1. No category

Iron and carbon isotope evidence for microbial iron

Earth and Planetary Science Letters 303 (2011) 121–132
Contents lists available at ScienceDirect
Earth and Planetary Science Letters
j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / e p s l
Iron and carbon isotope evidence for microbial iron respiration throughout
the Archean
Paul R. Craddock ⁎, Nicolas Dauphas
Origins Laboratory, Department of the Geophysical Sciences and Enrico Fermi Institute, The University of Chicago, 5734 South Ellis Avenue, Chicago, IL 60637, United States
a r t i c l e
i n f o
Article history:
Received 17 August 2010
Received in revised form 20 December 2010
Accepted 22 December 2010
Available online 22 January 2011
Editor: R.W. Carlson
Keywords:
iron-formation
Hamersley
Isua
iron carbonates
iron respiration
a b s t r a c t
Banded Iron-Formations (BIFs) are voluminous chemical sediments that are rich in iron-oxide, carbonate and
silica and whose occurrence is unique to the Precambrian. Their preservation in the geological record offers
insights to the surface chemical and biological cycling of iron and carbon on early Earth. However, many details
regarding the role of microbial activity in BIF deposition and diagenesis are unresolved. Laboratory studies have
shown that reaction between carbon and iron through microbial iron respiration [2Fe2O3∙nH2O + CH2O + 7H+ →
4Fe2+ + HCO−
3 + (2n + 4)H2O + chemical energy] can impart fractionation to the isotopic compositions of these
elements. Here, we report iron (δ56Fe, vs. IRMM-014) and carbon isotopic (δ13C, vs. V-PDB) compositions of
magnetite and of iron-rich and iron-poor carbonates in BIFs from the late Archean (~2.5 Ga) Hamersley Basin,
Australia and the early Archean (~3.8 Ga) Isua Supracrustal Belt (ISB), Greenland. The range of δ56Fe values
measured in the Hamersley Basin, including light values in magnetite and heavy values in iron-rich carbonates
(up to +1.2‰), are incompatible with their precipitation in equilibrium with seawater. Rather, the data together
with previously reported light δ13C values in iron-rich carbonates record evidence for diagenetic reduction of
ferric oxide precursors to magnetite and carbonate through microbial iron respiration (i.e., dissimilatory iron
reduction, DIR). Iron and carbon isotope data of iron-rich metacarbonates from the ISB are similar to those of late
Archean BIFs. The isotopic signatures of these metacarbonates are supportive of an early diagenetic origin despite
metasomatic overprint, and preserve evidence of microbial iron respiration within the oldest recognized
sedimentary rocks on Earth.
© 2010 Elsevier B.V. All rights reserved.
1. Introduction
Banded Iron-Formations (BIFs) are conspicuously laminated marine
chemical sediments that are characterized by high concentrations of ironbearing minerals (20–40 wt.% bulk Fe) commonly interbedded with
layers of silica, and whose occurrence is unique to the Precambrian
(James, 1954, 1983). The mineralogy of the best-preserved BIFs consists of
combinations of four dominant facies: oxide (magnetite, hematite),
carbonate (siderite, ankerite, Fe–dolomite and, less commonly, calcite),
chert and silicate (stilpnomelane, riebeckite, greenalite, minnesotaite),
and locally sulfide (pyrite) and phosphate (apatite). Most known BIFs
have ages in the range ~3.8 to 1.8 Ga, but these formations also occur to a
lesser extent in the Neoproterozoic at ~700 Ma (Klein, 2005). The study of
these formations preserved in the rock record offers critical insights to
surface geochemical cycles and chemical evolution of the ocean and
atmosphere in the Precambrian, and in particular the Archean (≥2.5 Ga).
Despite significant scientific interest and research, however, there is
no consensus on the origin of BIFs. The primary mechanism for oxidation
of Fe(II)aq in an Archean ocean that was purportedly anoxic (Canfield
⁎ Corresponding author. Tel.: + 1 773 834 3997.
E-mail address: [email protected] (P.R. Craddock).
0012-821X/$ – see front matter © 2010 Elsevier B.V. All rights reserved.
doi:10.1016/j.epsl.2010.12.045
et al., 2000; Farquhar et al., 2000; Kasting, 1987; Ono et al., 2003; Pavlov
and Kasting, 2002), is uncertain. Photochemical oxidation of Fe(II)
in surface ocean waters owing to interaction with incident UV radiation
has been proposed as an entirely abiological means of accounting for
ferric iron in BIFs (Braterman et al., 1983; Cairns–Smith, 1978). Oxidation
of Fe(II) by O2 produced via photosynthesis has also been suggested
(Cloud, 1965, 1973), implying an indirect biological influence on BIF
formation and hinting at the presence of free O2 oases in the Archean
surface ocean. Alternatively, direct biological activity has been implicated,
via anoxygenic photosynthesis that coupled oxidation of Fe(II) to
reduction of inorganic carbon to yield organic compounds (Garrels
et al., 1973; Kappler et al., 2005; Konhauser et al., 2007; Widdel et al.,
1993). It is also uncertain the extent to which the mineral assemblages
preserved in BIFs reflect either primary precipitates from seawater,
possibly in near-chemical equilibrium with the ocean and atmosphere, or
are authigenic minerals formed during early sedimentary diagenesis and
burial metamorphism. For example, the mineralogical, chemical (e.g., rare
earth element) and isotopic (e.g., δ13C, δ18O) characteristics of iron-rich
carbonates such as siderite [FeCO3] and ankerite [Ca0.5(Fe,Mg)0.5CO3] in
BIFs have been used to argue either for primary precipitation from an
anoxic and stratified water column (Beukes et al., 1990; Kaufman et al.,
1990; Klein and Beukes, 1989) or for an authigenic origin (Becker and
Clayton, 1972; Heimann et al., 2010; Walker, 1984).
122
Samplea
Hole a
Macroband
Mesoband Type
Dales Gorge Member, Brockman Iron Formation
Drill core from Wittenoom Gorge Area (118° 28′ E; 22° 25′ S)
13324 M1
27
BIF 12
Magnetite
13327 WC2
27
BIF 11
Chert
13327 WC1
27
BIF 11
Chert
13321 HC
27
BIF 10
Chert + hematite
13318 M3
27
BIF 7
Magnetite
13318 Pl
27
BIF 7
Chert
13318 FM1
27
BIF 7
Chert
13318 M1
27
BIF 7
Magnetite
13318 M2
27
BIF 7
Magnetite
13316 M1
27
BIF 6
Magnetite
13316 CA1
27
BIF 6
Fine−band combination
13316 CB1
27
BIF 6
Coarse-band combination
13309 FM1
27
BIF 1
Chert
13309 PC1
27
BIF 1
Chert
13309 QIO1
27
BIF 1
Chert-matrix
13309 M2
27
BIF 1
Magnetite
13326 M1
28
BIF 12
Magnetite
13328 WC2
28
BIF 11
Chert
13328 WC1
28
BIF 11
Chert
13322 WC1
28
BIF 10
Chert
13322 WCIA
28
BIF 10
Chert
13322 M1
28
BIF 10
Magnetite
13322 HC
28
BIF 10
Chert + hematite
13322 WHC
28
BIF 10
Chert + hematite
13322 FM1
28
BIF 10
Chert
13322 QIO1
28
BIF 10
Chert-matrix
13319 FM3
28
BIF 7
Chert
13319 FM2
28
BIF 7
Chert
13319 M1
28
BIF 7
Magnetite
13319 FM1
28
BIF 7
Chert
13313 M1
28
BIF 2
Magnetite
13310 PC1
28
BIF 1
Chert
M2
40
BIF 2
Magnetite
P2
40
BIF 2
Chert
P 2A
40
BIF 2
Chert
Depthb
(m)
85.2
98.4
98.4
99.5
122.2
122.2
122.2
122.3
122.3
128.1
128.2
128.2
162.2
162.3
162.4
162.4
110.9
114.9
115.0
125.2
125.2
125.2
125.2
125.2
125.3
125.3
146.5
146.6
146.6
146.6
174.5
186.7
196.8
196.8
196.8
δ13C
(‰)c
δ56Fe
(‰) d
Ankerite
Siderite
Ankerite/Siderite
− 11.41
− 15.05
− 6.98
− 9.24 ± 0.03
− 9.83 ± 0.04
− 9.72 ± 0.06
− 11.08
− 0.289 ± 0.031
0.013 ± 0.031
− 0.705 ± 0.031
− 9.01
− 9.87
− 10.06
− 7.45
− 0.584 ± 0.029
0.118 ± 0.031
− 9.31
− 9.41
− 11.07
− 12.86
− 9.73
− 9.61
− 12.48
− 6.76
− 6.50 ± 0.03
− 8.96
− 8.26
− 7.80
− 9.79
− 9.74 ± 0.03
− 10.18 ± 0.01
− 9.90
0.641 ± 0.031
0.178 ± 0.033
0.057 ± 0.031
− 0.161 ± 0.031
Magnetite
Hematite
0.424 ± 0.031
0.205 ± 0.031
0.465 ± 0.031
0.908 ± 0.038
0.006 ± 0.036
− 0.210 ± 0.034
0.159 ± 0.029
0.112 ± 0.031
0.650 ± 0.038
0.636 ± 0.033
0.196 ± 0.031
0.588 ± 0.031
0.490 ± 0.030
0.675 ± 0.031
0.488 ± 0.030
0.418 ± 0.031
− 0.555 ± 0.031
Ankerite/Siderite
− 0.442 ± 0.049
0.008 ± 0.046
− 1.023 ± 0.083
0.373 ± 0.029
− 0.865 ± 0.051
0.198 ± 0.046
0.933 ± 0.083
0.283 ± 0.060
0.070 ± 0.046
− 0.236 ± 0.043
0.790 ± 0.034
0.453 ± 0.035
0.294 ± 0.027
− 0.792 ± 0.038
− 1.081 ± 0.033
0.787 ± 0.033
0.566 ± 0.038
− 0.318 ± 0.029
0.086 ± 0.030
0.075 ± 0.030
Magnetite
− 0.298 ± 0.054
0.230 ± 0.051
0.179 ± 0.046
0.949 ± 0.058
0.914 ± 0.060
0.244 ± 0.049
0.852 ± 0.043
0.710 ± 0.046
0.990 ± 0.046
0.694 ± 0.044
0.610 ± 0.046
0.630 ± 0.044
0.662 ± 0.047
0.419 ± 0.040
0.730 ± 0.033
0.398 ± 0.030
1.194 ± 0.033
0.998 ± 0.038
− 0.287 ± 0.030
0.064 ± 0.030
− 0.022 ± 0.030
− 0.149 ± 0.029
1.085 ± 0.029
0.159 ± 0.035
0.551 ± 0.035
− 1.142 ± 0.044
− 1.601 ± 0.060
1.158 ± 0.044
0.847 ± 0.058
− 0.470 ± 0.043
0.156 ± 0.044
0.093 ± 0.044
− 0.303 ± 0.047
Hematite
0.630 ± 0.043
0.316 ± 0.049
0.699 ± 0.043
1.369 ± 0.044
0.005 ± 0.057
− 0.757 ± 0.049
0.426 ± 0.038
− 0.181 ± 0.028
− 9.70 ± 0.02
− 9.80 ± 0.06
δ57Fe
(‰)d
1.107 ± 0.060
0.584 ± 0.057
1.778 ± 0.044
1.477 ± 0.058
− 0.422 ± 0.044
0.077 ± 0.044
0.014 ± 0.044
− 0.212 ± 0.043
1.582 ± 0.043
0.241 ± 0.047
0.810 ± 0.047
0.561 ± 0.051
1.144 ± 0.054
P.R. Craddock, N. Dauphas / Earth and Planetary Science Letters 303 (2011) 121–132
Table 1
Iron and carbon isotopic compositions of oxides and carbonate from drill core samples of the Brockman Iron Formation, Hamersley Basin.
Table 1 (continued)
Samplea
Hole a
Macroband
Mesoband Type
Depthb
(m)
δ13C
(‰)c
Ankerite
a
196.8
104.5
155.4
155.6
155.6
− 10.28 ± 0.02
66.2
66.2
66.3
66.3
66.4
70.7
70.8
78.2
98.4
104.4
104.4
104.5
125.3
125.3
125.3
125.3
125.3
125.5
137.6
137.6
− 10.95 ± 0.02
Siderite
Ankerite/Siderite
δ57Fe
(‰)d
Magnetite
Hematite
Ankerite/Siderite
0.314 ± 0.028
− 0.189 ± 0.028
− 0.092 ± 0.035
Magnetite
Hematite
0.478 ± 0.047
− 0.282 ± 0.047
− 0.111 ± 0.047
− 8.65
− 0.305 ± 0.041
− 10.96
− 10.84 ± 0.01
0.301 ± 0.033
− 0.404 ± 0.029
− 12.59 ± 0.50
− 11.11
− 9.53
− 9.36
− 9.06
0.244 ± 0.033
− 9.34
− 9.31
0.013 ± 0.033
0.605 ± 0.038
0.456 ± 0.038
0.202 ± 0.029
0.351 ± 0.029
0.344 ± 0.051
0.297 ± 0.029
0.194 ± 0.029
− 0.005 ± 0.031
0.408 ± 0.033
0.727 ± 0.034
− 0.156 ± 0.029
− 0.448 ± 0.072
0.410 ± 0.083
− 0.588 ± 0.043
0.674 ± 0.051
0.395 ± 0.083
− 0.029 ± 0.083
0.870 ± 0.058
0.710 ± 0.044
0.665 ± 0.085
0.735 ± 0.029
0.877 ± 0.028
0.298 ± 0.041
0.515 ± 0.041
0.519 ± 0.133
0.433 ± 0.041
0.282 ± 0.133
− 0.017 ± 0.083
0.632 ± 0.083
1.029 ± 0.054
− 0.253 ± 0.041
0.887 ± 0.058
1.073 ± 0.041
1.283 ± 0.047
1.055 ± 0.033
− 9.48
− 9.75
1.212 ± 0.041
− 0.033 ± 0.029
0.096 ± 0.035
0.886 ± 0.041
0.963 ± 0.028
1.041 ± 0.041
0.600 ± 0.027
1.057 ± 0.133
1.568 ± 0.083
1.850 ± 0.072
− 0.053 ± 0.041
0.135 ± 0.047
1.301 ± 0.072
1.419 ± 0.047
1.549 ± 0.072
0.868 ± 0.040
Sample numbers follow the nomenclature adopted by Becker (1971). Core samples were recovered by drilling operations in two areas: Wittenoom Gorge and Yampire Gorge. Hole designates core number.
Depth in meters below modern surface.
c
Carbon isotopic (δ13C) ratios are per mil difference relative to PDB standard and are those published by Becker and Clayton (1972). Quoted uncertainties are the standard deviation of replicate analyses as reported by Becker and Clayton
(1972).
d
Iron isotopic (δ56Fe, δ57Fe) ratios are reported in per mil difference relative to iron metal standard IRMM-014 (δ57Fe ~ 1.5 x δ56Fe). Uncertainties are 95% confidence intervals.
b
P.R. Craddock, N. Dauphas / Earth and Planetary Science Letters 303 (2011) 121–132
Dales Gorge Member, Brockman Iron Formation
Drill core from Wittenoom Gorge Area (118° 28′ E; 22° 25′ S)
FM 2
40
BIF 2
Chert
M3
51
BIF 12
Magnetite
QIO 51
51
BIF 5
Chert-matrix
M1
51
BIF 5
Magnetite
Pl
51
BIF 5
Chert
Drill core from Yampire Gorge Area (118° 37′ E; 22° 27′ S)
13325 WC2
Y3
BIF 12
Chert
13325 M2
Y3
BIF 12
Magnetite
13325 M1
Y3
BIF 12
Magnetite
13325 FM1
Y3
BIF 12
Chert
13325 WC1
Y3
BIF 12
Chert
13329 M1
Y3
BIF 11
Magnetite
13329 FM1
Y3
BIF 11
Chert
13323 M1
Y3
BIF 10
Magnetite
13320 FM1
Y3
BIF 7
Chert
13317 CA1
Y3
BIF 6
Fine-band combination
13317 CBI
Y3
BIF 6
Coarse-band combination
13317 M1
Y3
BIF 6
Magnetite
13314 M3
Y3
BIF 3
Magnetite
13314 M2
Y3
BIF 3
Magnetite
13314 H1
Y3
BIF 3
Hematite
13314 M1
Y3
BIF 3
Magnetite
13314 FM1
Y3
BIF 3
Chert
13314 Carb 1
Y3
BIF 3
Chert + carbonate
13311 PC1
Y3
BIF 1
Chert
13311 PC1
Y3
BIF 1
Chert
δ56Fe
(‰) d
123
124
P.R. Craddock, N. Dauphas / Earth and Planetary Science Letters 303 (2011) 121–132
In order for BIFs to offer a faithful and critical perspective of the
Earth's early ocean and atmosphere biogeochemical cycles, an understanding of processes leading to their formation [i.e., primary Fe(II)
oxidation and precipitation versus diagenetic transformations] is
essential. Here, we report coupled iron, δ56Fe= [(56Fe/54Fe)sample/
(56Fe/54Fe)IRMM-014 − 1]× 103, and carbon, δ13C = [(13C/12C)sample/(13C/
12
C)V-PDB − 1] × 103, isotope compositions of iron oxides and carbonates
from one of the most well-preserved and least metamorphosed BIFs, the
~2.5 Ga Brockman Iron Formation of the Hamersley Basin, western
Australia, and in one of the oldest known Fe-rich chemical sedimentary
formations, that of the ~3.8 Ga Isua Supracrustal Belt (ISB), southern West
Greenland. Our data are used to constrain the probable chemical and
biological pathways for the precipitation of iron-bearing minerals in
Archean BIFs. A primary motivation for this research is the study of Becker
and Clayton (1972) that reported δ13C of iron-rich carbonates (siderite,
ankerite) in the Brockman Iron Formation, Hamersley Basin, mostly
between −8 and −11‰, significantly more depleted than that of typical
iron-poor marine carbonates (calcite, dolomite) with δ13C between −2
and +2‰. The source of light carbon in the iron-formation was suggested
to derive from organic carbon, possibly resulting from oxidation–
reduction reactions involving microbially metabolized Fe(III). We
examine specifically whether the range of iron and carbon isotope ratios
recorded in iron-rich carbonates and magnetite from the Brockman Iron
Formation is consistent with such a process and thus records evidence for
microbial iron respiration (i.e., dissimilatory iron reduction; DIR) in the
late Archean, a process that was originally suggested by Walker (1984)
and Baur et al. (1985). Johnson et al. (2008) have recently reported the
iron isotope compositions of magnetite and siderite from the Brockman
Iron Formation in order to distinguish between primary precipitation in
seawater and diagenetic iron transformations. These authors suggested
that the range of measured iron isotopic compositions could be consistent
with an authigenic origin, but they did not have access to carbon isotopic
data for the same carbonates to validate a DIR pathway. Complementary
iron and carbon isotopic studies of magnetite and carbonates from the 2.5
to 2.4 Ga Kuruman Iron Formation, Transvaal Craton, South Africa, have
also explored evidence for microbial iron respiration during this period
of Earth's history (Heimann et al., 2010; Johnson et al., 2003). Heimann
et al. (2010) concluded that the range of δ13C and δ56Fe values
documented in iron-rich carbonates was consistent with authigenic
pathways for carbonate formation through microbial DIR.
Further, we assess whether a record of microbial iron respiration
extends to the early Archean. Dauphas et al. (2007) have reported iron
isotopic compositions of iron-rich and iron-poor metacarbonates from
the ~3.8 Ga ISB. Metacarbonates from the ISB typically show a
metasomatic relationship with their host rocks (Rose et al., 1996; Rosing
et al., 1996). Rose et al. (1996) and Rosing et al. (1996) have argued that
metacarbonates were formed by metasomatic alteration of igneous
protoliths. However, the iron isotopic and trace element (e.g., REE, Fe/Ti)
characteristics of iron-rich metacarbonates are more similar to that of coexisting BIFs in the ISB and are more supportive of a chemical sedimentary
origin (Bolhar et al., 2004; Dauphas et al., 2007). Here, we integrate new
carbon and existing iron isotopic data of iron-rich and iron-poor
metacarbonates from the ISB to re-examine the origin of these rocks, in
particular evaluating whether these isotopic signatures are supportive of
a sedimentary origin prior to metasomatic overprint and record evidence
of microbial iron respiration within the oldest recognized sedimentary
rocks on Earth.
2. Materials and methods
2.1. Geology and sample descriptions
The geological setting, description and analytical procedures for
isotopic measurement of iron-formations in this study are described
in detail in the Supporting Online Material, and summarized below.
The Hamersley Basin, western Australia is a roughly elliptical basin in
which late Archean and Paleoproterozoic (~2.6 to 2.45 Ga) volcanics
and sediments of the Hamersley Group were deposited (Trendall and
Blockley, 1970). The Hamersley Group is of particular significance
owing to the presence of four major and several minor conspicuously
laminated, iron-rich stratigraphic units (Banded Iron Formation, BIF),
including the Brockman Iron Formation (MacLeod, 1966; Trendall,
2002; Trendall and Blockley, 1970). The Hamersley Group also
contains the Wittenoom Dolomite, which comprises between 200
and 300 m of interbedded carbonate, chert and shale with the
carbonate occurring typically as mostly massive calcite and dolomite.
Samples from the Hamersley Basin were selected both from an ironformation (Dales Gorge Member of the Brockman Iron Formation) and
from a carbonate formation (Wittenoom Dolomite). Samples of ironformation included iron oxide (magnetite) and iron-rich carbonates
(siderite, ankerite) and were from fresh drill core material that was
recovered by the Australian Blue Asbestos Co. as part of a prospecting
program under subsidy from the Western Australian Government
(Trendall and Blockley, 1970). Samples of carbonate formation were
hand specimens of massive, iron-poor carbonate (calcite, dolomite)
obtained from chipping out fresh surfaces of the Wittenoom Dolomite
from weathered exposures. All samples were originally collected and
prepared for carbon and oxygen isotope studies by Becker and Clayton
(1972,1976).
The Isua Supracrustal Belt (ISB) is part of the ~3.8 Ga high-grade
Itsaq Gneiss Complex and is one of the oldest metasedimentarybearing formations on Earth. The ISB assemblage comprises a variety
of metavolcanic, clastic and metasedimentary rock types, the
stratigraphy of which has been discussed in detail by previous
contributions (Boak and Dymek, 1982; Dymek and Klein, 1988;
Myers, 2001; Nutman and Friend, 2009; Nutman et al., 1984, 1996;
Rosing et al., 1996; van Zuilen et al., 2003). Protolith identification of
rocks in the ISB is complicated owing to the diversity of lithological
units observed in the ISB and metamorphic and/or metasomatic
overprint. For this study, metacarbonate sequences were targeted in
order to better identify their probable origin (chemical sedimentary
versus metasomatic) and possible association with genuine ironformations. Iron-poor metacarbonates occur typically at the contacts
with ultramafic intrusions in the southern ISB and have been
interpreted as originating by leaching of ultramafic protoliths by
metasomatic fluids (Rose et al., 1996; Rosing et al., 1996). Iron-rich
metacarbonates occur within a range of host rocks in the northern ISB,
including in association with banded magnetite–quartz rocks of a
chemical sedimentary origin. We report and contrast here the carbon
isotopic ratios of iron-rich and iron-poor metacarbonates and
combine these new data with existing iron isotopic data for the
same samples (Dauphas et al., 2007) to better constrain their primary
formation, in particular identifying whether the isotopic signatures
record a chemical sedimentary origin despite metasomatic overprint.
2.2. Sample preparation and analytical methods
Samples of magnetite and iron-rich carbonate from the Brockman
Iron Formation were originally taken from drill core material and
separated and crushed by Richard Becker to obtain milligram-sized
monomineralic powders for the analysis of carbon and oxygen
isotopes (Becker and Clayton, 1972, 1976). In addition, samples of
iron-poor carbonate from the Wittenoom Dolomite were cut from
unweathered faces of hand specimens and approximately one gram of
this material was crushed to a powder using an agate mortar and
pestle. Magnetite and iron-rich and iron-poor carbonates from the ISB
were taken from powders of hand specimens originally prepared for
iron isotopic analysis (see Dauphas et al., 2007).
Iron isotopic analyses of samples from the Brockman Iron Formation
and Wittenoom Dolomite were carried out following the procedures
developed in our laboratory (Craddock and Dauphas, 2010; Dauphas et
al., 2004a, 2009b). Iron isotopic measurements were performed on a
P.R. Craddock, N. Dauphas / Earth and Planetary Science Letters 303 (2011) 121–132
frequency
15
seawater
Fe2+aq
Fe-rich
carbonate
magnetite
10
Carbon isotopic analyses were carried out on iron-rich and ironpoor metacarbonates from the ISB at RSMAS, University of Miami. The
procedures followed those developed by Swart et al. (1991). Carbon
isotope compositions are reported relative to the V-PDB scale.
seawater
HCO3 , aq
Hamersley
~2.5 Ga
Fe-rich
carbonate
Fe-poor
carbonate
3. Results
Fe-poor
5 carbonate
0
Kuruman
~2.5 Ga
frequency
15
10
Fe-poor
carbonate
Fe-rich
carbonate
5
Fe-rich
carbonate
magnetite
Fe-poor
carbonate
0
-2.0 -1.5 -1.0 -0.5 0.0 0.5 1.0
δ56Fe,
-14 -12 -10 -8 -6 -4 -2 0
δ13C,
IRMM-014 (‰)
56
125
2
V-PDB (‰)
13
Fig. 1. Histograms of iron (δ Fe) and carbon (δ C) isotope ratios in iron-rich carbonates
(siderite, ankerite), iron-poor carbonates (calcite, dolomite) and magnetite from the ~2.5 Ga
Brockman Iron Formation, Hamersley Basin (top panel) and Kuruman Iron Formation,
Transvaal Craton (bottom panel). Iron isotope data for the Brockman Iron Formation are from
Table 1. Carbon isotope data for the Brockman Iron Formation are from Becker and Clayton
(1972). Carbon and iron isotope data for the Kuruman Iron Formation are from Johnson et al.
(2003) and Heimann et al. (2010). Bin widths for iron and carbon isotope values are 0.025‰
and 1.0‰, respectively. The vertical arrows at δ56Fe=−0.2‰ and δ13C=0‰ are the
estimated compositions of Fe(II)aq and dissolved inorganic carbon (DIC) in Archean seawater,
respectively. The range of iron and carbon isotopic ratios in these samples are incompatible
with formation of magnetite and iron-rich carbonates as direct precipitates from seawater,
but can be reconciled by a model invoking diagenetic mobilization of primary ferric oxides
and organic carbon through dissimilatory iron reduction (DIR).
Thermo Scientific Neptune MC-ICPMS at the University of Chicago. All
iron isotope data are calibrated relative to the IRMM scale (Craddock
and Dauphas, 2010; Taylor et al., 1992).
Iron isotopic (δ56Fe) compositions of magnetite samples from the
Brockman Iron Formation, Hamersley Basin, analyzed in this study range
from −0.3 to +1.2‰ (Table 1; Fig. 1). These data are similar to those
measured in magnetite samples from the same formation by Johnson
et al. (2008). In the two studies combined, magnetite from the Brockman
Iron Formation extends down to δ56Fe=−1.0‰ and has a mean iron
isotopic composition of +0.17‰ (n =94). Two types of carbonates from
the Hamersley Basin were studied: iron-rich carbonates (siderite,
ankerite) from the Brockman Iron Formation and iron-poor carbonates
(calcite, dolomite) from the underlying Wittenoom Dolomite [using
notations from van Zuilen et al. (2003) and Dauphas et al. (2007), ironrich carbonates have an Fe atomic ratio, Fe/(Fe+Mn +Ca+Mg), greater
than 0.40, whereas iron-poor carbonates have an Fe atomic ratio less than
0.15]. The δ56Fe values of iron-rich carbonates analyzed in this study range
from −1.1 to +1.2‰ (Table 1; Fig. 1). Considering the iron isotopic data
for siderite published by Johnson et al. (2008), the range extends down to
−2.0‰. The iron isotopic compositions of magnetite and iron-rich
carbonates sampled from adjacent bands in the Brockman Iron Formation
have a broadly positive correlation. The δ56Fe values of iron-poor
carbonates from the Wittenoom Dolomite analyzed in this study are all
very negative, between −0.5 and −1.0‰ (Table 2; Fig. 1). Carbon isotope
compositions (δ13C) of the carbonate samples analyzed in this study for
their iron isotopic compositions were previously reported by Becker and
Clayton (1972). All iron-rich carbonates have light δ13C compositions
(Fig. 1). Siderite δ13C range from −12.6 to −7.5‰ and cluster around the
average of −9.6‰ (n=11). Ankerite δ13C are slightly more variable,
ranging from −15.0 to −6.5‰, but cluster around a similar average of
−9.9‰ (n=32). Iron-poor carbonates from the Wittenoom Dolomite
have a limited range of heavier δ13C with an average −2‰ (n =15;
Fig. 1). For comparison, the iron and carbon isotope compositions of ironrich and iron-poor carbonates and of magnetite from the penecontemporaneous Kuruman Iron Formation, Transvaal Supergroup, South Africa
(Heimann et al., 2010; Johnson et al., 2003) are illustrated (Fig. 1). The
Table 2
Iron and carbon isotopic compositions of iron-poor carbonate hand specimens from the Wittenoom Dolomite, Hamersley Basin.
Samplea
Lithology
δ13C
(‰)b
Calcite
Wittenoom Dolomite Formation
12523
Dolomite
12523
vein Dolomite vein
13354
Calcite + quartz
13355 top
Calcite ± dolomite
13355 bottom
13356
Calcite
13357 chert
Calcite + quartz
13358 #1
Calcite
13358 #2
13359 chert
Black chert + calcite
13360
Calcite
RB-22a*
Calcite + silicate
RB-22b*
RB-23a*
Calcite + dolomite+
RB-23b*
silicate
a
δ56Fe
(‰)c
Dolomite
Calcite
− 1.06 ± 0.05
− 0.82 ± 0.06
− 4.91 ± 0.10
− 4.68 ± 0.02
− 6.36 ± 0.04
− 4.75 ± 0.03
− 1.56
− 1.84
Dolomite
Calcite
− 0.995 ± 0.068
− 0.774 ± 0.068
0.05 ± 0.01
0.38 ± 0.02
− 1.09 ± 0.02
− 1.80 ± 0.02
− 1.83
− 0.13 ± 0.05
− 1.22
− 0.30 ± 0.05
δ57Fe
(‰)c
Dolomite
− 1.427 ± 0.092
− 1.135 ± 0.092
− 0.728 ± 0.055
− 0.472 ± 0.055
− 1.059 ± 0.038
− 0.674 ± 0.038
− 0.921 ± 0.055
− 0.860 ± 0.055
− 0.882 ± 0.055
− 0.450 ± 0.055
− 0.887 ± 0.057
− 0.685 ± 0.057
− 1.313 ± 0.038
− 1.344 ± 0.038
− 1.338 ± 0.038
− 0.692 ± 0.038
− 1.366 ± 0.085
− 1.024 ± 0.085
− 5.67 ± 0.03
Sample numbers follow the nomenclature adopted by Becker (1971). Detailed descriptions given by Becker (1971).
Carbon isotopic (δ13C) ratios are per mil difference relative to PDB standard and are those published by Becker and Clayton (1972). Quoted uncertainties are the standard
deviation of replicate analyses as reported by Becker and Clayton (1972).
c
Iron isotopic (δ56Fe, δ57Fe) ratios are reported in per mil difference relative to iron metal standard IRMM-014 (δ57Fe~1.5 x δ56Fe). Uncertainties are 95% confidence intervals. * Indicates
iron-poor carbonates sampled from the base of the Wittenoom Dolomite that occur as fine-layers in apparently microbanded rocks containing silica and silicate and that are distinct in
appearance to more massive calcite and dolomite sampled elsewhere in the Wittenoom Dolomite (see Becker, 1971). These samples were not available for the analysis of iron isotopes.
b
126
P.R. Craddock, N. Dauphas / Earth and Planetary Science Letters 303 (2011) 121–132
Table 3
Iron and carbon isotopic compositions of bulk metacarbonates from Isua Supracrsutal Belt, southern West Greeland.
Sample
Description
Fe (wt.%)
δ13C (‰)
a
Fe-rich metacarbonates
IS-04-01
IS-04-02
IS-04-03
IS-04-07
IS-04-07*
IS-04-08
IS-04-09
IS-04-10
IS-04-11
Siderite, northwest of belt
Siderite, northwest of belt
Siderite, northwest of belt
Metacarbonate mix, east of belt
Replicate, this study
Siderite, east of belt
Graphite-rich metacarbonate
Siderite, east of belt
Graphite-rich metacarbonate
23.6
23.7
24.7
59.4
58.1
31.1
47.3
49.6
45.7
− 5.50 ± 0.06
− 4.52 ± 0.05
− 4.73 ± 0.02
− 4.10
Fe-poor metacarbonates
IS-04-05
IS-04-05*
AL-04-G18
Al-04-G23
Calcite/dolomite, west of belt
Replicate, this study
Carbonated Amitsoq gneiss
Carbonated Amitsoq gneiss
2.6
2.9
3.2
2.5
− 1.98 ± 0.07
− 4.52 ± 0.06
− 4.75 ± 0.05
− 5.94 ± 0.05
− 4.58 ± 0.04
− 1.52 ± 0.04
− 0.34 ± 0.06
δ56Fe (‰)
b
δ57Fe (‰)
b
0.236 ± 0.059
0.318 ± 0.125
0.448 ± 0.074
0.759 ± 0.038
0.743 ± 0.039
0.400 ± 0.034
0.294 ± 0.069
0.281 ± 0.084
0.204 ± 0.056
0.296 ± 0.093
0.500 ± 0.169
0.763 ± 0.151
1.129 ± 0.074
1.098 ± 0.045
0.587 ± 0.048
0.496 ± 0.165
0.400 ± 0.110
0.383 ± 0.114
− 0.742 ± 0.046
− 0.754 ± 0.037
− 0.901 ± 0.177
− 0.725 ± 0.177
− 1.096 ± 0.093
− 1.138 ± 0.043
− 1.456 ± 0.219
− 1.051 ± 0.219
a
Carbon isotopic (δ13C) ratios are reported in per mil difference relative to the V-PDB standard. Uncertainties are two standard deviation for replicate analyses of the same sample.
Iron isotopic (δ56Fe, δ57Fe) ratios are those previously published by Dauphas et al. (2007b), except for two samples IS-04-07 and IS-04-05 indicated by *, which were
independently measured for their iron isotopic composition as part of the present study for confirmation of accuracy. Iron isotopic ratios are reported in per mil difference relative to
iron metal standard IRMM-014. Quoted uncertainties are 95% confidence intervals.
b
data show a distribution of iron and carbon isotopic ratios that is similar to
that documented in the Hamersley Basin.
To compare the isotopic signatures of metacarbonates from the
~ 3.8 Ga Isua Supracrustal Belt (ISB) with those from genuine late
Archean BIFs, we combine new carbon isotope data of iron-rich and
iron-poor metacarbonates measured in this study with previously
reported iron isotope data for the same metacarbonates from Dauphas
et al. (2007). Iron-poor metacarbonates have light δ56Fe (− 0.90 to
− 0.73‰), whereas iron-rich metacarbonates have heavy δ56Fe
(+0.24 to +0.76‰; Table 3, Fig. 2). Iron-poor metacarbonates have
δ13C values between −2.0 and 0‰, whereas iron-rich metacarbonates
have distinctly lighter δ13C values between − 6.0 to − 4.1‰ (Table 3;
Fig. 2). Several studies have previously reported δ13C values of
metacarbonates from the ISB (Oehler and Smith, 1977; Perry and
Ahmad, 1977; Schidlowski et al., 1979; Ueno et al., 2002; van Zuilen
et al., 2003); only van Zuilen et al. (2003) have distinguished
geochemically between the carbon isotopic signatures of iron-poor
and iron-rich variants and none have iron isotopic data for the same
samples that enable a direct comparison with our study.
frequency
10
seawater
Fe2+aq
seawater
HCO3 , aq
Isua
~3.8 Ga
Fe-rich
metacarbonate
Fe-poor
metacarbonate
5
Fe-poor
metacarbonate
magnetite
0
-2.0 -1.5 -1.0 -0.5 0.0 0.5 1.0
δ56Fe,
IRMM-014 (‰)
Fe-rich
metacarbonate
-14 -12 -10 -8 -6 -4 -2 0
δ13C,
2
V-PDB (‰)
Fig. 2. Histograms of iron (δ56Fe) and carbon (δ13C) isotope ratios in iron-rich
metacarbonates (siderite, ankerite), iron-poor metacarbonates (calcite, dolomite) and
magnetite from the ~3.7 to 3.8 Ga Isua Supracrustal Belt, SW Greenland. Iron isotopic data
are from Dauphas et al. (2007). Bin widths for iron and carbon isotope values are 0.025‰ and
1.0‰, respectively. The vertical arrows are the same as those depicted in Fig. 1. Despite a
complex metamorphic history and metasomatic overprint, the iron and carbon isotopic data
are similar to those of genuine late-Archean BIFs and are compatible with derivation of ironrich metacarbonates in the ISB from reduction of ferric oxides (ferrihydrite), probably
coupled to oxidation of organic carbon.
4. Discussion
4.1. Evidence for an authigenic origin of iron-rich carbonates and
sedimentary microbial iron respiration at 2.5 Ga
Iron-rich carbonates in BIFs from the Hamersley Basin exhibit a
wide range of iron and carbon isotopic compositions. Can this isotopic
heterogeneity be explained entirely by precipitation of dissolved Fe
(II) and inorganic carbon in the water column in near-isotopic
equilibrium? Or, are the isotopic signatures consistent with an
authigenic origin in marine sediments, possibly involving microbial
iron and carbon respiration?
The δ13C of iron-poor carbonates of different ages ranging from
~3.8 to 2.5 Ga are remarkably uniform and similar to that of platform
carbonates deposited throughout the Phanerozoic (~ 0‰) (Becker and
Clayton, 1972; Schidlowski et al., 1975; Shields and Veizer, 2002;
Veizer et al., 1989, 1992). This argues for no significant changes (i.e.,
b10‰) in the bulk carbon isotopic reservoir of dissolved inorganic
carbon (DIC) in seawater (c.f. Beukes et al., 1990; Kaufman et al.,
1990). The uniform δ13C of iron-poor carbonates deposited synchronously in the late Archean across a range of water column depths from
continental shelves (platform) to abyssal plains (basinal) argues also
against vertical stratification with respect to carbon isotopes of DIC at
any given time (Fischer et al., 2009). Based on the experimentally
determined carbon isotopic fractionation between siderite and DIC,
Δ13CHCO3-Sid = −0.5 ± 0.2‰ (Jimenez-Lopez and Romanek, 2004),
siderite precipitated from the water column in isotopic equilibrium
with DIC should have near-uniform δ13C ~0 ± 2‰ (Fig. 3). The range
of δ13C between −6 and −15‰ documented in iron-rich carbonates
from the Brockman Iron Formation in the Hamersley Basin is clearly
different from that expected for, and argues against, formation of
these carbonates in isotopic equilibrium with Archean seawater (Baur
et al., 1985; Becker and Clayton, 1972).
Evidence from iron isotope data of the same iron-rich carbonates
further argues against primary precipitation in Archean seawater. The
prevailing consensus is that iron was delivered to the Archean ocean as
ferrous Fe(II)aq, primarily via hydrothermal activity (e.g., Bau and Möller,
1993; Dymek and Klein, 1988; Jacobsen and Pimental-Klose, 1988). The
δ56FeFe(II) composition of modern hydrothermal fluids ranges from −0.1
to −0.6‰, averaging −0.20‰ (n=19) (Beard et al., 2003; Severmann
et al., 2004; Sharma et al., 2001). Given that the range of iron isotopic
composition of Archean igneous rocks is limited and similar to modern
(e.g., Dauphas et al., 2009a), hydrothermal alteration of oceanic crust
in the Archean should contribute Fe(II) with an isotopic composition also
P.R. Craddock, N. Dauphas / Earth and Planetary Science Letters 303 (2011) 121–132
1.5
~2.5 Ga
carbonate δ56Fe (‰)
1.0
Kuruman
Fe-rich
0.5
0.0
-0.5
Brockman
Fe-rich
-1.0
Fe-poor carbonates
1.5
~3.8 Ga
carbonate δ56Fe (‰)
1.0
Isua, Fe-rich metacarbonates
0.5
seawater
0.0
-0.5
Isua, Fe-poor metacarbonates
-1.0
-1.5
-16
-14
-12
-10
-8
-6
-4
-2
0
carbonate δ13C (‰)
Fig. 3. Coupled δ56Fe and δ13C ratios iron-rich and iron-poor carbonates from the
Brockman Iron Formation (Becker and Clayton, 1972, this study) and Kuruman Iron
Formation (Heimann et al., 2010; Johnson et al., 2003) (top panel) and from the Isua
Supracrustal Belt (Dauphas et al., 2007, this study) (bottom panel). The white box
delineates the estimated isotopic composition of Archean seawater. The gray box is the
predicted range of isotopic compositions for iron-rich carbonates precipitated directly
from and in isotopic equilibrium with Archean seawater based on experimental (JimenezLopez and Romanek, 2004; Wiesli et al., 2004) and theoretical estimates (Anbar et al.,
2005; Blanchard et al., 2009; Schauble et al., 2001) of equilibrium isotope fractionation
between Fe(II)aq-siderite and HCO3-siderite. Iron-poor carbonates from both the ~2.5 Ga
Brockman and Kuruman Iron Formations and the ~3.8 Ga Isua Supracrustal Belt have iron
and carbon isotopic compositions very similar to that expected for carbonate precipitation
in Archean seawater. In contrast, iron-rich carbonates from these iron-formations have a
wide range of iron and carbon isotopic ratios that are inconsistent with their precipitation
in equilibrium with common seawater.
127
Iron Formation are heavier than −0.5‰ and extend up to +1.2‰ (Fig. 3).
These isotopic signatures cannot be explained by direct precipitation in
near isotopic equilibrium with bulk Archean seawater.
If iron-rich carbonates were not directly precipitated from seawater,
then what were the chemical and/or biological transformations of iron
and carbon that prevailed to their formation? The range of light δ13C
ratios measured in iron-rich carbonates from the Brockman Iron
Formation requires a source of carbon in the iron-formation with a
light isotopic composition. Becker and Clayton (1972) have previously
argued that the most reasonable source of this light carbon was organic
matter. The carbon isotopic compositions of fossil organic matter in the
Hamersley Basin and other Precambrian iron-formations fall in a
restricted range of very light δ13C between −30 and −33‰ (Barghoorn
et al., 1977; Brocks et al., 1999; Perry et al., 1973). Fossil biomarkers
preserved in the Marra Mamba Iron and Jeerinah Formations of the
Hamersley Basin (Brocks et al., 1999) and in iron- and carbonateformations of the penecontemporaneous Transvaal Supergroup (Waldbauer et al., 2009) provide support for oxygenic/anoxygenic photosynthesis and organic carbon production in surface seawater in the late
Archean. Diagenetic oxidation of organic carbon would deliver a pool
of isotopically-light carbon (as CO2 or dissolved carbonate) that
could dilute the existing pore water reservoir of DIC and accumulate
in marine sediments until saturation with respect to authigenic carbonates was obtained. Whereas carbonates with δ13C ratios ~−30‰
would have formed entirely from an organic-derived pool of carbonate
carbon, iron-rich carbonates in the Brockman Iron Formation (average
δ13C = −9.8 ± 0.2‰) precipitated from a pore water reservoir of
carbonate carbon contributed by both organic and DIC sources, the
latter possibly supported by partial exchange of DIC across the
sediment–seawater interface.
Oxidation of organic carbon must be coupled to reduction of an
appropriate electron acceptor. In modern marine sediments the order of
oxidant consumption is O2 N Mn-oxides N nitrate N Fe-oxides N sulfate
(Froelich et al., 1979). In anoxic marine sediments of the Archean,
oxygen would have been severely limited (if not absent) as the primary
electron acceptor, as would Mn-oxide and nitrate (e.g., Anbar and
Holland, 1992; Anbar et al., 2007; Chapman and Schopf, 1983; Walker,
1984). Banded iron-formations are anomalously rich in iron and
reduction of ferric oxides may have permitted oxidation of organic
carbon (Dimroth and Chauvel, 1973; Walker, 1984). Studies have
demonstrated that oxidation–reduction between ferric oxide and
organic carbon is effectively mediated by microbes via the process of
DIR (Lovley, 1993), which can be illustrated by the stoichiometric
reaction for ferrihydrite,
þ
2 +
2Fe2 O3 d nH2 O + CH2 O + 7H →4Fe
~−0.20‰. It is likely that the Archean ocean, at least below the surface
mixed layer (~100 m), was well-mixed and homogeneous with respect to
iron isotopes because the reservoir of Fe(II)aq required to form BIFs must
be considerable (20 to 100 mg Fe L− 1) and continually replenished
(Ewers and Morris, 1981; Morris, 1993). In effect, the residence time of
iron in the Archean ocean was significantly longer than the oceanic
mixing time (Johnson et al., 2003). Iron-rich carbonates precipitated in
equilibrium with Archean seawater are predicted to have light iron
isotopic ratios between ~−0.5 and −2‰ (Fig. 3), based on the net isotope
fractionation between carbonate (siderite) and Fe(II)aq determined by
experiments, Δ56FeSid–Fe(II) =−0.5±0.2‰ (Wiesli et al., 2004), and
theoretical calculations, Δ56FeSid–Fe(II) =−1.6 to −2.1‰ (Anbar et al.,
2005; Blanchard et al., 2009; Schauble et al., 2001). Such low δ56Fe values
are found in iron-poor carbonates from the Wittenoom Dolomite (Fig. 3),
a platform carbonate formation in the Hamersley Group underlying the
Brockman Iron Formation. Similar light δ56Fe values are observed in
platform carbonates associated with other late Archean and Paleoproterozoic BIFs (Heimann et al., 2010). In contrast, the iron isotopic
compositions measured in most iron-rich carbonates from the Brockman
–
+ HCO3 + ð2n + 4ÞH2 O:
ð1Þ
We write HCO−
3 as the carbonate product, reflecting the dominance of
this species in seawater at equilibrium (e.g., Walker, 1983). Previous
studies have applied simple mass balance models to estimate the
overall potential for Fe(III) reduction in the formation of iron oxides
and carbonates in BIFs and implicate a significant role for microbial
iron respiration in the late Archean (e.g., Konhauser et al., 2005;
Walker, 1984).
Our isotopic data for magnetite and iron-rich carbonates from the
Brockman Iron Formation provide an independent assessment of
these results and demonstrate that the isotopic signatures imprinted
in BIFs are consistent with extensive microbial iron respiration at
~ 2.5 Ga. Ferric iron in magnetite requires the oxidation and
precipitation of dissolved Fe(II)aq. The process by which large
amounts of Fe(II)aq were oxidized in the Archean ocean that was
anoxic is debated, but the primary Fe(III) precipitates in all cases were
likely amorphous ferric oxides (Ewers and Morris, 1981; Kappler and
Newman, 2004; Morris, 1993; Trendall and Blockley, 1970), such as
P.R. Craddock, N. Dauphas / Earth and Planetary Science Letters 303 (2011) 121–132
2.5
1:1
2.0
1.5
1.0
0.5
Δ Sid-Mgt
-0.5
‰
-2
.3
‰
0.0
-1
.8
ferrihydrite. Experimental studies have demonstrated that amorphous ferric oxides from partial Fe(II)aq oxidation are enriched in the
heavy isotopes of iron relative to starting Fe(II)aq, with an equilibrium
fractionation factor ranging from Δ56Fe ~ +0.9‰ for O2-mediated
oxidation (Bullen et al., 2001) to + 2.0‰ for direct microbial oxidation
(Balci et al., 2006; Beard et al., 2010; Croal et al., 2004). Thus, partial
oxidation of Fe(II)aq in Archean seawater should produce ferrihydrite
with positive δ56Fe values up to + 1.5‰. Near-complete Fe(II)
oxidation in surface seawater would produce ferrihydrite with a
δ56Fe value of ~ 0‰, similar to that of the starting reservoir of Fe(II)aq,
which is a lower limit on the range of δ56Fe in the primary ferric
oxides. Magnetite in the Brockman Iron Formation was likely formed
by the reaction of primary ferrihydrite with ferrous iron (Ewers and
Morris, 1981; Morris, 1993):
carbonate δ56Fe (‰)
128
equilibrium
2 +
Fe2 O3 d nH2 O + Fe
–
+ 2OH →Fe3 O4 + ðn + 1ÞH2 O:
ð2Þ
Magnetite formed by reaction between ferrihydrite and Fe(II)aq in the
water column can have positive δ56Fe values between ~0 and +1.0‰,
reflecting the contributions of iron from ferrihydrite (δ56Fe ~0 to +1.5‰)
and Fe(II)aq (δ56Fe~ −0.2‰) in the ratio 2:1. Measured δ56Fe ratios of
magnetite from the Brockman Iron Formation are as light as −1.0‰,
which is difficult to explain unless the pool of Fe(II)aq from which
magnetite formed was also very light, approaching −1 to −2‰. The most
likely source of Fe(II)aq with very light δ56Fe values is from partial
reduction of ferrihydrite in marine sediments via microbial DIR (Eq. (1);
also see Johnson et al. 2008; Heimann et al. 2010). Indeed, experimental
studies have documented a significant isotopic effect associated with
partial Fe(III) reduction via DIR that yields reduced Fe(II)aq depleted by
−1.0 to −2.5‰ in δ56Fe relative to precursor ferric oxides (Beard et al.,
1999, 2010; Crosby et al., 2007; Icopini et al., 2004; Johnson et al., 2005;
Tangalos et al., 2010).
Iron-rich carbonates in the Brockman Iron Formation exhibit a
wide range of δ56Fe values (Fig. 1). Iron-rich carbonates with heavy
δ56Fe compositions up to +1‰ could only have precipitated from a
pool of Fe(II)aq with positive δ56Fe. This reservoir of iron was not
common seawater, but was an authigenic reservoir contributed from
near-complete reduction in marine sediments of primary ferrihydrite
that had heavy δ56Fe values following partial Fe(II) oxidation in surface
seawater. Microbial DIR (Eq. (1)) produces both dissolved Fe(II) and
carbonate that would accumulate in sediment pore waters until
saturation with respect to iron-rich carbonates was attained:
2 +
Fe
–
þ
+ HCO3 →FeCO3 + H :
ð3Þ
Light δ13C values of the same iron-rich carbonates (Fig. 3) have
been shown to be consistent with derivation of carbonate ion from
oxidation of organic matter during microbial DIR (Baur et al., 1985;
Becker and Clayton, 1972; Walker, 1984).
Additional support for microbial iron respiration is provided by the
correlated iron isotopic ratios of magnetite and iron-rich carbonates
sampled from the same mesobands and microbands within the Brockman Iron Formation (Fig. 4). Following reduction of ferrihydrite,
authigenic magnetite and siderite can precipitate from the same
reservoir of Fe(II)aq. Thus, the δ56Fe signature of Fe(II)aq produced by
microbial iron reduction can be inherited by both magnetite and ironrich carbonate. Our interpretations are consistent with recent microanalytical studies of the Brockman Iron Formation that have documented
iron isotopic variations in magnetite and iron-rich carbonates on spatial
scales less than one millimeter, which are suggested to reflect postdepositional redistribution of iron within the sediment (Steinhoefel
et al., 2010). Geochemical studies have shown that for the penecontemparaneous Kuruman Iron Formation, carbonate facies contain higher
residual organic carbon contents up to an order of magnitude higher
than in oxide (magnetite) facies in the iron-formation (Beukes et al.,
1990; Fischer et al., 2009). These observations are consistent with the
-1.0
-1.0
-0.5
0.0
0.5
1.0
magnetite
δ56Fe
1.5
2.0
2.5
(‰)
Fig. 4. Iron isotopic compositions of paired magnetite and iron-rich carbonates (siderite,
ankerite) from the same laminations (chert-matrix mesobands) from the Brockman Iron
Formation. The isotopic fractionation between siderite and magnetite (light gray lines) is
estimated to be between −1.8‰, based on laboratory experiments (Johnson et al., 2005;
Wiesli et al., 2004), and −2.3‰, based on spectroscopy and theoretical calculations of
vibrational states of isotopic bonding (Blanchard et al., 2009; Polyakov et al., 2007). The
gray parallelogram delineates the range of iron isotopic ratios expected for iron-rich
carbonates and magnetite precipitated in isotopic equilibrium with Archean seawater. The
measured iron isotopic compositions of most magnetite and iron-rich carbonates are
incompatible with this model. Note the overall positive correlation between the iron
isotopic compositions of magnetite and siderite, which may reflect formation of most ironrich carbonates and magnetite in association during diagenetic iron and carbon cycling.
The blue band defines the 95% confidence interval of the regression (slope= 0.69± 0.33;
the two data with light carbonate δ56Fe are omitted from this regression because these data
may reflect near-isotopic equilibrium with seawater).
idea that the relative proportions of organic carbon and ferric iron
controlled the diagenetic fate of the sediment; preserving magnetite
only when ferric iron was in excess relative to organic matter.
Most, if not all, Archean and Paleoproterozoic BIFs have experienced varying degrees of metamorphism, which can affect the stable
mineral assemblage observed in iron-formations (e.g., French, 1973;
Klein, 1983, 2005). A possible burial metamorphic origin for carbonate
in the Brockman Iron Formation via inorganic reaction between
primary ferric iron, such as ferrihydrite, and organic carbon at
elevated temperature and pressure can, however, be discounted on
the basis of mass balance for iron and carbon (DIC) in BIFs. During
metamorphism, the sediment is isolated from exchange with the
ocean. The maximum concentration of inorganic carbonate in pore
water at the onset of metamorphism can be estimated by assuming
equilibrium with atmospheric CO2. The partial pressure, pCO2, of the
Archean atmosphere is debated (e.g., Hessler et al., 2004; Kasting,
1993; Lowe and Tice, 2004; Rosing et al., 2010; Rye et al., 1995;
Walker, 1985), but pCO2 between 0.1 and 1 bar (≫ 100 times present)
is a reasonable estimate. A kilogram of BIF in the Dales Gorge Member
of the Brockman Iron Formation contains on average 6 moles of iron
(46.37 wt.% Fe2O3; Ewers and Morris, 1981). This quantity of
sediment corresponds to a pore volume of ~ 3.4 × 103 cm3 assuming
a density of the iron-formation of ~3.4 g/cm3 (Ewers and Morris,
1981) and accounting for post-depositional compaction by up to
90% (Trendall and Blockley, 1970). This pore volume could contain
up to ~0.12 moles of total dissolved carbonate. According to reaction 4
(net reaction between ferrihydrite and organic matter to yield Ferich carbonate; note that 3 moles of DIC is needed for each mole of
organic C):
−
þ
2Fe2 O3 :nH2 O þ CH2 O þ 3HCO3 þ 3H →4FeCO3 þ ð2n þ 4ÞH2 O;
ð4Þ
P.R. Craddock, N. Dauphas / Earth and Planetary Science Letters 303 (2011) 121–132
4.2. Iron and carbon isotope evidence to support microbial iron
respiration in the early Archean
We now turn our attention to interpreting the iron and carbon
isotope signatures of metacarbonates from the ~3.8 Ga Isua Supracrustal
Belt (ISB). The origin of ISB metacarbonates is contentious, with
protolith identification ranging from sedimentary (Bolhar et al., 2004;
Dymek and Klein, 1988; Mojzsis et al., 1996; Nutman et al., 1984) to
entirely metasomatic (Myers, 2001; Rose et al., 1996; Rosing et al.,
1996). Iron-rich metacarbonates were initially interpreted as carbonate
facies iron-formations (Dymek and Klein, 1988; Nutman et al., 1984). In
this interpretation framework, metacherts and calc-silicates associated
with iron-rich metacarbonates were formed by reaction of sedimentary
carbonates and quartz during burial and high-grade metamorphism,
and the banding documented in metacherts was indicated to preserve
that of the sedimentary protolith (Dymek and Klein, 1988; Nutman
et al., 1984). Trace element characteristics (e.g., REE + Y) of these rocks
were interpreted as being consistent with their deposition as chemical
sediments in seawater (Bolhar et al., 2004; Frei and Polat, 2007).
Alternatively, iron-rich and iron-poor metacarbonates and calc-silicates
have been interpreted as metasomatic in origin, formed by carbonation
and desilication of igneous protoliths. A metasomatic origin for some
metacarbonates, in particular iron-poor variants, was indicated by the
discordant nature of calc-silicate and metacarbonate units, veining,
replacive textures and of the igneous lithologies within which these
units occur (Rose et al., 1996; Rosing et al., 1996).
In a recent publication, Dauphas et al. (2007) have reported the iron
isotopic compositions and trace element geochemistry of iron-rich and
iron-poor metacarbonates from the ISB in order to better distinguish a
metasomatic versus sedimentary origin. Iron-poor and iron-rich
metacarbonates have light and heavy δ56Fe ratios, respectively
(Fig. 2). The light δ56Fe of iron-poor metacarbonates was interpreted
as reflecting mobilization of isotopically light iron from pre-existing
mafic or ultramafic protoliths by fluid, consistent with a metasomatic
origin. In support, alteration of oceanic crust at modern seafloor
hydrothermal environments yields secondary mineral assemblages
with a heavy iron isotopic composition and an implied fluid with a
complementary light iron isotopic ratio (Rouxel et al., 2003). Heavy
δ56Fe ratios up to +0.80‰ in iron-rich metacarbonates, however, are
inconsistent with mobilization by metasomatic fluids of isotopically
light iron from an igneous protolith. Dauphas et al. (2007) showed that
iron-rich metacarbonates had similar heavy iron isotopic compositions
to magnetite from BIFs but the study was inconclusive as to the nature of
this relationship. New carbon isotopic data coupled to existing iron
isotopic data for the iron-rich and iron-poor metacarbonates from the
ISB provide further constraints on a possible chemical sedimentary
versus metasomatic origin for the metacarbonates. An authigenic origin,
similar to that indicated for iron-rich carbonates from the younger
Hamersley Basin, would implicate the evolution of microbial metabolic
pathways (oxygenic or anoxygenic photosynthesis, DIR) by ~3.8 Ga.
The ISB metacarbonates fall into two isotopically distinct groups
(Fig. 3). Iron-poor metacarbonates that have light δ56Fe have near-zero
δ13C ratios (−3 to 0‰), whereas Fe-rich metacarbonates that have
heavy δ56Fe have distinctly lighter δ13C ratios (mean of −4.8 ± 0.6‰).
An important consideration in interpreting these data is whether these
isotopic signatures are primary and record the conditions of precipitation, or if these have been disturbed by metamorphism. All rocks in the
ISB have been subject to amphibolite facies metamorphism, with peak
temperatures ~500 to 550 °C and pressures ~5 kbar (e.g., Boak and
Dymek, 1982). P–T phase relations of the reaction (Lamb, 2005),
FeCO3 + SiO2 → FeSiO3 + CO2;g
ð5Þ
suggest that siderite formed during diagenesis would have survived
peak metamorphic conditions (Fig. 5).
We interpret the different iron and carbon isotopic compositions of
iron-poor and iron-rich metacarbonates in the ISB as reflecting their
original formation through distinct pathways. The field relationship
between iron-poor metacarbonates and ultramafic host rocks are
supportive of metasomatic overprint by leaching of iron by a CO2-bearing
fluid from an ultramafic protolith (Rose et al., 1996; Rosing et al., 1996).
The light iron isotopic compositions of iron-poor metacarbonates are
possibly consistent with derivation from iron mobilized from ultramafic
rocks (Dauphas et al., 2007). The carbon and iron isotopic data combined,
however, reveal that the isotopic characteristics of iron-poor metacarbonates from the ISB are very similar to those of iron-poor carbonates
9
a
Siderite + Qtz = Ferrosilite + CO 2
b
8
Pressure (kbar)
this reservoir of dissolved carbonate would be sufficient to convert up
to 0.15 moles of iron in ferrihydrite to siderite, which is only ~2% of
the total inventory of Fe in the iron-formation. Magnetite and iron
carbonate constitute on average 30 wt.% and 15 wt.%, respectively, of
the Dales Gorge Member in the Brockman Iron Formation (Trendall
and Blockley, 1970). Assuming carbonate is siderite, the average mole
fraction of iron present as carbonate is 25% (assuming ankerite gives a
minimum of 10%), which is considerably greater than the 2% of Fe that
could be converted to carbonate through metamorphism. The
calculations imply that an external source of carbonate is required
to yield the large quantities of iron-rich carbonates observed in late
Archean iron-formations. We suggest that this carbonate was made
available both from oxidation of organic carbon and through partial
exchange of DIC between sediment pore water and the overlying
water column, and could only proceed during early diagenetic
transformation. This idea is consistent with several petrographic
studies that suggested an early diagenetic origin for iron-rich
carbonates in BIFs (e.g., Ewers and Morris, 1981; Morris, 1993; Pecoits
et al., 2009; Trendall and Blockley, 1970).
Finally, we note that an interpretation to explain the light δ56Fe
values of minerals from late Archean BIF sequences (e.g., pyrite in
shale units) has been proposed by Rouxel et al. (2005) whereby
partial oxidation and precipitation of ferric oxides in BIFs in the water
column leaves the residual Fe(II)aq reservoir enriched in the light
isotopes of iron. This reservoir is subsequently precipitated as ferrousbearing minerals. We have shown that the light δ56Fe values of ironbearing minerals in BIFs can be produced by iron cycling during
diagenesis and concur with Johnson et al. (2008) that the highlyfractionated and stratigraphically variable iron isotopic compositions
of iron oxides and carbonates in oxide facies BIFs primarily reflect
diagenetic pathways for their formation.
129
e
c
7
d
XCO2 = 1
f
6
Isua P-T
5
4
500
550
600
650
700
750
Temperature (˚C)
Fig. 5. Pressure–temperature phase relationship for the reaction siderite + quartz →
ferrosilite + CO2,g (shown by black line, a). Shown in the solid gray lines are equivalent
phase boundaries for magnesite + quartz → enstatite + CO2,g (b), dolomite + quartz →
diopside + CO2,g (c), ankerite+ quartz→hedenbergite+ CO2,g (d), calcite+ enstatite +
quartz→diopside+CO2,g (e) and calcite+ferrosilite+quartz →hedenbergite+ CO2,g (f).
Phase boundaries are reproduced from Lamb (2005). The peak metamorphic P–T conditions
(~5 kbar, 550 °C) to which rocks in the ISB were subjected (e.g., Boak and Dymek, 1982) are
shown by the gray box. The phase relations indicate that siderite (and other carbonates)
would have been stable during peak metamorphism.
130
P.R. Craddock, N. Dauphas / Earth and Planetary Science Letters 303 (2011) 121–132
preserved in association with late Archean and Paleoproterozoic BIFs
(Fig. 3). By analogy, these isotopic signatures can instead be interpreted as
reflecting formation of iron-poor metacarbonates as primary precipitates
in Archean seawater. The REE patterns of these carbonates also appear to
be supportive of their formation as chemical precipitates from seawater
(Dauphas et al., 2007). Thus, while iron-poor metacarbonates have
previously been suggested to be entirely metasomatic in origin, we
caution that metasomatic mobilization of a primary (i.e., chemical
sedimentary) carbonate cannot be excluded. Continued study of the
isotopic and chemical signatures of iron-poor metacarbonates is required
to unambiguously determine the primary derivation of these rocks.
Iron-rich metacarbonates from the ISB have iron and carbon isotopic
compositions similar to those of other iron-rich carbonates of known BIF
association (Fig. 3). To the extent that iron-rich metacarbonates of the
ISB share the same isotopic characteristics with those of late Archean
iron-formations, they may have formed by the same process. The heavy
δ56Fe values of iron-rich metacarbonates cannot be explained by
metasomatic mobilization of isotopically-light iron from igneous
protoliths. Instead, it is now recognized from studies of genuine BIFs
that iron-rich carbonates of a known authigenic origin can carry heavy
δ56Fe resulting from microbial iron respiration of primary ferric oxide
precipitates also with heavy δ56Fe (Heimann et al., 2010, this study).
Magnetite in BIFs from the ISB has heavy δ56Fe (Dauphas et al., 2004b,
2007; see also Whitehouse and Fedo, 2007). Microbial Fe(III) reduction
of this magnetite or other precursor ferric oxide would transfer the
heavy isotopic composition to the ferrous iron product and subsequently to iron-rich carbonates, which is exactly that observed. Further
evidence that microbial iron respiration (DIR) was involved in the
authigenic formation of iron-rich metacarbonates comes from the light
δ13C values of these samples. Direct evidence for reduced carbon of
biogenic origin in the ISB is disputed (e.g., Perry and Ahmad, 1977;
Mojzsis et al., 1996; Rosing, 1999; Schidlowski et al., 1979; van Zuilen
et al., 2002; 2003). Still, the light carbon isotopic compositions of ironrich metacarbonates from the ISB mirror those measured of late Archean
BIFs (Baur et al., 1985; Becker and Clayton, 1972; Heimann et al., 2010),
which are consistent with derivation of carbonate from oxidation of
organic carbon during iron respiration. We conclude that iron-rich
metacarbonates associated with iron-formation in the ISB formed by
similar microbially-mediated iron and carbon transformations as
documented in late Archean BIFs. The antiquity of microbial iron
respiration, as suggested by our iron and carbon isotope data from one of
the oldest known sedimentary sequences, is consistent with results of
phylogenetic studies. Ferric iron reduction has been documented as a
metabolic pathway within a large diversity of extent Bacteria and
Archaea including those most closely related to the last common
ancestor of modern life, which points toward iron respiration as one of
the earliest forms of microbial metabolism (Liu et al., 1997; Lovley,
1991; Vargas et al., 1998; Weber et al., 2006).
5. Summary and conclusions
Iron and carbon isotopic analyses of iron oxides and carbonates in
BIFs can be used to constrain the pathways of their formation. Here, we
report the iron (δ56Fe, vs. IRMM-014) and carbon (δ13C, vs. V-PDB)
isotopic compositions of magnetite and of iron-rich and iron-poor
carbonates from the 2.5 Ga Brockman Iron Formation in the Hamersley
Basin, Australia, and the ~3.8 Ga Isua Supracrustal Belt (ISB), West
Greenland. The key results and implications from this study are:
1. Magnetite and iron-rich carbonates (siderite, ankerite) from the
Hamersley Basin preserve a wide range of δ56Fe values that are
incompatible with direct precipitation in isotopic equilibrium with
Archean seawater.
2. Magnetite with light δ56Fe (≪0‰) must have precipitated from a
pool of Fe(II)aq with light δ56Fe between −1 and −2‰, which was
likely produced through microbial partial reduction of ferrihydrite.
3. Iron-rich carbonates with heavy δ56Fe (up to +1.2‰) must have
precipitated from a reservoir of Fe(II)aq with very positive δ56Fe.
The source of heavy δ56Fe–Fe(II)aq was likely from near-complete
microbial reduction of ferric oxides (ferrihydrite) in marine
sediments. The light δ13C of the same iron-rich carbonates was
derived from carbonate produced by the oxidation of organic
carbon that was probably coupled to reduction of iron.
4. The combined iron and carbon isotopic data support an authigenic
origin for iron-rich carbonates in the Hamersley Basin via coupled
organic carbon oxidation and ferrihydrite reduction. This process is
effectively mediated by microbes in marine sediments though
dissimilatory iron reduction (DIR), which implicates extensive
microbial iron respiration in the formation of late Archean BIFs.
5. Iron-rich metacarbonates from the ~3.8 Ga ISB have δ56Fe and δ13C
signatures similar to those of carbonates in late Archean BIFs. The
isotopic data are interpreted as reflecting formation of these ironrich metacarbonates as marine authigenic precipitates through
microbial iron respiration. Despite metasomatic overprint, ironrich metacarbonates in the ISB preserve primary isotopic characteristics supporting evolution of microbial iron catabolism by
~3.8 Ga during the formation of some of the oldest recognized
sedimentary-bearing rocks on Earth.
Acknowledgments
Drill core material of the Brockman Iron Formation from Hamersley
was made available by the Geological Survey of Western Australia. R. N.
Clayton provided hand specimens of the Wittenoom Dolomite from
Hamersley. We gratefully acknowledge contributions to this research by
R. H. Becker who previously carried out the mineral separation of the
Hamersley samples for carbon and oxygen isotope analysis. Metacarbonate samples from Isua were provided by M. van Zuilen and A. Lepland.
The manuscript benefited from discussions with R. N. Clayton, and from
constructive reviews by K. Konhauser and an anonymous reviewer. This
research was supported by National Science Foundation through grant
EAR-0820807 (Geobiology), National Aeronautics and Space Administration through grant NNX09AG59G (Cosmochemistry) and a Packard
Fellowship to N.D.
Appendix A. Supplementary data
Supplementary data to this article can be found online at
doi:10.1016/j.epsl.2010.12.045.
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