Geochimica et Cosmochimica Acta, Vol. 68, No. 3, pp. 573–589, 2004
Copyright © 2004 Elsevier Ltd
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Pergamon
doi:10.1016/S0016-7037(00)00462-9
Carbon isotopes and petrography of kerogens in ⬃3.5-Ga hydrothermal silica dikes in the
North Pole area, Western Australia
YUICHIRO UENO,1,3,* HIDEYOSHI YOSHIOKA,2 SHIGENORI MARUYAMA,3,† and YUKIO ISOZAKI1
1
Department of Earth Science and Astronomy, University of Tokyo, Meguro Tokyo 153-8902, Japan
2
Department of Chemistry, Tokyo Metropolitan University, Hachioji, Tokyo 192-0397, Japan
3
Department of Earth and Planetary Sciences, Tokyo Institute of Technology, Meguro, Tokyo 152-8551, Japan
(Received February 12, 2003; accepted in revised form June 24, 2003)
Abstract—More than 600 specimens of ⬃3.5 Ga-old hydrothermal silica dikes from the North Pole area,
Pilbara craton, Western Australia, have been studied petrographically. The kerogens in 44 samples have been
analyzed isotopically (C and N) and chemically (C, N, and H). The silica dikes are composed mainly of
fine-grained silica (modal abundance: ⬎97%) and are classified into two types by minor mineral assemblages:
B(black)-type and G(gray)-type. The B-type silica dikes contain kerogen (0.37 to 6.72 mgC/g; average 2.44
mgC/g, n ⫽ 21) and disseminated sulfides, dominantly pyrite and Fe-poor sphalerite. In some cases, carbonate
and apatite are also present. Their silica-dominated and sulfide-poor mineral assemblages suggest precipitation
from low-temperature reducing hydrothermal fluid (likely 100 –200°C). On the other hand, the G-type silica
dikes are sulfide-free and concentrations of kerogen are relatively low (0.05 to 0.41 mgC/g; average 0.17
mgC/g, n ⫽ 13). They typically contain Fe-oxide (mainly hematite) which commonly replaces cubic pyrite
and rhombic carbonate. Some G-types occur along secondary quartz veins. These textures indicate that the
G-type silica dikes were formed by postdepositional metasomatism (oxidation) of the B-types, and that the
B-types probably possess premetasomatic signatures. The ␦13C values of kerogen in the B-types are ⫺38.1 to
⫺33.1‰ (average ⫺35.9‰, n ⫽ 21), which are ⬃4‰ lower than those of the G-types (⫺34.5 to ⫺30.0‰;
average ⫺32.2‰, n ⫽ 19), and ⬃6‰ lower than bedded chert (⫺31.2 to ⫺29.4‰; average ⫺30.5‰, n ⫽ 4).
This indicates the preferential loss of 12C during the metasomatism (estimated fractionation factor: 0.9985).
Considering the metasomatic effect on carbon isotopes with probably minor diagenetic and metamorphic
overprints, we conclude that the original ␦13C values of the kerogen in the silica dikes would have been
heterogeneous (⬃5‰) and at least some material had initial ␦13C values of ⱕ ⫺38‰. The inferred
13
C-depletions of organic carbon could have been produced by anaerobic chemoautotrophs such as methanogen, but not by aerobic photoautotrophs. This is consistent with the estimated physical and chemical condition
of the hydrothermal fluid, which was probably habitable for anaerobic and thermophilic/hyperthermophilic
chemoautotrophs. Alternatively, the organic matter may have been possibly produced by abiological reaction
such as Fischer-Tropsch Type (FTT) synthesis under the hydrothermal condition. However, the estimated
condition is inconsistent with the presence of the effective catalysts for the FTT reaction (i.e., Fe-Ni alloy,
magnetite, and hematite). These lines of evidence suggest the possible existence of biosphere in the ⬃3.5 Ga
sub-seafloor hydrothermal system. Copyright © 2004 Elsevier Ltd
Woese, 1987; Stetter, 1998). To understand the role of hydrothermal systems for the origin of life and/or subsequent early
evolution, it is necessary to investigate the ancient hydrothermal systems preserved in the Early Archean terrane (before 3.0
Ga).
Biologic activities in the Early Archean have been inferred
mainly from 13C-depleted sedimentary organic matter (e.g.,
Mojzsis et al., 1996; Rosing, 1999; Schidlowski, 2001) and
bacterial microfossils (e.g., Awramik et al., 1983; Schopf,
1993). Although biological origins of some ⬃3.5 Ga microfossils and 3.8 Ga graphite are still a matter of debate (Buick,
1990; Brasier et al., 2002; Fedo and Whitehouse, 2002; Van
Zuilen et al., 2002), the existence of life in the Early Archean
has been a plausible hypothesis (Hayes et al., 1983; Rosing,
1999; Schidlowski, 2001; Mojzsis and Harrison, 2002).
The North Pole area, Western Australia, is well-known for
the occurrence of the Earth’s oldest (⬃3.5 Ga) microfossils in
bedded chert (Awramik et al., 1983; Schopf and Walter, 1983).
In this area, numerous hydrothermal silica dikes intruded into
pillowed basaltic greenstones below the fossil-bearing chert
beds (e.g., Nijman et al., 1999). Several independent field
1. INTRODUCTION
Hydrothermal systems are candidates for the birthplace of
life (e.g., Corliss et al., 1981; Holm and Andersson, 1998), and
are also candidates for sites of the first metabolic evolution
(e.g., Nisbet, 1995; Nisbet and Fowler, 1999). These ideas have
arisen for two reasons. One is that hydrothermal systems can
provide chemical potential for prebiotic organic synthesis (e.g.,
Shock, 1990; Shock and Schulte, 1998). The system would
have also provided various metastable chemical species for
primitive chemoautotrophic metabolisms (e.g., Shock et al.,
1995; McCollom and Shock, 1997). The other is a model of
microbial evolution based on the sequence comparison of the
ribosomal RNA gene (e.g., Woese, 1987). The model implies
that the last common ancestor was hyperthermophile, which
would adapt to high temperature environment (80 –120°C;
* Author to whom correspondence should be addressed
([email protected]).
†
Present address: Institute for Geo-Resources and Environment National Institute of Advanced Industrial Science and Technology Central
7, Higashi, Tukuba, 305-8567, Japan.
573
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Y. Ueno et al.
mappings and detailed field observations (Isozaki et al., 1997;
Nijman et al., 1999; Ueno et al., 2001a; Van Kranendonk et al.,
2001) suggested that the silica dikes were precipitated from
silica-rich hydrothermal fluid during the deposition of the ⬃3.5
Ga chert. The North Pole area is one of the best fields for
investigating the ancient hydrothermal system on the early
Earth.
The silica dikes contain considerable amounts of organic
matter (kerogen) as well as possible microfossils with 13Cdepleted isotopic compositions (Ueno et al., 2001a). Understanding the origin of the kerogen may provide important
implications for the origin of life and its earliest evolution,
because the organic compounds must have been produced
either by biological carbon fixation or by prebiotic organic
synthesis. However, detailed geological, petrological, and geochemical information of the kerogen in the silica dikes has so
far not been published. Here we report the result of isotopic (C
and N) and elemental (C, H, and N) analyses of the kerogen
together with petrography of the silica dikes. The detailed
analyses and observations enable us to recognize two distinct
types of silica dikes: One has been suffered from metasomatism, and the other has preserved premetasomatic signature
both for mineralogy and for carbon isotopic composition of the
kerogen. The results provide an important information to estimate the depositional environment of the silica dikes and the
original carbon isotopic compositions of the kerogens. Based
on these estimations, we finally discuss the origin of the kerogen.
2. GEOLOGICAL OUTLINE OF NORTH POLE AREA
The North Pole area is located in the central part of the
Pilbara Granite-Greenstone Terrane, Western Australia (Fig.
1). In the North Pole area, the lower part of the Warrawoona
Group crops out (Van Kranendonk et al., 2001), and consists of
ca. 6-km thick basaltic greenstones intercalated by bedded chert
in three horizons (Fig. 1). The lowermost chert unit is 1 to 70-m
thick and is intercalated with several barite beds of 0.1 to 5 m
thickness. This chert unit corresponds to the “Chert-Barite
Unit” previously described by Buick and Dunlop (1990). The
other two chert units are thinner (1–13 m) than the Chert-Barite
Unit, and scarcely associated with barite.
The precise age of the Chert-Barite Unit has never been
determined directly. Zircon U-Pb dating yields an age of 3458
⫾ 2 Ma for the felsic volcanics that overlie the cherts and
greenstones in the North Pole area (Thorpe et al., 1992b). A
model lead age of 3490 Ma (Thorpe et al., 1992a) was obtained
for galena from the Chert-Barite Unit. This may represent the
actual depositional age of the Chert-Barite Unit.
In the North Pole area, numerous (⬎2000 identified) silica
dikes characteristically intruded into pillowed basaltic greenstones (Figs. 1–3). They are 0.3–20 m wide and generally ⬎100
m long, with the longest one over 1 km (Fig. 4). The dikes are
massive and are composed mainly of chert-like fine-grained
silica (⬍10 ␮m). Some silica dikes show symmetrical pattern
along the dike axis, and sometimes has agate at the center (Figs.
4B,C), in which the silica shows fan-shape structure grown
from the hanging walls toward the center of the dike. This
suggests the precipitation of silica from a hydrothermal fluid.
The silica dike is distinguished from the bedded sedimentary
chert by the discordant relationships to adjacent strata and by
lack of internal bedding. These characteristics are also distinct
from the less common fissure-filling chert (Ueno et al., 2001b),
which has laminations parallel to those of the adjacent chert
beds, and is interpreted to have been formed by infiltration of
fine-grained clastics and sedimentary silica into an open fissure.
In addition, barite veins with 0.1 to 2-m widths also intruded
into basaltic greenstones. In the North Pole area, the distribution of barite veins is laterally discontinuous and is generally
restricted to the uppermost ⬃300 m of the pillow lava below
the Chert-Barite Unit, whereas silica dikes occur up to ⬃1000
m below the unit. Both the silica dikes and the barite veins
intrude into chert beds of the Chert-Barite Unit, but do not cut
through the entire unit, nor into the overlying pillow basalt. The
tops of the silica dikes show gradual transition into certain chert
beds, forming a clear T-junction. These relationships suggest
the silica dikes were formed intermittently during the deposition of chert beds of the Chert-Barite Unit (Isozaki et al., 1997;
Nijman et al., 1999; Ueno et al., 2001a; Van Kranendonk et al.,
2001).
3. SAMPLE LOCALITIES
Figures 1, 2, and 3 show sample localities of silica dikes. 601
specimens were examined. Among them, 40 silica dikes and
additional 4 bedded cherts were selected for isotope and elemental analyses. They cover the entire North Pole area (Fig. 1)
and are mainly from “Dresser domain” (Fig. 2) and from
“Dolerite Creek domain” (Fig. 3) except for six samples.
The Dresser domain is one of the most extensively studied
areas (e.g., Isozaki et al., 1997). In the domain, more than 200
silica dikes of several generations penetrated into pillowed
basaltic greenstone below the Chert-Barite Unit (Figs. 2 and 4).
They occur along the north-block-down listric normal fault
(D1; Nijman et al., 1999), and the conjugate synthetic and
antithetic extensional faults developed in the hanging-wall (Fig.
2). These structures were overprinted by layer-parallel thrusting
(D2; Nijman et al., 1999) and subsequent rotation and highangle strike-slip faulting due to the doming. The bedding of the
chert beds dip 30 – 40° to the east. To cover the entire Dresser
domain, 21 silica dike samples and 4 bedded chert samples
were selected for the analyses.
In the Dolerite Creek domain, several generations of silica
dikes penetrated into pillowed basaltic greenstone (Fig. 3).
They are stratigraphically 700 to 1200 m below the bottom
surface of the Chert-Barite Unit. The intrusive dolerite cuts the
greenstone and the silica dikes. Within ⬃50 m from the contact
of the dolerite, silica dikes were recrystallized and are composed of coarse-grained quartz (50 –200 ␮m). To evaluate
secondary fractionations associated with the contact metamorphism, 13 silica dike samples were selected for the analyses.
4. PETROGRAPHY OF THE SILICA DIKES
4.1. Material
For the petrological and mineralogical studies, 601 silica
dike samples were cut into petrographic thin sections and
examined by optical microscopy under transmitted and re-
Carbon isotopes and petrography of kerogens
575
Fig. 1. Geologic map of the North Pole area showing localities of silica dike samples (modified from Kitajima et al.,
2001a). Filled and open circles indicate kerogen-bearing and kerogen-free silica dikes, respectively.
flected light. To determine mineralogical compositions, a scanning electron microscope (JEOL JSM5310) with X-ray analysis
system (Oxford Link ISIS), and an electron probe micro-analyzer (JEOL JXA-8800A) both at Tokyo Institute of Technology were used for qualitative and quantitative analyses, respectively.
A laser Raman spectrophotometer (JASCO NRS-2000) at
Tokyo Institute of Technology was additionally used for characterization of opaque minerals, especially carbonaceous material. Excitation was provided by the 514.5-nm line of a
continuous-wave 20-mW Ar⫹ laser. The microscope objective
used is 50⫻; thus the analytical spot size is ⬃5 ␮m. Each point
was scanned typically for 30 s from 200 to 3800 cm⫺1 at a
spectral resolution of ⫾1 cm⫺1.
576
Y. Ueno et al.
Fig. 2. Lithologic map of the Dresser domain showing carbon isotopic distribution of the kerogens in the silica dikes. Note
that the contact between bedded chert and underlying pillowed basalt represents the ancient sea-floor surface at the time,
thus the silica dikes seem to have developed in about the uppermost 1000 m of the Archean oceanic crust.
4.2. Results
The silica dikes are classified into two types by their minor
mineral assemblages (modal abundance: ⬍3%): (1) B(black)type, which contains sulfides (Fig. 5); and (2) G(gray)-type,
which is sulfide-free (Fig. 6). Mineralogies are summarized in
Table 1. Among the 601 specimens, ⬃90% (n ⫽ 541) are
G-type, and remaining 10% (n ⫽ 60) are B-type. Note that the
classification pertains only to hand specimens. A single silica
dike, typically 50-cm wide and ⬎100-m long, occasionally
contains both B- and G-type domains. In a few composite
specimens, G-types occur typically along secondary quartz
micro-veins, whereas B-types occupy the groundmass (Figs.
6D,E). The contact between them is transitional and does not
show a cross-cutting relationship (Figs. 6D,E).
4.2.1. B-Type Silica Dike
B-type silica dikes are black and composed mainly of finegrained silica (⬍10 ␮m; modal abundance ⬎97%) with minor
amounts of fine-grained disseminated sulfides (typically ⬍50
␮m; modal abundance ⬍1%) and carbonaceous materials (Fig.
5).
Carbonaceous materials are black and typically occur as
kerogen clots (Figs. 5B,G,H). Raman spectra of the kerogen
(Fig. 7) are characterized by broad, first-order peaks (full width
at half maximum: ave. ⫽ 71.2 cm⫺1, SD ⫽ 5.5 cm⫺1, n ⫽ 70
for disordered peak near 1350 cm⫺1; ave. ⫽ 52.6 cm⫺1, SD ⫽
4.8 cm⫺1 for ordered peak near 1600 cm⫺1), high D(disordered-)/O (ordered-) peak intensity ratios (ave. ⫽ 1.42, SD ⫽
0.14), and high D/O area ratios (ave. ⫽ 1.92, SD ⫽ 0.11).
These Raman spectral features are similar to those of the
Fig. 3. Lithologic map of the Dolerite Creek domain showing localities of silica dike samples and carbon isotopic compositions of the
kerogens in them. Broken lines indicate the limits of coarse-grained
silica dikes around the intrusive dolerite.
Carbon isotopes and petrography of kerogens
577
Fig. 4. Photographs of the hydrothermal silica dikes. A) Annotated photograph of the Dresser domain, viewed to the east,
showing silica dikes developed in the Archean subseafloor. Bedded chert (top of the ridge) dip 30 – 40° to the east. Dashed
line shows contact surface between chert bed and underlying pillowed basaltic greenstone. Note that this surface represents
the ancient seafloor surface at the time of chert deposition. Width of the photo is ⬃500 m. B) Photograph of a ⬃1-m wide
silica dike. The box indicates the figure C. C) Central part of the silica dike. Width of the photo is ⬃30 cm.
kerogen from lower greenschist facies metasediments
(Wopenka and Pasteris, 1993; Yui et al., 1996).
Sulfides include pyrite and often sphalerite with or without
lesser amounts of chalcopyrite, galena, and Ni-Co-sulfide. Representative chemical compositions of the sulfides are summarized in Table 2. Pyrites are typically cubic euhedral with some
hexagonal shapes, suggesting possible replacement of pyrrhotite. Sphalerites are often euhedral with some irregular shape.
Concentrations of Fe in the sphalerites are generally low (⬍1
FeSmol%; Table 2). Some sphalerites contain pyrite dots (Fig.
5H), suggesting exsolution from more Fe-rich sphalerite.
Additionally, the B-type silica dikes often contain euhedral
or subhedral rhombic carbonate (dolomite-ankerite; Fig. 5C)
and secondary iron oxide and hydroxide (mainly hematite/
goethite). Some B-types also contain authigenic apatite, finegrained Ti-oxide (rutile/anatase), clay minerals, and rarely secondary barite.
The B-type silica dikes are generally massive without bedding and lamination. Kerogen clots and sulfides are uniformly
distributed. The B-types often contain ellipsoidal voids, which
filled with pure silica without kerogen and sulfide (Figs.
5A,E,G). Some are brecciated and filled with a second generation of silica dike (Fig. 5E). In addition, they rarely include
angular rock fragments, specifically silicified basalt or recycled
B-type silica dike. These entire structures were often overprinted by quartz and rarely barite micro-veins.
4.2.2. Low-Temperature Hydrothermal Origin of the B-Type
Silica Dike
The large-scale structures of the silica dikes, including symmetrical growth with agate at the center as described above,
clearly suggest that the dikes were precipitated from silica-rich
hydrothermal fluid. The brecciated texture in some B-type
dikes suggests the repeated infiltration of the hydrothermal
fluid. Their silica-dominated (⬎97%) and sulfide-poor (⬍1%)
mineral assemblages suggest that the temperature of the hydrothermal fluid was sufficiently low to have limited dissolution
and transportation of sulfides. These mineral assemblages are
comparable with modern low temperature (⬍200°C) hydrothermal vent mineralization, but dissimilar to high temperature
(200 to 350°C) massive sulfide deposits (Hannington et al.,
1995). This indicates that the silica dike probably formed at a
temperature less than 200°C.
A low-temperature hydrothermal origin of the dike is also
inferred from low concentration of FeS in sphalerite (⬍1
mol%; Fig. 8) and from the pyrite-sphalerite-dominated sulfide
mineral assemblage. A pyrite ⫹ Fe-poor sphalerite assemblage
is stable under lower temperature and/or higher fS2 condition
relative to Fe-rich sphalerite and pyrrhotite. In addition, the
scarcity of chalcopyrite is consistent with a low-temperature
origin. In other ancient massive sulfides, deposits further away
from the inferred high-temperature upflow are commonly
578
Y. Ueno et al.
Fig. 5. Optical photomicrographs (A-C, E, G, and H) and back scattered electron images (D and F), showing B-type silica
dike. See text.
sphalerite-rich and chalcopyrite-poor (i.e., low copper-to-zinc
ratio; Ohmoto et al., 1983), mainly due to higher solubility of
sphalerite relative to chalcopyrite.
Figure 9 shows the stability fields of pyrite and pyrrhotite
and the approximate ranges of FeS contents in sphalerite on the
T-log fO2 space. Some pyrite and sphalerite in B-type silica
Carbon isotopes and petrography of kerogens
579
Fig. 6. G-type silica dikes (A-C) and a composite specimen (D and E). (A, B, and E) Optical photomicrographs. (C) Back
scattered electron image. (D) Cut-surface of silica dike (NP111). Box in (D) indicates the area of (E). See text.
dikes show hexagonal shape and pyrite dots, respectively.
These textures suggest the existence of former pyrrhotite and
more iron-rich sphalerite. Thus, the cooling of higher-temperature hydrothermal fluid was responsible for the formation of
the B-type silica dike (Fig. 9).
4.2.3. G-Type Silica Dike
G-type silica dikes are gray and composed mainly of finegrained silica (⬍10 ␮m) with minor amounts of carbonaceous
materials (Fig. 6). Carbonaceous materials are scarce in Gtypes relatively to B-types. Raman spectral features of G- and
B-type dikes are similar (Fig. 7). The G-type dikes often
contain Fe-oxide (hematite) and Fe-oxyhydro-oxide (mainly
goethite). Some contain fine-grained Ti-oxide (rutile/anatase)
and clay minerals.
The G-type silica dikes are typically massive and have no
lamination or bedding. Similar to the B-types, they also contain
ellipsoidal voids (Fig. 6A). Some of the iron oxides are rhombic, cubic, or hexagonal (Figs. 6B,C), suggesting replacement
of former carbonate, pyrite, and pyrrhotite/sphalerite, respectively. These entire structures are overprinted by quartz microveins and rarely by barite micro-veins.
580
Y. Ueno et al.
Table 1. Mineralogy of silica dikes.
Minor minerals
Sample
B-type silic dike
NP069
NP070
NP071
NP111B
NP167
NP417
NP598
T005
T013B
T035B
T060
T065
T066
T070B
T205
96NP452
96NP759
Pano E296a
Pano E299aB
Pano E327
Pano G291
G-type silica dike
NP072
NP073
NP111G
T002
T009
T010
T012
T017
T021G
T027
T033
T067
T068
T070G
T253
T255
96NPT38
Pano E299aG
Pano E300a
a
b
Grain
sizea
Major mineral
silica
CMb
Pyrite
Sphalerite
Other sulfide
Carbonate
Fe-oxide
c
c
c
f
f
f
f
f
f
f
f
c
c
f
f
f
f
f
f
f
f
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
—
—
—
—
—
—
—
O
—
—
—
—
—
—
—
O
O
O
—
O
—
O
—
O
—
—
—
—
O
—
O
O
—
—
—
—
O
—
O
—
O
—
—
—
—
O
O
—
O
O
O
O
—
—
—
—
O
O
O
—
—
O
O
O
O
O
—
—
O
—
—
—
—
O
O
—
—
O
—
—
—
—
—
—
c
c
f
f
f
f
f
f
f
f
f
f
f
f
f
f
f
f
f
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
O
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
O
O
O
O
O
O
—
O
O
—
O
O
O
O
O
O
O
O
O
Other
TiO2
Apatite
Barite
Clay mineral
Clay mineral
Apatite, TiO2
Barite
Clay mineral
Clay mineral
TiO2
Clay mineral
TiO2
TiO2
Clay mineral
Zircon
Grain size of silica: f ⫽ fine grained (⬍50 ␮m); c ⫽ coarse grained (⬎50 ␮m).
CM ⫽ carbonaceous matter (kerogen).
4.2.4. Metasomatic Origin of the G-Type Silica Dike
The mineral assemblage of the G-type dike is more oxidized
than that of the B-type dikes. Pseudomorphs after pyrite and
carbonate, which are important components of the B-type,
suggest that the G-type silica dike was probably formed by
metasomatism from the B-type. This is consistent with textural
similarities between them (e.g., uniformly scattered kerogen
clots and ellipsoidal voids) and with the lower content of
organic carbon.
Note that the metasomatism does not represent weathering,
because G-types often occur along the secondary quartz microveins (Figs. 6D,E). Infiltration of oxidized fluid was probably
responsible for the metasomatism (Fig. 9). The temperature of
the metasomatism is poorly constrained. However, homogenization temperatures of fluid inclusions in secondary quartz
veins in the study area are around 150°C (Kitajima et al.,
2001b). This may possibly represent the temperature of the
metasomatic fluid.
The metasomatism significantly changed the chemical composition. Organic carbon, sulfide, and carbonate were consumed during the metasomatism. Approximately 90% of the
silica dikes are classified into G-type, but B-types occur sporadically. Thus, the metasomatism seems to have been pervasive, but heterogeneous. There is no marked correlation between depth below the bottom surface of the Chert-Barite Unit
and frequency of the occurrence of B-type (Fig. 2).
4.2.5. Coarse-Grained Silica Dikes
Coarse-grained silica dikes occur only around igneous intrusions, suggesting their contact metamorphic origin (Fig. 3).
Carbon isotopes and petrography of kerogens
Fig. 7. Representative Raman spectra of carbonaceous materials in
B-type (T005; top), G-type (T027; middle), and coarse-grained (T065;
bottom) silica dikes. O-peak (near 1600 cm⫺1) and D-peak (near 1350
cm⫺1) are from carbonaceous material.
They consist of coarse-grained quartz (⬎50 ␮m, typically
⬃100 ␮m) and more ordered carbonaceous material (Fig. 7)
with or without sulfide (i.e., B- or G-types, respectively).
The Raman spectra of the carbonaceous materials (Fig. 7) are
characterized by narrower first-order peaks (full width at half
maximum: ave. ⫽ 37.8 cm⫺1, SD ⫽ 4.2 cm⫺1, n ⫽ 8 for
⫺1
⫺1
D-peak; ave. ⫽ 21.6 cm , SD ⫽ 1.7 cm
for O-peak), lower
D/O intensity ratios (ave. ⫽ 0.26, SD ⫽ 0.03), and lower D/O
area ratios (ave. ⫽ 0.45, SD ⫽ 0.05). These Raman spectral
features clearly indicate that the carbonaceous materials are
more graphitized those in the fine-grained silica dikes, and are
similar to those in garnet-biotite schists (Wopenka and Pasteris,
1993).
Sulfides mainly consist of pyrite and pyrrhotite, suggesting
recrystallization at a higher temperature than the fine-grained
silica dikes, which are pyrrhotite-free.
5. ISOTOPIC AND ELEMENTAL ANALYSES
5.1. Analytical Procedure
Hand specimens were first cut into slabs ⬃10 cm across for
removal of the weathered surface, then the slabs were crushed
581
Fig. 8. FeS contents of sphalerite in the B-type silica dikes.
into small chips ⬃5 mm across. The chips were ultrasonically
cleaned with distilled water and subsequently with ethanol.
Clean and vein-free chips were picked and crushed using agate
ball mil.
The kerogen was isolated from the powdered samples by the
following process. The powder samples were demineralized
using 6 N HCl at 60°C for 1hr, followed by 46% HF and
concentrate HCl (92:8) dissolution at 60°C for 80 min. The
demineralization process was conducted in the sealed Teflon
bottle with mechanical shaking to promote complete acid dissolution of the minerals and avoid formation of fluoride phases.
The residue was finally washed with distilled water and dried.
Elemental analyses and gases for isotopic analyses were prepared by combustion of the samples at 1000°C in a Carlo Erba
EA-1108 on line to a Finnigan MAT ␦-S mass spectrometer at
Tokyo Metropolitan University. Nitrogen, carbon, and hydrogen contents in the kerogen were measured by N2, CO2, and
H2O produced through the combustion, respectively. While
H2O produced was identified by comparison of retention time
in the chromatograph with that of standard organic material
[2,5-bis (5-tert-butyl-benzoxazol-2-yl) thiophene], some samples showed another peak, which was close to the H2O peak but
had different retention time. In the latter cases, we did not
calculate the hydrogen content, because they may include hydrous or hygroscopic minerals (Table 3). Hydrogen and nitro-
Table 2. Representative compositions of sulfides in silica dikes.
Pyrite
Mineral
sample
96NP452
S
54.01
Fe
41.50
Co
0.30
Ni
2.27
Cu
0.31
Zn
1.12
As
0.00
Pb
0.00
Total
99.51
FeS mol% in sphalerite
Sphalerite
Chalcopyrite
Polydymite
96NP452
PanoE296
PanoE296
96NP452
96NP452
PanoE296
96NP452
96NP452
53.87
43.69
0.18
0.95
0.00
0.15
0.03
0.00
98.88
53.76
44.32
0.34
0.26
0.00
0.11
0.00
0.04
98.83
53.04
42.93
2.00
0.25
0.00
0.29
0.01
0.03
98.54
33.46
0.13
0.01
0.00
0.01
65.92
0.00
0.00
99.52
0.23
33.48
0.12
0.01
0.00
0.00
65.37
0.03
0.00
99.00
0.21
33.04
0.22
0.00
0.01
0.00
65.34
0.00
0.11
98.73
0.39
34.28
29.74
0.02
0.00
33.10
0.01
0.01
0.01
97.15
41.94
1.60
5.07
48.43
0.05
0.00
0.02
0.00
97.11
582
Y. Ueno et al.
6. DISCUSSION
To discuss the origin of the 13C-depleted kerogen in the
hydrothermal silica dikes, we first evaluate the degree of postdepositional effects on their carbon isotopic compositions, then
discuss the significance of their isotopic compositions in detail.
6.1. Postdepositional Effect on Carbon Isotopic
Composition
Fig. 9. T-logfO2 diagram showing hypothetical cooling pass of the
formation of the B-type silica dike, and presumable alteration pass to
the G-type dike. The stability fields of hematite, magnetite, pyrite and
pyrrhotite are calculated from Helgeson (1969) and Barton and Skinner
(1979), assuming following buffer reaction: 2H2S ⫹ O2 ⫽ S2 ⫹ 2H2O
Approximate range of FeS contents in sphalerite are from Czamanske
(1974) and Hannington and Scott (1989).
gen contents below the minimum values of the standard reagent
were not also used because they were too small values to have
enough reliability. For isotope analysis of nitrogen gas, carbon
dioxide was removed by calcium oxide and sodium hydroxide.
The standard deviation of replicate runs is 0.03‰ for carbon
and 0.06‰ for nitrogen.
The ␦13C value of carbonate was determined by conventional method. Carbonate powder was reacted in anhydrous
phosphoric acid (␳ ⭌ 1.89 g/mL) at 50°C for 48 h to produce
CO2. Isotopic composition of CO2 was determined with a
Finnigan MAT ␦-S mass spectrometer at Tokyo Metropolitan
University. The analytical reproducibility of the ␦13C value,
based on replicate analyses of NBS19 standards, is better than
⫾0.1‰.
Isotopic compositions are reported relative to PDB for carbon and relative to air for nitrogen.
5.2. Results
The results of isotopic and elemental analyses are summarized in Table 3. The organic carbon contents of the B-type,
G-type, and bedded chert are 0.37 to 6.72 mgC/g (average 2.44
mgC/g, n ⫽ 21), 0.05 to 0.41 mgC/g (average 0.17 mgC/g, n ⫽
13), and 0.06 to 0.10 mgC/g (average 0.08 mgC/g, n ⫽ 2),
respectively. Their carbon isotope compositions (␦13C) of the
B-type, G-type, and bedded chert are ⫺38.1 to ⫺33.1‰ (average ⫺35.9‰, n ⫽ 21), ⫺34.5 to ⫺30.0‰ (average ⫺32.2‰,
n ⫽ 19), and ⫺31.2 to ⫺29.4‰ (average ⫺30.5‰, n ⫽ 4),
respectively. There is no clear correlation between the ␦13C
value of the kerogen and its depth (Fig. 2). The H/C ratios are
0.09 to 0.23 (average 0.15, n ⫽ 19) for the B-type, and 0.07 (n
⫽ 1) for the G-type. The N/C ratios in the B-type silica dikes
are less than 0.005 and their nitrogen isotopic compositions
(␦15N) are ⫺4.1 to ⫹4.0‰ (average ⫺0.6‰, n ⫽ 6). The ␦13C
value of carbonate carbon in a B-type silica dike (96NP452) is
⫺2.11‰.
Organic carbon in the silica dikes was affected by several
geothermal processes after its deposition: diagenesis, metamorphism, and metasomatism. Among these processes, the abovementioned metasomatism would have caused significant loss of
organic carbon (⬃90%; Figs. 10 and 11). Accordingly, it was
probably the most important process affecting carbon isotopic
abundances.
During the metasomatism, the organic carbon seems to have
been enriched in 13C by ⬃4‰. Both the range and the average
of the ␦13C values of the G-types (⫺34.5 to ⫺30.0‰; average
⫺32.2‰) are ⬃4‰ higher than those of the B-types (⫺38.1 to
⫺33.1‰; average ⫺35.9‰). Concentrations of organic carbon
in G-types (average 0.17 mgC/g) are about ten times lower than
those in B-types (average 2.44 mgC/g). Given the assumption
that the G-type dikes are products of alteration of the B-type
dikes, this indicates loss of ⬃90% of the organic carbon.
Similar differences in isotopic compositions and organic-carbon contents are also recognized within individual hand specimens (⬃10 cm across; Figs. 6D,E, 11). Hence, the preferential
loss of 12C during the metasomatism probably caused 13Cenrichment of organic carbon in the silica dikes. The preferential loss of 12C can be examined using a Rayleigh fractionation
model:
␦final ⫽ (␦initial ⫹ 1000) ⫻ f共␣⫺1兲 ⫺ 1000
(1)
where ␦final, ␦initial, f, and ␣ are final and initial ␦13C values of
organic carbon, fraction of remaining organic carbon, and fractionation factor, respectively. The values ␦final, ␦initial, and f are
known for the two composite specimens (NP111 and
PanoE299a), which contain both B- and G-type materials
within a small domain (Table 4). The fractionation factors (␣)
in these specimens can be calculated from:
␣ ⫽ 1 ⫹ log关共 ␦ G ⫹ 1000兲/共 ␦ B ⫹ 1000兲兴/log(CG/CB)
(2)
where Cs and ␦s indicate concentrations and ␦13C values of
organic carbon, respectively, and subscripts B and G designate
B- and G-type dikes. Calculated fractionation factors for the
two specimens agree well (␣ ⫽ 0.9985; Table 4). This value is
also consistent with apparent fractionation factor calculated
from the averaged ␦13C and TOC values of silica dike specimens in the Dresser domain (Table 4). The estimated ␣ value
would be consistent with a ⬃1.5‰ kinetic isotope effect associated with the process of carbon loss.
In addition to the metasomatism, other four processes may
have possibly affected carbon isotopic composition: 1. contact
metamorphism, 2. isotopic exchange with carbonate carbon, 3.
long-term thermal maturation, and 4. very early diagenesis.
First, the organic carbon concentrations of the course-
Carbon isotopes and petrography of kerogens
583
Table 3. Results of the isotopic and elemental analyses.
Sample
B-type silic dike
NP069
NP070
NP071
NP111B
NP167
NP417
NP598
T005
T013B
T035B
T060
T065
T066
T070B
T205
96NP452
96NP759
Pano E296a
Pano E299aB
Pano E327
Pano G291
G-type silica dike
NP072
NP073
NP111G
T002
T009
T010
T012
T017
T021G
T027
T033
T067
T068
T070G
T253
T255
96NPT38
Pano E299aG
Pano E300a
Bedded chert
97NPT-A012
97NPT-A009
97NPT-A007
97NPT-A018
␦13Corg ⫾ SD (‰)
␦15Norg ⫾ SD (‰)
H/Catm
N/Catm
1.45
2.85
1.14
1.65
1.75
1.89
0.37
2.40
2.43
1.19
6.62
2.04
1.54
5.64
1.18
2.63
6.72
2.62
2.83
1.82
0.48
0.05
0.11
0.06
0.23
0.23
0.08
0.15
0.19
0.20
0.21
0.11
0.09
—
0.12
0.18
0.20
0.13
0.14
0.15
0.18
—
—
—
—
—
—
—
—
0.005
0.004
—
—
—
—
⬍0.005
—
⬍0.005
—
0.004
⬍0.005
—
—
⫺34.51 ⫾ 0.04
⫺33.65 ⫾ 0.82
⫺32.33 ⫾ 0.06
⫺31.87 ⫾ 0.62
⫺32.63 ⫾ 0.01
⫺30.49 ⫾ 0.66
⫺31.04 ⫾ 0.08
⫺31.56 ⫾ 0.05
⫺32.59 ⫾ 0.02
⫺34.30 ⫾ 0.02
⫺32.02 ⫾ 0.01
⫺32.52 ⫾ 0.09
⫺32.92 ⫾ 0.01
⫺33.94 ⫾ 0.09
⫺29.95 ⫾ 0.76
⫺32.88 ⫾ 0.01
⫺30.27 ⫾ 0.17
⫺31.32 ⫾ 0.20
⫺31.27 ⫾ 0.15
0.18
0.15
0.14
⬍0.5
0.17
0.21
0.21
0.13
0.08
0.18
⬍0.5
0.17
⬍0.3
⬍0.4
—
0.17
0.05
0.41
⬍0.5
—
—
—
—
—
—
—
—
—
0.07
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
⫺31.16 ⫾ 2.23
⫺30.92 ⫾ 0.25
⫺30.39 ⫾ 0.01
⫺29.43 ⫾ 0.09
⬍0.07
0.10
⬍0.5
0.06
—
—
—
—
—
—
—
—
⫺37.16 ⫾ 0.49
⫺36.36 ⫾ 0.04
⫺36.48 ⫾ 0.22
⫺35.92 ⫾ 0.16
⫺36.36 ⫾ 0.06
⫺35.56 ⫾ 0.07
⫺35.36 ⫾ 0.01
⫺34.80 ⫾ 0.54
⫺34.56 ⫾ 0.18
⫺38.08 ⫾ 0.08
⫺36.33 ⫾ 0.02
⫺37.19 ⫾ 0.04
⫺35.44 ⫾ 0.04
⫺35.87 ⫾ 0.04
⫺37.13 ⫾ 0.43
⫺37.18 ⫾ 0.06
⫺35.88 ⫾ 0.08
⫺33.10 ⫾ 0.05
⫺34.15 ⫾ 0.04
⫺37.21 ⫾ 0.10
⫺34.30 ⫾ 0.06
⫺4.10 ⫾ 0.20
1.78 ⫾ 1.52
3.98 ⫾ 1.63
0.98 ⫾ 1.20
⫺3.37 ⫾ 0.01
⫺2.78 ⫾ 1.27
grained silica dikes, which suffered from the contact metamorphism (Fig. 3), are within the range of other fine-grained silica
dikes (Fig. 11). Their carbon isotopic compositions (⫺37.2 to
⫺35.4‰ for B-type; ⫺34.5 to ⫺33.7‰ for G-type) are within
the range of other fine-grained silica dikes (Figs. 3, 11, and 12).
Thus, the contact metamorphism would not have caused any
marked isotopic modification of the kerogen.
Second, isotopic exchange between organic carbon and carbonate carbon mediated by CO2-rich metamorphic fluid could
have possibly caused 13C-enrichment under metamorphic condition (Valley and O’Neil, 1981; Wada and Suzuki, 1982;
Schidlowski et al., 1983; Ueno et al., 2002). However, the
metamorphic grade of basaltic greenstones surrounding the
silica dikes is generally below the greenschist facies (⬍350°C;
Dunlop and Buick, 1981; Kitajima et al., 2001a,b). This tem-
TOC (mg/g)
perature is consistent with the thermal maturity of kerogen in
silica dikes deduced from Raman spectroscopic features (Fig.
7). Thus, the maximum temperature they have experienced
might have been 350°C. Under such low temperatures
(⬍350°C), isotopic exchange between organic carbon and carbonate carbon is minimal (Hoefs and Frey, 1976; Scheele and
Hoefs, 1992). Therefore, the isotopic exchange with carbonate
carbon is probably negligible.
Third, it has been recognized that long-term thermal maturation of kerogen would have caused 13C-enrichment associated with decrease of H/C ratio (e.g., Hayes et al., 1983; Des
Marais, 1997). This shift may have overprinted the metasomatic effect, because both B- and G-type dikes have low H/C
ratios (⬃0.15). However, this shifts would have been smaller
than those resulting from the metasomatism (⬃4‰). According
584
Y. Ueno et al.
␦13C values of organic matter decreased by only 1.6‰ relative
to the initial values, probably due to selective preservation of
13
C-depleted organic compounds.
Consequently, the kinetic loss of 12C by the metasomatism
would have significantly increased the ␦13C values of the
kerogen (ⱖ4‰). The other effects may have possibly increased
and decreased the ␦13C values of the kerogen, though these
effects would have been smaller than the metasomatic effect.
This assumption enables further discussions on the origin of the
kerogen in the silica dikes.
6.2. Origin of the Kerogen in the Hydrothermal Silica
Dike
Fig. 10. Histograms showing reduced carbon concentrations of bedded chert (top), G-type silica dike (middle), and B-type silica dike
(bottom). Numbers with triangles denote averages. Data from Hayes et
al. (1983) are also compiled for chert from the North Pole Chart-Barite
Unit.
to Des Marais (1997), matured kerogen with H/C ratio of
⬃0.15 would have been enriched in 13C by 2 to 3‰ from a
hypothetical precursor with H/C ratio of 1.5.
Finally, the observed isotopic compositions provide little
information about early diagenesis. In the modern environment,
the early diagenesis is often mediated by micro-organisms,
though it is unknown whether the precursor of the kerogen in
the silica dike was utilized by microbes. Even if the microbial
mediation is assumed, there is no reason to consider that early
diagenetic fractionations would have exceeded a few permil.
Recent experiments and in situ observations of very early
diagenesis by Lehmann et al. (2002) suggested that the bulk
The present petrographic, elemental, and isotopic analyses
clearly suggest that the silica dike originally contained abundant 13C-depleted organic carbon. The ubiquitous distribution
of kerogen in silica dike indicates that the organic matter was
present in the hydrothermal fluid at the time of silica dike
formation. Therefore, the following two explanations are possible for the presence of the organic matter: (1) in-situ production of organic matter in fissures in the seafloor basalt that
would have acted as conduits for the upwelling hydrothermal
fluid; and (2) infiltration of organic matter into the fissures from
adjacent sources during the hydrothermal circulation. However,
there is no evidence to support the latter possibility. The wall
rock of the silica dike is kerogen-free pillowed basalt (Figs.
1– 4). No sedimentary rock has been recognized beneath the
Chert-Barite Unit in the North Pole area (Fig. 1). Transportation of organic matter from the overlying sediments is unlikely,
because there is no systematic correlation between organic
carbon concentrations of the dikes and the depths below the
Chert-Barite Unit. Additionally, overlying chert beds contain
less organic carbon (⬍0.1 mgC/g) than the silica dikes, suggesting that they could not be a source of the kerogen.
Consequently, we prefer the former possibility. Possible
mechanisms to have produced the 13C-depleted organic matter
in the conduit of hydrothermal fluid are 1) carbon fixation by
autotrophic organisms, and 2) abiotic synthesis of organic
compounds. The significant 13C-depletion of the kerogen may
suggest the former mechanism, but the second possibility cannot be dismissed completely. It has been widely suggested that
hydrothermal systems were possible sites for prebiotic synthesis of organic compounds (e.g., Corliss et al., 1981; Ferris,
1992). In the following sections, we will test the above-mentioned two possibilities.
6.2.1. Biological Origin
Fig. 11. Relationship between organic carbon concentration (TOC)
and carbon isotopic composition (␦13Corg) of the kerogen in the silica
dikes and bedded cherts. Dashed lines tie the values from the same rock
samples. Doted lines represent Rayleigh-fractionation trajectories with
fractionation factor (␣) of 0.9985. The three lines labeled 1, 10, and 100
are started with an initial ␦13C value of ⫺38‰, and with initial
concentrations of organic carbon of 1, 10, and 100 mgC/g, respectively.
See text and Table 4.
To test the biological origin, we should first consider whether
the chemical and physical conditions of the North Pole hydrothermal system were suitable for autotrophic activities. Given
the dark subseafloor environment, photosynthesis is unlikely,
whereas chemosynthesis is possible. Modern submarine hydrothermal systems generally provide various electron donors and
acceptors for energy yielding reactions. This is mainly due to
mixing between reducing hydrothermal fluid and more oxidizing seawater. Considering the North Pole hydrothermal system,
utilization of O2 for electron acceptor is unlikely because of the
estimated low fO2 condition (Fig. 9). The pyrite-dominated
Carbon isotopes and petrography of kerogens
585
Table 4. Calculated carbon isotope fractionation factor (␣).
B-type
G-type
Sample
␦13Corg (‰)
TOC.(mg/g)
␦13Corg (‰)
TOC.(mg/g)
⌬B-G (‰)
f
␣
NP111
Pano E299a
Dresser domain silica
dike (average)
⫺35.92
⫺34.15
⫺35.51
1.65
2.83
1.99
⫺32.33
⫺31.32
⫺32.12
0.14
0.41
0.19
3.59
2.83
3.39
0.09
0.14
0.10
0.99848
0.99849
0.99851
mineral assemblages of B-type dikes imply very reducing condition with sufficient H2 gas for H2-dependent autotrophy.
Therefore, anaerobic chemoautotrophs (e.g., methanogens)
could have survived in the hydrothermal condition, whereas
aerobes (e.g., thiotrophs and methanotrophs) could not. Even
though the temperatures were relatively high (⬎100°C), some
modern hyperthermophiles thrive at temperature up to 120°C
(e.g., Stetter, 1998). The silica dike probably formed from a
low temperature hydrothermal fluid (100 to 200°C), because of
their silica-dominated (modal abundance: ⬎97%) and sulfidepoor (modal abundance: ⬍1%) mineral assemblages. Note that
even the estimated temperature of up to 200°C does not eliminate the possibility for biologic activity. The temperature of
the fluid in the conduit may have varied, because precipitation
of silica requires cooling and/or mixing of hydrothermal fluid
with lower temperature seawater. Consequently, physical and
chemical conditions of the hydrothermal system could have
allowed biologic activity, for example anaerobic, thermophilic/
hyperthermophilic, and chemoautotrophic organisms such as
methanogens.
The chemoautotrophic origin of the kerogen can be tested by
their carbon isotopic compositions. Assuming that the metasomatism caused the main postdepositional fractionation (Fig.
11), the concentrations and/or ␦13C values of the parent material of the kerogen must have been heterogeneous. Examining
the Figure 11, there are three possibilities: (1) all samples had
the same initial ␦13C values, with concentrations ranging up to
80mgC/g; (2) all samples had the same initial concentration,
but initial ␦13C values varied over more than 5‰; or (3) both
initial concentrations and ␦13C values varied. However, the first
possibility is implausible, because it requires unlikely high
initial concentration (80mgC/g). Thus, the initial ␦13C values
were probably heterogeneous, and at least some of the material
would have been ⱕ ⫺38‰, because the processes of alteration
generally leads to 13C-enrichment.
The inferred heterogeneity of the initial ␦13C values is not
consistent with enzymatically catalyzed process operating in an
open system, which is expected to have produced an isotopically uniform product. However, some autotrophs especially
methanogenic bacteria rapidly consume surrounding CO2 (e.g.,
House 2003). In such cases, the system tends to be closed,
thereby producing isotopically heterogeneous organic carbon.
Hence, the inferred heterogeneity of the initial ␦13C values
(⬃5‰) could be consistent with the chemoautotrophic origin
of the kerogen.
Further, the inferred 13C-depletion of the initial material is
more decisive. The isotopic composition of carbonate carbon in
one B-type silica dike is ⫺2.1‰. This indicates that the equilibrated dissolved CO2 had ␦13C values of about ⫺4‰ at 100°C
(Mook et al., 1974). Thus, the fractionation between initial
organic carbon and dissolved CO2 (i.e., ␧) would have been
over 34‰ for at least some material. If the organic carbon was
produced by autotrophic organisms, then this fractionation is
too large to have been produced via the Calvin cycle (␧ ⱕ
30‰; e.g., House, 2003; Fig. 13) utilized by aerobic photoautotrophs. On the other hand, the large fractionation could have
been produced via the reductive acetyl-CoA pathway (␧ ⱕ
42‰; e.g., House, 2003; Fig. 13), which is utilized by H2dependent chemoautotrophs such as methanogen and acetogen.
In fact, the large fractionations up to 36‰ have been demonstrated by some anaerobic, thermophilic, and chemoautotrophic
bacteria such as Methanobacterium thermoautotrophicum
(Fuchs et al., 1979). Hence, the inferred large fractionation is
fully consistent with the estimated condition of the subseafloor
hydrothermal system.
In contrast with the carbon isotopes, the nitrogen isotopic
compositions of the kerogens provide little information about
their origin, because the observed ␦15N values are widely
variable (⫺4.1 to ⫹4.0‰; average ⫺0.6‰). The average ␦15N
value near 0‰ is close to present and possibly Early Archean
atmospheric N2 (Beaumont and Robert, 1999). This could be
consistent with a system, which involves N2-fixation.
In summary, the estimated physical and chemical conditions
and the carbon isotopic compositions of the kerogen can be
explained by in-situ production of organic matter in the subseafloor hydrothermal system by anaerobic chemoautotrophs
such as methanogenic bacteria.
6.2.2. Abiotic Synthesis of Organic Compounds
Fig. 12. Relationship between carbon isotopic compositions
(␦13Corg) and H/C atomic ratio of the kerogen in the silica dikes.
An alternative possibility is that the kerogen was produced
by abiological reactions. Brasier et al. (2002) observed kerogen
586
Y. Ueno et al.
Fig. 13. (A) Summary of carbon isotopic compositions of the kerogen in the bedded chert (black), B-type (gray) and
G-type (oblique line) silica dike. Open columns show carbon isotopic compositions of carbonates both in silica dike (this
study) and in bedded chert (Hayes et al., 1983). The range of the inferred initial ␦13C values of the kerogens in the dikes
is also shown. (B) Carbon isotopic variations of modern autotrophic bacteria utilizing four different fixation pathways,
calculated from previous culture experiments for a CO2 source with ␦13C ⫽ ⫺4‰, which are equilibrated with dissolved
inorganic carbon of ⫺2‰ at 100°C (Mook et al., 1974). The data for bacterial fractionations are compiled from Calder and
Parker (1973), Fuchs et al. (1979), Holo and Sirevåg (1986), House (2003), Mizutani and Wada (1982), Pardue et al. (1976),
Preuß et al. (1989), Quandt et al. (1977), Ruby et al. (1987), Sirevåg et al. (1977), and Whitman et al. (1992). Isotopic
variations of abiological carbon are also shown. Reduced carbon in mantle rocks are compiled from Fuex and Baker (1973),
Sugisaki and Mimura (1994), and Watanabe et al. (1983). Non-carbonate graphitic carbon (acid residue) in carbonaceous
chondrites are compiled from Belsky and Kaplan (1970), Bunch and Chang (1980), Krouse and Modzeleski (1970), Robert
and Epstein (1982), Smith and Kaplan (1970), Swart et al. (1983), and Yang and Epstein (1983). Data fields for inorganic
carbon are from Schidlowski et al. (1983).
in a similar hydrothermal silica dike at the ⬃3.5 Ga Chinaman
Creek, located ⬃50 km east of the North Pole area. They
proposed that the kerogen would have been produced by Fischer-Tropsch-Type (FTT) reactions under hydrothermal conditions.
Carbon isotope fractionation accompanied by abiological
synthesis of organic compound is poorly known, though FTT
synthesis under hydrothermal condition has been demonstrated
(e.g., Yanagawa and Kobayashi, 1992; Berndt et al., 1996;
Horita and Berndt, 1999; McCollom et al., 1999; Rushdi and
Simoneit, 2001). Berndt et al. (1996) performed experimental
serpentinization of ultramafic rocks at 300°C, and showed that
the CO2 in the hydrothermal fluid was converted to CH4,
hydrocarbon, and amorphous carbonaceous material by the
reaction with H2. They suggested that magnetite was produced
by the serpentinization of olivine, and catalyzed the FTT synthesis. Unfortunately, carbon isotope fractionation accompanied by aqueous FTT reaction has never been determined
except for CH4 (⌬CO2-CH4 ⫽ ⬃50‰ at 200°C and 20⫺30‰ at
300°C; Horita and Berndt, 1999). Thus, we can not compare
their results with our observations. It is unknown whether the
observed ␦13C values of the kerogens in the silica dikes can be
explained by the products of FTT reactions.
Moreover, it is questionable whether FTT synthesis took
place during the formation of the silica dike, because likely
catalysts for FTT reaction were probably absent. Only magnetite, hematite, and native metals are used as catalysts of commercial Fischer-Tropsch reaction. Thus, these minerals are
Carbon isotopes and petrography of kerogens
candidates for the most effective catalysts of FTT reaction.
However, all of them were probably unstable under the condition of the silica dike formation (Fig. 9). The apparent deficiency of the effective catalysts raises doubts as to whether
FTT reactions would have occurred in these hydrothermal
conditions.
In modern environments, hydrothermal systems around ultramafic rocks have been considered as possible sites of abiological organic synthesis (e.g., Holm and Andersson, 1998).
Holm and Charlou (2001) suggested that the ultramafic hydrothermal system is suitable for FTT synthesis relative to basaltic
setting, because (1) serpentinization of olivine provides magnetite as catalyst and H2 as reactant, and (2) native Fe-Ni
minerals, which is crucial catalysts for the reduction of CO2,
are more common in ultramafic rocks than basalts. In fact, high
emissions of apparently abiological methane and hydrocarbon
have been observed in ultramafic hydrothermal systems of
Mid-Atlantic Ridge (Charlou et al., 1998; Holm and Charlou,
2001), Zambales Ophiolite, Philippines (Abrajano et al., 1988;
Abrajano et al., 1990), and Oman Ophiolite (Neal and Stanger,
1983). These observations may suggest that the presence of
magnetite and/or native metals is crucial for the aqueous FTT
synthesis. If so, it is unlikely that FTT reaction produced the
organic matter during the silica dike formation, because the
wall rock of the silica dike is not ultramafic rocks but basalt
(Figs. 1– 4).
In summary, although we can not eliminate the possibility
that the 13C-depleted kerogen in the silica dike was produced
abiologically, it is premature to consider that FTT reactions
produced the kerogen in the Archean silica dikes. Further
experimental studies are necessary for clarifying the role of
other potential catalysts such as silica and sulfide, and for
determining carbon isotopic fractionations of kerogenous deposits through aqueous FTT synthesis.
7. CONCLUSIONS
Petrographic, isotopic (C and N), and elemental (C, H, and
N) analyses of kerogen in ⬃3.5 Ga hydrothermal silica dikes
provided the following new results and insights:
(1) All the silica dikes contain 13C-depleted kerogen
(⬍⫺30‰) and are classified into sulfide-bearing B-type
and sulfide-free G-type.
(2) The mineral assemblages, dominated by silica (⬎97%)
with minor pyrite, of the B-type silica dikes indicate that
they were probably deposited from reducing hydrothermal
fluid at a temperatures lower than 200°C. The estimated
condition would have been habitable for anaerobic and
thermophilic/hyperthermophilic chemoautotrophs.
(3) The G-type silica dikes often have Fe-oxide pseudomorphs
after carbonate and sulfide, suggesting that they were
formed by metasomatism (oxidation) from the B-types.
(4) The predominance of the G-type dike (⬃90%) indicates
that the metasomatism was pervasive. Nevertheless, the
sporadically occurring B-types preserve premetasomatic
mineralogical signature.
(5) Based on the comparison between B- and G-types, ⬎90%
of organic carbon was consumed during the metasomatism,
and residual carbon was isotopically enriched in 13C by
587
⬎4‰, probably due to kinetic isotope effect (estimated
fractionation factor: 0.9985).
(6) Assuming that the postdepositional shifts of the ␦13C values of the organic carbon were mainly due to the metasomatism, the parent material of the kerogen would have
been isotopically heterogeneous (⬃5‰) and at least some
of the materials would have possessed the ␦13C of
ⱕ⫺38‰.
(7) The inferred significant 13C-depletion and isotopic heterogeneity could have been produced by anaerobic chemoautotrophs such as methanogenic bacteria, but not by aerobic
autotrophs. This is consistent with the estimated condition
of the hydrothermal fluid, suggesting the possibility that the
organic matter was produced by chemoautotrophs in the
fissures developed in seafloor basalts, which acted as conduits of hydrothermal fluid.
(8) Alternatively, the 13C-depleted kerogen may have been
possibly produced by prebiotic reaction such as FTT synthesis. However, the effective catalysts of FTT reaction
(i.e., magnetite, hematite, and Fe-Ni alloy) were probably
unstable during the silica dike formation.
Based on these lines of evidence, we suggest that the ⬃3.5
Ga subseafloor hydrothermal system was probably inhabited by
anaerobic chemoautotrophs. The kerogen-bearing silica dikes
of this age may provide great insights into the early biosphere
on Earth.
Acknowledgments—We thank M. Terabayashi, Y. Kato, K. Okamoto,
T. Ota, T. Kabashima, K. Kitajima, and K. Shimizu for assistance in the
field work. The field collaboration with A. Thorne, K. J. McNamara,
and A. H. Hickman was helpful and much appreciated. We thank H.
Naraoka in Tokyo Metropolitan University for providing experimental
facilities and for discussion. We also thank G. Cody and S. Nakashima
for their helpful comments. Constructive reviews by J. Hayes and D.
Des Marais greatly improved the manuscript. This study was financially supported in part by the Ministry of Education, Science and
Culture, Japan (No. 11691119). Y.U. is grateful for the Research
Fellowships of the Japan Society for the Promotion of Science for
Young Scientists.
Associate editor: G. Cody
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