs, the OH sink term during the nineteenth century had to be increased by 40%, 35%, and 25% relative to today's value for the simulations in which anthropogenic biomass burning accounted for 20%, 30%, and 70% of the total.

[34] The three corresponding δ¹³CH₄ trends calculated with the BOSCAGE-8 model for the southern tropospheric box since 1790 A.D. are shown in Figure 3. The output from the BOSCAGE-8 scenarios with 20 and 30% of total biomass burning attributable to anthropogenic activities envelopes both the 1995 A.D. and 2001 A.D. Monte Carlo envelopes as well as the Siple Dome δ¹³CH₄ results. One interpretation of these results holds that between 20% (11 Tg/yr) and 30% (16 Tg/yr) of total CH₄ from biomass burning in 1985 AD is attributable to anthropogenic activities. It is, however, noteworthy that this result assumes (1) no anthropogenic biomass burning emissions at 1800 A.D. whereas savanna burning and wood use for cooking and heating were already significant contributors, and (2) that all natural CH₄ emissions have remained constant and the characteristic δ¹³CH₄ values that we have assigned to the various sources and sinks (Table 3) have also remained constant (thus assuming no shift of the anthropogenic biomass burning signature from “heavy” savanna burning to “lighter” forest burning).

[35] We finally utilized the δ¹³CH₄ output from the BOSCAGE-8 model as the input for the forward firn model to compare the predicted firn δ¹³CH₄ profile with our measurements. This test effectively evaluates the internal consistency between the BOSCAGE model assumptions and the most accurate estimates of the atmospheric δ¹³CH₄ record over the last century. We utilized the BOSCAGE model output that resulted from assuming that 30% of total biomass burning CH₄ emissions were anthropogenically produced. This BOSCAGE atmospheric record was used as the surface forcing along with the CH₄ loading record from Etheridge et al. [1998] in the Battle model ending in January 2001. The resulting δ¹³CH₄ profile is plotted in Figure 2b (black line). The difference between the BOSCAGE predicted δ¹³CH₄ profile and the measured data (omitting the 53 mbs sample) is 0.08‰ with the upper portion of the model curve having slightly lower δ¹³CH₄ values and the lower portion of the model curve having slightly higher δ¹³CH₄ values relative to the measured values. Overall, we consider the close agreement between the BOSCAGE model profile and the measured profile as support for the atmospheric δ¹³CH₄ history resulting from the chosen scenario.

[36] The BOSCAGE-8 model does not include a separate term for methane decomposition by stratospheric Cl. Rather, we have modeled the stratospheric uptake via both OH and Cl with a single fractionation factor (α_strato = 1.16). This tuned parameter is very close to the empirically determined value (α_strato = 1.153 ± 0.008) from Rice et al. [2003]. As the stratospheric fractionation factor is held constant throughout the model run and stratospheric Cl levels must have been lower during the eighteenth and nineteenth centuries, the model predicted tropospheric δ¹³CH₄ values during the nineteenth century are slightly higher than model runs in which the Cl effect had been treated separately. The magnitude of this oversimplification must be less than the total effect attributable to Cl today (0.2–0.5‰) [McCarthy et al., 2003; Platt et al., 2004; Wang et al., 2002].

[37] These preliminary tests using our reconstructed δ¹³CH₄ trends thus suggest that a significant contribution of natural biomass burning to the CH₄ budget must be invoked for the last 2 centuries. Our results also suggest that the source/sink scenario developed by Marik [1998] is compatible with our δ¹³CH₄ data over the last century. Testing other scenarios was beyond the scope of our study and will motivate future research.