Historical and Contemporary Global Methane Cycling

Abstract

Methane (CH₄) has been studied as an atmospheric constituent for more than 200 years. However, the first modern measurements of atmospheric CH₄ concentration were made by using the infrared absorption and estimated an atmospheric concentration of 2.0 ppmv in 1948. The development of gas chromatography (GC) and the flame ionization detector (FID) in the 1950s led to observations of vertical CH₄ distributions in the troposphere and stratosphere , and to establishment of time-series sampling programs starting from 1979. Results from these sampling programs led to suggestions that similar to atmospheric CO₂, the concentration of atmospheric CH₄ has been increasing. The data indicated that the atmospheric CH₄ concentration has increased almost exponentially from 722 ± 4 ppb at the onset of Industrial Revolution in 1750 to 1650 ppb by mid 1980s when global CH₄ emissions monitoring started. The current (2015) global annual surface mean abundance is 1845 ± 2 ppb, and a relative abundance of 256% compared to 1750. The mean annual absolute increase during the last 10 years is estimated at 6 ppb CH₄ yr⁻¹. The role of CH₄ as one of the greenhouse gas (GHG) causing global warming stimulated further research on sources and sinks of CH₄, which is emitted from a variety of sources, both of natural and anthropogenic origin. Natural sources account for about 40% of the total, while anthropogenic emissions contributing 60% of the global emissions. The anthropogenic sources fall under the main categories of agriculture, energy, waste, and industry. CH₄ is also of interest to microbiologists, but findings from microbiology have entered the larger context of the global CH₄ budget only recently. CH₄ is the most abundant hydrocarbon in the atmosphere, and its increase by a factor of 2.5 since the Industrial Era has raised concerns due to the potential effects on atmospheric chemistry and climate. It plays important roles in atmospheric chemistry and the radiative balance of the Earth. Stratospheric oxidation of CH₄ provides a means of introducing water vapor above the tropopause. The CH₄ also reacts with atomic chlorine in the stratosphere, forming HCl, a reservoir species for chlorine. Some 90% of the CH₄ entering the atmosphere is oxidized through reactions initiated by the OH radical, mostly in the troposphere. These reactions are important in controlling the oxidation state of the atmosphere. The CH₄ absorbs infrared (IR) radiation in the troposphere and is an important GHG. On per mole basis, CH₄ is more effective a GHG than CO₂. Current atmospheric burden is estimated at 4954 ± 10 Tg CH₄, with the annual increase estimated at 14 ± 3 Tg CH₄ yr⁻¹ from both natural (36.3%) and anthropogenic (63.7%) sources.

7.1 Introduction

Oxidation of CH₄ has been a subject of scientific inquiry since 1776, when Alexandro Volta (1745–1827), an Italian physicist described the first experiments of flammable air released by shallow sediments in Lake Maggiore in a letter to Father Carlo Campi (Wolfe 2004). Subsequently, CH₄ was exploited as a source of heat and light (King 1992). Despite its commercial significance, the biological and ecological aspects of CH₄ oxidation did not gain attention until the pioneering work of Söhngen, who isolated CH₄-oxidizing bacteria in 1906. Little additional progress was made until 1960s when systematic efforts of several groups provided the methodological tools and detail taxonomy, physiology and biochemistry of methanogenic metabolism (King 1992). Among other motivations, this work was stimulated by the potential role of methanotrophic bacteria in food spoilage and by possible use of methanotrophs in bioremediation of certain halogenated organic pollutants or agents of commercial biotransformation (Higgins et al. 1980). The critical role of CH₄ in atmospheric chemistry and Earth’s energy budget has stimulated environmental and ecological research since early 1990s (Crutzen 1995).

CH₄ is the most abundant hydrocarbon in the Earth’s atmosphere, which plays important roles in atmospheric chemistry and radiative balance of the Earth. Next to H₂O and CO₂, CH₄ is the most abundant long-lived greenhouse gas (GHG) which absorbs infrared radiation in the troposphere (Ramanathan et al. 1985). Both on molecule and mass basis, CH₄ is much more effective GHG than CO₂. These gases absorb long-wave radiation emitted by Earth’s surface in the 4–100 µm atmospheric window and hence affect atmospheric temperature directly by shielding heat loss and causing global warming (Hansen et al. 1988; Ramanathan 1988). Emission of CH₄ has caused radiative forcing of 0.97 W m⁻² since 1750, about twice as much as the direct CH₄ concentration based estimate of 0.48 ± 0.05 W m⁻² (Hofmann et al. 2006; Stocker et al. 2013). The difference in estimate is caused by the indirect impacts on CO₂, stratospheric water vapor, ozone (O₃), sulfate aerosol, and the change in lifetimes of hydrofluorocarbons (HFCs) and hydrochlorofluorocarbons (HCFCs) (Boucher et al. 2009). The anthropogenic CH₄ is contributing 17–20% to the total radiative forcing induced by the long-lived anthropogenic GHGs from 1750 to 2013 (Myhre et al. 2013; Kirschke et al. 2013). A recent estimate of atmospheric lifetime of CH₄ ranges from 9.8 ± 1.6 to 11.2 ± 1.3 years (Dlugokencky et al. 2009; Myhre et al. 2013), and its emission reduction would have immediate benefit to global climate. The global warming potential (GWP) of CH₄ is 28 (i.e., it is 28 times more efficient per mass as GHG) compared to CO₂ over a 100-year period (Myhre et al. 2013).

It is most the abundant chemically reactive trace gas in the troposphere important to chemical state of atmosphere due to its chemical destruction by reaction with the hydroxyl radical (OH) that influences the abundance of ozone (O₃) in the troposphere and stratosphere (Johnston 1984) and affect tropospheric air quality. It is also a major source of stratospheric water (Pollock et al. 1980). In the troposphere, CH₄ is oxidized to CO and ultimately to CO₂ and H₂O in oxidation sequence initiated by hydroxyl (OH) radical. These reactions constitute the main sink for atmospheric CH₄ (Ehhalt and Heidt 1973). By reacting with OH radicals, CH₄ reduces the oxidizing capacity of the atmosphere and generates O₃ in the troposphere. Stratospheric oxidation of CH₄ provides the means of introducing water vapor above the tropopause (Wofsy 1976), and therefore, long term increase in atmospheric CH₄ leads to an increased anthropogenic forcing from water vapor in the stratosphere. Stratospheric water vapor increased by 1.0 ± 0.2 ppm during 1980–2010, and CH₄ oxidation explains approximately 25% of this increase (Hurst et al. 2011). About 90% of the CH₄ entering the atmosphere is oxidized through reactions initiated by OH radical.

Formation of tropospheric OH is initiated by photo-dissociation of O₃, followed by reaction with water vapor which creates sensitivity to humidity, cloud cover, and solar radiation (Crutzen 1973). The other source of OH is through secondary reactions, especially oxidation of volatile organic compounds, notably isoprene in low-NO_x air (Lelieveld et al. 2008). Some of these reactions and chemistry underpinning them are still poorly understood (Taraborrelli et al. 2012). However, these reactions are important in controlling the oxidation state of the atmosphere. The chemical coupling between OH and CH₄ leads to significant amplification of CH₄ emission impacts such as increased CH₄ atmospheric lifetime and the atmospheric CH₄ burden. The CH₄ is also the primary sink for chlorine radicals (Cicerone and Oremland 1988).

In addition to changes in atmospheric chemistry and radiative balance of the Earth’s atmosphere, CH₄ plays a role in the global C cycle (Fig. 7.1), and is also an important component of natural biogeochemical cycles. CH₄ is also an important source of environmentally cleaner energy for modern society. Numerous direct roles of CH₄ to further impact climate, atmospheric chemistry, and biogeochemistry of land and ocean and also the energy needs have made the study of CH₄ biogeochemistry an urgency, far beyond the scientific curiosity. Therefore, understanding changes in its atmospheric concentration is an indication of understanding the underlying biogeochemistry and its reaction to external changes. A portion of organic matter (OM) that enter diverse O₂-free (e.g., anaerobic) environments such as wetlands, subsurface environments including hydrocarbon deposits and deep granitic aquifers and rumen of animals is recycled by methanogenic microorganisms to CH₄, which can be emitted to atmosphere or diffuses to aerobic zone and oxidized by methanotrophic organisms to CO₂.

Fig. 7.1

Methane sub-cycle of the carbon cycle showing estimated methane reservoirs (Tg or Pg CH₄) and fluxes (Tg CH₄ yr⁻¹) of the pre-industrial period. Data source Sundquist and Ackerman (2014), Paul and Dillon (2001)

The first quantitative measurements of atmospheric CH₄, both involving combustion and gravimetric determination of trapped oxidation products were reported in France by Boussingault (1834, 1864). Gautier (1901) estimated CH₄ concentrations of 10 and 0.28 ppmv on seashore and 95 ppmv in Paris. The first modern measurements of atmospheric CH₄ were the infrared absorption measurements which estimated an atmospheric concentration of 2.0 ppmv (Migeotte 1948). Development of gas chromatography (GC) and the flame ionization detector (FID) in 1950s led to observations of vertical CH₄ distributions in the troposphere and stratosphere, which led to an establishment of time series sampling programs. The first observation of a possible atmospheric CH₄ increase were reported by (Graedel and McRae 1980) when they analyzed 10-years continuous record of total hydrocarbon data collected at the New Jersey Department of Environmental Protection. However, an increasing trend in daily and annual minima was advanced somewhat tentatively due to concerns about the calibration technique, urban locations, and the fraction of total hydrocarbon attributable to CH₄. Rasmussen and Khalil (1981) used 22 months of an automated GC/FID measurements of atmospheric CH₄ at Cape Meares, Oregon to show a concentration increase of 2% yr⁻¹ and similar trend for northern hemisphere data from 1965 to 1980. Cape Meares station has operated since 1979, and the techniques used by Rasmussen and Khalil (1981) were used to establish the National Oceanic and Atmospheric Administration (NOAA) network of global sampling for analyzing atmospheric CH₄ concentrations.

The current CH₄ emissions are of both natural and anthropogenic sources. Natural sources account for about 40% of the global emissions, whereas 60% comes from anthropogenic sources (Ciais et al. 2013). Anthropogenic sources are the result of agriculture, energy, waste, and industrial production, which contribute 50.63, 28.65, 20.61 and 0.10% of the anthropogenic sources, respectively (Karakurt et al. 2012). The increase in number of ruminants and expansion of rice paddy agriculture are the dominant agricultural sources, while coal mining and leakage from natural gas wells and supply system are attributed to fossil fuels. The emissions from landfills and waste are the dominant urban sources. Total anthropogenic emissions contribute between 50 and 65% of the total CH₄ sources. Enteric fermentation (agriculture), landfills and energy are the dominant sources in the United States (EPA 2015).

CH₄ is produced by bacteria under anaerobic conditions in wet environments such as wetlands, swamps, bogs, ferns, tundra, rice paddies, and landfills. CH₄ emitted from inundated anoxic environments can substantially be reduced by bacterial oxidation of CH₄ in oxic layers above. A small sink can be attributed to bacterial oxidation on relatively dry soils (Born et al. 1990) which relies on consumption of atmospheric CH₄. It is also produced in the stomachs of ruminants (i.e., cattle and other cud chewing mammals) and termites. Most of this biogenic CH₄ is generated by two major bacterial pathways: (i) fermentation of acetate, and (ii) reduction of CO₂ with H₂ (Whiticar et al. 1986; Cicerone and Oremland 1988).

Some of the emitted CH₄ is exported to the stratosphere. The same oxidation reactions which occurs in the troposphere also occurs in the lower stratosphere, and is initiated by OH. At higher altitudes, reactions with exited O₂ atoms (O¹D) and with chlorine atoms lead to mixing ratio profiles that decrease with increase in altitude. The atmospheric increase in CH₄ during the industrial Era is mostly attributed to anthropogenic activities. Higher concentrations of CH₄ have been observed above and downwind of densely populated and intensive agriculture areas where anthropogenic emissions occur (Frankenberg et al. 2011).

Although CO₂ emission and its mitigation dominate much of climate change research and policies, there is an increasing recognition that CH₄ emissions may often provide cost effective and more efficient means to mitigate anthropogenic climate change. In addition, current projection of GHG concentrations and resultant climate forcing in the 21st century require an improved understanding of how natural CH₄ sources will respond to changes in climate. Therefore, this chapter provides a synthesis of scientific understanding of major global sources of CH₄ and projections of how these emissions may change in response to the projected climate change.

The objectives of this chapter are to examine the past trends in the CH₄ concentrations, determine the sources and sinks affecting its atmospheric growth rate, and elucidate factors that could affect its future atmospheric growth rate. The chapter also examines the current underlying chemistry and its role as sub-cycle of the global carbon cycle.

7.2 Atmospheric Methane

Systematic measurements of CH₄ in the atmosphere began rather recently compared to those of atmospheric CO₂ (i.e., 1978, Dlugokencky et al. 1994). Therefore, to establish the extent of recent increases in its atmospheric concentration, the natural variability, and the processes underlying natural changes, it is necessary to identify other ways to establish its evolution with time. Polar ice acts as a natural gas archive, which preserves samples of ancient atmospheric air from which the concentration of stable trace gases can be measured directly (Raynaud et al. 1993; Stauffer et al. 2002). On the polar plateaus of Antarctica and Greenland, snow generally does not melt, but is compressed by weight of overlying new snowfall, and eventually turning to solid ice with enclosed air bubbles that becomes isolated from the overlying atmosphere at a depth ranging from 60 to 100 m depending on the snow accumulation rate and temperature (Wolff 2011). Because the snow remains permeable to air for several years until the close-off, the air trapped in ice cores is older than the air above the snow surface but much younger than the ice that enclose it due to air diffusion to snow pack prior to close-off which causes smoothing (Joos and Spahni 2008; Köhler et al. 2011). As a result, a given depth contains bubbles with an age distribution. The age distribution within a given ice core layer (e.g., Holocene at Dome C in Antarctica) will be typically 400 years (Schwander and Stauffer 1984), and this is what controls the resolution that can be achieved from ice core CH₄ concentration analysis (Wolff 2011).

7.2.1 Variability of Methane Concentration During Glacial-Interglacial Cycles

The history of atmospheric CH₄ mixing ratios has been reconstructed from measurements of air bubbles trapped in ice cores, and ice core records of trace gases are currently available over the period extending as far back as 800,000 years, and covering eight glacial-interglacial cycles (Augustin et al. 2004; Loulergue et al. 2008b). The concentration of CH₄ determined from ice cores during this period stayed within natural limits, with glacial concentration ranging from 320 to 400 ppb which then increased to maximum values ranging from 600 to 800 ppb during interglacial periods (Fig. 7.2a; Loulergue et al. 2008b; Wolff 2011). During glacial and interglacial times, large natural variations are also observed over different timescales. Figure 7.2b shows the variability of the ratio of deuterium to H (i.e., δD) of CH₄ from ice cores during the glacial/interglacial period. Similar to CH₄ cycles, the climate record of the past 800,000 years is dominated by the alteration between relatively short warmer interglacial periods and much longer cold glacial periods for both Greenland and Antarctica as revealed by temperature reconstructed from deuterium profile (Fig. 7.2c; Jouzel et al. 2007). Both CH₄ and reconstructed temperature cycles recur at about 100,000 years intervals (Loulergue et al. 2008b). The CH₄ concentration also shows sharp millennial changes of the order of 200 ppbv (Flückiger et al. 2004; Loulergue et al. 2008a).

Fig. 7.2

The CH₄ (a), deuterium (b), and estimated temperature (c) fluctuations during glacial/interglacial cycles. Source Flückiger et al. (2004), Loulergue et al. (2008a, b), Parrenin et al. (2013)

The most conspicuous feature of the Vostok data is the close similarity among trends in CH₄ and local temperature reconstructed from the ratio of deuterium to H₂ in the ice (Fig. 7.2a, c). The most recent four cycles show long glacial period with short interglacial periods, and a temperature difference range for Antarctica of around 10 °C between warm and cold temperatures (Fig. 7.2c; Petit et al. 1999; Augustin et al. 2004). The patterns of Antarctic climate seen in the ice cores is also seen in other climate records around the world, (e.g., Lisiecki and Raymo 2005). The data revealed that maximum and minimum temperatures and CH₄ concentrations were nearly similar during each of eight glacial cycles, including the Holocene Epoch prior to human influence, which demonstrates the fact that similar to CO₂, CH₄ has contributed significantly to glacial/interglacial climate change. The abrupt climate change on timescales of 10–100 years resulting into CH₄ variability of 50–200 ppbv, and sharp climate jumps, which in Greenland can have an amplitude as high as 15 °C (Huber et al. 2006), termed as Dansgaard/Oeschger climatic events coinciding with summer insolation in the Northern Hemisphere, are also widespread throughout glacial/interglacial cycles (Severinghaus et al. 1998; Flückiger et al. 2004). The drivers of these changes are not well understood, but some indications suggest that they involve changes in ocean heat transport (Stocker and Johnsen 2003).

The isotopic ratio measurements from the ice cores (¹³CH₄) provide constraints on mechanisms responsible for temporal changes in CH₄ concentration. The δ¹³CH₄ is also useful for separating different sources of CH₄, since different CH₄ sources have characteristic δ¹³C signatures (Quay et al. 1999). The δ¹³CH₄ and deuterium (δD) ratio of CH₄ suggest that most of atmospheric CH₄ concentration increase during the interglacial period was the result of increased emissions from boreal and tropical wetlands and/or changing CH₄ sinks, most notably increased atmospheric CH₄ residence time due to reduced oxidative capacity of the atmosphere (Fischer et al. 2008). However, possible influence of other sources (i.e., marine CH₄ hydrates, biomass burning, and vegetation) must also be considered. Modeling studies suggest that temperature influence on tropical wetlands and global vegetation were the dominant controls of changes in atmospheric CH₄ concentration during glacial/interglacial period (Konijnendijk et al. 2011).

7.2.2 Methane Concentrations During Holocene

During the Holocene period, the interglacial which began ~11,700 years ago, the atmospheric CH₄ concentration decreased from near 700 ppb during the early periods of Holocene ~10,000 years, reaching the lowest concentrations of 570 ppb at about 5500 years, followed by an increase to 700 ppb in the pre-industrial time (Fig. 7.3a; Blunier et al. 1995; Flückiger et al. 2002). These changes are gradual, and not similar to Dansgaard-Oeschger event in Greenland climate and CH₄ or glacial/interglacial period. The mid-Holocene low atmospheric CH₄ concentrations might be mainly due to a reduction in tropical CH₄ sources (i.e., drying of the tropics, Flückiger et al. 2002), while the subsequent increase has been attributed to an expansion of boreal wetlands, although this has yet to be substantiated. It has also been suggested that the increase in atmospheric CH₄ concentration between 5000 and preindustrial period were the result of early agricultural activities in Eurasia, particularly forest clearance and rice cultivation (Ruddiman 2003). However, this early anthropogenic emissions hypothesis has been criticized and undergone significant modifications (Claussen et al. 2005; Broecker and Stocker 2006; Masson-Delmotte et al. 2006; Ruddiman 2007). It generally accepted that the reasons for late Holocene trends in CH₄ concentration is still not available (Wolff and Spahni 2007) due to findings that δ¹³CH₄ declined significantly from 1000 to 1800 AD (Fig. 7.3b) in absence of major changes in CH₄ concentrations (Ferretti et al. 2005).

Fig. 7.3

The methane concentration and δ¹³CH₄ record covering the Holocene period. Source Sowers (2010), Blunier et al. (1995), Brook et al. (1996, 2000), Buiron et al. (2011). Data available at https://www.ncdc.noaa.gov

Several mechanisms have been proposed to explain the late Holocene increase in atmospheric CH₄, including: (i) expansion of boreal wetland source areas (Blunier et al. 1995; Velichko et al. 1998), (ii) the CH₄ emission responses to changing climate (Indermuhle et al. 1999), and (iii) agricultural development, particularly rice paddy cultivation and widespread domestication of ruminants (Chappellaz et al. 1997; Ruddiman 2007). However, model simulations have not been able to simulate the anthropogenic or natural influence of CH₄ increase in the late Holocene. Part of the late Holocene CH₄ concentration increase could be anthropogenic (Kaplan et al. 2006) and part could be natural. However, natural wetland CH₄ models driven by simulated climate changes are able (Singarayer et al. 2011) or unable (Konijnendijk et al. 2011) to simulate the increase in CH₄ concentration observed in ice cores , reflecting the climate sensitivity of the modern CH₄ emissions from wetlands simulated by the models (Melton et al. 2013). The variations observed in CH₄ and δ¹³CH₄, are somewhat dissimilar (Fig. 7.3a, b) compared to those observed in CO₂ and δ¹³CO₂ during Holocene. Moreover, significant changes in atmospheric CH₄ can be caused by changes in specific sources and sinks that do not necessarily require the larger translocations of C needed to change atmospheric CH₄ and δ¹³CH₄, since CH₄ concentrations tends to be more susceptible to rapid variations resulting from changes in sources than CO₂ (Sundquist and Ackerman 2014). As a result, abrupt changes and spikes are generally more common in ice core records of CH₄ concentration than that of CO₂, which can be important links between climate and C cycle changes. Therefore, it is likely that a variety of processes reflecting the regional and global complexity of both C cycling and the climate contributed to the trends recorded in ice cores of the Holocene period CH₄. The atmospheric CH₄ concentration are more susceptible to rapid variations resulting from changes in sources than CO₂ concentrations, as the abrupt changes and spikes in CH₄ concentration record indicates.

The dominant source of natural CH₄ are wetlands which mainly occur in the tropics and northern hemisphere. The CH₄ produced in northern hemisphere is oxidized during its transit southward in the atmosphere, resulting in north to south decreasing atmospheric CH₄ concentration. This gradient appears to have varied during the recent millennium, with the CH₄ concentrations over Greenland exceeding those from Antarctica by 24 to 58 ± 10 ppb (Etheridge et al. 1998). The CH₄ measurements in ice from Law Dome site in Antarctica reveals a small variation that parallel the CO₂ trends over the last millennium. The CH₄ decrease by ~40 ppb in 16th and 17th century (Fig. 7.4; Etheridge et al. 1998; MacFarling Meure et al. 2006; Mitchell et al. 2011). Lower temperatures during this period may have reduced CH₄ emissions by wetlands. Alternative explanations that have been suggested include: (i) changes in biomass burning (Mischler et al. 2009), and (ii) changes in CH₄ emissions as a result of wars and plague (Mitchell et al. 2011). However, the δ¹³CH₄ measurements suggest pronounced variability in both anthropogenic and natural atmospheric CH₄ sources between 1000 and 1800 year-period (Ferretti et al. 2005; Sapart et al. 2012). Analysis of δ¹³CH₄ from Greenland ice cores and application of two-box model attributed the centennial variations in δ¹³CH₄ between 100 AD and 1600 AD to changes in pyrogenic and biogenic sources driven by both natural climate variability and changes in human population and land use (Ferretti et al. 2005; Sapart et al. 2012). Therefore, on the basis of δ¹³C isotope, the pre-industrial CH₄ sources can be divided into three categories: biogenic CH₄ (e.g., wetlands, rice paddies, and ruminants, mean δ¹³CH₄ = −60 ± 5‰), geological CH₄ (e.g., mud volcanoes, and micro-seepages, mean δ¹³CH₄ = −38 ± 7‰), and pyrogenic CH₄ (e.g., fires, biofuel, and coal burning, mean δ¹³CH₄ = −22 ± 3‰) (Quay et al. 1999; Ferretti et al. 200; Mischler et al. 2009; Houweling et al. 2008). Significant δ¹³CH₄ depletion during the last millennium (1000–1800) has been observed, which was attributed to changes in biomass burning emissions, probably influenced by both human activities and natural climate change (Ferretti et al. 2005). After 1800, an abrupt isotopic enrichment is observed, presumably resulting from increased fossil fuel emissions associated with the onset of industrialization (Ferretti et al. 2005; Mischler et al. 2009; Houweling et al. 2008).

Fig. 7.4

Law Dome ice core CH₄ concentrations over the last 2000 years superimposed by modern instrumental atmospheric CH₄ measurements (Etheridge et al. 1998; MacFarling Meure et al. 2006). Data available at https://www.ncdc.noaa.gov

7.2.3 Methane Concentration During the Industrial Era

The atmospheric CH₄ concentrations of the Industrial Era (i.e., after 1750) have been determined through reconstruction using air trapped in polar ice cores and firn air, the measurement record which has extended to recent decades (Etheridge et al. 1998; Ferretti et al. 2005; MacFarling Meure et al. 2006; Sapart et al. 2013). Measurements of CH₄ concentration from ice cores of Antarctica and Greenland suggest that after a very slow atmospheric CH₄ concentration increase averaged at 0.04 ppbv yr⁻¹ during the pre-industrial times from 1 to 1750, the atmospheric CH₄ concentration increased rapidly at an average rate of about 4.0 ppb yr⁻¹ during industrial times from 1750 to 1990 (Rasmussen and Khalil 1984; MacFarling Meure et al. 2006; Etheridge et al. 1998). Globally averaged atmospheric CH₄ concentration in 1750 was 722 ± 4 ppbv (Dlugokencky et al. 2005). After 1750, the atmospheric CH₄ levels rose gradually and reached 830 ppb by about 1850, then it increased exponentially with time reaching 1650 ppb by mid-1980s (Fig. 7.4; Etheridge et al. 1998). The global atmospheric abundance in 2015 is 1845 ± 2 ppb (WMO 2016). The extrapolated CH₄ trends from ice cores to the time periods with direct instrumental measurements are consistent with recent atmospheric measurements (Fig. 7.4; Stauffer et al. 1985; Etheridge et al. 1992), and provides a powerful support for the validity of ice core records extending back in time.

Direct measurements of CH₄ in the atmosphere at high precision began in 1978 (Blake et al. 1982; Rasmussen and Khalil 1984; Etheridge et al. 1992) and reached global coverage after 1983. Currently, the atmospheric CH₄ concentration is measured by a network of more than 100 discrete surface air samples collected regularly and with continuous measurements made at the surface (Etheridge et al. 1992; Blake et al. 1982; Francey et al. 1999; Cunnold et al. 2002; Dlugokencky et al. 2009, 2011), or in the troposphere profiles by aircraft (Brenninkmeijer et al. 2007), satellite (Wecht et al. 2012; Schuck et al. 2012), and remotely sensed measurements of atmospheric CH₄ columns retrieved from the surface or from space (Crevoisier et al. 2009; Frankenberg et al. 2011; Morino et al. 2011). Surface based observations from four networks—National Oceanic and Atmospheric Organization (NOAA), Advanced Global Atmospheric Gases Experiment (AGAGE), Commonwealth Scientific and Industrial Research Organization (CSIRO), University of California Irvin (UCI) have shown consistent changes in global growth rate of annual CH₄ concentrations since 1980s. Between the mid-1980s and 2000s the atmospheric growth rate of CH₄ declined from an average of 12 ± 6 ppb yr⁻¹ in 1980s to 6 ± 8 ppb yr⁻¹ in 1990s, and a general stabilization at nearly zero and concentration of 1773 ± 3 ppb from 1999 to 2006, a quasi-stable concentration, where the sum of emissions are in balance with the sum of sinks (Fig. 7.5; Dlugokencky et al. 2003; Khalil et al. 2007; Patra et al. 2011; Alexe et al. 2015). The atmospheric CH₄ increase picked up again from 2007 (Rigby et al. 2008). Since 2007, the atmospheric CH₄ concentrations have been increasing again at nearly steady state of ~6 ppb yr⁻¹ (Fig. 7.5, Rigby et al. 2008; Dlugokencky et al. 2009), and reached 1845 ± 2 ppb in 2015 (WMO 2016). The atmospheric CH₄ concentrations have increased by 2.53 times compared to pre-industrial concentrations of 722 ± 4 ppb (after conversion to NOAA 2004 standard scale (Dlugokencky et al. 2005). This increase reflects recent imbalance between CH₄ sources and sinks that is not fully understood (Dlugokencky et al. 2011), but is probably related to increased emissions from wetlands as a feedbacks from global climate change (Hoehler and Alperin 2014; Yvon-Durocher et al. 2014). There is very high level of confidence that anthropogenic activities caused the increase in atmospheric CH₄ concentration during the industrial era (Ciais et al. 2013). There are indications of increased anthropogenic and wetland emissions (Dlugokencky et al. 2009; Bousquet et al. 2011), but relative changes in anthropogenic and wetland emissions remains uncertain (Kirschke et al. 2013). Overall, the Arctic CH₄ concentration increased more than the global growth rate in 2007, but since then the Arctic growth rate has tracked the global growth rate, but in the southern tropics, the growth rate has remained above the global trends since 2007 (Nisbet et al. 2014).

Fig. 7.5

The atmospheric CH₄ mole fraction for the period 1983 to 2014 averaged from NOAA data (available online at: ftp://aftp.cmdl.noaa.gov/data/trace_gases/)

The reasons for the atmospheric CH₄ growth rate decline after the mid-1980s and the eventual stabilization from 1999 to 2006 are still debated but several scenarios quoted include: (i) reduction of anthropogenic emitting activities—including fossil-related emissions (coal mining, oil, and gas industry) estimated from AGAGE and NOAA data (Chen and Prinn 2006) or from ethane used as a proxy to fossil fuel-related CH₄ emissions (Aydin et al. 2011) and animal husbandry, especially in the countries of the former Soviet Union (Savolainen et al. 2009; Simpson et al. 2012), (ii) a compensation between increasing anthropogenic emissions and decreasing wetland emissions (Chen and Prinn 2006; Bousquet et al. 2006), (iii) significant to small changes in atmospheric hydroxyl ions (OH) concentration based on ¹³CH₄ datasets (Rigby et al. 2008; Montzka et al. 2011), (iv) reduced emissions from rice paddies attributed to changes in agricultural practices for paddy rice production such as water regime management (Kai et al. 2011), or (v) stable microbial and fossil fuel emissions from 1990 to 2005 (Levin et al. 2012). A synthesis of data suggests that a stable to increasing microbial emissions and a stable to decreasing fossil fuels emissions is more likely than other scenarios to explain the observed stability but it did not attribute those changes regionally (Kirschke et al. 2013).

The global network of instrumental measurements of CH₄ have improved the knowledge of CH₄ concentration in the atmosphere. The measured concentrations in the atmosphere are used to improve the CH₄ sources estimate both at country and regional level. Measurements from the global network of air sampling are also used in simulations of global atmospheric CH₄ chemistry and transport. Globally averaged monthly CH₄ mol fraction for surface flask measurements from NOAA/ESRL network indicates that atmospheric CH₄ concentration is the highest at the end of winter and decreases to the minimum during summer (Fig. 7.5) due to CH₄ oxidation in summer. The atmospheric CH₄ concentration increase since 1750 implies an anthropogenic emission of 340 ± 50 Tg CH₄ yr⁻¹. The concentration is higher than ever observed in ice core record over the past 800,000 years (Loulergue et al. 2008b).

The causes of renewed growth in atmospheric CH₄ burden since 2007 are still poorly understood and subject of intensive scientific discussion (Dlugokencky et al. 2009, 2011; Rigby et al. 2008). Recent studies have attributed the increased emissions of CH₄ in the Arctic as a result of higher temperatures in 2007 and greater rainfall in the tropics in 2008 (Dlugokencky et al. 2009). The temperature response is a snapshot of potentially positive feedback, with higher temperatures projected at high latitudes for the 21st century, increasing CH₄ emissions from wetlands, permafrost, and possibly CH₄ hydrates. Derived global CH₄ emissions for 2007–2010 estimated by top-down inverse modeling using climatological OH fields are 16–20 Tg CH₄ yr⁻¹ higher compared to 2003–2005 estimates, with most of inferred emission increase occurring in tropics and mid-latitudes of northern hemisphere (Bergamaschi et al. 2013).

7.3 Global Methane Budget

The global CH₄ budgets are overviews of sources and sinks of CH₄ derived from a priori emission estimates and long-term measurements of CH₄ concentrations at some background stations. The global CH₄ sources and sinks estimates are presented in Table 7.1. The global network of CH₄ measurement instruments have improved the knowledge on CH₄ concentrations in the atmosphere and the measured CH₄ concentrations are used to improve the understanding of source estimates at country, regional, and global level. The measurements from global air sampling network are also used in simulations of the global atmospheric chemistry and transport. Satellite information from monitoring instruments can also be used to validate model simulations and develop global maps of emission estimates. Inverse modeling which uses top-down approach to model the emission sources is also used to improve the priori emission estimates. For example, emission estimates from tropical wetlands and remote areas were improved through inverse modeling using satellite data (Bergamaschi et al. 2005, 2007, 2009, 2013; Alexe et al. 2015). Atmospheric CH₄ contribution to anthropogenic climate forcing is about half that from CO₂ when direct and indirect components of its forcing are summed (Hansen and Sato 2001). Therefore, better understanding of the global CH₄ budget and how it is changing with time is important for targeting specific sources for reduced CH₄ emission and mitigating future environmental impacts.

Table 7.1 Estimated Global of CH₄ sources and sinks for three decades of instrumental measurement (1980–2009)

The atmospheric CH₄ growth rate results from the imbalance between CH₄ emissions and CH₄ sinks. Increase in atmospheric CH₄ during the industrial era is driven mostly by anthropogenic CH₄ emissions from fossil fuels, industry, agricultural activities, biomass burning and waste. However, there are indications that meteorological feedbacks from changing climate may have also played a role by changing natural CH₄ emissions (Dlugokencky et al. 2011). Both paleorecords and current studies suggest that some sources of emissions such as wetlands and CH₄ hydrates, may be subject to strong positive feedbacks on warming (Nisbet and Chappellaz 2009). Therefore, targeting CH₄ emission is important strategy in efforts to mitigate climate change, and understanding the CH₄ budgets plays significant role in developing mitigation strategies. Global CH₄ budgets are generally quantified overview of sources and sinks for CH₄. They are a priori emission estimates and long-term measurements of CH₄ at background stations.

7.3.1 Methane Sources

CH₄ is emitted into the atmosphere by both natural and anthropogenic sources. Natural sources include wetlands, oceans, hydrates, geological sources, wild ruminant animals, termites, and wild fires. Important anthropogenic sources include coal mining, gas and oil industry, landfill and waste, domesticated ruminants, rice agriculture, and biomass burning. Natural sources are currently outweighed by anthropogenic sources. For example, decadal natural and anthropogenic CH₄ emissions estimated by bottom-up approach for 2000s are 347 ± 109 and 331 ± 27 Pg CH₄ yr⁻¹, respectively (Table 1, Kirschke et al. 2013; Ciais et al. 2013). Many significant sources of CH₄ on global scale, both natural and anthropogenic methanogenesis , except biomass burning, geological and fossil fuel sources.

Sources of CH₄ emissions can be broadly grouped into three categories: biogenic, thermogenic , and pyrogenic origin (Neef et al. 2010), and they can be the direct result of either anthropogenic activities and/or natural processes. Biogenic sources contain CH₄-generating microorganisms (i.e., methanogens), where CH₄ is generated from degradation of organic matter in anaerobic conditions (Cicerone and Oremland 1988). Such sources comprise natural wetlands and rice paddies, oxygen-poor fresh water reservoirs such as dams, digestive systems of ruminants and termites, and organic waste deposits such as manure, sewage, and landfills. Thermogenic CH₄ consists of CH₄ formed over millions of years through geological sources as a result of slow transformation of organic matter into fossil fuel. It can be vented from subsurface into the atmosphere through natural features such as terrestrial seeps, marine seeps, and mud volcanoes. It is also vented to the atmosphere through anthropogenic activities of exploitation of coal, oil, and natural gas. Pyrogenic CH₄ is produced by incomplete combustion of biomass, soil organic carbon (SOC), and other organic materials during wild fires, combustion of biofuels, and combustions of fossil fuels. Each of these three sources of CH₄ are characterized by ranges of isotopic δ¹³C composition of CH₄, typically δ¹³CH₄ range of −55 to −70‰ for biogenic emissions, −25 to −45‰ for thermogenic emissions, and −13 to −25‰ for pyrogenic emissions (Bousquet et al. 2006; Monteil et al. 2011; Neef et al. 2010). The isotopic composition of atmospheric CH₄ measured at subset of surface atmospheric monitoring stations has been used to constrain the atmospheric CH₄ source (Neef et al. 2010; Bousquet et al. 2006; Monteil et al. 2011; Fisher et al. 2011). The CH₄ emissions by living plants under aerobic conditions seems to play less significant role in global CH₄ budgets, contrary to very large estimates which have been reported (Keppler et al. 2006; Nisbet et al. 2009).

Since large-scale global or regional CH₄ fluxes cannot be directly measured, the estimation of these quantities are generally obtained through two complementary techniques—“bottom-up” emission inventories and “top-down” inverse modeling (Alexe et al. 2015). Top-down Bayesian inverse modeling approach operates under well-defined mathematical framework which combines a priori information on CH₄ emissions, atmospheric observations, atmospheric chemistry, and transport models to determine optimal surface fluxes that best fit the atmospheric CH₄ concentration observation over the period of interest (Kasibhatla et al. 2000), given the prior fluxes and their uncertainties. In this approach, the global concentration fields are used as input to estimate the sources and sinks. The results of inverse modeling are called ‘a posteriori’ or improved estimates (Van Amstel 2012).

The top-down estimates are generally those of regional emissions derived from global atmospheric concentration fields, are not easily interpreted, and are difficult to assign into individual sources. The quality of estimates obtained through inverse modeling depends in most part on the quality of the observational data available for spatial and temporal domain of interest, and on the quality of chemistry and transport models, however. The global atmospheric inversions provide a time varying distribution of CH₄ fluxes with limited insight into underlying biogeochemical processes controlling the emissions, especially where several processes and different sources overlap in the same region (Kirschke et al. 2013). The global to-down CH₄ estimates for 2011 are presented in Table 7.2.

Table 7.2 Estimated top-down CH₄ budget for year 2011

The decadal global top-down CH₄ emissions modeled based on the observation of the global CH₄ burden and estimate of its atmospheric lifetime of ~9.1 years for 1980s, 1990s, and 2000s are estimated at 551 ± 41, 554 ± 42, and 548 ± 21 Tg CH₄ yr⁻¹, respectively, of which, the global anthropogenic burden was 34,843, 37,282, and 33,562 Tg CH₄ yr⁻¹ for 1980s, 1990s, and 2000s, respectively (Kirschke et al. 2013; Ciais et al. 2013). The estimated atmospheric CH₄ growth rate by inverse models for 1980s, 1990s, and 2000s were 34, 17, and 6 Tg CH₄ yr⁻¹ (Kirschke et al. 2013). The two-year average emissions for 2010 and 2011 estimated by four top-down models is 538.4 Tg CH₄ yr⁻¹ (Alexe et al. 2015). However, models did not aggregate the emissions into natural and anthropogenic sources.

The bottom-up approach incorporate the knowledge of small-scale processes, and with additional information and constraints it can project local emissions to large scales compatible with atmospheric signals. The approach includes process based models estimating CH₄ emissions and chemistry climate models estimating the OH sink. Bottom-up emission inventories based on energy use, agriculture activities, and emission factors from different sectors are used to provide yearly or decadal mean estimates of anthropogenic emissions. The decadal bottom-up CH₄ emissions, (i.e., sum of natural and anthropogenic sources) for 1980s, 1990s, and 2000s are 663 ± 127, 699 ± 138, and 678 ± 136 Tg CH₄ yr⁻¹, respectively (Table 1, Kirschke et al. 2013; Ciais et al. 2013). Global CH₄ burden derived from bottom-up approach are generally much larger than those obtained from top-down estimates.

7.3.1.1 Overview of Inventories of Methane

The global sources of CH₄ is composed of wide range of sources, balanced by much smaller number of sinks, and any imbalance in these sources and sinks results in change in the atmospheric CH₄ concentration. Of the many significant sources of CH₄ on a global scale, both natural and anthropogenic, the bulk have a common basis—that of microbial methanogenesis . Although CH₄ from biomass burning, vegetation, and geological or fossil fuel sources may be largely non-microbial in nature, understanding the processes that underpin microbially mediated CH₄ fluxes is central to quantifying and, potentially reducing emissions from all major sources. The current understanding of microbial methanogenesis and the interactions between different microbial communities that result in the bulk of CH₄ emissions to the global atmosphere is discussed in the following section.

7.3.1.1.1 Natural Methane Sources

The major natural CH₄ sources include wetlands, termites, and release from onshore and offshore geological sources. Living vegetation has also been suggested as an important natural source of CH₄ (Keppler et al. 2006). However, the significance of aerobic plant sources have been questioned, and the earlier emission estimates have been significantly revised downwards (Kirschbaum et al. 2006; Parsons et al. 2006).

7.3.1.1.2 Wetlands

In wetlands, CH₄ is formed under anaerobic conditions by microbial decomposition of OM. The CH₄ production occurs under anaerobic soil conditions in natural wetlands as well as in rice paddies. The CH₄ is formed by methanogenic organisms. It is estimated that 2–7% of the net primary productivity (NPP) in wetlands is emitted as CH₄ (Aselmann and Crutzen 1989). The three key determinants of CH₄ emission from wetlands are: (i) temperature (Christensen et al. 2003), (ii) water table depth (Macdonald et al. 1998), and (iii) substrate availability (Christensen et al. 2003). The degree of sensitivity of emissions to changes in these determinants remains poorly resolved. Generally temperature tends to be the dominant factor. For example, in a number of northern wetland sites, soil temperature variation accounted for 84% of observed variation in CH₄ emissions with strong positive response to increased temperature (Christensen et al. 2003). Emissions from Arctic, in particular, have the potential to increase significantly as the global temperature increases and the vast stores of SOC thaw (Schuur et al. 2011; Harden et al. 2012). Therefore, the impact of climate change on CH₄ emissions from natural wetlands in the 21st century could be substantial. It has been suggested that an increase in global temperature by 3.4 °C could result in a 78% increase in wetland CH₄ emissions, which will amplify total anthropogenic radiative forcing by between 3.5 and 5% by 2100 (Shindell et al. 2004). Decadal emissions from natural wetlands estimated by bottom-up for 1880s, 1990s, and 2000s ranged from 183 to 266, 169 to 265, and 177–284 Tg CH₄ yr⁻¹, respectively, accounting for 57–75% of natural CH₄ emissions. The decadal emissions for the same period estimated by top-down inverse modeling ranged from 115 to 231, 144 to 160, and 142 to 208 Tg CH₄ yr⁻¹, respectively (Table 1, Kirschke et al. 2013; Ciais et al. 2013), reflecting increasing emission trends from global natural wetlands.

7.3.1.1.2.1 Geological Methane

The natural CH₄ emissions from geological sources have often focused on CH₄ hydrates, also called clathrates (i.e., ice-like mixtures of CH₄ and water found in ocean sediments). The CH₄ clathrates, discussed further in the following section. Because these hydrates have potential of climatic warming, destabilizing them has received significant attention in recent years (Westbrook et al. 2009). Etiope (2012) opined that the estimates of emissions from hydrates remains highly speculative, and that the overall geological sources of CH₄ emission to the atmosphere is higher than is commonly reported. In addition to emissions from CH₄ hydrates, CH₄ emissions also occurs through sedimentary seepage, mud volcanoes, other macro- and micro-seeps, geothermal and volcanic emissions. It is estimated that natural geological emissions could be responsible for as much as 33–80 Tg CH₄ yr⁻¹, (Etiope 2010; Kirschke et al. 2013; Ciais et al. 2013), and has remained constant over the last three decades. The determinants of geological CH₄ emissions include seismic activity, tectonics, and magmatism. Therefore, the global atmospheric CH₄ budget is also influenced by geophysical processes (Etiope 2010). Hydrate CH₄ is a potential sources of natural gas, but the method for its exploitation has not been established. Estimated of CH₄ emissions from hydrates range from 2 to 10 Tg CH₄ yr⁻¹ (Ciais et al. 2013; Westbrook et al. 2009), and overall emissions are increasing at an alarming rate (Westbrook et al. 2009).

CH₄ emissions from thawing permafrost and CH₄ hydrates in the northern circumpolar region will become potentially important in the 21st century because they could increase dramatically due to the rapid climate warming of the Arctic and the large C pools stored there (Tarnocai et al. 2009; Walter Anthony et al. 2012). Supersaturation of dissolved CH₄ at the bottom and surface waters in the East Siberian Arctic Shelf indicate some CH₄ activity across the region, with a net sea–air CH₄ flux of 10.5 Tg CH₄ yr⁻¹, which is similar in magnitude to that for the entire ocean (Shakhova et al. 2010). However, it is not possible to know whether these fluxes are the result of recent changes in Arctic or it has always been present. The ebullition of CH₄ from thawing and decomposing lake sediments in north Siberia with an estimated flux of ~4 Tg CH₄ yr⁻¹ also demonstrates the activity of this region, and of its potential importance in the future CH₄ emissions (Walter et al. 2006; van Huissteden et al. 2011; Walter Anthony et al. 2012).

Other natural sources of CH₄ include termite, freshwater and lakes, wild animals, permafrost, wild fires and vegetation. Some termite species produce no CH₄ at all. But the emission rarely exceed half a microgram per termite per day for those that do. However, the shear mass of termites globally gives rise to considerable estimate in global budget. With improved measurements and understanding of differences in CH₄ production rates between different termite species, and issues associated with upscaling and effects of land use change, that the importance of termites as a global source of CH₄ is probably overstated (Bignell 2010).

The CH₄ emissions from biomass burning are the result of incomplete combustion, and it encompasses a wide range of sources, including fuel wood, charcoal, peat, savannah, agricultural residues, agricultural waste and municipal waste combustion. Higher CH₄ emissions from biomass burning generally occur at high water content and low O₂. In developing countries, biomass burning is common energy source for heating, cooking and lighting purposes. Differentiating between natural and anthropogenic CH₄ source is inherently difficult, however. The IPCC Assessment Report (AR5) estimated that emissions from wild fires remained constant from 1980s to 2000s at 1–5 Pg CH₄ yr⁻¹, while those from biomass burning increased from 31 to 37 Tg CH₄ yr⁻¹ in 1980s, and to 32–39 Tg CH₄ yr⁻¹ in 2000s (Ciais et al. 2013). During the decade of 2000s, the natural CH₄ sources are estimated to account for 35–50% of the decadal global mean CH₄ emissions with the wetlands being the dominant source. The climate-driven changes in emissions from wetlands are the main drivers of the global inter-annual variability in CH₄ emissions. The sum of all natural emission estimates other than wetlands is still uncertain based on bottom-up studies and range of 238–484 Tg CH₄ yr⁻¹ for 2000–2009 (Kirschke et al. 2013; Ciais et al. 2013).

7.3.1.1.3 Anthropogenic Methane Sources

The anthropogenic emissions of CH₄ ranged between 50 and 65% of the global emissions since 2000s (Ciais et al. 2013), and the anthropogenic sources of CH₄ emissions can be grouped by sector into (i) agriculture, (ii) energy, (iii) waste, and (iv) industry. Figure 7.6 shows contribution of each sector to total anthropogenic CH₄ emission.

Fig. 7.6

Contribution of different sectors to anthropogenic CH₄ emissions calculated based on average emissions from 1990 to 2010

7.3.1.1.3.1 Agriculture

Agricultural sector is the largest contributor to global emissions of non-CO₂ GHG emissions, accounting for about 51% of anthropogenic CH₄ emissions (Fig. 7.6). The dominant source of agricultural CH₄ emission is the enteric fermentation which constitute nearly 60% of agricultural emissions, followed by emissions from rice cultivation, other agricultural activities, and manure management. China is the largest agricultural emission source, constituting 15.45% of global agricultural emissions. Other significant emitters are India, Brazil, and USA, among others. The top 10 emitters are responsible for 55% of agricultural emissions (Karakurt et al. 2012).

7.3.1.1.3.1.1 Enteric Fermentation

About 90% of annual anthropogenic CH₄ emissions in rural communities is due to rumen fermentation (Abberton et al. 2008). The enteric fermentation represents about 80% of CH₄ emission produced by livestock, but the emissions vary over time and between different regions of the world. The CH₄ is formed by methanogenic bacteria under anaerobic conditions in the rumen of domestic animals, a process which enables these animals to utilize the energy more efficiently from low-quality feeds like grass and fodder with high cellulose content. Pseudo-ruminants like horses and pigs also produces CH₄, but in much smaller quantities. Rumen methanogenesis results in the loss of 6–10% of gross energy intake or 8–14% of digestible energy intake of ruminants (Cottle et al. 2011). The bulk (>90%) of CH₄ produced in rumen is emitted through belching, and some dairy cattle emit several hundred liters of CH₄ per day. The CH₄ emissions produced by domestic ruminants depend on the feeding intake, energy consumption, rumen condition, and digestibility. The CH₄ emissions from ruminant livestock are highly dependent on demand pressure, and the global trend of increasing consumption of meat and dairy products with increased CH₄ emissions. In 2005, CH₄ emissions from ruminant livestock were estimated at 72 Tg CH₄ yr⁻¹, and projected to be as high as 100 Tg CH₄ yr⁻¹ in 2010 (Kelliher and Clark 2010). The IPCC AR5 estimated decadal changes in ruminant CH₄ emission for 1980s, 1990s, and 2000s at 81–90, 82–91, and 87–94 Tg CH₄ yr⁻¹, respectively (Ciais et al. 2013). CH₄ emissions from enteric fermentation are projected to increase by 22% by 2030, mostly due to increases in livestock production. The largest increases in CH₄ emission is expected to occur in Africa and Asia (EPA 2012).

7.3.1.1.3.1.2 Manure

CH₄ from animal manure is formed in anaerobic conditions when stored in lagoons or in manure tanks. Microbial methanogenesis in manure has the potential to produce significant amount of CH₄ because of high availability of OC substrates and anoxic conditions that tend to prevail. However, when manure is kept in contact with O₂ (i.e, spread on the fields), the CH₄ production is minimal. The quantity of CH₄ emitted from manure management operations is a function of three primary factors: (i) type of treatment or storage facility, (ii) the ambient climate, and (iii) the composition of manure. Higher ambient temperature and moisture conditions favor CH₄ production. The composition of manure is directly related to animal type and diet. Therefore, manure management system has significant influence on CH₄ emission.

7.3.1.1.3.1.3 Rice Paddies

CH₄ is formed by methanogenic organisms in flooded paddy soils during a growing season, escapes through bubbling and diffusion, and rice stems. Draining of the fields stops CH₄ formation due to aeration. Globally, about 80 million ha (Mha) of harvested wetland rice are potential sources of CH₄. Most paddies are submerged for about one third of the time, although practices vary widely around the world based on rice variety, culture, and water availability. The estimated emissions are 200–500 mg m⁻² during an average rice growing season of 130 days. Rice is a staple food for many parts of the world, and with the population growth, the CH₄ emissions from rice cultivation are expected to increase. The IPCC AR5 estimated that the CH₄ emission from rice production decreased from 41–47 Tg CH₄ yr⁻¹ in 1980s to 33–40 Tg CH₄ yr⁻¹ in 2000s (Ciais et al. 2013). However, CH₄ emission from rice cultivation will increase only marginally by 2% between 2005 and 2030, mainly due to decrease in per capita consumption (EPA 2012) and improved water management regime.

7.3.1.1.4 Waste

Emissions from waste account for 21% of the anthropogenic emissions (Fig. 7.6). Major sources are waste water (41%) and landfilling solid waste (59%). The largest global emitter under this category is USA (13.8%), followed by China and India.

7.3.1.1.4.1 Waste Water Treatment

CH₄ is emitted from anaerobically treated human and industrial liquid wastes, mainly agribusiness. The potential amount of CH₄ formed in different treatment systems depend on characteristics like chemical and biological oxygen demand in the degradation and the storage time and temperature. Industrial waste water treatment such as food processing and pulp and paper facilities with high organic material loading tend to be the main contributors of CH₄ emissions. CH₄ is emitted incidentally or deliberately during handling and treatment of municipal wastewater. The organic material in wastewater produces CH₄ when it decomposes anaerobically (Talyan et al. 2007). Estimated CH₄ emissions from domestic waste are 29 Tg CH₄ yr⁻¹ based on emission factor of 0.2–0.4 g CH₄ g⁻¹ of chemical oxygen demand (Doorn et al. 2000). Global emissions from wastewater have increased by 35% between 1990 and 2005, and the main driver for the increasing domestic wastewater emissions is the population increase (EPA 2012).

7.3.1.1.4.2 Landfills

CH₄ is emitted from anaerobic decomposition of organic material in landfills. The O₂ trapped in the landfill is rapidly consumed by microorganisms in the closed landfill solid waste. CH₄ constituting about 50% of storage gas is produced as a result of anaerobic decomposition of organic materials in landfilling solid wastes. The major drivers of emission are the amount of organic material in landfills, the extent of anaerobic decomposition, the thickness as well as physical and chemical properties of the landfill cover material, the seasonal variation in CH₄ oxidation rate, and the level of landfill CH₄ collection and combustion—such as energy use or flaring (Themelis and Ulloa 2007; Lou and Nair 2009). Early estimates for landfill CH₄ emissions were as high as 70 Tg CH₄ yr⁻¹, but successful implementation of mitigation strategies and capture of CH₄ for beneficial energy use has seen the reduction in emission from landfills over the years (Borgner and Spokas 2010). Estimated decadal CH₄ Emissions from waste and landfills for 1980s, 1990s and 2000s were 50–60, 63–68, and 67–90 Tg CH₄ yr⁻¹, respectively (Ciais et al. 2013; Kirschke et al. 2013). Emissions from wastewater are projected to increase by 28% by 2030, with the highest increase occurring from Africa, Middle East, and Central and South America (EPA 2012).

7.3.1.1.5 Fossil Energy

CH₄ is the principal component of natural gas, accounting for 95% of pipeline quality gas. Much of the CH₄ emissions attributable to natural gas each year is derived from release during extraction, and processing, and distribution of natural gas. Oil production and processing upstream of oil refineries can also emit CH₄ in significant quantities since natural gas is often found in conjunction with petroleum deposits (Mitchell 1994). In both natural gas and oil production systems, CH₄ is a fugitive emission from leaking equipment, system fault, and deliberate flaring and venting at production field, oil refinery facilities, natural gas transmission line, compressor station leakage, natural gas storage facilities and natural gas distribution lines (Karakurt et al. 2012). Some CH₄ is also emitted during incomplete fossil fuel combustion. The amount of CH₄ released is related to combustion conditions and the content of the fuel (Karakurt et al. 2012).

Coal mining and extraction also release CH₄, and constitute one of the largest individual source activities of anthropogenic CH₄ emission. CH₄ is formed as a part of the geological process of coal formation, where large deposits of CH₄ can remain trapped within coal seams and surrounding rock strata and are liberated when pressure above or surrounding the coal bed is reduced as a result of natural erosion, faulting, or mining operations. Because CH₄ is explosive, it must be removed from underground mines high in CH₄ as a safety precaution. CH₄ concentrations between 5 and 15% in the air of coal mines represent an explosion hazard, and so ventilation is commonly employed to get rid of this CH₄.

The quantity of CH₄ emitted from mining operation is a function of coal rank and coal depth. Pressure increases with depth, and prevents CH₄ from migrating to the surface, and as a result, underground mining operations typically release more CH₄ than surface or opencast mining due to lower pressure and coal rank (EPA 1993). Coal rank and permeability, depth of coal seam, geologic parameters and the amount of production are among the factors affecting the quantity of CH₄ released (Karakurt et al. 2012). In addition, post mining processing of coal and abandoned mines also release CH₄. Between 1990 and 2005, global CH₄ emission from natural gas and oil systems are estimated to have increased by 21%, and from 2005 to 2030 by 31% (EPA 2012). Between 1990 and 2005, CH₄ emissions from coal mining are estimated to decrease by 2%, but projected to increase by 50% from 2005 to 2030 (EPA 2012). Based on top-down inversion models, decadal emissions from fossil fuels for 1980s, 1990s, and 2000s are estimated at 94.19 ± 19, 95 ± 11, and 96 ± 19 Tg CH₄ yr⁻¹ respectively, while 89, 84 ± 18, and 96 ± 11 Tg CH₄ yr⁻¹ have been estimated for the same period based on bottom-up approach (Kirschke et al. 2013). Some indications suggest that emissions from human activities such as intensification of hydraulic fracturing for the shale natural gas and oil extraction in USA and the global coal mining and especially Chinese and Indian economies have increased since 2007 (BP 2015; Karion et al. 2013; Ohara et al. 2007; Cooper et al. 2016). However, a full understanding of the hydraulic fracturing impact on GHG emissions require monitoring over the gas well lifetime and also analysis of transport distribution system. Some studies have indicated that increase in natural wetlands emission in response to high temperatures in northern high latitudes and increased rainfall over the tropical wetlands could be responsible for increase in emissions after 2006 (Bloom et al. 2010; Dlugokencky et al. 2011; Kirschke et al. 2013). Since 2007, atmosphere has become more depleted in ¹³C an indication that the growth is dominated by ¹³C richer wetland and ruminants emissions (NOAA 2015; Sherwood et al. 2016).

7.3.2 Methane Sinks

There are three sinks for CH₄ emitted to the atmosphere: (i) destruction of CH₄ by hydroxyl (OH) radicals in the troposphere and stratosphere, (ii) destruction by oxidizing bacteria (methanotrophs ) in soils, and (iii) reactions with Cl in the troposphere and ocean surface. The primary sink for atmospheric CH₄ is the photochemical oxidation of CH₄ initiated by the reaction with OH radicals, mostly in the troposphere and stratosphere, which removes an amount equivalent to about 90% of the global CH₄ surface emission, and about 9% of the total atmospheric burden of 4700–4900 Tg CH₄ (Voulgarakis et al. 2013). The atmospheric lifetime of CH₄ with respect to OH and the atmospheric burden of 4800 Tg CH₄ is 7–11 years. Recent estimate of atmospheric CH₄ lifetime is 9.1 ± 0.9 years (Prather et al. 2012).

Additional oxidation sinks include methanotrophic bacteria in aerated soils which uses CH₄ as a source of C and energy (Curry 2007; Zhuang et al. 2004) which accounts for ~4% and reaction with chlorine radicals and atomic oxygen radical in the stratosphere (Cicerone and Oremland 1988; Neef et al. 2010; Nisbet et al. 2014) which oxidizes ~3%. A small sink of atmospheric CH₄ is also suspected but still debated, involves reactions with chlorine radicals from sea salt in the marine boundary layer (i.e., surface waters of the ocean) (Allan et al. 2007) estimated to oxidize ~3%. Decadal sink for atmospheric CH₄ for 1980s, 1990s and 2000s based on top-down atmospheric inversion models is estimated at 460–559, 518–579, and 514–560 Tg CH₄ yr⁻¹, respectively. The estimated decadal sink based on bottom-up for 1980s, 1990, and 2000s is 567 ± 147, 599 ± 69, and 632 ± 40 Tg CH₄ yr⁻¹, respectively (Table 1, Kirschke et al. 2013; Ciais et al. 2013).

7.4 Contemporary Global Methane Budget

Currently, source and sink processes of CH₄ are not accurately quantified despite several decades of research (Lassey and Ragnauth 2010), and there are still large uncertainties on estimated emissions and losses. The global sinks that remove tropospheric CH₄ has an uncertainty of about 20% (Ciais et al. 2013), anthropogenic sources have uncertainty range of ~30% (Kirschke et al. 2013; Ciais et al. 2013), while wetland emissions have uncertainty of 40% (Kirschke et al. 2013). The global atmospheric CH₄ budget is determined by many terrestrial and aquatic surface sources balanced primarily by one major sink in the atmosphere, and the atmospheric growth rate results from the balance between emissions and sinks (Kirschke et al. 2013).

Significant progress has been made using inverse modeling to quantitatively link regional sources and sinks, where inferred emissions are linked with photochemical sinks—tropospheric and stratospheric OH concentrations, and matched to atmospheric observations, and therefore, reducing the uncertainties (Bousquet et al. 2006). For example, using inversion techniques and dry air column averaged mole fractions of atmospheric CH₄ from the Thermal And Nuclear infrared Sensor for C Observations Fourier Transform Spectrometer (TANSO-FTS) on board Greenhouse Gases Observing Satellite (GOSAT), two-year average emissions based on monthly average CH₄ emission from January 2010 to December 2011 estimated at 537.3–540.4 Tg CH₄ yr⁻¹ (Alexe et al. 2015). Inversion techniques have also showed that global CH₄ emissions have increased significantly since 2006, with much higher emissions during 2007–2010 period compared to the average emissions during 2003–2005 period (Bergamaschi et al. 2013). Inversion techniques allowed for the attribution of the observed increases spatially, to tropics and mid-latitudes of the Northern Hemisphere, and largely to anthropogenic sources (Bergamaschi et al. 2013). In addition, significant interannual variations in CH₄ emissions from wetlands and biomass burning were also identified that were superimposed on increasing trends of anthropogenic CH₄ emissions (Bergamaschi et al. 2013). Inversion of atmospheric measurements of CH₄ from surface stations estimated that global CH₄ emissions for the 2000–2009 period range from 526 to 569 Tg CH₄ yr⁻¹ (Table 1). The total atmospheric oxidation range from 514 to 560 Tg CH₄ yr⁻¹, indicating a small imbalance of about 3 Tg CH₄ yr⁻¹, which is in line with the small atmospheric growth rate of 6 Tg CH₄ yr⁻¹ observed between 2000 and 2009. The bottom-up models and inventories showed a larger global CH₄ emissions, ranging from 542 to 852 Tg CH₄ yr⁻¹ (mean 678 Tg CH₄ yr⁻¹, Table 1). Reasons for the discrepancies between top-down inversion and bottom up inventories are still debated. Thus, upward re-evaluation of geological (Etiope et al. 2008) and freshwater (Walter et al. 2007; Bastviken et al. 2011 emission sources could account for the observed discrepancy. An ensemble of Atmospheric Chemistry models estimated an averaged total loss ranging from 592 to 785 Tg CH₄ yr⁻¹ (mean 632 Tg CH₄ yr⁻¹ during 2000–2009 decade (Lamarque et al. 2013), suggesting an imbalance of about 45 Tg CH₄ yr⁻¹. The observed average atmospheric growth rate during this period is estimated at 6 Tg CH₄ yr⁻¹ (Dlugokencky et al. 2011). Unlike top-down inversions which have constrained OH fields, no constraint applicable to the sum of emissions in the bottom-up approach. Although top-down inversions can help in constraining global CH₄ emissions in the global CH₄ budget, they do not provide detailed resolution of mix of sources that is achieved by the bottom-up approaches, and thus provide limited information about emission and sink processes.

Based on the pre-industrial atmospheric CH₄ concentration of 722 ppb, the atmospheric CH₄ burden was 2042 Tg CH₄ (Dlugokencky et al. 2005) compared to the current atmospheric burden estimated at 4800 Tg (Table 2, Voulgarakis et al. 2013).

7.5 Atmospheric Chemistry of Methane

In addition to its effects on radiation energy balance and Earths’ climate, CH₄ plays important chemical roles in the atmosphere, and its destruction with the OH radical in atmosphere initiates a chain of reactions that can affect tropospheric air quality and ozone formation. Therefore, changes in atmospheric CH₄ abundance are of substantial importance to atmospheric photochemistry . The chemical reactions that destructively oxidize atmospheric CH₄ affect the chemical state of the atmosphere through the products of reactions and through consumption of reactant species. The oxidation of CH₄ produces CO, CO₂, H₂O, H₂, and CH₂O, and also consume OH. The reaction pathways for CH₄ destructive oxidation affect tropospheric ozone (O₃) amounts, and they produce important quantities of stratospheric H₂O. Therefore, changes in atmospheric CH₄ concentration have caused changes in concentrations of O₃ and OH radical. In addition, stratospheric CH₄ reacts with chlorine atoms to form HCl, a reservoir species for chlorine atoms. Similarly, a portion of the flux of hydrogen atoms carried upward into the stratosphere in CH₄ escapes into space which represents a source of atmospheric O₂.

The OH is produced by the action of ultraviolet radiation (λ = 315 nm) on O₃ and water vapor (Eq. 7.1):

$${\text{O}}_{3} + {\text{hv}} \to {\text{O}}(^{1} {\text{D}}) + {\text{O}}_{2}$$

(7.1)

The electronically excited O₂ atoms (O(¹D)) produced quenched in collision with N₂ and O₂ (Eq. 7.2) and then ozone is reformed (Eq. 7.3):

$${\text{O}}(^{ 1} {\text{D}}) + {\text{N}}_{2} \to {\text{O}} + {\text{N}}_{2}$$

(7.2)

$${\text{O}} + {\text{O}}_{2} + {\text{M}} \to {\text{O}}_{3} + {\text{M}}$$

(7.3)

where, M = N₂, O₂, or any molecule which collides to stabilize O₃. Some O(¹D) about 1% react with water vapor to produce hydroxyl radical (OH) (Eq. 7.4).

$${\text{O}}(^{ 1} {\text{D}}) + {\text{H}}_{2} {\text{O}} \to 2{\text{OH}}$$

(7.4)

Net reaction:

$${\text{O}}_{3} + {\text{H}}_{2} {\text{O}} \to 2{\text{OH}} + {\text{O}}_{2}$$

(7.5)

The OH is responsible for the removal of almost all trace gases from the atmosphere. Despite the fact that the global average mixing ratio of OH in the atmosphere is 4 × 10⁻¹⁴, it is this ultra-minor constituent which is responsible for the first step in the removal of most natural and anthropogenic gases from the atmosphere (Crutzen 1995). The first oxidation step which utilizes OH and rapidly following reactions lead to compounds which are readily removed from the atmosphere by precipitation or uptake at the Earth’s surface.

About 90% of atmospheric O₃ is present in the stratosphere, and only about 10% of the atmospheric O₃ is located in the troposphere, with volume mixing ratio in unpolluted environment ranging from 15 ppbv at low altitudes in Southern Hemisphere to about 100 ppbv in the upper troposphere. But this relatively small amount of O₃ in the troposphere governs the oxidation processes in the Earth’s atmosphere through the formation of OH, and is responsible for absorption of solar ultraviolet (UV) radiation of wavelengths shorter than 310 nm. It leads to production of electronically excited O(¹D) atoms which have enough energy to react with water vapor to produce OH (Eqs. 7.1–7.5). OH controls the atmospheric lifetime of many gases. The production of O(¹D) is temperature and solar energy dependent. Thus, high levels of O₃ shortwave radiation and humidity favor production of OH. Abundance of water vapor is largely determined by the temperature. Thus, the primary formation of OH is controlled by solar ultraviolet radiation flux, dependent on overhead O₃ column as well as the local O₃ and water vapor concentrations (Lelieveld et al. 2004).

The OH has a short lifetime on the order of few seconds (Lelieveld et al. 2004), and its quantification is particularly challenging. Modeling becomes an essential tool to probe its spatial and temporal variability. About 40% of OH radicals react with CH₄ and its oxidation products, while 60% of OH react with CO. Therefore, oxidation of CH₄ has significant influence on tropospheric and stratospheric chemistry. The oxidation of CH₄ strongly affects the atmospheric budgets of OH and O₃ (Crutzen 1973). Chemistry—transport model studies have predicted OH decreases and atmospheric increases in CH₄ due to increases in CH₄ and CO emissions both of which consume OH. For example a 6% increase in CH₄ atmospheric lifetime from 1992 to 2050 (Lelieveld et al. 1998) and 13% increase in CH₄ lifetime in 2100 (Wild and Palmer 2008). The mechanism of CH₄ oxidation and the products that are formed are dependent on the concentration of nitrogen oxides (NO_x). In NO_x-poor environments, including marine areas, the free troposphere over the tropics, and most of the Southern Hemisphere, O₃ is destroyed and OH runs down (Crutzen 1979; Lelieveld et al. 2004). Therefore, in NO_x-depleted system, increase in CH₄ enhances OH loss, and atmospheric lifetimes of CH₄ increases. In NO_x-rich environments of Northern Hemisphere and planetary boundary layer of the tropics during dry season, where air is polluted or moderately dirty tropospheric air and all of the stratosphere, the oxidation of CH₄ by OH in the troposphere leads to the formation of formaldehyde (CH₂O), CO, and O₃ in the presence of sufficiently high levels NO_x (Wuebbles and Hayhoe 2002) and yield hydrogen oxides (HO and HO₂) (Cicerone and Oremland 1988; Crutzen and Zimmermann 1991). Therefore, in NO_x-rich environments, O₃ is formed, and OH radicals are recycled (Lelieveld et al. 2004). NO_x is an important component of local pollution in urban areas, and its increased presence results in O₃ formation (Eq. 7.1) that in turn increases OH production. Modeling has shown that increasing NO_x emission causes increase in global OH, and hence increased CH₄ sink (Fuglestvedt et al. 2000). Most current models suggest that an increase in CH₄ will decrease OH and increase O₃ throughout the troposphere (Wuebbles and Hayhoe 2002).

Destructive oxidation reactions with OH radical which removes CH₄ from the atmosphere are outlined in Eqs. 7.6–7.14:

$${\text{CH}}_{4} + {\text{OH}} \to {\text{CH}}_{3} + {\text{H}}_{2} {\text{O}}$$

(7.6)

$${\text{CH}}_{3} + {\text{O}}_{2} + {\text{M}} \to {\text{CH}}_{3} {\text{O}}_{2} + {\text{M}}$$

(7.7)

$${\text{CH}}_{3} {\text{O}}_{2} + {\text{HO}}_{2} \to {\text{CH}}_{3} {\text{O}}_{2} {\text{H}} + {\text{O}}_{2}$$

(7.8)

$${\text{CH}}_{3} {\text{O}}_{2} {\text{H}} + {\text{hv}}\left( { {\le} 330\;{\text{nm}}} \right) \to {\text{CH}}_{2} {\text{O}} + {\text{OH}}$$

(7.9)

$${\text{Net:}}\,{\text{CH}}_{4} + {\text{O}}_{2} \to {\text{CH}}_{2} {\text{O}} + {\text{H}}_{2} {\text{O}}$$

(7.10)

The CH₂O is an important intermediate in the removal processes of CH₄ and hydrocarbons as well as being important in general chemical reactivity of the atmosphere. Regardless of NO_x levels, CH₄ oxidation is responsible for the formation of much of CH₂O in the atmosphere. It is estimated that production of CH₂O in the atmosphere is in the order of 10¹⁴ g yr⁻¹ (Lowe and Schmidt 1983; Wuebbles and Hayhoe 2002). Once generated, CH₂O is quickly destroyed by direct insolation. The photochemical breakdown of formaldehyde to CO occurs as follows:

$${\text{CH}}_{2} {\text{O}} + {\text{hv}}\left( { {\le} 350\,{\text{nm}}} \right) \to {\text{H}} + {\text{CHO}}$$

(7.11)

$${\text{H}} + {\text{O}}_{2} + {\text{M}} \to {\text{HO}}_{2} + {\text{M}}$$

(7.12)

$${\text{CHO}} + {\text{O}}_{2} \to {\text{CO}} + {\text{HO}}_{2}$$

(7.13)

$${\text{Net:}}\,{\text{CH}}_{2} {\text{O}} + hv + 2{\text{O}}_{2} \to {\text{CO}} + 2{\text{HO}}_{2}$$

(7.14)

During the daytime, photolysis is the dominant loss process for CH₂O (Eqs. 7.11–7.14), while in the polluted atmospheres reactions with OH and hydroperoxyl are similar to photolysis degradation (Wuebbles and Hayhoe 2002) At night, CH₂O reacts with NO₃, however, it is a slow process compared to daytime (Wuebbles and Hayhoe 2002). The theoretical maximum yield of O₃ from CH₄ molecule oxidized all the way to CO₂ and H₂O is 5 O₃ molecules. Nevertheless, this maximum yield in never realized under actual atmospheric conditions because of competing reactions. When levels of NO_x are sufficiently high, peroxy radicals (HO₂, CH₃O₂) react exclusively with NO and all the CH₂O formed is photolyzed by the radical path (Eq. 7.14). The CH₂O and CO are important intermediates in the atmospheric process.

Based on the reaction sequences for CH₄, CH₂O, and CO in polluted and unpolluted environments, the effect of the complete oxidation of 1 mol of CH₄ results into increase of 0.4–0.5 and 3.6–3.8 mol OH and O₃, respectively in high NO_x (Crutzen 1973; Wuebbles and Hayhoe 2002), and decrease in OH ranging from −3.5 to −3.9 mol of OH in low NO_x environment (Wuebbles and Hayhoe 2002). The destruction of OH depends on the OH concentration, methyl peroxide concentration, the methyl peroxide chemical reaction pathways and the heterogeneous removal rates of important intermediate species.

The alternative pathways are:

$${\text{CH}}_{2} {\text{O}} + {\text{hv}}\left( { {\le} 350\,{\text{nm}}} \right) \to {\text{CO}} + {\text{H}}_{2}$$

(7.15)

$${\text{CH}}_{2} {\text{O}} + {\text{OH}} \to {\text{CHO}} + {\text{H}}_{2} {\text{O}}$$

(7.16)

$${\text{CHO}} + {\text{O}}_{2} \to {\text{CO}} + {\text{HO}}_{2}$$

(7.17)

$${\text{Net:}}\,{\text{CH}}_{2} {\text{O}} + {\text{OH}} + {\text{O}}_{2} \to {\text{CO}} + {\text{H}}_{2} {\text{O}} + {\text{HO}}_{2}$$

(7.18)

The remaining CH₄ is destroyed by OH, Cl or O(¹D) atoms in the stratosphere. A small fraction of CH₄ goes through the stratosphere to the mesosphere where a very short ultraviolet (UV) light, i.e., Lyman alpha radiation—121.6 nm) destroys CH₄ photolytically. The largest single sink of atmospheric CH₄ is its reaction with OH. It is estimated that 85–90% of CH₄ in the troposphere is destroyed by oxidation with OH. A complete oxidation of CH₄ yields CO₂ and H₂O:

$${\text{CH}}_{4} + 2{\text{O}}_{2} \to {\text{CO}}_{2} + 2{\text{H}}_{2} {\text{O}}$$

(7.19)

The atmospheric lifetime of CH₄ ranges from 9 to 11 years with CO₂ and H₂O the eventual oxidation products (Prather et al. 2012; Holmes et al. 2013; Myhre et al. 2013). Although Eq. 7.19 is a simple and clear oxidation reaction, it does not describe the mechanism through which the atmosphere oxidizes CH₄. In the atmosphere, the process is initiated by OH radicals and not O₂ and it requires light. Altogether, the oxidation of CH₄, CH₂O and CO to CO₂ leads to loss of OH and HO₂ in NO_x-poor environments and a gain of O₃ in NO_x-rich environments. Generally, increases in NO_x emissions have been associated with more OH generation because NO_x generally leads to O₃ production which is the main source of OH, and NO_x-rich environments favor more secondary OH production through conversion of HO₂ to OH (Voulgarakis et al. 2013; Crutzen 1973).

In the troposphere, especially in the altitude range of 0–6 km, CH₄ oxidation consumes O₃, OH, and HO₂ in producing CO₂, H₂O, and H₂. The flow of CH₄ in the atmosphere and the products of its oxidation is illustrated in Fig. 7.7. Reaction with stratospheric OH is the dominant sink of CH₄, followed by that of Cl atoms. A fraction of H atoms that are released in CH₄ oxidation and from photochemical decomposition of H₂O in the stratosphere are subsequently released to space. The H atoms in the thermosphere and exosphere are furnished in large part by CH₄ oxidation. Therefore, temporal increase in CH₄ concentrations is also causing an increase in the rate of escape of H to space (Ehhalt 1986). Oxidation processes removes a range of environmentally important species from the atmosphere. Besides its role in CH₄ oxidation, OH is involved in removing other trace gases such as nitrogen oxides (NO_x), CO, non-CH₄ volatile organic compounds, and ozone-depleting substances such as hydrofluorocarbons (HFCs) (DeMore 1996; Voulgarakis et al. 2013). Tropospheric oxidation depends heavily on the level of OH radical and its geographical distribution. A potentially important consequence of CH₄ oxidation is that of CH₄, CO, and OH perturbations. Reaction with OH is also the major sink of atmospheric CO, and the reactions of OH with CH₄ and CO suppress OH concentrations. The chemical coupling of CH₄ and OH leads to significant amplification of CH₄ emissions because increased CH₄ emissions decrease tropospheric OH which increases the CH₄ lifetime and thereby its concentration (Myhre et al. 2013; Holmes et al. 2013). Current increase in CH₄ is probably decreasing OH radical concentration, resulting into further increase in atmospheric CH₄ concentrations.

Fig. 7.7

Schematic illustration of the flow of CH₄ from several representative biological and abiogenic sources. The CH₄ oxidation occurs in sediments, waters, and the atmosphere. In the atmosphere, most methane is destroyed in the troposphere by OH radical, but 10–20% is transported upward to the stratosphere where the remainder is destroyed

7.6 Biological Aspects of Methanogenesis

About 70% of the atmospheric CH₄ originates from biogenic sources (Conrad 2009), and biological processes under anoxic environment are the main biogenic source of atmospheric CH₄ emission. Commonly described methanogens microorganisms belong to domain archaea, kingdom Archaebacteria and phylum Euryarchaeota (Nazaries et al. 2013; Thauer et al. 2008) and share a set of bidirectional enzymes responsible for their respective metabolisms (Thauer 2011). However, archaea members of phylum Bathyarchaeota from coal beds with genetic potential to generate CH₄ been identified by DNA sequencing (Evans et al. 2015; Lloyd 2015). Coal bed methanogens have a potential to increase natural gas productivity from existing natural gas wells.

Anaerobic archaea are major contributors to global CH₄ cycling, and estimated to produce 1.0 Pg CH₄ yr⁻¹, with equal amount estimated to be oxidized archaeal methanotrophs (Reeburgh 2014). They produce CH₄ as the major product of their energy conserving metabolism in anaerobic respiration. All methanogens archaea characterized so far are strictly anaerobic respiration. They are a large and phylogenetically diverse group, which are classified into two classes: Methanobacteria and Methanococci, and six identified orders: Methanobacteriales, Methanococcales, Methanomicrobiales, Methanosarcinales, Methanopyrales, and Methanocellaces, which are further divided into 12 families (Liu and Whitman 2008; Offre et al. 2013; Nazaries et al. 2013). All of methanogenic archaea are obligate CH₄-producers that obtain all or most of their energy from methanogenesis . Methanogens are the only Archaea currently cultivated that are cosmopolitan, and are found in wide variety of the anaerobic environments on Earth. They are usually abundant where electron acceptors such as NO₃ ⁻, Fe⁺, and SO₄ ²⁻ are in short supply. In addition to temperate habitats, they are also common in environments of extreme temperatures, salinity, and pH. The common habitats of methanogens include marine and fresh water sediments, flooded soils, human and animal gastrointestinal tracts, termites, anaerobic digesters, landfills, geothermal systems, polar systems and heartwood of trees (Liu and Whitman 2008). Although obligate anaerobes, a recent study confirmed that methanogens are ubiquitous in aerobic soils and can actively produce CH₄ as soon as conditions changes to anoxic (Angel et al. 2012).

The methanogenesis pathway is a complex, and requires a number of unique co-enzymes and membrane-bound enzyme complexes for coupling to the proton motive force. The complexity and uniqueness of methanogenesis as a form of anaerobic respiration resides in requirement of six unusual co-enzymes—ferredoxin, methanofuran, tetrahydromethanopterin, coenzyme F₄₂₀, coenzyme M, and coenzyme B; a multistep pathway and several unique membrane-bound enzyme complexes coupled to the generation of a proton gradient driving ATP synthesis (Ferry 2010a; Nazaries et al. 2013). The three main methanogenic substrates are (i) CO₂, (ii) acetate, and (iii) methyl group containing compounds such as methanol, methylated amines, and methylated sulfides. Several methanogenic pathways that rely on these substrates have been described for CH₄ production: (i) hydrogentrophic and formatrophic methanogens, (ii) acetotrophic methanogens, and (iii) methylotrophic methanogens. In hydrogenotrophic and formatrophic methanogenesis , CO₂ is reduced with H₂ or formate as electron donors, to form CH₄ through formyl, methylene and methyl levels (i.e., methanol reduction with H₂). In acetotrophic methanogenesis, acetate is split into CH₄ and CO₂ after activation to acetyl-CoA and then split into methyl-CoM and CO (i.e., fermentation of acetate). In methylotrophic methanogen, methyl is reduced to CH₄ and CO is oxidized to CO₂ (i.e., dismutation of methylated compounds such as methanol, methylamines, dimethylsulfide (DMS) or methanethiol) (Liu and Whitman 2008; Ferry 2010b). In methyl containing compounds, Methanosarcinales and Methanosphaera convert methylated compounds by first transferring methyl group to protein (corrinoid) and then to coenzyme M. Methyl-CoM then enters the methanoganesis pathway and is reduced to CH₄ and CO₂ (Ferry 2010a).

Whereas most cultivated methanogens reduce CO₂ with H₂, only members of Methanosarcinales have the ability to produce CH₄ from the fermentation of acetate and dismutation of methylated compounds. The recently discovered methanogenic Thermoplasmata reduce methanol with H₂ (Dridi et al. 2012; Paul et al. 2012), and might also use methylamines as methanogenic substrate (Poulsen et al. 2013). Currently one of these methanogenic metabolisms have been found in bacteria or eukaryotes (Offre et al. 2013). Acetate is the major intermediate in anaerobic food chain, and about two-thirds of biologically generated CH₄ and released to the atmosphere by archeal methanogenesis is derived from acetate. However, only two genera—Methanosacrcina and Methanosaeta are known to use acetate (Jetten et al. 1992). The reduction of CO₂ accounts for the rest of the archeal contribution to atmospheric CH₄, with minor amounts of CH₄ produced by the dismutation of methyl compounds (Ferry 2010b).

Methanogenic archaea have also been found in oxic environments including various aerated soils (Angel et al. 2012) and the oxygenated water column of oligotrophic lake (Grossart et al. 2011). The methanogens in aerated soils become active under wet anoxic conditions (Angel et al. 2012) and those in oxygenated lake waters are attached to photoautotrophs, which might enable anaerobic growth by supply of methanogenic substrates (Grossart et al. 2011). Several other processes, such as the microbial decomposition of methylphosphonate could be responsible for CH₄ production in oxygenated waters (Keppler et al. 2009), and fungi (Lenhart et al. 2012) and plants (Bruhn et al. 2012) are possible sources of CH₄ in aerated soils. The global significance of these alternative aerobic methanogenic pathways remains to be assessed.

Methanogenic archaea engage in various syntropic partnerships with anaerobic bacteria, protozoa, and fungi (Stams and Plugge 2009) that involve the transfer of electrons from a fermentative organism to the methanogen through carrier molecule such as H₂ or acetate (Sieber et al. 2012). Syntrophy is a tightly coupled mutualistic interaction between H₂/formate-producing and H₂/formate using microorganisms where the pool size of intermediates that are exchanged between the partners must be kept very low for efficient cooperation among the microorganisms to occur. It occurs throughout the microbial community. The methanogens use the carrier molecule as electron donor for energy conservation, and the fermentative organism gains energy from the redox reaction that produces the electron carrier only if the methanogens oxidize the carrier molecule, keeping the carrier at a low concentration. Syntropic interactions enable methanogenesis when methanogenic substrates are limiting, and their establishment can also lead to increased CH₄ production rates (Kato et al. 2012). The global biogeochemical impact of syntropic interaction involving methanogenic Euryarchaeota is considerable as they enable the complete degradation of complex organic molecules to CO₂ and CH₄ in methanogenic habitat (Sieber et al. 2012). Acetotrophic methanogens can also produce H₂ and support the H₂-dependent dechlorination of xenobiotic compounds by dehalorespiring microorganisms (Heimann et al. 2006), further broadening the biogeochemical significance of the interactions involving methanogenic archaea.

7.6.1 Methane Oxidation in Soils

The oxidation of CH₄ is carried out by bacteria and archaea. The bacterial methanotrophs are gram-negative bacteria and generally only grow on CH₄ or methanol as a source of C and energy. Methanotrophic bacteria are often found at the anoxic/oxic interphase of various habitats such as geothermal reservoirs, landfills, soils, peat bogs, wetlands or aquatic environments and sediments where they consume the CH₄ arising from methanogenesis and are able to lower the CH₄ emissions (Conrad 1996). These methanotrophs can oxidize high concentrations of CH₄ (>100 ppm) and are sometimes referred to as low affinity methanotrophs. Other methanotrophs have the ability to oxidize CH₄ at atmospheric concentrations (~1.8 ppm) and are referred to as high affinity methanotrophs (Bender and Conrad 1992). Methanogenic bacteria belong to kingdom Eubacteria and two phyla—Proteobacteria and Verrucomcrobia. A total of 21 genera and at least 56 species have been identified (Nazaries et al. 2013).

The atmospheric CH₄ sink occurs in upland forest soils, and this fraction is attributed to high-affinity methanotrophs (Kolb 2009). In temperate soils, moisture ranging from 60 to 100% water-filled pore space (WFPS), decreases the CH₄ oxidation rates. This is attributed to limited O₂ availability and soil gas diffusivity, particularly because CH₄ transport in water is 10,000 times slower than in air (Castro et al. 1995). In contrast, when moisture is low (<12% v/v), CH₄ consumption is also inhibited by desiccation and increased osmotic stress (Jäckel et al. 2001). The effects of temperature on CH₄ oxidation are inconsistent, suggesting that methanotrophic bacteria can adapt to different temperatures. Overall temperatures <10 and >40 °C seem to decrease methanotrophic activity significantly in forest and landfill cover soils possibly as a result of inhibition of mesophilic methanotrophic activity (Semrau et al. 2010). Methanotrophs are also adaptive to wide range of pH—from pH 3.5 to pH 8, probably due to existence of acido-, neutro- and alkaliphiles. Other soil characteristics such as bulk density, porosity, WFPS are closely linked to water content and therefore influence O₂, availability and diffusion of CH₄ in soils. In addition, soil texture is also a strong regulator of CH₄ flux (Castro et al. 1995).

Unlike methanotrophic bacteria, methanotrophic archaea are strict anaerobes that gain energy by coupling the oxidation of CH₄ to the reduction of SO ⁻²₄ (Thauer 2011). Methanotrophic archaea belongs to Euryarchaeota and all are representatives of a single taxonomic class of Methanomicrobia along with various methanogenic archaea (Offre et al. 2013). They thrive in anoxic environment where both CH₄ and SO₄ ²⁻ are present, which occurs in marine sediments, where their habitat is restricted to the SO₄ ²⁻—CH₄ transition zone. CH₄ oxidizing archaea have also been detected in anoxic water columns, hydrothermal vents, soils, aquifers, and freshwater habitats. Their physiological settings differ widely which indicates that archaeal methanotrophs are physiologically diverse.

7.6.2 Methane Oxidation in Marine Environments

Anaerobic oxidation of CH₄ exerts strong control over ocean CH₄ emissions, which accounts for only 2% of the CH₄ released to the atmosphere. The process consumes an estimated 20–300 Tg CH₄ yr⁻¹, which is equivalent to 5–60% of the global annual CH₄ flux to the atmosphere (Smemo and Yavitt 2011). Global estimates of anaerobic oxidation of CH₄ in oceanic environments suggest that large fraction (>50%) of gross annual production in marine systems is consumed by anaerobic methanotrophs before CH₄ is even released to the ocean waters and about 90% of CH₄ produced in marine sediments is oxidized by microorganisms before it reaches the surface (Reeburgh 2007). Many aspects of the physiological mechanisms of anaerobic oxidation of CH₄ under marine environments remains mysterious and difficult to interpret, however.

The anaerobic CH₄ oxidation in the ocean has been mostly attributed to the CH₄-oxidizing activity of anaerobic methanotrophic archaea, but other microorganisms may also be involved (Beal et al. 2009), their identity and contribution to the global CH₄ has not been determined. The ability of anaerobic methanotrophs to respond to potential increase in oceanic CH₄ production, which could result from an accelerated melting of the large reservoir of CH₄ hydrates present in the seabed (Boswell and Collett 2011) is uncertain however. Coastal regions comprise small area of the Earth’s surface, but coastal wetlands are estimated to emit 40–60 Tg CH₄ yr⁻¹ to the atmosphere, accounting for 7–30% of the global annual fluxes of CH₄ (Dentener et al. 2003). Coastal wetlands regions also play significant role in cycling CH₄, however, few direct measurements of CH₄ consumption are available (Van der Nat et al. 1997; King et al. 1990), and particularly anaerobic oxidation of methane in coastal habitats is unknown. The projected rise in sea level may cause a dramatic shifts in biogeochemical functioning of coastal sediments through changes in salinity and availability of electron acceptors including sulfate and iron and cause increase in CH₄ fluxes from these regions (Weston et al. 2011).

7.6.3 Methane Oxidation in the Atmosphere

In addition, oxidation of CH₄ occurs in the atmosphere worldwide, but the sources of CH₄ are not uniformly distributed. For example, over the most recent 1000 years, it has been shown that the Greenland CH₄ exceeded that of Antarctica by 24–58 ± 10 ppbv (Etheridge et al. 1998), and over 10,000 years the difference between Greenland CH₄ concentrations and Antarctica has been estimated at 33 ± 7 ppbv. Therefore, varieties of processes at regional and global complexities of both C cycle and climate system may have contributed to the trends observed in ice core record of Holocene CO₂ and CH₄.

The CH₄ measurements from the Law Dome site in Antarctica show small variations that parallel the CO₂ trends over the most recent millennium including a decrease of 40 ppbv coinciding with the CO₂ decrease in 16th and 17th centuries (Etheridge et al. 1998). Correlations between decrease in CH₄ and lower temperatures reconstructed during 15th and 16th centuries suggests that climate change may have reduced CH₄ emissions by wetlands during this period. The maximum natural rates of change in atmospheric CO₂ and CH₄ during the most recent millennium were at least an order of magnitude smaller than current anthropogenic rates of change, and the natural variations were much less persistent than the increasing trends of the most recent two centuries.

Extrapolated trends in CH₄ measurements at these sites and at the Dryas 3 site in Greenland are likewise consistent with recent atmospheric measurements (Stauffer et al. 1985; Etheridge et al. 1992). Similarly, atmospheric CH₄ concentration decreased from early Holocene to about 5000 years and then increased gradually by about 100 ppbv between 5000 and 1750 (Fig. 7.3). Both Greenland and Antarctica ice cores indicate that atmospheric CH₄ concentrations varied much wider during Holocene Epoch, with its concentrations decreasing from near 700 ppbv at 10,000 years to concentrations >600 ppbv at 5000 years, followed by gradual return to near the 700 ppbv by 1000 years BP (Flückiger et al. 2002; Chappellaz et al. 1997; Blunier et al. 1995). Alternative hypothesis attributes atmospheric CH₄ increase between 8000 and 1000 years to expansion of boreal wetland source areas (Blunier et al. 1995; Chappellaz et al. 1997; Velichko et al. 1998). Others have attributed late Holocene atmospheric CH₄ increase partly to onset of early human cultivation of rice (Ruddiman 2007; Subak 1994; Chappellaz et al. 1997) and widespread domestication of ruminants after 5000 years (Ruddiman 2007; Fuller et al. 2011). A study by Kaplan et al. (2006) suggested that part of the late Holocene CH₄ increase could be explained by anthropogenic sources.

7.7 Methane Clathrate Hydrates

Gas hydrates are ice-like crystalline solids that form from mixtures of water and light natural gases such as CH₄, CO₂, ethane and butane. The CH₄ hydrates are nonstoichiometric solid structures composed of cages of water molecules surrounding CH₄ molecules in high pressure and low temperature conditions stabilized by van der Waal type interactions (Hester and Brewer 2009; Xu and Li 2015). The water molecules are arranged in a rigid framework of cages called clathrates, many of which are occupied and stabilized by a molecule of CH₄. The occurrence of CH₄ hydrates is controlled by an interaction of temperature, pressure, gas saturation, and water; and local chemical conditions combine these factors to make them stable (Kvenvolden 1993b; Demirbas 2010a). Most marine gas hydrate that have been analyzed are CH₄ hydrates. Such conditions can exist in ocean bottom sediments at water depths below 500 m (Kvenvolden 1993a) or in the Arctic, where hydrate forms beneath the permafrost (Demirbas 2010a). Hydrates occur in these two type of settings because the temperature-pressure conditions in these settings are within the hydrate stability field (Lerche and Bagirov 1998). Gas hydrates occur in pore spaces of sediments, and form cements, nodules, veins, or layers. Given a favorable combination of pressure and temperature, and availability of free CH₄ and water, gas hydrates can form and remain stable (Kvenvolden 1993b). The hydrate deposits themselves may be several hundred meters thick (Demirbas 2010a). In the oceans, gas hydrate is generally found at depths between 300 and 4000 m and at temperatures between 2.5 and 25 °C (Max and Dillon 1998). The existence of CH₄-rich natural hydrate in large quantities was recognized in 1960s (Makogon 1965), and currently over 90 sites have been directly or indirectly identified to contain natural gas hydrates.

The term CH₄ hydrate and gas hydrate are often used interchangeably and refers to these CH₄-water crystalline structure called clathrate. Clathrates are crystalline solids which look like ice and occur when water molecules form a cage-like structure around smaller guest molecules. The most common guest molecules are CH₄, ethane, propane, isobutene normal butane, N₂, CO₂ and H₂S, of which CH₄ occurs most abundantly in natural hydrates (Mahajan et al. 2007). Clathrates are stable under specific pressure-temperature conditions. Under appropriate pressure, they can exist at temperatures significantly above the freezing point of water. The maximum temperature at which a clathrate can exist depends on pressure and the gas composition. For example, CH₄ and water at 41 atmosphere pressure forms a hydrate at 278 K, whereas, the same pressure with 1% propane forms hydrate at 281 K (Demirbas 2010b). Hydrate stability can also be influenced by other factors such as salinity (Archer 2007). Hydrate formation is bound by temperature and pressure. Hydrate contains large amount of gas. For example, 1 m³ of hydrate disassociates at atmospheric temperature and pressure to form 164 m³ of CH₄ and 0.8 m³ of water (Kvenvolden 1993b). With pressurization, hydrates remain stable at temperatures up to 291 K. Because of the requirements of pressure and temperature, and also the requirement of relatively large amounts of OM for bacterial methanogenesis , clathrates are mainly restricted to two regions: (i) high latitudes, and (ii) along continental margins in the oceans (Kvenvolden 1993a, 1998).

Estimates of the size of the CH₄ reservoir sequestered as hydrates are enormous. Because these CH₄ hydrates are frequently close to the sediment–water interface and because they represent such a large reservoir, hydrate stability has been a major concern. There is extensive evidence of slides and craters on ocean shelves, presumably caused by destabilization and release of gas-phase CH₄ (Hovland and Judd 1988), but little information on the magnitudes of CH₄ release needed to form these features exist. While CH₄ is an important part of the natural C cycle, little is known about the role of CH₄ hydrates in the global budget (Cicerone and Oremland 1988). Specifically, much remains unknown about the processes leading to CH₄ generation, its transport, and transformations within sediments and its significance in global C budgets (Hovland and Judd 1988; Dickens 2003a; Torres et al. 2004). The CH₄ hydrates have a latent heat of fusion quite similar to that of water (Sloan 1998). Therefore, the decomposition of hydrates is a matter of heat flux and not a rise in temperature. Hydrate stability is not a matter of pressure–temperature conditions alone, as CH₄ concentrations substantially higher than saturation values are required for hydrate formation and stability (Sloan 1998).

7.7.1 Origin and Formation of Methane Hydrates

The CH₄ in gas hydrates is dominantly generated by bacterial anaerobic degradation of OM in low O₂ environments. In ocean, CH₄ could be produced by bacteria near the sea floor that are decomposing organic sediments or originate from oil deposits leak in the sea floor bottom through faults and cracks. Bacterial CH₄ formed in early diagenesis of OM and thermogenic gas leaking to the surface from a deep thermogenic gas accumulation can form a gas hydrate in the same continental shelf sediment. There are hydrates forms in both primary and secondary pore spaces and fractures in sediments as a diagenetic mineral (Sloan 1998). Sub-seabed CH₄ within the continental margin sediment is produced primarily by microbial or thermogenic processes. In the thermogienic process, thermal cracking of organically derived materials forms petroleum hydrocarbons, including CH₄, generally at a considerable depths (>2 km) in sedimentary basins where temperatures exceed 273 K, or thermal degradation of oil at even greater depths, and also by maturation of coal (Desa 2001). However, it is unknown how CH₄ is trapped in conditions favorable for CH₄ retention within the lattice.

7.7.2 Methane Hydrate Reservoir

CH₄ is an important component of the global C cycle, however, much remains unknown about the processes leading to CH₄ generation, transport, and transformation within the sediments, and its significance in global C budgets (Hovland and Judd 1988; Dickens 2003b; Torres et al. 2004). CH₄ comprises more than 99% of the hydrocarbon gases in most natural gas hydrate. Minor quantities of CO₂ and H₂S are often present, but because the main component is CH₄ the gas hydrate is called CH₄ hydrate (Sloan 1998; Kvenvolden and Lorenson 2001). The CH₄ hydrate recovered from the Gulf of Mexico show that CH₄ hydrates are often accompanied by significant amounts of ethane and propane and higher molecular weight hydrocarbon gases (Sloan 1998; Sassen and Macdonald 1994). The information on molecular composition of hydrocarbon gases in gas hydrate samples, coupled with isotopic composition (δ¹³C and deuterium) of CH₄ provide basis for interpreting the origin of gas. In most of CH₄ hydrate samples CH₄ has C isotopic composition lighter than −60‰, suggesting that the CH₄ is mainly microbial in origin. Microbial CH₄ in CH₄ hydrate may have either been generated in place from methanogenic processes taking place in shallow sediment, migrated short distances, or recycled during processes of hydrate formation and dissociation accompanying sedimentation (Paull et al. 1994). In contrast, CH₄ with C isotopic compositions heavier than −60‰, accompanied by significant amounts of higher molecular weight hydrocarbons, such as ethane, propane and larger hydrocarbon gases is of thermal origin. Most natural gas hydrate contains CH₄ that was generated by microbial processes. The CH₄ with thermal origin occurs in gas hydrates only where deeper deposits of thermogenic natural gas provide a source of CH₄ (Kvenvolden and Lorenson 2001). The estimates of CH₄ content of CH₄ hydrate are highly speculative, but they suggest that CH₄ quantities are very large (Kvenvolden 1999). The potential size of the ocean CH₄ hydrate reservoir estimated based on two models for CH₄ hydrate formation: (i) in situ bacterial production, and (ii) a fluid expulsion model suggested CH₄ hydrate reservoirs of 14,000 and 75,000 Pg C, respectively (Gornitz and Fung 1994). A consensus value of 10,000 Pg C for the amount of CH₄ hydrate has been suggested (Kvenvolden 1993a, b; Kvenvolden and Lorenson 2001; Archer et al. 2009; Burwicz et al. 2011). This is a large quantity, nearly 2000-fold larger than the atmospheric CH₄ inventory and ~30% of the ocean dissolved inorganic carbon reservoir. It is estimated as one of the largest C reservoir on Earth (Milkov 2004; Reeburgh 2007), containing about ten times more than the atmosphere. The need for energy is driving much of the current natural hydrate research and clathrate exploration programs (Collett 2002; Demirbas 2010b). Conversely, the vast quantities of hydrates in marine sediments pose a risk as geo-hazard, which have been implicated in past climate change event (Kennett 2003). This section focuses on its role on global C cycle and the possible impact on the global climate change. When warmed or depressurized, CH₄ dissociates into water and CH₄. Discharge of large amount of CH₄ into the atmosphere causes global warming. Direct transfer of even a small fraction of the ocean hydrate CH₄ to the atmosphere would have a catastrophic effects on climate, a process proposed as both an initial driver of positive feedback to climate change (Dickens et al. 1995; Archer 2007).

7.7.3 Methane Hydrate Decomposition Climate Change

The possible connection between climate and hydrates have been of interest to geoscientists. Natural hydrates may affect climate because when they are warmed or depressurized they decompose and dissociate into water and CH₄. Discharge of large quantities of CH₄ into the atmosphere will cause global warming. The rapid climate change in the past from cold glacial period to warmer interglacial periods has been indeed attributed to large release of CH₄ from gas hydrates. Some models have linked these observations with climate change caused by hydrates dissociation and vice versa (Blunier 2000; Kennett et al. 2000). It is suggested that glaciation and/or sea level fall will reduce hydrostatic pressure on sediments of shelf and slope at mid-latitudes, thereby destabilizing hydrates and consequently releasing CH₄ into the atmosphere. The CH₄ released will cause rapid warming, which will lead to dissociation of hydrates in the permafrost regions and shallow continental margins of high latitude regions, a positive feedback causing further warming. As a result, ice in the higher latitudes will melt and increase sea level. The rise in sea level will increase the hydrostatic pressure and the sediments shelf and slope at mid-latitudes initiating negative feedback that inhibit further release of CH₄ from hydrates. The geologic evidence, primarily from the C isotopic record of deep-sea cores, for CH₄ releases in the past exist (Dickens et al. 1995; Katz et al. 1999).

The CH₄ hydrates are located in shallow submarine geosphere, which is a finely balanced system in equilibrium with all its components, including sediment, pore water, fluid flows, pressure, temperature, overlying water, and hydrate. Removal of any one component of this equilibrium may destabilize the whole system leading to irreversible damage. The CH₄ Hydrates may be in steady state, but they cannot be stable because of differences in the chemical potential of CH₄ in the hydrate phase and the adjacent dissolved interstitial CH₄ (Dickens 2003a). Thus, CH₄ hydrates must be viewed as dynamic, with large and unknown fluxes to and from the ocean. The mechanisms involved in destabilizing CH₄ hydrate include deep-water warming (Hesselbo et al. 2000; Norris and Rohl 1999) and mass wasting of continental slopes (Vogt and Jung 2002). The destabilizing factors may either be natural, anthropogenic or combination of both. The CH₄ hydrates are susceptible to the influence of temperature (Glasby 2003). Due to enormous amount of CH₄ stored in CH₄ hydrates worldwide, their destabilization have the potential to affect global climate and geological environment at a catastrophic scale.

The rate and pathway by which CH₄ gas is released from hydrate on sea floor and may be transferred to the atmosphere is still a matter of debate. An understanding of the rate of contemporary hydrate-derived CH₄ additions to the ocean is essential to evaluating the role of hydrates in the global CH₄ budget. However, accurate measurements which could permit this estimate to be made is challenging because: (i) no means of discriminating between hydrate CH₄ and diagenetically produced CH₄, (ii) CH₄ released by dissociating hydrates is effectively oxidized by microbially mediated processes in anoxic sediments (Reeburgh 1980; Alperin and Reeburgh 1984), as well as in the adjacent oxic water column (Valentine et al. 2001), therefore, elevated CH₄ concentrations or “hot spots” are rarely observed, and (iii) hydrates collected by coring or dredging decompose rapidly, and laboratory studies are difficult if not impossible to conduct at sea.

7.7.4 How Can We Estimate the Rate of Methane Clathrate Decomposition?

Methane clathrate decomposition has been implicated in the Latest Paleocene Thermal Maximum (~55 Ma) by an extraordinary injection of isotopically light carbon into the carbon cycle (Dickens 2000, 2001) and in Quaternary interstadials as indicated by observations of isotopically light foraminifera in Santa Barbara Basin sediments (Kennett et al. 2000). Research on CH₄ clathrates has been advanced by using remotely operated vehicles (ROVs), which can be equipped not only to synthesize hydrates in situ under appropriate pressure–temperature conditions (Brewer et al. 1997, 1998), but to locate seeps for sampling and to place acoustic beacons for longer-term stability experiments as well. Direct measurements of the decomposition rates of pure CH₄ and CO₂ hydrates in a seafloor setting nominally within the gas hydrate pressure-temperature stability zone have been made recently (Rehder et al. 2004). Their rates represent end members because of the flow field and undersaturation. The water-column CH₄ oxidation rate measurements of (Valentine et al. 2001) were conducted adjacent to seeps in the Eel River Basin that were previously located with an ROV. These oxidation rate measurements were conducted with tracer additions of 3 H–CH₄, and showed the highest rates near the bottom. These instantaneous rate measurements are difficult to interpret in a dynamic advecting coastal environment, but the fractional turnover rates can be used to infer a CH₄ turnover time of ~1.5 years in the deep waters. The depth-integrated rates (5.2 mmol CH₄ m⁻² yr⁻¹) can be used to estimate total oxidation for the 25 km² vent field. The amount of CH₄ emitted is unknown, so it is not possible to estimate the oxidized fraction, but results from a closed system oxidation model applied to d¹³CH₄ measurements suggested that ~45% of the added CH₄ was oxidized. Grant and Whiticar (2002) concluded from observations of isotopically heavy d¹³CH₄ at Hydrate Ridge that substantial aerobic CH₄ oxidation occurred there. These isotopic enrichment measurements indicate cumulative CH₄ oxidation, but provide no rate information. Clearly, many more direct CH₄ oxidation rate measurements similar to those of (Valentine et al. 2001) are needed.

Recent work indicates that it may be possible to discriminate between hydrate–CH₄ and diagenetic–CH₄ with measurements of ¹⁴CH₄. If hydrates are formed by fluid expulsion, they will probably involve CH₄ from deep sources and will probably contain no radiocarbon; if they are formed by in situ bacterial production, the radiocarbon age will depend on the age of the substrate Winckler et al. 2002) reported radiocarbon measurements on CH₄ from massive CH₄ hydrates collected at Hydrate Ridge, Cascadia margin, and showed that they, as well as the unpublished results, are devoid of radiocarbon, indicating that there are no recent (¹⁴C-active) contributions to the hydrate reservoir. Measurements of noble gases in hydrates (Winckler et al. 2002) showed that light noble gases are not incorporated into the hydrate structure, but that significant quantities of the heavier noble gases—argon, krypton, and xenon—are present. The heavy noble gases are potential secondary tracers of hydrate decomposition.

7.8 Conclusions

The atmospheric concentration of CH₄ has increased dramatically over the last century and continues to increase. The atmospheric concentration increase is a result of change in both global natural and anthropogenic emission sources. Emission of CH₄ is mainly the result of environmental processes such as methanogenesis in wetlands and other anoxic environments such as rice fields, rumen of ruminant animals and landfills. Additional sources include geological sources such as fossil fuel production and CH₄ hydrates. The global CH₄ emissions based on inversion of atmospheric measurement of CH₄ from surface stations for 2000–2009 are 553 ± 27 Tg CH₄ yr⁻¹. The total loss of atmospheric CH₄ is 550 ± 36 Tg yr⁻¹ indicating imbalance of 3 Tg CH₄ yr⁻¹, the global total emissions based on bottom-up models and inventories is 678 ± 136 Tg CH₄ yr⁻¹. The total loss is 632 ± 40 Tg CH₄ yr⁻¹, with an imbalance of about 45 Tg CH₄ yr⁻¹ compared to annual atmospheric growth rate of 6 Tg CH₄ yr⁻¹, indicating that bottom-up models may overestimate the global CH₄ budget. The top-down inversion have constrained OH fields, but there is no constraint applicable for sum of emissions in the bottom-up approach. Budget analysis have confirmed the causal role of human activities in this increase, however, significant uncertainties remain in understanding the factors that affect emissions from diverse sources and how these will change over time as a result of global climate change. Uncertainties in how the sources and sinks of CH₄ will change in future limit the ability to develop meaningful climate change policies aimed at controlling CH₄ emissions. The renewed increase in 2007 after a period of stabilization between 1999 and 2006 has created a renewed the debate about the future levels of CH₄. The increase is consistent with higher emissions from wetlands and fossil fuel burning, although the relative contributions remains uncertain. This increase is considered the first sign of the important biogenic positive feedback to climate change—the response of the wetlands to increases in global temperatures projected to occur in this century. Given the potential importance of CH₄ emissions as a positive feedback to anthropogenic climate change and other perturbations, it is of paramount importance that research addresses the knowledge gaps in CH₄ sources, sinks, and biogeochemical processes responsible for changes in emission. More research is needed to fully understand the CH₄ growth rate and its future trajectory. Complete analysis requires increased coordination between programs measuring atmospheric abundances, isotopic ratios, and modeling effort that couple biosphere and ocean biogeochemistry with atmospheric processes.