Annu. Rev. Earth Planet. Sci : Copyright 2000 by Annual Reviews. All rights reserved

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1 Annu. Rev. Earth Planet. Sci : Copyright 2000 by Annual Reviews. All rights reserved CLATHRATE HYDRATES Bruce A. Buffett Department of Earth and Ocean Sciences, University of British Columbia, Vancouver, Canada; V6T 1Z4 Key Words gas hydrates, methane source, carbon cycle, climate change, glacial cycles Abstract Substantial volumes of methane gas are trapped below the seafloor and in permafrost by an ice-like solid called clathrate hydrate. Global estimates of the methane in clathrate hydrate may exceed kg, which represents one of the largest sources of hydrocarbon on Earth. Speculations about large releases of methane from clathrate hydrate have raised serious but unresolved questions about its possible role in climate change. Progress in our understanding of clathrate hydrate has been made through integrated geophysical and geochemical surveys of known clathrate occurrences. Details from these surveys have motivated new investigations of the physical, chemical, and biological processes that contribute to growth and breakdown of clathrate hydrate in natural settings. In this article, I give an overview of recent advances and future challenges. INTRODUCTION Conditions encountered in many regions of the solar system permit mixtures of water and various gases to crystallize into an icy solid known as clathrate hydrate (Miller 1961). The crystalline structure and physical properties of clathrate hydrate are similar in many ways to those of ice. As in ice, the crystal structure in clathrate hydrate is composed of hydrogen-bounded water molecules, but their open arrangement accommodates small gas molecules in nearly spherical cavities (Jeffrey & McMullan 1967). Weak interactions between individual gas molecules and the surrounding cage of water molecules help to stabilize the structure. In fact, a significant fraction of the cavities must be occupied with gas molecules to ensure stability. Interest in clathrate hydrate arises from its capacity to store large volumes of gas. Clathrates probably formed in the solar nebula, trapping volatile gases into icy solids (Lunine & Stevenson 1985). Methane and other gases are subsequently added to planets and satellites by accretion of icy planetesimals or by infalling cometary material (Zahnle et al 1992). Although the existence of clathrates in comets is still debated (Klinger 1991), there are strong indications that clathrates /00/ $

2 478 BUFFETT play an important role in the atmospheres and interiors of satellites such as Titan (Lunine & Stevenson 1987). Clathrates are also known to occur at or near the present-day surfaces of Mars and Earth. Clathrates of carbon dioxide are stable in the Martian ice caps (Miller & Symthe 1970), where they sequester carbon dioxide and water vapor from the atmosphere. Geological evidence of past flooding on Mars is sometimes linked to a catastrophe breakdown of the clathrate hydrate and an associated release of greenhouse gases (Milton 1974, Baker et al 1991). Clathrates on Earth are found in seafloor sediments along deep continental margins and in permafrost regions (see Figure 1). The gas component is mainly methane, although small amounts of heavier hydrocarbons, carbon dioxide, and hydrogen sulfide have also been recovered (Brooks et al 1986, Kastner et al 1998). The amount of methane trapped in clathrate hydrate on Earth is uncertain, but even conservative estimates are substantial. Kvenvolden (1988a) estimates that kg of carbon are trapped in oceanic sediments in the form of methane hydrate. Smaller contributions are expected in permafrost regions (MacDonald 1990a). If these estimates are reliable, then clathrates represent the largest source of hydrocarbon on Earth. Such vast supplies of methane motivate interest in clathrate hydrate as an energy resource. The proximity of large volumes of the methane gas to the surface also raises questions about the role of clathrate hydrate in past and present-day climate change (Kvenvolden 1988b, MacDonald 1990b). Release of methane gas into the atmosphere following the breakdown (dissociation) of clathrate hydrate alters the radiative energy balance in the atmosphere because methane is a strong greenhouse gas (Lashof & Ahuja 1990). Increases Figure 1 Map showing the locations of known or inferred gas hydrate. The circles indicate sites in marine sediments, and the triangles denote permafrost sites. Adapted from Kvenvolden et al (1993).

3 CLATHRATE HYDRATES 479 in atmospheric temperature may drive further dissociation of clathrate, which provides a positive feedback for global warming. Clathrate hydrate may also play a role in past climate change. Nisbet (1990) suggests that dissociation of clathrates contributed to the rapid increase in atmospheric methane at the end of the last glaciation, which accelerated the retreat of the continental ice sheets. Alternatively, Paull et al (1991) suggest that clathrates are responsible for limiting the extent of glaciation. Because water is removed from the oceans during the growth of the ice sheets, pressure at the seafloor is lowered and clathrates are expected to become unstable. Many of these speculations await further quantitative testing, but the implications are far-reaching. The goal of this review is to assess our current understanding of clathrate hydrates in natural environments and to identify the challenges that remain in assessing their role in climate change and other geological processes. Before we proceed to more detailed discussions of clathrate hydrate in the geological context, it is useful to make a few remarks regarding terminology. Clathrate hydrate is distinguished from several other types of hydrate by the nature of the interaction between gas and water molecules. In clathrate hydrate the gas and water interact through weak van der Waals forces. Other types of hydrate exist where the gas combines chemically with water molecules in fixed proportions (ammonia hydrate being one example). We exclude these types of hydrates from consideration in this review and follow customary practice by referring to clathrate hydrate as gas hydrate. For the purposes of this review, the terms clathrate hydrate, gas hydrate, or simply hydrate are used interchangeably. CRYSTAL STRUCTURE AND STABILITY Thermodynamic conditions for the stability of clathrate hydrate are strongly dependent on the size and shape of the gas component. The gas molecules must be small enough to fit into the cavities of the lattice but large enough to lend stability to the structure. Early crystallographic studies of the clathrate structure (Pauling & Marsh 1952) initiated the development of a statistical thermodynamic model by van der Waals & Platteuw (1959). This model remains in use today for phase equilibria calculations, although more recent advances have refined some of the original approximations (Rodger 1990, Sparks & Tester 1992, Hwang et al 1993, Tanaka & Kiyohara 1993). Clathrate hydrate can possess many different crystal structures (Belosludov et al 1991), but only three structures are known to occur in natural environments. Structure I and Structure II each have two cavity sizes (small and large), whereas the recently discovered Structure H has three different types of cavities (Ripmeester et al 1987). Structure I is the most common form of clathrate in natural settings where methane is the main hydrate-forming gas. Structures II and H have been reported when the gas mixture includes larger molecules (Brooks et al 1986, Sassen & MacDonald 1994).

4 480 BUFFETT TABLE 1 The shape of the small and large cavities of Structure I are illustrated in Figure 2. The small cavity is composed of 20 water molecules arranged to form 12 pentagonal faces. The resulting polyhedra is known as a dodecahedron, although a more convenient nomenclature is 5 12 (Jeffrey & McMullan 1967). The large cavity contains 24 water molecules, which form 12 pentagonal and 2 hexagonal faces (e.g ). In Structure I the small cavities are located at the center and the four corners of the unit cell to form a body-centered cubic structure. Six additional water molecules inside the unit cell (for a total of 46) link the small cavities to form the large cavities. Each unit cell contains 2 small and 6 large cavities. Structure II is also constructed from 5 12 cavities, though a different arrangement of the small cavities creates large cavities of type and introduces a small distortion into the small 5 12 cavities (Sloan 1998). Table 1 summarizes the key structural features of Structures I and II. Additional information on Structure H is given by Ripmeester et al (1994). Clathrate structures are determined mainly by the size of the gas molecules. Methane gas is small enough to enter the small and large cavities of both Structure I and II, but Structure I is preferred because methane contributes more to the stability of Structure I. Because the two structures differ most in the size and number of their large cavities (see Table 1), the preferred structure results from a more favorable fit of the methane molecule into the large cavity of Structure I. Figure 2 The shape of the small (left) and large (right) cavities in Structure I hydrate. Structure of hydrate crystals Property Structure I Structure II Cavity Small Large Small Large Description Cavities/unit cell Average cavity radius (nm) Water molecules/unit cell Source: Sloan 1998

5 CLATHRATE HYDRATES 481 However, mixtures of methane and larger gas molecules can promote the formation of Structure II. For example, propane is too large to fit into either cavity of Structure I, but it fits comfortably into the large cavity of Structure II. Even small quantities of propane can enhance the stability of Structure II over Structure I, lowering the formation pressure relative to that of pure methane clathrate. Reductions in the formation pressure are also observed when methane is mixed with other gases, including ethane, carbon dioxide, and hydrogen sulfide (Sloan 1998). The stability of clathrate hydrate is typically predicted using the statistical thermodynamic model of van der Waals & Platteuw (1959). Clathrates are viewed as a solid solution of gas molecules in a metastable lattice of water molecules. Each gas molecule is confined to a single cage, where it interacts with the surrounding water molecules through weak van der Waals (dispersive) forces. These interactions lower the free energy of the water molecules, making the clathrate structure stable when a sufficient number of cavities are filled with gas. (We denote the occupied fraction of small and large cavities by h s and h l, respectively.) h The change Dl w in the free energy of the water molecules per mole (e.g. chemical potential) is expressed by van der Waals & Platteuw (1959) in the form h Dlw RT[msln(1 h s) mlln(1 h l)], (1) where m s and m l are the number of cavities per water molecule in the clathrate lattice, R is the gas constant, and T is the temperature. For Structure I, m s 1/23 and m l 3/23 (see Table 1). According to Equation 1, the reduction in the free energy of the water molecules depends only on the occupancy of the small and large cavities; there is no explicit dependence on the type of gas molecules occupying the cavities. Differences resulting from the size and shape of the gas molecules enter the thermodynamic model through their influence on h s and h l. For a single gas component, the occupancy is given by Cf s Cf l hs ; hl, (2) 1 Cf 1 Cf s where C s and C l are constants that depend on the interaction between the gas and water molecules in the small and large cavities, and f is the fugacity of the gas. Extensions to the theory for mixtures of gases are straightforward (Sloan 1998). The occupancies in Equation 2 are expressed in a form that is identical to Langmuir s expression for the adsorption of gas onto a surface. In Langmuir s theory, C s and C l are called Langmuir constants, and f reduces to the partial pressure for an ideal gas. This analogy with Langmuir s theory provides a useful conceptual model for describing the stability of clathrate hydrate. High gas pressure (corresponding to high f) increases the occupancy and lowers the free energy of the water molecules in the clathrate structure. A second contribution to the free b energy of the water molecules, Dl w, represents the excess free energy needed to l

6 482 BUFFETT assemble the clathrate lattice from either liquid water or ice, depending on temperature. Clathrate becomes stable when the energy reduction Dl w resulting from h b occupation overcomes the energy excess Dl w associated with reorganizing water molecules into the clathrate lattice. Estimates of b h Dlw and Dlw cannot be uniquely obtained from experiments because it is not possible to construct an empty clathrate lattice. In practice, Dl w is obtained by fitting experimental measurements of b clathrate formation temperatures and pressures for a number of different gases. b The resulting estimates of Dl w are assumed to be independent of the gas com- ponent, with the assumption that the gas molecules do not distort the clathrate lattice. Although this assumption is reasonable for small molecules such as methane, corrections have been proposed for some larger gases (Hwang et al 1993). Corrections must also be applied to account for the influences of pressure and b temperature on Dl w (see Holder et al 1988 for a review). Phase Equilibria Establishing the stability conditions for clathrate hydrate is essential for understanding their role in geological processes. An important condition for stability is defined by requiring the clathrate phase to coexist with both the liquid and vapor phases in a three-phase equilibrium. This particular equilibrium state occurs at temperature T 3 (P), which is solely a function of pressure P. Estimates of T 3 (P) for methane gas and pure water are shown in Figure 3, based on the phase equilibrium calculations of Sloan (1998). Pressure is expressed as an equivalent depth, assuming a hydrostatic pressure gradient of 10 4 Pa m 1. Methane hydrate is stable when the temperature is less than or equal to T 3 (P). When T T 3 (P), the clathrate phase coexists with either the liquid or vapor phases, depending on the relative abundance of water and gas. In the seafloor we expect water to be in greater abundance, so the relevant two-phase equilibrium occurs between clathrate and Figure 3 Phase diagram for a methanewater mixture as a function of pressure and temperature. Pressure is plotted as an equivalent depth assuming a hydrostatic gradient of 10 4 Pa m 1. Temperature T 3 (P) bounds the region where hydrate is stable and T m is the melting temperature of pure water. Depth (m) ice + vapor vapor + hydrate liquid + hydrate T m liquid + vapor T (P) Temperature (K)

7 CLATHRATE HYDRATES 483 Temperature (K) liquid solution. In permafrost regions, where temperatures fall below 0 C, clathrate coexists with the vapor phase. Ice cannot coexist with clathrate in two-phase equilibrium because the concentration of gas in ice is too low. Gas molecules may occupy defects in the ice crystal, but the low gas concentration implies a low fugacity, f, which is too small to stabilize clathrate. Ice should convert into clathrate in the presence of methane gas, but the low diffusivity of gas in ice (Stern et al 1996) may permit ice to persist under nonequilibrium conditions. Although the conditions for three-phase equilibrium have been studied extensively and are relatively well known (Englezos et al 1990, Sloan 1998), comparatively few studies have addressed the conditions for equilibrium between clathrate and liquid solution. Calculations by Miller (1974) and Handa (1990) predict a modest decrease in the concentration of dissolved gas as pressure increases at constant temperature. In effect, gas is squeezed from the liquid solution, increasing the relative abundance of the clathrate phase. More recently, Zatsepina & Buffett (1997) have emphasized the role of temperature when clathrate coexists with liquid solution. Their results show how the composition of the liquid solution changes as a function of temperature (see Figure 4). The solid line in Figure 4 represents the solubility of methane in pure water. We observe that warming a mixture of methane and water above T 3 (P) lowers the solubility; the opposite behavior occurs at temperatures below T 3 (P) when hydrate is present. The reduction in solubility with cooling below T 3 (P) has two important consequences. First, cooling causes growth of the clathrate phase because gas in excess of the solubility is transferred out of the liquid phase. Second, the temperature at which clathrate becomes stable depends on both pressure and concentration of dissolved gas. Indeed, two degrees of freedom are predicted by Gibbs phase rule in a two-component mixture containing only two phases. For a given pressure, the maximum temperature for hydrate stability [e.g. T 3 (P)] coincides with the P = 20 MPa liquid liquid + hydrate liquid + vapor T (P) Figure 4 Phase diagram for a methane-water mixture as a function of temperature and gas concentration. The solid line represents the solubility of gas in equilibrium with either vapor or hydrate. Gas Concentration (mole fraction)

8 484 BUFFETT largest concentration of dissolved gas in Figure 4. At lower gas concentrations the temperature for stability decreases. Other factors that influence the stability of clathrate hydrate in natural environments include the effects of dissolved salts and sediment properties. Dissolved salts lower T 3 (P) by1 2 C for a plausible range of salinities in the seafloor sediments (Englezos & Bishnoi 1988, Dickens & Quinby-Hunt 1997). From a thermodynamic point of view, the main effect of salt on the three-phase equilibrium is to lower the free energy of the water molecules in the liquid, which inhibits the formation of clathrate. Salt also causes an increase in the free energy of gas molecules in solution, but this has little influence on the stability of clathrate; gas is simply driven from solution into the vapor phase. When the vapor phase is absent, the effect of salt on the phase equilibrium is more subtle. Salt continues to lower the free energy of the water molecules in solution, but the increase in the free energy of the dissolved gas molecules tends to drive gas into the clathrate phase. The net effect is a 4 5% reduction in the solubility of methane gas when hydrate is present (Zatsepina & Buffett 1998). Less gas is required in solution to stabilize the clathrate phase, which means that salt actually promotes the formation of clathrate from liquid water in the absence of gas bubbles. Because clathrate excludes salts during crystallization, we might expect the resulting increases in salinity to provide a positive feedback for further crystallization from solution. However, depletion of gas from solution as a result of clathrate formation is a much stronger negative feedback, which maintains the stability of the system in response to small changes in thermodynamic conditions. Sediment properties may also contribute to the stability of clathrate hydrate, especially when pore volumes in the sediments are very small. Geometric restrictions imposed by the sediments on the size of crystal nucleii make very small crystals energetically unfavorable. The surface energy of nucleii in restricted volumes lowers the temperature required to promote crystallization compared with bulk conditions (Handa et al 1992). Sediments may also lower the free energy of liquid water in the pore because of interactions with the pore wall. Water molecules nearest the wall can remain bound to the surface and do not freeze during cooling, whereas water molecules farther from the wall experience a more modest reduction in free energy. Reductions in the freezing temperature of pure water in small pores are well documented (Rennie & Clifford 1977). Analogous effects were observed by Handa & Stupin (1992) in the formation of methane hydrate. In their study, methane hydrate was formed in silica gel with a typical pore radius of m. The measured formation temperature was approximately 6 C below bulk predictions. The measurements of Handa & Stupin were shown to be in good agreement with theoretical predictions, using the solid solution model of van der Waals & Platteuw (1959), when the free energy of the pore water was inferred from the freezing depression of pure water in the same medium (Handa et al 1992). This suggests that relatively simple experiments on the freezing temperature of pure water in sediments can be used to estimate shifts in the formation temperature of

9 CLATHRATE HYDRATES 485 clathrate hydrate. In natural sediments, the depression of T 3 (P) compared with bulk equilibrium predictions is expected to be small because the typical radii of pores in fine-grain clays are roughly 10 7 m, which is large compared with the pore sizes used in the study of Handa & Stupin (1992). Clennell et al (1999) estimate that changes in T 3 (P) should not exceed a few tenths of a degree, based on sediment samples collected at clathrate occurrences on the Blake Ridge. Application to Geological Environment Clathrate stability near the Earth s surface is established by applying phase equilibrium predictions to geological conditions. Low temperatures in permafrost regions cause ice to form in the shallow sediments where the pressure is too low to stabilize hydrate (see Figure 5a). At greater depths, ice is converted into hydrate in the presence of methane gas. According to equilibrium predictions, ice should vanish below 200 m, but nonequilibrium conditions may persist because of the slow diffusion of gas into ice. The base of the clathrate stability zone extends to the depth where temperature reaches T 3 (P). Below this depth we expect liquid water and bubbles of methane gas. In marine sediments along deep continental margins, the temperature at the seafloor is a few degrees above 0 C, but the pressure of the overlying water is sufficient to stabilize methane hydrates (see Figure 5b). The depth of the hydrate stability zone is limited by the increase in temperature below the seafloor. If the abundance of gas is sufficient, then the base of the stability zone is defined by Depth Temperature surface T base of permafrost T base of hydrate stability (a) T (P) 3 Depth Temperature T (P) 3 T seafloor T base of hydrate stability (b) Figure 5 A schematic profile of temperature T in (a) continental permafrost and (b) marine sediments. Hydrate stability in the sediments (shaded region) is limited to depths where T T 3 (P).

10 486 BUFFETT the intersection of the temperature profile with the temperature for three-phase equilibrium. In the shallower sediments, clathrate coexists with seawater in a twophase equilibrium, whereas at greater depths we expect to find seawater in equilibrium with gas bubbles. When the solubility calculations of Zatsepina & Buffett (1998) are applied to seafloor conditions, we obtain estimates of the dissolved gas concentrations necessary for hydrate stability (see Figure 6). Below the stability zone, the solubility of gas is nearly constant because the changes in pressure and temperature have counteracting effects. Inside the stability zone, however, the solubility is largely controlled by temperature. As a result, the solubility decreases sharply toward the seafloor. If the concentration of dissolved gas falls below the solubility, then hydrate dissociates to reestablish the equilibrium concentration. On the other hand, hydrate crystallizes when the gas concentration exceeds the solubility in order to restore the dissolved concentration to equilibrium. Thermodynamic predictions provide powerful constraints on the structure of hydrate occurrences in the seafloor. For example, the base of the stability zone coincides with the temperature for three-phase equilibrium only when gas bubbles are present in the underlying sediments. In the absence of gas bubbles we expect the concentration of dissolved gas to be less than the local solubility at the depth coinciding with T 3 (P). Consequently, hydrate is not stable at this depth, and the base of the stability zone shifts toward the seafloor until the actual gas concentration equals the local solubility (Xu & Ruppel 1999). Indeed, hydrate should be absent from any region inside the stability zone where the concentration of dissolved gas is persistently less than the local solubility. Conversely, the concentration of gas in the pore fluid should equal the local solubility where gas hydrate is present. Figure 6 Solubility of methane in seawater as a function of depth below the seafloor. Pressure in the sediments increases hydrostatically through the seafloor below a water depth of 2 km. Temperature at the seafloor is 2 C and the temperature gradient is 0.05 Cm 1. Depth below seafloor (m) liquid base of hydrate stability zone liquid liquid + hydrate liquid + vapor Gas Concentration X (mole fraction)

11 CLATHRATE HYDRATES 487 Thermodynamic considerations may also explain why gas bubbles in the sediments appear to accumulate at the base of the hydrate stability zone. Gas in the vapor phase is thermodynamically unstable inside the stability zone, so the occurrence of gas bubbles should terminate abruptly at the depth for three-phase equilibrium without requiring a permeability barrier. As sedimentation buries gas hydrate to greater depths, the temperature eventually exceeds T 3 (P) and the hydrates dissociate (Kvenvolden & Barnard 1983). Gas bubbles appear below the stability zone because the concentration of gas in the pore water prior to dissociation is fixed by the local solubility. Subsequent addition of gas resulting from hydrate dissociation inevitably exceeds the solubility and produces gas bubbles. Nucleation of gas bubbles from solution requires a modest supersaturation of gas in the pore fluid, which can influence the depth at which gas bubbles first appear. Theoretical calculations show that gas hydrate can persist metastably in the supersaturated fluid at temperatures above T 3 (P) prior to the formation of gas bubbles. Once the gas bubbles appear, however, the dissolved gas concentration returns to equilibrium levels, and any remaining hydrate rapidly dissociates. Experiments in natural porous media indicate that the overheating of hydrate can exceed several degrees (Buffett & Zatsepina 1999). OBSERVATIONS FROM NATURAL OCCURRENCES Evidence of gas hydrate in the seafloor and in polar environments is compiled from a variety of sources. In some instances, samples of hydrate have been recovered by deep sea drilling or sediment coring (Kvenvolden et al 1993). More often, the presence of hydrate is inferred indirectly from geophysical and geochemical observations. The most detailed information is obtained from borehole measurements collected at sites of deep sea drilling. Geophysical measurements of seismic velocity, electrical resistivity, and porosity provide information about the physical characteristics of hydrate occurrences. Geochemical analysis of the pore fluids and exsolved gases constrains the abundance of hydrate or the conditions under which the hydrate formed. Information from borehole measurements is necessarily confined to the immediate vicinity of the drilling site. More regional inferences about areal distribution or lateral variations of hydrate abundance are usually obtained by seismic methods. Although several alternative approaches have been proposed (e.g. Willoughby & Edwards 1997, Edwards 1997), these methods have not seen widespread use. In this section we summarize the observations that are available to study hydrate occurrences, along with the conclusions that may be drawn from these observations. Geophysical Observations Most occurrences of gas hydrate in marine sediments are associated with a seismic reflection that parallels the reflection from the seafloor. This bottom-simulating reflector (BSR) is observed near the predicted base of the hydrate stability zone,

12 488 BUFFETT although small discrepancies have been reported (Ruppel 1997). Reflection amplitudes from many BSRs imply an abrupt and significant change in seismic velocity. Estimates of the velocity change are obtained by comparing reflection amplitudes from the BSR with those from the seafloor. Vertical reflection coefficients of 10 15% have been reported at many locations where strong BSRs are observed (Shipley et al 1979, Miller et al 1991, Hyndman & Spence 1992, Wood et al 1994). The polarity of the BSR reflection is opposite to that from the seafloor, which means that the BSR marks an interface between high-velocity sediments inside the stability zone and low-velocity sediments immediately below the stability zone. High velocities in the overlying layer are generally attributed to the presence of hydrate in the pore volume (Stoll et al 1971). Increases in velocity occur because the pore fluid is replaced with a higher velocity icy solid. Further increases may result if hydrate cements and strengthens the sediment matrix (Dvorkin et al 1991). Low velocities in the underlying sediments are usually attributed to the presence of gas bubbles; even small amounts of gas bubbles can dramatically reduce P-wave velocities (Murphy 1984). Debate has arisen over the relative importance of hydrate and gas bubbles in explaining the BSR reflection. The study of Hyndman & Spence (1992) on the Cascadia margin attributed most of the velocity contrast at the BSR to the presence of gas hydrate in the overlying sediments. They required 30% of the pore volume to be filled with hydrate to account for the amplitude of the BSR reflection. In the same study area, Singh et al (1993) argued for gas bubbles in the underlying sediments as the main cause of the BSR reflection. Their best fit to the observed seismic waveforms was obtained with a small volume of gas bubbles (a few percent of the pore volume) in a 30-m-thick layer below the BSR. More detailed information about the seismic structure of hydrate occurrences is available from recent sites of the Ocean Drilling Program (ODP). Vertical seismic profiling (VSP) employs seismic sources at the sea surface and receivers at a series of locations in the borehole. Narrow spacings between receiver locations (typically a few meters) permit detailed estimates of the seismic velocity as a function of depth. MacKay et al (1994) reported VSP data from ODP Leg 146 on the Cascadia margin near the site of the earlier seismic surveys. The VSP data revealed anomalously fast velocities above the BSR and an abrupt drop in velocity below the BSR (see Figure 7a). The low velocity below the BSR indicated a thin layer of gas bubbles, but the anomalously fast velocities in the overlying layer suggest that both hydrate and gas bubbles contribute to the amplitude of the BSR reflection (Yuan et al 1996, 1999). Quantitative interpretation of the seismic data requires (a) estimates of the sediment velocity in the absence of hydrate and (b) a relationship for the effects of hydrate on seismic velocity. Yuan et al (1996) addressed the question of estimating the velocity of unhydrated sediments using multichannel seismic (MSC) data. Trends in the seismic velocities below the BSR were used to extrapolate the velocity of unhydrated sediment into the stability zone. The result was significantly lower than the observed velocity in the stability region. The task of infer-

13 CLATHRATE HYDRATES 489 Figure 7 (a) Estimates of seismic velocity in the shallow sediments at or near Site 889, ODP Leg 146. The circles indicates estimates from multichannel seismic (MCS) data. Results from vertical seismic profiling (VSP) are indicated by the irregular solid line, and the two thin dashed lines are acoustic logs at Sites 889A and 889B. The smooth trend is the estimate of velocity in unhydrated sediments. (b) Hydrate volumes inferred from the seismic velocity estimates. Adapted from Yuan et al (1996). ring the amount of hydrate in the sediments based on the anomalous seismic velocity is more difficult. Several simple approaches have been adopted by various investigators (Hyndman & Spence 1992, Lee et al 1996). Although the reliability of these methods is not firmly established, the similarity of estimates obtained using these different approaches is encouraging (Hobro et al 1998). Estimates of the hydrate distribution from the study of Yuan et al (1996), shown in Figure 7b, use the porosity-reduction method described by Hyndman & Spence (1992). Hydrate occurrences on passive continental margins share some similarities with those on active margins, but a number of differences are also apparent. VSP data from ODP Leg 164 on the Blake Ridge revealed a thick layer of gas bubbles extending more than 200 m below the BSR at sites 995 and 997 (Holbrook et al 1996). Small increases in the seismic velocity above the BSR suggest that 5 7% of the pore volume is filled with hydrate. Such low volume fractions suggest that the BSR is produced mostly by the top of the gas layer. No discernable variations in the abundance of hydrate were evident through the stability zone because of the low volume fraction. However, in situ measurements of methane quantities provide some indication of the hydrate distribution (Dickens et al 1997b). The

14 490 BUFFETT recovered quantities of methane greatly exceed the solubility through most of the hydrate stability zone and are maximum near the base. Assuming that any gas in excess of the solubility is stored in hydrate, Dickens et al (1997b) estimate that as much as 9% of the pore space in the stability zone is filled with methane hydrate. Gas quantities recovered from below the stability zone suggest that bubbles of methane gas occupy 12% of the pore volume. Indeed, the gas trapped in bubbles below the BSR may exceed the gas stored in the overlying hydrate. Regional variations in the prominence of the BSR are reported in most study areas. Large abundances of hydrate often correlate with strong BSR amplitudes (Hobro et al 1998), although this correspondence does not hold everywhere. For example, evidence of gas hydrate was reported on Blake Ridge in a location where there was no indication of a BSR (Holbrook et al 1996). Variations in the BSR amplitude also appear to correlate with seafloor topography in many locations, where the strongest reflectors occur under topographic highs (Shipley et al 1979, Minshull et al 1994). In other cases the lateral variations are connected with tectonic features (Rowe & Gettrust 1993, Pecher et al 1998). Rowe & Gettrust (1993) reported vertical offsets in the BSR on the Blake Ridge that were clearly associated with faults. Because the BSR is expected to adjust to local thermodynamic conditions, the persistence of these offsets suggests that faulting has altered the temperature field, possibly by altering fluid circulation. Organic Geochemistry The molecular composition and isotopic signature of the gases recovered from hydrate samples or from hydrate-bearing sediment cores are consistent with gases generated by biological processes. In most locations the hydrocarbon gas is nearly pure methane, and the isotopic composition of the carbon (d 13 C) is usually lighter than 60 relative to the Peedee Belemnite standard (Claypool & Kvenvolden 1983). By comparison, thermal conversion of organic matter into hydrocarbon gases at temperatures in excess of 80 C produces larger quantities of ethane and propane, and the isotopic composition of the carbon is relatively heavy. Differences in the composition of the gas supply should be reflected in the composition of the hydrate. Although it is possible that methane gas is preferentially incorporated into the hydrate during crystallization (Thiery et al 1998), these effects are not likely to explain the abundance of methane gas (relative to ethane and propane) in most natural hydrates. In fact, occasional reports of hydrate samples containing gas mixtures with less than 60% methane (Ginsburg et al 1992) suggest that fractionation does not erase evidence of thermogenic gases when the source is clearly thermal. Thus the vast majority of clathrate hydrate near the Earth s surface is probably the result of biogenic conversion of organic matter into methane gas at shallow depths below the seafloor. Biological generation of methane gas occurs through a series of reactions. A typical reaction sequence produces one mole of methane from one mole of organic carbon. The amount of organic carbon in sediments is conventionally expressed

15 CLATHRATE HYDRATES 491 as a mass fraction of organic carbon in dry sediments. Sediments on the Blake Ridge contain about 1 1.5% organic carbon (Kvenvolden & Barnard 1983). Not all of this organic carbon is available to microorganisms. Some studies suggest that only one half of the organic carbon is converted to methane (Paull et al 1994), although lower efficiencies have also been reported (Claypool & Kvenvolden 1983). Paull et al (1994) estimate that 1% organic carbon yields 260 moles of CH 4 from 1 m 3 of sediment (assuming a porosity of 0.4 and an efficiency of 50%). If this amount of gas were dissolved in the pore fluid, it would yield a concentration of 0.6 M (or 0.01 mole fraction). Compared to the solubility shown in Figure 6, we observe that the available gas exceeds the solubility, but only by a factor of 3 4. Lower efficiencies reduce the available gas, and further losses occur through diffusion of gas toward the seafloor (see Figure 6). Even if the gas concentration exceeds the solubility by a factor of 3, this supply of gas is insufficient to fill more than a few percent of the pore volume with hydrate. To accumulate large volumes of hydrate, we require an influx of methane or organic matter into the stability region (Ginsburg & Soloviev 1997). A more quantitative discussion of the methane supply is deferred until the section on hydrate formation. Inorganic Geochemistry Pore fluids are routinely extracted from sediment cores during deep sea drilling. Measurements of chemical components and isotopic tracers provide a wealth of information about hydrate occurrences. The most commonly reported observation is the chlorinity of the pore fluid. Anomalously low measurements of chlorinity were first reported by Hesse & Harrison (1981), who attributed these observations to hydrate dissociation during the recovery of the sediments. Because the clathrate structure excludes salts during crystallization, the pore fluid becomes anomalously salty. However, the effects of diffusion and fluid flow can gradually restore the chlorinity concentrations to values more characteristic of seawater (approximately 544 mm). When hydrate dissociates in the sediments during recovery, the freshening is proportional to the in situ abundance of hydrate. Estimates of the hydrate abundances depend on the initial chlorinity concentration in the pore fluid prior to hydrate dissociation. Most studies adopt an average seawater chlorinity as a reasonable initial concentration throughout the sediments. However, deviations from seawater chlorinities can arise in several ways. Persistent formation of hydrate in the sediments can maintain anomalously high chlorinities in the stability zone, whereas low chlorinities occur below the stability zone if sedimentation causes dissociation of hydrate at the base of the stability zone (see next section). Factors unrelated to gas hydrate may also be important for interpreting chlorinity measurements. Migration of low chlorinity water from continents or from a deeper source owing to dewatering of clay minerals (Kastner et al 1995) can influence the observed chlorinities.

16 492 BUFFETT Profiles of chlorinity measurements from hydrate occurrences around the world are remarkably similar. One example from ODP Leg 146 on the Cascadia margin is shown in Figure 8a. Estimates of the inferred hydrate abundance are indicated in Figure 8b, assuming an initial chlorinity concentration of 544 mm. The fraction of the pore volume filled with hydrate increases from 0% at the seafloor to nearly 40% at the BSR. Lower hydrate volumes were inferred from seismic data at the same site, which may indicate an additional source of freshening besides hydrate dissociation (Kastner et al 1995). On the other hand, the differences are probably consistent with the uncertainties in the estimates. A similar level of agreement between geophysical and geochemical estimates of hydrate volume has been obtained at other sites (e.g. Brown et al 1996), but this is not always the case. Chlorinity measurements from ODP Leg 164 on the Blake Ridge suggest that hydrate occupies as much as 20% of the pore volume, whereas 5 7% volume fractions are estimated from VSP data by Holbrook et al (1996), and a 9% volume fraction is inferred by Dickens et al (1997b) from in situ quantities of methane gas. Paull et al (1998) questioned whether all of the freshening at sites on the Blake Ridge was due to hydrate. They fit low-degree polynomials to the observed chlorinity profiles, assuming that the average trends were due to an unknown source of relatively fresh water. Deviations from these smooth trends were attrib- Depth below seafloor (m) Chlorinity (mm) (a) Hydrate (% of pore volume) (b) Figure 8 (a) Measurements of pore fluid chlorinity at Sites 889/890 from ODP Leg 146. (b) Estimates of the hydrate volume inferred from the measured freshening compared with seawater chlorinity (544 mm).

17 CLATHRATE HYDRATES 493 uted to the presence of hydrate. This procedure had the desired effect of reconciling the estimates of hydrate abundance with other inferences, but it raised questions about the interpretation of chlorinity measurements if most of the signal is unrelated to the presence of hydrate. Two issues are important for interpreting the chlorinity measurements. First, the characteristic profile of chlorinity in Figure 8a is characteristic of regions where other indications of gas hydrate are evident (e.g. BSRs). Second, freshening below the BSR is inevitable when sedimentation or climate change causes hydrate to dissociate at the base of the stability zone. The typical timescale for chemical diffusion in the top few hundred meters of marine sediments is roughly 10 6 years, so dissociation of hydrate near the BSR should leave observable signatures in the chlorinity profiles. We show in the next section that the smooth trend observed on the Blake Ridge is probably a consequence of sedimentation. Several other geochemical indicators are suitable for detecting gas hydrate or constraining the formation environment. Measurements of the isotopic composition of oxygen in water (d 18 O relative to Standard Mean Ocean Water) provide another means of estimating the hydrate abundance. Fractionation of oxygen isotopes during hydrate formation increases d 18 O by 2.9% in the hydrated water and makes the pore water isotopically light (Kastner et al 1998). Mixing pore water with seawater or deeper fluids tends to restore the values of d 18 O to zero. Subsequent dissociation of hydrate in the sediment cores during recovery yields positive d 18 O values that may be interpreted as a measure of the hydrate volume, analogous to the use of chlorinity measurements. At drill sites in Guatemala (Hesse & Harrison 1981) and Peru (Kvenvolden & Kastner 1990) the profiles of d 18 O increase with depth, indicating a gradual increase in hydrate abundance below the seafloor. However, at sites on the Cascadia margin (Kastner et al 1995) and the Chile Triple Junction region (Brown et al 1996), the d 18 O profiles decrease with depth, presumably because of diagenetic reactions in the sediments. More recent studies of Paull et al (1998) and Kastner et al (1998) have exploited systematic variations of 87 Sr/ 86 Sr in seawater over the past few million years to quantify the extent of mixing in the pore fluid. Known variations of 87 Sr/ 86 Sr in seawater establish a timescale that may be used to estimate the average age of pore water in the sediments (Farrell et al 1995). Departures from the age of the host sediments indicate mixing in the pore fluid. The ages inferred for the pore fluids on the Blake Ridge are in good agreement with the age of host sediments below the BSR, whereas more modern isotopic signatures were recorded inside the stability region (Paull et al 1998). Gradients in the strontium concentration inside the stability region imply downward diffusion of strontium into the sediments, and may explain the more modern isotopic signatures. This suggests that diffusion is sufficient to carry strontium into the sediments and argues against a significant upward transport of fluid from below. On the other hand, Egeberg & Dickens (1999) argue for an upward vertical flow of 0.2 mm year 1 based on profiles of Br and I concentrations in the pore fluids.

18 494 BUFFETT Further constraints on fluid transport are derived from profiles of sulfate concentration in the top 10 or 20 m of sediments. Depletion of sulfate is caused by sulfate-reducing microorganisms and/or by anaerobic oxidation of methane. The latter case prevails when methane from below is brought into contact with sulfate near the seafloor. The reaction 2 SO4 CH4 HS HCO3 HO 2 consumes one mole of sulfate for each for each mole of methane. Concentrations 2 of both SO 4 and CH 4 vanish at the reaction front in the shallow sediments, which causes concentration gradients in the surrounding sediments. Diffusion brings sulfate and methane into the reaction front at equal rates. Borowski et al (1996) argues for anaerobic oxidation of methane as the principal cause of sulfate depletion on the Blake Ridge. Measurements of gradients in the sulfate concentration establish the diffusive flux of sulfate into the floor and provide a proxy for the methane flux from below. The inferred methane flux varies considerably across the study area, but peak values were mmol cm 2 year 1, corresponding to a methane gradient of 2 mm m 1. We show in the next section that such steep gradients in the methane concentration are not easily maintained without fluid flow from below. Variability of the methane flux suggests that fluid flow is not uniform across the area of study. FORMATION IN THE SEAFLOOR The amount of methane gas sequestered in marine gas hydrate often exceeds the gas that is available from direct biogenic conversion of organic material. Accumulations of hydrate in excess of a few percent of the pore volume require either an influx of methane or new organic material. Indeed, the total supply of methane must be sufficient to overcome persistent losses of methane due to diffusion along the concentration gradient imposed by thermodynamic equilibrium (see Figure 6). Without a continual supply of methane, losses of gas hydrate become significant over 10 6 years (Zatsepina & Buffett 1998). Differences between active and passive margins have motivated two different models for hydrate accumulation. The model proposed by Hyndman & Davis (1992) for active margins relies on fluid expulsion from the accretionary prism (Carson & Screaton 1998). Large volumes of hydrate accumulate by focusing dissolved methane into the stability region from an extensive source region across the prism. As methane-bearing fluid migrates into the stability region, the drop in solubility toward the seafloor causes methane to exsolve from solution and crystallize into gas hydrate. The rate at which hydrate forms depends on the transport of methane into the stability zone and on the decrease in solubility. Quantitative predictions for the distribution of hydrate and its rate of accumulation have been obtained by assuming a pervasive fluid flow through a uniform porous medium (Rempel & Buffett 1997, 1998; Xu & Ruppel 1999). Results from the

19 CLATHRATE HYDRATES 495 Depth below seafloor (m) calculation of Rempel & Buffett (1998) are shown in Figure 9. The calculated distribution of hydrate is similar to that inferred from borehole measurements on Cascadia margin (see Figure 7b). For a nominal velocity of 1 mm year 1,we expect roughly 1% of the pore volume to fill with hydrate in 10 5 years. At this rate of accumulation, hydrate volumes of 20 30% on the Cascadia margin could develop over several million years. Dissolved concentrations of methane gas are constrained by the local solubility anywhere hydrate or gas bubbles are present. In Figure 9 we assume that the gas concentration at the base of the hydrate stability zone is maintained by gas bubbles. Otherwise, the base of the hydrate zone can shift to a shallower position in the sediments (Zatsepina & Buffett 1997, Xu & Ruppel 1999). Hydrate does not extend to the seafloor because the gas concentration is less than the solubility in the shallowest sediments. A steep gradient in the methane concentration develops as the concentration adjusts to the low value imposed at the seafloor. The gradient in the methane concentration near the seafloor is controlled by a competition between fluid transport and diffusion of dissolve gas. Larger fluid velocities yield steeper methane gradients at the seafloor. The methane gradient achieved in Figure 9 using a velocity of 1 mm year 1 is more than sufficient to explain the largest sulfate gradients observed by Borowski et al (1996) on the Blake Ridge. In fact, estimates of the methane fluxes inferred from profiles of sulfate concentrations can provide valuable information about the vertical fluid flow. This information has not been widely exploited in previous studies fluid velocity = 1 mm yr t = 100 kyr t = 200 kyr (a) Hydrate volume fraction (%) Depth below seafloor (m) solubility (b) Gas concentration X / X o Figure 9 (a) Accumulation of the hydrate predicted using a vertical fluid velocity of 1 mm year 1. The hydrate volume is expressed as a percentage of the pore volume, assuming a porosity of 0.5. (b) The profile of dissolved gas concentration X in the stability zone, normalized by the solubility X 0 immediately below the stability zone. We assume that gas bubbles immediately below the hydrate stability zone maintain the concentration of dissolved gas at X 0 at a depth of 200 mbsf.