In the following sections, we first review background information on gas hydrates, including their global distribution and the amount of sequestered CH 4 , along with an overview of the impact of climate change processes on gas hydrates. We then assess the Intergovernmental Panel on Climate Change's assumptions about the emission of hydrate‐derived CH 4 to the atmosphere, explore the key challenges associated with distinguishing hydrate‐derived CH 4 from emissions originating with non‐hydrate sources, and review the sinks that prevent most CH 4 released by gas hydrate dissociation from reaching the atmosphere. We briefly consider major pre‐Holocene climate episodes for which researchers have inferred large‐scale gas hydrate dissociation events and provide an in‐depth assessment of each physiographic province that hosts gas hydrates, discussing the impact of climate change processes on those deposits and the existing data on CH 4 emissions. We also review modeling efforts that have been used to assess the interaction of gas hydrates and global climate and offer recommendations about the key knowledge gaps to be addressed by researchers in future observational and numerical modeling studies. Finally, we briefly explore the potential climate impact of inadvertent CH 4 leakage during hypothetical production of gas from hydrate deposits.

This paper reviews the current state of knowledge on the interactions between methane hydrates and the global climate system and addresses misconceptions about posited runaway dissociation, the contemporary input of hydrate‐derived methane to the atmosphere, the potential for massive methane releases, and the distribution of gas hydrates in high‐latitude regions. While there have been numerous studies of marine gas hydrate reservoir changes in response to climate events that occurred millions of years ago (e.g., the Paleocene‐Eocene Thermal Maximum at ~55.5 Ma) [ Dickens et al ., 1997 ] or during the late Quaternary [ Kennett et al ., 2000 , 2003 ], less consideration has been given to assessing observational data and other evidence about climate‐hydrate interactions over the contemporary period [ Archer , 2007 ; Archer et al ., 2009 ; Nisbet , 1990a , 1990b ; Ruppel , 2011a ]. O'Connor et al . [ 2010 ] considered the fate of methane hydrates under different future climate change scenarios but did not include factors (sinks) that strongly mitigate the impact that hydrate‐derived CH 4 has on the ocean‐atmosphere system. James et al . [ 2016 ] recently explored the interaction between climate change and gas hydrates with full acknowledgement of sinks but with a focus on the Arctic Ocean.

Taken together, the dependence of gas hydrate stability on pressure‐temperature (P‐T) conditions, the relatively shallow depths of hydrate occurrence beneath the seafloor or in permafrost areas (i.e., relative to conventional natural gas), hydrate's tendency to concentrate gas, and the large amount of carbon trapped in global gas hydrates contribute to the perception that gas hydrate breakdown, termed “dissociation”, is a potential threat associated with global warming. In addition, large‐scale gas hydrate dissociation is sometimes portrayed not only as a consequence of warming but also as a potential synergistic driver for enhanced warming if the CH 4 released from gas hydrates reaches the atmosphere. These dual roles of gas hydrate dissociation—as both an effect and possible contributor to global warming—have led some to adopt a catastrophic perspective on the interaction of the climate system with the global gas hydrate reservoir [e.g., Bohannon , 2008 ; Krey et al ., 2009 ; MacDonald , 1990 ; Mascarelli , 2009 ; Whiteman et al ., 2013 ].

The estimated amount of carbon trapped in global gas hydrate deposits has shrunk over the past few decades, as documented by]. With other sources remaining the same, this comparison shows the relative proportion of mobile (exogenic) carbon in gas hydrates from (a) the 11,000 Gt C estimate [] made soon after the first dedicated drilling programs to (b) the 1800 Gt C adopted here, which is a value close to the estimates of] and] and within the range given by. []. (c) To place the methane emission rate from dissociating gas hydrates as assumed by several IPCC reports in perspective, the assumed emission of 6 Tg yrCHis plotted with other estimates of atmospheric methane sources from the bottom‐up attribution of. [].

The impact of (a) temperature and (b) pressure increases on the marine gas hydrate stability field and methane solubility in pore space assuming initial seafloor depth of 750 m, background geothermal gradient of 35°C km, sediment thermal diffusivity of 3 × 10, and homogeneous sediments. Large changes were imposed here for the sake of demonstration. In Figure 5 a, seafloor temperature increased by 2.5°C over 1000 years, and the BGHS shoals as the geotherm reequilibrates. Solubility decreases slightly beneath the BGHS, squeezing more dissolved gas out of solution into free gas, but increases measurably within the GHSZ, meaning that gas hydrate may dissolve. The original GHSZ encompasses the entire gray zone, with the GHSZ after 1000 yr is reduced to the darker gray zone. In Figure 5 b, a pressure change corresponding to a 100 m increase in sea level over 1000 years is imposed. Since these plots are relative to the seafloor, they show the relative change in the position of the stability curve (red) as pressure increases and the BGHS deepens. The geotherm (blue) remains unchanged. Solubility increases below the BGHS, meaning bubbles may dissolve at these depths. Above the BGHS, solubility barely changes. New hydrate will only form near the new BGHS if methane (e.g., from preexisting underlying free gas) is available in excess of the increased solubility curve. The original GHSZ is the light gray zone, while GHSZ after 1000 yr also includes the darker gray zone. Solubility calculations carried out using Matlab code provided by W. Waite (personal communication, 2016).

Gas hydrate stability constraints for nominal (a) permafrost and (b) deepwater marine settings after]. Note that gas hydrate stability calculations usually use hydrostatic pressure in shallow subseafloor sediments owing to their high, water‐filled porosity. BSR refers to the bottom‐simulating reflector that marks the boundary between free gas and overlying hydrate‐bearing sediments in some marine reservoirs. The purple zones correspond to gas hydrate with dissolved gas, while yellow regions may have coexisting dissolved gas and free gas (gas bubbles). (c) An expansion of the seafloor part of Figure 4 b, demonstrating the relationship among the gas hydrate stability zone (GHSZ) and the zone where gas hydrate actually occurs based on pressure‐temperature and methane solubility (activity) constraints. Blue solubility curves are read on the bottom axis and red temperature curves on the top axis. The green areas have dissolved gas as their only methane phase but are within the theoretical GHSZ. The white region has dissolved gas but lies beneath the base of the GHSZ. Figure modified from] after]. SRZ refers to the nominal sulfate reduction zone, where methane is consumed via anaerobic oxidation of methane (AOM).

Thickness of the theoretical gas hydrate stability zone (GHSZ) as calculated by. [] as the base map, with locations of known gas hydrate (recovered samples/photographs; blue circles) and inferred gas hydrate (based on well logs or geophysical markers like bottom‐simulating seismic reflectors at the base of gas hydrate stability; red circles) superposed. Gas hydrate distribution is typically heterogeneous within the stability zone, and gas hydrate only rarely occurs through the full thickness of the stability zone. The map also shows place names used in the paper. GOM refers to the Gulf of Mexico. Boxes on the Canadian Beaufort margin [.,], northwest U.S. Pacific margin [.,], northern U.S. Atlantic margin [.,], and West Spitsbergen margin [.,] show locations where upper continental slope methane seeps may be linked to gas hydrate dissociation processes.

Methane hydrate is stable over a range of intermediate‐pressure and low‐temperature conditions found close to the seafloor in the sediments of deepwater (greater than 300–600 m) continental slopes and also within and beneath permafrost at high northern latitudes. These relationships are shown in map form in Figure 3 and on pressure‐temperature stability diagrams in Figure 4 . These stability conditions and the global distribution of gas hydrate make it susceptible to the key perturbations associated with global warming, namely changes in sea level (pressure) and increases in ocean/air temperatures. Simple calculations demonstrating the impact of changes in pressure and temperature on the gas hydrate reservoir are shown in Figure 5 . Gas hydrates sequester large amounts of carbon, with estimates ranging from more than half [ Kvenvolden , 1988a ] to ~15% [ Boswell and Collett , 2011 ; Milkov , 2004 ] of Earth's total mobile carbon, which includes that in soils, land biota, fossil fuels, peat, and other reservoirs. The relative importance of gas hydrates within the global carbon reservoir based on older analyses [ Kvenvolden , 1988a ] and more modern perspectives [ Boswell and Collett , 2011 ] can be deduced from Figure 6 .

Gas hydrate (white material) formed beneath mussel‐coated carbonate rock on the seafloor of the Gulf of Mexico as photographed by NOAA's Deep Discoverer remotely operated vehicle in 2014. Image courtesy of the NOAA Ocean Exploration and Research Program. Seafloor gas hydrate is not volumetrically important within the global reservoir, but the existence of such gas hydrates illustrates that methane and other hydrate‐forming gases migrate across the sediment‐water interface and interact with ocean waters. Methane seeps, authigenic carbonates, and seafloor gas hydrates can be important habitats for chemosynthetic organisms reliant on methane or sulfide for their metabolic processes [e.g.,].

Among the large CH 4 carbon reservoirs that naturally interact with the ocean‐atmosphere system and thus global climate, gas hydrates (Figure 2 ) have special relevance. Gas hydrate is an ice‐like substance formed when water and low‐molecular‐weight gas (CO 2 , H 2 S, CH 4 , and higher‐order hydrocarbons) [ Sloan and Koh , 2008 ] combine in a clathrate structure. Methane is by far the predominant gas within natural gas hydrates that have been directly sampled, an observation that may partially reflect the abundance of CH 4 (as opposed to higher‐order thermogenic gases) at the shallow depths typically accessed by coring and drilling. Gas hydrate concentrates CH 4 within its cage‐like molecules, with 1 m 3 of gas hydrate sequestering a maximum of 180 m 3 of methane as measured at standard temperature and pressure (STP). Owing to the predominance of CH 4 as the guest molecule, gas hydrate is often referred to as methane ice, and the terms “gas hydrate” or “methane hydrate” and just “hydrate” will be used interchangeably in this paper.

The absolute concentration of CO 2 in the atmosphere (~400 ppm) is ~220 times more than the concentration of CH 4 (~1.83 ppm), yet the radiative forcing of anthropogenic CH 4 is ~25% that of anthropogenic CO 2 . CO 2 concentrations have increased less than 50% since the preindustrial age, while CH 4 concentrations have increased by ~150% (Figure 1 ). Although the rate of change in atmospheric CH 4 concentrations has been tempered by some periods of slower increases [ Dlugokencky et al ., 1994 , 2003 ], the rise in absolute concentrations since the middle of the twentieth century and the strong radiative warming associated with this gas justify the prominent role that CH 4 has been given in discussions of greenhouse warming.

As evidence mounts for sustained global warming during the last half of the 20th century and the start of the 21st century, there is increased awareness of the relative importance of methane (CH 4 ) to greenhouse warming. In the most recent assessment of the Intergovernmental Panel on Climate Change [ Intergovernmental Panel on Climate Change ( IPCC ), 2013 ], methane (CH 4 ) was deemed 84 times more potent than carbon dioxide (CO 2 ) as a greenhouse gas over a 20‐year timeframe and 25 times more potent over a century on a per unit mass basis. Recent years have seen greater scrutiny of global sources of CH 4 emissions and, in some places, new regulation of anthropogenic activities that enhance these emissions.

The fundamental impact of climate‐related changes on the gas hydrate reservoir is relatively straightforward, but several points bear mentioning: First, the gas hydrate reservoir tends to be viewed as relatively static during periods of climate stability or in certain settings (e.g., deep oceans). In fact, gas hydrates near the BGHS are nearly always undergoing dissociation due to normal sedimentation, small perturbations in pressure, and propagation of past temperature changes to depth [e.g., Dickens , 2001a ]. Second, the impact of past climate change events is often used to frame how the gas hydrate reservoir may respond to future climate change. As noted by Archer and Buffett [ 2005 ], the predicted changes associated with anthropogenically‐driven global warming may be far larger than those of many past climate events, affecting even the stability of deep ocean hydrates. Finally, just as snow piles do not instantaneously melt on a hot day, gas hydrate dissociation is also not instantaneous merely because pressure or temperature conditions in the sediments lie outside those required for hydrate stability. The endothermic heat of reaction (~439 J g −1 of methane hydrate) [ Gupta et al ., 2008 ] hinders rapid dissociation and makes the dissociation process self‐regulating [ Circone et al ., 2005 ]. Without the delivery of additional heat to the hydrate deposits, a scenario of runaway dissociation (dissociation that self‐perpetuates once initiated) is unlikely.

While not widespread on present‐day Earth, one population of gas hydrates does experience higher temperatures and decreased pressures during the same climate event [e.g., Nisbet , 1990b ]. Gas hydrates formed beneath continental ice sheets or grounded ice on continental margins may have constituted a significant part of the global gas hydrate inventory during some periods of Earth's history (section 5 ). The depresssurization effect associated with thawing ice sheets that may have reached thicknesses greater than 1 km greatly exceeds that related to the total range in sea level changes (~300 m) [ Miller et al ., 2005 ] since 50 Ma. Thawing ice sheets may be a primary driver for gas hydrate dissociation early in deglacial cycles, followed later by dissociation due to warming temperatures.

To set the stage for the rest of this paper, we briefly explore the impact of global climate change processes on gas hydrate deposits, with more in‐depth discussion left to section 6 . Gas hydrates are destabilized by increasing temperature or decreased pressure (Figure 5 ), conditions rarely associated with the same climate event. For example, global warming leads to higher average ocean and air temperatures but also increased sea level (pressure). For hydrate conditions close to the phase boundary, the morphology of the gas hydrate stability curve means that the destabilizing impact of slightly higher temperatures usually overwhelms the small stabilizing effect associated with increased sea level (1 m increase in sea level yields <0.01 MPa increase in pressure) but on different time scales. For elastic sediments, the pressure perturbation associated with rising sea level would be relatively instantaneous. In contrast, the impact of temperature changes at the tundra surface (permafrost hydrates) or seafloor (marine hydrates or those associated with subsea permafrost) on hydrates buried at depth may be delayed by hundreds or thousands of years, depending on the thickness and thermal diffusivity of the overlying sediments. This lag means that gas hydrate that remains stable over the scale of centuries in response to climate perturbations [e.g., Kretschmer et al ., 2015 ; Ruppel , 2011a ] may liberate significant gas on millennial scales.

While we primarily focus here on the amount of CH 4 sequestered in gas hydrates, it is critical to note that especially in the marine environment, hydrates are spatially associated with large amounts of CH 4 that occurs as free gas below the BGHS or sometimes within the GHSZ [e.g., Flemings et al ., 2003 ; Gorman et al ., 2002 ]. Hornbach et al . [ 2004 ] estimated that the amount of free gas that could potentially be released due to perturbations to the GHSZ in gas hydrate provinces could be as large as two‐thirds of the estimated gas in place in hydrate.

The global estimate preferred by Boswell and Collett [ 2011 ] based on an exhaustive review of other assessments and considering the lessons of many drilling programs is 3 × 10 15 m 3 of methane gas in place (calculated at STP) in global gas hydrates, corresponding to ~1500 gigatons (Gt or 10 15 g) of carbon or ~2.0 million Tg (10 12 g) CH 4 . This estimate is similar that of the lower end of the range determined by Milkov [ 2004 ], 9–20% higher than that of Archer et al . [ 2009 ], and more than 30 times smaller than the 1 × 10 17 m 3 gas in place estimated by Klauda and Sandler [ 2005 ] at about the same time. A global estimate that apportions gas in place in methane hydrate among different countries yields a value slightly higher (~1800 Gt C) [ Johnson , 2011 ] than the Boswell and Collett [ 2011 ] estimate. Another modern estimate made by Dickens [ 2011 ] gives an upper bound of 12,400 Gt C, ~8.3 times larger than that of Boswell and Collett [ 2011 ] and also larger than an independent calculation 500–2300 Gt C for global methane hydrates [ Piñero et al ., 2013 ]. This Dickens [ 2011 ] value is close to the 11,000 Gt C given by Kvenvolden [ 1988b ] two decades earlier, an estimate that was criticized by Laherrere [ 1999 ] as too high. Several independent assessments use various observations, models, and methodology to converge on values that are mostly less than 2000 Gt C sequestered in methane hydrates [e.g., the lower range of Milkov , 2004 ; Archer et al ., 2009 ; Boswell and Collett , 2011 ; Piñero et al ., 2013 ]. We here adopt a value of ~1800 Gt C (~2400 Gt CH 4 ) gas in place in methane hydrates for the global system, excluding Antarctica. To render the discussion more generic, some explanations are also cast in terms of the percentages of the global gas hydrate reservoir, not absolute estimates.

Besides the distribution of gas hydrates in sediments, the other important issue related to assessing climate‐hydrate interactions is establishing the amount of CH 4 sequestered in climate‐sensitive hydrate deposits, which requires a priori constraints on the global CH 4 in place in all gas hydrates. As documented by Boswell and Collett [ 2011 ], estimates have ranged over several orders of magnitude in the past four decades (Figure 6 ), but they have mostly fallen since the 1990s [ Klauda and Sandler , 2005 ; Milkov , 2004 ]. One reason for the high early estimates is that researchers typically assumed that all potential porosity within the GHSZ was filled with hydrate. After the advent of dedicated drilling programs in the mid‐1990s [ Boswell and Collett , 2011 ; Ruppel , 2011b ], samples of hydrate‐bearing sediments and quantitative borehole logs began to reveal that only a small fraction of formation porosity typically hosts gas hydrate [e.g., Trehu et al ., 2006 ], except in some high‐permeability, coarse‐grained reservoirs.

The remaining ~1% or more of global gas hydrates occurs in high northern latitude permafrost areas [ McIver , 1981 ; Ruppel , 2015 ], both onshore beneath tundra (e.g., Russia, Canada, and the U.S.) and on continental shelves on circum‐Arctic Ocean margins whose permafrost has been inundated by sea level rise since ~15 ka [e.g., Kvenvolden , 1993 ; Lachenbruch et al ., 1982 ; Rachold et al ., 2007 ]. The shallowest permafrost‐associated gas hydrates (PAGH) are predicted to lie a few hundred meters deep but still within the permafrost zone (Figure 4 a). For permafrost that is several hundreds of meters thick, gas hydrate should also be stable beneath the base of permafrost, depending on the prevailing geothermal gradient. Many PAGH formed by a process that can be described in the vernacular as “freezing in place” of gaseous CH 4 that has presumably migrated to shallower depths from underlying conventional gas reservoirs containing thermogenic gas [ Collett et al ., 2011 ; Judge and Majorowicz , 1992 ; Ruppel , 2015 ].

When considering the vertical distribution of gas hydrates in marine sediments, a critical distinction is made between the GHSZ (Figure 4 b) and the zone of actual gas hydrate occurrence [e.g., Xu and Ruppel , 1999 ; Zatsepina and Buffett , 1998 ], as shown in Figure 4 c. For gas hydrate to form from dissolved methane, as much of the gas hydrate in marine sediments likely does, requires not only appropriate P‐T conditions but also sufficient water and the presence of methane in excess of its solubility in surrounding pore waters [ Klauda and Sandler , 2005 ; Makogon , 1997 ; Xu and Ruppel , 1999 ; Zatsepina and Buffett , 1998 ]. Thus, based only on thermodynamics and physical chemistry, the zone where gas hydrate can potentially occur will have both a top and a bottom that are controlled by the local solubility of CH 4 in pore space and the availability of CH 4 at each location in the sedimentary section (Figure 4 c). Hydrate can form as deep as the base of the gas hydrate stability (BGHS) only when enough gas is available in excess of local solubility, which typically requires elevated flux or long periods of hydrate formation [ Nimblett and Ruppel , 2003 ; Xu and Ruppel , 1999 ]. In practice, the top of the gas hydrate zone in marine sediments is controlled not only by CH 4 flux but also by a biogeochemical condition, namely the presence of a sulfate reduction zone in which methane is consumed by microbial processes (section 4 ). Except in certain high‐flux cold seep environments or fractured formations, high saturations of gas hydrate are usually not present throughout the GHSZ. Even within homogeneous sediments with sufficient CH 4 , gas hydrate will rarely saturate pore space completely, and subtle heterogeneities (e.g., diatom‐rich layers) [ Kraemer et al ., 2000 ] cause preferential permeability that affects the final distribution of gas hydrates. In addition, clays, fine‐grained sediments, and higher‐salinity pore water inhibit gas hydrate formation [ Clennell et al ., 1999 ; Ruppel et al ., 2005 ], while high‐permeability coarse‐grained sediments like sand [ Boswell et al ., 2009b ; Dai et al ., 2012 ] and areas with rapid advection of fluids and gas [ Xu and Ruppel , 1999 ] tend to host higher saturations of hydrate [ Ginsburg et al ., 2000 ]. Overall, the amount of CH 4 sequestered in the GHSZ will always be less than the amount of available pore space.

About 99% of gas hydrates form in marine sediments [e.g., McIver , 1981 ] on continental slopes at water depths of greater than ~500 m in temperate latitudes and ~300 m at high latitudes, where bottom waters are colder. These depths mark the shallowest P‐T limit for the gas hydrate stability zone (GHSZ) on continental slopes, where the stability zone vanishes. Downslope, the GHSZ in the sedimentary section thickens as pressure increases, and it may eventually encompass the uppermost few hundred meters of sediments in waters deeper than 1000 m. For example, the GHSZ beneath the Blake Ridge sediment drift deposit (>2000 m water depth) reaches ~550 m thickness [ Paull et al ., 1996 ]. Owing to the concentration of organic carbon on continental margins, these locations are where most gas hydrates occur (Figure 3 ), and gas hydrates are largely absent beneath abyssal plains. The organic carbon is delivered to the sediment both by the rain of phytoplankton to the seafloor in highly productive continental margin waters and by export of terrestrial sediment from the continents. Remineralization of sedimentary organic carbon produces CO 2 , and most CH 4 formed in sediments by microbial processes is the result of reducing this CO 2 . Microbial CH 4 , instead of thermogenic CH 4 formed at higher temperatures via the same processes responsible for conventional natural gas, is the type most often found in recovered gas hydrates.

Among the most important issues related to assessing climate‐hydrates interactions are establishing where gas hydrates occur and estimating the size of the potential CH 4 source associated with them. These elements define the role of methane hydrates within the global carbon cycle. Here we review the general distribution of gas hydrates, leaving specifics for section 6 , where we assess the vulnerability of each population of gas hydrate to climate forcing processes. This overview does not include the Antarctic continent, for which only one assessment has been made for subglacial hydrates [ Wadham et al ., 2012 ], nor the minor populations of gas hydrate present at inland locations (e.g., Lake Baikal or Tibetan Plateau permafrost) [ Scholz et al ., 1993 ; Yang et al ., 2010 ].

Numerous studies have constrained the upward flux of CHthrough soils and sediments and to seawater and the atmosphere, but fingerprinting CHin the environment as having originated with the dissociation of gas hydrates faces two major challenges. First, a measurable tracer must exist that uniquely identifies CHas once having been encapsulated by hydrates. Second, the true environmental impacts of CHreleased from dissociating hydrates may not be limited to only the recently encapsulated CHsince emissions associated with dissociation of gas hydrate may also release free gas previously stored in the sediments [e.g.,.,]. Attempts to fingerprint direct and indirect emissions associated with dissociating CHhydrates have focused on three main approaches: age dating, chemical tracers, and stable isotopic fractionation.

Measurements of δD‐CH 4 have been used in paleoclimate investigations to distinguish marine (e.g., methane hydrate dissociation) from terrestrial methane sources. For example, Sowers [ 2006 ] used δD‐CH 4 from Greenland ice cores to search for isotopic signals that might indicate that methane from dissociating marine hydrates had reached the atmosphere during late Pleistocene to Holocene warm periods (section 5.4 ). While δD‐CH 4 is useful as a first‐order discriminant between marine and terrestrial sources of CH 4 , it cannot distinguish between CH 4 emitted from dissociating gas hydrates and that emitted from other marine sources.

While it is common practice to measure the stable carbon isotopes of CH 4 (δ 13 C‐CH 4 ), these have not been shown to fractionate during formation or dissociation of gas hydrates [ Lapham et al ., 2012 ]. Methane is formed by both biological and geological (thermogenic) processes, each displaying a relatively unique value for δ 13 C‐CH 4 . Since gas hydrate can incorporate either microbial (light δ 13 C) or thermogenic (heavy δ 13 C) CH 4 , the gas produced during hydrate dissociation lacks a unique signature with respect to other CH 4 in the environment. In addition, anaerobic and aerobic methane oxidation, which can both affect CH 4 released by gas hydrates (section 4 ), have been shown to aggressively fractionate δ 13 C‐CH 4 . This further complicates source discrimination after CH 4 is released from dissociating hydrates [ Alperin et al ., 1988 ; Whiticar , 1999 ]. Although δ 13 C‐CH 4 cannot uniquely fingerprint CH 4 released from dissociating gas hydrates, paleoclimate studies often interpret negative carbon isotopic excursions in carbonates or paleosols as indicating large‐scale dissociation of the gas hydrate reservoir. These dissociation events have been assumed to release methane with a δ 13 C‐CH 4 isotopic signature of −60‰ [e.g., Dickens et al ., 1997 ]. The interpretation of negative carbon isotopic excursions in terms of dissociation events is justified based on the large size and widespread distribution of the gas hydrate reservoir, as well as its relatively unique δ 13 C signature compared to other global carbon reservoirs. However, while gas hydrate methane is isotopically‐light compared to other global carbon reservoirs, no systematic global surveys have been conducted to constrain the global average value for δ 13 C‐CH 4 encapsulated by hydrates. It should also be noted that wetlands also produce large amounts of isotopically light δ 13 C‐CH 4 (< −60‰) and are frequently implicated in methane releases investigated in paleoclimate studies that rely on nonmarine records [e.g., Chappellaz et al ., 2013 ; Petrenko et al ., 2009 ; Raynaud et al ., 1998 ; Sowers , 2006 ].

Stable isotopic fractionation. Physical and chemical processes often modify the naturally‐occurring ratio of stable isotopes of various molecules. For example, the dissolution and oxidation of methane [ Alperin et al ., 1988 ; Harting et al ., 1976 ; Kessler et al ., 2006 ] systematically change the ratio of 13 CH 4 / 12 CH 4 as well as 12 CDH 3 / 12 CH 4 . Thus, it has been hypothesized that the formation and/or dissolution of gas hydrates may also modify the stable isotopes, producing an isotopic tracer. The formation of hydrates has been shown to fractionate the stable isotopes of water forming the hydrate cages, as noticed in the pore waters of ocean sediment cores [ Hesse and Harrison , 1981 ; Kvenvolden and Kastner , 1990 ; Matsumoto and Borowski , 2000 ]. However, if the goal is to trace CH 4 that is emitted from the seafloor to hydrate decomposition in the sediments, any changes in water isotopes in the seep gas stream will be overwhelmed by the overlying aqueous environment.

Another promising approach for fingerprinting released gas as possibly derived from recently dissociated CH 4 hydrate exploits noble gas concentrations. In CH 4 hydrates, noble gas molecules enter the lattice in order by molecular weight (MW), meaning that gas hydrate should contain relatively more helium (MW = 4.003 g/mol) than krypton (MW = 83.300 g/mol). Such an ordering of noble gases would not be characteristic of generic gas streams originating with shallowly produced gas, coal beds, and leaking deep reservoirs. Analyses of the noble gas characteristics of natural gas hydrate samples was attempted, but encountered air contamination and other problems, for Blake Ridge and Hydrate Ridge [ Dickens and Kennedy , 2000 ; Winckler et al ., 2002 ]. Experiments on synthetic CH 4 hydrates and one set of natural subseafloor gas hydrate samples that were subjected to controlled dissociation proved that that light gases are preferentially released early in the process [ Hunt et al ., 2013b ], probably because their atoms are too small to contribute to stabilizing the gas hydrate lattice [ Hunt et al ., 2013a ]. This approach must now be tested on a greater number of natural gas hydrate samples and on gas streams that may originate with gas hydrate dissociation before it can be rendered a true fingerprinting method for identifying CH 4 gas emissions from degrading natural hydrates. However, even if this fingerprinting technique proves to be a reliable marker of methane that was once encapsulated in hydrate, it still will not identify coexisting free gas that was coreleased to the ocean‐atmosphere system during emission events associated with gas hydrate dissociation.

Chemical tracers. If gas hydrate formation preferentially incorporates specific molecules from the source gas mixture, a unique chemical fingerprint may be retained in these structures. Upon dissociation, this chemical signature would be detectable and enable identification of gas streams originating with gas hydrate dissociation. For example, Dallimore et al . [ 2008 ] found a mixture of CH 4 and higher‐order hydrocarbons in permafrost gas hydrates of the Mackenzie River Delta, and this same mixture was detected seeping from the tundra. This observation could imply that gas hydrate dissociation directly feeds this seep, but an alternate possibility is that a deeper conventional reservoir supplies both the seep and the gas for the hydrate reservoir. More recent investigations of the widespread CH 4 seeps on the West Spitsbergen margin [ Westbrook et al ., 2009 ] have discovered a similar chemical composition in the seep gas and the gas supplying the underlying hydrates [ Berndt et al ., 2014 ].

The inclusion of hydrate dissociation as a possible source of atmospheric CH 4 in the IPCC reports is a rightful acknowledgement of the fact that the amount of CH 4 sequestered in this reservoir dwarfs that in some other parts of the Earth system. On the other hand, the IPCC reports cite no direct sources that constrain emissions of CH 4 to the atmosphere as a result of gas hydrate dissociation. Indeed, while gas hydrate deposits are likely dissociating and releasing CH 4 to sedimentary sections and the ocean on contemporary Earth, there remains no evidence that this hydrate‐derived CH 4 reaches the atmosphere or that the amounts that could potentially reach the atmosphere are significant enough to affect the overall CH 4 budget. In the following sections, we discuss some of the difficulties in discerning methane released from gas hydrates from other populations of methane in the ocean and atmosphere and also underscore the powerful role of sinks in mitigating the transfer to the atmosphere of methane released by dissociating gas hydrates.

The values quoted by the various IPCC reports have never been based on observational evidence for CH 4 emissions derived from gas hydrate dissociation since no such measurements exist. A few examples underscore this point: The clearly identified assumption of Cicerone and Oremland [ 1988 ] that 5 Tg yr −1 CH 4 reached the atmosphere from gas hydrate dissociation has set the stage for the subsequent quarter century. The third IPCC [ 2001 ] cites the Fung et al . [ 1991 ] forward modeling study, which merely assigned a value for the contribution to atmospheric CH 4 emissions from gas hydrate dissociation. This was also the case with Lelieveld et al . [ 1998 ], which assumed 10 Tg yr −1 CH 4 emissions from gas hydrate for one scenario, a number that was then adopted by the third IPCC [ 2001 ]. The Wuebbles and Hayhoe [ 2002 ] study cited by the fourth IPCC [ 2007 ] is sometimes considered the observationally based source for the now oft‐used 5 Tg yr −1 CH 4 estimate for atmospheric CH 4 flux from gas hydrates. That study in turn cites Judd [ 2000 ], which is a geologic methane seepage study that does not provide an independent estimate for emissions derived from gas hydrate dissociation. Cranston [ 1994 ], on which Judd [ 2000 ] relies for his hydrate‐related flux discussion, estimates the sum of global diffusive and ebullitive fluxes from marine sediments to the atmosphere to be ~1.3 Tg yr −1 to 13 Tg yr −1 CH 4 considering all sources, including shallow‐water seeps and deepwater gas hydrates. The Denman et al . [ 2007 ] study cited in the fifth IPCC [ 2013 ] is the climate coupling chapter from the fourth IPCC [ 2007 ] and not an independent source of information. The fifth IPCC [ 2013 ] also refers to Dickens [ 2003b ], which is a book review of Kennett et al . [ 2003 ] that did not provide an estimate for CH 4 flux to the atmosphere from dissociating gas hydrates, as Dickens [ 2003a ] also did not. Shakhova et al . [ 2010a ], also given as a source for the hydrate‐derived atmospheric CH 4 flux in the fifth IPCC [ 2013 ], did not attribute the 7.98 Tg yr −1 CH 4 flux that they calculated for the East Siberian Arctic shelf to gas hydrate degradation, rather considering a range of potential sources.

The first published estimates of CH 4 flux to the atmosphere due to gas hydrate dissociation ranged from 0.12 Gt yr −1 to 8 Gt yr −1 C [ Bell , 1983 ; Revelle , 1983 ] as described by Kvenvolden [ 1988b ]. Later, Kvenvolden and Rogers [ 2005 ] reported these same numerical values in units of Tg yr −1 C (1000 times smaller) and cited Kvenvolden [ 1991 ] for an estimate of 4 Tg yr −1 CH 4 (3 Tg yr −1 C) emitted to the atmosphere by gas hydrate dissociation. From the first IPCC [ 1990 ] through the most recent IPCC reports (with the exception of the second IPCC [ 1996 ]), CH 4 released by gas hydrate dissociation has appeared on the list of sources contributing to atmospheric CH 4 , usually at the level of ~5 Tg yr −1 (Table 1 ). The range in the first IPCC [ 1990 ] was 0–100 Tg yr −1 CH 4 emitted from gas hydrates, while that in the fifth IPCC [ 2013 ] is 2–9 Tg yr −1 . At the rate of 5 Tg yr −1 CH 4 emissions, 400 ky would hypothetically be required for wholesale transfer of the estimated methane in place in the global gas hydrate reservoir to the atmosphere based on the Boswell and Collett [ 2011 ] estimate. Given the depth of burial of some gas hydrate populations and the strong sinks that upward migrating CH 4 would encounter both in the sediments and the overlying ocean on contemporary Earth, it is likely that little of this CH 4 could reach the atmosphere. However, dissociating hydrates could be a major source of atmospheric CH 4 under certain catastrophic, but unlikely, circumstances. Rapid, large‐volume CH 4 emissions might bypass strong sediment [e.g., Knittel and Boetius , 2009 ; Martens and Klump , 1980 ] or water column [e.g., Elliott et al ., 2011 ] sinks, injecting more methane into the atmosphere. Ruppel [ 2011a ] calculated that instantaneous release of 1.8 Gt C from the gas hydrate reservoir (~0.1% of estimated gas in place) would temporarily increase atmospheric CH 4 concentrations by more than 60% if all of the gas reached the atmosphere.

The annual flux of CH 4 to the atmosphere from all sources is estimated at ~555 Tg yr −1 . As updated by Kirschke et al . [ 2013 ] and summarized in the report of the fifth IPCC [ 2013 ], the total top‐down global methane emissions estimate is ~130 Tg yr −1 smaller than the bottom‐up estimate. At present, the top‐down estimation methods do not have the capacity to attribute CH 4 to individual geologic sources like gas hydrates. Kirschke et al . [ 2013 ] note that both the flux of methane to the atmosphere and the global sink for methane are likely overestimated in the bottom‐up assessment. The commonly adopted assumption for bottom‐up assessment is that ~5 Tg yr −1 CH 4 is emitted to the atmosphere from gas hydrate dissociation (e.g., IPCC [ 2001 ] and subsequent IPCC panels). The only CH 4 source that routinely appears in bottom‐up studies and that has a smaller contribution than gas hydrate dissociation is wildfires [e.g., Kirschke et al ., 2013 ].

A bottom‐up assessment scales measurements of CH 4 emissions made at or near the Earth's surface up to a larger area. The contributions of the different emission sources are then summed in order to derive a global estimate of CH 4 flux to the atmosphere. For example, CH 4 emissions measured at a group of terrestrial seeps or at rice paddies in a specific location may be upscaled to estimate the total flux from all such features. Relatively small overestimates in the contributions from different CH 4 sources can become magnified in the upscaling process, leading to a global emissions estimate that cannot be reconciled with the known amount and rate of increase of atmospheric CH 4 . Another challenge is that measurements of CH 4 emissions on the Earth's surface are inherently biased. For example, scientists quantifying geologic methane emissions tend to measure the most active seep sites, and upscaling of the biased sample can lead to attribution of too much methane flux to a particular source.

A top‐down assessment uses as its starting point the total amount of methane in the atmosphere on a global or regional scale. Applying several methods, including sophisticated inversion techniques, the total CH 4 concentration is then attributed to various sources. By its very nature, the top‐down assessment accounts for all CH 4 present in the atmosphere at a given spatial scale; however, the attribution to sources is typically not very detailed since several sources are typically lumped together.

Even before early observational studies attempted to link atmospheric methane records to CH 4 emission processes that might include gas hydrate dissociation [e.g., Kvenvolden , 1993 ], researchers had attempted to include gas hydrate dynamics in their evaluation of global methane budgets [e.g., Cicerone and Oremland , 1988 ; Revelle , 1983 ]. The two approaches applied to evaluating the CH 4 sources in the atmosphere are termed top‐down and bottom‐up, which are described below and explored in detail by Nisbet and Weiss [ 2010 ] and Kirschke et al . [ 2013 ]. Both the fifth IPCC [ 2013 ] and Kirschke et al . [ 2013 ] provide up‐to‐date comparisons of top‐down and bottom‐up atmospheric CH 4 assessments. Each approach has different implications for the potential role of gas hydrate dissociation in releasing CH 4 to the atmosphere.

If CH 4 from dissociating gas hydrate reaches the atmosphere, it encounters a strong photolytic sink in the troposphere. An estimated 90% of CH 4 emitted to the atmosphere by all sources is removed by oxidation with the hydroxyl (OH) radical within about a decade [e.g., Cicerone and Oremland , 1988 ]. Numerous intermediate species, each with its own global warming potential, are generated during the process of tropospheric oxidation. However, the ultimate result of this CH 4 oxidation is CO 2 . A small proportion of CH 4 that reaches the atmosphere is transported to the stratosphere. Oxidation there results in the production of water vapor, itself a strong greenhouse species and a significant participant in the catalytic reaction for stratospheric ozone destruction. Forster et al . [ 2007 ] estimate that the combined products of CH 4 oxidation in the troposphere and stratosphere have about the same radiative forcing as the CH 4 itself, further underscoring the importance of CH 4 in the climate system. This rough equivalence between the radiative forcing associated with CH 4 and its oxidative products means that precisely determining the lifetime of atmospheric methane in the range of 8.9 ± 0.6 years [ Prinn et al ., 1995 ] to 12.4 yr [ IPCC , 2013 ] may not be critical for models that track the impact of CH 4 release at time scales of centuries to millennia.

A discussion of cold region terrestrial CH 4 sinks [ Jorgensen et al ., 2015 ; Walter Anthony et al ., 2014 ] is largely beyond the scope of this paper and is most relevant where CH 4 from dissociating deeply buried permafrost‐associated gas hydrates (estimated ~1% or more of the global inventory) [ Ruppel , 2015 ] migrates upward and encounters microbial soil sinks. Aerobic and anaerobic CH 4 oxidation in permafrost settings depends on a variety of factors [ Lee et al ., 2012 ; Preuss et al ., 2013 ]. The paucity of sulfate in terrestrial sediments means that oxidation pathways and rates are substantially different in tundra settings relative to deep marine sediments. Still, if CH 4 from dissociating gas hydrates were to pass through sediments in dissolved (bioavailable) form, some oxidation processes might be expected to mitigate the fraction of this CH 4 that reaches the surface [ Lau et al ., 2015 ]. Like marine sediments, tundra may also have low‐permeability traps that act as a kind of physical sink, retarding upward migration of CH 4 . A unique physical barrier in high‐latitude settings is pore‐filling ice, which could trap CH 4 released during subpermafrost gas hydrate dissociation.

Bubble stripping and the water column oxidation sink may prevent much of the CH 4 emitted at the seafloor from reaching the sea‐air interface, but these processes could still potentially have an important impact on ocean chemistry and habitats. As noted above, aggressive oxidation following massive CH 4 releases can deoxygenate waters and deplete certain chemical species while also increasing dissolved CO 2 that could enhance acidification [e.g., Archer et al ., 2009 ; Biastoch et al ., 2011 ; Dickens , 2001a ; Elliott et al ., 2011 ; Kessler et al ., 2011 ]. Outside of semienclosed anoxic basins, typical CH 4 concentrations range from 2 to 300 n M [ Heintz et al ., 2012 ; Mau et al ., 2013 ; Rona et al ., 2015 ], while the background concentrations of dissolved CO 2 and dissolved inorganic carbon (DIC) are approximately 30 μM and 2260 μM, respectively. The 4 to 6 orders of magnitude difference between dissolved methane concentration and DIC concentration suggests that even complete oxidation of CH 4 emitted at the seafloor from natural seeps will have an insignificant influence on inorganic carbon chemistry and, by extension, seawater pH, unless CH 4 and CH 4 ‐derived DIC can substantially accumulate in seawater.

Several aspects of the aerobic MOx sink remain poorly understood. For example, the chemical stoichiometry of this reaction, especially how it changes as the microbial population is aggressively growing in response to a seafloor injection of CH 4 , is unknown [ Chan et al ., 2016 ]. In addition, several investigations have documented depletions in dissolved oxygen, trace metals, and nutrients following the aerobic oxidation of CH 4 and oil in the deep waters of the Gulf of Mexico after the Deepwater Horizon incident [ Du and Kessler , 2012 ; Joung and Shiller , 2013 ; Schiller and Joung , 2012 ], leading to questions about whether these chemical species may be limiting factors for oxidation of seafloor CH 4 emissions. Also unknown is whether currents can disperse elevated concentrations of CH 4 to levels below the threshold to initiate MOx before microbial populations can increase in response to CH 4 injection or whether transport of dissolved CH 4 to the surface mixed layer and hence to the sea‐air interface can outcompete oxidation. Measurements of the natural radiocarbon content of CH 4 at a few locations have suggested that CH 4 in the surface mixed layer is not originating with seafloor emissions [ Kessler et al ., 2008 ] but rather with unique methane generation processes active in the near‐surface ocean [ Damm et al ., 2008 ; Karl et al ., 2008 ; Metcalf et al ., 2012 ; Repeta et al ., 2016 ].

Once dissolved in seawater, aerobic microbial oxidation (MOx) is a strong sink that can limit the flux of dissolved methane to the atmosphere [e.g., Mau et al ., 2007 ; Ward et al ., 1987 ]. MOx was traditionally thought to be a relatively slow process [e.g., Valentine et al ., 2001 ], but more recent studies show that oxidation rates can increase when seawater is perturbed with elevated concentrations of CH 4 [e.g., Crespo‐Medina et al ., 2014 ; de Angelis and Scranton , 1993 ; Du and Kessler , 2012 ; Heintz et al ., 2012 ; Kessler et al ., 2011 ; Mau et al ., 2013 ; Pack et al ., 2015 ]. Most of these studies inoculate seawater samples with radioisotope tracers to measure MOx rates in discrete parcels of water at the specific locations and times of sampling. The true strength of this sink can only be assessed by measuring the total integrated amount of MOx in the water column, either regionally [e.g., Du and Kessler , 2012 ; Leonte et al ., 2017 ] or globally. For example, assuming that the open ocean is in steady state with a dissolved methane concentration of 2 n M , this would equate to an open‐ocean methane burden of 43.2 Tg CH 4 [ Reeburgh , 2007 ]. If methane emissions to the water column are balanced by aerobic methane oxidation with first‐order MOx rate constants ranging from 0.001 to 0.2 d −1 [ de Angelis and Scranton , 1993 ; Mau et al ., 2013 ; Pack et al ., 2011 ; Valentine et al ., 2001 ], then both the rates of MOx and seafloor CH 4 emission to the water column probably lie in the range of 16 to 3200 Tg yr −1 . While we make no attempt to constrain the most likely value within this large range, combining even the low end‐member with the global sink associated with CH 4 oxidation in marine sediments (80–90% of 400 Tg yr −1 ) [ Reeburgh , 2007 ] suggests that oceanic sinks of methane are at least equal to, and potentially greater than, all atmospheric sinks of CH 4 (~540 Tg yr −1 from top‐down calculations in Kirschke et al . [ 2013 ]).

(a) From. []. The original caption is “Contour plot of the percentage of the initial methane mass reaching the atmosphere as a function of initial bubble diameter and release depth (methane reaching the surface is read at the point where the bubble diameter and release depth intersect on the plot). Environmental conditions were those from the Black Sea; however, these results are also valid as a first approximation for “normal” open ocean (e.g., Monterey Bay) or lake/reservoir conditions.” (b) Hydroacoustic image of seeping methane gas collected with a 38 kHz split‐beam sensor on the Baltimore Canyon seep field of the U.S. Atlantic margin in 2015, after. [].

As an example, for 50% of the CH 4 contained in the bubble at the seafloor to reach the atmosphere requires 14 and 20 mm diameter bubbles to be emitted at 50 and 100 m water depths, respectively, and even larger bubbles at greater water depths. (Figure 8 a). While quantification of bubble sizes at marine seeps is in its infancy, the bubble sizes measured to date are far smaller than would be necessary to ensure that CH 4 reaches the atmosphere [e.g., Römer et al ., 2012 ; Skarke et al ., 2014 ; Wang et al ., 2016 ]. Once dissolved in ocean waters, CH 4 can eventually be emitted to the atmosphere by gas exchange, which can be enhanced by certain conditions (e.g., high winds and storminess [ Shakhova et al ., 2014 ; Wanninkhof , 1992 ]). In deeper waters, CH 4 could remain in the oceans for centuries, depending on the nature of ocean circulation and the depth below the surface mixed layer at which the CH 4 is dissolved.

Once CH 4 is emitted at the seafloor, two major factors may prevent its reaching the atmosphere. First, despite methane's low solubility in seawater [ Wiesenburg and Guinasso , 1979 ], the concentrations of CH 4 in most ocean waters are still so low that the gas diffuses rapidly from rising CH 4 ‐filled bubbles following emission at the seafloor. During this bubble‐stripping process, CH 4 is replaced by oxygen and nitrogen [ McGinnis et al ., 2006 ; Vielstädte et al ., 2015 ]. Bubbles emitted deeper than the shallowest extent of gas hydrate stability in the water column may develop an armor of gas hydrate [ Chen et al ., 2014 , 2016 ; Graves et al ., 2015 ; Rehder et al ., 2002 , 2009 ; Sauter et al ., 2006 ; Topham , 1984 ; Wang et al ., 2016 ; Warzinski et al ., 2014 ; Zhang , 2003 ], but such armoring may or may not reduce the rate at which CH 4 leaves the rising bubbles [ Rehder et al ., 2002 ; Wang et al ., 2016 ]. Overall, most CH 4 emitted from the seafloor either above or below the top of the GHSZ and whether originating with gas hydrate dissociation or other processes will be dissolved relatively deep in the water column.

The physical characteristics of marine sediments also constitute a type of sink for CH 4 released from dissociating gas hydrates. These physical sinks do not transform CH 4 in the way that AOM does, but they can prevent it from interacting with the ocean‐atmosphere system for thousands of years. The most important physical factors are those that impede the migration of methane through the sedimentary section. Examples include low‐permeability sediments, structural traps, and hydrate‐ and/or gas‐saturated sediments that impede fluid advection [e.g., Bhatnagar et al ., 2007 ; Chatterjee et al ., 2014 ; Davies and Clarke , 2010 ; Davies et al ., 2014 ; Garg et al ., 2008 ; Liu and Flemings , 2007 ; Nimblett and Ruppel , 2003 ].

AOM is also closely coupled to carbonate precipitation at shallow depths beneath the seafloor since AOM produces bicarbonate, which increases alkalinity. The resulting authigenic carbonates, which are often exhumed at the seafloor at methane seeps, effectively remove carbon from the mobile carbon cycle. The carbonates also carry a record of CH 4 emissions in both their absolute age [e.g., Bayon et al ., 2009 ; Crémière et al ., 2016 ; Liebetrau et al ., 2010 ; Teichert et al ., 2003 ; Watanabe et al ., 2008 ] and in their changing carbon isotopic characteristics [e.g., Aloisi et al ., 2000 ]. In some places, the isotopic characteristics of these carbonates have been interpreted as reflecting local destabilization of gas hydrate [ Bohrmann et al ., 1998 ].

In the shallow marine sedimentary section, the strongest biochemical sink is the anaerobic oxidation of methane (AOM) [ Barnes and Goldberg , 1976 ; Martens and Berner , 1977 ; Reeburgh , 1976 ], which is carried out by a consortium of microbes [ Knittel and Boetius , 2009 ] and is strongly coupled to sulfate reduction [e.g., Malinverno and Pohlman , 2011 ], particularly in diffusion‐dominated provinces lacking additional hydrocarbon sources [ Joye et al ., 2004 ]. The sulfate reduction zone (SRZ) occupies the centimeters to meters just below the seafloor, with a thicker SRZ corresponding to areas of lower upward methane flux [ Borowski et al ., 1997 ]. The microbial consortium that carries out AOM [ Boetius et al ., 2000 ] has been termed a biofilter that prevents upwardly‐migrating CH 4 from reaching the seafloor, where it could be emitted to the ocean. A summary by Reeburgh [ 2007 ] concluded that up to 80% to 90% of the estimated 400 Tg yr −1 CH 4 that reaches the SRZ via upward migration through the sediments is consumed by AOM [ Hinrichs and Boetius , 2003 ]. At sites with vigorous seepage, AOM has sometimes been found to be highly efficient [ Joye et al ., 2004 ], while only about 20% of the CH 4 is consumed by AOM at other locations [ Boetius and Wenzhofer , 2013 ]. The reduced efficiency of AOM in some higher flux settings leads to the possibility that rapidly ascending gas in the form of bubbles may bypass the sediment biofilter [e.g., Martens and Klump , 1980 ] and be injected into the overlying ocean without major alteration by the AOM sediment sink.

Even when CH 4 is released from gas hydrate during dissociation, physical, chemical, and biological sinks in the sediments and ocean waters mitigate the amount that reaches the atmosphere (Figure 7 ). These sinks are so strong that there are likely very few locations on present‐day Earth where gas hydrate dissociation could release CH 4 that reaches the atmosphere in any significant quantities.

An evaluation of the clathrate gun hypothesis, although during a more recent time period, can be made using terrestrial ice core records. Sowers [ 2006 ] used Greenland ice cores to search for isotopic signals that might indicate that CH 4 from dissociating gas hydrates had reached the atmosphere during key late Pleistocene to Holocene warm periods. Given that microbial CH 4 released from wetlands and from dissociating gas hydrates cannot be distinguished on the basis of carbon isotopes alone (section 3 ), Sowers [ 2006 ] instead relied on hydrogen isotopic signatures (δD) that should provide a clear marker for deep ocean CH 4 hydrate dissociation (heavy deuterium) versus CH 4 expulsion from wetlands (light deuterium). The correlation between warmer climates and increases in atmospheric methane is well‐established, but the hydrogen isotope analyses by Sowers [ 2006 ] and some associated numerical modeling demonstrated that the atmospheric CH 4 did not originate with deep marine gas hydrate dissociation. A later ice core analysis by Petrenko et al . [ 2009 ] focused on the rise in atmospheric methane at ~11.6 ka warming event (Younger Dryas termination), explaining the increased methane in terms of wetland processes. Melton et al . [ 2012 ] used carbon isotopic evidence from a Greenland ice core to confirm the low likelihood that methane hydrate dissociation was a key source for atmospheric methane emissions at this time. More recently, Chappellaz et al . [ 2013 ] have used analytical methods with subdecadal resolution to study CH 4 in a Greenland ice core, revealing atmospheric CH 4 signals that correlate with local temperature. This implies that a marine gas hydrate dissociation source is not necessary to explain this high‐resolution ice core CH 4 record.

A secondary aspect of the clathrate gun hypothesis is the conjecture that warming ocean waters above sediments hosting gas hydrates leads to an increased incidence of submarine slope failures. Some researchers posit that large‐scale submarine slope failures or erosional episodes are an efficient means to rapidly dissociate gas hydrate and release large amounts of gas that can reach the atmosphere without being lost to sediment or water column sinks [ Bangs et al ., 2010 ; Nixon and Grozic , 2006 ; Paull et al ., 2002 ]. Through the careful analysis of large‐scale submarine slides in the late Quaternary, Maslin et al . [ 2004 ] showed that these slope failures most often originated during Heinrich events, not D/O events. Heinrich events are associated with massive rafting of icebergs from glacial ice sheets, leading to the addition of cool, dense waters to the oceans, not warming of intermediate waters.

Paleoceanographic records [e.g., Kennett and Ingram , 1995 ] and biomarker evidence [ Hinrichs et al ., 2003 ] have also been used to infer repeated dissociation of marine gas hydrates during late Quaternary climate events, including during late Pleistocene glaciations [ de Garidel‐Thoron et al ., 2004 ]. On the basis of a high‐resolution carbon isotopic record obtained in Santa Barbara Basin, Kennett et al . [ 2000 , 2003 ] advanced the clathrate gun hypothesis for the late Quaternary. This idea postulates that interstadials, which are temporary warming episodes within glacial periods (often called Dansgaard‐Oeschger or D/O events), are associated with warming of intermediate ocean waters and dissociation of continental slope gas hydrates that released CH 4 to the overlying waters and atmosphere. More recent studies use δD‐CH 4 (section 3.2 ) to show that changing atmospheric CH 4 concentrations at and between some D/O events cannot be linked to dissociation of the marine gas hydrate reservoir and more likely represent the influence of wetlands methane [ Bock et al ., 2010 ].

To explain observations during the PETM, researchers must rely on some form of an isotope mass balance in order to infer the source of carbon causing the isotopic excursion. Hypotheses as disparate as comet impact [ Kent et al ., 2003 ], permafrost thaw [ DeConto et al ., 2012 ], biomass burning [ Kurtz et al ., 2003 ], and volcanic heating that causes contact metamorphism of organic carbon [ Svensen et al ., 2004 ] have been advanced to explain the PETM. A few recent studies give credence to some of these mechanisms and increase the estimate of the amount of carbon released, as is necessary to satisfy the isotope mass balance [e.g., Wright and Schaller , 2013 ]. The methane hydrate dissociation hypothesis continues to garner the most attention, although the triggering mechanism for the dissociation remains uncertain. In addition, the amount of CH 4 release required to produce the observed carbon isotopic anomaly is unlikely to have caused sustained warming of the atmosphere [ Higgins and Schrag , 2006 ; Zachos et al ., 2003 ]. Another quandary is that terrestrial records reveal a consistently larger isotopic excursion [ Bowen et al ., 2001 ] than deep ocean sediment cores, which may have poorer preservation of the event [ McCarren et al ., 2008 ]. If the global carbon isotopic anomaly is related to a marine source, then the marked change in terrestrial carbon isotopes during the PETM implies that the water column oxidation sink did not prevent emissions of methane carbon to the atmosphere [e.g., Zhang , 2003 ]. Interestingly, a recent study constrained the release rate of carbon during the PETM to <1.1 Pg yr −1 [ Zeebe et al ., 2016 ]. Since the current rate of CH 4 release from the seafloor ranges from 0.016 to 3.2 Pg yr −1 (section 4.2 ; 1 Pg = 1000 Tg), the modern seafloor CH 4 emission regime is potentially comparable to that during the PETM. Recent modeling has also shown that multiple carbon (CH 4 and/or CO 2 ) emission events and possibly multiple sources may be required to explain observational data [ Carozza et al ., 2011 ] for the PETM.

The most widely studied hyperthermal event was first recognized by Kennett and Stott [ 1991 ] and is termed the Paleocene‐Eocene Thermal Maximum (PETM) or the Late Paleocene Thermal Maximum (LPTM) in older literature. This abrupt, multiphase extreme warming event [e.g., Bowen et al ., 2015 ] produced ~5°C increase in global temperature [ Zachos et al ., 2007 ] and even more heating in deep ocean basins [ Tripati and Elderfield , 2005 ] and also coincided with a 3.5 to 5‰ negative excursion in marine δ 13 C [ McCarren et al ., 2008 ]. The carbon isotopic data are supplemented by oxygen isotopic analyses of benthic forams, which provide a temperature proxy for the deep ocean [ Zachos et al ., 2001 , 2010 ]. Dickens et al . [ 1995 , 1997 ] hypothesized that large‐scale dissociation of marine gas hydrates released 1100 to 2100 Gt C of isotopically‐light carbon to the ocean‐atmosphere system during the PETM.

At the end of the Early Jurassic, three distinct negative carbon isotopic shifts during the Toarcian (~183 Ma) collectively produced an excursion of ~5–7‰ in δ 13 C [ Kemp et al ., 2005 ]. This carbon isotope signature is larger than that associated with the end of the snowball Earth episodes in the Neoproterozoic and also with the one that characterizes the well‐studied Paleocene‐Eocene Thermal Maximum (section 5.3 ). Astronomical changes (Milankovitch cycles) superposed on a longer‐term warming event are postulated to have repeatedly triggered the release of isotopically‐light carbon [ Kemp et al ., 2005 ] during the Toarcian, and large‐scale methane hydrate dissociation [ Hesselbo et al ., 2000 ] is the primary hypothesis for injection of ~5000 Gt C to the ocean‐atmosphere system [ Beerling et al ., 2002 ]. Several major ocean geochemical changes coincided with this inferred dissociation event [e.g., Jenkyns et al ., 2001 ; Schouten et al ., 2000 ], including ocean anoxia that would have affected the strength of the water column methane sink. Whereas dissociation events in the Neoproterozoic were precursors to expanding biodiversity, at least two of the events in the Toarcian were associated with widespread extinctions [e.g., Harries and Little , 1999 ; Kemp et al ., 2005 ].

The Neoproterozoic was a period of at least two extreme glaciations (snowball Earth) [ Hoffman et al ., 1998 ] that brought ice to low latitudes [ Sohl et al ., 1999 ]. Overlying the siliciclastic glacial sediments deposited in conjunction with these global cooling events are authigenic cap carbonates that occur in thin (up to 5 m thick) layers. The cap carbonates have textures and features similar to those produced by contemporary CH 4 cold seep processes and a strongly negative (up to 5‰) δ 13 C excursion and related δ 34 S anomaly, as would be expected if a large reservoir of light δ 13 C methane had been released and oxidized during the formation of the carbonates [e.g., Jiang et al ., 2003 ; Kennedy et al ., 2001 ]. The major hypothesis for these observations is widespread degassing of terrestrial permafrost‐associated gas hydrates [ Kennedy et al ., 2001 ], subsea permafrost hydrates [ Kennedy et al ., 2008 ], and/or most of the marine gas hydrate reservoir [ Jiang et al ., 2003 ] during ice sheet retreat and other deglacial events. Kennedy et al . [ 2008 ] estimate that ~3000 Gt C may have been released from nonmarine gas hydrates during warming following the Marinoan (~635 Ma) deglaciation and postulate that the deglaciation event set the stage for biogeochemical changes that were necessary for the explosion of life in the Cambrian. Bjerrum and Canfield [ 2011 ] focus on a younger climate event (~551 Ma) in the Neoproterozoic and conclude that large‐scale dissociation of marine gas hydrates beneath an anoxic ocean led to CH 4 being injected into an oxygen‐poor atmosphere during a short‐lived warming episode.

Hypotheses suggesting positive feedback between modern climate change and the release of methane from dissociating gas hydrates are based on the susceptibility of gas hydrates to pressure and temperature changes, the large amount of CH 4 sequestered in gas hydrates, and paleoclimate studies that have implicated such a feedback during certain periods of Earth's history. In this section, we explore some of the key paleoclimate evidence for the interaction of gas hydrates and the global climate system during past climate events. Throughout the discussion, it should be recognized that (1) paleoclimate interpretations that connect warming events and hydrate destabilization remain controversial; (2) the inferred release rate of CH 4 from hydrates in some paleoclimate interpretations is a small fraction of the current anthropogenic CO 2 release rate [ Archer and Buffett , 2005 ; Zeebe et al ., 2016 ]; and (3) globally significant releases of methane from hydrates, either in the past or future, generally require millennia or at least centuries.

6 Climate Susceptibility of Gas Hydrates by Physiographic Province

Ruppel [2011a] briefly explored the potential for dissociation of contemporary gas hydrate deposits over the next several centuries by considering the distinct physiographic settings where gas hydrate occurs and the susceptibility of each hydrate population to global warming processes. Figure 9 is an updated version of the hypothetical cross section from Ruppel [2011a], showing CH 4 sources, gas hydrate distributions, and possible leakage points. The following sections examine climate‐hydrates interactions in each setting, as summarized in Table 2.

Figure 9 Open in figure viewer PowerPoint Ruppel [ 2011a Ruppel [ 2011a Schematic of methane hydrate dynamics and methane distributions in different physiographic provinces. This diagram is updated from] with the addition of subglacial hydrates and methane accumulation under ice. The most climate‐susceptible hydrates are associated with (1) thawing subsea permafrost beneath Arctic Ocean shelves that were unglaciated at the Last Glacial Maximum (LGM) and (2) dissociating gas hydrates on upper continental slopes, respectively corresponding to Sectors 2 and 3 of].

Table 2. Assessment of Climate‐Hydrate Synergies in Characteristic Locations for Gas Hydrates Geographic Setting Estimated Methane in Place in Hydrates (%)a Predominant Type of Methane Nominal Shallowest Depth Beneath Land Surface/Seafloorb Susceptibility to Climate Change Pre‐atmospheric Sinks for Liberated Methane Onshore permafrost‐associated gas hydrates 20 Gt Cc (1.11%) Thermogenic ~200 m (for pure methane); as shallow as 100 m for mixed gases Intermediate: substantial climate warming and permafrost thaw expected at high latitudes, but hydrates are deeply buried Sediment permeability, ice‐blocked migration pathways, anaerobic/aerobic processes in sediments Subsea permafrost‐associated gas hydrates (PAGH) on arctic continental shelves Thermogenic ~200 m Intermediate: hydrates are deeply buried but will continue to dissociate as warming from inundation propagates to depth on Arctic Ocean shelves Sediment permeability and some ice‐blocked migration paths, AOM in near‐seafloor sediments, aerobic oxidation/dissolution of methane in the water column Subglacial Unknown at LGM; contemporary Antarctica: ~80–400 Gt C d, e Microbial and thermogenic in Antarctica; could have been primarily shallow microbial at LGM in other areas Very shallow High at LGM due to widespread thawing of ice sheets on land and grounded ice sheets on shelves; low in Antarctica over next centuries For onshore settings, few sinks except anaerobic/aerobic processes in sediments; for inundated settings, see subsea PAGH Upper continental slopes 63 Gt Cf (~3.5%) Microbial except in thermogenic basins (e.g., Gulf of Mexico) Base of sulfate reduction zone (centimeters to tens of meters, depending on methane flux rate) High, owing to warming of intermediate ocean waters and very small magnitude of offsetting stabilization effect associated with rising sea level AOM in sediments, sediment permeability, aerobic oxidation/dissolution of methane in water column Deep marine ~1717 Gt C (95.3%) Microbial except in thermogenic basins Base of sulfate reduction zone ncentimeters to tens of meters, depending on methane flux rate) Low due to long‐term stability of deep ocean temperatures and the fact that the hydrates are generally far inside the stability zone AOM in sediments, sediment permeability, aerobic oxidation/dissolution of methane in water column

6.1 Gas Hydrates Associated With Terrestrial Permafrost For high‐latitude regions, discussions of climate‐hydrate interactions often focus on hydrates associated with terrestrial permafrost, not marine settings. As for the global assessment of methane trapped in gas hydrate (section 2), we here consider only permafrost at high northern latitudes given the paucity of data on Antarctic gas hydrates [Wadham et al., 2012] and the relatively minor amount of permafrost in settings like the Tibetan Plateau [Yang et al., 2010]. Warming climate, which is more pronounced at polar latitudes than at more temperate locations (Arctic amplification), is already causing significant thawing of terrestrial permafrost [e.g., Oberman, 2008; Romanovsky et al., 2010] that sometimes cap gas hydrate deposits. Such thawing in turn leads to the liberation of carbon long bound in ice‐bearing strata [Schuur et al., 2009] and the consequent release of CO 2 and CH 4 from the shallow sediments as a result of microbial processes. In some areas arctic tundra is net sink for CH 4 [e.g., Jorgensen et al., 2015], and thermokarst lakes may also have become a net sink for carbon during the Holocene [Walter Anthony et al., 2014]. Where emissions dominate, CO 2 [Schadel et al., 2016] and CH 4 [McCalley et al., 2014] have alternately been identified as the major greenhouse gas released from thawing permafrost. Local geology plays a critical role in CH 4 emissions, which are enhanced near the boundaries of continuous permafrost and where faults intersect thawing permafrost [Walter Anthony et al., 2012]. The primary factors mitigating the potential impact of climate change on PAGH are their depth of occurrence and their limited extent relative to better studied marine gas hydrates [e.g., Ruppel, 2015]. In unglaciated areas, the shallowest PAGH should be found at depths of 200 m or more for pure CH 4 as the guest gas and possibly as shallow as 100 m for some mixtures of CH 4 and higher‐order thermogenic gases. The GHSZ can extend hundreds of meters beneath the base of permafrost (Figures 4 and 9). With sustained climate warming, permafrost thaws and CH 4 hydrate dissociates at both the top and bottom of its stability fields [e.g., Ruppel, 2011a]. As in marine sediments, CH 4 released from gas hydrate dissociation at depth has to overcome numerous physical and chemical sinks to reach the tundra surface, and ice‐bearing permafrost can be an effective permeability cap for upward migration of CH 4 liberated during dissociation. While conditions are suitable for gas hydrate stability in areas with thick permafrost, whether hydrates actually occur there is a more complicated issue that directly bears on the amount of methane that could be released from these deposits. As outlined by Ruppel [2015], gas hydrates are not as ubiquitous in permafrost areas as deepwater marine gas hydrates are in continental margin sediments. PAGH at high northern latitudes probably formed as gas bubbles froze in place during profound cooling of the land surface in unglaciated areas [Majorowicz et al., 2008], with the most recent episode being during the late Pleistocene [Collett et al., 2011]. The emphasis on unglaciated regions is related to their having experienced very low sustained temperatures at the land surface (e.g., −10 to −30°C), which can lead to the development of thick, continuous permafrost and thus a thick GHSZ. The origin of PAGH deposits as bubbles implies that migration pathways may link conventional gas reservoirs to the loci for hydrate formation. Thus, spatial coincidence with and/or geologic links to conventional reservoirs and the existence of a structural or permeability trap for the gas may be a requirement for the formation of PAGH [Ruppel, 2015]. In addition, contemporary PAGH settings are characterized by variegated clastic sedimentary sequences (e.g., intercalated silts, sands, and clays), and the high‐permeability, coarse‐grained layers critical to formation of high‐saturation gas hydrates [e.g., Boswell et al., 2009a; Winters et al., 2011] may be limited in thickness and regional distribution or occur in a part of the section that was inaccessible to a source of gas [Ruppel, 2015]. To date, numerous direct measurements have been made to quantify bulk CH 4 emissions from tundra, high‐latitude wetlands, and thermokarst lakes [e.g., Christensen, 1993; Walter et al., 2006; Whalen and Reeburgh, 1988; Wille et al., 2008; Zona et al., 2016], but attributing a fraction of this CH 4 stream to gas hydrate dissociation is not currently possible. Furthermore, the observed increase in atmospheric CH 4 concentrations since about 2007 [Dlugokencky et al., 2009] cannot be attributed to arctic emissions, which are expected to continue rising as global warming leads to enhanced methane production and/or release from several sources [World Meteorological Organization, 2013]. Even under a possible future scenario of rising arctic CH 4 emissions, which are expected to lag warming, discerning the component related to gas hydrate dissociation may always remain more challenging at high northern latitudes due to the number of methane sources in these settings and their overlapping depths of origin (Figure 10). Figure 10 Open in figure viewer PowerPoint Ruppel, 2015 Lecher et al., 2016 Bussmann, 2013 Walter Anthony et al., 2014 4 emissions [Walter et al., 2006 2 versus CH 4 emissions from tundra due to microbial processes in shallow sediments. Not shown is methanogenesis in near‐surface ocean waters [e.g., Damm et al., 2008 Terrestrial and continental shelf stores of methane and potential emissions to the ocean and atmosphere at high northern latitudes. The primary onshore and offshore methane sources to the atmosphere include wetlands, in situ methane generation in organic‐rich sediments (e.g., thermokarst lakes, and deltaic sediments), methane generation from carbon newly thawed from former permafrost zones, coalbeds, leakage from deep thermogenic reservoirs, and possible gas hydrate dissociation. Gas hydrate is shown concentrated in purple layers in the sediments, with the sparsity of gas hydrate portrayed here thought to be representative of these hydrates in nature []. SGD refers to relatively freshwater submarine groundwater discharge from thawing permafrost [.,]. A cryopeg is a layer of thawed sediment set within permafrost, and these features could contain methane bubbles. Not shown is riverine transport of methane, which delivers methane to coastal waters in some areas [e.g.,]. As described in the text, thermokarst lakes have been interpreted as a net sink of carbon in the Holocene [.,], despite their well‐characterized CHemissions [.,]. There is conflicting information about the predominance of COversus CHemissions from tundra due to microbial processes in shallow sediments. Not shown is methanogenesis in near‐surface ocean waters [e.g.,]. Except for the observation described earlier, in which Dallimore et al. [2008] detected tundra gas emissions having the same composition as the gas mixture in underlying gas hydrates, there has so far been no evidence to link tundra CH 4 emissions to dissociation of gas hydrates at depth. Heterogeneous thawing of permafrost may deliver substantially more heat to great depths in one location than another, implying that there may be locations where intrapermafrost or subpermafrost gas hydrate could be affected as climate warming continues. For example, thick thaw bulbs (taliks) beneath thermokarst lakes may thermodynamically perturb or even intersect the gas hydrate stability zone [e.g., Nicolsky et al., 2012; Frederick and Buffett, 2014], and a Mackenzie Delta seismic study by Bellefleur et al. [2009] images one such candidate talik. In recent years the discovery of deep, rapidly developed Yamal Peninsula craters that emit CH 4 has been attributed by some to thawing gas hydrates, although recent analyses of high‐resolution satellite imagery implies pingo collapse as a more likely cause [Kizyakov et al., 2015]. Although PAGH dissociation likely played an outsized role in CH 4 emissions from the deglaciating snowball Earth episodes during the late Neoproterozoic [Kennedy et al., 2008], even unanticipated, centuries‐scale outgassing of a portion of the 20 Gt C (26,600 Tg CH 4 ) estimated to be sequestered in PAGH deposits onshore and offshore [Ruppel, 2015] would have little impact on atmospheric CH 4 given current annual emissions of ~555 Tg CH 4 .

6.2 Shallow Water: Gas Hydrates Associated With Subsea Permafrost A special class of PAGH are those associated with subsea permafrost (Figures 9 and 10). Warming after the Last Glacial Maximum (LGM; ~20 ka) thawed large continental ice sheets (e.g., Laurentide and Fennoscandian) and raised global sea level up to ~125 m [e.g., Fairbanks, 1989], resulting in the inundation of permafrost tundra at the edges of the Arctic Ocean. The inundation replaced frigid average annual air temperatures with ocean water that was warmer by as much as 10–15°C [Kvenvolden, 1988b; Shakhova et al., 2010a]. On Arctic Ocean continental shelves, the inundation‐associated warming led to thawing of some of the continuous permafrost, with the remainder comprising contemporary subsea permafrost. Early maps [Brown et al., 1997] placed the edge of subsea permafrost at the shelf break (~100 m water depth) in the Arctic Ocean, but geophysical studies reveal that contemporary ice‐bearing subsea permafrost does not extend much beyond the 20 or 30 m isobath in the U.S. Beaufort and Kara Seas [Brothers et al., 2012, 2016; Portnov et al., 2013; Ruppel et al., 2016], may remain in patches farther offshore in the Laptev Sea [Rekant et al., 2005, 2015], and could reach the 100 m isobath in parts of the Canadian Beaufort Sea [Hu et al., 2013; Hunter and Hobson, 1974]. Relative to the Brown et al. [1997] map, the more restricted distribution of relict permafrost beneath Arctic Ocean continental shelves likely means that some PAGH have already dissociated since the onset of post‐LGM inundation. Gas hydrates associated with subsea permafrost lie in water‐covered areas, but they have some characteristics that distinguish them from both marine gas hydrates and their terrestrial PAGH counterparts. As noted above, marine gas hydrates cannot form in the contemporary ocean beneath waters shallower than ~300 to 600 m. However, PAGH that now lie offshore on high‐latitude continental shelves are in sediments covered by a maximum of 100 to 120 m of water. This adds up to ~1 MPa of additional hydrostatic pressure, may cause formation of new gas hydrate near the top of the stability zone, and serves as a slightly stabilizing influence offsetting the profound destabilization caused by warming of the sedimentary section during and after inundation. Dissociating gas hydrates associated with subsea permafrost liberate CH 4 that is subject to the full suite of marine sedimentary and water column sinks that affect marine gas hydrates [Overduin et al., 2015; Thornton and Crill, 2015]. However, sulfate, which is necessary for marine sedimentary AOM sink processes, may not fully intrude sediments inundated only since ~15 ka, and the shallow water depths mean that gas emitted at the seafloor is more likely to reach the sea surface before bubble stripping [McGinnis et al., 2006], CH 4 dissolution, and/or microbial oxidation can have a large mitigating impact. As in other settings, CH 4 liberated from dissociating gas hydrates at depth must also navigate overlying sediments to reach the seafloor. In addition, ice‐related processes have contributed to the widespread development of indurated (low‐permeability) sediments that could be particularly effective at trapping CH 4 beneath some Arctic Ocean shelves. Like other PAGH, those that ended up within submerged shelves are unlikely to be widely distributed or sequester large amounts of CH 4 [Ruppel, 2015]. Some researchers do infer large amounts of PAGH beneath arctic continental shelves (e.g., 35 Gt C in hydrate beneath the Laptev Sea shelf) [Shakhova et al., 2010a], but several assumptions used in making this estimate may not fully account for the complexity of PAGH systems. Shakhova et al. [2010a] also invoked anomalous shallow gas hydrates beneath the East Siberian Arctic shelf as a potential CH 4 source and to explain elevated estimates of CH 4 sequestered in gas hydrates. This area was not glaciated at the LGM, as is usually required for shallow gas hydrates to occur, and the origin and existence of possible anomalous gas hydrate deposits remain controversial and require further examination. Intact arctic continental shelf gas hydrate certainly remains today within or beneath subsea permafrost, but distinguishing hydrate‐ from ice‐bearing sediments based on geophysical data is nearly impossible without direct sampling. During the top‐down warming associated with Holocene inundation, gas hydrate deposits require longer to leave their stability field and to degrade than the associated permafrost takes to thaw (Figure 4a). Even without anomalous preservation processes such as submarine groundwater discharge [e.g., Frederick and Buffett, 2016], hydrate can remain where ice‐bearing permafrost is no longer detectable [e.g., Pokrovsky, 2003]. Where ice‐bearing subsea permafrost has now thawed on the U.S. Beaufort margin, Collett et al. [2011] identify possible gas hydrate at 530 m below the seafloor, and drilling of an outer shelf well yielded gas hydrate from the formation at a maximum depth of 754 m below seafloor [Ruppel et al., 2016]. The fact that hydrate can remain in the section even where the subsea permafrost has completely thawed means that (a) methane may be released by hydrate dissociation over a region that extends beyond the seaward edge of subsea permafrost [Paull et al., 2007; Portnov et al., 2013; Serov et al., 2015; Shakhova et al., 2010a, 2010b], sometimes at pingo‐like features; and (b) methane emissions from these dissociating hydrates could lag permafrost thaw by hundreds or thousands of years. Circum‐Arctic Ocean continental shelves have long been presumed as a source of atmospheric CH 4 emissions [e.g., Kvenvolden et al., 1993], and attention in recent years has focused on the Siberian shelves, where [Shakhova et al., 2014] estimate annual atmospheric CH 4 emissions of up to 17 Tg CH 4 when ebullitive and diffusive fluxes are combined. Thornton et al. [2016] described a continuous shipboard survey of CH 4 concentrations in the atmosphere and near‐surface waters in much of this same area. They conclude that ebullition does not substantially contribute to the sea‐air CH 4 flux, which they calculate to be less than 2.9 Tg yr−1 CH 4. They also note that some of the previously reported atmospheric CH 4 concentrations on the East Siberian Arctic shelf may be unrealistic. Like Kvenvolden et al. [1993] and Kort et al. [2012], Thornton et al. [2016] underscored the critical role of sea ice in trapping CH 4 until leads or ice‐out conditions render possible the diffusive release across the sea‐air interface. Continuous sea‐air CH 4 flux surveys like that of Thornton et al. [2016] have also been conducted on the U.S. and Canadian Beaufort shelf [Pohlman et al., 2012]. Measurements there reveal high nearshore CH 4 concentrations inferred to be produced in organic‐rich sediments, but regional annual flux is several orders of magnitude lower than the Thornton et al. [2016] estimate for the Siberian shelves and comparable to that recorded for the North Sea [Bange et al., 1994]. Several key lessons emerge from these studies: (1) Accurate constraints on CH 4 emissions to the atmosphere are particularly critical in settings where gas hydrates may be the most susceptible to global warming processes, such as Arctic Ocean continental shelves. Refining sea‐air CH 4 fluxes in these settings may address concerns raised about catastrophic methane releases accompanying continued climate warming [Whiteman et al., 2013] and help to reconcile bottom‐up CH 4 measurements with top‐down time series of atmospheric CH 4 observations, which show no recent increase in arctic atmospheric CH 4 emissions [Bruhwiler et al., 2015; World Meteorological Organization, 2013]. (2) Methane fluxes to the atmosphere vary significantly on high‐latitude continental shelves, and measurements made in one area cannot be upscaled for application to the entire Circum‐Arctic Ocean. (3) Numerous sources—most of them shallower and more directly connected to the seafloor than dissociating gas hydrates—contribute to CH 4 concentrations in the seawater of Arctic Ocean continental shelves and CH 4 emissions to the atmosphere (Figure 10). Only by systematically eliminating the contributions of most of these sources will it be possible to identify emissions that could be related to gas hydrates.

6.3 Gas Hydrates in Glaciated Areas Ruppel [2011a] identified only five geographic sectors for gas hydrates in assessing their interaction with the climate system, and a missing category was gas hydrates that formed beneath ice sheets during the late Pleistocene [e.g., Nisbet, 1990b] or that currently occur beneath ice sheets (Figures 9 and 11). Subglacial hydrates in Antarctica could presently sequester enough methane to increase global estimates of gas in place [Wadham et al., 2012], and Portnov et al. [2016] recently described PAGH that could have been stable under the West Spitsbergen ice sheet at the LGM for either warm‐base or cold‐base conditions. They postulate that degassing of these gas hydrates during post‐LGM ice sheet thaw may have occurred in shallow waters at the ice sheet's terminus as shelfal inundation got underway. Portnov et al. [2016] also propose that these degassing processes could have preconditioned upper continental slope sediments to be the loci of contemporary ebullitive CH 4 flux [e.g., Westbrook et al., 2009], although their extrapolation to the northern U.S. Atlantic margin seeps described by Skarke et al. [2014] does not appropriately consider well‐established LGM ice distributions on that margin. Figure 11 Open in figure viewer PowerPoint Sloan and Koh [ 2008 −1, with constant thermal properties assumed for both the permafrost and subpermafrost zones, a substantial simplification. Ice density for pressure calculations is 910 kg m−3; using a higher density will result in a thinner GHSZ. Subglacial gas hydrates are likely widespread in Antarctica [Wadham et al., 2012 Nisbet, 1990b Portnov et al., 2016 Maslin et al., 2010 Hunter and Hobson [ 1974 The theoretical distribution of permafrost (ground with temperature less than 0°C whose base is delineated by the thick white line) and the pure methane hydrate GHSZ (calculated with 0% pore water salinity from]) for warm‐base (0.5°C; purple represents GHSZ) and cold‐base (−5°C; blue denotes GHSZ) ice sheets that are 200 m, 500 m, and 1000 m thick. The thick white line marks the base of permafrost. The assumed conductive geothermal gradient is 30°C km, with constant thermal properties assumed for both the permafrost and subpermafrost zones, a substantial simplification. Ice density for pressure calculations is 910 kg m; using a higher density will result in a thinner GHSZ. Subglacial gas hydrates are likely widespread in Antarctica [.,], and thawing ice sheets and rising sea levels could have destabilized significant subglacial deposits at the end of the LGM [.,] or during other time periods [.,]. In practice,] found that ice‐bearing permafrost in the Beaufort Sea corresponds to temperatures less than −1.8°C, so the extent of permafrost is likely overestimated in this calculation. Figure 11 shows nominal conditions for permafrost evolution and gas hydrate stability beneath cold and warm‐base ice sheets. Even where permafrost is lacking beneath warm‐base ice, gas hydrate is stable at shallow depths in the sedimentary section for ice sheets a mere 500 m thick. Such shallow hydrates could form from microbial gas instead of the thermogenic gas thought to be sourcing many contemporary PAGH [Ruppel, 2015]. Anomalously shallow gas hydrates have been postulated for the Yamal Peninsula [Chuvilin et al., 2002] and invoked to explain some observations on the East Siberian Arctic Shelf [Shakhova et al., 2010a], as discussed above. Neither area was glaciated at the LGM, and the shallow gas releases on which the anomalous hydrate interpretation is based [Chuvilin et al., 2002] are common in permafrost areas during drilling and thought to be unrelated to gas hydrate dynamics. Even if proof for anomalous gas hydrates is eventually found, it remains uncertain how the pressure and temperature conditions at shallow depths (e.g., less than 100 m) could have been within the gas hydrate stability field absent recent glacial loading or a highly unusual mixture of hydrocarbons. While the possible existence of relict subglacial gas hydrates at high northern latitudes is not likely to greatly expand the estimate of global gas in place, a recent review of Pleistocene glacial extents [Jakobsson et al., 2014] could provide guidance for locating other margins where these unusual hydrates may have existed at the LGM and may have degassed during the Holocene. This category of gas hydrates should probably also be included in millennial scale models of global warming impacts on the Greenland and Antarctic ice sheets [Maslin et al., 2010].

6.4 Upper Continental Slopes As described by Kvenvolden [1988b] and considered in numerous observational and modeling studies [e.g., Berndt et al., 2014; Brothers et al., 2014; Davies et al., 2015; Gorman and Senger, 2010; Johnson et al., 2015; Kennett et al., 2003; Marín‐Moreno et al., 2015; Marín‐Moreno et al., 2013; Mienert et al., 2005; Pecher et al., 2005; Phrampus and Hornbach, 2012; Phrampus et al., 2014; Reagan and Moridis, 2009; Reagan et al., 2011; Ruppel, 2011a; Skarke et al., 2014; Stranne et al., 2016a; Stranne et al., 2016b; Weinstein et al., 2016; Westbrook et al., 2009] gas hydrates within upper continental slope sediments constitute the key marine hydrate population that is susceptible to degradation during ocean warming (Figure 9). The GHSZ vanishes on upper slopes (i.e., the “feather edge” of hydrate stability in Ruppel [2011a]). Small perturbations in the temperatures of impinging intermediate ocean waters or even small pressure perturbations associated with tides or other oceanographic phenomena can affect the stability of these deposits. Dissociation driven by these processes may also condition submarine slopes to failure, particularly at the shallower water depths (300–800 m) close to the landward limit of gas hydrate stability [Nixon and Grozic, 2006]. Based on potential distributions of upper slope gas hydrates on marine continental margins and conservative assumptions about gas hydrate saturations, but ignoring any biogeochemical sinks in the sediments, Ruppel [2011a] estimated that ~3.5% of the global gas hydrate inventory (~63 Gt C or ~83,790 Tg CH 4 based on an assumed inventory of 1800 Gt C) might be susceptible to climate change on a time scale of centuries. While upper continental slope gas hydrates are generally viewed as being in a net dissociation regime in light of contemporary climate warming, the details are certainly more complicated, with gas hydrate dissociating and re‐forming at the shallowly buried BGHS in response to oscillating temperatures and pressures on the slopes. Key questions include the degree to which gas hydrates remain out of equilibrium with local P‐T conditions, the rate at which upper slope gas hydrates respond to climate forcing, and whether upper slope hydrate dissociation processes that do not produce seafloor seepage can be detected by geophysical surveys. Despite the expectation that upper continental slopes host the most climate‐susceptible gas hydrate populations, widespread upper slope seepage has so far only been recognized on the West Spitsbergen margin [Westbrook et al., 2009], the U.S. Atlantic margin [Skarke et al., 2014], and the northwestern U.S. Pacific margin [Johnson et al., 2015]. Given the role of salt tectonism (section 6.5) and the lack of a systematic catalog for northern Gulf of Mexico seeps associated with hydrates, these are not considered here, although upper slope dissociation processes are clearly active [e.g., MacDonald et al., 1994; Weber et al., 2014]. Upper continental slope seepage on the other margins has been interpreted in terms of warming of intermediate waters on time scales of years to centuries [Berndt et al., 2014; Biastoch et al., 2011; Brothers et al., 2014; Ruppel, 2011a; Stranne et al., 2016b], but so far only the West Spitsbergen margin seepage has been firmly linked to dissociating gas hydrate [Berndt et al., 2014]. On the northwestern U.S. Pacific margin, many of the seeps are close to the feather edge of hydrate stability on the upper continental slope, providing circumstantial evidence for their origin in methane hydrate destabilization [Johnson et al., 2015]. In the Arctic Ocean, some researchers have used thermal infrared (TIR) data from satellites to infer methane emissions linked to seasonal ocean warming that may drive hydrate dissociation on continental slopes and nearby shelves [Leifer et al., 2014]. However, the TIR technology has low sensitivity in the lower troposphere, rendering the data poor at distinguishing local and regional methane sources [Jacob et al., 2016]. Furthermore, the strength of water column sinks makes it unlikely that upper continental slope methane is reaching the ocean‐atmosphere interface in any case, leading to questions about the origin of the signal being detected in the satellite data. A full discussion of the approaches for estimating methane fluxes from deepwater marine gas hydrate settings is included in section 6.5, and here we focus only on the results related specifically to upper continental slopes and choose the West Spitsbergen margin [e.g., Westbrook et al., 2009] as the focus. Sahling et al. [2014] estimated flux of 9 to 118 × 106 mol yr−1 CH 4 from seeps in a depth range mostly updip of the landward edge of gas hydrate stability and note that several methane sources likely contribute to this flux. Graves et al. [2015] recorded near‐bottom methane concentrations as high as 825 nM over the upper slope seeps but show that most of this methane remains near the bottom and that the methane at shallower depths in the water column does not originate with in situ seafloor emissions. Steinle et al. [2015] documented the strength of the MOx sink in these cold waters, demonstrating that changing ocean currents have a profound effect on the efficiency of the sink. Myhre et al. [2016] and Graves et al. [2015] used direct measurements and indirect arguments to demonstrate that sea‐air flux over the upper continental slopes is not elevated, while Fisher et al. [2011] concluded that the atmospheric methane in this area lacks a signal related to gas hydrate dissociation. Nonetheless, at least for the CH 4 detected near the seafloor, there is strong evidence that it is being sourced in gas hydrate dissociation [Berndt et al., 2014]. Establishing the link between upper slope seepage and gas hydrate dissociation on the U.S. Atlantic margin will be more difficult since the gas appears to have no thermogenic component like that which assisted with fingerprinting on the West Spitsbergen margin. Furthermore, as noted by Skarke et al. [2014], the U.S. Atlantic margin seepage occurs mostly at water depths upslope from the contemporary updip limit of methane hydrate stability, except in Hudson Canyon [Weinstein et al., 2016]. The large percentage of excess heat absorbed by the Atlantic Ocean over the past few decades [Lee et al., 2011; Levitus et al., 2012] may imply greater dynamism for the GHSZ on upper slopes here than in other ocean basins and may also lead to more rapid downdip migration of the gas hydrate stability field [e.g., Brothers et al., 2014]. Still, in situ gas hydrate dissociation has been ruled out as a gas source for one prominent seep cluster [Prouty et al., 2016] that may have been active starting after the LGM, and seepage at other upper slope locations would require updip migration of gas through permeable strata [Brothers et al., 2014; Skarke et al., 2014]. In addition to seepage, certain erosional and other features have been described at the upper feather edge of gas hydrate stability on global margins [Davies et al., 2015; Pecher et al., 2005]. These merit further attention as potential markers for gas hydrate dissociation. The types of exhaustive studies of seafloor CH 4 flux rates that are available for the West Spitsbergen margin have not yet been completed for the U.S. Atlantic or northwestern U.S. Pacific margin seeps. Based on limited bubble observations, Skarke et al. [2014] gave a conservative estimate of 15–90 Mg yr−1 CH 4 for seafloor flux for the ~570 seeps they describe along 950 km of the Atlantic margin, while Weinstein et al. [2016] used indirect methods to infer 70–280 Mg yr−1 CH 4 in Hudson Canyon alone. How much, if any, of these emissions originate in dissociating gas hydrate remains unknown. Preliminary continuous sea‐air flux measurements indicate that the Atlantic margin seeps are unlikely to be contributing CH 4 to the atmosphere [Ruppel et al., 2015].