A distinctive type of sedimentary
rock that formed predominantly during the Precambrian and
is the major source of the world’s iron reserves. Banded iron
formations (BIFs) are a thinly bedded, chemically precipitated,
iron-rich rock, with layers of iron ore minerals typically
interbedded with thin layers of chert or microcrystalline silica.
Many are completely devoid of detrital or clastic sedimentary
input. Most banded iron formations formed between 2.6
billion and 1.8 billion years ago, and only a few very small
similar types of deposits have been discovered in younger
mountain belts. This observation suggests that the conditions
necessary to form the BIFs were present on Earth in early
(Precambrian) time, but largely disappeared by 1.8 billion
years ago. The chemical composition and reduced state of
much of the iron of BIFs suggest that they may have formed
in an oxygen-poor atmosphere/ocean system, explaining their
disappearance around the time that atmospheric oxygen was
on the rise. BIFs may also be intimately associated with early
biological activity and may preserve the record of the development
of life on Earth. The world’s oldest BIF is located in
the 3.8-billion-year-old Isua belt in southwestern Greenland,
and some geologists have suggested that this formation contains
chemical signatures that indicate biological activity was
involved in its formation.
Banded iron formations can be divided into two main
types based on the geometric characteristics of the deposits.
Algoma-type BIFs are lenticular bodies that are closely associated
with volcanic rocks, typically basalts. Most are several
hundred meters to kilometers in scale. In contrast, Superiortype
BIFs are very large in scale, many initially covering tens
of thousands of square kilometers. Superior-type BIFs are
closely associated with shallow marine shelf types of sedimentary
rocks including carbonates, quartzites, and shales.
Banded iron formations are also divisible into four types
based on their mineralogy. Oxide iron formations contain
layers of hematite, magnetite, and chert (or cryptocrystalline
silica). Silicate iron formations contain hydrous silicate minerals,
including chlorite, amphibole, greenalite, stilpnomelande,
and minnesotaite. Carbonate iron formations contain
siderite, ferrodolomite, and calcite. Sulfide iron formations
contain pyrite.
In addition to being rich in iron, BIFs are ubiquitously
silica-rich, indicating that the water from which they precipitated
was saturated in silica as well as iron. Other chemical
characteristics of BIFs include low alumina and titanium, elements
that are generally increased by erosion of the continents.
Therefore, BIFs are thought to have been deposited in
environments away from any detrital sediment input. Some
BIFs, especially the sulfide facies Algoma-type iron formation,
have chemical signatures compatible with formation
near black smoker types of seafloor hydrothermal vents,
whereas others may have been deposited on quiet marine
platforms. In particular, many of the Superior-types of
deposits have many characteristics of deposition on a shallow
shelf, including their association with shallow water sediments,
their chemical and mineralogical constituency, and the
very thin and laterally continuous nature of their layering.
For instance, in the Archean Hamersley Basin of Western
Australia, millimeter-thick layers in the BIF can be traced for
hundreds of kilometers.
The environments that BIFs formed in and the mechanism
responsible for the deposition of the iron and silica in
BIFs prior to 1.8 Ga ago is still being debated. Any model
must explain the large-scale transport and deposition of iron
and silica in thin layers, in some cases over large areas, for a
limited time period of Earth’s history. Some observations are
pertinent. First, to form such thin layers, the iron and silica
must have been dissolved in solution. For iron to be in solution,
it needs to be in the ferrous (reduced) state, in turn suggesting
that the Earth’s early oceans and atmosphere had
little if any free oxygen and were reducing. The source of the
iron and silica is also problematic; it may have come from
weathering of continents, or from hydrothermal vents on the
seafloor. There is currently evidence to support both ideas for
individual and different kinds of BIFs, although the scales
seem to be tipped in favor of hydrothermal origins for Algoma-
types of deposits, and weathering of continents for Superior-
type deposits.
The mechanisms responsible for causing dissolved iron to
precipitate from the seawater to form the layers in BIFs have
also proven elusive and problematical. It seems likely that
changes in pH and acidity of seawater may have induced the
iron precipitation, with periods of heavy iron deposition occurring
during a steady background rate of silica deposition. Periods
of nondeposition of iron would then be marked by
deposition of silica layers. Prior to 1.8 Ga the oceans did not
have organisms (e.g., diatoms) that removed silica from the
oceans to make their shells, so the oceans would have been
close to saturated in silica at this time, easing its deposition.
Several models have attempted to bring together the
observations and requirements for the formation of BIFs, but
none appear completely satisfactory at present. Perhaps
there is no unifying model or environment of deposition, and
multiple origins are possible. One model calls on alternating
periods of evaporation and recharge to a restricted basin
(such as a lake or playa), with changes in pH and acidity
being induced by the evaporation. This would cause deposition
of alternating layers of silica and iron. However, most
BIFs do not appear to have been deposited in lakes. Another
model calls on biological activity to induce the precipitation
of iron, but fossils and other traces of life are generally rare
in BIFs, although present in some. In this model, the layers
would represent daily or seasonal variations in biological
activity. Another model suggests that the layering may have
been induced by periodic mixing of an early stratified ocean,
where a shallow surface layer may have had some free oxygen
resulting from near-surface photosynthesis, and a deeper
layer would be made of reducing waters, containing dissolved
elements produced at hydrothermal seafloor vents. In
this model, precipitation and deposition of iron would occur
when deep reducing water upwelled onto continental shelves
and mixed with oxidized surface waters. The layers in this
model would then represent the seasonal (or other cycle)
variation in the strength of the coastal upwelling. This last
model seems most capable of explaining features of the
Superior-types of deposits, such as those of the Hamersley
Basin in Western Australia. Variations in the exhalations of
deep-sea vents may be responsible for the layering in the
Algoma-type deposits. Other variations in these environments,
such as oxidation, acidity, and amount of organic
material, may explain the mineralogical differences between
different banded iron formations. For instance, sulfide-facies
iron formations have high amounts of organic carbon (especially
in associated black shales and cherts) and were therefore
probably deposited in shallow basins with enhanced
biological activity. Carbonate-facies BIFs have lower
amounts of organic carbon, and sedimentary structures
indicative of shallow water deposition, so these probably
were deposited on shallow shelves but further from the sites
of major biological activity than the sulfide-facies BIFs.
Oxide-facies BIFs have low contents of organic carbon but
have a range of sedimentary structures indicating deposition
in a variety of environments.
The essential disappearance of banded iron formation
from the geological record at 1.8 billion years ago is thought
to represent a major transition on the planet from an essentially
reducing atmosphere to an oxygenated atmosphere. The
exact amounts and rate of change of oxygen dissolved in the
atmosphere and oceans would have changed gradually, but
the sudden disappearance of BIFs at 1.8 Ga seems to mark
the time when the rate of supply of biologically produced
oxygen overwhelmed the ability of chemical reactions in the
oceans to oxidize and consume the free oxygen. The end of
BIFs therefore marks the new dominance of photosynthesis as
one of the main factors controlling the composition of the
atmosphere and oceans.
See also ATMOSPHERE; PRECAMBRIAN.
Further Reading
Morris, R. C. “Genetic Modeling for Banded Iron Formation of the
Hamersley Group, Pilbara Craton, Western Australia.” Precambrian
Research 60 (1993): 243–286.
Simonson, Bruce M. “Sedimentological Constraints on the Origins of
Precambrian Banded Iron Formations.” Geological Society of
America Bulletin 96 (1985): 244–252.
rock that formed predominantly during the Precambrian and
is the major source of the world’s iron reserves. Banded iron
formations (BIFs) are a thinly bedded, chemically precipitated,
iron-rich rock, with layers of iron ore minerals typically
interbedded with thin layers of chert or microcrystalline silica.
Many are completely devoid of detrital or clastic sedimentary
input. Most banded iron formations formed between 2.6
billion and 1.8 billion years ago, and only a few very small
similar types of deposits have been discovered in younger
mountain belts. This observation suggests that the conditions
necessary to form the BIFs were present on Earth in early
(Precambrian) time, but largely disappeared by 1.8 billion
years ago. The chemical composition and reduced state of
much of the iron of BIFs suggest that they may have formed
in an oxygen-poor atmosphere/ocean system, explaining their
disappearance around the time that atmospheric oxygen was
on the rise. BIFs may also be intimately associated with early
biological activity and may preserve the record of the development
of life on Earth. The world’s oldest BIF is located in
the 3.8-billion-year-old Isua belt in southwestern Greenland,
and some geologists have suggested that this formation contains
chemical signatures that indicate biological activity was
involved in its formation.
Banded iron formations can be divided into two main
types based on the geometric characteristics of the deposits.
Algoma-type BIFs are lenticular bodies that are closely associated
with volcanic rocks, typically basalts. Most are several
hundred meters to kilometers in scale. In contrast, Superiortype
BIFs are very large in scale, many initially covering tens
of thousands of square kilometers. Superior-type BIFs are
closely associated with shallow marine shelf types of sedimentary
rocks including carbonates, quartzites, and shales.
Banded iron formations are also divisible into four types
based on their mineralogy. Oxide iron formations contain
layers of hematite, magnetite, and chert (or cryptocrystalline
silica). Silicate iron formations contain hydrous silicate minerals,
including chlorite, amphibole, greenalite, stilpnomelande,
and minnesotaite. Carbonate iron formations contain
siderite, ferrodolomite, and calcite. Sulfide iron formations
contain pyrite.
In addition to being rich in iron, BIFs are ubiquitously
silica-rich, indicating that the water from which they precipitated
was saturated in silica as well as iron. Other chemical
characteristics of BIFs include low alumina and titanium, elements
that are generally increased by erosion of the continents.
Therefore, BIFs are thought to have been deposited in
environments away from any detrital sediment input. Some
BIFs, especially the sulfide facies Algoma-type iron formation,
have chemical signatures compatible with formation
near black smoker types of seafloor hydrothermal vents,
whereas others may have been deposited on quiet marine
platforms. In particular, many of the Superior-types of
deposits have many characteristics of deposition on a shallow
shelf, including their association with shallow water sediments,
their chemical and mineralogical constituency, and the
very thin and laterally continuous nature of their layering.
For instance, in the Archean Hamersley Basin of Western
Australia, millimeter-thick layers in the BIF can be traced for
hundreds of kilometers.
The environments that BIFs formed in and the mechanism
responsible for the deposition of the iron and silica in
BIFs prior to 1.8 Ga ago is still being debated. Any model
must explain the large-scale transport and deposition of iron
and silica in thin layers, in some cases over large areas, for a
limited time period of Earth’s history. Some observations are
pertinent. First, to form such thin layers, the iron and silica
must have been dissolved in solution. For iron to be in solution,
it needs to be in the ferrous (reduced) state, in turn suggesting
that the Earth’s early oceans and atmosphere had
little if any free oxygen and were reducing. The source of the
iron and silica is also problematic; it may have come from
weathering of continents, or from hydrothermal vents on the
seafloor. There is currently evidence to support both ideas for
individual and different kinds of BIFs, although the scales
seem to be tipped in favor of hydrothermal origins for Algoma-
types of deposits, and weathering of continents for Superior-
type deposits.
The mechanisms responsible for causing dissolved iron to
precipitate from the seawater to form the layers in BIFs have
also proven elusive and problematical. It seems likely that
changes in pH and acidity of seawater may have induced the
iron precipitation, with periods of heavy iron deposition occurring
during a steady background rate of silica deposition. Periods
of nondeposition of iron would then be marked by
deposition of silica layers. Prior to 1.8 Ga the oceans did not
have organisms (e.g., diatoms) that removed silica from the
oceans to make their shells, so the oceans would have been
close to saturated in silica at this time, easing its deposition.
Several models have attempted to bring together the
observations and requirements for the formation of BIFs, but
none appear completely satisfactory at present. Perhaps
there is no unifying model or environment of deposition, and
multiple origins are possible. One model calls on alternating
periods of evaporation and recharge to a restricted basin
(such as a lake or playa), with changes in pH and acidity
being induced by the evaporation. This would cause deposition
of alternating layers of silica and iron. However, most
BIFs do not appear to have been deposited in lakes. Another
model calls on biological activity to induce the precipitation
of iron, but fossils and other traces of life are generally rare
in BIFs, although present in some. In this model, the layers
would represent daily or seasonal variations in biological
activity. Another model suggests that the layering may have
been induced by periodic mixing of an early stratified ocean,
where a shallow surface layer may have had some free oxygen
resulting from near-surface photosynthesis, and a deeper
layer would be made of reducing waters, containing dissolved
elements produced at hydrothermal seafloor vents. In
this model, precipitation and deposition of iron would occur
when deep reducing water upwelled onto continental shelves
and mixed with oxidized surface waters. The layers in this
model would then represent the seasonal (or other cycle)
variation in the strength of the coastal upwelling. This last
model seems most capable of explaining features of the
Superior-types of deposits, such as those of the Hamersley
Basin in Western Australia. Variations in the exhalations of
deep-sea vents may be responsible for the layering in the
Algoma-type deposits. Other variations in these environments,
such as oxidation, acidity, and amount of organic
material, may explain the mineralogical differences between
different banded iron formations. For instance, sulfide-facies
iron formations have high amounts of organic carbon (especially
in associated black shales and cherts) and were therefore
probably deposited in shallow basins with enhanced
biological activity. Carbonate-facies BIFs have lower
amounts of organic carbon, and sedimentary structures
indicative of shallow water deposition, so these probably
were deposited on shallow shelves but further from the sites
of major biological activity than the sulfide-facies BIFs.
Oxide-facies BIFs have low contents of organic carbon but
have a range of sedimentary structures indicating deposition
in a variety of environments.
The essential disappearance of banded iron formation
from the geological record at 1.8 billion years ago is thought
to represent a major transition on the planet from an essentially
reducing atmosphere to an oxygenated atmosphere. The
exact amounts and rate of change of oxygen dissolved in the
atmosphere and oceans would have changed gradually, but
the sudden disappearance of BIFs at 1.8 Ga seems to mark
the time when the rate of supply of biologically produced
oxygen overwhelmed the ability of chemical reactions in the
oceans to oxidize and consume the free oxygen. The end of
BIFs therefore marks the new dominance of photosynthesis as
one of the main factors controlling the composition of the
atmosphere and oceans.
See also ATMOSPHERE; PRECAMBRIAN.
Further Reading
Morris, R. C. “Genetic Modeling for Banded Iron Formation of the
Hamersley Group, Pilbara Craton, Western Australia.” Precambrian
Research 60 (1993): 243–286.
Simonson, Bruce M. “Sedimentological Constraints on the Origins of
Precambrian Banded Iron Formations.” Geological Society of
America Bulletin 96 (1985): 244–252.
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