Earth and Planetary Science Letters 254 (2007) 358 – 376
www.elsevier.com/locate/epsl
Identification of chemical sedimentary protoliths using iron isotopes
in the N 3750 Ma Nuvvuagittuq supracrustal belt, Canada
Nicolas Dauphas a,b,⁎, Nicole L. Cates c , Stephen J. Mojzsis c , Vincent Busigny a,b
a
c
Origins Laboratory, Department of the Geophysical Sciences and Enrico Fermi Institute, The University of Chicago,
5734 South Ellis Avenue, Chicago IL 60637, USA
b
Department of Geology, The Field Museum, 1400 South Lake Shore Drive, Chicago IL 60605, USA
Department of Geological Sciences, Center for Astrobiology, University of Colorado, Boulder, Colorado 80309-0399, USA
Received 15 June 2006; received in revised form 31 October 2006; accepted 25 November 2006
Available online 16 January 2007
Editor: H. Elderfield
Abstract
An Eoarchean supracrustal belt dated at ca. 3750 Ma was recently identified in the Innuksuac Complex, northern Québec
(Canada). Rocks from the Nuvvuagittuq locality include mafic and ultramafic amphibolites, quartz–biotite and pelitic schists,
orthogneisses, and banded quartz–magnetite–amphibole/pyroxene rocks of probable chemical sedimentary origin. The purported
metasediments are enriched in the heavy isotopes of Fe by approximately 0.3‰/amu relative to IRMM-014. They also have high
Fe/Ti ratios, up to 100× that of associated amphibolite units. These signatures demonstrate that quartz–magnetite–amphibole/
pyroxene rocks from Nuvvuagittuq are chemical sediments (e.g. banded iron-formations, BIFs) formed by precipitation of
dissolved ferrous iron in a marine setting. All units were metamorphosed to upper amphibolite facies, which partly homogenized Fe
isotopes. Variable Fe isotope compositions of bulk quartz–magnetite rocks are interpreted to reflect binary mixing between primary
oxides and carbonates. Mixing relationships with major element chemistry (Ca/Fe, Mg/Fe, and Mn/Fe) are used to estimate the Fe
isotope composition of the primary Fe-oxide phase (0.3 to 0.4‰/amu) and the chemistry of the carbonate (siderite and ankerite).
Iron isotopes can thus be used to constrain the primary mineralogy of Fe-rich chemical sedimentary precipitates before
metamorphism. The possible presence of siderite in the primary mineral assemblage supports deposition under high PCO2. We
developed an isotope distillation model that includes two possible abiotic oxidation paths, homogeneous and heterogeneous. The
isotopic composition of Fe in the precursor phase of magnetite in BIFs can be explained by partial oxidation through oxygenic or
anoxygenic photosynthesis of Fe from a hydrothermal source.
© 2006 Elsevier B.V. All rights reserved.
Keywords: Archean; Sediment; BIF; Iron; Isotopes; Metamorphism
1. Introduction
⁎ Corresponding author. Origins Laboratory, Department of the
Geophysical Sciences and Enrico Fermi Institute, The University of
Chicago, 5734 South Ellis Avenue, Chicago IL 60637, USA. Tel.: +1
773 7022930.
E-mail address: [email protected] (N. Dauphas).
0012-821X/$ - see front matter © 2006 Elsevier B.V. All rights reserved.
doi:10.1016/j.epsl.2006.11.042
No terrestrial rocks formed before ca. 4.03 Ga are
known to have survived subsequent crustal recycling
processes [1]. The only direct witnesses of the earliest
times are tiny zircon grains accumulated in younger
detrital sediments [2]. In this respect, the large exposures
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
of Eoarchean (3.6–3.8 Ga) supracrustal rocks preserved in the North Atlantic province of the North
American craton (Isua supracrustal belt, ISB; and Akilia
association, AA of southern West Greenland; Nulliak
assemblage of northeast Labrador, Canada) are important [3–9]. The protoliths of some of these formed as
marine sediments and convey crucial information on the
early history of our planet. Was the Earth habitable by
3.8 Ga? Had life already emerged? If so, what was the
nature of the biosphere? How was the chemistry of the
atmosphere–ocean system different? What was the flux
of extraterrestrial matter to the Earth? A principal issue
with these samples is that until recently, they were
thought to be exclusive to West Greenland and parts of
Labrador, and it was unclear whether they offered an
unbiased perspective of the Earth at that time.
359
In 2002, David et al. reported the discovery of a new
Eoarchean volcano–sedimentary (supracrustal) sequence in northern Québec along the eastern shore of
Hudson Bay (Ungava Peninsula, Canada) [10]. The
Nuvvuagittuq supracrustal belt (NSB) is located in
the Inukjuak lithotectonic domain of the Northeastern
Superior Province (Minto block, Canadian Shield,
Fig. 1). Other supracrustals in the area comprise the
Innuksuac Complex. David and coworkers [10,11]
determined the age of a felsic unit from the NSB as
3825 ± 16 Ma based on U–Pb TIMS geochronology of
zircons, after rejection of an outlier (3805 ± 77 Ma if it is
included [12]). Zircon morphologies and trace element
geochemistry were used to suggest the host rock was a
volcanic tuff and that the zircons were not xenocrystic
[13]. More recently, Cates and Mojzsis [12] reported an
Fig. 1. Map of the Ungava Peninsula, Northern Québec. The Nuvvuagittuq supracrustal belt (NSB, marked with a white star) is located in the
Inukjuak lithotectonic domain of the Northeastern Superior Province (Canadian Shield). The inset map also shows (marked with black stars) other
exposures of Eoarchean supracrustal rocks in SW Greenland (Isua supracrustal belt ISB and Akilia Association AA) and NE Labrador (Nulliak
assemblage). Map modified from [73,74].
360
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
age of 3758 + 51/− 47 Ma for the same unit based on ion
microprobe U–Pb zircon geochronology. For comparison, in West Greenland the ISB is also 3.7–3.8 Ga [3,4]
and some components of the AA are older than 3.83 Ga
[4–7]. The mineral phases observed in the NSB indicate
that it was metamorphosed to upper-amphibolite/lowergranulite facies [12,13], similar to or higher than what is
documented for the ISB (N 550 °C 5 kb) [14].
Archean sedimentary rocks show variable Fe isotope
compositions at mineral and bulk sample scales (from
− 1.75 to +0.75‰/amu relative to the reference material
IRMM-014) [15–19]. This is in contrast with values
measured in modern igneous rocks formed in different
tectonic settings, which cluster narrowly around
+0.04‰/amu [20]. The measured variations may be
due to isotope fractionation during chemical precipitation in the oceans, biological reworking of Fe after
sediment deposition, and abiotic reactions during
diagenesis. The banded quartz–magnetite and quartz–
amphibole/pyroxene rocks (purported BIFs) from the
ca. 3.8 Ga ISB and AA supracrustal units in West
Greenland are enriched in the heavy isotopes of Fe
relative to surrounding igneous rocks (from + 0.1 to
+0.5‰/amu) [17]. This feature may be the imprint of
partial oxidation in the water column, which is known
experimentally to enrich the oxidized species in the
heavy isotopes of Fe [21–24]. Units of likely chemical
sedimentary origin have been identified in the NSB
[10–13]. The presence of such rocks offers us the
exciting possibility to probe Earth's surface chemistry at
3.7–3.8 Ga from a different perspective than that
provided by the West Greenland samples. The goals of
this study are to: (i) establish the origin of the banded
quartz ± magnetite ± amphibole/pyroxene rocks found in
the NSB; (ii) search for possible petrogenetic relationships with supracrustal rocks from the coeval West
Greenland units; (iii) evaluate the effect of metamorphism on Fe isotopes; and (iv) place further constrains
on the mechanisms responsible for Fe precipitation in
the Archean oceans.
2. Sample selection and description
Sample selection was guided by detailed (1:50 scale)
mapping of a 130 × 40 m shoreline exposure in the NSB
[12]. The documented supracrustal sequence includes
mafic and ultramafic amphibolite units, polymict
(conglomeratic) quartz–biotite schists, pelitic to psammitic schists, orthogneisses, and banded quartz–magnetite + ferruginous quartz–amphibole/pyroxene rocks.
The supracrustal lithologies have been multiply deformed in several metamorphic events, and are intruded
by tonalite–trondhjemite–granodiorite (TTG) composition orthogneisses, including 3.76 Ga trondhjemitic
gneisses, some with clear cross-cutting relationships to
the paragneisses (see [12] for detailed descriptions of
these units).
2.1. Mafic and ultramafic units (IN05013, 19, 45, 47)
Mafic amphibolites have tholeiitic compositions with
flat chondrite-normalized rare earth element (REE)
patterns [12] and radiogenic initial Nd isotope compositions indicative of a depleted mantle source [25]. Major
mineral phases are amphibole, plagioclase, garnet, and
quartz. Textures range from massive to banded.
Ultramafic lenses of basaltic komatiite composition are
intercalated with the amphibolite units (IN05047).
2.2. Quartz–biotite schists (polymict conglomerate?)
(IN05004, 05, 20, 37)
Rocks that contain highly-strained polycrystalline
quartz embedded in a biotite–clinozoisite matrix have
trace element geochemistry consistent with a detrital
origin. Due to the high strain, protolith assignment to
this unit is uncertain.
2.3. Pelitic schists (IN05050, 53)
A pelitic–psammitic quartzite unit preserves fine
laminations a few millimetres to a few centimetres thick.
Pelitic layers are dominantly biotite, with minor
clinozoisite, carbonate and disseminated quartz, whereas quartzite layers are dominantly fine-grained quartz in
aggregates of grains with (annealed) triple-junctions
+ minor biotite, clinozoisite and carbonate.
2.4. Quartz–magnetite and ferruginous quartz–amphibole/
pyroxene units (IN05007, 08, 09, 10, 48)
The best preserved examples that escaped silica
mobility and strain retain bands of magnetite alternating
with bands of quartz and amphibole which resemble
banded iron-formation. In more metamorphosed samples (IN05048), pyroxenes rather than amphiboles are
present. REE patterns, including positive Eu anomalies,
are consistent with an origin from hydrothermal sediments for this unit [12,26]. Because sample IN05007 is
the best-preserved quartz–magnetite rock analyzed in
this study, it was selected for detailed petrographic and
Fe isotope characterizations (Figs. 2 and 3). The handspecimen (approximately 2.8 × 3.7 × 1.8 cm in size),
consists of parallel, partly anastomosing, bands of
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
361
Fig. 2. Petrography and chemistry of silicates and carbonates in banded quartz–magnetite sample IN05007. Top panel: SEM photomicrograph
showing textural relationships between ankerite (ank), quartz (qtz), actinolite (act), and cummingtonite (cum). Bottom panel: chemical compositions of
carbonates and amphiboles plotted in ternary diagrams. The grey areas are the convex hulls enclosing values measured in BIFs from the ISB [38,44].
The metamorphic reaction responsible for the observed paragenesis is, primary carbonates + qtz + H2O → cum + act + ank + cal + CO2 (Eq. (4)) [41].
magnetite (1–4 mm in thickness) alternating with bands
of quartz and amphibole. Although the rock has been
strongly deformed, these layers may represent relict
sedimentary banding. Carbonates are also present and
crosscut in some places the layering. Image analysis of a
0.5 mm2 area reveals that these carbonates consist of
20% ankerite (Ca0.495Mg0.239Fe0.191Mn0.075) and 80%
calcite (Ca0.910Mg0.015Fe0.030Mn0.045). The width of the
solvus between these two phases can be used as a
thermometer [27] and generates a temperature of 380 °C,
much lower than inferred peak metamorphic conditions
of N 550 °C. This can be explained by equilibration on
the retrograde path or mobilization by fluids. The
amphiboles consist predominantly (N99%) of cummingtonite (Na2O 0.64, MgO 11.28, Al2O3 0.29, SiO2 53.77,
CaO 0.96, MnO 2.30, FeO 30.78 wt.%) with minor
actinolite (Na2O 0.55, MgO 12.02, Al2O3 0.99, SiO2
54.50, CaO 11.53, MnO 0.75, FeO 19.61 wt.%).
Disseminated pyrite is also present in IN05007. Overall,
the petrography and chemistry of this sample (also see
[12]) is very similar to what has been described for
banded iron-formations from the ISB [28].
3. Analytical methods
Large specimens were cut or broken to offer fresh
surfaces and minimize the influence of alteration. After
cleaning with acetone, the bulk samples (typically ∼ 1 g)
were reduced to powder in an agate mortar. Sampling of
individual magnetite bands in IN05007 was done using a
micromill apparatus (MicroMill, New Wave Research)
equipped with a tungsten carbide mill bit (Brasseler,
scriber H1621.11). The section was extensively characterized before microsampling using X-ray and BSE
imaging on a JEOL JSM-5800LV SEM (Figs. 2 and 3).
At each of the 25 sample points, the following sequence
was followed. A drop of MilliQ water was deposited at
the surface of the polished sample. The computercontrolled bit was moved within the water droplet to mill
an array of holes (∼ 18), each approximately 100 μm in
362
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
Fig. 3. Chemical and isotopic mapping of IN05007. Left panel: False color X-ray mosaic (assembled from ∼ 4 × 1000 frames), with Ca in cyan (C),
S in magenta (M), Fe in yellow (Y), and Si in black (K). The whole section is 10 cm2. Magnetite is concentrated in parallel layers (yellow). The vein
(cyan) that crosscuts the layering in the top-left corner consists of calcite and ankerite. Pyrite (magenta) is disseminated in magnetite layers.
Cummingtonite (light dark yellow–grey) and quartz (dark grey) are the most abundant silicate phases. Actinolite is present as tiny domains (green–
grey) associated with cummingtonite. Right panel: FFe values (‰/amu relative to IRMM-014) in individual bands of magnetite (the mill locations are
shown on the schematic map) see Table 1. Distance is taken from the bottom of the section. The grey band is the 95% confidence interval for a bulk
sample measurement. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)
diameter and 500 μm in depth (assuming a cylindrical
geometry, this corresponds to a total of ∼ 270 μg Fe
milled in magnetite). The sample slurry containing 100–
500 μg of Fe was transferred in a Teflon beaker using a
micropipette. It was then digested following the same
protocol as for bulk samples described hereafter.
The protocol for sample dissolution, purification, and
isotope analysis of Fe has been described elsewhere
[29,30], and is only briefly reviewed here. Powder
aliquots weighing less than 10 mg were dissolved in
Teflon beakers in 1 mL HF, 0.5 mL HNO3, and a few
drops of HClO4 at ∼ 100 °C for 5–10 h (unless otherwise
noted, all acids listed are concentrated solutions).
The solutions were then evaporated to dryness on hot
plates under heat lamps and the salts were dissolved in
0.25 mL HNO3, 0.75 mL HCl, and a few drops of
HClO4, evaporated, dissolved in 0.5 mL HNO3, 1 mL
HCl, and a few drops of HClO4, evaporated again, and
finally dissolved in 0.5 mL HCl 6 M. At this stage, the
samples were ready for loading on anion exchange resin
(typically 50–500 μg Fe).
Iron was separated from matrix elements and direct
isobars (54Cr and 58Ni can interfere on 54Fe and 58Fe,
respectively) by anion exchange chromatography. After
loading of sample solutions onto Bio-Rad Poly-Prep
columns filled with 1 mL of AG1-X8 200–400 mesh
resin, matrix elements were eluted with 8 mL of HCl 6 M,
and Fe was recovered using 9 mL of HCl 0.4 M. This
sequence was repeated. The yield is close to 100% and no
isotope fractionation is introduced by the chemistry [29].
The blank of the whole procedure is b40 ng Fe (b0.1% of
the amount of Fe loaded on the column).
Iron isotopic compositions were measured on an
Isoprobe (Micromass) multi-collector inductively coupled plasma mass spectrometer (MC-ICPMS) installed at
the Field Museum. The samples (2 ppm Fe in HNO3
0.45 M) were introduced into the mass spectrometer
using a Cetac Aridus desolvating system. Instrumental
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
mass fractionation (∼50‰/amu) was corrected for using
standard bracketing and time interpolation with the
reference material IRMM-014 (5.845% 54Fe, 91.754%
56
Fe, 2.1191% 57Fe, and 0.2819% 58Fe) [31]. Isotopic
compositions are calculated using the δ notation, averaged
over 2–8 replicate analyses of the same solution,
di=j Fe ¼ ½ði Fe =j FeÞSample =ði Fe=j FeÞIRMM014 −1103 :
ð1Þ
Because different laboratories measure different isotope ratios, we prefer to use the FFe notation [30], which is
expressed in permil per atomic mass unit (‰/amu)
deviation relative to the composition of the standard,
FFe ¼ di=j Fe=ði−jÞ:
ð2Þ
If isotope ratios do not depart significantly from
linear mass fractionation over the range of variations
investigated, then the FFe notation should give identical
values for any pair of isotopes considered (in this
study 54Fe and 56Fe). Error bars are calculated based on
multiple standard analyses and are reported as 95%
confidence intervals.
At the beginning of each session of analyses, several
geostandards and igneous rocks were measured to test
the stability of the instrumental mass bias and confirm
that there had been no problem during sample digestion
and chromatography (Table 1). The samples T4D2#1
(basalt from Loihi, Hawaii), Payun (pseudomorph of
hematite after magnetite from Payun Matra volcano,
Argentina), and Naoned#1 (Eclogite from St-Philbertde-Bouaine, France) have Fe isotope compositions
(0.041 ± 0.041, − 0.027 ± 0.032, and 0.045 ± 0.088‰/
amu, respectively) identical to earlier results reported
for these samples (0.072 ± 0.051‰/amu for T4D2#1
[17] and 0.004 ± 0.030‰/amu for Payun [32]) or for
igneous rocks (0.039 ± 0.008‰/amu [20,30]). Eight
replicate analyses (including dissolution and chromatography) of the geostandard IF-G (a 3.7–3.8 Ga BIF
from the ISB [33]), give a weighted average of 0.288 ±
0.019‰/amu, indistinguishable from the recommended
value of 0.316 ± 0.010‰/amu [30]. The external reproducibility of these 8 replicate measurements is 0.09‰/
amu (2σ), which is in good agreement with internal
precisions obtained for individual analyses (from 0.04 to
0.10‰/amu).
4. Results
Mafic and ultramafic samples from the Nuvvuagittuq
supracrustal belt (n = 4) have Fe isotopic compositions
363
very close to IRMM-014 (1/σ2 weighted average
− 0.030 ± 0.018‰/amu or − 0.006 ± 0.022‰/amu if
IN05045, a banded amphibolite with a low Fe isotopic
composition, is excluded, Table 1). This is similar to
values reported for ≥ 3.83 Ga amphibolite rocks [7]
from the Akilia association (+ 0.011 ± 0.010‰/amu)
[17] and carbonaceous chondrites (− 0.012 ± 0.010‰/
amu, Table 1) [30]. Beard et al. [20] showed that modern
(b0.1 Ga) igneous rocks sampled in a variety of tectonic
settings have almost homogeneous Fe isotopic compositions, around +0.04‰/amu. This is slightly heavier
than the values found for Eoarchean samples. Variations
of − 0.15 to + 0.20‰/amu have been documented in
carbonatites (Dauphas et al., in prep) and granitoids
[34]. The values measured in mafic and ultramafic rocks
from NSB fall well within this range. There are few
measurements published in the literature with which
3.8 Ga mafic and ultramafic rocks from the NSB can be
compared. Poitrasson et al. [35] reported a value of
+ 0.019 ± 0.008‰/amu for a 3.5 Ga komatiite from
Barberton (South Africa), which is close to the values
reported here.
All quartz–biotite schists of possible conglomeratic
(detrital) origin (n = 4) have indistinguishable Fe isotopic
compositions, averaging − 0.026 ± 0.025‰/amu. This is
identical to the signature of igneous rocks from the NSB.
The pelitic schists have light Fe isotopic compositions
(n = 2, −0.118 ± 0.070 and –0.240 ± 0.070‰/amu) relative to IRMM-014 and igneous rocks. Similar signatures
have been documented for other Archean pelitic rocks.
Yamaguchi et al. [16] measured the Fe isotope compositions of well-preserved shales and greywackes from drill
cores ranging in age from 2.20 to 3.25 Ga. Samples with
small amounts of organic carbon, carbonates, and
sulfides, have Fe isotopic compositions close to the
signature measured in igneous rocks and suspended river
loads. In contrast, most of the samples that contain
significant siderite, organic carbon, and magnetite have
negative isotopic compositions, down to −1‰/amu.
Negative FFe values are also found in well-preserved
BIFs from the Kaapvaal craton, especially in pyrite and
Fe-carbonate rich layers (down to −1.20‰/amu [15]).
The banded quartz–magnetite rocks of likely BIF
protolith (n = 5) are, for the majority, enriched in the
heavy isotope of Fe (around + 0.3‰/amu). The only
exception is IN05048, a highly deformed quartz–
pyroxene rock within the finely banded quartz–magnetite unit, with a FFe of + 0.038 ± 0.023‰/amu. These
heavy Fe values are close to those reported for bulk
ferruginous quartzites from supracrustals rocks of the
Akilia association and Isua (from +0.1 to + 0.5‰/amu
[17]). The Mean Square Weighted Deviate (MSWD, also
364
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
Table 1
Chemical and isotopic compositions of geostandards, meteorites, and supracrustal rocks from the NSB
Type
Sample
Description
Fe
(mol.g− 1)
T4D2#1 Basalt (Loihi, HI)
Naoned#1 Eclogite (St-Philbertde-Bouaine, France)
Payun
Hmt, pseudomorph
after mgt (Argentina)
Average Payun
Vigarano CV3 chondrite
Ornans
CO3 chondrite
Lance
CO3 chondrite
BCR-2
Basalt (Columbia
River, OR)
IF-G
BIF geostandard
(Isua, Greenland)
N Québec
igneous
rocks
IN05019
IN05047
IN05013
IN05045
N Québec BIFs
IN05048
IN05007
IN05009
IN05008
IN05010
IN1S1
IN1S2
IN2S1
IN2S2
IN2S3
0.081 ± 0.082
0.089 ± 0.175
Average IN05048
Relatively well
preserved BIF
Average IN05007
Intermediate
between IN05007
and IN05048
Average IN05009
Relatively well
preserved BIF
Relatively well
preserved BIF
Microsampling
IN05007 (0.17 cm)
Microsampling
IN05007 (0.17 cm)
Microsampling
IN05007 (0.17 cm)
Microsampling
IN05007 (0.19 cm)
Microsampling
IN05007 (0.21 cm)
δ57Fe
(‰)
0.059 ± 0.158
0.206 ± 0.152
FFe
(‰/amu)
Fe/
Ti
Ca/
Fe
Mg/
Fe
Mn/
Fe
0.041 ± 0.041
0.045 ± 0.088
− 0.039 ± 0.080 − 0.045 ± 0.167 −0.020 ± 0.040
0.001730
− 0.077 ± 0.110
− 0.052 ± 0.065
0.034 ± 0.110
− 0.094 ± 0.082
0.030 ± 0.157
0.022 ± 0.080
0.469 ± 0.087
Average IF-G
0.006995
Massive amphibolite 0.001628
Ultramatic near the
BIFs (mgt-rich)
Amphibolite within
the BIFs (plg-rich)
Banded amphibolite
25 m from the BIFs
Coarse Qtz–Px rock
δ56Fe
(‰)
0.002505
0.001618
0.001155
0.000901
− 0.140 ± 0.074 −0.039 ± 0.055
− 0.124 ± 0.068 −0.026 ± 0.032
− 0.060 ± 0.587 0.017 ± 0.055
− 0.224 ± 0.212 −0.047 ± 0.041
0.191 ± 0.135 0.015 ± 0.079
0.089 ± 0.167 0.011 ± 0.040 6.133
0.763 ± 0.211
0.734 0.514 0.0160
0.235 ± 0.044
0.648 ± 0.087 0.997 ± 0.177 0.324 ± 0.044
0.496 ± 0.194 0.760 ± 0.332 0.248 ± 0.097
0.608 ± 0.098 0.897 ± 0.154 0.304 ± 0.049
0.577 ± 0.139 0.834 ± 0.268 0.289 ± 0.070
0.622 ± 0.083 0.934 ± 0.125 0.311 ± 0.042
0.592 ± 0.091 0.816 ± 0.097 0.296 ± 0.046
0.379 ± 0.193 0.625 ± 0.256 0.190 ± 0.097
0.576 ± 0.037 0.857 ± 0.057 0.288 ± 0.018 10346
− 0.024 ± 0.06 − 0.078 ± 0.078 −0.012 ± 0.030 15.695
0.040 0.067 0.0008
1.012 1.399 0.0164
− 0.025 ± 0.151 − 0.044 ± 0.243 −0.013 ± 0.076 22.431
0.074 2.230 0.0124
0.007 ± 0.071
− 0.148 ± 0.06
0.029 ± 0.087
0.004 ± 0.036 18.495
− 0.246 ± 0.078 −0.074 ± 0.030 8.496
0.691 1.490 0.0157
0.958 0.527 0.0196
0.004145
0.083 ± 0.06
0.066 ± 0.071
0.076 ± 0.046
0.436 ± 0.151
0.056 ± 0.078
0.081 ± 0.087
0.067 ± 0.058
0.634 ± 0.243
0.042 ± 0.030 720.886 0.358 0.501 0.0346
0.033 ± 0.036
0.038 ± 0.023
0.218 ± 0.076 3300.864 0.021 0.125 0.0126
0.004328
0.507 ± 0.151
0.472 ± 0.107
0.552 ± 0.067
0.742 ± 0.243
0.688 ± 0.172
0.758 ± 0.092
0.254 ± 0.076
0.236 ± 0.053
0.276 ± 0.034 153.159 0.035 0.198 0.0080
0.003960
0.654 ± 0.067
0.603 ± 0.047
0.572 ± 0.139
0.886 ± 0.092
0.822 ± 0.065
0.91 ± 0.229
0.327 ± 0.034
0.302 ± 0.024
0.286 ± 0.070 3153.561 0.058 0.167 0.0123
0.004937
0.749 ± 0.092
1.062 ± 0.145
0.375 ± 0.046 3931.905 0.033 0.135 0.0070
0.492 ± 0.107
0.822 ± 0.222
0.246 ± 0.054
0.556 ± 0.139
0.809 ± 0.232
0.278 ± 0.070
0.436 ± 0.088
0.659 ± 0.182
0.218 ± 0.044
0.588 ± 0.113
0.867 ± 0.160
0.294 ± 0.056
0.616 ± 0.165
0.893 ± 0.264
0.308 ± 0.083
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
365
Table 1 (continued)
Type
Sample
Description
IN2S4
IN05050
Microsampling
IN05007 (0.21 cm)
Microsampling
IN05007 (0.39 cm)
Microsampling
IN05007 (0.44 cm)
Microsampling
IN05007 (0.83 cm)
Microsampling
IN05007 (0.79 cm)
Microsampling
IN05007 (0.77 cm)
Microsampling
IN05007 (1.03 cm)
Microsampling
IN05007 (1.03 cm)
Microsampling
IN05007 (1.00 cm)
Microsampling
IN05007 (0.96 cm)
Microsampling
IN05007 (1.05 cm)
Microsampling
IN05007 (1.26 cm)
Microsampling
IN05007 (1.50 cm)
Microsampling
IN05007 (1.45 cm)
Microsampling
IN05007 (1.41 cm)
Microsampling
IN05007 (1.44 cm)
Microsampling
IN05007 (1.52 cm)
Microsampling
IN05007 (1.94 cm)
Microsampling
IN05007 (2.27 cm)
Microsampling
IN05007 (2.40 cm)
Qtz–bt polymict
conglomerate
Qtz–bt polymict
conglomerate
Qtz–bt polymict
conglomerate
Qtz–bt polymict
conglomerate
Bt–Qtz schist
IN05053
Bt–Qtz schist
IN3S1
IN4S1
IN6S1
IN6S2
IN6S3
IN7S1
IN7S2
IN7S3
IN7S4
IN7S5
IN8S1
IN9S1
IN9S2
IN9S3
IN9S4
IN9S4P
IN10S1
IN11S1
IN12S1
N Québec
conglomerate
IN05004
IN05005
IN05020
IN05037
N Québec
schists
Fe
(mol.g− 1)
δ56Fe
(‰)
δ57Fe
(‰)
FFe
(‰/amu)
0.628 ± 0.159
1.082 ± 0.268
0.314 ± 0.080
0.566 ± 0.088
0.934 ± 0.182
0.283 ± 0.044
0.373 ± 0.107
0.606 ± 0.222
0.187 ± 0.054
0.614 ± 0.113
0.890 ± 0.160
0.307 ± 0.056
0.510 ± 0.173
0.790 ± 0.245
0.255 ± 0.087
0.524 ± 0.173
0.787 ± 0.245
0.262 ± 0.087
0.757 ± 0.159
1.037 ± 0.268
0.379 ± 0.080
0.511 ± 0.091
0.669 ± 0.097
0.256 ± 0.045
0.510 ± 0.091
0.797 ± 0.097
0.255 ± 0.045
0.509 ± 0.193
0.830 ± 0.256
0.255 ± 0.097
0.597 ± 0.193
0.787 ± 0.256
0.299 ± 0.097
0.569 ± 0.193
0.741 ± 0.256
0.285 ± 0.097
0.638 ± 0.091
0.940 ± 0.097
0.319 ± 0.045
0.758 ± 0.125
1.123 ± 0.174
0.379 ± 0.063
0.716 ± 0.125
1.073 ± 0.174
0.358 ± 0.063
0.455 ± 0.135
0.720 ± 0.202
0.228 ± 0.068
0.703 ± 0.125
0.993 ± 0.174
0.352 ± 0.063
0.639 ± 0.091
0.976 ± 0.097
0.319 ± 0.045
0.570 ± 0.179
0.883 ± 0.291
0.285 ± 0.090
0.591 ± 0.165
0.917 ± 0.264
0.296 ± 0.083
Fe/
Ti
Ca/
Fe
Mg/
Fe
Mn/
Fe
0.001750
− 0.104 ± 0.092 − 0.138 ± 0.145 − 0.052 ± 0.046 28.008
0.057 0.836 0.0040
0.002109
− 0.013 ± 0.139
0.008 ± 0.229 − 0.007 ± 0.070 41.861
0.193 0.779 0.0074
0.001279
− 0.061 ± 0.092
− 0.17 ± 0.145 − 0.031 ± 0.046 18.821
0.888 1.231 0.0144
0.000914
− 0.008 ± 0.092
− 0.14 ± 0.145 − 0.004 ± 0.046 29.248
–
0.002370
− 0.235 ± 0.139 − 0.355 ± 0.229 − 0.118 ± 0.070 72.800
0.010 0.870 0.0088
0.000695
− 0.479 ± 0.139 − 0.655 ± 0.229 − 0.240 ± 0.070 21.360
0.090 1.180 0.0099
0.881 0.0062
Isotopic compositions are expressed relative to IRMM-014 (Eqs. (1) and (2)). The chemical compositions of geostandards are from [17] for Ti in IF-G
and [33] for others. The chemical compositions of supracrustal rocks from the NSB are from [12]. Uncertainties are 95% confidence intervals. They
correspond to 2–8 replicate analyses of the same solution and were calculated using student's t-distribution. For the samples that have been
micromilled from the section of IN05007, the numbers in parentheses are the distances from the bottom (Fig. 3).
366
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
known as the χ2 per degree of freedom or reduced χ2)
for the microdrill analyses of individual magnetite bands
in sample IN05007 is 2.64. This is outside the 95%
confidence interval of 0.52 to 1.64 given by χ2 statistics.
At the scale of the section studied (10 cm2), the Fe
isotopic composition appears to vary more than predicted
from analytical uncertainty (Table 1). This may be a relict
depositional feature [19] or may indicate the presence of
inter-mineral equilibrium fractionation in BIFs at N 550 °C
[36] if the micro-drilled magnetite was mixed with other
phases. The average of the 25 microsample analyses
(+0.284 ± 0.012‰/amu) is identical within error to the
average of the 2 bulk measurements performed on the
same sample (+0.236 ± 0.054‰/amu).
The Fe/Ti ratio has proven to be useful for tracing Fe
mobilization in altered basalts [37] and distinguishing
between metasomatized igneous rocks and rocks of
chemical sedimentary origin such as BIFs [17]. The
mafic and ultramafic rocks, quartz–biotite (polymict
conglomerate?) schists, and pelitic schists from the NSB
have identical Fe/Ti ratios (18–42) as metaigneous rocks
from the Akilia association. The only exception is
IN05045, a banded amphibolite with a low Fe/Ti ratio of
8.50 and low Fe isotopic composition (− 0.15‰/amu).
For comparison, the bulk silicate Earth (BSE) Fe/Ti ratio
is 44.5. All candidate rocks of BIF origin investigated in
Fig. 4. Comparison between FFe (‰/amu relative to IRMM-014) and the
atomic ratio of Fe to Ti in amphibolites, quartz–biotite conglomeratic (?)
schists, psammitic schists, and quartz± magnetite± amphibole/pyroxene
rocks from the NSB. The Fe/Ti ratio of the bulk silicate Earth is 44.53
[40] and the FFe of modern (b0.1 Ga) magmatic rocks (continental
basalts, OIBs, and MORBs) is ∼0.05‰/amu [30,54]. The hatched field
is the convex hull enclosing measurements of BIFs and ferruginous
quartz–amphibole/pyroxene rocks from the ISB and AA (SW Greenland) [17]. The heavy line is the trajectory for mixing between IF-G (BIF
geostandard from the ISB) and igneous material.
this study have extremely high Fe/Ti ratios compared to
surrounding igneous rocks (Fig. 4). For some, Ti was
below detection level and only a lower limit could be
derived for the Fe/Ti ratio. In those cases, the Fe/Ti ratio
is at least 3 orders of magnitude higher than the ratio
measured in mafic and ultramafic units. This is similar to
what has been observed for the metamorphosed chemical
sedimentary precipitates documented from the ISB and
Akilia association [6,7,17,38,39]. IN05048, a quartz–
pyroxene rock with Fe isotopic composition indistinguishable to igneous rocks has high Fe/Ti ratio (720.9).
5. Discussion
5.1. Protolith identification
Quartz–biotite schists of possible conglomeratic origin
have Fe isotopic compositions that resemble mafic and
ultramafic rocks from the NSB, but are distinct from
modern basalts (Fig. 4). Cates and Mojzsis [12] showed
that bulk element compositions normalized to NASC
(North American Shale Composite) are consistent with a
detrital origin for this unit. Modern clastic sediments
formed in a variety of environments (aerosols, loess, soils,
marine sediments, suspended river loads) have Fe isotopic
compositions similar to igneous rocks [20]. Thus, Fe
isotopes are not useful to distinguish between different
sources of clastic sediments.
Two interpretations on the possible origin of the
pelitic schist unit have been proposed in the literature.
Nadeau suggested that they might derive from metasomatic alteration of a massive amphibolite unit [13] while
Stevenson and David described them as “metapelites”
[25]. We find that the FFe value of this unit (− 0.2‰/amu)
is significantly lower than that of amphibolites (∼0‰/
amu) while the Fe/Ti ratio is close to that of surrounding
igneous rocks. The later observation suggests that Fe was
not extensively mobilized and that the low FFe values
measured in these rocks are likely primary depositional
features. Such low values have been interpreted in
Archean shales and greywackes to represent the imprint
of authigenic phases (carbonates, oxides, and sulfides)
carrying isotopically fractionated Fe [16]. Our results
support a sedimentary rather than a metasomatic origin
for this unit.
Iron isotopes have proven to be very useful to
distinguish between rocks of true chemical sedimentary
origin and metasomatized igneous rocks in SW Greenland
[17]. This is an important step in early Earth studies
because all known Eoarchean supracrustal rocks have
been metamorphosed to at least amphibolite facies and
recognition of their protoliths is rarely straightforward.
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
Our results show that most banded quartz–magnetite
rocks from the NSB have heavy Fe isotopic compositions
and high Fe/Ti (Fig. 4). Many samples fall in the field
covered by BIFs from the ISB and Fe-bearing quartz–
pyroxene units from the Akilia association [17]. High Fe/
Ti in these rocks probably reflects the low solubility of Ti
in seawater, a feature that is transferred to the chemical
sediment at precipitation. Furthermore, the high FFe values
of these rocks may be the imprint of partial Fe oxidation
[15,17], which is known to enrich the precipitate in the
heavy isotopes [21–24]. These results clearly establish the
presence of rocks of BIF parentage in the NSB. Sample
IN05048 (a quartz–pyroxene rock) has an Fe isotopic
composition (+0.038 ± 0.023‰/amu) similar to igneous
rocks. This signature does not reflect simple admixture of
igneous material because the Fe/Ti ratio of this rock is high
(721). Fig. 4 shows a simple two-component mixing curve
between IF-G, a BIF from Isua with the highest Fe/Ti ratio
measured thus far (10343) and heavy Fe isotopic
composition (+0.316‰/amu) [17,18], and igneous material (Fe/Ti = 45 and FFe = +0.04‰/amu) [20,40]. At an Fe/
Ti ratio of 721, the FFe value of the mixture would still be
+0.300‰/amu. We conclude that the range of Fe isotopic
compositions measured in BIFs from the NSB (from 0 to
+0.4‰/amu) is a primary feature of the rock.
5.2. Metamorphic overprint
In this section, we examine the effect of upper
amphibolite/lower granulite facies metamorphism on
the Fe isotopic composition of BIFs in more detail.
Comparison with the more extensively characterized
ISB provides some insights on what may have been the
metamorphic reactions involved. Magnetite-bearing
quartzite units from the NSB contain amphiboles (and
sometimes pyroxenes) that presumably formed by
reaction between primary carbonates and quartz following the general reaction [41],
primary carbonates þ quartz
þ water→secondary carbonates
þ amphiboles=pyroxenes þ carbon dioxide:
ð3Þ
The primary carbonates may have had different
origins. They could have been deposited as primary
chemical sediments together with Fe-hydroxides, as is
observed in younger, better preserved BIFs such as those
of the Hamersley Range (Western Australia) or the
Transvaal Supergroup (South Africa) [15,41]. They could
have been introduced in the rock through carbonate
metasomatism as has been extensively documented in the
ISB [42–44]. It is also possible that they formed during
367
diagenesis by microbial dissimilatory iron reduction of
hydrous ferric oxide [19,45]. Carbonates were found and
characterized in IN05007 (Figs. 2 and 3) but given the
degree of metamorphism of the rock, it is impossible to
distinguish between sedimentary, diagenetic, or metasomatic origins based solely on petrography or chemistry.
With amphiboles and pyroxenes derived from reaction
between carbonates and quartz, magnetite (Fe3O4) is a
major carrier of Fe in quartzitic supracrustals (including
BIFs) from the NSB (Figs. 2 and 3). Magnetite is not an
authigenic phase but was probably formed diagenetically
shortly after deposition from the transformation of ferro–
ferric hydroxide precursors [41,46]. After formation,
magnetite is not involved in any metamorphic reaction
[19,41]. There is little doubt that magnetite was ultimately
derived from chemical precipitation in the water column
and its Fe isotope composition may help us understand the
mechanism that caused Fe oxidation and BIF precipitation
in the Eoarchean oceans. The FFe values of individual
phases were presumably equilibrated at some scale and
one cannot rely on single-phase measurements to infer
their compositions before metamorphism.
Two approaches can be used to estimate the initial
composition of magnetite. In the element map of IN05007
(Fig. 3, 10.00 cm2), the surface area of magnetite is
1.38 cm2 and that of amphibole (mainly cummingtonite)
is 2.15 cm2 (quartz and other phases occupy 6.34 cm2 and
0.14 cm2, respectively). Assuming that the surface
distribution is representative of the volume distribution,
this analysis shows that approximately 74% of Fe is in
magnetite and 26% is in cummingtonite (using densities
of 5.1 and 3.4 g/cm3 and Fe concentrations of 0.724 and
0.246 g/g for magnetite and cummingtonite, respectively).
Dideriksen et al. [47] and Markl et al. [48] analyzed the Fe
isotopic compositions of natural carbonate minerals
formed in a variety of geological settings (magmatic,
metamorphic, hydrothermal, and sedimentary). All samples, without exception, have light Fe isotope compositions relative to IRMM-014, from −0.6 to 0‰/amu. This
is similar to the range of FFe values reported for
carbonates in well-preserved BIFs from the Kaapvaal
Craton (South Africa) and the Biwabik formation
(Minnesota), from −0.5 to 0.1‰/amu [15,19]. One can
reasonably assume that the carbonates that were involved
in the formation of amphiboles had light Fe isotopic
compositions. If the rock behaved as a closed system for
Fe, it is straightforward to make the mass balance
(0.26 ⁎ FFe carbonate + 0.74 ⁎ FFe magnetite = FFe bulk)
and calculate that the initial FFe value of Fe-oxides in
IN05007 must have been in the range 0.3 to 0.5‰/amu to
explain the present value of 0.24‰/amu (for FFe values of
carbonates of 0 and −0.5‰/amu, respectively).
368
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
Another way to tackle this question is to recognize
that carbonates should leave an imprint on the bulk
chemistry of the rock if they were present at the time of
formation. The possible cations present in appreciable
quantities in the carbonates from rocks of the composition explored in this study are Fe, Mn, Ca, and Mg. In
Fig. 5, the FFe values of bulk ferruginous quartzites from
the NSB are plotted as a function of Ca/Fe, Mg/Fe, and
Mn/Fe ratios. Simple binary mixing is represented as
straight lines in such diagrams. As discussed in Section
5.2, the low FFe value of IN05048 (0.038 ± 0.023‰/
amu) cannot be attributed to admixture of an igneous
component. Instead, this sample has high Ca/Fe, Mg/Fe,
and Mn/Fe and its isotopic composition may reflect the
contribution of carbonates with low FFe values (Fig. 5).
Note that similar relationships are obtained for ratios of
transition metals with ionic radii similar to Fe (Ni/Fe, Cu/
Fe, and Zn/Fe, see source data in [12]). If the mixing
interpretation is correct, then one can extrapolate mixing
lines shown in Fig. 5 to Ca/Fe or Mg/Fe equal to 0
(magnetite contains insignificant Ca or Mg) and estimate
the FFe value of an Fe-oxide end-member. The result of
this analysis is that the initial FFe value of the primary
Fe-oxide phase must have been around 0.3–0.4‰/amu,
in agreement with the value inferred previously (0.3 to
0.5‰/amu). The relationship between FFe and Mn/Fe is
less useful because the Mn/Fe ratio of the Fe-oxide endmember is not well defined. It is interesting to note that
IF-G, a BIF geostandard from Isua [33] falls close to the
mixing lines defined by quartz–magnetite rocks from the
NSB on all diagrams [17,30,33]. This analysis also
suggests that the carbonate end-member had light Fe
isotopic composition, lower than 0‰/amu. Theoretically, it should be possible to follow the same approach
using a trivalent ion that can partition into magnetite (Al,
Cr, V) but not in carbonate to derive the isotopic
composition of the carbonate end-member. However, no
clear correlation was found between FFe values and Al/
Fe, Cr/Fe, and V/Fe ratios, possibly because these
elements are present at trace levels and their concentrations in the source fluid may have been variable. If we
assume that the primary carbonates had FFe values
between − 0.5 and 0‰/amu, then it is possible to
calculate the Ca/Fe, Mg/Fe, and Mn/Fe ratios of this endmember. The composition of the primary carbonate endmember is thus estimated to be between Ca0.2Mg0.3Fe0.5
and Ca0.3Mg0.4Fe0.3. It is a virtual component that must
reflect mixing between two or more carbonate phases
(e.g., ankerite 0.5 × Ca0.5Mg0.3Fe0.2 + siderite 0.5 × Mg0.3
Fe0.7 = Ca0.25Mg0.3Fe0.45). In a ternary diagram, this
composition falls on a line that joins amphiboles and
carbonates now present in IN05007, which is required by
Fig. 5. Binary mixing diagrams between hypothetical end-members Feoxide (Fe-ox) and carbonate (Carb). The filled symbols are measurements of quartz ± magnetite ± amphibole/pyroxene rocks from the
NSB (Table 1) and the empty diamond corresponds to the composition
of IF-G, a BIF geostandard from the ISB [17,33]. Extrapolating
the trends to Ca/Fe = 0 or Mg/Fe = 0 gives the composition of the Feoxide end-member, 0.3–0.4‰/amu. The equations of the best-fit lines
are y = − 0.8016x + 0.3279 for Ca, y = −0.7285x + 0.4112 for Mg, and
y = −10.917x + 0.4095 for Mn.
mass balance if these phases formed from reaction
between primary carbonates and quartz (Fig. 6). The
analysis presented here allows us to better define what
may have been the metamorphic reactions involved in
the formation of the mineral assemblage preserved in
IN05007,
1=2Ca0:25 Mg0:35 Fe0:15 þ1=2Mg0:3 Fe0:7
zfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflffl}|fflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflffl{
Ca0:25 Mg0:32 Fe0:43 CO3
ankeriteþsiderite
þ SiO2 þ H2 O Y Ca0:2 Mg2:7 Fe4:1 ½Si8 O22 ðOHÞ2
quartz
water
cummingtonite
þ Ca1:8 Mg2:7 Fe2:5 ½Si8 O22 ðOHÞ2 þ Ca0:95 Mg0:02 Fe0:03 CO3
calcite
actinolite
þ Ca0:54 Mg0:26 Fe0:21 CO3 þ
ankerite
CO2
carbon dioxide
:
ð4Þ
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
369
There are 7 free coefficients and 7 elements and this
metamorphic reaction can be balanced (Fig. 6). The
possible presence of siderite in the primary mineral
assemblage supports deposition under high PCO2
[49,50]. Whether this high PCO2 translates into high
PCO2 in the Eoarchean atmosphere or reflects local
conditions of precipitation is open to debate [51,52].
Because magnetite does not participate in these
metamorphic reactions [41], its Fe isotopic composition
may be less affected than other phases by metamorphism. This feature was documented during contact
metamorphism in the Biwabik iron-formation (Minnesota), where primary Fe isotope heterogeneity of
magnetite was preserved, up to temperatures in excess
of 500 °C [19]. The self-diffusion coefficient of Fe in
magnetite at 500 °C depends on the activity of O2 but is
greater than ∼10− 16 cm2/sp[53].
ffiffiffiffiffiffiffiffiffiffi The corresponding
diffusion length for 1 My ð 4DtÞ is ∼ 0.1 cm. Thus,
over timescales relevant to regional metamorphism, the
Fe isotope signature of magnetite in rocks of BIF
parentage may have been homogenized at the scale of a
hand specimen. The best-preserved quartz–magnetite
rocks in the NSB have well-preserved bands of
Fig. 7. Possible reaction paths for O2-mediated abiotic oxidation.
During homogeneous oxidation, oxidation of Fe(II)aq followed by
polymerisation of Fe(III)aq allows formation of a solid precipitate.
During heterogeneous oxidation, part of Fe(II)aq is first adsorbed on
the precipitate to then be oxidized. The second reaction path is
autocatalytic in the sense that the rate of oxidation increases with the
concentration of ferrihydrite in suspension. Assuming that Fe(III)aq
and Fe(II)ad have very short lifetimes, the system can be modelled
using Rayleigh transfer of Fe isotopes from Fe(II)aq to Fe(III)s (see text
and Appendix for details).
Fig. 6. Carbonate–quartz metamorphic reaction in the Ca–Mg–Fe
ternary diagram. The measured compositions of amphiboles (cummingtonite and actinolite, green squares) and carbonates (ankerite and calcite,
yellow squares) are from Fig. 2. The possible composition of the primary
carbonate end-member was calculated from Fig. 5 assuming an FFe
value between −0.5 and 0‰/amu [15,47] (heavy yellow line, l1). It
must also lie along the line that connects the average compositions of
amphiboles and carbonates, the products of the metamorphic reaction
carbonate + qtz + H2O → cum + act + ank + cal + CO2 (solid line l2).
The calculated composition of the primary carbonate endmember (Ca0.25Mg0.325Fe0.425) may reflect mixing between ankerite
and siderite (e.g., 1/2Ca0.5Mg0.35Fe0.15 + 1/2Mg0.3Fe0.7, dashed line l3).
(For interpretation of the references to color in this figure legend, the
reader is referred to the web version of this article.)
magnetite, quartz, and amphibole. Although deformed,
these bands most likely reflect original sedimentary
strata. Multiple measurements of IN05007 using a
micromill apparatus (Fig. 3) show that magnetite has
almost homogeneous Fe isotopic composition at the
scale of the sample (small variations may be present but
it is unclear whether they represent a depositional
feature or a problem with sampling single phases, see
Section 4). Overall, it seems that in the Nuvvuagittuq
supracrustal samples we studied, precipitated Fe (hydr)
oxides had almost constant Fe isotope compositions,
around +0.3‰/amu.
5.3. Homogeneous/heterogeneous oxidation, source of
Fe and mechanism of BIF precipitation
The aim of this section is to integrate the results
obtained on the Fe isotopic composition of magnetite in
370
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
a feasible scenario for banded iron-formation deposition. As discussed previously, various chemical pathways may have led to the formation of magnetite
[19,41]. Below, we shall make the simple assumption
that Fe(II)aq was oxidized and precipitated as Fe(III)s.
Upon diagenesis, part of Fe(III)s was reduced to form
mixed valence magnetite. Beard and Johnson [54] and
Dauphas and Rouxel [30] developed a two-stage
distillation model to calculate Fe isotope fractionation
during oxidation–precipitation of Fe(II)aq. This model
may be relevant to Fe(II)aq oxidation at low pH, where
significant Fe(III)aq can form. At circum-neutral pH,
polymerization of ferric hydroxide occurs on a short
time-scale so one would not expect large amounts of Fe
(III)aq to be present at any time [55].
Walker [56] reviewed the possible conditions in the
Archean ocean when BIFs formed. For PCO2 of
∼ 0.1 atm, the pH would have been around 7 and the
concentration of Fe(II)aq around 10− 3 mol/L (also see
[46]). At these conditions of circum-neutral pH and
high Fe concentrations, it is well documented that Fe
oxidation can proceed along two paths [57–59] (Fig. 7).
For low concentrations of solids (b 3 mg/L), precipitation proceeds predominantly through oxidation of Fe
(II)aq and subsequent polymerisation of ferric hydroxide
to form colloids which eventually precipitate. In this
case (homogeneous oxidation) the rate is,
d½FeðIIÞaq dt
¼ −k1 ½O2 ½OH− 2 ½FeðIIÞaq :
ð5Þ
This formula applies to most natural waters at present
[60]. When the concentration of solids in suspension
increases to above ∼ 3 mg/L, Tamura et al. [57] showed
that an autocatalytic effect occurs whereby the rate of
oxidation increases with the concentration of Fe(III)
hydroxide present. In this case (heterogeneous oxidation) the rate is,
d½FeðIIÞaq dt
¼ −k2 ½O2 ½FeðIIÞaq ½FeðIIIÞs =½Hþ :
ð6Þ
The two processes can happen concurrently. At
constant pH and PO2 we can write,
d½FeðIIÞaq dt
¼ −k½FeðIIÞaq −k V½FeðIIÞaq ½FeðIIIÞs :
ð7Þ
The ratio k′/k only depends on the nature of the
phase precipitated and pH (∝[H+]). It is equal to
∼ 5 × 103 mol− 1 L at pH 7 for an amorphous ferric
hydroxide formed by hydrolysis of Fe(III) perchlorate [57,58]. The dimensionless parameter (k is in s− 1
and k′ in mol− 1 s− 1) that governs the evolution of the
system is,
c¼
k V½FeðIIÞaq 0
k
ð8Þ
where [Fe(II)aq]0 is the initial concentration of dissolved
ferrous iron. For an initial concentration of 10− 3 mol
L− 1, c is equal to 5. This is higher than 1 and
heterogeneous oxidation could clearly be relevant to
BIF precipitation. In a closed system with only Fe(II)aq
present at the beginning, the fraction of Fe precipitated
is given by,
FeðIIIÞs
FeðIIÞ0aq
¼
eð1þcÞu −1
;
c þ eð1þcÞu
ð9Þ
where u is the dimensionless variable kt [57,59]. Iron
isotopes could be fractionated to different extents along
the two oxidation paths. Assuming that Fe oxidation
results in Rayleigh-type transfer of isotopes (i.e., after
precipitation there is no isotopic exchange between the
precipitate and the fluid), then the composition of the
precipitate can be calculated,
IIaq
IIIs
FFe
¼ FFe
ð0Þ
þ
1þc
1þc
III he
III ho
III he
;
ðD
−D
Þu
þ
D
ln
IIaq
IIaq
IIaq
1−eð1þcÞu
c þ eð1þcÞu
ð10Þ
IIaq
where FFe
(0) is the initial composition of ferrous
III he
III ho
iron, ΔIIaq and ΔIIaq
are the isotopic fractionations in
‰/amu between the instantaneous precipitate and the
pool of Fe(II)aq for heterogeneous and homogeneous
oxidations, respectively (see Appendix for details).
The three processes most commonly advocated for
Fe oxidation in an anoxic atmosphere are, (i) oxidation
from O2 generated by oxygenic photosynthesis, (ii)
oxidation during anoxygenic photosynthesis where Fe
(II)aq is used as an electron donor in place of H2O, and
(iii) photo-oxidation of Fe at the surface of oceans by
energetic photons penetrating through an atmosphere
transparent to UV radiation. As yet, the effect of photooxidation (iii) on the isotopic composition of Fe has
not been documented. The other processes (i, ii)
are known to enrich Fe(III)s in the heavy isotopes of
Fe relative to Fe(II)aq. Welch et al. [22] and Balci et al.
[24] studied Fe isotope fractionation during oxidation
at low pH (b 3), which may not be relevant to BIF
formation. The isotopic fractionation between Fe(II)aq
and Fe(III)aq at room temperature under sterile conditions is + 1.4‰/amu [22]. Bullen et al. [21] measured a
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
lower fractionation between Fe(II)aq and Fe(III)s of
around 0.5‰/amu during precipitation of ferrihydrite at
pH around 6, which may partly reflect kinetic isotope
fractionation during precipitation of Fe(III)aq into Fe
(III)s [54]. These values would be relevant to oxidation
from O2 generated by oxygenic photosynthesis. Croal
et al. [23] measured the Fe isotopic fractionation between Fe(II)aq and Fe(III)s during anoxygenic photosynthesis at pH ∼ 7 and found values of around 0.7‰/
amu. Again, this value may be the superposition of
effects associated with Fe(II)aq–Fe(III)aq oxidation and
Fe(III)aq–Fe(III)s precipitation. Precipitation of Fe(III)aq
into hematite is associated with little equilibrium
fractionation but with a significant kinetic isotope effect
that enriches the precipitate in the light isotopes [61].
This fractionation increases as the rate of precipitation
increases and can reach values as low as − 0.5‰/amu.
The isotopic fractionation associated with heterogeneous (O2-mediated) oxidation is unknown. Adsorption
may play a role in this process, which can potentially
create isotopic fractionation, with Fe(II)ad being
enriched in the heavy isotopes of Fe relative to the
pool of Fe(II)aq. Icopini et al. [62] argued that the
fractionation may be as large as 1.5‰/amu but values of
around 0.2 to 0.4‰/amu may be more reasonable
[63,64]. Multiple lines of evidence suggest that the
primary source of Fe found in BIFs was hydrothermal
[12,26,65]. Present hydrothermal fluids sampled at
vents along ocean-ridges have negative FFe values,
centred on − 0.15‰/amu [66–68]. The net isotopic
fractionation between Fe-oxide precursors in banded
quartz–magnetite rocks in the NSB (FFe = 0.35‰/amu)
and the possible hydrothermal source of Fe (FFe =
− 0.15‰/amu) is therefore ∼ 0.5‰/amu.
For the purpose of illustration, we shall adopt the
following set of parameters to model Fe isotope fractionation during BIF precipitation using Eqs. (9) and (10),
IIaq
(i) O2-mediated abiotic oxidation: FFe
(0) = −0.15‰/
III ho
III he
amu, c = 5, ΔIIaq = 0.5‰/amu, and ΔIIaq = 0.2‰/amu.
(ii) Anoxygenic photosynthetic oxidation: standard
IIaq
Rayleigh distillation, equivalent to FFe
(0) = −0.15‰/
III ho
amu, c = 0, ΔIIaq = 0.7‰/amu.
In Fig. 8, we report the calculated Fe isotopic composition of the precipitate as a function of the fraction of
Fe precipitated. As shown, the Fe isotopic composition
of the precursor of magnetite can be explained by both
scenarios (i) and (ii). The fact that precipitated Fe-(hydr)
oxide in Isua and Nuvvuagittuq had almost constant Fe
isotopic composition, around 0.3–0.4‰/amu, suggests
that the fraction of Fe precipitated remained constant
and probably represented a small fraction of the total
inventory of dissolved Fe, presumably less than 10%. To
371
Fig. 8. Iron isotopic composition of Fe precipitate as a function of the
fraction of Fe precipitated (from Eqs. (9) and (10), also see Appendix).
The parameters used in the model calculation are for O2-mediated
homogeneous/heterogeneous oxidation, FIIaq
Fe (0) = −0.15‰/amu,
ho
III he
c = 5, ΔIII
IIaq = 0.5‰/amu, and ΔIIaq = 0.2‰/amu and for anoxygenic
III ho
photosynthetic oxidation, F IIaq
Fe (0) = −0.15‰/amu, c = 0, ΔIIaq = 0.7‰/
amu (see text for details). The composition of the precursor of magnetite
in chemical sediments from Nuvvuagittuq (∼0.35‰/amu, Fig. 5) is
shown for comparison (grey band).
summarize, the heavy Fe isotopic composition found in
magnetite of quartz–magnetite rocks (banded ironformations) from the Nuvvuagittuq supracrustal belt
can be explained if Fe was derived from a hydrothermal source and was only partially oxidized in the water
column.
5.4. The Fe isotopic composition of Eoarchean mafic
and ultramafic magmas
Mantle-derived magmas and peridotites are the only
samples available for estimating the Fe isotopic
composition of the Bulk Silicate Earth (BSE). Beard et
al. [20] showed that igneous rocks sampled in a variety
of geological settings have almost uniform isotopic
compositions. However, further detailed studies have
revealed the presence of large-scale Fe isotope variations in mantle and magmatic rocks [34,69–72]. For
instance, peridotites and granitoids have a large range of
FFe values, from − 0.2 to +0.2‰/amu [34,71]. Most
mantle peridotites analyzed so far have light Fe isotopic
compositions, down to − 0.2‰/amu [70–72]. In some
cases, FFe values of bulk peridotites correlate with
chemical indicators of depletion and oxidation [71]. A
possible interpretation is that during partial melting, the
melt does not have the same isotopic composition as the
starting material because there is mineral/melt isotope
fractionation and/or there is inter-mineral isotope
fractionation and melting is non-modal. For degrees of
melting between 0 and 40% of a spinel lherzolite
composition, Williams et al. [71] showed that the
372
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
magmas produced could be fractionated by as much as
0.07‰/amu relative to the source and yet show little
variation from one to another over a large melting
interval. Thus, modern continental basalts, MORBs, and
OIBs could be fractionated relative to the mantle but still
show almost constant Fe isotopic compositions, around
+0.05‰/amu (Fig. 9). Lherzolites may be a better proxy
of the composition of the BSE [72]. The 7 bulk rock
lherzolite measurements published so far range from 0
to + 0.03‰/amu, with a weighted average of about
+0.02 [17,71,72] (dominated by a very precise measurement of a single xenolith from Kilbourne Hole
[71]). This is lower than the composition of modern
basalts but even the most pristine peridotites have
experienced melt depletion/metasomatism and it is
crucial at this stage to extend the database of otherwise
well-characterized fertile lherzolites.
Although this is at the limit of resolution with present
analytical uncertainties, mafic and ultramafic magmatic
rocks from the ca. 3.83 Ga Akilia island [7] and ca.
3.75 Ga NSB [12] localities may have light Fe isotopic
compositions (+ 0.004 ± 0.016 and − 0.006 ± 0.022‰/
amu respectively, Table 1, Fig. 9) compared to modern
basalts [17]. This could reflect modification of initial Fe
isotopic compositions after emplacement of the rocks.
Indeed, pervasive carbonate and alkali metasomatism
has been documented in SW Greenland [42–44] and its
effect on Fe isotopes has not been explored. However, it
is unlikely that the diverse suite of samples analyzed in
this study (Table 1) and in [17] would have experienced
the same degree of metasomatic alteration and would all
show the same shift in FFe values. This suggests that the
light Fe isotopic compositions measured in Akilia and
the NSB may represent primary magmatic signatures.
Poitrasson et al. [35] measured a 3.5 Ga komatiite from
Barberton (geostandard WITS-1) and also found a low
FFe value of 0.019 ± 0.008‰/amu. In contrast, Weyer
et al. [72] measured a 2.7 Ga komatiite from Alexo
(geostandard KAL-1, Ontario, Canada) and found a
value that is indistinguishable from modern basalts
(0.036 ± 0.007‰/amu). The possible change in Fe
isotopic composition of magmas formed by partial
melting of the mantle (from ∼ 0‰/amu in the early
Archean to ∼ 0.05‰/amu at present) could reflect
changes in the thermal regime of the mantle and the
conditions of magma generation (e.g., fraction of partial
melting, temperature, buoyancy). Clearly, the database
of Fe isotope analyses of Archean magmas must be
extended to ascertain whether the secular variation
discussed here is real and if it is, address its origin.
6. Conclusions
Fig. 9. Comparison between FFe values (‰/amu relative to IRMM014) of carbonaceous chondrites (Table 1), Vesta and Mars (HED and
SNC achondrite meteorites), N3.5 Ga mafic and ultramafic magmatic
rocks (Table 1), lherzolites, modern (b0.1 Ga) continental basalts,
OIBs (Table 1), and MORBs (see text and [30] for an extensive
reference list). The curves are kernel density estimates with automatic
bandwidth selection (all curves have the same surface area) [75,76].
The recently discovered 3.75 Ga Nuvvuagittuq
supracrustal belt provides us with a means to investigate
the surface chemistry of the Eoarchean Earth from the
perspective of a “new” terrane contemporaneous with
the Isua locality. The best-preserved banded quartz–
magnetite rocks (BIFs) contain alternating bands of
magnetite and quartz/amphibole (actinolite and cummingtonite) with minor sulfide (pyrite) and carbonate
(calcite and ankerite). The mineral paragenesis is best
explained by amphibolite facies metamorphism of a BIF
protolith consisting of magnetite, quartz, and carbonate
as major constituents.
The heavy Fe isotopic compositions (around +0.3‰/
amu relative to IRMM-014) and high Fe/Ti ratios (up to
∼ 100× the ratio of surrounding igneous rocks)
measured in some BIFs demonstrate that these rocks
are true chemical sediments precipitated from seawater.
These samples can therefore be used in future studies to
document the chemistry of Eoarchean oceans.
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
We interpret the dispersion in Fe isotope compositions
at a bulk sample scale in banded iron-formations (from 0
to +0.3‰/amu) to reflect binary mixing between Feoxides and carbonates. By extrapolating mixing lines in
Ca/Fe vs FFe and Mg/Fe vs FFe space to Ca/Fe = 0 and
Mg/Fe = 0, one can estimate the Fe isotopic composition
of the pure Fe-oxide end-member, which is between +0.3
to +0.4‰/amu. On the other hand, if it is assumed that the
primary carbonates had Fe isotope compositions between
−0.5 and 0‰/amu, which covers the range of values
measured so far in carbonates from a variety of geological
settings, one can derive the chemistry of the carbonate
end-member. In a Fe–Mg–Ca ternary diagram, the
calculated composition falls along a line connecting
carbonates and amphiboles now present in the rock,
which is consistent with derivation of these phases from
reaction between quartz and primary carbonate. The
composition of the primary carbonate does not correspond
to a real phase but must reflect a mixture, possibly
between ankerite and siderite. The possible presence of
siderite supports deposition under high PCO2. Iron
isotopes can thus be used as an independent means to
infer the protolith mineralogy of heavily metamorphosed
rocks that may otherwise be inaccessible, especially if the
rocks did not behave as a perfect closed-systems.
A distillation model involving two possible oxidation
paths (homogeneous and heterogeneous) was developed. It is used to demonstrate that the heavy Fe isotopic
composition inferred for the precursor phase(s) of magnetite can be explained by partial oxidation of Fe(II)aq
deriving from a hydrothermal source through oxygenic
or anoxygenic photosynthesis.
Preliminary data suggest that mafic and ultramafic
magmatic rocks from the NSB have lower Fe isotopic
compositions than modern basalts. A more extensive
characterization of magmas generated during the early
history of the Earth is required to ascertain whether this
effect is real or not, work that is currently in progress.
As far as Fe isotopes and bulk chemistry are
concerned, rocks of likely BIF protolith from the
Nuvvuagittuq supracrustal belt share many similarities
with banded iron-formations from the Isua supracrustal
belt. Based on the results of this study, we find that the
similarity is such that the two localities may in fact be
genetically related and provide us with a markedly
consistent picture of the chemistry of the early Earth's
hydrosphere.
373
the fruitful discussions and help with MC-ICPMS
and SEM analyses. Constructive comments by 3
anonymous reviewers substantially improved the
manuscript and were greatly appreciated. We are
grateful to the Pituvik Landholding Corporation and
the people of Inukjuak, Québec for the logistical help
and J.D. Adam for the field assistance. This work
was supported by the National Aeronautics and
Space Administration through grants NNG06GG75G
(to ND), NAG5-13497 and NASA Astrobiology
Institute (to SJM) and NAG5-13497. Correspondence
and requests for materials should be addressed to
[email protected].
Appendix A. Homogeneous/heterogeneous
oxidation
At neutral pH, ferrous iron in solution can be oxidized
through the homogeneous and heterogeneous paths,
homogenous
heterogenous
zfflfflffl}|fflfflffl{ zfflfflfflfflfflfflfflffl}|fflfflfflfflfflfflfflffl{
dFeIIaq
¼ − kFeIIaq − k VFeIIaq FeIIIs :
dt
ðA1Þ
For a system with all iron as FeIIaq at time 0, we have,
dFeIIaq
IIaq
¼ −kFeIIaq −k VFeIIaq ðFeIIaq
Þ:
0 −Fe
dt
ðA2Þ
Using the dimensionless variables and parameters,
u = kt, f = FeIIaq/Fe0IIaq, and c = k′Fe0IIaq /k, one can write
Eq. (A2) as,
df
¼ −f −cf ð1−f Þ:
du
ðA3Þ
Integration is straightforward (pose h = f − 1 and use
the method of variation of the constant),
f ¼
1þc
:
c þ eð1þcÞu
ðA4Þ
For the total amount of iron precipitated (g = FeIIIs/
) we have,
Fe0IIaq = 1 − f
g¼
eð1þcÞu −1
:
c þ eð1þcÞu
ðA5Þ
Acknowledgments
The expressions had been derived in Tamura et al.
[57]. We now turn to isotopes which is a new
development of this work. Let us denote a and b two
isotopes of Fe and ϕai → j the flux of a from reservoir i to
j. We have for the pool of ferrous iron in solution,
We wish to thank M. Wadhwa, P.E. Janney, A.M.
Davis, M. van Zuilen, J. Levine, and R. Yokochi, for
dFeIIaq
ho
a
¼ −/IIaqYIII
−/aIIaqYIII he :
a
dt
ðA6Þ
374
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
We note R the ratio of the abundances of isotopes a
and b, with Ri the ratio in reservoir i and Ri→j the ratio
in the flux from i to j. Eq. (A6) can be written,
IIaq
dFeIIaq
b R
dt
¼ −/bIIaqYIII ho RIIaqYIII ho −/bIIaqYIII he RIIaqYIII he :
ðA7Þ
We impose Rayleigh conditions between Fe IIaq–
Fe
and FeIIaq –FeIII he . Mathematically, this means
III ho
III he
that the ratios αIIaq
= RIIaq → III ho /RIIaq and αIIaq
=
IIaq → III he IIaq
R
/R
must remain constant.
After some rearrangement of Eq. (A7), it can be
shown that,
III ho
dFeIIaq
dlnRIIaq
þ FeIIaq
dt
dt
ho
IIaq
he
IIaq
IIaq
¼ −aIII
kFe
−aIII
ðFeIIaq
Þ:
IIaq
IIaq k VFe
0 −Fe
ðA8Þ
In dimensionless coordinates, this takes the form,
df
dlnRIIaq
ho
III he
þf
¼ −aIII
IIaq f −aIIaq cf ð1−f Þ:
du
du
ðA9Þ
Isotope variations are in most cases small and we use
the following notation, δi = (Ri/Rstd − 1) × 103. The δ
notation is used here to improve readability. We also
write Δij = (αij − 1) × 103. Using these notations and the
fact that δ/103 ≈ 0, Eq. (A9) can be rearranged,
ddIIaq
dlnf
ho
3
−ðDIII
¼ −103
IIaq þ 10 Þ
du
du
he
3
−cðDIII
IIaq þ 10 Þð1−f Þ:
ðA10Þ
This equation can be integrated (inject the relationship between dln f/du and f from Eq. (A3) into Eq.
(A10)),
III he
ho
dIIaq ¼ dIIaq
−DIII
IIaq Þu
0 þ ðDIIaq
1þc
he
þ DIII
ln
IIaq
c þ eð1þcÞu
ðA11Þ
The isotopic composition of the precipitate can easily
be calculated from mass-balance considerations,
¼ f dIIaq þ ð1−f ÞdIIIs :
dIIaq
0
ðA12Þ
It follows that,
dIIIs ¼ dIIaq
0
þ
1þc
1þc
III he
III ho
III he
:
ðD
−D
Þu
þ
D
ln
IIaq
IIaq
IIaq
1−eð1þcÞu
c þ eð1þcÞu
ðA13Þ
References
[1] S.A. Bowring, I.S. Williams, Priscoan (4.00–4.03 Ga) orthogneisses from northwestern Canada, Contrib. Mineral. Petrol. 134
(1999) 3–16.
[2] D.O. Froude, T.R. Ireland, P.D. Kinny, I.S. Williams, W.
Compston, I.R. Williams, J.S. Myers, Ion microprobe identification of 4100–4200 Myr-old terrestrial zircons, Nature 304
(1983) 616–618.
[3] S. Moorbath, R.K. O'Nions, R.J. Pankhurst, Early Archaean age
for the Isua iron formation, West Greenland, Nature 245 (1973)
138–139.
[4] A.P. Nutman, V.R. McGregor, C.R.L. Friend, V.C. Bennett, P.D.
Kinny, The Itsaq Gneiss Complex of southern West Greenland;
the world's most extensive record of early crustal evolution
(3900–3600 Ma), Precambrian Res. 78 (1996) 1–39.
[5] A.P. Nutman, S.J. Mojzsis, C.R.L. Friend, Recognition of
≥3850 Ma water-lain sediments in West Greenland and their
significance for the early Archaean Earth, Geochim. Cosmochim.
Acta 61 (1997) 2475–2484.
[6] N. Cates, S.J. Mojzsis, Chemical and isotope evidence for
widespread Eoarchean (3750 Ma) metasedimentary enclaves in
southern West Greenland, Geochim. Cosmochim. Acta 70 (2006)
4229–4275.
[7] C.E. Manning, S.J. Mojzsis, T.M. Harrison, Geology, age and
origin of supracrustal rocks at Akilia, West Greenland, Am. J.
Sci. 306 (2006) 303–366.
[8] A.P. Nutman, B. Fryer, D. Bridgwater, The early Archaean
Nulliak (supracrustal) assemblage, northern Labrador, Can. J.
Earth Sci. 26 (1989) 2159–2168.
[9] L. Schiøtte, W. Compston, D. Bridgwater, U–Th–Pb ages of
single zircons in Archaean supracrustals from the Nain Province,
Labrador, Canada, Can. J. Earth Sci. 26 (1989) 2636–2644 (m).
[10] J. David, M. Parent, R. Stevenson, P. Nadeau, L. Godin, La
séquence supracrustale de Porpoise Cove, région d'Inukjuak: un
exemple unique de croûte paléo-archéenne (ca. 3.8 Ga) dans la
Province du Supérieur, L'exploration minérale au Québec, notre
savoir, vos découvertes, vol. 2002-10, Ministère des Ressources
Naturelles, Québec, DV, 2002, p. 17.
[11] J. David, M. Parent, R. Stevenson, P. Nadeau, L. Godin, The
Porpoise Cove supracrustal sequence, Inukjuak area: a unique
example of Paleoarchean crust (ca. 3.8 Ga) in the Superior
Province, GAC–MAC Annual Meeting, Vancouver, Canada, 2003.
[12] N.L. Cates, S.J. Mojzsis, Pre-3750 Ma supracrustal rocks from
the Nuvvuagittuq supracrustal belt, northern Québec. Earth
Planet. Sci. Lett. (in press).
[13] P. Nadeau, Structural investigation of the Porpoise Cove area,
Northeastern Superior Province, Northern Québec, M.Sc. thesis,
Simon Fraser University (2003).
[14] J.L. Boak, R.F. Dymek, Metamorphism of the ca. 3800 Ma
supracrustal rocks at Isua, West Greenland: implications for early
Archaean crustal evolution, Earth Planet. Sci. Lett. 59 (1982)
155–176.
[15] C.M. Johnson, B.L. Beard, N.J. Beukes, C. Klein, J.M. O'Leary,
Ancient geochemical cycling in the Earth as inferred from Fe
isotope studies of banded iron formations from the Transvaal
Craton, Contrib. Mineral. Petrol. 144 (2003) 523–547.
[16] K.E. Yamaguchi, C.M. Johnson, B.L. Beard, H. Ohmoto,
Biogeochemical cycling of iron in the Archean–Paleoproterozoic
Earth: constraints from iron isotope variations in sedimentary
rocks from the Kaapvaal and Pilbara Cratons, Chem. Geol. 218
(2005) 135–169.
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
[17] N. Dauphas, M. van Zuilen, M. Wadhwa, A.M. Davis, B. Marty,
P.E. Janney, Clues from Fe isotope variations on the origin
of Early Archean BIFs from Greenland, Science 306 (2004)
2077–2080.
[18] O. Rouxel, A. Bekker, K. Edwards, Iron isotope constraints on
the Archean and Paleoproterozoic ocean redox state, Science 307
(2005) 1088–1091.
[19] C.D. Frost, F. von Blanckenburg, R. Schoenberg, B.R. Frost, S.M.
Swapp, Preservation of Fe isotope heterogeneities during
diagenesis and metamorphism of banded iron-formation, Contrib.
Mineral. Petrol. 153 (2007) 211–235.
[20] B.L. Beard, C.M. Johnson, J.L. Skulan, K.H. Nealson, L. Cox, H.
Sun, Application of Fe isotopes to tracing the geochemical and
biological cycling of Fe, Chem. Geol. 195 (2003) 87–117.
[21] T.D. Bullen, A.F. White, C.W. Childs, D.V. Vivit, M.S. Schulz,
Demonstration of significant abiotic iron isotope fractionation,
Geology 29 (2001) 699–702.
[22] S.A. Welch, B.L. Beard, C.M. Johnson, P.S. Braterman, Kinetic
and equilibrium Fe isotope fractionation between aqueous Fe(II)
and Fe(III), Geochim. Cosmochim. Acta 67 (2003) 4231–4250.
[23] L.R. Croal, C.M. Johnson, B.L. Beard, D.K. Newman, Iron
isotope fractionation by Fe(II)-oxidizing photoautotrophic bacteria, Geochim. Cosmochim. Acta 68 (2004) 1227–1242.
[24] N. Balci, T.D. Bullen, K. Witte-Lien, W.C. Shanks, M. Motelica,
K.W. Mandernack, Iron isotope fractionation during microbially
stimulated Fe(II) oxidation and Fe(III) precipitation, Geochim.
Cosmochim. Acta 70 (2006) 622–639.
[25] R.K. Stevenson, J. David, Geochemistry and isotopic evolution
(Nd, Hf) of the 3.825 Ga Porpoise Cove sequence, Northeastern
Superior Province, Québec, GAC–MAC Annual Meeting,
Vancouver, Canada, 2003.
[26] W. Ciolkiewicz, J. Larocque, D. Francis, A habitable Earth at
3.8 Ga? Annual Undergraduate Research Conference, Faculty of
Science, McGill University.
[27] L.M. Anovitz, E.J. Essene, Phase equilibria in the system
CaCO3–MgCO3–FeCO3, J. Petrol. 28 (1987) 389–414.
[28] R.F. Dymek, C. Klein, Chemistry, petrology, and origin of
banded iron formation lithologies from the 3800 Ma Isua
Supracrustal Belt, West Greenland, Precambrian Res. 39 (1988)
247–302.
[29] N. Dauphas, P.E. Janney, R.A. Mendybaev, M. Wadhwa, F.M.
Richter, A.M. Davis, M. van Zuilen, R. Hines, C.N. Foley,
Chromatographic separation and multicollection—ICPMS analysis of iron. Investigating mass-dependent and -independent
isotope effects, Anal. Chem. 76 (2004) 5855–5863.
[30] N. Dauphas, O. Rouxel, Mass spectrometry and natural variations of iron isotopes, Mass Spectrom. Rev. 25 (2006) 515–550
(Erratum 25:831–832).
[31] P.D.P. Taylor, R. Maeck, P. De Bièvre, Determination of the
absolute isotopic composition and atomic weight of a reference
sample of natural iron, Int. J. Mass Spectrom. Ion Process. 121
(1992) 111–125.
[32] V. Busigny, N. Dauphas, Tracing paleofluid circulations using
iron isotopes: A study of hematite and goethite concretions from
the Navajo sandstone (Utah, USA). Earth Planet. Sci. Lett.
(in press).
[33] K. Govindaraju, 1994 compilation of working values and sample
description for 383 geostandard, Geostand. Newsl. 18 (1994)
1–158.
[34] F. Poitrasson, R. Freydier, Heavy iron isotope composition of
granites determined by high resolution MC-ICP-MS, Chem.
Geol. 222 (2005) 132–147.
375
[35] F. Poitrasson, A.N. Halliday, D.-C. Lee, S. Levasseur, N.
Teutsch, Iron isotope differences between Earth, Mars and
Vesta as possible records of contrasted accretion mechanisms,
Earth Planet. Sci. Lett. 223 (2004) 253–266.
[36] V.B. Polyakov, S.D. Mineev, The use of Mössbauer spectroscopy
in stable isotope geochemistry, Geochim. Cosmochim. Acta 64
(2000) 849–865.
[37] O. Rouxel, N. Dobbek, J. Ludden, Y. Fouquet, Iron isotope
fractionation during oceanic crust alteration, Chem. Geol. 202
(2003) 155–182.
[38] R.F. Dymek, C. Klein, Chemistry, petrology, and origin of
banded iron formation lithologies from the 3800 Ma Isua
Supracrustal Belt, West Greenland, Precambrian Res. 39 (1988)
247–302.
[39] V.R. McGregor, B. Mason, Petrogenesis and geochemistry of
metabasaltic and metasedimentary enclaves in the Amitsoq
gneisses, West Greenland, Am. Mineral. 62 (1977) 887–904.
[40] W.F. McDonough, S.-S. Sun, The composition of the Earth,
Chem. Geol. 120 (1995) 223–253.
[41] C. Klein, Some Precambrian banded iron-formations (BIFs) from
around the world: their age, geologic setting, mineralogy,
metamorphism, geochemistry, and origin, Am. Mineral. 90
(2005) 1473–1499.
[42] M.T. Rosing, N.M. Rose, D. Bridgwater, H.S. Thomsen, Earliest
part of Earth's stratigraphic record: a reappraisal of the N3.7 Ga
Isua (Greenland) supracrustal sequence, Geology 24 (1996)
43–46.
[43] N.M. Rose, M.T. Rosing, D. Bridgwater, The origin of
metacarbonate rocks in the Archean Isua supracrustal belt,
West Greenland, Am. J. Sci. 296 (1996) 1004–1044.
[44] M.A. van Zuilen, A. Lepland, J. Teranes, J. Finarelli, M. Wahlen,
G. Arrhenius, Graphite and carbonates in the 3.8 Ga old Isua
Supracrustal Belt, southern West Greenland, Precambrian Res.
126 (2003) 331–348.
[45] C.M. Johnson, E.E. Roden, S.A. Welch, B. Beard, Experimental
constraints on Fe isotope fractionation during magnetite and Fe
carbonate formation coupled to dissimilatory hydrous ferric oxide
reduction, Geochim. Cosmochim. Acta 69 (2005) 963–993.
[46] W.E. Ewers, Chemical factors in the deposition and diagenesis of
banded iron-formation, in: A.F. Trendall, R.C. Morris (Eds.),
Iron-Formation: Facts and Problems, Elsevier, Amsterdam, 1983,
pp. 491–512.
[47] K. Dideriksen, J.A. Baker, S.L.S. Stipp, Iron isotopes in natural
carbonate minerals determined by MC-ICP-MS with a 58Fe–54Fe
double spike, Geochim. Cosmochim. Acta 70 (2006) 118–132.
[48] G. Markl, F. von Blanckenburg, T. Wagner, Iron isotope
fractionation during hydrothermal ore deposition and alteration,
Geochim. Cosmochim. Acta 70 (2006) 3011–3030.
[49] R. Rye, P.H. Kuo, H.D. Holland, Atmospheric carbon dioxide
concentrations before 2 billion years ago, Nature 378 (1995)
603–605.
[50] H. Ohmoto, Y. Watanabe, K. Kumazawa, Evidence from massive
siderite beds for a CO2-rich atmosphere before ∼ 1.8 billion years
ago, Nature 429 (2004) 395–399.
[51] R.H. Becker, R.N. Clayton, Carbon isotopic evidence for the
origin of a banded iron-formation in Western Australia, Geochim.
Cosmochim. Acta 36 (1972) 577–595.
[52] J.F. Kasting, Archaean atmosphere and climate, Comment on «
H. Ohmoto, Y. Watanabe, K. Kumazawa, Evidence from massive
siderite beds for a CO2-rich atmosphere before ∼ 1.8 billion
years ago. Nature 429 (2004) 395–399 », Nature 432 (2004).
doi:10.1038/nature03166.
376
N. Dauphas et al. / Earth and Planetary Science Letters 254 (2007) 358–376
[53] A. Atkinson, M.L. O'Dwyer, R.I. Taylor, 55Fe diffusion in
magnetite crystals at 500 °C and its relevance to oxidation of
iron, J. Mater. Sci. 18 (1983) 2371–2379.
[54] B.L. Beard, C.M. Johnson, Fe isotope variations in the modern
and ancient Earth and other planetary bodies, Rev. Mineral. 55
(2004) 319–357.
[55] A. Ninh Pham, A.L. Rose, A.J. Feitz, T.D. Waite, Kinetics of Fe
(III) precipitation in aqueous solutions at pH 6.0–9.5 and 25 °C,
Geochim. Cosmochim. Acta 70 (2006) 640–650.
[56] J.C.G. Walker, Possible limits on the composition of the Archean
ocean, Nature 302 (1983) 518–520.
[57] H. Tamura, K. Goto, M. Nagayama, The effect of ferric
hydroxide on the oxygenation of ferrous ions in neutral solutions,
Corros. Sci. 16 (1976) 197–207.
[58] H. Tamura, S. Kawamura, M. Hagayama, Acceleration of the
oxidation of Fe2+ ions by Fe(III)-oxyhydroxides, Corros. Sci. 20
(1980) 963–971.
[59] W. Sung, J.J. Morgan, Kinetics and product of ferrous iron
oxygenation in aqueous systems, Environ. Sci. Technol. 14
(1980) 561–568.
[60] W. Davison, G. Seed, The kinetics of the oxidation of ferrous iron
in synthetic and natural waters, Geochim. Cosmochim. Acta 47
(1983) 67–79.
[61] J.L. Skulan, B.L. Beard, C.M. Johnson, Kinetic and equilibrium
Fe isotope fractionation between aqueous Fe(III) and hematite,
Geochim. Cosmochim. Acta 66 (2002) 2995–3015.
[62] G.A. Icopini, A.D. Anbar, S.S. Ruebush, M. Tien, S.L. Brantley,
Iron isotope fractionation during microbial reduction of iron: the
importance of adsorption, Geology 32 (2004) 205–208.
[63] N. Teutsch, U. von Gunten, D. Porcelli, O.A. Cirpka, A.N.
Halliday, Adsorption as a cause for iron isotope fractionation in
reduced groundwater, Geochim. Cosmochim. Acta 69 (2005)
4175–4185.
[64] H.A. Crosby, C.M. Johnson, E.E. Roden, B.L. Beard, Coupled
Fe(II)–Fe(III) electron and atom exchange as a mechanism for Fe
isotope fractionation during dissimilatory iron oxide reduction,
Environ. Sci. Technol. 39 (2005) 6698–6704.
[65] S.B. Jacobsen, M.R. Pimentel-Klose, A Nd isotopic study of the
Hamersley and Michipicoten banded iron formations: the source
of REE and Fe in Archean oceans, Earth Planet. Sci. Lett. 87
(1988) 29–44.
[66] M. Sharma, M. Polizzotto, A.D. Anbar, Iron isotopes in hot
springs along the Juan de Fuca Ridge, Earth Planet. Sci. Lett. 194
(2001) 39–51.
[67] S. Severmann, C.M. Johnson, B.L. Beard, C.R. German, H.N.
Edmonds, H. Chiba, D.R.H. Green, The effect of plume
processes on the Fe isotope composition of hydrothermally
derived Fe in the deep ocean as inferred from the Rainbow vent
site, Mid-Atlantic Ridge, 36°14′N, Earth Planet. Sci. Lett. 225
(2004) 63–76.
[68] B.L. Beard, C.M. Johnson, K.L. von Damm, R.L. Poulson, Iron
isotope constraints on Fe cycling and mass balance in oxygenated
Earth oceans, Geology 31 (2003) 629–632.
[69] X.K. Zhu, Y. Guo, R.J.P. Williams, R.K. O'Nions, A. Matthews,
N.S. Belshaw, G.W. Canters, E.C. de Waal, U. Weser, B.K.
Burgess, B. Salvato, Mass fractionation processes of transition
metal isotopes, Earth Planet. Sci. Lett. 200 (2002) 47–62.
[70] B.L. Beard, C.M. Johnson, Inter-mineral Fe isotope variations in
mantle-derived rocks and implications for the Fe geochemical
cycle, Geochim. Cosmochim. Acta 68 (2004) 47274743.
[71] H.M. Williams, A.H. Peslier, C. McCammon, A.N. Halliday, S.
Levasseur, N. Teutsch, J.-P. Burg, Systematic iron isotope
variations in mantle rocks and minerals: the effects of partial
melting and oxygen fugacity, Earth Planet. Sci. Lett. 235 (2005)
435–452.
[72] S. Weyer, A.D. Anbar, G.P. Brey, C. Münker, K. Mezger, A.B.
Woodland, Iron isotope fractionation during planetary differentiation, Earth Planet. Sci. Lett. 240 (2005) 251–264.
[73] K.D. Card, K.H. Poulsen, Geology and mineral deposits of the
Superior Province of the Canadian Shield, in: S.B. Lucas, M.R.
St-Onge (Eds.), Geology of the Precambrian Superior and
Grenville Provinces and Precambrian Fossils in North America,
Geology of Canada Series, vol. 7, Geological Survey of Canada,
1998, pp. 13–194.
[74] M. Simard, M. Parent, J. David, K.N.M. Sharma, Géologie de la
région de la rivière Innuksuac (SNRC 34K et 34L), vol. 2002-10,
Ministère des Ressources Naturelles, Québec, RG, 2003 43 pp.
[75] S.J. Sheather, M.C. Jones, A reliable data-based bandwidth
selection method for kernel density estimation, J. R. Stat. Soc., B
53 (1991) 683–690.
[76] A.W. Bowman, A. Azzalini, Applied Smoothing Techniques for
Data Analysis: the Kernal Approach With S-Plus Illustrations,
Oxford University Press, Oxford, 1997.