International Geology Review, Vol. 47, 2005, p. 952–970.
Copyright © 2005 by V. H. Winston & Son, Inc. All rights reserved.

Rare-Earth Element, Lead, Carbon, and Nitrogen Geochemistryof Apatite-Bearing Metasediments from the ~3.8 Ga IsuaSupracrustal Belt, West Greenland
MANABU NISHIZAWA,1 NAOTO TAKAHATA,
Center for Advanced Marine Research, Ocean Research Institute, University of Tokyo, Minamidai 1-15-1,
Nakanoku, Tokyo 164-8639, JapanKENTARO TERADA,
Department of Earth and Planetary Sciences, Hiroshima University, Kagami-yama 1-3-1, Higashi-Hiroshima, 739-8526, JapanTSUYOSHI KOMIYA,
Department of Earth and Planetary Sciences, Tokyo Institute of Technology, Ookayama 2-12-1,
Meguroku, Tokyo, 152-8551, Japan

YUICHIRO UENO,
Research Center for the Evolving Earth and Planets, Department of Environmental Science and Technology,
Tokyo Institute of Technology, Post No. S2-17, Midori-ku, Yokohama 226-8503, JapanAND YUJI SANO
Center for Advanced Marine Research, Ocean Research Institute, University of Tokyo, Minamidai 1-15-1,
Nakanoku, Tokyo 164-8639, Japan

Abstract
We performed rare-earth element (REE) geochemistry and U-Pb geochronology on apatites inmetasediments from the ~3.8 Ga Isua supracrustal belt (ISB) and Akilia Island, West Greenland,
together with stepwise combustion isotopic investigation of carbon and nitrogen for the apatite-bearing quartz-magnetite BIF of uncontested sedimentary origin from northeastern ISB.
Ion microprobe analyses reveal that apatites in psammitic schist from the ISB show a U-Pb
isochron of 1.5 ± 0.3 Ga. This age is similar to those of Akilia apatite and the Rb-Sr age of 1.6 Ga
for the pegmatitic gneiss in the Isukasia area in literature, suggesting a late (~1.5 Ga) metamorphic
event (≥400°C). Pb isotopic ratios of apatite in the quartz-magnetite BIF are also affected by the latemetamorphic event around 1.5 Ga. Chondrite-normalized REE patterns of apatites in the BIF showflat patterns with a significant positive Eu anomaly, suggesting hydrothermal influence; this isconsistent with a primary depositional origin. In contrast with the quartz-magnetite BIF, apatites in
the psammitic schist from the ISB and those in the Akilia BIF show different REE patterns, which
resemble those of apatites from secondary mafic and felsic rocks, respectively.
Carbon isotopic ratios for the quartz-magnetite BIF by stepwise combustion suggest that twocomponents of reduced carbon are present. One is released below 1000°C (mainly 200–400°C; lowtemperature carbon = LTC), and the other above 1000°C (high-temperature carbon = HTC). δ13C
values of the LTC are about –24‰. The LTC is clearly contaminant incorporated after metamorphism, because such a low-temperature component could not have survived the ≥400°C metamorphic event. On the other hand, δ13C values of the HTC are –30‰ for one aliquot and –19‰ foranother. The HTC is probably sequestered within magnetite in the BIF, because the decrepitation
temperature of magnetite is ~1200°C. The HTC could not exist within quartz and apatite (decrepitation temperatures: 400–600°C and 600–800°C, respectively), or along grain boundaries. Because
the magnetite is concordant with bedding surfaces, it is plausible that the HTC was incorporated in
the magnetite during diagenesis. Thus, HTC is the most important candidate for primary carbon
1Corresponding author; email: nishizawa@ori.u-tokyo.ac.jp

0020-6814/05/825/952-19 $25.00

952

ISUA SUPRACRUSTAL BELT, WEST GREENLAND

953

preserved in the BIF. δ13C values of HTC cannot be explained as those of Isua carbonate. On theother hand,that the very low δ13C values (–30‰), negative δ15N values (–3‰), and low C/N elemental ratios (86) for the >1000°C fraction of one aliquot are comparable to those of kerogen in Archeanmetasediments. Therefore, despite the presence of secondary carbon (i.e., LTC), the BIF is suggested to possibly contain highly 13C-depleted kerogenous material, which is unlikely to have been
incorporated after metamorphism. Although carbon isotopic change of the kerogenous material due
to metamorphic effects cannot be precisely determined from the present data, this study shows thatfurther analysis of magnetite from the Isua BIF is a key to the search for the early life.

Introduction
THE OLDEST METASEDIMENTS exposed in the ~3.8 Ga
Isua supracrustal belt (ISB) and >3.8 Ga Akilia
Island, West Greenland, have been intensively studied to search for the oldest trace of life on Earth(e.g., Schidlowski et al., 1979; Mojzsis et al., 1996;
Rosing, 1999). Although the severe metamorphismand deformation would not allow preservation of
microfossil structures, 13C-depleted graphite occurs
in these oldest metasediments (–50 to –6‰; Oehler
and Smith, 1977; Perry and Ahmad, 1977; Schidlowski et al., 1979; Hayes et al., 1983; Shimoyama
and Matsubaya, 1992; Naraoka et al., 1996; Mojzsis
et al., 1996; Rosing, 1999; Ueno et al., 2002), someof which have comparable δ13C values to those of
biologically produced sedimentary organic carbonof the present Earth. Among these graphites, themost 13C-depleted ones are enclosed within apatite(–20 to –50‰; Mojzsis et al., 1996); thus the apatite-graphite association is a candidate for biomarker. Recently, this idea was challenged by theobservations of Fedo and Whitehouse (2002a) and
Lepland et al. (2002), suggesting that graphite-bearing apatite occurs only in metasomatized mafic/
ultramafic volcanic rocks, whereas apatite in quartzmagnetite BIF of uncontested sedimentary origin isgraphite-free. This indicates that some apatites aswell as the graphite inclusions would have been
incorporated during the metasomatism or metamorphism, thus not of primary depositional origin.
Therefore, understanding the geochemistry andgeochronology of the apatite is necessary to evaluate
the origin and age of the apatite.
Here we report results of in situ analyses of rareearth elements (REEs) of apatites from two metasediments of different lithologies (i.e., quartz-magnetite
BIF, and psammitic schist) in the ISB to evaluate theorigin of these apatites. The REE pattern of apatite
may be used to infer the crystallization environments (Gromlet and Silver, 1983), the chemistry of
fluids related to sedimentary environments (Elisabeth and Alain, 1986; Toyoda and Tokonami, 1990),

and the effects of metamorphism (Bingen et al.,1996). In addition, results of ion microprobe U-Pb
geochronological study of these apatites are also
reported, in order to search for primary apatite
deposited at ~3.8 Ga. Results of this study suggestthat apatites in the quartz-magnetite BIF are ofprimary depositional origin.
Another question is where 13C-depleted graphite
exists in the BIF. Unfortunately, the apatite of ourstudy does not contain visible graphite (>0.1 µm)
under the optical microscope. Despite thoroughobservations of more than 300 thin sections (Ueno etal., 2002; Lepland et al., 2002), graphite has not yet
been reported from quartz-magnetite BIF in the ISB.
Thus, petrographic study is not always useful tolocate the site of 13C-depleted graphite, as demonstrated for Isua turbidite by Rosing (1999). However, this does not prove the absence of graphite in
the BIF, because smaller graphite particles and/or
inclusions within opaque minerals such as magnetite possibly exist in the BIF. In order to identify
the site of the graphite, extraction of carbon was
conducted from the sample by stepwise combustion,
because graphites enclosed within different minerals are expected to be degassed at distinctive
temperature steps. The method has also potential toseparate original carbon from secondary contaminant (e.g. Pinti et al., 2001). Recently, Van Zuilen etal. (2002) reported that the reduced carbon in Isua
metasediment is almost all released below 500°C,
suggesting a post-metamorphic origin. They concluded that there is no evidence for the existence oforiginal reduced carbon in the Isua metasediments.
However, carbon isotopic measurements of thehigher-temperature fraction (>1000°C) has not been
performed so far, and has potential to obtain theoriginal signature. In addition to carbon, measurement of nitrogen is important, because nitrogen isalso a bio-essential element.
We also report the results of stepwise combustion
measurements for carbon and nitrogen over a wide
range of temperature steps (200°–1200°C) for thequartz-magnetite BIF. Based on the release patterns

954

NISHIZAWA ET AL.

and their isotopic ratios, we evaluate the site andphase of carbon and nitrogen released at various
temperature steps, and discuss timing of their incorporation in the rock and their origin.

Sample and Analytical Method
Sample description
The ~3.8 Ga Isua supracrustal belt (ISB) of West
Greenland is the oldest known succession of volcanic and sedimentary rocks on Earth (Nutman, 1986;
Appel et al., 1998; Komiya et al., 1999; Myers,
2001; Nutman et al., 2002). Supracrustal rocks are
exposed in about a 35 km long arcuate belt (Fig. 1).
Most of the rocks underwent metamorphism up toamphibolite facies (e.g., Boak and Dymek, 1982;
Nutman, 1986; Hayashi et al., 2000; Appel et al.,
2001; Komiya et al., 2002), deformation (e.g., Appelet al., 1998; Myers, 2001; Nutman et al., 2002), andmetasomatism (Rose et al., 1996; Rosing et al.,1996), which makes recognition of the protoliths
difficult. However, in the northeastern part of the
ISB (nearly the same area as the low-strain domain
of Appel et al., 1998), primary volcanic and sedimentary structures including pillow lava, pillow
breccia, graded bedding, and polymictic conglomerate (Nutman, 1986; Appel et al., 1998; Komiya etal., 1999; Fedo, 2000; Myers, 2001) are exceptionally well preserved. Recently, Appel et al. (2001)
reported quartz globules with primary fluid inclusions derived from the ~3.8 Ga hydrothermal metamorphism in the low-strain domain, because thedeformation and metasomatism of the domain are
relatively weak. Thus, for analyses, we chose two
apatite-bearing metasediments from the northeastern ISB.
One sample is quartz-magnetite BIF (43-44A) of
uncontested sedimentary origin from the northeastern end of the ISB (Fig. 1). About 20 m thick BIF
overlies basaltic pillow lava, and is overlain by
bedded chert. Detailed field occurrence of the BIF
is also given in Figure 9B of Komiya et al. (1999).
The BIF (43-44A) mainly consists of alternating
quartz and magnetite layers, with minor amounts ofapatite and actinolite (Fig. 2). The apatites areconcentrated in a thin (<1 mm) layer, which is concordant with the original bedding surface. Note that
under the microscope (>0.1 µm) graphite is not
visible in the specimen.
Apatite-bearing psammitic schist (K485) fromthe northeastern ISB was also selected for analysis(Fig. 1). The sample mainly consists of quartz, withminor amounts of actinolite, anthophyllite, andcalcite. Trace amounts of pyrrhotite, chalcopyrite,
magnetite, and graphite are also present. Graphite
particles occur on grain boundaries of quartz, andoccur within mineral such as quartz and garnet.
Detailed descriptions are given in Ueno et al.
(2002).
For comparison, apatite in the granulite-facies
BIF from Akilia Island (Mojzsis et al., 1996;
provided by A. P. Nutman) was also studied. 13Cdepleted graphite in the apatite of the sample has
been claimed as the oldest trace of life (Mojzsis etal., 1996). However, recent field observations andtrace element geochemistry by Fedo and Whitehouse (2002a) suggested the possibility that “theAkilia BIF” of Mojzsis et al. (1996) is metasomatized volcanic rock, and the protolith of the rock is
still debated (e.g., Fedo and Whitehouse, 2002b;
Friend et al., 2002; Mojzsis and Harrison, 2002;
Mojzsis et al., 2003; Bolhar et al., 2004). The U-Pb
age of 1.5 Ga for the Akilia apatite was reported bySano et al. (1999b).

Analytical techniques
For ion microprobe analysis, sample chips were
cast into epoxy-resin disks with several grains ofstandard apatite, and polished until mid-sections
were exposed. The samples were evacuated in thesample lock overnight, and introduced into the sample stage of the ion source chamber of the Sensitive
High Resolution Ion Micro Probe (SHRIMP)
installed at Hiroshima University. A primary beamof about 2.5 nA O2 – was focused to sputter an areaof 20 µm diameter on individual apatite grains andthe positive secondary ions were accelerated at 10
kV. There were no isobaric interferences in the mass
range over 204Pb and 208Pb at a mass resolution of
5800. Mercury interference on 204Pb was negligible,
which was verified by 200Hg and 202Hg measurements (Sano and Terada, 2002). Concentrations ofU, Th, and 206 Pb were obtained by the relative
intensity of each element to the matrix beam of
40Ca 31P16O +. Experimental details were given in
2
3
Sano et al. (1999a). Analytical data are presented inTables 1 and 2, and Figures 3 and 4.
Identical primary beam conditions were adopted
in rare-earth element (REE) measurements, while
the mass resolution of ion microprobe was enhanced
to 9300 at 1% peak height in order to separate
oxides of light REEs from heavy REEs in cases
where we did not use the energy filtering technique.
The magnet was cyclically peak-stepped from mass

ISUA SUPRACRUSTAL BELT, WEST GREENLAND

955

FIG. 1. Geological map of the Isua supracrustal belt, West Greenland (modified after Komiya et al., 1999) showing
sample locations.

139 (139La+) to 175 (175Lu+), including the background and all significant REE isotopes (140Ce+,
141 Pr + , 145 Nd + , 146 Nd + , 147 Sm + , 149 Sm + , 151 Eu + ,153 Eu + , 155 Gd + , 157 Gd + , 159 Tb + , 161 Dy + , 163 Dy + ,165Ho+, 166Er+, 167Er+, 169Tm+, 171Yb+, and 172Yb+)
and the matrix peak ( 40 Ca2 31 P 16 O3 + ). Observed
ratios of REE + / 40 Ca 2 31 P 16 O 3 + were calibrated
against those of standard apatite, whose REE

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NISHIZAWA ET AL.

The carbonaceous gas was converted to CO2, thenpurified and trapped in Pyrex glass tube by traps
held at iced pentane and liquid nitrogen temperatures, respectively. Each CO2-filled tube was cut
and sealed by standard glass-blowing techniques.
Nitrogen isotopic ratios were measured using a modified noble gas mass spectrometer (VG3600, VG
Isotopes Co.) at the Ocean Research Institute, University of Tokyo, by static operation mode. Prior toisotopic analysis, the amount of nitrogen was
adjusted to that of air standard gas (0.3–0.6 nmol) inorder to compensate for any pressure effect on theisotopic ratio. Experimental details such as abundance measurements of carbon, nitrogen, and argon
gases, and blank correction were given by Takahata
et al. (1998). Carbon isotopic ratios of the purified
CO2 were measured by a conventional stable isotope
mass spectrometer (Finnigan MAT delta S) at
Shimane University. Isotopic ratios are reported relative to PDB for carbon and relative to atmosphericN2 for nitrogen. Analytical data are presented inTable 4 and Figure 6.

Results and Discussion
FIG. 2. A. Cut-surface of apatite-bearing quartz-magnetite
BIF (43-44A) from the northeastern ISB. Vertical to original
bedding surface. Scale bar represents 1 cm. B. Photograph of
apatite-rich layer separated from the BIF and mounted on
epoxy resin for microprobe analyses. Open circles indicate
spots analyzed by SHRIMP. Scale bar represents 300 µm.

abundances were determined by ICP-MS after
chemical dissolution (Sano et al., 2002). Analytical
data are presented in Table 3 and Figure 5.
For stepwise combustion experiments of BIF 4344A, the samples were cut into tips (~125 mm3) bya water-cooled saw. The rock tips contain both
quartz and magnetite layers. The tips were ultrasonically washed with distilled water and subsequently
acetone, then loaded into a vacuum vessel andbaked under vacuum at 200°C for ~8 hours before
the measurements. The combustion was carried out
in steps of 200°C in order to degas carbon, nitrogen,and argon phases included in quartz, apatite, andmagnetite separately. At each step, using two hotcopper oxides, two hot platinum foils, and four cryogenic traps, nitrogen and argon were separated,
purified from carbonaceous (hydrocarbon, CO2) gas.

Rare-earth element abundance in apatite
Table 3 lists concentrations of rare-earth elements for the apatite in the quartz-magnetite BIF(43-44A), the psammitic schist (K485) from ISB,
and the Akilia BIF, together with Eu anomalies
expressed as Eu/Eu*. Figure 5 shows chondritenormalized REE patterns for these apatites. Akilia
apatite shows a pattern of LREE enrichment and asmall positive Eu anomaly. On the other hand,
apatite in the psammitic schist (K485) shows a
LREE-depleted REE pattern. Apatite in the quartzmagnetite BIF shows a flat pattern with a significant
positive Eu anomaly.
The REE pattern of the Akilia apatite is similar
to those of apatite in granodiorite from the eastern
Peninsular Ranges batholith (Gromlet and Silver,
1983), suggesting that the Akilia apatite wasderived from felsic igneous magma (i.e., detrital
origin). The REE pattern of the apatite in the
psammitic schist (K485) is similar to that of apatite
in a mafic dike from the ISB (AL1-2 of Lepland etal., 2002), suggesting that this apatite was derived
from mafic igneous magma (i.e. detrital origin).
The REE pattern of apatite in the quartz-magnetite BIF (43-44A) is different from that of secondary
apatite in carbonate rocks and mafic dikes from the
ISB (i.e., a MREE-enriched convex pattern without

957

ISUA SUPRACRUSTAL BELT, WEST GREENLANDTABLE 1. U and Th Contents, 238U/204Pb Ratios, and Pb Isotopic Compositionsof Apatite in Psammitic Schist (K485) from the ISB1
238U/204Pb

206Pb/204Pb

13.6

194.7 ± 61.2

69.7 ± 19.4

0.367 ± 0.022

4.5

15.5 ± 3.9

24.6 ± 3.3

0.758 ± 0.060

U,

Th,

ppm

ppm

K485.a32.9

K485.b

2.0

207Pb/206Pb

K485.c1.9

2.3

47.8 ± 17.0

26.9 ± 5.8

0.621 ± 0.047

K485.d

2.2

15.1

29.6 ± 7.8

27.2 ± 6.1

0.708 ± 0.040

K485.e3.5

17.8

76.5 ± 27.7

36.9 ± 10.2

0.728 ± 0.070

K485.f

1.9

10.7

28.6 ± 7.8

33.5 ± 8.2

0.779 ± 0.045

K485.g35.5

30.5

410.3 ± 112.7

123.7 ± 30.90.240 ± 0.011

K485.h55.0

56.4

666.0 ± 152.3

212.8 ± 35.9

0.155 ± 0.011

K485.i

3.8

11.2

131.6 ± 44.6

55.6 ± 21.5

0.487 ± 0.093

K485.j4.9

1.2

149.7 ± 55.5

64.3 ± 23.4

0.341 ± 0.046

K485.k

18.9

12.6

34.2 ± 4.6

23.7 ± 2.7

0.622 ± 0.051

1Uncertainty associated with the ratios is 1 sigma. Those of U and Th contents are ± 30% as estimated by repeat measurements of apatite standard.

a Eu anomaly and a LREE-depleted pattern without
a positive Eu anomaly, respectively; Lepland et al.,2002). On the other hand, this pattern is similar tothat of primary apatite in adjacent BIF from ISB (aflat pattern with a significant positive Eu anomaly;
Lepland et al., 2002). A positive Eu anomaly was
also measured for quartz-magnetite BIF from thenortheastern ISB (Shimizu et al., 1990; Dymek andKlein, 1988). Their REE patterns with positive Eu
anomalies suggest that they were likely depositedfrom hydrothermally influenced contemporaneous
seawater. This is because positive Eu anomalies
currently are observed for hydrothermal fluids suchas at the East Pacific Rise and the Mid-Atlantic
Ridge (Eric et al., 1999) and REE characteristics ofapatite and BIF are thought to contain signatures ofthe solution from which they precipitated (Elisabethand Alain, 1986; Toyoda and Tokonami, 1990;
Derry and Jacobsen, 1990). Recently, Appel et al.
(2001) showed that metabasites in the Isuasupracrustal belt underwent hydrothermal metamorphism at ~3.8 Ga. Therefore, the REE patternsuggests that apatite in the quartz-magnetite BIF
was likely deposited from hydrothermally influenced early Archean seawater. This is consistent
with a primary depositional origin.

U-Pb dating of apatite
Apatite in psammitic schist (K485). Table 1 lists
results of ion microprobe U-Pb dating of apatite
separated from the psammitic schist (K485). Figure 3
shows a positive correlation between 238U/204Pb and
206Pb/204Pb ratios. A least-squares fitting using theYork method (York, 1969) yields a 238U-206Pb isochron
age of 1533 ± 260 Ma (95% confidence level,
MSWD = 0.8, error correlation = 0.789). This age ismuch younger than the sedimentary age of ~3.8 Ga(Nutman et al., 1997). This suggests that the U-Pb
system of apatite in K485 was open at about 1.5 Ga.
This age is comparable with a ~1.5 Ga apatite in the
Akilia BIF of Mojzsis et al. (1996), dated by Sano et al.
(1999b). It is well documented that the Isuasupracrustal rocks and the Amîtsoq gneiss were subjected to a series of metamorphic processes (Pankhurst
et al., 1973). The most recent thermal event, as suggested by Rb-Sr systematics of muscovite-phengite
samples from pegmatitic gneiss in the Isukasia area,
was at 1623 ± 65 Ma (Baadsgaard et al., 1986), which
may have a close relation with the ~1.5 Ga apatite.
Using Dodson’s (1973) equation, an activationenergy of 229 kJ/mol, and a frequency factor of 2 ×
10 –8 m2/sec (Cherniak et al., 1991), the closure

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NISHIZAWA ET AL.

TABLE 2. U, Th, and Pb contents, and Pb Isotopic Compositionsof Apatite in Quartz-Magnetite BIF (43-44A) from the ISB1
206Pb/204Pb

207Pb/204Pb

208Pb/204Pb

13.12 ± 0.22

13.66 ± 0.38

32.97 ± 0.78

43-44A.420.b

12.53 ± 0.24

13.88 ± 0.31

32.19 ± 0.84

43-44A.420.c

13.13 ± 0.33

13.46 ± 0.44

31.83 ± 1.08

43-44A.420.d

13.63 ± 0.27

14.60 ± 0.35

34.74 ± 0.94

43-44A.420.e

12.87 ± 0.38

14.19 ± 0.47

32.05 ± 1.49

Sample43-44A.420.a

43-44A.420.f

U,

Th,

206Pb,

ppm

ppm

ppm

0.020

0.004

0.98

12.06 ± 0.19

13.02 ± 0.28

30.33 ± 0.87

43-44A.420.g

11.58 ± 0.44

12.78 ± 0.82

29.87 ± 1.75

43-44A.420.h

12.48 ± 0.28

12.70 ± 0.40

30.74 ± 1.12

43-44A.420.i

12.91 ± 0.27

13.68 ± 0.33

31.68 ± 0.88

43-44A.420.j

13.04 ± 0.37

14.17 ± 0.47

33.56 ± 1.26

43-44A.420.k

12.35 ± 0.36

13.16 ± 0.46

32.15 ± 1.33

43-44A.420.l

12.94 ± 0.27

14.19 ± 0.42

32.91 ± 1.15

43-44A.420.m

12.01 ± 0.23

13.34 ± 0.30

31.02 ± 0.92

43-44A.420.n

0.051

13.85 ± 0.11

14.56 ± 0.33

34.90 ± 0.68

12.29 ± 0.12

13.29 ± 0.25

30.72 ± 0.54

43-44A.420.p

11.81 ± 0.14

13.04 ± 0.24

30.50 ± 0.66

43-44A.420.q

12.16 ± 0.27

13.72 ± 0.82

32.35 ± 1.78

43-44A.420.r

11.55 ± 0.19

13.21 ± 0.40

31.07 ± 1.80

43-44A.420.s

12.12 ± 0.23

14.02 ± 0.69

32.67 ± 1.31

43-44A.420.t

14.06 ± 0.29

14.03 ± 0.45

32.62 ± 1.23

43-44A.420.u

16.40 ± 0.48

15.33 ± 0.68

36.88 ± 1.79

12.19 ± 0.15

13.18 ± 0.24

30.52 ± 0.70

12.00 ± 0.20

13.73 ± 0.80

33.06 ± 1.60

43-44A.420.x

12.53 ± 0.57

13.80 ± 1.26

32.98 ± 2.20

12.61 ± 0.04

13.56 ± 0.08

31.98 ± 0.20

0.050

0.019

2.88

43-44A.420.w

Average

0.057

0.033

1.24

43-44A.420.o

43-44A.420.v

0.072

0.020

0.019

1.80

1.72

1Uncertainty associated with the ratios is 1 sigma. Those of contents are ± 30%. Because contents of U and Th were very
low, measurements were carried out for 4 spots.

temperature is estimated to be 384–419°C for theapatite U-Pb system with a 30 µm grain size and at aslow cooling rate of 1–10°C/Ma. The estimate suggests that the K485 underwent metamorphism above
400°C at 1.5 Ga.
Apatite in quartz-magnetite BIF (43-44A). Table
2 lists U, 206Pb concentrations, and Pb isotopic
ratios of the apatite in the quartz-magnetite BIF (43-44A). Inasmuch as U contents are substantiallylower than Pb contents, in situ radiogenic production of 206Pb and 207Pb during geological time issuggested to have been negligibly small, probably
within the analytical uncertainty of the observed Pb
isotopic ratio at the 2σ level. We compared thePb isotopic ratios of the apatite in the BIF with two
Pb growth models (Fig. 4). One is the model for

ISUA SUPRACRUSTAL BELT, WEST GREENLAND

959

FIG. 3. Correlation diagram of 238U/204Pb–206Pb/204Pb ratios of apatite in psammitic schist (K485) from ISB. Uncertainty associated with the ratios is 1 sigma. Solid line is the best fit using the York method (York, 1969). In the fitting
calculation, we take error correlation of Rec = 0.872, which is calculated by the correlation coefficient between δ(238U/
204Pb)/(238U/204Pb) and δ(206Pb/204Pb)/(206Pb/204Pb).

FIG. 4. Pb isotopic ratios of apatite (open circles) in quartz-magnetite BIF (43-44A) from northeastern ISB. Uncertainty assigned to the ratios is 1 sigma. A black square indicates the initial Pb isotopic ratio of the ~1.5 Ga apatite inK485, for comparison. Two model growth curves for terrestrial Pb (SK = Stacey and Kramers, 1975) and Early Archean
Pb (FR = Frei and Rosing, 2001) are also shown. Using the York (1969) method, best fit of the Pb isotopic ratios of theapatite (43-44A) intersects the growth curves at model Pb ages of 3.9 ± 0.2 and 1.5 (+0.6, –1.1) Ga for the SK model,
and at 4.0 ± 0.2 and 1.2 (+ 0.6, –1.2) Ga for the FR model, respectively.

terrestrial Pb (Stacey and Kramers, 1975; named SK
model), and the other for Early Archean Pb definedby galena from the western part of ISB (Frei and

Rosing, 2001; named FR model). Figure 4 also
shows the initial Pb isotopic ratio of the apatite
in adjacent psammitic schist (K485), which is

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NISHIZAWA ET AL.

TABLE 3. Rare-Earth Element Abundances (ppm) of Apatites from ISB and Akilia Island, West Greenland1
43-44A.420.y43-44A.421.a

K485.l

K485.mAkilia.a

Akilia.b

La

5.0

3.7

5.1

6.5

495.0

456.0

Ce

31.0

19.7

27.6

26.6

1030.0

870.0

Pr

5.0

3.9

5.5

4.7

89.5

80.0

Nd

21.2

19.9

23.3

22.2

330.0

285.0

Sm

7.2

7.1

8.0

7.3

58.0

50.5

Eu

6.5

6.3

5.6

4.9

25.7

24.0

Gd

9.2

10.1

27.6

24.1

79.6

64.1

Tb

1.6

1.7

8.6

7.2

12.2

10.4

Dy

8.9

10.0

48.8

43.1

63.6

55.7

Ho

2.1

2.4

11.2

9.1

13.6

11.5

Er

7.1

7.1

28.4

26.8

37.1

29.9

Tm

1.2

1.2

4.1

3.8

4.8

3.6

Yb

8.5

8.7

18.4

16.6

24.2

17.9

Lu

1.5

1.6

3.1

2.9

3.7

3.5

sum REE

120.0

100.0

225.0

206.0

2270.01960.0(Eu/Eu*)CN

2.4

2.3

1.2

1.1

1.1

1.3

1Uncertainty in REE abundances is ± 20% based on repeat measurements of apatite standard. (Eu/Eu*)
CN = (Eu)CN/{(Sm)CN*(GdCN)}0.5; CN means chondrite normalization.

calculated by a 3D linear regression for the total U/
Pb isochron, constrained to intersect the TeraWasserburg concordia (Ludwig, 1998).
As shown in Figure 4, the apatite in the BIF hasvarious Pb isotopic ratios. Within analytical uncertainty, one of the data points corresponds to the initial Pb isotopic ratio for apatite in K485 with a ~1.5
Ga U-Pb age, suggesting incorporation of Pb derived
from the ~1.5 Ga metamorphism. However, most of
the other data points are scattered toward ~3.8 Ga.
This may indicate a two-component mixing trend.
Least-squares fitting for the Pb isotopic ratios of theapatite by York (1969) method gives a slope of 0.52
± 0.15 and a Y-intercept of 7.0 ± 1.9 (95% confidence level, MSWD = 1.3, error correlation =0.742). The fitting line has two intersects with the
Pb growth curve at model Pb ages of 3.9 ± 0.2 and1.5 (+0.6, –1.1) Ga for SK model, and at 4.0 ± 0.2
and 1.2 (+0.6, –1.2) Ga for FR model. The older
model Pb age of the intersection is consistent withdepositional age of the BIF (Nutman et al., 1997),
even though either model is adopted.
Because the quartz-magnetite BIF contains
extremely unradiogenic Pb (Moorbath et al., 1973;
Frei et al., 1999), it cannot be ruled out that apatite
incorporated ancient Pb from the BIF during metamorphic event(s) above 400°C. In this case, theolder model age is not a direct estimate of age of theapatite, but corresponds to the age of the BIF (4344A). On the other hand, the REE pattern of theapatite suggests that the apatite had a primary andhydrothermal depositional origin, as discussed in
the previous section. This is consistent with the factthat closure temperature of REE in apatite is muchhigher than that of Pb; a closure temperature for Nd
is 612–655°C, whereas that for Pb is 384–419°C for
apatite with a 30 µm grain size at a cooling rate of
1–10°C/Ma; (Cherniak et al., 1991; Cherniak 2000).
Therefore, the apatite was also formed at ~3.8 Ga
even in the latter case, because the formation of

ISUA SUPRACRUSTAL BELT, WEST GREENLAND

961

FIG. 5. Chondrite-normalized rare-earth element abundance patterns of apatites in the quartz-magnetite BIF (4344A) and psammitic schist (K485) from the ISB. Data for graphite-bearing apatites in granulite-facies BIF from the
Akilia Island (Mojzsis et al., 1996) are also shown. Uncertainty assigned to the abundance is 1 sigma.

apatite is simultaneous with the deposition of the
BIF. Thus, it is important to locate the carbon andnitrogen within the primary apatite in the BIF. In thefollowing sections, we discuss the site (e.g., within
apatite, magnetite), age, and the origin of carbon
and nitrogen in the BIF, based on results of stepwise
combustion of the BIF.

Stepwise combustion experiments of carbon
and nitrogen for 43-44A BIF
For the quartz-magnetite BIF (43-44A) from thenortheastern ISB, carbon and nitrogen isotopic
ratios were measured by stepwise combustion. Contents, isotopic and elemental ratios of carbon andnitrogen released from each temperature step areshown in Table 4. Figure 6 shows release patterns of
two aliquots from the same BIF. Carbon and nitrogen
abundances show two major peaks, with distinctive
isotopic ratios. For carbon, two major peaks are
identified at temperature steps of 200–400°C and
1000–1200°C, whereas for nitrogen, two peaks
appear at steps of 400–600°C and 1000–1200°C. In

the following section, we discuss two groups ofreleased gases, separately. One is released below
1000°C, and the other above 1000°C.
Carbon and nitrogen released at temperature
below 1000°C. Below 1000°C, there are abundance
peaks at a step of 200–400°C for carbon, and 400–
600°C for nitrogen. The δ13C values of the peakfractions are –23.55 ± 0.02‰ for one aliquot (4c),and –25.08 ± 0.03‰ for another (5c), whereas theδ15N value is +6.9 ± 0.8 and +1.7 ± 0.6‰ for 4c and5c, respectively.
The carbon released below 400°C probably is asecondary contaminant incorporated after metamorphism, because the BIF probably experienced~400°C metamorphism (Hayashi et al., 2000). Sucha low-temperature fraction is unlikely to have survived the metamorphic events. The carbon released
below 400°C probably exists along grain boundaries, because decrepitation temperatures of quartz(400–600°C; Sano and Pillinger, 1990), apatite(600–800°C; Nadeau et al., 1999), and magnetite(~1200°C; Tolstikhin et al., 2002), which are the962

NISHIZAWA ET AL.

TABLE 4. Summary of Stepwise Combustion Experiments, Showing N and C Contents, C/N Atomic Ratios,and N and C Isotopic Compositions of the Quartz-Magnetite BIF (43-44A) from the ISB1
Temp.

N

C

C/N

δ15N

δ13C

°C

ppm

ppm

atomic ratio

‰

‰

200

–

3.9

–

–

–

400

0.056

33.0

688

2.6 ± 0.6

–23.55 ± 0.02

600

0.210

12.0

67

6.9 ± 0.8

–24.03 ± 0.03

800

0.075

3.5

54

11.4 ± 0.6

–

1000

0.030

6.8

264

7.8 ± 0.8

–20.73 ± 0.04

1200

0.019

11.0

675

0.8 ± 7.4

–18.82 ± 0.02

Total

0.390

70.2

210

6.9

–21.3

200

–

2.1

–

–

–

400

0.122

20.0

191

–0.2 ± 0.6–25.08 ± 0.03

600

0.419

14.0

39

1.7 ± 0.6

–25.31 ± 0.02

800

0.110

2.0

21

7.1 ± 0.6

–

1000

0.054

2.5

54

5.3 ± 0.6

–

1200

0.258

19.0

86

–3.1 ± 0.7–30.25 ± 0.02

Total

0.963

59.6

72

1.0

–27.0

4c, 300.6 mg

5c, 288.2 mg

1Uncertainty assigned to δ15N values and δ13C values are 1 sigma; “ – ” = blank level or below detection limit.

main components of the BIF (43-44A), are higher
than 400°C; thus the low-temperature fraction is notexpected to be released from these minerals. δ13C
values of the carbon released below 400°C (–23 to–25‰) are similar to those of sedimentary organiccarbon on the present Earth, but are different from
Isua carbonates (–7 to +5‰; Oehler and Smith,
1977; Perry and Ahmad, 1977; Schidlowski et al.,1979). Thus, secondary reduced carbon probably
exists along grain boundaries of the BIF. This result
indicates that we should be careful to identify andseparate potential primary carbon from the secondary carbon.
In addition, a smaller amount of carbon isreleased at 400–1000°C. This carbon is possibly
derived from inclusions within quartz and apatite,which release their inclusions at these temperature
steps. However, we cannot evaluate their contributions to the amount of released carbon because
reduced carbon along grain boundaries could also

be degassed over a wide range of temperatures
below 1000°C, by analogy with the combustion
temperature for kerogen (e.g., Wedeking et al.,
1983; Beaumont and Robert, 1999).
It is important that no release peak was observed
at 600–800°C, which corresponds to the decrepitation temperature of apatite. Thus, there is no evidence to show a close relationship between apatite
and 13C-depleted reduced carbon. The presence of
apatite does not necessarily mean the existence of a“biomarker”.
On the other hand, the phases of nitrogen
released at 400–600°C are problematic. One possible phase is N2-fluid enclosed within quartz. Pinti etal. (2001) also identified nitrogen with δ15N valuesof about +7‰ released at ~600°C, associated with amajor peak of primordial 36Ar abundance, by stepwise combustion of Isua metacherts. They concluded that the nitrogen was derived from fluidinclusions enclosed in quartz. δ15N values of ourISUA SUPRACRUSTAL BELT, WEST GREENLAND

963

FIG. 6. Results of stepwise combustion experiments for two aliquots (4c and 5c) of the quartz-magnetite BIF (4344A). Stepwise release patterns of abundances and isotopic ratios of carbon and nitrogen are shown. One sigma error bars
are shown for nitrogen isotope ratios. For carbon isotope ratios, one sigma error bars are within the symbol size.

sample 4c (6.9 ± 0.8‰) are closely similar to thefluid nitrogen of Pinti et al. (2001). The other possible phase for the nitrogen is kerogenous material on
grain boundaries and/or within quartz.
Carbon and nitrogen released above 1000°C.
Above 1000°C, there are release peaks both forcarbon and nitrogen (Fig. 6). δ13C values and δ15N
values are –18.82 ± 0.02‰ and 0.8 ± 7.4‰ for one
aliquot (4c), and –30.25 ± 0.02‰, –3.1 ± 0.7‰ foranother (5c), respectively. These high-temperature
fractions are very important as a candidate forprimary organic matter, because their high releasetemperature makes their post-metamorphic origin
unlikely, as discussed in the following sections. We
named the carbon and nitrogen released above
1000°C HTC and HTN, respectively. In the following sections, we discuss site, phase, and origin of the
HTC and HTN.
Site of HTC and HTN, and the timing of theirincorporation into the BIF: The BIF (43-44A)
consists mainly of quartz and magnetite with minor
apatite and actinolite. Among these minerals, only
magnetite has a decrepitation temperature over
1000°C (~1200°C; Tolstikhin et al., 2002), suggesting that HTC exists within the magnetite. In addi-tion, it is unlikely that HTC is on grain boundaries
and/or on the sample surface, because the combustion temperature of kerogen is below 1000°C (e.g.
Wedeking et al., 1983; Beaumont and Robert,1999). Thus, HTC probably exists within the magnetite. This idea is also supported by an earlier
study which showed that peaks of release abundance of carbon and nitrogen occur above 1000°C
by stepwise combustion of magnetite separated from
Isua BIF (Pinti et al., 2001). Therefore, the precursors of HTC and HTN were probably present at thetime of magnetite formation.
Magnetite of the analyzed specimen occurs along
bedding surfaces, and not as secondary veins (Fig.
2). Thus, it is plausible that the HTC and HTN
would have been entrapped by magnetite during
diagenesis of the quartz-magnetite BIF. In addition,
the metamorphic age of bedded magnetite in BIF
from the same northeastern end of the ISB is 3.69
Ga, estimated by Pb-Pb systematics (Frei et al.,1999). This indicates that the carbon and nitrogenwere incorporated in the BIF before 3.69 Ga.
Phase of HTC and HTN: Possible candidates forthe phase of HTC are graphite (including kerogenous material), fluid, and carbonate. The δ13C

964

NISHIZAWA ET AL.

values of the HTC are much lower than those of Isua
carbonate (–7 to +5‰; Oehler and Smith, 1977;
Perry and Ahmad, 1977; Schidlowski et al., 1979),
and those of CO2 gas (likely ≥10‰). Thus, it isunlikely that the HTC was derived from carbonate or
CO2 fluid. Significant 13C depletions of HTC (up to–30‰) suggest that the HTC is derived fromgraphitic or kerogenous material.
The kerogenous origin of HTC is consistent with
the negative δ15N value of HTN (–3‰) and their
C/N ratio (86) for one aliquot (5c), because thesevalues are within the range of those of Archean
kerogen in metasediments (Beaumont and Robert,1999). This suggests that HTN would be derived
from kerogenous material. In summary, the phaseof HTC and HTN is suggested to be kerogenous
material.
Origin of HTC: On the basis of the above discussion, we conclude that 13C-depleted (up to –30‰)
kerogenous material exists within magnetite of thequartz-magnetite BIF. In contrast with secondary
reduced carbon released mainly below 400°C, theprecursor of the HTC would have been incorporated
into magnetite during diagenesis. Previous studies
reported δ13C values of ≥–19‰ for graphite of premetamorphic origin from Isua metasediments (Rosing 1999; Ueno et al., 2002). In contrast, this study
first suggests that bedded magnetite in BIF from thenortheastern end of the ISB contains more 13Cdepleted carbon (up to –30‰) of primary origin.
Thus, it is important to discuss the origin of HTC as
well as HTN to test for early life.
In order to discuss the origin of the HTC, secondary isotope fractionation of the precursor of the HTC
during metamorphism was calculated. Attending the
stepwise combustion experiments of the BIF, trace
amounts of H2O (ppm) was also detected above
1000°C. Thus, it is possible that the precursor of
HTC (kerogenous material) reacted with H2O inmagnetite and lost CO2- and CH4-bearing fluid during metamorphism. Secondary isotope fractionationof the precursor of HTC due to loss of CO2- andCH 4 -bearing fluid was estimated employing a
Rayleigh distillation model. In the model, we usedfollowing equation:
δ13Cfinal – δ13Cinitial = (F α–1 – 1) × (δ13Cinitial + 1000)≈ (F α– 1 – 1) × 1000
Values of δ13Cfinal and δ13Cinitial refer to the δ13C
value of kerogenous material after metamorphism
and before metamorphism, respectively. F refers to
the fraction of kerogenous material remaining, andα is a carbon isotope fractionation factor of the Cbearing fluid relative to the remaining kerogenous
material. Detailed explanation of the metamorphic
temperature, total fluid pressure, oxygen fugacity,and α value are given in the Appendix.
As shown in Figure 7B, the direction of carbon
isotopic change of the kerogenous material duringmetamorphism depends on the CO2/CH4 molar ratioof the fluid (i.e., oxygen fugacity within magnetite).
It is possible that the δ13Cinitial value was more negative than the δ13Cfinal value in some cases (i.e.,
CO2/CH4 molar ratio ≤ 0.73 at T = 400°C). Here, theprecursor of HTC would have been more 13 Cdepleted than HTC, and HTC (13C-depleted up
to –30‰) is possibly derived from organic carbon
produced by early life. On the other hand, it is also
possible that the δ13Cinitial value is more positive
than the δ13Cfinal value (i.e., CO2/CH4 molar ratio >0.73 at T = 400°C). In these cases, the precursor of
HTC would have been more 13C-enriched thanHTC, and HTC (13C-depleted up to –30‰) possibly
would have been derived from abiological graphite
produced by siderite decomposition (δ13C value of –10 to –12‰; Van Zuilen et al., 2002). For example,
the value of δ13Cfinal – δ13Cinitial can be –18‰ when
the F value decreases to 0.20, due to loss of pure
CO2 at 400°C.
Although we cannot exclude the possibility ofabiological graphite as the origin of the HTC withthe present data set, the fact that the HTC and HTN
have δ13C values, δ15N values, and C/N elemental
ratios consistent with Archean kerogen (Beaumont
and Robert, 1999) indicates that HTC was notderived from abiological graphite. For precise
estimation of the isotopic change of the precursorof HTC and HTN during metamorphism, further
study is important to elucidate the relation
between content and isotopic ratio of HTC andHTN, respectively.

Conclusions
Ion microprobe REE geochemistry and U-Pb
geochronology of apatite in ~3.8 Ga metasediments,
together with carbon and nitrogen isotopic analysesof the quartz-magnetite BIF by stepwise combustion, provide the following new information.
1. The chondrite-normalized REE pattern of apatite in the quartz-magnetite BIF (43-44A) is flat witha positive Eu anomaly. This pattern is different fromthat of secondary apatite in carbonate rocks andISUA SUPRACRUSTAL BELT, WEST GREENLAND

965

FIG. 7. Calculated change in δ13C value of precursor of HTC (kerogenous material) due to loss of CO2- and CH4bearing fluid during metamorphism. The precursor of HTC is suggested to have coexisted with H2O, both within magnetite in BIF 43-44A. Rayleigh distillation model was used for the calculation. A. Relation between oxygen fugacity andmetamorphic temperature of the kerogenous material–H2O system within magnetite in the BIF. We assumed that oxygen
fugacity of the kerogenous material–H2O system is higher than that determined by FMQ buffer and lower than that
determined by graphite buffer. B. Calculated change of δ13C value of the kerogenous material due to loss of CO2- andCH4-bearing fluid at 400°C. The range of CO2/CH4 molar ratio of the C-bearing fluid was estimated within the range ofoxygen fugacity of the kerogenous material–H2O system determined by Figure 7A.

mafic dikes from the ISB, respectively (Lepland etal., 2002). On the other hand, the REE pattern suggests that the apatite was likely deposited fromhydrothermally influenced Early Archean seawater.

This is consistent with the primary depositional origin of the apatite and the BIF.
2. Apatite in the quartz-magnetite BIF is U-poor,and records various Pb isotopic ratios, which may be

966

NISHIZAWA ET AL.

explained by mixing of two model age components(~3.8 Ga and ~1.5 Ga). This suggests that thequartz-magnetite BIF was affected by metamorphism at 1.5 Ga.
3. Apatite in the psammitic schist (K485) showsa U-Pb isochron of 1.5 ± 0.3 Ga, supporting metamorphism at 1.5 Ga. This age is also consistent witha Rb-Sr age of 1.6 Ga for muscovite-phengite samples from pegmatitic gneiss in the Isukasia area
(Baadsgaard et al., 1986).
4. Stepwise combustion experiments suggest thatquartz-magnetite BIF (43-44A) contains two distinctive components of carbon. One is released
below 1000°C (mostly released between 200° and
400°C); it probably exists along grain boundaries,
and is clearly contaminant incorporated after themetamorphism. The other is released above 1000°C,
and is probably included in magnetite which occurs
concordantly with bedding. Thus, it is plausible thatthe precursor of the high-temperature carbon wasincorporated into magnetite during diagenesis of the
BIF (~3.8 Ga).
5. Significant 13C depletion (–30‰), as well as
negative δ15N values (–3‰), and C/N ratios (86) of
the high-temperature fraction for one aliquot of
the BIF are within the range of those of Archean
kerogen in metasediments. This suggests that theseelements were derived from kerogenous material.
These lines of evidence indicate that the apatitebearing quartz-magnetite BIF may preserve a
primary depositional signature at ~3.8 Ga. Despite
the presence of secondary carbon (i.e., LTC), theBIF is suggested to contain pre-metamorphic 13Cdepleted carbon (up to –30‰), entrapped bymagnetite during diagenesis.

Acknowledgments
We thank R. Yokochi and T. Monde for analyticalassistance; T. K. Dalai for improvement of the
English; S. Maruyama and H. Shimizu for discussions; A. P. Nutman and H. Yoshioka for providing
sedimentary rock samples; and K. Ludwig for the
Isoplot/Ex. computer software. We thank M. T. Rosing for a critical review of an early draft of the manuscript. Constructive review by E. G. Nisbet greatly
improved the manuscript. YU is grateful for the
Research Fellowships of the Japan Society for the
Promotion of Science for Young Scientists. This is acontribution of a joint project between the SHRIMP
laboratory at Hiroshima University and the Center

for Advanced Marine Research at the Ocean
Research Institute, University of Tokyo.

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Appendix. Calculation of Carbon Isotopic Change of Precusor of HTC due to loss of
CO2- and CH4-Bearing Fluid During Metamorphism
In the paper, we calculated carbon isotopic
change of precursor of HTC (kerogenous material)
due to loss of CO2- and CH4-bearing fluid, in order
to discuss secondary isotope fractionation of the precursor of HTC during metamorphism. The precursorof HTC is suggested to have coexisted with H2O,
both within magnetite in quartz-magnetite BIF (4344A) from the ISB. A Rayleigh distillation model
was used for the calculation. In the model, we usedfollowing equation:
δ13Cfinal – δ13Cinitial = (Fα–1 –1) × (δ13Cinitial + 1000)≈ (Fα–1 –1) × 1000.
δ13Cfinal and δ13Cinitial refer to the δ13C values of kerogenous material after metamorphism and before metamorphism, respectively. F refers to the fraction of
kerogenous material remaining, and α is the isotope
fractionation factor of the C-bearing fluid relative tokerogenous material remaining. The calculation was
performed under following conditions.

Metamorphic temperature
It had been long considered that the ISB underwent uniform regional metamorphism (Boak andDymek, 1982). These authors concluded that the
ISB uniformly experienced ~550°C during prograde
metamorphism and later was subjected to ~460°C
during retrograde metamorphism. This is because
metamorphic temperature of metapelite from the
ISB, calculated by garnet-biotite geothermometry,suggests equilibration at ~550°C for garnet cores,
and at ~460°C for garnet rims.
However, it would be inappropriate to use themetamorphic conditions to calculate the carbon isotopic change of the kerogenous material within the
BIF during metamorphism. This is because petrochemical study of garnet from each structural
domain in the ISB suggests that metamorphic history of the domains that host the metapelite (Domain
II, III, and IV of Rollinson, 2003) is different fromthat of the domain that hosts the BIF (Domain I of
Rollinson, 2003).
Based on petrochemical and geothermobarometric studies of more than 1500 rock samples in
Domain I, progressive metamorphic zonation was
discovered in the domain (Hayashi et al., 2000;
Komiya et al., 2002). The grade and temperature ofeach metamorphic zone in the domain is as follows:(1) Zone A (greenschist facies); (2) Zone B (albiteepidote-amphibolite facies; 400–420°C; Hayashi etal., 2000); (3) Zone C (lower amphibolite facies: up
to 480°C; Appel et al., 2001); (4) Zone D (upper
amphibolite facies: up to 550°C; Hayashi et al.,2000).
The metamorphic temperature of each zone wasestimated from garnet-biotite geothermometry(Hayashi et al., 2000, Appel et al., 2001). Because
most garnets from Domain I show normal zoning anddo not contain evidence of retrogression (Hayashi etal., 2000; Rollinson, 2003), the estimated temperature of each zone probably reflects the temperature
at peak metamorphic conditions. Because the BIF isfrom Zone A, the metamorphic temperature of the
BIF is suggested to be ~400°C. Thus, we calculated
the carbon isotopic change of the kerogenous material within the BIF at 400°C.

Total fluid pressure
We assumed that the total fluid pressure equals
the metamorphic pressure. A pressure of ~5 kbar for
Domain I was estimated from garnet-hornblendeplagioclase-quartz geobarometry (Hayashi et al.,2000). Thus, we used 5 kbar for our calculation.
Oxygen fugacity and α value
We assumed that oxygen fugacity of the kerogenous material-H 2 O system within magnetite ishigher than that determined by the FMQ buffer, andlower than that determined by the graphite buffer(Fig. 7A). Within the range of oxygen fugacity of the
kerogenous material–H2 O system, the CO 2/CH 4

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molar ratio of fluid released from the kerogenous
material was estimated (Ohmoto and Kerrick, 1977).
To calculate the α value, we used αCO2-graphite andαCH4-graphite values from Chacko et al. (1991) andHorita (2001), respectively.