Download Sedimentary iron deposits

3 Sedimentary iron deposits
3.1 Introduction
Iron is present in all types of sedimentary rocks having concentration of a few per cent,
but in some cases it forms ironstones and iron-formations, where the iron content exceeds
15%.
The element iron occurs in two valence states: a divalent form, ferrous iron (Fe2+) and
trivalent form, ferric iron (Fe3+). As a result of this, the behavior of iron and the
precipitation of its minerals are strongly controlled by the chemistry of the surface and
diagenitic environments. The common iron minerals in sedimentary rocks are given in
Tab. 3.1.
Tab. 3.1: The iron minerals of sedimentary rocks
The majority of sedimentary iron deposits were formed in marine environments and
many of the Paleozoic and Mesozoic examples contain normal marine fossils.
In the early middle Precambrian, thick units of various iron minerals interbedded with
chert, called banded iron formations (BIF) were deposited in large intracratonic basins.
On the other hand, in the Phanerozoic, thin units of ironstones, mainly oolitic, were
deposited in localized small areas.
3.2 Sources and transportation of iron
There are two sources of iron for formation of ironstones: continental weathering and
volcanism. Intense chemical weathering under humid tropical climate releases the iron
from mafic and heavy minerals in igneous and other rocks and produces iron-charged
ground water and iron-rich lateritic soils. Through erosion of these soils, the iron is
transported to the sea by rivers.
Volcanic and hydrothermal activity can supply considerable amounts of iron, as the case
in the Precambrian.
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Iron in true solution in river water and ground water is in very low concentration (less
than 1 ppm) and in sea water the concentration is around 0.003 ppm. This is because in
natural surface water the Ph and Eh ranges make iron to be stable only as insoluble ferric
hydroxide, not as soluble ferrous phase as Fig. 3.1 shows.
Fig. 3.1: Eh-pH diagram showing the fields of some naturally occurring waters. Redox
(reduction-oxidation) potentials of natural solutions are limited by reactions involving
water, which are dependent on pH. The upper limit of Eh (is a measure of the oxidizing
or reducing nature of the solution) is determined by the oxidation of water to oxygen (the
upper diagonal line) and the lower limit of Eh is the reduction of water to hydrogen
(lower diagonal). The pH is a measure of the acidity or alkalinity of the solution.
Iron can be transported by three mechanisms:
1- by rivers in the form of colloidal suspension of ferric hydroxide which is stabilized in
the presence of organic matter. This colloidal iron can be precipitated in the sea though
flocculation of the colloidal suspension;
2- by adsorption and chelation onto organic matter;
3- by clay minerals, either as part of the clay structure, or of more important, as oxide
films on the surface of clays.
When the clays and organic matter are deposited, iron is released from them into the pore
water of sediments if the Eh-pH conditions are appropriate, and then reprecipitated to
form iron minerals.
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Ironstone formation is favored when there are low rates of sedimentation, both
siliciclastic material and carbonates.
Regarding the quantity of iron in the Archean and early Proterozoic iron formations, it is
too high to be explained by continental weathering and volcanic-hydrothermal activities.
Therefore, it is suggested that the Precambrian atmosphere was different than present
atmosphere, in having more carbon dioxide and less oxygen in order to allow leaching
and transportation of iron more efficiently, particularly as Fe2+ in solution.
3.3 The formation of the principal iron minerals
Fig. 3.1 shows that Fe3+ is stable under more oxidizing and more alkaline conditions,
whereas Fe2+ is stable under more reducing and more acidic conditions.
In fact, in the pH-Eh range of natural environments, Fe3+ is present as the highly
insoluble Fe(OH), whereas Fe2+ is present in solution.
Other two factors control the precipitation of iron minerals:
1- The activity (the effective concentration) of carbonate ions that can be measured by
partial pressure of CO2 to be expressed as PCO2.
2- The activity of sulphur, represented by pS2-, the negative logarithm of the activity of
the sulphide ion.
The stability fields of the common iron minerals are plotted on Eh-pH, Eh- pS2-, and EhPCO2 diagrams in Figs. 3.2, 3.3 and 3.4.
Fig. 3.2: Eh-pH diagram showing the stability fields of ferrous and ferric iron, hematite,
siderite, pyrite and magnetite.
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Fig. 3.3: The stability fields of iron minerals plotted on an Eh- pS2- diagram for a pH of
7.4.
Fig. 3.4: The stability fields of iron minerals plotted on Eh-log PCO2.
It can be seen from Figs 3.3 and 3.4 that hematite is the most stable mineral under
moderately to strongly oxidizing conditions.
Generally, it is believed that hematite forms diagenetically from a hydrated ferric oxide
precursor similar to goethite, by an aeging process involving dehydration.
For the ferrous minerals, pyrite, siderite and magnetite, they are stable under conditions
of negative Eh. Thus these minerals are precipitated within sediments (not at sediment
surface) during early diagenesis, where reducing conditions have developed through
bacterial decomposition of organic matter.
The iron for these minerals is Fe2+ in the pore waters, mostly liberated by bacterial
reduction of iron oxides/hydroxides on clays and organic matter in the sediments.
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Precipitation of pyrite requires also sulphide that comes from bacterial reduction of
dissolved sulphate in pore water, which produces H2S that reacts with the Fe2+ in
solution. This dissolved sulphate is present in seawater, so that pyrite is atypical
authigenic mineral of organic-rich marine muds, forming in the anoxic (free of oxygen)
sulphidic diagenetic environment. Dissolved sulphate has very low concentration in fresh
water, thus pyrite is not common in non-marine sediments. Pyrite is black and finely
crystalline.
Siderite is precipitated where the carbonate activity is high and the sulphide activity is
low. The latter is rarely attained in marine sediment pore waters because of the abundant
dissolved sulphate. However, if all the SO42- is reduced, as the case in anoxic nonsulphidic methane –rich diagenetic environment, then siderite will precipitate. Therefore,
siderite is more common in non-marine sediments.
If there is insufficient Fe2+ relative to Ca2+ and Mg2+ in pore waters, then ankerite,
CaMg0.5Fe0.5(CO3)2 or ferroan calcite or ferroan dolomite may form in preference to
siderite.
The formation of magnetite is favored by low activities of both sulphide and carbonate,
together with negative Eh and neutral pH (Figs. 3.3 and 3.4). Such conditions are rare in
nature and so magnetite precipitation is not common.
3.4 Occurrence and petrography of iron minerals
Hematite is present in both Precambrian iron formations and Phanerozoic ironstones.
In the Precambrian, hematite occurs as thin beds and laminae, alternating with chert, and
also as massive, peloidal and oolitic form. On the other hand, hematite in the Phanerozoic
ironstones occurs mainly as ooids and impregnations and replacement of fossils (Fig.
3.5).
Fig. 3.5: Hematite replacing crinoidal fragments and carbonate ooids in Jurassic
sediment, plane polarized light.
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Hematite in thin sections is opaque and typically cryptocrystalline, but could by coarse
crystalline, deep red, translucent if it occurs as cement in sandstones.
Goethite is absent from Precambrian ironstones, but abundant in Phanerozoic ones. It
forms ooids that are quite spherical. Goethite ooids and pisoids form in lateritic soils of
tropical regions. Goethite in thin sections is a yellow- to brown-colored and generally
appears isotropic.
Limonite is a poorly defined hydrated form of iron oxide, containing goethite, other
materials such as clays and adsorbed water. The term is best restricted to the yellowbrown amorphous product of subaerial weathering of iron oxides and other minerals.
Magnetite is abundant in Precambrian iron formations, where it is interlaminated with
chert and rare in Phanerozoic ones. It occurs as small replacement crystals or granules
within oolitic ironstones.
Siderite is a major constituent of both Precambrian and Phanerozoic iron-rich sediments.
It occurs as cement to many Phanerozoic berthierine-chamosite oolites and it can replace
ooids and skeletal grains. Aggregates of fibrous spherulitic siderite form a rock called
sphaerosiderite (Fig. 3.6).
Fig. 3.6: Sphaerosiderite consisting of spherulites of fibrous siderite, plane polarized
light.
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Pyrite is a minor mineral in organic-rich sediments, such as estuarine and tidal flat ones.
It occurs as disseminated grains and crystals (cubic), and could also replace skeletal
fragments. Aggregates of spherical microconcretions of pyrite are known as framboids.
Iron silicate minerals include berthierine-chamosite, greenalite and glauconite.
Berthierine-chamosite typically occurs as ooids in Phanerozoic ironstones, within a
cement of siderite or calcite (Figs. 3.7, 3.8)
Fig. 3.7: Berthierine-chamosite ooids in Jurassic sediment showing distorted shapes and
elephantine features in calcite cement, plane polarized light.
Fig. 3.8: Berthierine-chamosite ooids in Jurassic sediment with some shape distortion, in
a siderite cement partly altered to goethite (brown), plane polarized light.
Greenalite is a hydrous ferrous silicate interbedded with chert and constitutes beds and
lenses in Precambrian sedimentary iron deposits. It occurs as rounded to subangular
pellets, with little internal structure.
Galuconite is a K-Fe aluminosilicate with a high Fe3+/Fe3+ ratio. It occurs typically as a
light to dark green pellets and aggregates up to 1 mm in diameter (Fig. 3.9).
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Fig. 3.9: Glauconite (four green grains) set in iron-rich dolomite cement of a sandstone.
Galuconite is characterized by green color, microcrystalline nature, PPL.
Glauconite occurs in many sandstones and may be a major constituent forming the
greensands. It is of great significance to the Bir Fa‘as Formation of the Early Cretaceous
Kurnub Group of Jordan since it is employed through K/Ar radiometric dating to
constrain the Aptian Age of this formation.
3.5 Precambrian iron-formations
From studies of iron formations in Canada, two group types are recognized:
1- Algoma-type deposits which are lenticular, relatively thin and narrow across strike,
usually closely associated with volcanic rocks and greywackes. They are Archaean in age
(2500-3000 Ma).
2- Superior-type deposits that are thicker and much more regionally extensive, deposited
on stable shelves and broad basins. They are of early middle Proterozoic age (1900-2500
Ma).
Four sedimentary facies are recognized within these Precambrian ironstone deposits:
oxides (hematite, magnetite), silicate (greenalite), carbonate (siderite) and sulphide
(pyrite) facies.
All these ironstone deposits are characterized by distinctive bedding structure consisting
of alternating beds or laminae or bands of ironstone (hematite, or magnetite, and less
commonly siderite or greenalite) with chert (Figs. 3.10, 3.11). Thus these deposits are
called banded iron formations (BIF).
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Fig. 3.10: A polished hand-specimen of banded iron formation (BIF) showing fine
lamination of light colored hematitic chert laminae alternating with darker hematite
laminae.
Fig. 3.11: Thin section illustrating microcrystalline quartz band (chert band) underlain by
a hematite band and grading upwardly into hematitic chert band, crossed polarized light.
In certain cases, the bands can be traced over 30,000 km2. Such large aerial distribution
indicated that the iron facies were deposited in deep-water basins, on shelves below wave
base, and in lagoons. The lamination indicates seasonal changes in the depositional
environment.
What was the source of iron employed in deposition of the BIF? A volcanic source has
been suggested, particularly for the Archaean Algoma-types which are clearly associated
with contemporaneous volcanic rocks. The iron may derive from hydrothermal vents on
the sea floor. The REE and Nd isotopes data suggest a hydrothermal source for BIF.
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On the other hand, for many Proterozoic iron formations, there appears to have been no
contemporaneous volcanicity, so that deep weathering of continental rocks probably
provided the iron.
To facilitate leaching of iron and its transportation, it is suggested that the Earth‘s
atmosphere had a higher carbon dioxide content at that time, and little or no oxygen. A
higher partial pressure of CO2 would have the effect of lowering the pH of surface
waters, leading to greater efficiency in iron leaching and transportation. Sea water with a
lower pH could itself have been a major reservoir of Fe2+ in solution.
3.6 Phanerozoic ironstones
Phanerozoic ironstones vary in grade, lithology and iron minerals present. The most
important are the oolitic ironstone, which generally are hematite-chamosite in the
Paleozoic and goethite-berthierine in the Mesozoic. Less significant ones are sideritc
mudstones and sulphidic ironstones.
The oolitic ironstone formed mainly in the Ordovician Period and the Jurassic Period.
These times were both intervals of global high sea level and warm humid climate
dominating over the continents which facilitated intensive chemical weathering.
Sea level changes are important in controlling the amount of iron that can be supplied to
the marine environment by reworking terrestrial weathering products.
Oolitic hematitic limestone occurs in Bir Fa‘as Formation of the Early Cretaceous
Kurnub Group of Jordan.
3.7 Recent iron deposits
The only modern environment where significant iron ores are forming at present time is
in swamps and lakes of the mid- to high latitudes such as northern America, Europe and
Asia.
The ores range from hard oolitic, pisolitic and concretionary forms to earthy and soft
types. The iron ores consist mainly of goethite and less common siderite and vivianite
(Fe3P2)8.H2O). Manganese oxides constitute a few per cents, but may attain in certain
cases up to 40%.
3.8 Ferromagnesian nodules and crusts, and metalliferous sediments
Ferromanganese nodules, crusts and metalliferous sediment occur on sea floors,
particularly, Atlantic, Pacific and Indian oceans.
Nodules of manganese and iron form in areas of low sedimentation, where there are
strong bottom currents, and at depths of several thousand meters. In oceanic settings, they
commonly form on the flanks of active mid ocean ridges, on seamounts and abyssal
plains.
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Fe-Mn nodules vary considerably in chemistry and mineralogy. Both Mn-rich-Fe-poor
and Mn-poor-Fe-rich varieties occur. Generally, they are rich in the metals: Co, Ni, Cu,
Cr and V (Tab. 3.3).
Tab. 3.3: Average concentration of Fe, Mn, Cu, Co and Ni (in percentages) in shallow
and deep water sediments and ferromanganese nodules from three sea-floor settings.
According to this concentration of base metals, submarine mining of these deposits is
now considered a viability.
Although the minerals constituting the nodules are X-ray amorphous, todorokite (Mn2+R+
R2+)(Mn4+Mn2+)O6.3H2O (where R represents other metals) is one manganese oxide that
commonly occurs, together with the hydrated iron oxide goethite.
The origin of ferromanganese nodules is attributed to hydrothermal-volcanic activities,
since they are associated with mid-oceanic ridge volcanism. But in some cases, the
nodules are present in areas far away from volcanism, so that some sort of direct or
indirect precipitation from sea water is needed.
The majority of ferromanganese nodules form at slow rates of around 1mm per 106 years.
In addition to nodules and crusts, metalliferous sediments also occur in the vicinity of
oceanic active spreading ridges at top of ocean floor basalts. These sediments are rich
with Fe, Mn, Cu, Pb, Zn, Ni, Co, Cr and V.
The fluids causing this metal enrichment are thought to derive from mantle magmatic
sources or from the interaction of basalt with sea water.
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