18.104.22.168 Effects of Climate
Effects of climate on CH4 biogeochemistry are investigated by examining records of the past and from model simulations under various climate change scenarios. Ice core records going back 650 ka (Petit et al., 1999; Spahni et al., 2005) reveal that the atmospheric concentration of CH4 is closely tied to atmospheric temperature, falling and rising in phase with temperature at the inception and termination of glacial episodes (Wuebbles and Hayhoe, 2002). Brook et al. (2000) show that, following each transition, temperature increased more rapidly than CH4 concentration. Since biogenic CH4 production and emission from major sources (wetlands, landfills, rice agriculture and biomass burning) are influenced by climate variables such as temperature and moisture, the effect of climate on emissions from these sources is significant.
Several studies indicate a high sensitivity of wetland CH4 emissions to temperature and water table. Before the 1990s, elevated surface temperature and emissions from wetlands were believed to contribute to the increase in global CH4 emissions (Walter and Heimann, 2001a,b; Christensen et al., 2003; Zhuang et al., 2004). Observations indicate substantial increases in CH4 released from northern peatlands that are experiencing permafrost melt (Christensen et al., 2004; Wickland et al., 2006). Based on the relationship between emissions and temperature at two wetland sites in Scotland, Chapman and Thurlow (1996) predicted that CH4 emissions would increase by 17, 30 and 60% for warmings of 1.5°C, 2.5°C and 4.5°C (warming above the site’s mean temperature during 1951 to 1980), respectively. A model simulation by Cao et al. (1998) yielded a 19% emission increase under a uniform 2°C warming. The combined effects of a 2°C warming and a 10% increase in precipitation yielded an increase of 21% in emissions. In most cases, the net emission depends on how an increase in temperature affects net ecosystem production (NEP), as this is the source of methanogenic substrates (Christensen et al., 2003), and on the moisture regime of wetlands, which determines if decomposition is aerobic or anaerobic. Emissions increase under a scenario where an increase in temperature is associated with increases in precipitation and NEP, but emissions decrease if elevated temperature results in either reduced precipitation or reduced NEP.
For a doubling in atmospheric CO2 concentration, the GCM of Shindell et al. (2004) simulates a 3.4°C warming. Changes in the hydrological cycle due to this CO2 doubling cause CH4 emissions from wetlands to increase by 78%. Gedney et al. (2004) also simulate an increase in CH4 emissions from northern wetlands due to an increase in wetland area and an increase in CH4 production due to higher temperatures. Zhuang et al. (2004) use a terrestrial ecosystem model based on emission data for the 1990s to study how rates of CH4 emission and consumption in high-latitude soils of the NH (north of 45°N) have changed over the past century (1900–2000) in response to observed change in the region’s climate. They estimate that average net emissions of CH4 increased by 0.08 Tg yr–1 over the 20th century. Their decadal net CH4 emission rate correlates with soil temperature and water table depth.
In rice agriculture, climate factors that will likely influence CH4 emission are those associated with plant growth. Plant growth controls net emissions by determining how much substrate will be available for either methanogenesis or methanotrophy (Matthews and Wassmann, 2003). Sass et al. (2002) show that CH4 emissions correlate strongly with plant growth (height) in a Texas rice field. Any climate change scenario that results in an increase in plant biomass in rice agriculture is likely to increase CH4 emissions (Xu et al., 2004). However, the magnitude of increased emission depends largely on water management. For example, field drainage could significantly reduce emission due to aeration of the soil (i.e., influx of air into anaerobic zones that subsequently suppresses methanogenesis, Li et al., 2002).
Past observations indicate large interannual variations in CH4 growth rates (Dlugokencky et al., 2001). The mechanisms causing these variations are poorly understood and the role of climate is not well known. Emissions from wetlands and biomass burning may have contributed to emission peaks in 1993 to 1994 and 1997 to 1998 (Langenfelds et al., 2002; Butler et al., 2004). Unusually warm and dry conditions in the NH during ENSO periods increase biomass burning. Kasischke and Bruhwiler (2002) attribute CH4 releases of 3 to 5 Tg in 1998 to boreal forest fires in Eastern Siberia resulting from unusually warm and dry conditions.
Meteorological conditions can affect global mean removal rates (Warwick et al., 2002; Dentener et al., 2003a). Dentener et al. find that over the period 1979 to 1993, the primary effect resulted from changes in OH distribution caused by variations in tropical tropospheric water vapour. Johnson et al. (2001) studied predictions of the CH4 evolution over the 21st century and found that there is also a substantial increase in CH4 destruction due to increases in the CH4 + OH rate coefficient in a warming climate. There also appear to be significant interannual variations in the active Cl sink, but a climate influence has yet to be identified (Allan et al., 2005). On the other hand, several model studies indicate that CH4 oxidation in soil is relatively insensitive to temperature increase (Ridgwell at al., 1999; Zhuang et al., 2004). A doubling of atmospheric CO2 would likely change the sink strength only marginally (in the range of –1 to +3 Tg(CH4) yr–1; Ridgwell et al., 1999). However, any change in climate that alters the amount and pattern of precipitation may significantly affect the CH4 oxidation capacity of soils. A process-based model simulation indicated that CH4 oxidation strongly depends on soil gas diffusivity, which is a function of soil bulk density and soil moisture content (Bogner et al., 2000; Del Grosso et al., 2000).
Climate also affects the stability of CH4 hydrates beneath the ocean, where large amounts of CH4 are stored (~4 ×106 Tg; Buffett and Archer, 2004). The δ13C values of ancient seafloor carbonates reveal several hydrate dissociation events that appear to have occurred in connection with rapid warming episodes in the Earth’s history (Dickens et al., 1997; Dickens, 2001). Model results indicate that these hydrate decomposition events occurred too fast to be controlled by the propagation of the temperature change into the sediments (Katz et al., 1999; Paull et al., 2003). Additional studies infer other indirect and inherently more rapid mechanisms such as enhanced migration of free gas, or reordering of gas hydrates due to slump slides (Hesselbo et al., 2000; Jahren et al., 2001; Kirschvink et al., 2003; Ryskin, 2003). Recent modelling suggests that today’s seafloor CH4 inventory would be diminished by 85% with a warming of bottom water temperatures by 3°C (Buffett and Archer, 2004). Based on this inventory, the time-dependent feedback of hydrate destabilisation to global warming has been addressed using different assumptions for the time constant of destabilisation: an anthropogenic release of 2,000 GtC to the atmosphere could cause an additional release of CH4 from gas hydrates of a similar magnitude (∼2,000 Gt(CH4)) over a period of 1 to 100 kyr (Archer and Buffett, 2005). Thus, gas hydrate decomposition represents an important positive CH4 feedback to be considered in global warming scenarios on longer time scales.
In summary, advances have been made since the TAR in constraining estimates of CH4 source strengths and in understanding emission variations. These improvements are attributed to increasing availability of worldwide observations and improved modelling techniques. Emissions from anthropogenic sources remain the major contributor to atmospheric CH4 budgets. Global emissions are likely not to have increased since the time of the TAR, as nearly zero growth rates in atmospheric CH4 concentrations have been observed with no significant change in the sink strengths.