Anoxic dissolution processes of biotite: implications for Fe behavior

Earth and Planetary Science Letters 224 (2004) 117 – 129
www.elsevier.com/locate/epsl
Anoxic dissolution processes of biotite: implications for Fe
behavior during Archean weathering
Takashi Murakami a,*, Jun-Ichi Ito a, Satoshi Utsunomiya a,1, Takeshi Kasama a,2,
Naofumi Kozai b,3, Toshihiko Ohnuki b,3
a
Department of Earth and Planetary Science, The University of Tokyo, 7-3-1 Hongo, Tokyo 113-0033, Japan
b
Japan Atomic Energy Research Institute, Tokai-mura, Ibaraki-ken 319-1195, Japan
Received 17 December 2003; received in revised form 16 April 2004; accepted 29 April 2004
Abstract
Iron-rich biotite (Fe/Mg = 5) dissolution experiments were carried out in a batch system under O2-deficient conditions
( PO2 < 3 10 5 atm; referred to as ‘anoxic’ conditions) at 1 atm of PCO2, pH 4.6, and 100 jC for 7 – 120 days for a better
understanding of ‘anoxic’ weathering processes and Fe behavior during weathering before 2.2 Ga. For comparison, oxic Fe-rich
biotite dissolution experiments were conducted under present atmospheric conditions at pH 4.7 and 100 jC for 7 – 80 days
(referred to as oxic conditions) by using the same starting biotite as that for the ‘anoxic’ experiments. The concentrations of Fe
in solution after the dissolution experiments were larger by one to more than two orders of magnitude under ‘anoxic’ conditions
than under oxic conditions. High-resolution scanning and transmission electron microscopy revealed that Fe(II)-rich vermiculite
or smectite was precipitated as a secondary mineral at the edge of biotite under ‘anoxic’ conditions, in contrast to the formation
of Fe(III)- and Al-(hydr)oxides under oxic conditions. The results suggest that part of Fe(II) is released to water as ‘anoxic’
weathering proceeds, explaining the decrease of Fe content in pre-2.2 Ga paleosols relative to their parent rocks. The Fe/Mg
molar ratio of the secondary vermiculite or smectite was more than 7 while the starting Fe-rich biotite had a value of about 5.
The Fe/Mg molar ratio was less than 2.5 in solution. The results of the ‘anoxic’ experiments suggest that Fe(II)-rich vermiculite
or smectite could be the precursor of the chlorite preserved in pre-2.2 Ga paleosols. This is further corroborated by the increase
in Fe/Mg molar ratios in chlorite with decreasing depth in some Precambrian paleosols.
D 2004 Elsevier B.V. All rights reserved.
Keywords: Precambrian; atmospheric evolution; biotite dissolution; paleosol; anoxic weathering
* Corresponding author. Tel.: +81-3-58414541; fax: +81-3-58414555.
E-mail addresses: [email protected] (T. Murakami), [email protected] (J.-I. Ito), [email protected] (S. Utsunomiya),
[email protected] (T. Kasama), [email protected] (N. Kozai), [email protected] (T. Ohnuki).
1
Present address: Geological Sciences, University of Michigan, 425 East University, Ann Arbor, MI 48109-1063, USA.
2
Present address: Department of Materials Science and Metallurgy, University of Cambridge, Pembroke Street, Cambridge CB2 3QZ, UK.
3
Tel.: +81-29-2825361; fax: +81-29-2825927.
0012-821X/$ - see front matter D 2004 Elsevier B.V. All rights reserved.
doi:10.1016/j.epsl.2004.04.040
118
T. Murakami et al. / Earth and Planetary Science Letters 224 (2004) 117–129
1. Introduction
Atmospheric evolution during the Precambrian is
still an ongoing topic including the ‘‘faint young Sun
paradox’’ [1]. Two of the most important atmospheric
components, CO2 and O2, have been closely related to
water – rock interactions on Earth’s surface (e.g., [2]).
Paleosols, ancient soils formed by weathering, can
provide data regarding the composition of the atmosphere, especially O2 content at the time they formed
[3]. Despite numerous studies on Precambrian paleosols, the evolutionary history of O2 is not fully
deduced. Such a comprehension is essential because
O2 evolution has significant bearing on early evolution of life on Earth (e.g., [4,5]). Two contrasting
models of atmospheric O2 evolution prevail: One
proposes that atmospheric PO2 was less than approximately 10 3 atm before 2.2 Ga and changed dramatically to more than approximately 10 2 atm sometime
between 2.2 and 2.0 Ga (e.g., [2,3,6 – 8]); the other
insists that PO2 has essentially remained constant since
3.0– 4.0 Ga (e.g., [9,10]).
Rye and Holland [3] recently reviewed data of
more than 50 reported paleosols and concluded that
the oxidation state of paleosols changed significantly
at about 2.2 Ga. The most important criteria for
determining the oxidation state of paleosols are Fe
contents and ferrous to ferric ratios in weathering
profiles [11]. However, Ohmoto [10,12] has proposed
that oxidized profiles can be reduced readily by postweathering hydrothermal alteration or organic acids.
These two contrasting interpretations arise because
post-weathering diagenesis and metamorphism commonly have obscured the geochemical and mineralogical features of weathering profiles. For instance, Kaddition resulting in the formation of sericite is
ubiquitous in paleosols formed before early Proterozoic time [13 –18]. If weathering processes before 2.2
Ga can be reconstructed, a better understanding of
paleosols, and thus, the evolution of atmospheric O2,
will result.
With this in the backdrop, we carried out Fe-rich
biotite (Fe/Mg = 5) dissolution experiments under ‘anoxic’ conditions ( PO2 < 3 10 5 atm) to examine Fe
behavior during weathering before 2.2 Ga. Note the
term ‘anoxic’ used in the present study implies PO2 of
approximately < 10 5 atm. Although the regional
abundance of biotite is not large (7.6% of the exposed
continental crust surface [19]), biotite is a major
source of Fe in ground water (e.g., [20]). In modern
weathering, biotite has been observed to be altered to
other sheet silicates depending on the alteration conditions in nature ([21] and references therein), and in
particular, to vermiculite and kaolinite ([22] and
references therein). Murakami et al. [23] have compared Fe-rich biotite dissolution processes between
nature and the laboratory under present atmospheric
conditions. It has been found that almost all Fe(II)
dissolved from biotite forms Fe(III)-(hydr)oxides and
very little vermiculite is formed because of less
availability of Mg in solution. In contrast to extensive
studies on biotite weathering under oxic conditions
[24], only a few studies have been carried out to
investigate the behavior of Fe(II) in solution during
biotite weathering under ‘anoxic’ conditions [25]. Rye
et al. [26] have proposed that siderite was not a
secondary Fe-bearing mineral during 2.2 – 2.8 Ga
weathering, and that the atmospheric PCO2 was therefore less than about 10 2 atm. Rye and Holland [27]
have suggested that Fe(II)-rich trioctahedral smectite
could be a common weathering product at about 2.2
Ga, and provide a source for chlorite in paleosols.
Because weathering products reflect the nature of the
ambient atmosphere as demonstrated by Murakami et
al. [28], an understanding of weathering processes
under ‘anoxic’ conditions is necessary to evaluate O2
evolution. Iron-rich biotite dissolution experiments
under ‘anoxic’ conditions were undertaken to gain a
better understanding of paleosols, and thus of Precambrian atmospheric O2 evolution.
2. Experimental and analytical methods
2.1. Samples
The biotite samples used for the present study were
from a coarse-grained granite in Inada, central Japan
[29], having a chemical formula of (K0.91Na0.01)
(Al 0 . 1 4 Mg 0 . 4 0 Fe 2 . 0 7 Mn 0 . 0 5 Ti 0 . 1 9 )(Si 2 . 8 2 Al 1 . 1 8 )
O10(OH)2 [23]. The molar ratio of Fe(II) to Fe(III) of
the biotite was about 9:1 as determined by X-ray
absorption near-edge structure analysis using synchrotron radiation at beam line BL-4A, Photon Factory,
National Laboratory for High Energy Physics, Tsukuba, Japan (A. Monkawa, personal communication).
T. Murakami et al. / Earth and Planetary Science Letters 224 (2004) 117–129
The biotite contained no interstratified alteration products except for 1– 2% of chlorite that was distributed
randomly in the biotite as observed by high-resolution
transmission electron microscopy (HRTEM) [23]. The
biotite samples were collected by crushing the granite
and handpicking. The biotite samples were crushed
into 55- to 106-Am-sized grains, and washed ultrasonically with acetone for 5 min to remove fine particles on
the surface of the biotite grains prior to the dissolution
experiments described below. The surface of the biotite
grains after ultrasonication was smooth and almost free
of fine particles [23].
2.2. Anoxic experiments
‘Anoxic’ biotite dissolution experiments were carried out considering the atmosphere before the proposed dramatic increase in atmospheric oxygen
content (e.g., [8]). The biotite dissolution was done
in a glove box where the O2 concentration was kept at
less than 1 ppm at room temperature. The O2-deficient
conditions achieved in the glove box are hereafter
called ‘anoxic’ conditions. Fig. 1 shows a schematic
setup for the ‘anoxic’ experiment; biotite grains were
reacted with water in reaction vessels placed in an
Fig. 1. Schematic setup for ‘anoxic’ experiments. A close-up in the
oven is given in the right bottom.
119
oven that was set in the glove box. Four reaction
vessels at maximum can be placed in the oven. The
glove box was equipped with a path box that enabled
us to bring materials in and out without affecting the
O2 concentration significantly.
The glove box was constructed to allow Ar gas to
be circulated through metal Cu and a molecular sieve
to remove O2 and other gases. The variation of O2
concentration in the glove box was monitored during
first 30 days of the ‘anoxic’ experiments, and ranged
from 0.35 to 0.92 ppm (equivalent to 4.2 10 7 to
1.1 10 6 atm of PO2 at room temperature). We
introduced CO2 to Teflon reaction vessels [30,31] in
stainless steel containers that were placed in the
electric oven kept at 100 ± 5 jC. The value of PCO2
was set at 1 atm at 100 jC for the ‘anoxic’ experiments (e.g., [8]), calculated from the pressure shown
by the regulator of a CO2 cylinder by subtracting the
value of the saturated vapor pressure. Impurities in the
CO2 were less than 20 ppm of O2, 50 ppm of N2, 10
ppm of H2, and 10 ppm of CH4. O2, 20 ppm, was
equivalent to about 3 10 5 atm of PO2 at 100 jC.
Thus, the actual PO2 and PCO2 in the Teflon reaction
vessels were 3 10 5 atm at maximum and 1 atm,
respectively, at the experimental temperature of 100
jC. The value of PO2, 3 10 5 atm, is lower than the
value of 5 10 4 atm estimated for the atmosphere
prior to 2.4 Ga [3].
All solutions used for the present study were
prepared with deionized water (18.2 MV) and
reagent-grade chemicals. The subsequent processes
of ‘anoxic’ biotite dissolution were done in the glove
box. Deionized water was brought into the glove box,
and bubbled with Ar gas in the glove box for 1 day in
advance of dissolution experiments to be equilibrated
with the atmosphere in the glove box. The resultant
deionized water was used as a reactant solution. For
each run, a 10-mg batch of crushed biotite grains was
reacted with 10 ml of deionized water in the Teflon
reaction vessel in the glove box. The ‘anoxic’ experiments were carried out at 100 F 5 jC for 7, 8, 14, 40,
and 120 days (five runs for total). The mass loss of the
solution during the experiments was less than 5%.
After each ‘anoxic’ experiment, the vessel was cooled
to room temperature in 30 min by an electric fan to
avoid possible precipitation (e.g., Si and Al products)
from the solution during quench. Murakami et al. [32]
carried out anorthite dissolution experiments at 90,
120
T. Murakami et al. / Earth and Planetary Science Letters 224 (2004) 117–129
150, and 210 jC using an experimental procedure
similar to that described above, and confirmed by
TEM no precipitation of Si or Al products during the
quench. The solution was then separated from the
solids by filtration through a 0.22-Am filter, and a
solution with 10 wt.% nitric acid was added so that the
final solution contained 1 wt.% nitric acid to lower the
pH and preserve the metals for analysis. The solids
were then washed gently in acetone and dried. The
solutions and solids were finally taken out of the
glove box for analysis.
Most CO2 introduced in the reaction vessel and
solution would be released into the atmosphere of the
glove box when the vessel was opened, which must
change the pH of the solution. We, therefore, did not
measure the pH of the solutions after the ‘anoxic’
experiments. Instead, we calculated the values of pH
by EQ3NR [33] assuming H+ was balanced with other
cations and anions in the solutions during the ‘anoxic’
experiments. The pH of a pure CO2 solution at 1 atm
of PCO2 and 100 jC was about 4.2 (e.g., [34]). The pH
of the solutions after the ‘anoxic’ experiments for 7 to
120 days was calculated to be about 4.6 at 1 atm of
PCO2 and 100 jC, when dissolved cation concentrations were taken into account (Table 1).
2.3. Oxic experiments
The conditions under which the following dissolution experiments were carried out outside the glove
box are hereafter referred to as oxic conditions.
Deionized water with a buffer of sodium acetate
(0.03 mol/l) and acetate was used as a reactant
solution for the oxic experiments. Acetate is often
used as a pH buffer (e.g., [35]), and has no significant
effect on dissolution [36]. The ionic strength of the
reactant solution was 0.067. The pH of the solution
was adjusted to 4.5 by acetate at room temperature.
The pH, 4.5, was increased to 4.7 at 100 jC, which
was similar to 4.6 for the ‘anoxic’ experiments. A 10mg batch of the crushed biotite grains was reacted
with 10 ml of the solution in a Teflon reaction vessel
in an oven. The oven was not placed in the glove box.
The oxic experiments were carried out at 100 F 5 jC
for 7, 21, and 80 days without CO2 introduction (three
runs for total). Treatment of the solids and solutions in
the oxic experiments was the same as that of the
‘anoxic’ experiments, except that they were done
under ambient atmosphere conditions.
2.4. Analytical methods
The biotite samples were subjected to powder Xray diffraction analysis (monochromatized CuKa radiation at 40 kV and 20 mA, Rigaku RINT 2000)
before and after the ‘anoxic’ and oxic experiments, to
check the formation of secondary minerals. However,
no change in X-ray diffraction patterns was observed
before and after the experiments. Biotite grains and
their polished samples after the 40- and 120-day
‘anoxic’ and 80-day oxic experiments were observed
and analyzed by field emission scanning electron
microscopy (FESEM, Hitachi S4500) equipped with
an energy dispersive X-ray spectrometer (EDS, Kevex
Table 1
Cation concentrations (mol/l) and Fe/Mg molar ratios in solutions after the ‘anoxic’ dissolution experiments
Duration
(days)
Si
7
8
14
40
120
7a
8a
14a
40a
120a
3.3 10
4.4 10
4.5 10
4.5 10
4.8 10
1.2 10
1.6 10
1.6 10
1.6 10
1.7 10
a
Al
4
4
4
4
4
4
4
4
4
4
7.7 10
1.5 10
1.2 10
1.2 10
1.2 10
5.8 10
1.1 10
9.1 10
9.1 10
9.1 10
Fe
6
5
4
4
4
6
5
5
5
5
2.3 10
3.7 10
4.1 10
5.2 10
5.8 10
1.1 10
1.8 10
2.0 10
2.5 10
2.8 10
Mg
5
5
5
5
5
5
5
5
5
5
2.6 10
4.3 10
1.8 10
2.1 10
2.5 10
6.5 10
1.1 10
4.5 10
5.3 10
6.3 10
K
5
5
5
5
5
5
4
5
5
5
1.2 10
1.8 10
2.4 10
9.4 10
3.6 10
1.3 10
2.0 10
2.6 10
1.0 10
4.0 10
Fe/Mg
4
4
4
5
4
4
4
4
4
4
0.88
0.86
2.3
2.5
2.3
0.17
0.16
0.44
0.47
0.44
Normalized cation concentrations based on the biotite chemical formula, (K0.91Na0.01)(Al0.14Mg0.40Fe2.07Mn0.05Ti0.19)(Si2.82Al1.18)O10
(OH)2: a cation concentration in the upper half of the table was divided by the number of moles of the corresponding cation per O10(OH)2.
T. Murakami et al. / Earth and Planetary Science Letters 224 (2004) 117–129
system). Morphological changes and secondary mineral formation were observed by FESEM, and the
compositions of secondary minerals were determined
qualitatively by FESEM-EDS. Polished samples were
made using the following procedure: some of biotite
grains were embedded in epoxy resin, and mechanically polished by diamond paste. The accelerating
voltage applied was 3 – 15 kV for secondary and
backscattered electron imaging and elemental analysis. The lower accelerating voltages were employed to
obtain images of secondary minerals of submicron
size, and to analyze the minerals. The Fe/Mg molar
ratios of secondary minerals were semiquantitatively
analyzed by EDS operated at 10 kV, using the fresh,
starting biotite, of which the chemical composition
was determined by electron probe microanalysis [23],
as a standard, and using software MAGIC V installed
in the Kevex system. The semiquantitative analysis
was made for the polished samples.
For the 40- and120-day ‘anoxic’ experiments, the
biotite grains were further examined by HRTEM
(JEOL 2010) equipped with EDS (Kevex system) to
observe secondary minerals formed at the surface and
within the biotite grains. The HRTEM had a point
resolution of 0.2 nm, and was operated at 200 kV. The
TEM specimens for the 40-day ‘anoxic’ experiment
were prepared by producing suspension of secondary
products ultrasonically from the biotite grains and
putting the suspension on a TEM grid. The TEM
specimens for the 120-day ‘anoxic’ experiment were
prepared by impregnating the biotite grains in epoxy
resin, squeezing them between two glass slides, slicing and polishing them mechanically, and thinning
them to electron transparency by Ar ion milling
(Gatan Dual Ion Mill). Some TEM negatives were
digitized and processed to remove both noise and any
images of amorphous material by rotational filtering
[37]. The use of rotationally filtered images, utilizing
the software Digital Micrograph V. 2.5 (Gatan), has
been described in detail by Banfield and Murakami
[38]. The Fe/Mg molar ratios of secondary minerals
were semiquantitatively analyzed by EDS operated at
200 kV, using MAGIC V (standardless, thin-film
analysis) installed in the Kevex system.
Silicon, Al, K, Fe, and Mg concentrations in the
solutions after the ‘anoxic’ experiments were measured by inductively coupled plasma atomic emission
spectrometry (ICP-AES, Seiko SPS7700). Silicon and
121
Fe concentrations after the oxic experiments were also
measured by ICP-AES to compare them with those of
the ‘anoxic’ experiments. The quantitative detection
limit of the present ICP-AES was about 10 6 mol/l for
Al and about 10 7 mol/l for Fe.
3. Results
Fig. 2 compares the variations of Si and Fe concentrations with time between the ‘anoxic’ and oxic
experiments. Although the Si concentrations in the
first 20 days were slightly different between the
‘anoxic’ and oxic experiments, which leads to an
apparent difference in dissolution rate, the Si concentration variations were similar between the ‘anoxic’
and oxic experiments. The Si concentrations already
increased to 1 – 3 10 4 mol/l for the first 7-day
experiments, and then, remained almost constant. In
contrast to the Si variations, the Fe variations were
different: for the ‘anoxic’ experiments, the Fe variation was similar to that of Si, although the Fe
concentration was lower by one order of magnitude
than the Si concentration. For the oxic experiments,
the Fe concentration was very low (about 10 7 mol/l)
or below the quantitative detection limit.
Ion concentrations of the solutions after the ‘anoxic’
experiments are presented in Table 1. Biotite showed
incongruent dissolution apparently for the ‘anoxic’
experiments, and the dissolution rates based on the
biotite chemical formula were mostly in the order of
Fig. 2. Variations of Si and Fe concentrations in solution with time,
under ‘anoxic’ and oxic conditions. ND stands for a value under the
detection limit.
122
T. Murakami et al. / Earth and Planetary Science Letters 224 (2004) 117–129
K>Si>Al>Mg>Fe after 14 to 120 days (Table 1). The
apparent incongruent dissolution occurred mainly because of the secondary mineralization described below.
It has been reported that incongruent dissolution of
biotite occurs, even with flow-through type reactors
under the present atmosphere [25,39,40]. The Fe/Mg
molar ratio was about 5 in the fresh biotite, whereas it
was less than 2.5 in the solutions, or less than 0.5 after
normalization based on the biotite chemical formula
(Table 1). This indicates that more Mg was redistributed into solution than Fe by ‘anoxic’ dissolution.
Fresh biotite before the oxic and ‘anoxic’ experiments had a distinctive texture without any secondary
minerals at the edge (Fig. 3A). The basal surface did
not show any significant changes after both the
experiments as reported in the previous biotite dissolution study [23]. Edges of biotite grains were covered
with a layer of two types of fine particles of submicron in diameter after 80 days for the oxic experiment
(Fig. 3B and C). Both morphology and components
(only Fe and O) indicate that one type of the fine
particles (Fig. 3B) is crystalline Fe(III)-(hydr)oxides
[23]. The other type of the fine particles (Fig. 3C)
consisted of aggregates that contained Fe with less
amount of Al (Fig. 3D). The aggregates are probably
amorphous Fe(III)- and Al-(hydr)oxides. Silicon and
some of the Al in Fig. 3D are inferred to be from the
underlying biotite because the layer of the aggregates
is submicron thick and the X-ray can be generated
from the underlying biotite. The presence of a large
amount of Fe(III)-(hydr)oxides is consistent with the
near absence of Fe in the solution ( < 10 7 mol/l):
most Fe(II) dissolved from the biotite was oxidized to
Fe(III) and precipitated as Fe(III)-(hydr)oxides during
the oxic experiments. Similar observations for biotite
dissolution were made by Murakami et al. [23], who
reported boehmite as a secondary mineral in addition
to hematite.
Edges of biotite grains were covered with a layer of
materials irregular in shape after 40 days for the
‘anoxic’ experiment (Fig. 4A). The materials contained
Fe, Al, and Si (Fig. 4B); some of Al and Si can be
generated from the underlying biotite. Fig. 4C shows a
TEM image of the materials, indicating the presence of
a mineral with 1.4-nm fringes. Further identification of
the mineral was not made at this point.
X-ray diffraction analysis did not show any alteration of the biotite after 120 days of the ‘anoxic’
experiment. We observed more than 30 biotite grains
by FESEM, and found that the edges of biotite
grains were covered with clay-like layers 0.2– 0.5
Am thick (Fig. 5A and B). The secondary mineral-
Fig. 3. Secondary electron images at the edges of biotite grains before experiment (A) and after 80-day, oxic experiment (B and C), and an
energy dispersive X-ray spectrometry (EDS) profile of the secondary products in C (D). The operating voltages were 5 kV for A and B, 3 kV for
C, and 10 kV for D.
T. Murakami et al. / Earth and Planetary Science Letters 224 (2004) 117–129
123
Fig. 4. Secondary electron image at the edge of a biotite grain after 40-day, ‘anoxic’ experiment (A), its EDS spectrum (B), and transmission
electron microscopy (TEM) image of secondary products formed at the edge (C). The operating voltages were 3 and 5 kV for A and B,
respectively.
ization occurred only at the edges and not at the
basal surface. Fig. 5B shows a backscattered electron
image of a polished specimen of such biotite grains;
the gray contrast (arrows in Fig. 5B) at the edge of
biotite corresponds to the clay-like materials in Fig.
5A. The lower contrast of the clay-like materials in
Fig. 5B, although they have a higher Fe content (Fig.
6), is due to a lower density than that of biotite.
EDS spectra of the clay-like materials and fresh
biotite are compared in Fig. 6A and B, respectively; the
analytical point of the clay-like materials is shown by
the thick arrow in Fig. 5B. Because we analyzed the
materials at an operating voltage of 10 kV, electron
penetration in the sample was about 1 Am. Consequently, analysis of the clay-like materials could be affected
by the neighboring biotite because of the thickness of
the clay-like layers. The low operating voltage of 10 kV
lowers apparent intensities at higher energies. Therefore, FeL at about 0.7 keV should be compared for the
Fe concentrations between the two EDS spectra. The
124
T. Murakami et al. / Earth and Planetary Science Letters 224 (2004) 117–129
Fig. 6A. The semiquantitative analysis revealed that
the clay-like materials are richer in Fe than the fresh
biotite; the Fe/Mg molar ratio of the clay-like materials, 7.2, is larger than that of the fresh biotite, 5.2. The
actual Fe concentration of the clay-like materials may
be higher, because of the larger interaction volume of
the electron beam with the neighboring biotite.
The clay-like materials were further observed by
HRTEM. Fig. 7A shows an example of the precipitation of the clay-like materials (light-gray contrast in
the left of Fig. 7A) on the edge of a biotite grain
(darker contrast in the middle and the right of Fig.
7A). The upper and lower areas with light-gray
contrast are resin. Fig. 7B is a lattice-fringe image
of the clay-like materials on the edge of the biotite
grain in Fig. 7A. The interplanar-spacing of the claylike materials is 1.4 nm (Fig. 7B). The cross fringes
(inset in Fig. 7B) indicate that these clay-like materials have a 0.5-nm periodicity approximately normal to
the 1.4-nm fringes. The interplanar-spacings (Fig. 7B)
Fig. 5. Secondary (A) and backscattered (B) electron images of
secondary products after 120-day, ‘anoxic’ experiment, at the edge
of biotite and for polished thin section, respectively. Arrows show
clay-like materials (of which the layer is as thin as 0.2 – 0.5 Am, with
dark contrast) formed on the edge of a biotite grain (white contrast
with black lines that are parallel to the sheet structure of biotite).
The analytical point of the clay-like materials by EDS is shown by
the thicker arrow (see Fig. 6).
composition of the clay-like materials is characterized
by higher Fe and lower K than those of the fresh biotite.
Potassium could be contained in the clay-like materials,
or could result from the neighboring biotite.
We obtained more than 10 EDS spectra of the claylike materials by SEM-EDS. Most clay-like layers
were as thin as 0.2 Am, and the spectra were significantly affected by neighboring biotite. The area
shown by the thick arrow in Fig. 5B was the thickest
part (as thick as 0.5 Am) of the clay-like layers, and
thus, the EDS analysis of this area was least affected
by neighboring biotite. Therefore, we made semiquantitative analysis of this area, i.e., the EDS spectrum in
Fig. 6. EDS spectra of clay-like materials after 120-day, ‘anoxic’
experiment (A) and fresh biotite (B). See for the analytical point of
the clay-like materials in Fig. 5B. The analysis was done at an
operating voltage of 10 kV by SEM-EDS.
T. Murakami et al. / Earth and Planetary Science Letters 224 (2004) 117–129
Fig. 7. TEM image of clay-like materials formed on the edge of a
biotite grain (A) and a close-up of the clay-like materials (B) after
120-day, ‘anoxic’ experiment. The interplanar-spacings, 1.4 and 0.5
nm, are those of the clay-like materials.
and the chemical components (Fig. 6A) indicate that
the clay-like materials are either vermiculite or smectite with a higher Fe/Mg molar ratio than that of the
fresh biotite. Semiquantitative analysis of vermiculite
or smectite domains showed that the Fe/Mg molar
ratio was about 9 (8.9 F 1.4 for seven analytical
points). The value of the Fe/Mg molar ratio is consistent with that obtained by SEM-EDS.
We cannot distinguish vermiculite from smectite
without a quantitative chemical composition. The 1.4nm fringes in Fig. 7B are different from those of
chlorite that displays a periodicity of one thin and
one thick dark fringes; the thick dark fringe can be split
125
into two or three dark fringes (e.g., [41]). No Fe(III)(hydr)oxides were observed, even by HRTEM. The
identification described above strongly suggests that
the mineral with 1.4-nm fringes formed after the 40day, ‘anoxic’ experiment (Fig. 4) is vermiculite or
smectite with a high Fe/Mg molar ratio.
In addition to the precipitation of vermiculite or
smectite at the edge of biotite after the 120-day
‘anoxic’ experiment (Fig. 5), secondary mineralization also occurred within biotite grains. Fig. 8A and
B shows rotationally filtered HRTEM images of
biotite in the atomic scale after the 120-day ‘anoxic’
experiment, and Fig. 8C gives simulated images of
biotite, chlorite, and vermiculite (or smectite) for
comparison. The detailed description of calculations,
and the interpretation of the simulated images have
been presented in Murakami et al. [23]. Interlayers
with white contrast were observed between biotite
layers with and without impurity chlorite layers
(VSs in Fig. 8A and B, respectively). Such interlayers were scattered between biotite layers as
shown in Fig. 8A and B. Vermiculite or smectite
is distinguished from biotite by the contrast of the
interlayers; white lines for vermiculitic or smectitic
interlayers and gray spots alternating with white
spots for biotite interlayers (Fig. 8C). The interlayers with white lines (Fig. 8A and B) clearly
indicate the formation of vermiculite or smectite
within biotite grains. Such formation occurs as a
result of layer-by-layer formation of vermiculite or
smectite from biotite [42], i.e., K cations in one
biotitic interlayer are replaced by other cations such
as Mg bound to water molecules without changing
the silicate layers drastically, which forms one
vermiculitic or smectitic interlayer. The direct transition from biotite to vermiculite or smectite along
one interlayer is observed between the two arrows
in Fig. 8B; the interlayer at the top part is still
biotitic, i.e., an interlayer with gray spots alternating
white spots, while the interlayer at the bottom
vermiculitic or smectitic, i.e., an interlayer with a
white line. Thus, secondary vermiculite or smectite
was formed not only outside, but also within, the
biotite grains. However, the amount of the secondary mineral within biotite grains was much less than
at the edge, and the secondary mineral forming at
the edges is not necessarily identical in crystal
structure or chemistry to that replacing biotite
126
T. Murakami et al. / Earth and Planetary Science Letters 224 (2004) 117–129
4. Discussion
4.1. Dissolution and weathering of biotite under
anoxic conditions
Fig. 8. [11̄0] rotationally filtered high resolution TEM images of a
biotite sample after 120-day, ‘anoxic’ experiment showing vermiculite or smectite formation in biotite layers with and without a
chloritic interlayer (A and B, respectively), and [11̄0] simulated
images of vermiculite or trioctahedral smectite (VS), chlorite (Ch),
and biotite (Bi) (C). The detailed description of simulation is given
in Murakami et al. [23]. A chloritic interlayer is shown by Ch with
an arrow, and a vermiculitic or smectitic interlayer by VS. O, T, I,
and H in C, respectively, denote an octahedral sheet, a tetrahedral
sheet, an interlayer, and a hydroxide sheet. The columns of
tetrahedral cations (black spots near VS with an arrow in A) are
shifted with one another across the interlayer for the observed image
of vermiculite or smectite, which is slightly different from the
calculated image. Such shift is commonly observed for biotite-tovermiculite transformations (e.g., [42]). The vermiculitic interlayer
shown by two arrows in B indicates the direct transition from biotite
to vermiculite or smectite along one interlayer; gray spots
alternating white spots characteristic of the biotitic interlayer (I of
biotite (Bi) in C) are still visible in the top of the interlayer while
they are replaced by a white line characteristic of the vermiculitic
interlayer (I of vermiculite or smectite (VS) in C) in the bottom. The
d001-spacing of vermiculite or smectite is 1.0 nm in this figure,
which is compared to 1.4 nm in Figs. 4 and 7. This is due to the
collapse of the interlayers of vermiculite or smectite at high
magnification, i.e., high electron beam concentration of TEM.
layers. Note that fresh biotite was already observed
intensively by HRTEM and no single layer of
vermiculite or smectite was found within the fresh
biotite grains [23].
The concentrations of Fe in solution after the
dissolution experiments were larger by one to more
than two orders of magnitude under ‘anoxic’ conditions than under oxic conditions (Fig. 2). Under oxic
conditions, Fe(II) oxidation to Fe(III) and the subsequent formation of Fe(III)-(hydr)oxides are fast [43],
as has been observed in the current experiments (Figs.
2 and 3). The long-time persistence of Fe in solution
under ‘anoxic’ conditions (Fig. 2) strongly suggests
that dissolved Fe(II) from biotite is not oxidized to
Fe(III) and remains mostly as Fe(II) in solution under
‘anoxic’ conditions. Thus, as expected previously, Fe
flows out of a weathered zone under ‘anoxic’ conditions more than under oxic conditions. The Fe/Mg
molar ratio was about 5 in the fresh biotite, whereas it
was less than 2.5 in the solutions. This suggests that
the Fe/Mg ratio flowing out from a weathered zone is
lower than that in primary biotite under ‘anoxic’
conditions.
Vermiculite or smectite with high Fe was precipitated at the edge of biotite as a secondary mineral
under ‘anoxic’ conditions (Figs. 5 and 6). Although
Fe(II) is available in solution as mentioned above,
little Fe(III) is present in solution. If Fe(III) is present
in solution, it is quickly consumed to form Fe(III)(hydr)oxides as we observed under oxic conditions
(Figs. 2 and 3). Indeed, Fe(III)-(hydr)oxides were not
detected under ‘anoxic’ conditions, even by HRTEM.
The precipitation takes place consuming Fe(II) in
solution to result in the formation of Fe(II)-rich
vermiculite or smectite under ‘anoxic’ conditions,
which is also supported by the presence of Fe(II)-rich
vermiculite and smectite formed by diagenetic and
hydrothermal alteration [44 – 46]. Note that naturally
occurring vermiculite and smectite are usually quite
low in Fe(II)/Fe(III) and low in Fe/Mg [47]. Thus,
Fe(II)-rich vermiculite or smectite is most likely to
occur under ‘anoxic’ conditions although Fe(II) in
vermiculite or smectite was not directly measured for
the present study.
In addition to the formation of vermiculite or
smectite on the edges of biotite grains, layer-by-layer
formation of vermiculite or smectite within biotite
T. Murakami et al. / Earth and Planetary Science Letters 224 (2004) 117–129
grains also occurred (Fig. 8). Murakami et al. [23]
pointed out that layer-by-layer formation of vermiculite does not occur in the early stage of Fe-rich biotite
dissolution/weathering under oxic conditions, because
dissolved Fe(II) quickly forms Fe(III)-(hydr)oxides
and Mg in solution is not available enough to form
the interlayer cations of vermiculite. In contrast to
oxic dissolution/weathering, Fe(II), in addition to Mg,
is available in solution (Table 1) to form vermiculite
or smectite within biotite grains during ‘anoxic’
dissolution/weathering. Note that the starting, fresh
biotite used by Murakami et al. [23] was the same as
ours. Kozai et al. [48] exchanged interlayer cations of
montmorillonite by Fe(II) under ‘anoxic’ conditions,
and confirmed the presence of Fe(II) in the interlayers
by Mössbauer spectrometry. This suggests that Fe(II)
behaves like Mg for the formation of vermiculitic or
smectitic interlayers, i.e., Fe(II) is accommodated in
the interlayers of vermiculite or smectite under ‘anoxic’ conditions, as is Mg.
4.2. Weathering processes under anoxic conditions
and implication for Fe behavior during pre-2.2 Ga
weathering
Because weathering conditions such as temperature, pH, and solution chemistry before 2.2 Ga are not
known, it is not possible to understand the whole
aspect of weathering before 2.2 Ga by using the
present results. However, it is useful to apply the
present results to weathering before 2.2 Ga for a better
understanding of atmospheric evolution, implying
weathering processes and Fe behavior.
The edges of biotite grains were covered with a
thin film (0.2 – 0.5 Am) of vermiculite or smectite
under ‘anoxic’ conditions. This occurs without significant breakdown of biotite except for dissolution of a
small part of biotite (Fig. 5B). At this stage, Fe(II) is
released into water during biotite weathering under
‘anoxic’ conditions, and part of the Fe(II) is used to
form Fe(II)-rich vermiculite or smectite. Rye and
Holland [27] predicted that vermiculite or smectite
decomposes into kaolinite, releasing additional Fe(II)
to the water, and then, kaolinite further breaks down
to form gibbsite, as ‘anoxic’ weathering proceeds.
Such ‘anoxic’ weathering processes are quite similar
to those of modern weathering of biotite described by
Velde [49] except for the Fe behavior. Although the
127
present ‘anoxic’ experiments were carried out at 100
jC, Fe(II)-rich vermiculite or smectite was certainly
formed. Fe(II)-rich vermiculite or smectite finally
decomposes as weathering proceeds [27,49]. The
‘anoxic’ weathering processes mentioned above adequately explain the loss of Fe from pre-2.2 Ga
paleosols, which is consistent with the low-PO2 model
for the pre-2.2 Ga atmosphere.
The Fe/Mg molar ratio in the secondary vermiculite or smectite, more than 7, was larger than that in
the fresh biotite, about 5, which is consistent with the
smaller Fe/Mg molar ratio in solution, less than 2.5,
than that in the fresh biotite. Chlorite, a post-weathering product formed by diagenesis or metamorphism,
is a major Fe-bearing mineral in pre-2.2 Ga paleosols.
If vermiculite or smectite is a product of pre-2.2 Ga
weathering, chlorite should be formed consuming
vermiculite or smectite and primary Fe-bearing minerals such as biotite. Consequently, the chlorite is
characterized by a higher Fe/Mg molar ratio than
those of the primary Fe-bearing minerals because of
the higher Fe/Mg molar ratio of vermiculite or smectite. In addition, the abundance ratio of vermiculite or
smectite to the primary Fe-bearing minerals increases
toward the top of weathering profiles or increases with
increase in weathering until vermiculite or smectite
finally disappears from the weathering profiles. Consequently, the Fe/Mg molar ratio of chlorite increases
toward the top of weathering profiles in which vermiculite or smectite have reacted to form chlorite. Our
prediction of the increase in Fe/Mg molar ratio is
confirmed by recent studies that have revealed that the
Fe/Mg molar ratios of chlorites increase toward the
top of paleosols in Hekpoort, South Africa [27], and
Cooper Lake, Canada [50]. A decrease in Fe/Mg
molar ratio of chlorite is reported for the Lauzon
Bay paleosol, Canada [51]. However, it is most likely
that Mg was added in later events at Lauzon Bay (e.g.,
[3]), which resulted in the decrease in Fe/Mg molar
ratio. Our results indicate that Fe(II)-rich vermiculite
or smectite was formed during ‘anoxic’ weathering
and was a precursor to chlorite in paleosols.
Another important aspect of the secondary mineralization is that Fe(II)-rich vermiculite or smectite was
formed under high PCO2 of 1 atm. Rye et al. [26] and
Rye and Holland [27] suggested that siderite is thermodynamically stable at more than about 10 2 atm of
PCO2. However, our ‘anoxic’ experiments formed
128
T. Murakami et al. / Earth and Planetary Science Letters 224 (2004) 117–129
Fe(II)-rich vermiculite or smectite at high PCO2, 1 atm,
and low pH, 4.6, and produced no other Fe-bearing
minerals. Because no thermodynamic database for
vermiculite, or Fe(II)-rich vermiculite or smectite are
available, we cannot calculate stability relationships
between siderite and Fe(II)-rich vermiculite or smectite. It may be possible that Fe(II)-rich vermiculite or
smectite was formed but not siderite during ’anoxic’
weathering because of kinetics. Although it is evident
that Fe(II)-rich vermiculite or smectite was formed at
1 atm of PCO2, dissolution experiments with various
combinations of PCO2 and PO2 should be carried out to
examine the effects of PCO2 and PO2 on the formation
of secondary weathering products.
Acknowledgements
The authors are grateful to G. Kamei at the Japan
Nuclear Cycle Development Institute for the use of
the glove boxes and to B. Sreenivas for discussion.
We thank A. Monkawa for the XANES data, and T.
Tachikawa and T. Takeshige for technical assistance.
The reviews by D.R. Peacor and M.-P. Turpault
improved the original version of the manuscript. The
electron microscopy was performed in the Electron
Microbeam Analysis Facility for Mineralogy at the
Department of Earth and Planetary Science, the
University of Tokyo. This work was supported by
the Science Grant of the Ministry of Education,
Science and Culture. [EB]
References
[1] A.A. Pavlov, J.F. Kasting, L.L. Brown, K.A. Rages, R.
Freedman, Greenhouse warming by CH4 in the atmosphere of
early Earth, J. Geophys. Res. 105 (2000) 11981 – 11990.
[2] H.D. Holland, The chemical evolution of the atmosphere and
oceans, Princeton Series in Geochemistry, Princeton Univ.
Press, Princeton, NJ, 1984, 582 pp.
[3] R. Rye, H.D. Holland, Paleosols and the evolution of atmospheric oxygen, Am. J. Sci. 298 (1998) 621 – 672.
[4] J.W. Schopf, A.B. Kudryavtsev, D.G. Agresti, T.J. Wdowiak,
A.D. Czaja, Laser-Raman imagery of Earth’s earliest fossils,
Nature 416 (2002) 73 – 76.
[5] M.D. Brasier, O.R. Green, A.P. Jephcoat, A.K. Kleppe, M.J.
Van Kranendonk, J.F. Lindsay, A. Steele, N.V. Grassineau,
Questioning the evidence for Earth’s oldest fossils, Nature
416 (2002) 76 – 81.
[6] P.E. Cloud, Atmospheric and hydrospheric evolution on the
primitive Earth, Science 160 (1968) 729 – 736.
[7] J.C.G. Walker, Evolution of the Atmosphere, Macmillan Publishing, New York, 1977, 318 pp.
[8] J.F. Kasting, Earth’s early atmosphere, Science 259 (1993)
920 – 926.
[9] E. Dimroth, M.M. Kimberley, Precambrian atmospheric oxygen: evidence in the sedimentary distributions of carbon, sulfur,
uranium, and iron, Can. J. Earth Sci. 13 (1976) 1161 – 1185.
[10] H. Ohmoto, Evidence in pre-2.2 Ga paleosols for the early
evolution of atmospheric oxygen and terrestrial biota, Geology 24 (1996) 1135 – 1138.
[11] H.D. Holland, R. Rye, Evidence in pre-2.2 Ga paleosols for
the early evolution of atmospheric oxygen and terrestrial biota:
comment and reply—comment, Geology 25 (1997) 857 – 858.
[12] H. Ohmoto, Evidence in pre-2.2 Ga paleosols for the early
evolution of atmospheric oxygen and terrestrial biota: comment and reply—reply, Geology 25 (1997) 858 – 859.
[13] H.W. Nesbitt, G.M. Young, Formation and diagenesis of
weathering profiles, J. Geol. 97 (1989) 129 – 147.
[14] H.W. Nesbitt, Diagenesis and metasomatism of weathering profiles with emphasis on Precambrian paleosols, in: I.P. Martini,
W. Chesworth (Eds.), Weathering, Soils and Paleosols, Elsevier, New York, 1992, pp. 127 – 152.
[15] R.H. Rainbird, H.W. Nesbitt, J.A. Donaldson, Formation and
diagenesis of a sub-Huronian saprolith: comparison with a
modern weathering profile, J. Geol. 98 (1990) 801 – 822.
[16] A.W. Macfarlane, H.D. Holland, The timing of alkali metasomatism in paleosols, Can. Miner. 29 (1991) 1043 – 1050.
[17] C.M. Fedo, H.W. Nesbitt, G.M. Young, Unraveling the effects
of potassium metasomatism in sedimentary rocks and paleosols with implications for paleoweathering conditions and
provenance, Geology 23 (1995) 921 – 924.
[18] C.M. Fedo, G.M. Young, H.W. Nesbitt, J.M. Hanchar, Potassic and sodic metasomatism in the Southern Province of the
Canadian Shield: evidence from the Paleoproterozoic Serpent
Formation, Huronian Supergroup, Canada, Precambrian Res.
84 (1997) 17 – 36.
[19] H.W. Nesbitt, G.M. Young, Prediction of some weathering
trends of plutonic and volcanic rocks based on thermodynamic
and kinetic considerations, Geochim. Cosmochim. Acta 48
(1984) 1523 – 1534.
[20] E.K. Berner, R.A. Berner, Global Environment: Water, Air,
and Geochemical Cycles, Prentice Hall, Upper Saddle River,
NJ, 1996, 376 pp.
[21] H. Dong, D.R. Peacor, S.F. Murphy, TEM study of progressive alteration of igneous biotite to kaolinite throughout a
weathered soil profile, Geochim. Cosmochim. Acta 62
(1998) 1881 – 1887.
[22] D.S. Fanning, V.Z. Keramidas, M.A. El-Desoky, Micas,
in: J.B. Dixon, S.B. Weed (Eds.), Minerals in Soil Environments, Soil Science Society of America, Madison, WI,
1989, pp. 551 – 634.
[23] T. Murakami, S. Utsunomiya, T. Yokoyama, T. Kasama, Biotite dissolution processes and mechanisms in the laboratory
and in nature: early stage weathering environment and vermiculitization, Am. Mineral. 88 (2003) 377 – 386.
T. Murakami et al. / Earth and Planetary Science Letters 224 (2004) 117–129
[24] K.L. Nagy, Dissolution and precipitation kinetics of sheet
silicates, in: A.F. White, S.L. Brantley (Eds.), Chemical
Weathering Rates of Silicate Minerals, Reviews in Mineralogy, vol. 31, Mineralogical Society of America, Washington,
DC, 1995, pp. 173 – 233.
[25] J.B. Acker, O.P. Bricker, The influence of pH on biotite dissolution and alteration kinetics at low temperature, Geochim.
Cosmochim. Acta 56 (1992) 3073 – 3092.
[26] R. Rye, P.H. Kuo, H.D. Holland, Atmospheric carbon dioxide
concentrations before 2.2 billion years ago, Nature 378 (1995)
603 – 605.
[27] R. Rye, H.D. Holland, Geology and geochemistry of paleosols
developed on the Hekpoort basalt, Pretoria Group, South
Africa, Am. J. Sci. 300 (2000) 85 – 141.
[28] T. Murakami, S. Utsunomiya, Y. Imazu, N. Prasad, Direct
evidence of late Archean to early Proterozoic anoxic atmosphere from a product of 2.5 Ga old weathering, Earth Planet.
Sci. Lett. 184 (2001) 523 – 528.
[29] Y. Takahashi, Geology of the granitic rocks in the Tsukuba
area (in Japanese), Geol. Mag. (Japan) 88 (1982) 177 – 184.
[30] K. Tsukimura, Hydrothermal solution equipment capable of
easy extraction of liquid up to 200 jC, Min. J. 17 (1995)
322 – 329.
[31] S. Utsunomiya, T. Murakami, H. Kadohara, K. Tsukimura,
The effect of partial pressure of carbon dioxide on anorthite
dissolution, Min. J. 21 (1999) 1 – 8.
[32] T. Murakami, T. Kogure, H. Kadohara, T. Ohnuki, Formation
of secondary minerals and its effect on anorthite dissolution,
Am. Mineral. 83 (1998) 1209 – 1219.
[33] T.J. Wolery, EQ3NR, a computer program for geochemical
aqueous speciation-solubility calculations: theoretical manual,
user’s guide, and related documentation (version 7.0) (UCRLMA-110662 PT III), Lawrence Livermore National Laboratory, University of California, Lawrence Livermore, CA, 1992,
246 pp.
[34] W. Stumm, J.J. Morgan, Aquatic Chemistry, 2nd ed., Wiley,
New York, 1981, 780 pp.
[35] S.P. Franklin, A. Hajash, T.A. Dewers, T.T. Tieh, The role of
carboxylic acids in albite and quartz dissolution: an experimental study under diagenetic conditions, Geochim. Cosmochim. Acta 58 (1994) 4259 – 4279.
[36] S.A. Welch, W.J. Ullman, The effect of organic acids on plagioclase dissolution rates and stoichiometry, Geochim. Cosmochim. Acta 57 (1993) 2725 – 2736.
[37] R. Kilaas, Optical and near-optical filters in high-resolution
electron microscopy, J. Microsc. 190 (1998) 45 – 51.
129
[38] J.F. Banfield, T. Murakami, Atomic-resolution transmission
electron microscope evidence for the mechanism by which
chlorite weathers to 1:1 semi-regular chlorite – vermiculite,
Am. Mineral. 83 (1998) 348 – 357.
[39] B.E. Kalinowski, P. Schweda, Kinetics of muscovite, phlogopite, and biotite dissolution and alteration at pH 1 – 4, room
temperature, Geochim. Cosmochim. Acta 60 (1996) 367 – 385.
[40] M. Malmström, S. Banwart, Biotite dissolution at 25 jC: the
pH dependence of dissolution rate and stoichiometry, Geochim. Cosmochim. Acta 61 (1997) 2779 – 2799.
[41] T. Murakami, T. Sato, A. Inoue, HRTEM evidence for the
process and mechanism of saponite-to-chlorite conversion
through corrensite, Am. Mineral. 84 (1999) 1080 – 1087.
[42] T. Kogure, T. Murakami, Direct identification of biotite/vermiculite layers in hydrobiotite using high-resolution TEM,
Min. J. 18 (1996) 131 – 137.
[43] W. Stumm, G.F. Lee, Oxygenation of ferrous iron, Ind. Eng.
Chem. 53 (1961) 143 – 146.
[44] D. Craw, Ferrous-iron-bearing vermiculite – smectite series
formed during alteration of chlorite to kaolinite, Otago schist,
New Zealand, Clay Miner. 19 (1984) 509 – 520.
[45] R.H. April, D.M. Keller, Saponite and vermiculite in amygdales of the Granby Basaltic Tuff, Connecticut Valley, Clays
Clay Miner. 40 (1992) 22 – 31.
[46] D. Craw, D.W. Smith, J.H. Youngson, Formation of authigenic
Fe2 +-bearing smectite – vermiculite during terrestrial diagenesis, southern New Zealand, N.Z. J. Geol. Geophys. 38 (1995)
151 – 158.
[47] A.C.D. Newman, G. Brown, The chemical constitution of
clays, in: A.C.D. Newman (Ed.), Chemistry of Clays and
Clay Minerals, Mineralogical Society, London, UK, 1987,
pp. 1 – 128.
[48] N. Kozai, Y. Adachi, S. Kawamura, K. Inada, T. Kozai, S. Sato,
H. Ohashi, T. Ohnuki, T. Banba, Characterization of Fe-montmorillonite: a simulant of buffer materials accommodating
overpack corrosion product, J. Nucl. Sci. Technol. 38 (2001)
1141 – 1143.
[49] B. Velde, Introduction to Clay Minerals, Chapman & Hall,
London, UK, 1992, 198 pp.
[50] S. Utsunomiya, T. Murakami, M. Nakada, T. Kasama, Iron
oxidation state of a 2.45-Byr-old paleosol developed on mafic
volcanics, Geochim. Cosmochim. Acta 67 (2003) 213 – 221.
[51] S.J. Sutton, J.B. Maynard, Multiple alteration events in the
history of a sub-Huronian regolith at Lauzon Bay, Ontario,
Can. J. Earth Sci. 29 (1992) 432 – 445.