14 Gas Hydrates in Marine Sediments
502
higher than that of seawater. Most commonly
these brines are associated with regions where
the presence of old evaporites (e.g. Milano Dome,
ODP Site 970 in the eastern Mediterranean,
DeLange and Brumsack 1998), or salt-diapir
intrusions (e.g. Blake Ridge Diapir ODP Site 996,
Egeberg and Dickens 1999; mud volcanoes in the
Northern Gulf of Mexico, Ruppel et al. 2005) leads
to the enhanced chloride content. In addition to
these settings, high dissolved chloride concentration associated with hydration reactions in the
vicinity of an active spreading ridge was reported
from ODP Sites 859 and 860 in the accretionary
wedge at the Chile Triple Junction (Froelich et al.
1995).
In contrast to these regions in which brines
are produced by geological processes, at the
Hydrate Ridge summit, the high chloride brines
observed (Haeckel et al. 2004; Torres et al. 2004)
are generated by the rapid formation of gas
hydrate deposits near the seafloor (Fig. 14.18). A
one-dimensional transport-reaction model was
used to simulate this chloride enrichment and
place constrains on the mechanisms and time
frames necessary to produce the observed
concomitant massive hydrate deposition at the
ridge summit. The models of Torres et al. (2004)
and Haeckel et al. (2004) demonstrate the need for
the presence of a fluid-gas mixture through the
GHSZ, since the observed chloride enrichment
cannot be generated exclusively from the
transport of methane dissolved in the pore fluids.
These massive hydrate deposits are forming very
rapidly, and the continuous supply of methane
gas maintains the pore water brines and the
shallow gas hydrate deposits in contact with the
methane-poor bottom seawater.
14.4.2 Gas Hydrate and Water Isotope
Anomalies
The water sequestered in the hydrate lattice is
preferentially enriched in
18
O and deuterium (D), thus
the isotopic composition of the water in the pore
spaces collected from gas hydrate bearing sediment
can provide additional information on the abundance
and the characteristics of these deposits. Pore fluid
samples that had been modified by hydrate
decomposition upon core recovery during ODP Legs
146 (Kastner et al. 1998), and 164 (Matsumoto and
Borowski 2000) provided the first field data to derive
the oxygen isotope fractionation factor for in situ
hydrate formation. A more comprehensive sampling
protocol was subsequently conducted during Leg 204
(Tomaru et al. submitted). These calculations are based
on the percent variation of Cl
-
relative to background
(∆Cl
-
):
100
)
1
(
×
−
=
∆
−
f
Cl
(3)
where f is a fraction of formation water in sampled
water given by:
0
−
−
= Cl
Cl
f
S
(4)
Cl
-
S
and Cl
-
0
are the Cl
-
concentrations of sampled and
formation water (i.e., in situ interstitial water)
determined as background, respectively. The
fractionation factors for oxygen (α O ) and hydrogen
(α H ) can be determined from equilibrium equation, such
that:
)
1
(
ln
1000
0
f
GH
−
⋅
⋅
=
−
=
∆
α
δ
δ
δ
(5)
where α GH and α 0 are is δ
18
O or δD values for gas
hydrate and formation (background) water.
Figure 14.19 illustrates how the fractionation of
18
O to
16
O and H to D between pore water, and water
derived from hydrate dissociation is related to the
fractionation under in situ conditions, assuming a
closed system. The average values of α O and α H from
Leg 204 samples with negative ∆Cl
-
are calculated to
be 1.0025 and 1.022. These fractionation factors agree
with previously estimated α O values from Leg 146
(Kastner et al. 1998) and Leg 164 (Matsumoto and
Borowski 2000), and with the extrapolated α H value
from Leg 112 (Kvenvolden and Kastner 1990).
Figure 14.19 illustrates the fractionation factors for
in situ hydrate formation that correspond to
experimentally obtained values for oxygen (α ο : 1.0023
to 1.0032) and for hydrogen (α H :1.014 to 1.022)
(Maekawa 2004). Analyses of pore water samples from
a pore water brine sampled during Leg 204 reveal the
oxygen and hydrogen isotopic fractionation during
hydrate formation in natural systems. There are special
challenges in fully constraining these values, since
the dissolved chloride data from these brines reflects
a mixture of the in situ fluids, with an unknown amount
of fresh water added by hydrate dissociation during
sample recovery. Nevertheless, Tomaru et al. (submitted) show that the isotopic fractionation in these
massive deposits departs significantly from experimental data. More research is needed to fully understand these deviations.
502
higher than that of seawater. Most commonly
these brines are associated with regions where
the presence of old evaporites (e.g. Milano Dome,
ODP Site 970 in the eastern Mediterranean,
DeLange and Brumsack 1998), or salt-diapir
intrusions (e.g. Blake Ridge Diapir ODP Site 996,
Egeberg and Dickens 1999; mud volcanoes in the
Northern Gulf of Mexico, Ruppel et al. 2005) leads
to the enhanced chloride content. In addition to
these settings, high dissolved chloride concentration associated with hydration reactions in the
vicinity of an active spreading ridge was reported
from ODP Sites 859 and 860 in the accretionary
wedge at the Chile Triple Junction (Froelich et al.
1995).
In contrast to these regions in which brines
are produced by geological processes, at the
Hydrate Ridge summit, the high chloride brines
observed (Haeckel et al. 2004; Torres et al. 2004)
are generated by the rapid formation of gas
hydrate deposits near the seafloor (Fig. 14.18). A
one-dimensional transport-reaction model was
used to simulate this chloride enrichment and
place constrains on the mechanisms and time
frames necessary to produce the observed
concomitant massive hydrate deposition at the
ridge summit. The models of Torres et al. (2004)
and Haeckel et al. (2004) demonstrate the need for
the presence of a fluid-gas mixture through the
GHSZ, since the observed chloride enrichment
cannot be generated exclusively from the
transport of methane dissolved in the pore fluids.
These massive hydrate deposits are forming very
rapidly, and the continuous supply of methane
gas maintains the pore water brines and the
shallow gas hydrate deposits in contact with the
methane-poor bottom seawater.
14.4.2 Gas Hydrate and Water Isotope
Anomalies
The water sequestered in the hydrate lattice is
preferentially enriched in
18
O and deuterium (D), thus
the isotopic composition of the water in the pore
spaces collected from gas hydrate bearing sediment
can provide additional information on the abundance
and the characteristics of these deposits. Pore fluid
samples that had been modified by hydrate
decomposition upon core recovery during ODP Legs
146 (Kastner et al. 1998), and 164 (Matsumoto and
Borowski 2000) provided the first field data to derive
the oxygen isotope fractionation factor for in situ
hydrate formation. A more comprehensive sampling
protocol was subsequently conducted during Leg 204
(Tomaru et al. submitted). These calculations are based
on the percent variation of Cl
-
relative to background
(∆Cl
-
):
100
)
1
(
×
−
=
∆
−
f
Cl
(3)
where f is a fraction of formation water in sampled
water given by:
0
−
−
= Cl
Cl
f
S
(4)
Cl
-
S
and Cl
-
0
are the Cl
-
concentrations of sampled and
formation water (i.e., in situ interstitial water)
determined as background, respectively. The
fractionation factors for oxygen (α O ) and hydrogen
(α H ) can be determined from equilibrium equation, such
that:
)
1
(
ln
1000
0
f
GH
−
⋅
⋅
=
−
=
∆
α
δ
δ
δ
(5)
where α GH and α 0 are is δ
18
O or δD values for gas
hydrate and formation (background) water.
Figure 14.19 illustrates how the fractionation of
18
O to
16
O and H to D between pore water, and water
derived from hydrate dissociation is related to the
fractionation under in situ conditions, assuming a
closed system. The average values of α O and α H from
Leg 204 samples with negative ∆Cl
-
are calculated to
be 1.0025 and 1.022. These fractionation factors agree
with previously estimated α O values from Leg 146
(Kastner et al. 1998) and Leg 164 (Matsumoto and
Borowski 2000), and with the extrapolated α H value
from Leg 112 (Kvenvolden and Kastner 1990).
Figure 14.19 illustrates the fractionation factors for
in situ hydrate formation that correspond to
experimentally obtained values for oxygen (α ο : 1.0023
to 1.0032) and for hydrogen (α H :1.014 to 1.022)
(Maekawa 2004). Analyses of pore water samples from
a pore water brine sampled during Leg 204 reveal the
oxygen and hydrogen isotopic fractionation during
hydrate formation in natural systems. There are special
challenges in fully constraining these values, since
the dissolved chloride data from these brines reflects
a mixture of the in situ fluids, with an unknown amount
of fresh water added by hydrate dissociation during
sample recovery. Nevertheless, Tomaru et al. (submitted) show that the isotopic fractionation in these
massive deposits departs significantly from experimental data. More research is needed to fully understand these deviations.