FRONTIER RESEARCH ON EARTH EVOLUTION, VOL. 2
Evolution of the geochemical cycle of Fe through geologic time:
Iron isotope perspective
Kosei E. Yamaguchi1, 2
1 Research
2 NASA
Program for Paleoenvironments, Institute for Research on Earth Evolution (IFREE)
Astrobiology Institute
1. Introduction
how, when, and why anoxic environments prevailed and how life
responded, survived, and evolved during the last 200 Ma.
Motivation of this paper comes from attempts to apply Fe isotope
geochemistry as a “redox probe” to geologic samples of critical
time intervals in the history of the Earth, where there were important environmental changes. As such, Archean-Paleoproterozoic
samples were chosen for this study. On-going projects of Fe isotope geochemistry of sediments / sedimentary rocks in younger
ages, such as those of modern anoxic lake, Cenomanian-Turonian
OAE2, K/T boundary, and Paleocene-Eocene Thermal Maximum
(PETM), will be reported later.
Detailed description of used samples, their geologic settings,
and analytical method are not the scope of this paper and thus not
presented. See Yamaguchi [2005a] for the discussion of the rise of
pO2 on which this paper is based, Yamaguchi et al. [2005a] for
descriptions of the samples, their geologic settings, and analytical
method, and Yamaguchi [2005b] for sample preparation method.
In the Earth’s crust, Fe exists as 4th most abundant element. In
the surface environments of the Earth, Fe occurs as either ferrous
Fe (Fe2+; reduced) in O2-poor environments or as ferric Fe (Fe3+;
oxidized) in O2-rich environments. As apparent from Eh-pH diagram for Fe, Fe3+ is essentially insoluble but Fe2+ is soluble in
aqueous solutions under normal pH conditions. Although controversy continues [e.g., Ohmoto, 1997; Holland, 1999], it has been
believed that the Earth could have been relatively O2-poor >2 billion years (Ga) ago [e.g., Kasting,, 1987; Holland, 1994;
Yamaguchi, 2005a;]. Therefore, redox-sensitive nature of Fe has
been used by many investigators to track the redox evolution of
the Earth’s atmosphere/oceans, and thus the evolution of Fe geochemical cycle in the history of the Earth. As indirect indicators of
the atmospheric / oceanic redox conditions in the distant past,
geochemistry of Fe has been applied to Fe-bearing geologic materials such as paleosols, shales, sandstones, conglomerates, red
beds, and banded iron-formations (BIFs) [e.g., Holland, 1984,
1994]. However, consensus on the interpretation of the obtained
data sets has not been reached, and controversy still continues vigorously [e.g., Ohmoto, 1997; Holland, 1999; Phillips et al., 2001;
Yamaguchi, 2005a]. To resolve such controversy, new analytical
approaches have been awaited among geoscientists.
There are four stable isotopes of Fe that occur naturally: 54Fe
(5.8%), 56Fe (91.8%), 57Fe (2.1%), and 58Fe (0.3%). High precision
measurements of Fe isotope compositions have been made possible
only recently, with the technical development of an analytical
instrument called MC-ICPMS (multi-collector inductively coupled
plasma mass spectrometer). Published studies have shown that there
is ~4 per mil (‰) range (in 56Fe/54Fe ratios) in isotope variations
preserved in geologic records [Beard and Johnson, 2004]. It has
also been shown that isotope fractionation of Fe is relatively large
during inorganic redox reactions at low temperatures and during
biological processing of Fe. Iron isotope geochemistry, an emerging
new field of geochemistry, is rapidly growing because of great
potentials to be used as a tool to trace geochemical cycle of Fe.
Application of Fe isotope geochemistry to ancient rock records in
the Precambrian is likely to open a door for better understanding of
the evolution of the geochemical cycle of Fe.
This contribution consists of three parts. A brief review of how
Fe geochemistry has been applied to constrain the redox evolution
of the Earth’s surface environments is given first, followed by that
of associated problems and points of discussions. Then introduced
is how Fe isotope geochemistry has been applied to ArcheanPaleoproterozoic rocks. Lastly, importance of Fe isotope geochemistry using younger samples is emphasized to advance our understanding of environmental changes in recent and distant past.
One of the IFREE’s major objectives is to better understand
2. Geochemistry of Fe in Archean-Paleoproterozoic
rocks as an indicator of pO2 level
Paleosols: As ancient weathering profiles of continental rocks,
paleosols are presumably in direct contact with the overlying
atmosphere. Chemistry of weathering fluids is influenced by
redox state of the atmosphere. Thus, because of redox-sensitive
nature of Fe, its loss or retention from upper part of paleosols has
been used as an indicator of the redox state of the atmosphere
[e.g., Holland, 1994]. However, there are many shortcomings in
using Fe geochemistry of paleosols as a redox indicator.
Hydrothermal alteration of paleosols, often recognized in many
sections, tends to erase the original chemical characteristics.
Correct recognition of paleosols is difficult [e.g., Rye and
Holland, 1998; Beukes et al., 2002]. Moreover, detailed geochemistry of many paleosols using Fe3+/2+/Ti ratios has shown that
essentially all paleosols retain characteristics of soils formed in
oxygenated environments [Ohmoto, 1996].
Shales: Pyrite in >2.2 Ga shales is thought to have formed by
magmatic H2S because of the common belief that oceanic sulfate
level was very low and sulfate-reducing bacteria (SRB) were
absent [e.g., Hattori et al., 1983]. Such interpretation is based on
bulk-rock sulfur isotope analyses of pyrite in shales. However,
micro-scale sulfur isotope analysis suggest that SRB were active
and sulfate level in Archean oceans was >10 mM [e.g., Ohmoto et
al., 1993].
Sandstones: Presence and assemblages of “detrital” heavy
unstable minerals such as siderite (FeCO3) in 3.2~2.8 Ga sandstones have been used to indicate that the atmosphere was reducing [Rasmussen and Buick, 1999]; however, it has been ques1
FRONTIER RESEARCH ON EARTH EVOLUTION, VOL. 2
tioned, using theoretical geochemistry, from conditions to stabilize of such minerals [Ohmoto, 1999]. No other discovery of such
minerals has been reported since 1999.
Conglomerates: Uraninites and pyrites, unstable (i.e., very soluble) in oxygenated environments, in quartz-pebble conglomerates of >2.4 Ga such as those in Witwatersrand in South Africa
and Elliot Lake in Canada have been thought to be detrital in origin, and thus the atmosphere at the time of deposition has been
thought to have been anoxic [e.g., Holland, 1994]. However,
detailed chemical and isotopic microanalysis has shown that
pyrite grains in Elliot Lake district have multiple crystal growth
history involving sulfate reduction [Yamaguchi and Ohmoto,
1998, 2005] and some pyrites in Witwatersrand district are pyritized BIF fragments [Ohtake et al., 2004].
BIFs: A popular model for the formation of BIFs postulates
that the distal deepwater rich in Fe2+ supplied from submarine
hydrothermal venting in/near mid oceanic-ridges was transported
by upwelling to the shallow basins near continental margins,
where precipitation of Fe as Fe3+-hydroxides occurred by reacting
with O2 produced locally and seasonally by oxygenic photosynthesizers such as cyanobacteria [e.g., Beukes and Klein, 1992]. In
order for deepwater to be rich in dissolved Fe2+ and to be transported to great distances, the deepwater is required to have been
globally and permanently anoxic and thus the overlying atmosphere needs to have been also anoxic. However, only Superiortype BIFs are considered to construct such model and Algomatype BIFs are not considered. The latter are typically deep facies
and always associated with volcanic (basaltic) rocks and do not
normally contain clastic sedimentary components. A question
remains how to form Fe-oxide in deep oceans to form Algomatype BIFs.
Ga black shales in South Africa and Australia, suggesting that Fereducing bacteria were active [Yamaguchi et al., in prep]; (4) Fe
isotope compositions of 3.2 Ga greywacke in South Africa
[Yamaguchi et al., 2005a]; (5) Fe isotope compositions of pyrite
in 2.9 Ga uraniferous quartz-pebble conglomerates in the
Witwatersrand district of South Africa, possibly suggesting nondetrital origin of the pyrite [Yamaguchi et al., in prep]; (6) Fe isotope compositions of 2.7 Ga red beds in Canada [Yamaguchi et
al., in prep]; (7) Fe isotope compositions of 2.2 Ga red shales in
South Africa [Yamaguchi et al., 2005a]; and (8) Fe isotope compositions of 2.5 Ga BIFs in South Africa [Johnson et al., 2003].
These data sets, although much more data are necessary, will
require revisions of previous ideas concerning the origins of Febearing minerals in the above samples and environmental factors
controlling those Fe mineralogy.
4. Iron isotope geochemistry: Secular changes?
When the Fe isotope compositions of sedimentary materials
together with those literature data are plotted against depositional
ages, an interesting picture will emerge. Figure 1 summarizes the
Fe isotope data for sediments and sedimentary rocks throughout
geologic time, since 3.5 Ga to present [Yamaguchi et al., 2005a].
Wide ranges in the δ56Fe values are observed for sedimentary
rocks over the last 3.3 Ga (Fig. 1). Figure 1 apparently illustrates a
secular change of Fe isotope compositions and thus evolution of
geochemical cycle of Fe through geologic time. However, it needs
to be seen with cautions. Although it is striking that the largest
range in the δ56Fe values seems to occur in the Archean, it should
be noted that the Archean sample suite is biased toward samples
that are rich in organic carbon (Corg), and many contain high Fe
and carbonate contents. On the other hand, samples of similar
composition in terms of C org and Fe contents from the
Mesoproterozoic, Neoproterozoic and early Phanerozoic have not
been analyzed.
Nevertheless, sedimentary rocks that are poor in Corg, Ccarfb (carbonate carbon), and Fe appear to have relatively constant δ56Fe values near zero over much of Earth’s history. Such δ56Fe values deviate little from the average for igneous rocks (0.00 ± 0.05 ‰), or the
compositions of modern river sediments, aerosols, loess, and modern clastic marine rocks ((0.00 ± 0.05 ‰; Fig. 1). The largest range
in the δ56Fe values are found in Corg-rich rocks or those that are rich
in carbonate, magnetite, or pyrite, although, so far, Fe isotope variations for such rocks of Archean in age seem to be larger than those
of Mesozoic rocks (Fig. 1). An important exception is the large
range measured for Fe2+-rich pore fluids and pyrite in relatively
anoxic, modern marine sediments (Fig. 1).
3. Iron isotope geochemistry: Constraining ArcheanPaleoproterozoic environments
All of the materials presented above can be good targets of Fe
isotope geochemistry, not simply because the samples contain lots
of Fe for high quality measurements, but because redox chemistry
of Fe plays a crucial role to form Fe-bearing minerals in the above
geologic materials and associated isotope fractionations are
expected to have occurred and recorded in the samples. For example, isotope compositions of Fe in paleosols may tell us about its
geochemical history as to whether a paleosol suffered oxidation
after Fe-leaching at the time of formation or it suffered reduction
first such as by hydrothermal alteration after Fe oxidation at the
time of formation, or how much Fe was added to or removed from
the original paleosol section. Iron isotope compositions of pyritebound-Fe in shales, sandstones, and conglomerates may also tell
us about how pyrite was formed, i.e., detrital (magmatic) or biogenic in origin. Isotope compositions of Fe in various Fe-bearing
phases in BIFs [e.g., hematite, magnetite, siderite, and pyrite] may
tell us about their origins and the degree of involvement of biological activity to form Fe-bearing minerals.
Such data sets do exist. They include: (1) Fe isotope compositions of 2.2 Ga lateritic paleosols in Botswana, suggesting extensive water-rock interaction during leteritization and development
of terrestrial biota [Yamaguchi, 2005c; Yamaguchi et al., 2005b];
(2) Fe isotope compositions of 3.2 ~ 2.2 Ga black shales in South
Africa and Australia [See Yamaguchi, 2005d; Yamaguchi et al.,
2005a]; (3) Fe isotope compositions of pyrite grains in 3.2 ~ 2.2
5. Iron isotope geochemistry: Future directions
Present is the key to the past. In order to better understand the
Precambrian evolution of Fe geochemical cycle from Fe isotope
perspective, we need to understand Fe geochemical cycle in relatively younger time using modern-day, Cenozoic, or Phanerozoic
samples. As emphasized above, Fe isotope compositions of sedimentary rock with low Corg and Fe contents have not been analyzed for Mesoproterozoic, Neoproterozoic and Paleozoic. The
data for those of Mesozoic and Cenozoic are still very scarce.
Much work needs to be done. One of the author’s research activities at IFREE, Fe isotope geochemistry of various modern sedi2
FRONTIER RESEARCH ON EARTH EVOLUTION, VOL. 2
ments and Cenozoic and Mesozoic sedimentary rocks, aims to fulfill the gap presented above toward the proposed goal. A vast variety of experimental studies involving (micro)biology needs also to
be done in order to investigate how a variety of microbes fractionate Fe isotopes during microbial (assimilatory and dissimilatory)
processing of Fe at different environmental conditions [e.g., pressure, temperature, pH, Eh, nutrient, light, UV, humidity, etc.].
Together, an equal variety of inorganic experimental studies needs
to be done, with an equal importance to above-mentioned biological experiments, to investigate a variety of kinetic and equilibrium
fractionation of Fe isotopes among various Fe-bearing solid phases liquid phases at different experimental conditions (pressure,
temperature, pH, Eh, UV, etc.). Then, we will be able to establish
a firm basis for interpretation of Fe isotope data throughout geologic time. There are so many future topics for Fe isotope studies,
and door is wide-open.
Geochemical News, 93, 12-13 and 26-27, 1997.
Ohmoto, H., T. Kakegawa and D.R. Lowe, 3.4-billion-year-old biogenic pyrites from Barberton, South Africa: sulfur isotope evidence, Science, 262, 555-557, 1993.
Ohtake, T., Y. Watanabe, W. Alterman and H. Ohmoto, “Detrital”
pyrites in the Archean Witwatersrand Basin (South Africa) are not
detrital, Int’l. J. Astrobiol., 1, 38-39, 2004.
Phillips, G.N., J.D.M. Law and R.E. Myers, Is the redox state of the
Archean atmosphere constrained? Soc. Econ. Geol. Newsletter, 47,
2001.
Rasmussen, B. and R. Buick, Redox state of the Archean atmosphere:
Evidence from detrital heavy minerals in ca. 3250-2750 Ma sandstones from the Pilbara Craton, Australia, Geology, 27, 115-118,
1999.
Rye, R. and H.D. Holland, Paleosols and the evolution of atmospheric
oxygen: a critical review, Am. J. Sci., 88, 621-672, 1998.
Yamaguchi, K.E., Evolution of the atmospheric oxygen in the early
Precambrian: An updated review of geological “evidence.” In
Frontier Research on Earth Evolution (IFREE Report for 20032004), Ed. Y. Fukao, 2 (this volume), 2005a.
Yamaguchi, K.E., Iron isotope analysis by multi-collector inductively
coupled plasma mass spectrometer: I. Sample preparation method,
In Frontier Research on Earth Evolution (IFREE Report for 20032004), Ed. Y. Fukao, 2 (this volume), 2005b.
Yamaguchi, K.E., Iron isotope compositions of Fe-oxide as a measure
of water-rock interaction: An example from Precambrian tropical
laterite in Botswana, In Frontier Research on Earth Evolution
(IFREE Report for 2003-2004), Ed. Y. Fukao, 2 (this volume),
2005c.
Yamaguchi, K.E., Isotope evidence for 3 billion years of bacterial
redox cycling of iron. In Frontier Research on Earth Evolution
(IFREE Report for 2003-2004), Ed. Y. Fukao, 2 (this volume),
2005d.
Yamaguchi, K.E. and H. Ohmoto, Diverse origins of pyrites in
Paleoproterozoic uraniferous quartz-pebble conglomerate, Elliot
Lake, Canada: Evidence from laser-microprobe sulfur isotope
analyses, Mineral. Mag., 62A, 1673-1674, 1998.
Yamaguchi, K.E., and H. Ohmoto, Evidence from sulfur isotope and
trace elements in pyrites for their multiple post-depositional
processes in uranium ores at the Stanleigh Mine, Elliot Lake,
Ontario, Canada. In Evolution of the Early Atmosphere,
Hydrosphere, and Biosphere: Constraints from Ore Deposits, Ed.,
S. Kesler and H. Ohmoto, Geological Society of America, Special
Publication, in press, 2005.
Yamaguchi, K.E., C.M. Johnson, B.L. Beardand and 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,
135-169, 2005a.
Yamaguchi, K.E., C.M. Johnson, B.L. Beard, N.J. Beukes, J. Gutzmer
and H. Ohmoto, Iron isotope fractionation during Paleoproterozoic
lateritization of the Hekpoort paleosol profile from Gaborone,
Botswana, Earth Planet. Sci. Lett., Submitted, 2005b.
Acknowledgements. Profs. Hiroshi Ohmoto, Clark Johnson, and
Dr. Brian Beard are thanked for discussion regarding Fe isotope geochemistry and atmospheric evolution. Discussions with members of
IFREE4 were also beneficial.
References
Beard, B.L. and C.M. Johnson, Fe isotope variations in the modern
and ancient Earth and other planetary bodies, In Geochemistry of
Non-Traditional Stable Isotopes, Ed., C.M. Johnson, B.L. Beard,
F. Albarde, Rev. Mineral. Geochem., 55, 319-357, 2004.
Beukes, N.J. and C. Klein, Models for iron-formation deposition, In
The Proterozoic Biosphere: A Multidisciplinary Study, Ed., J.W.
Schopf, and C. Klein, Cambridge University Press, Cambridge,
England, 147-152, 1992.
Beukes, N.J., H. Dorland, J. Gutzmer, M. Nedachi and H. Ohmoto,
Tropical laterites, life on land, and the history of atmospheric oxygen in the Paleoproterozoic. Geology, 30, 491-494, 2002.
Hattori, K., F.A. Campbell and H.R. Krouse, Sulphur isotope abundances in Aphebian clastic rocks: implications for the coeval
atmosphere, Nature, 302, 323-326, 1983.
Holland, H.D., Chemical evolution of the atmosphere and ocean,
Princeton University Press, Princeton, 582p, 1984.
Holland, H.D., Early Proterozoic atmospheric change, In Early Life on
Earth. Nobel Symposium No. 84, Ed., S.Bengston, Columbia
University Press, New York, 237-244, 1994.
Holland, H.D., When did the Earth's atmosphere become oxic? A
Reply, Geochemical News, 100, 20-22, 1999.
Johnson, C.M., B.L. Beard, N.J. Beukes, C. Klein and J.M. O’Leary,
Ancient geochemical cycling in the earth as inferred from Fe isotope studies of banded iron formations from the Transvaal Craton,
Cont. Mineral. Petrol., 144, 523-547, 2003.
Kasting, J.F., Theoretical constraints on oxygen and carbon dioxide
concentrations in the Precambrian atmosphere, Precam. Res., 34,
205-228, 1987.
Matthews, A., H.S. Morgans-Bell, S. Emmanuel, H.C. Jenkyns, Y.
Erel and L. Halicz, Controls of iron-isotope fractionation in organic-rich sediments (Kimmeridge Clay, Upper Jurassic, southern
England), Geochim. Cosmochim. Acta, 68, 3107-3123, 2004.
Ohmoto, H., Evidence in pre-2.2 Ga paleosols for the early evolution
of atmospheric oxygen and terrestrial biota, Geology, 24, 11351138, 1996.
Ohmoto, H., When did the Earth's atmosphere become oxic? The
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FRONTIER RESEARCH ON EARTH EVOLUTION, VOL. 2
Figure 1. Iron isotope compositions of the Archean-Paleoproterozoic sedimentary rocks of
this study versus their depositional ages. Literature data are also included for comparison:
magnetite, hematite, siderite, and pyrite separates from BIF of the 2.5 Ga Kuruman Iron
Formation [Johnson et al., 2003]; late Jurassic Kimmeridge Clay Formation from UK
[Matthews et al., 2004], modern aerosol, suspended river load, loess, and marine sediments [Beard et al., 2003a]. Also included are data for supergene Fe ore from the
Griqualand West, South Africa [Yamaguchi et al., unpub. data], and for the Cretaceous
(Cenomanian-Turonian) black shales and their adjacent carbonate rocks from Italy
[Yamaguchi et al., unpub. data]. Modified after Yamaguchi et al. [2005a].
4