DEFINITION OF BANDED IRON FORMATION

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.

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