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What are the key findings and results, important concepts and background context of the research in 'Kita et al 1980 - D-H ratios of H2 in soil gases as an indicator of fault movements'?
|
Based on the research paper 'Kita et al 1980 - D-H ratios of H2 in soil gases as an indicator of fault movements', the key findings and results include (JAPAN) Kita et al 1980 - D-H ratios of H2 in soil gases as an indicator of fault movements.txt
Geochemical Journal, Vol. 14, pp. 317 to 320, 1980317
D/H ratios of H2 in soil gases as an indicator of
fault movements
ITSURO KITA,I SADAO MATSUO,r HIROSHI WAKITA' and Yuji NAKAMURA'
Department of Chemistry, Faculty of Science,
Tokyo Institute of Technology, O-okayama, Meguro-ku, Tokyo 152'
and Laboratory for Earthquake Chemistry, Faculty of Science,
University of Tokyo, Hongo 7-3-1, Bunkyo-ku, Tokyo 113?Japan
/Received December 8, 1980; Accepted January 22, 1981.... The background context for this study involves the motivation and circumstances that led to this investigation.
|
(JAPAN) Kita et al 1980 - D-H ratios of H2 in soil gases as an indicator of fault movements.txt
Geochemical Journal, Vol. 14, pp. 317 to 320, 1980317
D/H ratios of H2 in soil gases as an indicator of
fault movements
ITSURO KITA,I SADAO MATSUO,r HIROSHI WAKITA' and Yuji NAKAMURA'
Department of Chemistry, Faculty of Science,
Tokyo Institute of Technology, O-okayama, Meguro-ku, Tokyo 152'
and Laboratory for Earthquake Chemistry, Faculty of Science,
University of Tokyo, Hongo 7-3-1, Bunkyo-ku, Tokyo 113?Japan
/Received December 8, 1980; Accepted January 22, 1981)
The SD values of molecular hydrogen in soil gases obtained from sites along the Yamasaki fault,
one of active faults in Southwest Japan were measured. The SD values were in the range between -770
and -470°/00, which suggests that molecular hydrogen is formed by chemical reactions between ground
water and the fresh surface of basement rocks formed by the fault movement.
INTRODUCTION locations of sampling sites are shown in Fig. 1.
The H2 concentration in soil gases from sites
The soil near the ground surface contains away from the fault zone was about 0.5ppm
gases. The chemical compositions of these gases and was almost the same as that found in the
are mainly atmospheric, but small portion of atmosphere. The spatial distribution of sites
the gases originated from great depths in the with high H, concentrations is quite systematic
earth may also be involved. Through an in (WAKITA et al., 1980).
vestigation of these gases, information about This study reports the first data on D/H
processes occurring at depths and the tectonic ratios of molecular H, in soil gases in relation
structure of the area may be obtained. Degas to the origin of H2,
sing from deep levels of the earth is more
pronounced in active fault zones than in regions EXPERIMENTAL
where tectonic activity is low. Measurements
of the concentrations and fluxes of gases from A series of field experiments was carried out
depths are useful for the estimation of the fault in the area around the Yamasaki fault in Hyogo
activity. Hydrogen and helium are inferred to Prefecture, Japan, beginning in November, 1978.
be the most ideal elements for this purpose Shallow holes 2.5 and 5.0cm in diameter and
because of their low abundance in the atmo of 50-100 cm depth were drilled with either an
sphere, high mobility and chemical inertness. electric power drill or a hand auger. The soil
Isotopic ratios of these elements are particularly gas samples for isotopic measurement were col
useful in elucidating the origin of the gases. lected at site C along the fault zone. Site C
For example, migration of He from the upper is one of the most representative sites where
mantle has been demonstrated on the basis of significant amounts of H, were found (WAKITA
the isotopic ratio of 'He/'He (CRAIG et at, et al., 1980).
1975; TOLSTIKHIN, 1978). Evacuated vessels (120m1) equipped with a
It was reported that a significant amount stopcock were used for gas sampling. Hydrogen
of H,, up to more than 3% v/v, has been ob in the soil gases was separated from other gases
served in the soil gases from sites along the such as N,, O,, At, CO, and water vapor for
Yamasaki fault. The Yamasaki fault and the isotopic measurement. H, was separated and
318 1. KITA et al
Yemasekl-3
3-R Pt
_,/ -2
-3
yam
N,M-f
mil
~ F
uY
K a KU02Table 1.SD values of H2 gas and concentrations of H2,
CH4 and CO, in the soil gases at site C
Site C-1 C-2 C-1-9 C-2-2
Date
SD (%o)
H2 (%)
CH4 (PPM)
CO2 (%)(1)
-470
1.0
<150
0.7(1)
-510
1.0
<150
0.6(2)
-770
3.0
800
0.9(2)
-590
0.7
n.d.
Fig. 1. Locations of the sampling sites of soil gases in
the area around the Yamasaki fault. Closed circles
indicate the sites where significant amounts of H2 were
measured. Open circles indicate the sites where H2 con
centrations were almost the same as that of the
atmosphere.
purified by the following method. First, water
vapor and CO, were separated from the other
gases by immersing the sampling reservoir into
liquid N2 for more than 6 hours. Hydrogen
was then separated from the remaining gases by
being converted to H2O through CuO-Cu2O at
450°C and H2 was separated by circulating the
gases for more than 9 hours in a closed system.
After trapping H2O thus formed completely, the
noncondensable gases were pumped off. The
H2O was then converted again to H2 by reaction
with uranium metal at 750°C and the yield was
measured manometrically. Finally, the D/H
ratio measurements were made on this H2 using
a mass spectrometer with a dual inlet and col
lector system (Hitachi RMD).
RESULTS AND DtscussioN
The SD values for H2 in soil gases collected
from the Yamasaki fault are -470%0 (C-1) and
-510%o(C-2) in November, 1978 and -590%,
(C-2-2) and -770%o(C-1-9) in March, 1979.
The SD value is presented relative to SMOW as
follows,
SD(°/oo) _ (D/H) sample (D/H)SMow X 103. (D/H)sMowl1) November 14, 1978.
(2) March 21, 1979.
These results are shown together with the
concentrations of CH4 and CO, in the soil gases
in Table 1. The SD value of H2 in soil gases
has not been previously determined. Observed
SD values are much lower than those reported
for various hydrogen bearing materials on the
earth. No measurement was made of soil gases
from sites away from the fault zone because of
extremely low abundance of H, (WAKiTA et al.,
1980).
Most hydrogen on the earth exists in the
chemical forms H2O, OH bearing minerals, or
ganic materials, CH4 and H2. The SD values for
terrestrial water and OH-minerals have been
extensively measured and are generally in the
range from -350 to +50%0. The lowest values
of -460 and -420%o were reported for the
Antarctic ice (EPSTEIN and SHARP, 1965) and
pectolite, a hydrous silicate (KURODA et at,
1979), respectively. The mean SD value of
atmospheric methane is -103%0, while those of
CH4 from natural gases and geothermal sources
are -188 and -261 %o, respectively (BEGEMANN
and FRIEDMAN, 1968). The SD value of molec
ular H2 has seldom been reported. The SD
values of atmospheric H2 is about +70%o with
the heaviest value of +180%0 (FRIEDMAN and
ScHOLz, 1974). The SD values of H2 in volcanic
gases have been reported to be in the range
between -158 and -144%0 (ARNASON and
SIGURGEIRSSON, 1968). The SD values of H2 in
volcanic gases were inferred to be the result of
isotopic exchange equilibrium between H2 and
water at 800 to 1,100°C. From the study of
the hydrothermal area, a value of about -660%o
has been obtained for H2 in fumarolic gases
D/H ratios of H2 in soilgases 319
from Yellowstone Park (GUNTER and MUSGRAVE,
1971). These gases contain significant amounts
of CO2 (83%) and CH, (14%) and the low value
of -660%o for H,, was interpreted to arise
from isotopic exchange equilibrium between
CH4 and H2. It is to be noted that the soil
gases measured in this study were not collected
in either a hydrothermal or volcanic area but
rather in areas along a fault zone.
Molecular H2 can be generated in soil by
biological activity. In this case, larger amounts
of CH, and CO, than that of H2 are generally
released (ALEXANDER, 1961). It has been known
that some marine bacteria produce H2 highly
depleted in deuterium in cultivating experi
ments. The 8D values of this H2 range between
-814 and -763%o (KRICHEVSKY et al., 1961),
in the SMOW scale. Concentrations of CH, and
CO, in the soil gases we analyzed are quite low
as shown in Table 1. Some amount of soil was col
lected from the site, where a significant amount
of H2 is released and brought to the laboratory.
The soil sample was divided into two aliquots
in flasks. One sample was stored in a room at
25°C for four months (118 days), and the other
was kept in a dark room at 15°C for the same
period. It was found that H2 concentrations
in the gas phase of these flasks were the same
as that of the normal air (0.5ppm). There was
no remarkable difference in both cases with
respect to gas composition ; O, decreased
to the level lower than 0.6%, N2 and CO, in
creased up to 86 and 6% respectively and CH4
concentration was 13ppm at the end of storage.
On the basis of the above findings and field
observations we may conclude that bacterial
production of H2 is negligible in this particular
case.
It is also probable that H2 gas originated
from the upper mantle is released through the
fault. In this case, H, gas is expected to be
generated by the decomposition of H2O at high
temperature (?1,000°C) and the D/H ratio of
H2 gas must be quite close to that of the source
H2O because of the high temperature. The
SD value of mantle water fixed in hydrous
silicates has been estimated to be around-100%o (MATsuo et al., 1978). The actually
measured SD values of H, in soil gases are too
low to conclude that a significant portion of
H, of mantle origin is admixed.
A hypothesis for the origin of H2 in soil
gases along the Yamasaki fault has been pro
posed by WAKiTA et al. (1980). According to
this hypothesis, H2 is produced by chemical
reactions between groundwater and fresh rock
surface of basement rocks formed by the fault
movement.
The reaction responsible for H2 production
is basically the reduction reaction of water on
surface of the crushed rocks. Even when a
kinetic isotope effect is involved, the hydrogen
isotopic fractionation factor between H,O and
H, is a function of the reaction temperature.
Isotopic fractionation factor(a) depends on
the temperature of reaction. Fractionation fac
tor is defined as follows,
a = (D/H)e2o /(D/H)H 2
=(1 + 10-'SDH2o)/(1 + 10-3SDH2).
Isotopic fractionation factors between H2O
and H2 have been reported with the temperature
dependence by SUESS (1949) and BOTTINGA
(1969). The temperature dependence of frac
tionation factor was expressed by SUESS as,
1031n a = 467.6 X (10'T-') 303.9.
On the basis of the temperature dependence of
fractionation factor reported by SuESS, an at
tempt to infer the reaction temperature was
made using the 6D values of H2 gas found in
sites along the Yamasaki fault.
SDH,O of groundwater in the Yamasaki
region is about -50%o. On the other hand,
SDH2 values found are -470%o (C-1, 1978)
-510%o (C-2, 1978), -590%, (C-2-2, 1979)
and -770%o (C-1-9, 1979). Hence, a can be
respectively calculated to be 1.79, 1.94, 2.32
and 4.13.
On the basis of the relationship between
fractionation factor and the temperature re
320 I. KITA et al.
ported by SuEss (1949), the temperatures of
apparent isotopic exchange equilibrium were
calculated to be 255, 210, 135°C and about
room temperature for 1.79, 1.94, 2.32 and 4.13,
respectively.
Assuming the lapse rate of underground tem
perature to be 3°C/100m, we may obtain the
depths at which the reactions took place to be
8km (C-1), 6km (C-2) in November, 1978 and
4km (C-2-2) and near ground surface (C-1-9) in
March, 1979. It has been reported that the foci
of microearthquakes were concentrated in the
depths of a few km to 20km with the average
of 12km (DIKE, 1977).
If the lapse rate of underground temperature
is assumed to be 2°C/100m, we can obtain a
better agreement between the estimated depths
and the foci of microearthquakes (12km(C-1)
and 9km(C-2) in November, 1978 and 6km(C
2-2) and near ground surface (C-1-9) in March,
1979).
In general, kinetic isotope fractionation
factor for the reaction, H2O -> H2 is larger than
the isotopic fractionation factor under equi
librium condition. Therefore, the depth esti
mated on the basis of kinetic fractionation
factor may be deeper than that estimated by the
equilibrium fractionation factor. The depth
estimated by the equilibrium fractionation fac
tor may be regarded as the shallowest values
for the depth where the reaction takes place.
Although the number of data is not yet suf
ficient at the present time to give a quantitative
correlation between SD values and seismicity
(fault movement), we believe that SD values of
H2 in soil gases may be a sensitive indicator of
the depth where the reaction between ground
water and fresh surface of crushed rock takes
place.
Acknowledgements-We thank Dr. N. FUIII of Kobe
University and Dr. K. NOTSU of Tsukuba University for
their help in sample collection and for extensive discus
sions. This work was supported in part by a grant for
the Yamasaki Fault Test Site Project for Earthquake
Prediction Research of the Ministry of Education, Sci
ence and Culture, Japan. REFERENCES
ALEXANDER, M. (1961) Introduction to soil micro
biology, John Wiley and Sons. Inc., New York, 159p.
ARNASON, B. and SIGURGEIRSSON, T. (1968) Deu
terium content of water vapour and hydrogen in vol
canic gas at Surtsey, Iceland. Geochim. Cosmochim.
Acta 32, 807-813.
BEGEMANN, F. and FRIEDMAN, I. (1968) Tritium
and deuterium in atmospheric methane. J. Geophys.
Res. 73, 1149-1153.
BOTTINGA, Y. (1969) Calculated fractionation fac
tors for carbon and hydrogen isotope exchange in
the system calcite-carbon dioxide graphite methane
hydrogen water vapor. Geochim. Cosmochim. Acta
33, 49-64.
CRAIG, H., CLARKE, W. B. and BEG, M. A. (1975)
Excess 'He in deep water on the east Pacific rise.
Earth Planet. Sci Lett. 26, 125-132.
EPSTEIN, S. and SHARP, R. S. (1965) Six-year record
of oxygen and hydrogen isotope variations in South
Pole Fim. J. Geophys Res. 70, 1809-1814.
FRIEDMAN, I. and SCHOLZ, T. G. (1974) Isotopic
composition of atmospheric hydrogen. J. Geophys.
Res. 79, 785-788.
GUNTER, B. D. and MUSGRAVE, B. C. (1971) New
evidence on the origin of methane in hydrothermal
gases. Geochim. Cosmochim. Acta 35, 113-118.
KRICHEVSKY, M. I., FRIEDMAN, I., NEWELL, M. F. and
SISLER, F. D. (1961) Deuterium fractionation
during molecular hydrogen formation in a marine
pseudomonad. J. Biol. Chem. 236, 2520-2525.
KURODA, Y., SUZUOKI, T. and MATSUO, S. (1979)
The lowest SD value found in a hydrous silicate,
pectolite. Nature 279, 227-228.
MATSUO, S., KURODA, Y., SUZUOKI, T., AOKI, K. and
HARIYA, Y. (1978) Mantle water based on the
hydrogen isotopic ratios of hydrous silicates in the
mantle. Short Papers of the Fourth International
Conference, Geochronology, Cosmochronology, Iso
tope Geology. 278-280.
OIKE, K. (1977) Seismic activities and crustal move
ments at the Yamasaki fault and surrounding regions
in the Southwest Japan. J. Phys. Earth, 25, Suppl.,
S31-S41.
SUESS, H. E. (1949) Das Gleichgewicht H2 + HDO
'.HD+H2O and die weiteren Austauschgleichge. wichte im System H2, D2 and H2O. Z. Naturforsch.
4,328-332.
TOLSTIKHIN, 1. N. (1978) Terrestrial rare gases, ed.
E. C ALEXANDER, JR. and M. OZIMA (Japan Sci
entific Societies Press, Tokyo), 33-62.
WAKITA, H., NAKAMURA, Y., KITA, I., FUJII, N. and
NOTSU, K. (1980) Hydrogen release: New indicator
of fault activity. Science 210, 188-190.
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What are the key findings and results, important concepts and background context of the research in 'Kiyosu 1982 - d2H-H2 and d2H-CH4 from volcanic areas in northeastern Japan'?
|
Based on the research paper 'Kiyosu 1982 - d2H-H2 and d2H-CH4 from volcanic areas in northeastern Japan', the key findings and results include (JAPAN) Kiyosu 1982 - d2H-H2 and d2H-CH4 from volcanic areas in northeastern Japan.txt
Earth and Planetary Science Letters, 62 (1983) 41-52 41
Elsevier Scientific Publishing Company, Amsterdam - Printed in The Netherlands
[31
Hydrogen isotopic compositions of hydrogen and methane from
some volcanic areas in northeastern Japan
Yasuhiro Kiyosu
Department of Earth Sciences, Faculty of Science, Nagoya University, Chikusa, Nagoya 464 (Japan)
Received April 21, 1982
Revised version received September 14, 1982
D/H ratios of volcanic condensates, hydrogen and methane collected from Hachimanta.... The background context for this study involves the motivation and circumstances that led to this investigation.
|
(JAPAN) Kiyosu 1982 - d2H-H2 and d2H-CH4 from volcanic areas in northeastern Japan.txt
Earth and Planetary Science Letters, 62 (1983) 41-52 41
Elsevier Scientific Publishing Company, Amsterdam - Printed in The Netherlands
[31
Hydrogen isotopic compositions of hydrogen and methane from
some volcanic areas in northeastern Japan
Yasuhiro Kiyosu
Department of Earth Sciences, Faculty of Science, Nagoya University, Chikusa, Nagoya 464 (Japan)
Received April 21, 1982
Revised version received September 14, 1982
D/H ratios of volcanic condensates, hydrogen and methane collected from Hachimantai, Za~, Azuma, Nasu and
Kusatsu shirane volcanic areas in northeastern Japan were determined together with chemical components. On the basis
of the isotopic ratios for volcanic condensates, it is concluded that most of the water vapor is essentially local surface
water which has been heated and slightly enriched in 180 by exchange with silicates. 8D values of hydrogen and
methane range from -220 to -515%0 (SMOW) and from - 180 to -487%0 (SMOW), respectively. The D/H ratio in
volcanic hydrogen indicates that hydrogen may have isotopically equilibrated with water vapor at temperatures between
200 and 400°C. The 8D value of methane also suggests that volcanic methane has been in isotopic disequilibrium with
the water vapor and hydrogen and that most of the methanes may be thermogenic methanes produced by pyrolysis of
carbonaceous materials.
1. Introduction
Volcanic gases usually contain alkali-unab-
sorbed gases such as nitrogen, hydrogen, methane
and inert gases.
High-temperature (> 400°C) volcanic gas of
magmatic origin, which has not passed through a
groundwater body generally has high concentra-
tions of hydrogen. Presumably, volcanic hydrogen
is produced by the high-temperature reaction of
water with ferrous oxides and silicates. Whereas
data on hydrogen and oxygen isotopes in volcanic
condensates have been reported, only little is
known about the deuterium concentration of
volcanic hydrogen. Arnason and Sigurgeirsson [1]
analyzed the deuterium contents of the water vapor
and hydrogen from the oceanic volcano Surtsey
near Iceland and showed that total hydrogen con-
tained in the volcanic gases has 8D value of
-55.3~ vs. SMOW. Arnason [2] estimated the
bottom temperature of drill holes from the ob-
served D/H fractionation between hydrogen and water in geothermal areas of Iceland; the calcu-
lated temperatures were generally equal to or
greater than the observed temperatures.
The gas boiling from heated groundwater (<
400°C) in geothermal systems often contains more
methane than hydrogen. The concentration of
methane may be controlled by the chemical reac-
tion of carbon dioxide and hydrogen [3]. This
reaction equilibrium shifts to the side of enrich-
ment of methane as temperature decreases and
pressure increases. The methane observed in
volcanic gases may be created by this reaction,
and/or by the near-surface reaction of volcanic
gases and heat with organic material [4]. Very few
data have been reported on the variation of D/H
ratios in volcanic methane because this gas is
relatively uncommon as compared to components
like hydrogen.
In order to clarify the origins of volcanic hydro-
gen and methane, the present isotopic investiga-
tions were conducted. In some volcanic areas, an
attempt was also made to compare measured tem-
0012-821/83/0000-0000/$03.00 © 1983 Elsevier Scientific Publishing Company
42
peratures and isotopic temperatures calculated
from the isotopic fractionation between water
vapor and hydrogen.
2. Sample localities
Gas samples were collected from fumaroles of
Za~, Azuma (Issaikyo), Nasu and Kusatsu shirane
in 1980 to 1981, and from six dry-steam wells in
the Matsukawa geothermal area (Hachimantai) in
1977 and 1981. The sampling localities are shown
in Fig. 1.
Za~, ,4zuma and Nasu volcanoes
Za~, Azuma (Issaikyo) and Nasu are strato-
volcanos made of pyroxene andesite, belonging to
the northern subzone of the Nasu volcanic zone
[5]. The volcanic rocks overlie the Miocene base-
ment of green tuff, volcanics and sediments, and
pre-Tertiary granodiorite. The present activity
Hachimantai
(usatsu
e
0 1 O0 200 Krn
i i I
Fig. 1. Sampling localities. consists of acid hot springs and fumaroles con-
centrated around these volcanic areas. The Nasu
samples discussed here are gas samples collected
from Oana and Mugen fumaroles in this area.
Kusatsu shirane volcano
Kusatsu shirane, located at the intersection of
the extension of Nasu and Fuji volcanic zones, is a
stratovolcano built from pyroxene andesite [6,7].
The most recent activity was a minor steam erup-
tion which occurred around the pit of Mizugama
at the summit crater, in March 1976. There are
four principal fumaroles and acid hot spring areas:
Kusatsu shirane (the northern foot of the volcano),
Sesshogawara, Kusatsu yubatake and Manza
karabuki.
Matsukawa geothermal fieM
Matsukawa is northwest of Mt. Iwate in the
Hachimantai volcanic groups which belongs to the
Nasu volcanic zone. Natural activity at Matsukawa
consists only of hot springs. The area is one of
andesitic volcanism, and the volcanic activity is
divided into the Quaternary volcanic complex and
the lower Pleistocene Tamagawa welded tuff com-
plex, with the dacitic or andesitic tuff. Underlying
the Tamagawa complex is the Miocene Yamatsuda
formation, consisting of sandstone, siltstone and
tuff. The Tamagawa welded tuff is the main res-
ervoir for production wells. The hydrothermal al-
teration of the andesite has several zones with
increasing depth. At depth there are radially dis-
tributed zones of alteration consisting of chlorite,
montmorillonite, kaolinite and alunite. A relic al-
teration pattern of pyrophyllite is recognized in
the deeper wells [8-10].
3. Experimental
3.1. Sample collection
The sampling method was modified after that
of Mizutani [11] and Ozawa [12] for chemical and
isotopic analyses, as follows: Volcanic and geo-
thermal gas samples were collected in a water-
cooled KOH solution by introducing the gases
from a fumarole through a glass tube and from a
steam vent on the steam pipeline through rubber
tubing, respectively. The hydrogen, methane and
other alkali-unabsorbed gases were collected in a
pyrex gas collector with a volume of about 100 ml.
When the alkali-unabsorbed gases were collected
to a sufficient amount, both ends of the collector
were sealed off with a torch. The condensates, hot
spring waters and river waters were sampled for
isotopic analysis. The outlet temperature of the
fumarole was measured with a mercury thermome-
ter.
3.2. Analytical methods
Chemical composition of gas condensates in
alkaline solutions was analyzed by the method of
Ozawa [12], except for H 2 S determination. Carbon
dioxide was determined by the micro-diffusion
method. Sulfur dioxide and hydrogen sulfide were
determined gravimetrically as BaSO 4 and Ag2S,
respectively. The chloride was determined photo-
metrically with mercuric thiocyanate and iron
alum. The residual gases such as hydrogen, nitro-
gen, methane and argon were analyzed by gas
chromatography.
Deuterium analysis was carried out by passing
5 mg of water over hot uranium and comparing
the resulting hydrogen gas with a standard of
known deuterium content in a dual collector mass
spectrometer [13,14]. For 180 analysis, 2 ml of
water sample was equilibrated isotopically with 4
ml of CO 2 at 25°C [15]. This CO 2 was isotopically
analyzed with a mass spectrometer.
In order to determine the hydrogen isotopic
compositions of volcanic hydrogen and methane,
these gases must be separated from other alkali-
unabsorbed gases. The separation of hydrogen and
methane from other gas components was made by
a gas chromatograph on a 3 m molecular sieve 5 A
column with helium carrier gas [4]. After separa-
tion, the hydrogen was passed over a CuO fur-
nance, heated to about 500°C and the combustion
product, H20, was trapped from the helium car-
rier gas. Similarly, the purified methane was con-
verted to CO 2 and H20 by combustion over hot
copper oxide at 830°C. The H20 was separated
from the CO 2 in a dry ice-cooled trap. These H20
samples were reduced to hydrogen gas over hot 43
uranium and trapped on an active charcoal at
liquid nitrogen temperatures [16]. The D/H ratios
of gases were then analyzed on a mass spectrome-
ter.
The isotopic ratios were expressed as 8 values
relative to the SMOW standard. Overall accuracies
of water samples are estimated to be _ 2%o for 8D
and _+0.2%0 for 8180. The overall reproducibility
for hydrogen and methane D/H analyses is _+ 5
and _ 3%o, respectively.
4. Results and discussion
4.1. Chemical composition of volcanic and geother-
mal gases
The analytical results (in vol.%) of the gases
sampled from volcanic and geothermal areas are
given in Table 1 and 2. Fig. 2a is a triangular
diagram showing the relative CO2, H2S + SO 2 and
N 2 contents exclusive of the major constituent
H20. In most of the areas, carbon dioxide con-
stitutes a large proportion of the gases. Hydrogen
sulfide is a predominant sulfur species except for
Za~. The Manza karabuki fumarolic gases in
Kusatsu shirane are characterized by exceptionally
high H2S contents relative to other components.
Little or no HCI is contained in these volcanic and
geothermal gases. The residual non-acidic gases of
Nasu and Issaiky~ (Azuma) have appreciable con-
centrations of hydrogen and methane, while the
gas in other volcanic areas has a high proportion
of nitrogen (Fig. 2b). In the Matsukawa geother-
mal system, hydrogen and methane concentrations
are much higher than those in other volcanic re-
gions. In most of the gas samples from these
volcanic and geothermal areas, the content of hy-
drogen is higher than that of methane.
High-temperature volcanic gases predominantly
contain SO 2, HC1 and H 2 except for H20 and
CO 2, and little CH 4 [17-19]. When these deep-
seated volcanic gases are cooled and condensed
through groundwater, sulfate and hydrogen sulfide
are formed by the disproportionation of SO 2, and
hydrogen chloride is dissolved in water [20]. Up-
ward movement of this thermal system containing
gases is accompanied by boiling and phase separa-
tion. Consequently, sulfate and chloride are dis-
44
TABLE 1
Chemical composition of volcanic gases collected 1981 at Zag, Azuma (Issaiky~), Nasu and Kusatsu shirane from June 1980 to
October 1981 (in vol.%).
Location Outlet H20 CO 2 H2S SO 2 HCI N 2 H 2 CH 4 Ar
T(°C) (xl0 -4) (xl0 -4) (xl0 -4)
Za~ 94.3 96.8 2.51 0.018 0.27 tr. 0.398 3.9 21 25
96.0 97.7 1.98 0.012 0.24 tr. 0.088 4.3 4.1 4.0
Issaiky~ 87.0 - 86.6 7.33 0.30 - 5.78 607 63.6 231
94.0 97.0 2.30 0.537 0.099 tr. 0.054 61.2 17.0 1.7
Nasu Oana 110 99.6 0.224 0.064 0.009 0.016 0.053 130 71.7 4.7
112 99.7 0.210 0.065 0.013 0.007 0.008 27.4 4.5 0.7
113 99.7 0.213 0.066 0.014 0.012 0.006 14.3 2.2 0.3
Nasu Mugen 149 99.6 0.298 0.052 0.012 0.006 0.017 26.2 15.1 3.2
169 99.8 0.156 0.037 0.002 0.011 0.005 7.0 1.1 0.6
154 99.7 0.243 0.046 0.005 0.011 0.004 15.8 2.0 0.3
Kusatsu shirane 96.0 98.2 1.35 0.196 0.009 tr. 0.439 18.9 1.2 28.1
Kusatsu
seshogawara 95.0 94.1 3.55 2.11 0.01 tr. 0.225 5.1 39.5 16.3
Kusatsu yubatake 64.0 - 49.6 1.89 0.12 - 47.5 29.5 987 8320
Manza karabuki 96.0 99.2 0.051 0.702 0.008 tr. 0.039 5.3 2.8 5.2
94.0 99.2 0.146 0.629 0.022 0.002 0.004 0.54 0.15 0.36
tr. = trace;- = not determined.
tributed in the liquid phase and H2S, CO 2 mainly
into the vapor. Hydrogen remains together with
methane in the gas phase because these gases are
hardly dissolved in the liquid phase. This liquid
phase may correspond to acid chloride-sulfate type
waters which exist in northeastern Japanese
volcanic areas such as Tamagawa, Za~ and Kusatsu Manza. Thus, a mixture of the high-temperature
gases with groundwater is indicated by little or no
HC1 and SO 2 contents, and high H2//CH4 ratio in
most gas analyses. The high water content is also
again suggestive of meteoric water contamination.
Therefore, these facts suggest that most of the
samples collected in northeastern Japan represent
TABLE 2
Results of chemical and isotopic analyses of collected gases at Matsukawa geothermal areas (in vol.% and %o, respectively)
Well No.: 1 2 2 5 6 8 8 9
Year: 1977 1977 1981 1981 1977 1977 1981 1981
H20 99.5 99.6 99.8 99.7 99.6 99.7 99.6
CO 2 0.415 0.319 0.201 0.296 0.305 0.199 0.320
H 2 S 0.043 0.051 0.031 0.036 0.055 0.065 0.069
N 2 ( × 10 -4) 49.9 25.2 22.7 37.2 47.2 20.5 43.2
H 2 ( X 10 -4) 24.1 63.6 38.5 14.4 14.0 29.7 38.6
CH 4 (X l0 -4) 43.2 10.9 10.8 15.3 35.1 5.5 22.2
DH2 -- 364 - 419 - 465 - 399 - 385 - - 406
8 DCH 4 -- 274 -- 386 -- 410 -- 351 -- 274 -- 320 -- 384
8 DH2 o -- 76.3 -- 76.8 -- 74.1 -- 78.3 -- 75.8 -- 72.8 -- 79.5
818OH2 o -- 4.8 -- 4.1 -- 7.8 -- 8.4 -- 5.0 -- 5.3 -- 8.8 99.4
0.538
0.033
25.0
47.6
20.0
- 466
- 395
- 76.6
-8.1
- = not determined.
45
(b) N2
20 o 80
®
CO 2 " ).,p~.. 80 s0 t0 20 ~s" r'°2
L \
,0.,7".'.." 80 60 40 20 N2
b the composition of gas phase equilibrated with
heated groundwater.
4.2. Hydrogen isotopic composition of volcanic hy-
drogen
The results of isotopic analyses of volcanic hy-
drogen and methane are listed in Table 2 and 3,
and shown in Fig. 3. The variation of 8D in
volcanic hydrogen is considerable; -220 to
-515%o. Volcanic hydrogen from Nasu fumarolic
gases, in which HOI is found in low concentration
at temperatures in the range from 110 to 170°C, is
depleted in deuterium compared to other Japanese
volcanic gases. On the other.hand, the 8D value of
the hydrogen at Kusatsu shirane, in which the
ratio of hydrogen to methane is large, is -2207~
Fig. 2. (a) Relative proportions of carbon dioxide, and total
sulfur and nitrogen in volcanic and geothermal gas sa~aples,
northeastern Japan. (b) Relative H2, CH 4 and N 2. • = Zao,
Q= Azuma (IssaikyS); ® = Nasu; O = Kusatsu shirane; @=
Matsukawa.
TABLE 3
Isotopic analyses of hydrogen, methane and waters collected at Za~, Issaiky~, Nasu and
1981 (in 96o) Kusatsu shirane from June 1980 to October
Location 8DHz 8DcH 4 8OH: O 818OHIo
Za~ - - - 87.0 - 13.6
- 67.3 - 9.8 (h.w.)
-370 - 181 -77.1 - 10.8
- 58.9 - 7.3 (h.w.)
Issaiky~ - - - 91.4 - 12.8
- 486 - 378 - 69.7 - 8.0
Nasu Oana -488 -400 -62.3 -7.1
- -394 -59.5 -6.8
-515 -430 -59.1 -6.8
Nasu Mugen - - 457 - 63.9 - 6.7
-31.4 - 1.3 (h.w.)
- -407 -61.4 -6.8
- - 487 - 60.5 - 6.7
Kusatsu shirane - 220 - - 75.9 - 9.3
Kusatsu seshogawara - -220 - 102 - 16.6
Manza karabuki -300 - - 101 - 14.5
- 72.7 - 8.7 (h.w.)
- - - 103 - 15.0
-73.1 -8.7 (h.w.)
h.w. = hot water; -: not determined.
46
Surtsey
Zo5
Azuma
NOsu
Kusotsu shirane
Matsukawa
IceLand
Yellow stone t
0
i
I
i
• o I
I
I
I
I
o
I
i
i
n j.J-L o IF.=.
I I I I I I
- 600 - 400 - 200 ~ D (%,)
Fig. 3. Hydrogen isotopic ratios of hydrogen and methane in
volcanic and geothermal gases. Also included are data on
hydrogen from Surtsey volcano [1] and Iceland geothermal
areas [2], and hydrogen and methane from Yellowstone hot
springs [4]. Large solid and small open circles indicate hydro-
gen and methane, respectively.
and much higher than those of other volcanic
areas except Surtsey [1]. The average 8D value of
volcanic hydrogen from northeastern Japan is
much lower than those of hydrogen from the
Surtsey volcano in Iceland. The hydrogen in hot-
water-dominated geothermal system in Iceland is
significantly depleted in D compared with that
from volcanic gases in Surtsey and the northeast-
ern Japanese volcanic areas, but the isotopic val-
ues are higher than those of the high-temperature
geothermal area at Yellowstone Park (SD about
-700%o [4]). On the other hand, the 8D value of
hydrogen from several wells in Matsukawa, which
range from -360 to -470%~, are higher than
other geothermal areas and are similar to those
from volcanic gases in northeastern Japan except
for Kusatsu shirane.
4.3. Isotopic composition of waters from volcanic
and geothermal areas
On the basis of isotopic studies, it has been
concluded that most waters associated with
volcanic and geothermal areas are of local meteoric
origin [21,22]. Kusakabe et al. [22] and Mizutani
[23] gave isotopic data on fumarolic condensates j/ -40
~-60
t-~
oo
-80
4
Hot ~k:lt er
Local Meteoric
o Water -i oo
I
-1 5 -10 ~ls 0 (,/,,) -5 0
Fig. 4. Isotopic composition of waters from volcanic and geo-
thermal areas. M.W.L = Meteoric water line (gD = 8 8n80+
10). • = Za6; Q= Azuma (Issaiky~), • = Nasu; © = Kusatsu
shirane; ~ = Matsukawa.
from several volcanoes of Japan. The results indi-
cate that Japanese volcanic condensates are de-
rived from local surface water which has been
heated and enriched in ~80 by exchange with
silicates and/or a mixture of local meteoric waters
with magmatic water recycled from hydrous crustal
rocks. The former situation best describes the
low-temperature fumarolic condensates from most
of the volcanoes in this study, the latter, the
high-temperature (> 400°C) steam condensates
such as Showashinzan [18,23] and Satsuma-
Iwojima [19].
The relationship between 8D and 180 of vari-
ous type of waters collected in this study is shown
in Fig. 4. The pattern found in this figure is simi-
lar to those in the low-temperature fumarole in
other volcanoes as above. Volcanic condensates
from low-temperature fumaroles of -100°C in
Za6 and Manza karabuki (Kusatsu shirane) are
accompanied by boiling water and liquid-vapor
separation. The temperatures of liquid-vapor sep-
aration in these fumaroles estimated from the ex-
tent of fractionation between steam and hot water
agree with the fumarolic temperature. Most steam
condensates with small D and 180 shifts are essen-
tially local meteoric in origin. It will also be seen
from Fig. 4 that the deuterium content of steam at
Matsukawa vapor-dominated geothermal systems
is approximately equal to that of the local meteoric
water. However, the enrichment in 180 of steam
can be attributed to high-temperature exchange
with silicates.
4.4. Isotopic geothermometry
Volcanic hydrogen is significantly depleted in
deuterium compared to water, due to hydrogen
isotope exchange. The 8 D values of coexisting H 2
and H20 have been used as a geothermometers,
assuming that the gases are in isotopic equilibrium
through the following reaction:
H20 + HD = HDO + H 2
The isotope distribution in a gas phase is given by
the fractionation factor a for the exchange reac-
tion:
a = (D/H)w/(D/H)H
where "W" and "H" indicate water vapor and
TABLE 4
Comparison of isotopic temperatures with temperatures mea-
sured in volcanic and geothermal areas
Location Otw-H 2 Isotope Outlet
temper- temper-
ature ature
(°c) (°c)
Za~ 1.4938 390 96
lssaiky~ 1.8097 248 94
NasuOana (1) 1.8315 241 110
(2) 1.9400 211 113
Kusatsu shirane 1.1847 714 95
Manza karabuki 1.3247 526 96
Matsukawa Year Otw_H2 Isotope Bottom
Well No. temper- temper-
ature ature
(oc) (oc)
1 1977 1.4523 415 250
2 1977 1.5890 342 240
2 1981 1.7307 270 240
5 1981 1.5336 370 220
6 1977 1.5028 382 240
8 1981 1.5497 370 260
9 1981 1.7292 270 230
a = fractionation factor for the hydrogen isotope exchange
between water vapor and hydrogen gas. 47
hydrogen, respectively. The equilibrium fractiona-
tion factor between water vapor and hydrogen is
taken from both theoretical and experimental
studies [24-26]. According to Rolston et al. [26],
the fractionation factor, a, over the range
273-473°K can be expressed as:
In a = -0.2735 + 449.2/T+ 2380/T 2 (1)
where T is the absolute temperature. This experi-
mental scale is nearly consistent with the theoreti-
cal scale by Bottinga [24]. Hydrogen isotopic ex-
change could also taken place between hydrogen
and other hydrogen compounds such as HzS and
CH 4. However, since water vapor constitutes more
than 95% of the volcanic gases in northeastern
Japan, the isotopic behavior of hydrogen would be
controlled by isotopic equilibrium with water. Tak-
ing the 8D value of meteoric water in northeastern
Japan [27] and the isotopic fractionation factors
between water vapor and hydrogen at a maximum
temperature range of 100 - 400°C [26], hydrogen
of 8D = - 360 to - 650%o in volcanic and geother-
mal areas would be estimated (Fig. 3). The 8D
value of hydrogen becomes lower when the tem-
perature of exchange reaction falls. Except for the
fumarolic gases in Kusatsu shirane, the observed
values of -364 to -515%o are compatible with
the model if the uncertainties in temperature is
taken into account. This fact suggests that volcanic
hydrogen and water are in isotopic equilibrium
with each other in a hydrothermal temperature
range (100-400°C).
Isotopic temperatures calculated for isotopic ex-
change between water vapor and hydrogen using
equation (1) are summarized in Table 4. The iso-
topic temperatures in Za~ and Manza karabuki
(Kusatsu shirane) are calculated from the deu-
terium content of local surface water because in
both areas the exchange is accompanied by
liquid-vapor separation. The isotopic temperatures
obtained range from 200 to 700°C, and are much
higher than the outlet temperature of the fumaroles
of around 100°C. The highest outlet temperature
of Mugen fumaroles in Nasu volcano of 180°C
nearly agrees with the isotopic temperature. Arna-
son [2] obtained reasonable agreement between the
measured temperatures in wells of Iceland and the
isotopic temperatures from the hydrogen-water ex-
48
change reaction. Similar results have been ob-
tained for Wairakei, New Zealand [28]. Therefore,
the discrepancy between observed and calculated
temperatures indicates that the high equilibrium
temperature at depth has been frozen-in as tem-
peratures decreased in the rising gas, or that iso-
topic equilibrium was never established. In the
Matsukawa geothermal areas, the isotopic temper-
atures from drill holes 2 and 9 are approximately
the same as the temperatures measured. However,
the isotopic temperatures from other drill holes are
higher than the measured temperatures. It is possi-
ble that the equilibrium isotopic ratio of hydrogen
changes due to cooling during upward flow. Kiyosu
[29] reported that the isotopic temperatures ob-
tained from 8348 of hydrogen sulfide and anhydrite
in the Matsukawa area ranged from 250 to 350°C.
These isotopic temperatures are similar to the hy-
drogen isotopic temperatures. This fact suggests
that in this area the H20, H 2 and H2S have
chemically and isotopically equilibrated with the
deeper mineral components at a temperature of at
least 250°C.
4.5. D / H ratios of volcanic methanes
The 8D values of volcanic methane are more
variable than those of hydrogen; - 180 to - 487%o,
as shown in Fig. 3. Gunter and Musgrave [4] re-
ported 8D values of -225 to -292%o for hot
spring gases of yellowstone Park. On the other
hand, the range of geothermal methanes from
Matsukawa is -247 to -410%o. The variation in
D/H ratios is remarkable and comprises a range
of 300%o. The increase or decrease of deuterium in
volcanic methane could be due to the fractionation
of hydrogen isotopes and/or the mixing of various
methanes (e.g., biogenic or thermogenic CH4).
If methane is in isotopic equilibrium with water
vapor, the H20-CH 4 hydrogen exchange reaction
is responsible for the D/H variation of volcanic
methane. At isotopic equilibrium between CH 4
and H20, for example at a temperature range of
100-400°C, the methanes should be 80-100%o de-
pleted in deuterium as compared to cogenetic water
vapor [30]. Most waters associated with volcanic
gases lie within the range - 50 to - 90%o. Methanes
in equilibrium with these waters should range in -40
- 6O
-80
-I O0 @
I I
-400 -300 /I .u
/fl,
°ti!'
,"I! ° I
-200
Fig. 5. Relationship between 8D values of methane and water.
• = Za~, Q= Azuma (Issaiky~); • = Nasu; © = Kusatsu
shirane; O = Matsukawa. Solid and dashed lines: equilibrium
isotherms of liquid water and water vapor [30].
their 8D values from -150 to -200%o. On the
other hand, the 8D values of methanes based on
the theoretical scale of Bottinga [24] range from
-96 to -155%o. However, the range of the pre-
sented volcanic gases is - 300 to - - 450%0 except
for Za~ and Kusatsu shirane, as shown in Fig. 5.
The data presented here thus show that hydrogen
isotopic equilibrium has not been achieved in the
methanes from volcanic areas.
Fig. 6 shows the relationship between the hy-
drogen isotopic composition of hydrogen and
methane from volcanic gases. The equilibrium iso-
therms based on the experimental scale of Horibe
and Craig [31] are given in this figure. The arrow
indicates the direction of the change in 8D values
in hydrogen and methane after equilibration with
the associated H20 with 8D values of -50 to
-90%o. Most data plot within the field deviated
from this arrow. This suggests that volcanic
methane has not isotopically equilibrated with
water vapor and hydrogen. On the other hand, the
higher the 8D value of volcanic hydrogen, the
/ /
I DeCrt, a$ing //H20(V ) , , / lncr~tsmg .~
-2ooL ..",
/ / / J /¢'-~/ / I I i I //'71."
4 °°t ,' ," ," ,7,'I" ~=~ I I I" , "/',7 I I I C~ / / ' (250"C) ~/// ~_'e / ~o I I i i l ® /i /" o/(3oo-4oo'o
-400 I ii ii I 111¢~7e/~
I II il i I /(~3~1
V lilt/Ill/lit ,
-600 -500 -400 -300
DR2 (%*)
Fig. 6. Plot of 8D values for hydrogen versus methane. O =
Yellowstone [4]; • = ZaB; Q= Azuma (Issaiky~); Q = Nasu;
O = Matsukawa. Bracket represents the isotopic temperature
based on the observed HzO-H 2 hydrogen isotope fractiona-
tions. Dashed line: equilibrium isotherms [31]. H20(1 ) = liquid
water. H20(v ) = water vapor.
higher the deuterium content of methane is. Fur-
thermore, the D concentration in the methane
increases with increasing isotopic temperature
calculated from H20-H 2 hydrogen isotope frac-
tionations as shown in Fig. 6. Thus, D/H isotope
analyses of these methanes reveal that their hydro-
gen isotopic composition are dependent on the 8D
value of hydrogen and isotopic temperature.
Variation of 8D values in volcanic methanes
may be caused by admixture of natural gases of
different origin. According to Schoell [32], the 8D
values of biogenic methane from natural gases of
worldwide occurrences, range from -180 to
-2807o0. It should be noted that the D/H varia-
tions in volcanic methanes of northeastern Japan
are not solely due to the mixture of biogenic
methanes, since most of the 8D values for volcanic
methanes are more negative than -300~. Al-
though the hydrogen isotopic composition of ther-
mogenic methane ranges from a 8D value of - 150
to -260700 [32], their D/H variations are qualita-
tively comparable to the variations which were
observed in laboratory pyrolysis experiments [16].
At 400 ° and 500°C there are hydrogen isotope
fractionations during methane production of 170
and 140%o, respectively, with deuterium depletion
in the methane caused by the by pyrolysis of 49
hydrocarbons. Although there is an uncertainty, a
temperature effect of 30%o per 100°C in the pre-
dictable direction is obtained, If this fact is cor-
rect, then hydrogen isotopic fractionation during
methane production below 300°C would be larger
than 200%0. Assuming that hydrocarbons of a
mean isotopic composition such as that of crude
oils [33] and kerogens [34] were the parent materials
to all these gases, the deuterium concentrations of
produced methane range from - 350 to - 4007~ at
300°C, -230 to -320%o at 400°C and -180 to
- 3007o0 at 500°C. Thus, the 8D values in thermo-
genic methane increase with increasing tempera-
ture. This variation pattern is similar to that of the
values observed. Fig. 7 shows a plot of 8D value
for methane vs. the CH4/H: ratio. The deuterium
concentrations are related to the CH4/H 2 ratio
and would also increase with increasing isotopic
temperature as above. The production of methane
by pyrolysis of hydrocarbon occurs rapidly as
temperature increases [16]. Therefore, D/H varia-
tion of volcanic methane may be due to the kinetic
effect during the production of thermogenic
methane.
The 8D values of methane may also be affected
by the kinetic effect during inorganic production,
assuming that the chemical reaction proceeds as
-200
~ -300
t"~
t,,o
- 400 ee '1
I
e/
I
0.1 1
CHdH 2 I
10
Fig. 7. Plot of 8D values for methane versus CH4/H: ratios.
• = Za~; ®= Azuma (IssaikyB); (D = Nasu; O = Kusatsu
shirane; O = Matsukawa.
50
follows:
CO 2 + 4 H 2-- CH 4+ 2 H20
This chemical reaction is important for the forma-
tion of methane [3,35] and has been suggested in
early carbon isotope studies [36,37]. Schoell [32]
and Coleman et al. [38] reported D/H fractiona-
tion to occur when methane is produced and/or
oxidized by microbial processes. Isotopic
fractionation in the inorganic formation or oxida-
tion of methane may also be controlled by kinet-
ics. This reaction shifts to the right with decreasing
temperatures. Since hydrogen readily exchanges
isotopically with the associated water vapor and
methane, the 8D values of hydrogen and methane
decrease as the CHa/H 2 ratio increases in re-
sponse to a decrease in temperature. However,
lower CH4/H 2 ratio should result in even more
negative 8D values of hydrogen and methane pre-
sented here as shown in Figs. 6 and 7. Therefore,
the volcanic methanes may not be formed by the
reaction of carbon dioxide and hydrogen. Thus,
this formation process is obviously of minor im-
portance for the origin of volcanic methane.
Gunter and Musgrave [4] suggested that the
methane of hydrothermal gases from Yellowstone
Park is formed by the thermal decomposition of
organic matter, not by the reaction of hydrogen
with carbon dioxide. Variations in 8D values of
volcanic methane in northeastern Japan may be
caused by a mixture of thermogenic methane. Fur-
ther discussion may become possible when a sys-
tematic carbon isotope study of the volcanic
methane is completed.
5. Conclusions
The hydrogen isotopic composition of hydrogen
from low-temperature fumarolic gases (- 100°C)
in northeastern Japan is lighter than that of high-
temperature (> 1000°C) fumarolic gases such as
those from the Surtsey volcano in Iceland, and
havier than hydrogen of hot spring gases in Yel-
lowstone Park. The D/H ratio of hydrogen from
geothermal gases in the Matsukawa vapor- dominated system is higher than that of the Ice-
land hot-water-dominated geothermal areas and
similar to that of fumarolic gases in Nasu volcano.
The hydrogen isotope data suggest that the hydro-
gen of volcanic gases in northeastern Japan may
be isotopically equilibrated with water vapor at a
temperature of 200-400°C at depth.
In volcanic regions other than Nasu, the iso-
topic temperature, evaluated from the isotopic
fractionation factor between water vapor and hy-
drogen, is much higher than the outlet tempera-
tures of the fumaroles. The discordance between
observed and isotopic temperatures indicates the
freezing-in of a high-temperature equilibrium at
depth as temperatures decrease in the rising gas, or
the lack of isotope exchange between water vapor
and hydrogen at low temperatures. In the Nasu
volcano, the isotopic temperatures are similar to
the fumarolic temperatures, suggesting that the
water vapor equilibrates isotopically with hydro-
gen near the surface. The isotopic temperatures
from Matsukawa geothermal area are higher by
30 - 105°C than the bottom temperatures in drill
holes.
The hydrogen isotope ratio of volcanic methane
is more variable than that of hydrogen. The D/H
isotope data also suggest that most volcanic
methane has isotopically disequilibrated with water
vapor and hydrogen. Variation of 8D values in
these methanes may be due to a mixture of ther-
mogenic methane. Future investigations of carbon
and hydrogen isotopes will be most useful in order
to clarify the origin of volcanic methane.
Acknowledgements
The author wishes to thank N. Nakai and M.
Kurahashi (Nagoya University) and Y. Yoshida
(Japan Metals-Chemicals Co.) for their help in the
collection and analysis of the gas samples, and H.
Craig (University of California), who reviewed the
manuscript, for helpful comments.
Part of the cost of this study was defrayed by a
Grant in Aid for Scientic Research, No. 554168,
from the Ministry of Education.
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What are the key findings and results, important concepts and background context of the research in 'Suda et al 2014 - origin of CH4 in serpentinite-hosted hydrothermal systems, Hakuba Happo hot spring, Japan'?
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Based on the research paper 'Suda et al 2014 - origin of CH4 in serpentinite-hosted hydrothermal systems, Hakuba Happo hot spring, Japan', the key findings and results include (JAPAN) Suda et al 2014 - origin of CH4 in serpentinite-hosted hydrothermal systems, Hakuba Happo hot spring, Japan.txt
Earth and Planetary Science Letters 386 (2014) 112-125 Contents lists available at ScienceDirect EARTH Earth and Planetary Science Letters ELSEVIER www.elsevier.com/locate/epsl Origin of methane in serpentinite-hosted hydrothermal systems: rossMark The CH4-H2-H2O hydrogen isotope systematics of the Hakuba Happo hot spring Soichi Omori f, Keita Yamada , Naohiro Yoshida Cg, Shigenori Maruyama a,c aDepartment fEarthandPlanetarycincTokyoIstituteofTechlogy,11kaymaMur-uToky58551an.... The background context for this study involves the motivation and circumstances that led to this investigation.
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(JAPAN) Suda et al 2014 - origin of CH4 in serpentinite-hosted hydrothermal systems, Hakuba Happo hot spring, Japan.txt
Earth and Planetary Science Letters 386 (2014) 112-125 Contents lists available at ScienceDirect EARTH Earth and Planetary Science Letters ELSEVIER www.elsevier.com/locate/epsl Origin of methane in serpentinite-hosted hydrothermal systems: rossMark The CH4-H2-H2O hydrogen isotope systematics of the Hakuba Happo hot spring Soichi Omori f, Keita Yamada , Naohiro Yoshida Cg, Shigenori Maruyama a,c aDepartment fEarthandPlanetarycincTokyoIstituteofTechlogy,11kaymaMur-uToky58551an bPrecambrianEcosystemLaboratoryJapanAgencyforMarin-EarthcienceandTechnologyJAMSTEC)2-15Natsushima-cho,okosuka23-0061,Jaan Earth-ifcincIstitutokyIstitutfTehnlogy1kayamaMgurToky55n dDepartmentofionformation,TokyoInstituteofTechnology459NagatsudaMidorikuYokhama2268502,Jaan DepartmentfiologicalcinceokystituteofTehlgy1kayama,Mgurkoky550an fFaculty of Liberal Arts,heOpenUniversity of Japan-1Wakaba,Mihama-ku,hiba261-8586,Japan DepartmentofEnvironmentalChemistryandEngineeringInterdisciplinaryGraduatechoolofScienceandEngineerinTokyoInstituteofTechnology 4259 Nagatsuda,Midoriku,Yokohama 226-8502,Japan ARTICLE INFＯ ABSTRACT Article history: Serpentinite-hosted hydrothermal systems have attracted considerable attention as sites of abiotic Received 6 June 2013 organic synthesis and as habitats for the earliest microbial communities. Here, we report a systematic Received in revised form 31 October 2013 isotopic study of a new serpentinite-hosted system: the Hakuba Happo hot spring in the Shiroumadake Accepted 1 November 2013 area, Japan (36°42'N, 137°48'E). We collected water directly from the hot spring from two drilling Available online 22 November 2013 wells more than 500 m deep; all water samples were strongly alkaline (pH > 10) and rich in H2 Editor: B. Marty (201-664 μmol/L) and CH4 (124-201 μmol/L). Despite the relatively low temperatures (50-60°C), Keywords: thermodynamic calculations suggest that the Hz was likely derived from serpentinization reactions. abiotic methane Hydrogen isotope compositions for Happo #1 (Happo #3) were found to be as follows: 8D-H2 = serpentinization -700% (-710%o), 8D-CH4 = -210%o (-300%o), and 8D-H20 = -85%o (-84%o). The carbon isotope hydrogen isotope compositions of methane from Happo #1 and #3 were found to be 813c = -34.5%o and -33.9%, respectively. The CH4-H2-HzO hydrogen isotope systematics indicate that at least two different mechanisms were responsible for methane formation. Happo #1 has a similar hydrogen isotope compositions to other serpentinite-hosted systems reported previously. The elevated SD-CH4 (with s a # o n n sisns (n a on n sourced from molecular hydrogen but was derived directly from water. This implies that the methane may not have been produced via the Fischer-Tropsch-type (FTT) synthesis but possibly by the hydration of olivine. Conversely, the depleted SD-CH4 (with respect to the equilibrium relationship) in Happo #3 suggests the incorporation of biological methane. Based on a comparison of the hydrogen isotope systematics of our results with those of other serpentinite-hosted hydrothermal systems, we suggest that abiotic CH4 production directly from H2O (without mediation by Hz） may be more common in serpentinite-hosted systems. Hydration of olivine may play a more significant role in abiotic methane production than previously thought. 2013 Elsevier B.V. All rights reserved. 1. Introduction were exposed more extensively in the Hadean and early Archean ocean owing to the higher potential temperatures compared to the Serpentinite-hosted hydrothermal systems are thought to be present-day mantle (Komiya et al., 2004). Accordingly, abiotic gas important for prebiotic synthesis and have been suggested as habi- production associated with serpentinization is expected to have tats for some of the earliest life. Although ultramafic rocks are played a more significant role in the early Earth environment. Hy- rarely exposed at the modern seafloor, it is possible that they drothermal fluids derived from ultramafic rocks are characterized by high concentrations of H2 and CH4 (e.g., Neal and Stanger, 1983; Abrajano et al., 1988; Charlou et al., 2002; Takai et al., 2004; E-mail address: [email protected] (K. Suda). Proskurowski et al., 2006). During water-rock reactions, Fe(Il) in 0012-821X/S - see front matter ? 2013 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.epsl.2013.11.001 K.Suda et al./ Earth and Planetary Science Letters 386(2014)112-125 113 olivine of ultramafic rock is oxidized to Fe(I), which accompa- 813C alone may not improve understanding of methane production nies the reduction of water to yield molecular hydrogen (H2). The mechanisms under serpentinite-hosted hydrothermal conditions. resulting high concentrations of H2 are important: they support In the present study, we focus on hydrogen isotopic systematics chemolithoautotrophic metabolic pathways (including methano- to consider the origin of methane in serpentinite-hosted systems. genesis) that may have represented Earth's earliest microbial com- In particular, it is important to characterize the isotopic relation- munity (Nisbet and Sleep, 2001; Takai et al., 2006; Ueno et al. ships between CH4 and substrates and to estimate the effects of 2006; Sleep and Bird, 2007). methane production mechanisms on isotopic content. Compari- Furthermore, the abiotic formation of organic compounds in son between measured and equilibrium hydrogen isotopic rela- serpentinite-hosted systems has long been thought possible; such tionships for the H2-CH4-H2O system is essential for understand- compounds were likely crucial for the chemical evolution of life ing the origins of methane in serpentinite-hosted hydrothermal (e.g, Holm et al., 2006). Methane and hydrocarbons are often systems (Bradley and Summons, 2010). As hydrogen sources, H2 observed in serpentinite-hosted hydrothermal systems and are and HzO are potential substrates for methane production reac- thought to be produced from H2 and CO2 via Fischer-Tropsch- tions in such systems. However, full hydrogen isotopic analysis type (FTT) reactions (e.g., Holm and Charlou, 2001; Proskurowski of CH4, H2, and H2O from samples from serpentinite-hosted sys- et al., 2008). FTT reactions are defined as the surface-catalyzed re- tems has been reported in only a few studies (Fritz et al., 1992; duction and polymerization of oxidized single carbon compounds Proskurowski et al., 2006). (McCollom and Seewald, 2007), including the Sabatier reaction (CO2 + 4H2 = CH4 + 2H2O), which yields CH4 but not higher sporadically in the serpentinite-bearing Shiroumadake area, Japan, and have discovered a new serpentinite-hosted system in Hakuba hydrocarbons. Previous laboratory experiments under hydrother- mal conditions (>2o0°C) have demonstrated that minerals such Happo. Here, we report the results of chemical and C-H isotopic as Fe-Ni alloy and magnetite catalyze FTT reactions (Horita and analyses and discuss the processes of methane generation based Berndt, 1999; Fu et al., 2007). Such minerals can be produced dur- on CH4-H2-H2O hydrogen isotope systematics, which led us to ing the serpentinization of ultramafic rocks. However, in recent reconsider the origins of methane in typical serpentinite-hosted hydrothermal systems. years, increasing field investigation of the serpentinization process has suggested that low-temperature (<150°C) abiotic methane 2. Geological setting synthesis is more common than previously thought (Etiope and Sherwood Lollar, 2013). In particular, thermodynamic calculations The Shiroumadake area (36°N, 137°E) is situated in central have suggested one possible mechanism for abiotic methane for- mation at lower temperatures: hydration of olivine or orthopyrox- Honshu, Japan (Fig. 1). The region to the west of the Itoigawa- Shizuoka Tectonic Line (ISTL) belongs to the Hida marginal belt, ene in the presence of a carbon source (Oze and Sharma, 2005; which consists of Paleozoic and Mesozoic basement rocks, Late Miura et al., 2011). Unlike FTT reactions, this reaction does not re- quire molecular hydrogen to produce methane. However, to date, (Nakano et al., 2002). The Paleozoic basement consists of ultra- the production mechanism of methane in serpentinite-hosted hy- mafic, sedimentary, and metamorphic rocks. The ultramafic rocks drothermal systems remains poorly understood. are mostly serpentinized and sheared, whereas relatively fresh Carbon stable isotope ratios (813C) have been used previously to olivine remains in the Happo and Iwatake-Yama areas. discriminate the origins of methane; for example, the 813C values In the Shiroumadake area, active hot springs are distributed sporadically along the ISTL owing to recent volcanic activity conditions have been investigated experimentally. Previous labora- (Homma and Tsukahara, 2008). The unique location of the ul- a e e s a tramafic body in this region provides an ideal opportunity to 1999; McCollom et al., 2010) have shown carbon isotopic frac- ' p oio u uo us o s tionation between carbon substrates and generated methane to be e ps ss o s 8 = -68%o to -30%o. Thus, if methane was produced from car- mafic rock body. Fig. 1 illustrates the localities studied: the Hakuba bon dioxide in seawater (813C = 0%o), its 813C-CH4 value would Happo, Kurashita, Renge, Wakakuri, and Tsugaike hot springs. be expected to be -30%o to -68%o. However, the 813C values Hakuba Happo (36°42'N, 137°48'E) is only hot spring that over- of methane observed in seafloor serpentinite-hosted hydrother- lies the ultramafic rock body and is one of the most alkaline hot mal systems range from -19%o to -8%o (Proskurowski et al., springs (pH > 10) in Japan. The water temperature is approxi- 2008; Konn et al., 2009; Keir et al., 2009; Takai et al., 2004), o s on (o) e o whereas those in continental serpentinite-hosted systems range H2 gas. Happo hot spring water does not well up spontaneously from -37%o to -6%o (Abrajano et al., 1988; Fritz et al., 1992; but is artificially pumped from two drilling wells (Happo #1 and Lyon and Giggenbach, 1994; Etiope et al., 2011, 2013a, 2013b; Happo #3). Water from these two wells is then mixed in a tank Boschetti et al., 2013; Szponar et al., 2013). The observed 13C- and degassed to supply hot water for commercial spa facilities enriched CH4 (i.e., up to -6%o) has not yet been reproduced in around Happo Town (Fig. 2). The Happo #1 and #3 wells are laboratory experiments, which instead demonstrated 13C-depleted 515 m and 7o0 m deep, respectively, and are constructed from CH4 (i.e., below -30%o; Horita and Berndt, 1999; McCollom et al., tight steel casing pipes including strainers at 291-509 m and 2010). Thus, an FTT origin for the CH4 in serpentinite-hosted sys- 240-700 m, respectively. The hot water is pumped from the deep tems has not yet been demonstrated experimentally. Moreover, well and is supplied laterally through the strainers continuously dissolved inorganic carbon is virtually absent in the hyper-alkaline (Fig. 2). Without pumping, the water table occurs at 4.5 m and conditions that prevail in serpentinite-hosted systems (Hoehler 22.7 m for #1 and #3, respectively. During sampling, the pumping et al., 2008), although carbon sources should have been present rates of Happo #1 and #3 were 320 L/min and 710 L/min, respec- during methane generation. Under such carbon-poor conditions, it tively. The subsurface layer is composed entirely of serpentinized is difficult to compare the 13C contents of CH4 with those of poten- ultramafic rock, up to at least 700 m depth from the surface. The tial substrate carbon. If methane is produced in carbon-poor envi- geothermal gradient is illustrated in Fig. 2 based on the measured ronments, isotopic fractionation between methane and substrates temperature in the two wells. may be affected by the Rayleigh process and could be small ow- Kurashita is located only 4 km east of Hakuba Happo and is ing to near-complete conversion. Therefore, using information from a natural hot spring that has been drilled to a depth of 1oo0 m. 114 K.Suda et al./Earth and Planetary Science Letters 386(2014)112-125 Quaternary Sediment [ Igneous rock Neogene JAPAN Pyroclastic rock Jurassic Wakaku Sedimentary rock Permian Ultramafic rock Sedimentary rock Metamorphic rock Hakuba Happo Fault (broken where inferred) ☆O Hot spring Fig.1. Geological map of the wes ernShiroumadakearea.Circlesandstarindicatethestudiedhotsprings.Thestarrepr ents the location of the Hakuba Happo hot spring. Ha Happo#3 BMixingtank-E A Mixing tank 150m 100 200 50° 500 600 700 ※Data from coring log degassed in the mixing tank to supply hot water to commercial spa facilities. Star symbols indicate sampling points. According to the coring log, the static water level is 4.5 mbs for Happo #1 and 22.7 mbs for Happo #3. Serpentinized ultramafic rocks occur down to at least 700 m depth. The geothermal gradient was predicted on the basis of the coring log. The pH and water temperature are approximately 7 and 43-45 °C, and 1200 m, respectively, at the eastern foot of Norikuradake; both respectively. At present, the hot spring water flows out at a rate of 1500 L/min, accompanied by large quantities of gas. More than mately 58 °C and 44°C, respectively (Tsukuda, 2000). 90% of the gas component consists of CO2, with minor CH4, C2H6, and H2 (Tsukuda, 2000). The Renge hot spring group is located in 3. Sample collection and analytical methods the northern part of the Norikuradake zone, at a height of 1,475 m above sea level, and consists of fumaroles and bubbling and muddy Hot spring water samples in the Shiroumadake area were ob- warm springs. The hot spring waters in the Renge group exhibit a n i an u l a ti a u wide range of temperatures (30-89°C) and chemistries (Homma pH, dissolved oxygen level (DO), electrical conductivity (EC), and and Tsukahara, 2008). The pH values of thermal fluids are vari- salinity were measured at the sampling points using portable wa- able, ranging from strong acidic to nearly neutral. The Wakakuri ter quality meters (P30 series, DKK-TOA Corporation). Fluid pH was and Tsugaike hot springs have been drilled to depths of 1100 m measured at in situ temperature using a sensor calibrated with K.Suda et al./ Earth and Planetary Science Letters 386(2014)112-125 115 commercial pH buffers (pH = 4.01 and 6.86). For comparison, we 0.3 also sampled river water from the Matsu River, which runs through H H ±± the Shiroumadake area. 3 3.0 8么 In July 2010, the Hakuba Happo samples were collected at both the drilling well (Happo #3) and the mixing tank (mixing tanks 50年8 5 15.4 06 5151515 A and B) (Fig. 2). In the Wakakuri spring, water samples were carbonate 1 干9 + obtained directly from the drilling well. In the Kurashita spring, 80 2.9.4 838 9 77986 direct sampling from the drilling well was impossible owing to ac- tive eruption. Therefore, we obtained hot spring waters from the drainage outlet of a storage tank that stored erupted spring wa- ±0.2 20 2222 干开干干 ter after degassing; these water samples underwent degassing in 27 12.7 217 the tank. The Renge samples were collected directly from muddy bubbling hot springs. In June, 2011, we sampled the hot spring water of Hakuba pelagic sediment 2 3 Happo again, when we obtained samples directly from two drilling 2038 wells (Happo #1 and Happo #3) and from the mixing tank (mixing tank B) (Fig. 2). sition We collected several samples to achieve the desired analyses: 1000 0.2 (I) concentration and stable isotopic composition of dissolved gas, H (Il) stable isotopic composition of CO2, (Ill) stable isotopic composi- 333 tion of H2OLiq and (IV) dissolved ion concentration of water. For (I1), opic 50 cc water samples were syringed into 125 cc glass vials sealed isoto with butyl rubber stoppers and aluminum crimp seals. These sample vials had been thermally pretreated (at 450°C for 1 h) 1 68 8 and evacuated. About 0.5 g of cadmium acetate [Cd(CH3COO)2] had been contained in each vial to prevent biological activity 6.3.Accretion in sample water and to precipitate corrosive H2S. In the labora- b.d.l. tory, dissolved gas compositions were determined according to a headspace method by using gas chromatography. Stable hydrogen and carbon isotope analyses were performed according to con- 1067 1298 2 s s s s isotope ratio mass spectrometers. For (Il), 50 cc water samples were syringed into the same glass vials as for (l) without cad- 5.5 2220 mium acetate. Dissolved inorganic carbon was extracted as CO2 according to a headspace method similar to (I) with additional 77209159 b.d.l. acid treatment. For (Ill), water samples were collected using a sy- 5 ringe with a filter (pore size: 0.45 μm) and stored in 2 cc glass conc 201 钇切8 positions of HzO were measured by cavity ring-down spectroscopy b. ved (CRDS, Picarro, L2120-i). For (IV), water samples filtered using a 0 soly carb syringe filter (pore size: 0.20 μm) were filled into 30 cc glass 152087 vials and sealed with butyl rubber stoppers and aluminum crimp seals. Cation measurements were performed by inductively cou- inorga pled plasma-mass spectrometry (ICP-MS, ThermoFisher Scientific). 0.03 2000m 000 3. 00 The dissolved anion concentrations were determined by ion chro- total oceanicplate matography. - Detailed description of the analytical techniques used in the 70.3 199 5 15.2 study is given in Supplementary material. salinity, 4. Results 0.29 90 70 The results of in situ site measurement, dissolved gas concen- .9 tion, s n n s s si s /ses (/1 are presented in Table 1. Table 2 presents the concentrations of analy irem anions and cations from the Hakuba Happo hot spring. 52 85 3 36044 tope mea 10.8 8以89 4.1. Chemistry of hot spring fluids 00 8 Istable 11 For the Hakuba Happo site, the hot spring waters from both and 50°C. The primary components of the dissolved gas were N2, H2, chemical and CH4, with minor amounts of O2 and C2H6. Of the five hot odde eanieh re uo aop sm s 7 'pns ss I-river W carbonate concentration of CH4 in the Happo water was found to be 10-100 品 atsu lgail times higher than that of the other hot springs. Dissolved inorganic SP 品 M阳W limestone 明明明 W 12 carbon (DIC) was low for both Happo #1 and #3, likely owing to 116 K.Suda et al./ Earth and Planetary Science Letters 386(2014)112-125 either the alkaline conditions or biological consumption. Charac- teristics such as the strongly alkaline water, H2- and CH4-bearing gas, and depleted inorganic carbon are similar to conditions ob- served for other serpentinite-hosted low-temperature hydrother- mal systems such as Lost City, Oman, Genova, Othrys, Cabeco de Vide and Tablelands (Kelley et al., 2001; Neal and Stanger, 1983; Boschetti et al., 2013; Etiope et al., 2013a, 2013b; Szponar et al., 2013). The Happo #1 water was found to be richer in gas com- ponents that the Happo #3 water. Table 2 presents the results of dissolved ion analyses and demonstrates that Happo #1 and #3 ex- hibit almost identical ion compositions. The concentrations of most ions, particularly Mg, are lower at Hakuba Happo than at other hot 00 springs in the Shiroumadake area. The other four hot springs exhibit significantly different fluid chemistries to that of the Hakuba Happo samples. The Kurashita 6 p an o p sm ( 8'g = H) s s no 00019 0000 salinity of all the samples and high CO2 concentration (>5 mM). Although CH4 (9.1 μmol/L) and C2H6 (0.4 μmol/L) were detected in the dissolved gas, H2 was below the detection limit. The Ci/C2+ ratio for Kurashita was found to be 24; this low C1/C2+ ratio (i.e., <102) is similar to that found for hydrocarbons of thermal decom- 15.18 position origin (Berner et al., 1993). Since the Kurashita sample was sampled after degassing, the actual gas concentration was likely higher than that measured. The Renge hot spring water was the most acidic and exhibited the highest temperature (pH = 1.8 e ) i sn on nnd o o mo o< ne 000 Tsukahara, 2008). A relatively low CH4 concentration (0.3 μmol/L) was also observed at the Renge hot spring. Both Wakakuri and Tsugaike were found to exhibit near-neutral pH at temperatures 000 of 54°C and 41 °C, respectively. The Wakakuri sample was found A to contain 8.0 μmol/L of CH4, whereas the CH4 concentration at Tsugaike was below the detection limit. 8.68 4.2. Stable isotope compositions Comparison of the 2010 and 2011 stable isotope data for Hap- 3 po #3 demonstrates that no isotopic change occurred within the year. The 8D values of hydrogen and water for Happo #1 were 000 found to be similar to those for Happo #3, although the 8D-CH4 value of Happo #1 is enriched in deuterium by approximately 80%0 0.16 relative to Happo #3. The hydrogen isotope composition of the mixing tank methane lies between those of Happo #1 and #3. The pumping rates at Happo #1 and Happo #3 were 320 L/min and 400 710 L/min, respectively. Assuming the mixing ratio calculated from these pumping rates, the 8D-CH4 value of the mixing tank was es- 0.48 timated to be -272%o. The measured mixing value (SD = -261 7. 16. 24 13.7.3. 220 99 15 12 to -248%o） indicates slight deuterium enrichment relative to the estimated value, probably owing to the effects of degassing. The 813C values of CH4 for Happo #1 and for Happo #3 were found to 9 be -34.5%o and -33.8%o, respectively. In contrast to the hydrogen 18108 1112 isotope, almost no difference in carbon isotope ratios was observed between the two samples. Except at Kurashita, all hot spring waters in the Shiroumadake 8 (2008). area were found to exhibit similar 8180 and 8D values, ranging 28Ｂ.2２７43９ - 0 8- - 1- .n ara values are close to the isotopic composition of adjacent river wa- anion ter collected from the Matsu River (8180 = -14%o, 8D = -92%), 8.20 providing clear evidence for a local meteoric origin for all of the pue Year hot springs. In contrast to HzO, the isotopic composition of CH4 fcation lard differs for each hot spring. In particular, the Wakakuri CH4 ex- hibits extremely heavy C and H isotope compositions. The methane f (8D = +8%o) appears to be enriched in deuterium relative to wa- ter (8D = -87%o). Furthermore, the methane appears to be en- riched in 13c (813c = -7.3%), similar to total inorganic carbon ec (813C = -7.2%o). These 13C and D enrichments could be due to microbial oxidation. K.Suda et al./Earth and Planetary Science Letters 386(2014)112-125 117 5. Discussion (a) 5.1. Origin of Hakuba Happo H2 otinite-hosted ☆Happo#1 500 ★Happo #3 In the Hakuba area, high concentrations of H2 were observed 400 Lost City ℃ O Logatchev only at the Hakuba Happo hot spring, which is located within the 300- Oman serpentinite body. The relationship observed between H2 concen- Sediment-hosted tration, pH, and temperature in Happo samples is roughly consis- 200 Basalt-hosted tent with the thermodynamic calculation of McCollom and Bach Continental (2009), suggesting that the serpentinization reaction at 50-60°C ￥ Volcanic gas +Shield borehole should be responsible for the production of H2 in the Happo water. 100 McCollom and Bach (2009) performed thermodynamic calculations in which fluid chemistry depended on mineral composition dur- ing serpentinization using a reaction path model. The predicted pH and H2 concentrations for the reaction of harzburgite with seawater at 50-60°C were 11.2-10.7 and 7-10 mM, respectively (McCollom and Bach, 2009). These values are similar to those for the Happo water, suggesting that the H2 in the Happo hot spring was generated from serpentinization reactions at low temperatures -1600 -1400 -1200 -1000 -800 -600 -400-200 (~60°C). However, the predicted concentration of H2 is approxi- Measured fractionation ε(H2-H2Oaq) %o mately ten to thirty times that observed in Happo water. The cause of this discrepancy is unknown; it may be that H2 was lost through 500 (b) degassing and/or microbial process, although the application of the thermodynamic calculation to such a low-temperature regime is 250- associated with many uncertainties. ℃ Moreover, a recent laboratory experiment demonstrated that H2 ure was produced by the hydration of mafic and ultramafic rocks at low temperatures of 55 °C and 100°C (Mayhew et al., 2013). This further supports the assertion that the Hz in the Hakuba Happo hot spring was generated by serpentinization reactions at low tem- peratures (i.e., below 100 °C). Additionally, the observed hydrogen pa ☆ isotopic fractionation between H2 and H20 (-1120%o to -1150%o) is similar to the equilibrium fractionation at about 60°C; this is > consistent with a low-temperature origin for the H2, although the 8D values of H2 may have been altered owing to isotopic re- △ equilibrium, as discussed in the following section. 0 5.2. Origin of Hakuba Happo CH4 -600 -500 -400 -300 -200 -100 Measured fractionation &(CH4-H2Oaq) % 5.2.1. Preservation of hydrogen isotopic composition of H2 and CH4 Comparison of observed and equilibrium fractionations can pro- vide valuable information regarding the origin of methane in hy- dicted equilibrium fractionation at a given temperature for (a) the H2-HzOaq system and (b) the CH4-HzOaq system. The equilibrium lines were calculated based on drothermal systems (Proskurowski et al., 2006; Bradley and Sum- experimentally calibrated equations for the equilibrium D/H fractionation factors mons, 2010). In this section, isotopic re-equilibration among CH4, (Horibe and Craig, 1995). Open and flled stars represent Happo #1 and #3, re- H2, and H2O is first evaluated. We compare our Happo data with spectively. Other open symbols represent serpentinite-hosted systems: Lost City, Lo- the thermodynamically predicted equilibrium fractionations be- gatchev and Oman. Filled gray symbols indicate the seafloor hydrothermal systems associated with other host rocks. 8D-HzO was assumed to be 0%o for Logatchev and tween H2 and H2O (Fig. 3a) and between CH4 and H2O (Fig. 3b). sediment- and basalt-hosted seafloor hydrothermal sites. Crossed symbols indicate continental gas data from volcanoes and shield boreholes. Hydrogen isotope data at a given temperature according to Horibe and Craig (1995): sources are listed in Table 3. α(H20aq-H2) = 1.0473 + 201 036/T2 + 2.060 × 109 /T4 This equation was calculated from previous experiments for de- termining α(CH4-H2), α(H2Oaq-H2Ovapor), and α(H2Ovapor-H2) + 0.180 × 1015 /T6 (Horibe and Craig, 1995; Horita and Wesolowski, 1994; Suess, where T represents temperature in K. This equation was de- 1949). Then, the enrichment factor 8(CH4-H2Oaq) was defined as veloped from the experimentally determined fractionation factor follows: α(H2Oaq-H20vapor) (Horita and Wesolowski, 1994) and the the0- retical fractionation factor α(HzOvapor-H2) (Bardo and Wolfsberg, ε(CH4 - H2Oaq) = 1000 lnα(CH4 - H2Oaq) 1976). We defined the enrichment factor 8(H2-H2Oaq) as follows: In the H2-H2O system, the equilibrium temperatures for Happo 8(H2-H20aq)= 1000 Inα(H2-H2Oaq) #1 and #3 are 67°C and 60 °C, respectively (Fig. 3a). These equilib- rium temperatures are similar to (but approximately 10°C greater On the other hand, the equilibrium CH4-H2Oaq fractionation was than) the measured temperatures, which were found to 52 °C and defined according to the following equation: 48 °C, respectively. In relatively high-temperature hydrothermal α(H2Oaq-CH4) = 1.0997 + 8456/T2 + 0.9611 × 109 /T4 the measured temperatures. In contrast, the predicted tempera- - 27.82 × 1012 /T6 tures for lower temperature systems (<2o0°C) are consistently Table 3 Hydrogen isotope datasets from various natural environments. Location Geology Sample pH Temp DVSMOW 813CVPDB Reference (°C) SD-Hz &D-CH4 D-HO 1000lnα 1000lnα 1000Inα 813C-CH4 (%) (%) (%) HHO CH4HO CH4H (%) Serpentinite-hosted hydrothermal systems Continental Hakuba Happo_#1 Serpentinite Fluid 10.8 52.3 700 210 84 1117 148 969 34.5 Apns s Hakuba Happo_#3 Serpentinite Fluid 10.7 48.0 710 299 84 1149 267 882 This study Oman Serpentinite Fluid > 11.5 3242a 733 to 712 251 to 210 2.95.2 1326 to1248 294 to 239 993-1031 14.7 to 12.0 [1], °[2] Deep seafloor Lost City Serpentinite + gabbro Fluid 10.510.7b 2890 689 to 605 141 to 99 27 1174 to 933 157 to 109 804-1060 13.6 to 9.4 [3], b[4], ≤[5] Logatchev Serpentinite + gabbro Fluid 3.3d 350 372 601- 0 466 116 350 13.6d [3], °[6] Suda et Deep seafloor hydrothermal systemts Sediment-hosted 户 Endeavor Sediment Fluid 3.46.2e 375 345 to 332 125 to 107 0 423 to 403 134 to 113 289-290 [3]. [7] /Earth Guaymas Pelagic sediment Fluid 5.9f 317 379 96 .0 477 101 376 [3]. [(8] Middle Valley Terrigenous sediment Fluid 5.15.58 274285 424 to 416 107 to 96 0 551 to 538 113 to 101 425-450 [6]s“[e] and Iheya North field Pelagic sediment Fluid 4.85.0 10 635 113 0 1008 120 888 53.9 Pelagic sediment [01] Iheya North field Fluid 4.85.0 60 430 132 0 562 142 421 54.0 [01] Basalt-hosted 21°N EPR Basalt Fluid 350 401 to 396 126 to 102 512 to 504 135 to 108 370405 [11], [12], [8] Continental systems Matsukawa (Japan) Volcanic gas Andesite + Gas 220260 466 to 364 410 to 274 80 to 74 548 to 373 451 to 241 36166 [13] Zao (apan) welded tuff Pyroxene andesite Gas 96 370 181 77 401 139 262 [13] Issaikyo (Japan) Gas 94 486 378 70* 403 191 (2014) 112 Nasu Oana (Japan) Pyroxene andesite [1] Pyroxene andesite Gas 112 515 to 488 430 to 400 62 to 59" S09 01 E99 501 to 447 159-161 Silicic peralkaline [13] Socorro Island Gas 5699 580 to556 117 to 70 53 to 12" 820 to 765 112 to 21 708791 [14] (Mexico) rock 125 Shield borehole gas Pori region Sedimentary rock Fluid <1015 659 245 46 1076 18z 795 48.5 [15], [16] (Finland) Juuka region Serpentinite lense Fluid 649 to 619 284 to 281 8 1039 to 957 326 to 322 L1129 29.0 to 28.6 [16] Sudbury. (Finland) Quartz diorite + Gas 637 to 138 335 to 274 64 to 47 965 to 82 345 to 259 259620 35.0 to 29.7 [17] Ontario serpentinite dike (Canada) [1] Fritz et al. (1992): [2] Sano et al. (1993): [3] Proskurowski et al. (2006): [4] Lang et al. (2012): [5] Proskurowski et al. (2008): [6] Charlou et al. (2002): [7] Seyfried et al. (2003): [8] Von Damm (1990): [9] Butterfield et al. (1994): [10] Kawagucci et al. (2011): [11] Welhan and Craig (1983): [12] Horibe and Craig (1995): [13] Kiyosu (1983): [14] Taran et al. (1992): [15] Sherwood Lollar et al. (1993a): [16] Sherwood Lollar et al. (1993b): [17] Sherwood Lollar et al. (1988). AssumeD-H2O=0%e AD-HO value of water vapor. K.Suda et al./Earth and Planetary Science Letters 386(2014)112-125 119 greater than the measured values (Fig. 3a), likely owing to inhibi- produced by hydrogenotrophic methanogens typically exhibits a tion of isotopic exchange between H2 and H2Oaq by slow reaction hydrogen isotopic composition that is not controlled by 8D-H2, a a r s n but depends mainly on 8D-H2O (Valentine et al., 2004). The lower cooling of the fluid compared to the rate of isotopic exchange reac- 8(CH4-H2Oaq) value relative to the equilibrium value is probably tions (Proskurowski et al., 2006). However, for volcanic gas (Kiyosu, due to the kinetic isotope effect during enzymatic reaction of the 1983; Taran et al., 1992) and shield gas (Sherwood Lollar et al., methanogens. Conversely, the enzyme hydrogenase efficiently cat- 1993b), the correlation between predicted and measured temper- alyzes isotopic exchange between H2 and H2O (Valentine et al., atures is relatively poor, probably because the gases are separated 2004); thus, 8(H2-H2Oaq) tends to approach the equilibrium value from the liquid water reservoir in these gas-rich/water-poor condi- of the growth temperature. These characteristics are consistent tions, prohibiting complete exchange with Hzo. In the case of the with the Happo #3 fluid. It is likely that the microbial methane is Happo spring, H2 can be isotopically re-equilibrated with H2O ow- o- 'so # o ing to the relatively rapid (i.e., compared to geologic timescales) tematics produced by FTT reactions result in a SD-CH4 value that H2-H20 isotopic exchange (Sherwood Lollar et al., 2007). Thus, it is nearly equal to the 8D-H2 value, indicating that little or no frac- is possible that the SD value of H2 may have been modified from tionation occurs when CH4 is produced from H2 (Fu et al., 2007; its original value. Taran et al., 2010; McCollom et al., 2010). The three experiments In contrast to the H2-H2O system, the measured fractionation with different catalysts and temperatures share a fundamental between CH4 and H2O shows little correlation with the measured characteristic in that hydrogen isotopic fractionation between the temperature (Fig. 3b). The CH4-H2Oaq equilibrium temperatures product CH4 and the reactant H2 is minimal (see Table 4). How- for Happo #1 and #3 are 183°C and 7°C, respectively. Both val- ever, the 8D-CH4 value of the Happo #3 methane is far from ues are far from the measured temperatures of 52°C and 48°C. the associated SD-H2 value (Fig. 4). Thus, the FTT origin of the Isotopic exchange between CH4 and HzO is known to be slower Happo #3 methane is questionable when compared with the pre- than that between H2 and HzO. Thus, the hydrogen isotopic com- vious experimental results (Fu et al., 2007; Taran et al., 2010; o n s McCollom et al., 2010). re-equilibrated with H2O. The Happo #1 methane exhibits less fractionation between In general, the equilibrium temperatures of 8D-CH4-H2Oaq CH4 and HzOaq than the equilibrium value, as seen in typical n n serpentinite-hosted hydrothermal systems (Fig. 4a). In serpentinite- mal systems (>200°C) (Fig. 4b). Conversely, lower temperature hosted hydrothermal systems with high H2 concentrations, FTT serpentinite-hosted systems can be categorized into two groups: synthesis reactions and methanogen metabolism may be the one exhibits less fractionation than the equilibrium fractionation most likely processes for methane production. However, the (i.e., Happo #1 and Lost City), whereas the other exhibits more CH4-H2-H2O systematics of Happo #1 exhibit considerable dif- fractionation than the equilibrium value (i.e. Happo #3 and Oman). ferences to those of FTT reactions or microbial methane (Fig. 4b). The latter cannot be explained by cooling of formerly equilibrated It is typically argued that the abiogenic methane in serpentinite- hosted hydrothermal systems is produced by FTT reactions (e.g., gating the heating of Ci-C5 n-alkanes in aqueous solution (323 °C Holm and Charlou, 2001), which produce CH4 and hydrocar- - (%o>） xa a o uos y ( 9- p bons using H2 as a source of hydrogen. However, the Happo #1 tween CH4 and water during experimental times of over 100 days fluid exhibits large differences between SD-CH4 and 8D-H2 val- (Reeves et al., 2012). In the case of the low-temperature Happo ues that is markedly contrast with previous experiments in- spring (50-60°C), isotopic exchange is expected to be much more vestigating FTT synthesis (Fu et al, 2007; Taran et al., 2010; sluggish. Thus, the hydrogen isotopic composition of the CH4 can McCollom et al., 2010). Thus, the Happo #1 methane could have be assumed to retain information relating to the original produc- been produced directly from HzO but not from H2. tion process. Previously, hydrogen isotope data for FTT synthesis had been 5.2.2. CH4-H2-H2O hydrogen isotope systematics et al., 2007; Taran et al., 2010; McCollom et al., 2010). At lower Fig. 4 illustrates the correlation between &(Hz-H2Oaq) and temperature, the isotopic fractionation may possibly differ from 8(CH4-H2Oaq) values. The 8(CH4-H2Oaq) value for Happo #1 lies the laboratory experiments owing to differences in kinetic effects. above the isotopic equilibrium curve and is similar to the values However, the Happo CH4 is enriched in deuterium by 490%o rela- obtained for other serpentinite-hosted systems, excluding Oman. tive to co-existing H2. If the D-enrichment of CH4 were responsible In contrast, the &(CH4-H2Oaq) value for Happo #3 plots below the for kinetic fractionation, the reaction would have caused a revers equilibrium curve and is similar to that of Oman. This opposite iso- isotope effect (i.e., isotopically heavy H2 would have reacted faster topic relationship with respect to the equilibrium value suggests to yield CH4). No known kinetic process can cause this 490% that the Happo #1 and #3 methanes have different origins, or that reverse isotope effect from H2 to CH4, irrespective of catalyst, tem- secondary isotopic modification has overprinted differently for the perature, and time. two fluids. These two wells are only 150 m apart, although their On the other hand, the much longer reaction times in natu- ranges of sampling depths are different. Thus, it is likely that the ral systems could have resulted in greater fractionation between two different source fluids may have been mixed at a different CH4 and H2 than that in the experiments. Abiotic reactions accom- mixing ratio for each well. We argue that at least two different panying isotopic exchange in the H2-CH4-H2O system could have sources of CH4 are required to explain the observed &(CH4-H2Oaq) resulted in the observed D-enrichment of CH4 relative to H2. In relationship. this case, 8(CH4-H2Oaq) and 8(H2-H2Oaq) may gradually approach First, we discuss the origin of Happo #3 methane. A CH4-H2Oaq the equilibrium line given sufficient time (Fig. 4), although they fractionation larger than the equilibrium value has been re- may not be expected to crosscut the line (Fig. 4c). Thus, the ob- served small fractionation between CH4 and H2O is unlikely to drogenotrophic methanogens (Balabane et al., 1987; Valentine be attained if the CH4 is produced from co-existing H2, even for et al., 2004) and (2) experiments investigating abiotic CH4 syn- much longer reaction times. We cannot discard the possibility that thesis via FTT reactions (Fu et al., 2007; McCollom et al., 2010; H2 may have been an intermediate in the production of CH4 from Taran et al., 2010) (Fig. 4b). Both of these mechanisms are poten- H2O, although the primary source of hydrogen for both H2 and CH4 tial sources of the CH4 included in the Happo #3 fluid. Methane should be H2O in an aqueous system such as the Happo spring. 120 K.Sudaetal./EarthandPlanetaryScienceLetters386(2014)112-125 ε(H2-H2O) % -1600 -1400 -1200 -1000 -800 -600 -400 -200 (a) ** 类 100 ￥ +4306460 500 200 100 -200 lsotopic equilbium curve ￥* 「☆Happo #1 ★ 300 ★Happo #3 CH Lost City -400 OLogatchev ? ** L Oman -500 Seafloorhyo △Sediment-hosted Basalt-hosted 600 ￥Volcanic gas otinental +Shield borehole -700 H2/CO2 Methanoge -800 FTT exp. (Gas) FTT exp.(Aque -900 ε(H2-H2O) %0 1600 -1400 -1200 -1000 -800 -600 400 200 (b) 100 ☆ 300 400 500 200 100 -200 ￥ Isotopic equibrium curve -300 3 400 500 % 600 -700 -800 -900 E(H2-H2O) %0 -1600-1400-1200 -1000-800 -600 400 200 (c) AbioticCH4productiondirectlyfromH2O 100 PS7 200 100 -200 Isotopic equilibium curve 口口 ★ 0C 300 网 400 Biogenic -500 FTTsynthesis 600 -700 -800 -900 Fig. 4. CH4-H2-H2O hydrogen isotope systematics: correlation between observed &(H2-H2Oaq) and e(CH4-H2Oaq) values. The epsilon (E) value is defined as & = 1000 Inα. The gray belt represents the range of previously estimated hydrogen isotopic exchange equilibrium curves based on both experiments and theoretical calculations (Horita and Wesolowski, 1994; Bardo and Wolfsberg, 1976; Cerrai et al., 1954; Suess, 1949; Horibe and Craig, 1995). The bold line is the curve recommended by Horibe and Craig (1995). The thin line indicates the CHa-H2 equilibrium fractionation at a given temperature. Open and flled stars represent Happo #1 and #3, respectively. The other hydrogen isotope data sources are listed in Table 3.(a) Happo data are compared with hydrogen isotope compositions observed in natural environments. (b) Happo data are compared with results from laboratory experiments of biological and abiological methanogenesis.“"H2/CO2 Methanogens" refers to an incubation experiment investigating CO2-reducing methanogens (Balabane et al., 1987; Valentine et al., 2004). “FTT exp. (gas)" and “FTT exp. (aqueous)” refer to experiments investigating abiotic methane production via FTT synthesis in the gas phase (Taran et al., 2010) and aqueous phase (Fu et al, 2007; McCollom et al, 2010), respectively. (c) Interpretation. Each dashed zone indicates the area within which methane produced by each mechanism is expected to fall Abiotic CH4 production directly from Hz0. Biogenic: biological methane generation by methanogens. FTT synthesis: abiotic methane synthesis via FTT reaction. Table 4 Hydrogen isotope datasets from laboratory experiments: pure cultures of the CO2 reduction methanogens and abiotic synthesis of CH4 by FTT reaction. Methanogens incubation experiment Exp. ID Organism Substrate System type Temp. 1000lnα 1000lnα 1000lnα Reference H2-H20 CH4-H20 CH4-H2 K (°℃) Suda A Methanobacterium formicicum H2/CO2 34 -1353 to -1330 -568to-537 762-817 Balabane et al. (1987) B Methanobacterium formicicum H2/CO2 closed 34 -1356 to -1316 -481 to -470 835-887 Balabane et al. (1987) C Methanobacteriumformicicum H2/CO2 closed 34 -1340 -524 816 Balabane et al. (1987) D Methanobacteriumformicicum H2/CO2 closed 34 -1359 to -1306 -513to-488 818-852 Balabane et al. (1987) E Methanobacterium formicicum H2/CO2 34 -1369 to -1357 -499 to-496 861-870 Balabane et al. (1987) D-1 Methanothermobacter marburgensis H2/CO2 open 65 -1121 to -1110 -371 to-360 750 Valentine et al. (2004) Fischer-Tropschtype(FTT)experiment Exp. ID Catalyst C-source H-source Temp. 1000Inα 1000Inα 1000Inα Reference (5。) H2-H20 CH4-H20 CH4-H2 Aqueous phase Exp. CO#1 Fe co Fe + HCI → H2 250 -865to-595a -825 to -767a 179-83 McCollom et al. (2010) Exp.CO#3 Fe CO Fe + HCI → H2 250 -812to-587b -790 to-750b 162-22 McCollom et al. (2010) Exp. #1(HCOOH) magnetite HCOOH HCOOH 400 -388 to -382 -215 to-214 168-174 Fu et al. (2007) 8 Exp.#2(H2/CO2) magnetite CO2 gas H2 gas 400 -389 -229 160 Fu et al. (2007) (2014)112- Exp.#3(H2/CO2) magnetite CO2 gas H2 gas 400 -402to-398 -228 to-214 174-184 Fu et al. (2007) Gas phase FTCO2-Fe1 Fe CO2 gas H2 gas 350 -379 -411 Taran et al. (2010) FTCO2-Fe2 Fe CO2 gas H2 gas 350 -352 -387 -35 Taran et al. (2010) 1.5 FTCO2-C01 Co CO2 gas H2 gas 245 -563 -439 124 Taran et al. (2010) FTCO2-C02 CO2 gas H2 gas 245 -535 -422 113 Taran et al. (2010) a Assuming that 8D-H2O is attained isotopic exchange equilibrium with H2 at 68 h of experiment time. b Assuming that 8D-H20 is attained isotopic exchange equilibrium with H2 at 240 h of experiment time. 12 122 K.Suda et al./ Earth and Planetary Science Letters 386(2014)112-125 Volcano distribution in the SW Japan arc. Contrasting active volcano spacing is a characteristic in the H20 be investigated in future experiments. The 813C-CH4 value of -34.5%o for Happo #1 was found to CH4 be depleted in i3C by roughly 20%o relative to the values typi- cally observed in seafloor ultramafic-hosted systems such as Lost City (813c = -13.6% to -9.4%o; Proskurowski et al., 2008). In % 8 a continental setting, isotopically depleted CO2 can be available in groundwater owing to respiration, thus providing a source of 13C-depleted methane. Similarly 13C-depleted methane has been found in other continental systems, such as the Othrys ophio- lite (813C-CH4 = -36.8%o to -27.0%o; Etiope et al., 2013a). Thus, H2 the differences in 813C-CH4 between seafloor and continental set- tings may be due to carbon isotopic variations of the initial Time able for CO2 owing to the low CO2 concentrations at Hakuba Fig. 5. A schematic showing the temporal evolution of 8D values for H2, CH4, and H20. The 8D-HzO value was assumed to be constant. Deuterium-enriched CH4 is Happo. Moreover, the carbon source and carbon isotope fraction- produced from heavy H2; subsequently, H2 is re-equilibrated with H20. dotite hydration reaction occurred, CO2 should have been present Four possible mechanisms can explain the hydrogen isotope when the CH4 was formed (probably before the pH increased compositions with the 8(CH4-H2Oaq) higher than the equilibrium value: (1) methane formation directly from H2O, (2) methane pro- Figure 2 shows the along-arc distribution of Quaternary volcanoes and the volume of each volcano. It should duction from D-enriched H2, (3) a high-temperature origin of CH4 and H2 with subsequent isotopic re-equilibrium during cooling, methane. and (4) microbial secondary modification of the stable isotope ra- tio of CH4. 5.2.2.2. Methane production from D-enriched H2 The observed high 8(CH4-H2Oaq) value may be explained by the original production 5.2.2.1. Methane formation directly from H2O Methane from Hakuba that the contrasting volcano spacing in the SW Japan arc during Holocene has been continued from 2.6 Ma. isotopically with HzO (Fig. 5). In this case, the D-enriched CH4 tion at low temperatures (about 50 °C). As described in Section 5.1, could have been produced from H2 with minimum isotope ef- the primary 8D-CH4 value is likely preserved if methane produc- fect. After production of CH4, the remaining H2 may have been tion occurs at low temperatures owing to sluggish CH4-H2O iso- isotopically modified by isotope exchange with H2O. However, topic exchange at 50°C. The 8(CH4-H2O) values from Happo #1 such production from original D-enriched H2 is unlikely. Assuming and the majority of serpentinite-hosted samples are relatively that no isotopic fractionation occurs between CH4 and H2 during constant (—157%o to -109%o), suggesting that 8D-CH4 depends methane production, the original 8(H2-H2Oaq) value must be at n s ( ) a o ou o-a least higher than -160%o to explain the fact that 8(CH4-H2Oaq) CH4 production from HzO with relatively small fractionation be- > -160%o. The small 8(H2-H2Oaq) value would require high tem- tween CH4 and H2O is possible. One possible reaction in this peratures that are inconsistent with the water temperature of regard is the hydration of olivine or orthopyroxene that is asso- the Happo spring (50-60 °C). Isotopic fractionation between gen- ciated with serpentinization in the presence of a carbon source, erated H2 and H2O during low-temperature serpentinization is as described by the following reactions (Oze and Sharma, 2005; largely unknown. Using radiolytic H2 as an example of abio- Miura et al., 2011): genic H2, laboratory experiments have demonstrated that the H2-H2O fractionation that occurs during H2 production by radi- 24(Mg, Fe)2SiO4 +26H2O + CO2 olytic decomposition of water at room temperature is between Olivine -723%o and -383%o (Lin et al., 2005). These 8(H2-H2Oaq) val- > 12(Mg, Fe)3Si2O5(OH)4 + 4Fe3O4 + CH4 ues are higher than the equilibrium values at the experimen- Serpentine Magnetite tal temperature but are still far from the values required, i.e. &(H2-H2Oaq) > -160%o. Thus, as far as low-temperature origins 36MgSiO3 + 12FeSiO3 +3H2O + CO2 are concerned, it is difficult to ascertain a source for the extremely Orthopyroxene D-enriched H2. → 12Mg3Si4O10(OH)4 + 4Fe304 + CH4 5.2.2.3. High-temperature origin of CH4 and H2 with subsequent iso- Talc Magnetite topic re-equilibration during cooling The Happo methane and hy- Methane production in this manner is thermodynamically favored drogen could have been produced originally at much higher at low temperatures, i.e., below 300°C (Oze and Sharma, 2005). In temperatures than those measured for the hot spring water such cases, H2O provides a source of hydrogen for methane for- (~50°C). The hydrogen isotopes of CH4-H2-H2O systems are as- mation that is compatible with the observed isotopic relationships serpentinized mantle. As indicated in Fig. 2, however, the volcanic gap may not be so clearly observed when (Fig. 4c). Previous laboratory experiments have reported CH4 to the life span of arc composite volcanoes may be several hundreds of kilo years. In order to examine the linkage be yielded in association with H2 generation and increasing pH ensure isotopic exchange. If cooling of the fluid were to occur during the interaction between olivine (or peridotite) and water more rapidly than the rate of hydrogen isotopic exchange reac- in the presence of a carbon source (Janecky and Seyfried, 1986; tions, the CH4-H2-H2O system would be far removed from that Jones et al., 2010). Hydrogen isotopic fractionation during CH4 for- at equilibrium. It has been shown empirically that the H2-H20 mation through the hydration of olivine or orthopyroxene is largely exchange reaction occurs more rapidly than that of CH4-H2O. unknown, although it is expected that the 8D-CH4 value would be Lécluse and Robert (1994) determined rate constants experimen- controlled primarily by 8D-H2O because H2O-derived hydrogen is incorporated into CH4. The &(CH4-H2Oaq) value might be similar to the equilibrium value or slightly higher owing to kinetic effects K.Suda et al./Earth and Planetary Science Letters 386(2014)112-125 123 constants, the time required to approach 99% of isotopic equilib- Happo water suggests that H2 is produced by serpentinization at S o o- S o n 1 relatively low temperatures (~50°C). In Happo #1, the observed for H2-CH3D at 50°C (Giggenbach, 1982). The exchange between high deuterium enrichment of CH4 relative to H2 and the low CH4 and H2O is believed to be slower than or equal to the CH4-H2 levels of isotopic fractionation between CH4 and HzO relative to paleo-position of the KPR at 3 Ma could be reconstructed based on the current motion of the PHS plate (Fig. 3); isotopic equilibrium relationships suggest that the methane in this curs at 50 °C, 8(CH4-H2Oaq) may plot above the equilibrium curve, system is likely produced from H2O but not from H2. The hydra- as observed in Happo #1 (Fig. 4). However, high-temperature ori- gins for CH4 and H2 are somewhat problematic. The extrapolated is one possible mechanism for the production of the Happo #1 geothermal gradient suggests a temperature of 200 °C at approxi- methane. Since Happo #1 exhibits a similar CH4-H2-H2O hydro- mately 3.8 km depth. It is unlikely that meteoric water penetrated gen isotopic relationship to those observed in other serpentinite- to this depth and then upwelled to depths shallower than 1 km in hosted hydrothermal systems, it can be concluded that abiotic the Happo hydrological cycle. CH4 production directly from H2O (without mediation by H2) may Alternatively, CH4 and H2 can be produced by the thermal de- be more common in serpentinite-hosted systems, regardless of composition of organic matter. The SD value of the Happo CH4 whether such systems are continental or on the seafloor. However, (-300%o to -210%o） lies within the typical range for thermo- this hypothesis remains speculative at this stage. Further experi- last 14 my, the slab depth beneath these forearc or near trench magmatic belts would have been<50 km, much mental study at lower temperatures will be required to confirm -33.8%o) is somewhat heavy compared to that of typical thermo- the occurrence of abiotic CH4 production through hydration reac- genic CH4. Thermal processes commonly yield not only CH4 but tions. also other hydrocarbons (such as C2H6 and C3Hg), yielding high C2+ concentrations compared to those of CH4 (Berner et al., 1993). Acknowledgements The C1/C2+ ratios for the Happo gas ranged from 387 to 727; these We thank Mr. Sejima (Happoone Development Company) for 102; e.g., Whiticar, 1999). Moreover, the high H2/CH4 ratio of the his cooperation and assistance in the field study. We also thank Happo gas is in stark contrast to that obtained through thermo- the managers of the Kurashita, Renge, Wakakuri and Tsugaike hot genic processes (H2/CH4 < 1) based on a previous hydrothermal springs for their cooperation. This research was supported by a experiment using natural seafloor sediments (Seewald et al., 1994). grant for the Global COE Program entitled “From the Earth to Therefore, the observed low hydrocarbon and high H2 concentra- "Earths"", MEXT, Japan. Y.U. is supported partly by the Funding n e s Program for Next Generation World-Leading Researchers from the methane. Japanese Society for Promotion of Science (JSPS). Magmatic input is also unlikely to have affected the Happo CH4 considerably. The hydrogen and oxygen isotope compositions Appendix A. Supplementary material of the water samples indicate that the Happo hot spring water is almost meteoric in origin and entirely distinct from magmatic water (8180 = +6 to +8%o, 8D = -20 to -30%o) (Sakai and Mat- Supplementary material related to this article can be found on- suhisa, 1996). The relatively high 3He/4He ratio (4.5 RA/R) found line at http: //dx.doi.org/10.1016/j.epsl.2013.11.001. for the Hakuba Happo hot spring suggests some contribution of high-resolution thermal structure suitable for examining the behavior of water in the subducted slab and the References cally present in low concentrations (i.e., on the order of ppm) in magmatic gases (Kiyosu, 1983). Therefore, we believe the contri- Abrajano, TA., Sturchio, N.C., Bohlke, J.K, Lyon, G.L., Poreda, RJ, Stevens, C.M., 1988. Methane-hydrogen gas seeps, Zambales Ophiolite, Philippines: deep or shallow bution of magmatic CH4 to Happo water to be minimal. origin?. Chem. Geol. 71, 211-222. Balabane, M., Galimov, E., Hermann, M., Letolle, R, 1987. Hydrogen and carbon iso- 5.2.2.4. Microbial secondary modification of stable isotope ratio of CH4 tope fractionation during experimental production of bacterial methane. Org. Biological methane consumption can modify the isotopic compo- Geochem. 11, 115-119. sition of CH4. During methane oxidation, the remaining methane Bardo, R.D., Wolfsberg, M., 1976. A theoretical calculation of equilibrium con- I# oddeh au h 'a pue oei yoq u Aaissaioid paua s! stant for isotopic exchange reaction between HzO and HD. J. Phys. Chem. 80, 1068-1071. methane was produced originally by methanogenesis, as proposed Berner, U., Scheeder, G., Panten, D., Hufnagel, H., Faber, F, 1993. Primary crack- for Happo #3, subsequent oxidation could explain the observed ing of Type-I/Il and Type-Ill kerogens - kinetic-models of isotope variations in high SD-CH4 value of #1. However, we found almost no differ- methane, ethane, and propane. Abstr. Pap. - Am. 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