Formations of Banded Iron in World | Metals | Metallurgy

In this article we will discuss about:- 1. Introduction to Formations of Banded Iron 2. Geologic Setting and Depositional Environments of Banded Iron-Formations 3. Sedimentary Facies 4. World Distribution.

Introduction to Formations of Banded Iron:

Estimated iron ore resources in the world by the end of 2008 are about over 8,00, 000 MT. (MCS, 2008). Iron ore deposits are distributed in different regions of the world under varied geological conditions and in different geological formations. The largest concentration of iron ores is formed in banded sedimentary iron formations in the Precambrian age.

These formations constitute the bulk of world’s iron ore resources. The top ten countries in the world as per their iron ore deposits are the Commonwealth of Independent States (CIS), who belonged to the former Soviet Union, China, Brazil, Australia, USA, India, Sweden, Canada, South Africa and Venezuela.

Among the identified resources about 15 per cent are classified as Economic Demonstrated Resources (EDR) by the US Bureau of Mines. The EDR though small, when compared with the total identified resources are sufficient to sustain present products for more than 160 years.

The Precambrian iron-formations, of which banded ferruginous or iron cherts and jaspilites are the most important component, are known on all the continents. The rocks constituting these formations are very diverse in lithologic composition and origin: from typical volcanic (both basic and acid) to purely sedimentary (clastic, chemogenic). These iron-formations are usually related to complexly deformed geosynclinal complexes -eugeosynclinal or miogeosynclinal – although they are also encountered in slightly altered sediments on all platforms.

The largest areas where Precambrian iron-formations are developed are the East European platform: e.g. Krivoy Rog, Kremenchug, Belozerka, and other districts of the Ukrainian shield, the Voronezh, Baltic, and Central Kazakhstan shields, the Lake Superior area on the North American platform, Minas Gerais in Brazil, Hamersley in Australia, Singhbum-Orissa and Dharwar Region in India, etc.

 

BIF have been located in the Precambrians beginning with the oldest strata, 3,500 m.y. old. The rocks of the Konka-Belozerka zone of the Ukrainian shield and of the Pilbara Block of the Australian platform have an age of 2,700-3,500 m.y. An age of 2,000 m.y. has been established for the iron cherts of the Bazaviuk zone of the Ukrainian shield, the Olenogorsk deposit on the Baltic shield, the Singhbum, Orissa deposit in India, and the Algoma-type BIF of the Lake Superior area.

The rocks of the Krivoy Rog – Kremenchug zone, of the Kursk group in the Kursk Magnetic Anomaly (KMA), of the Animikie group of Lake Superior, and of the Hamersley formation in Australia are dated to 1,700-2,000 m.y. Younger BIF are also known, including post-Precambrian ones, related chiefly to volcanogenic rocks.

There exists no accepted classification based on lithologic association of these rocks, though four types of geosynclinal cherty iron-formations were identified Semenenko et al., (1967). Gross (1961,1965, 1973) distinguishes a sedimentary cherty iron-formation of Lake Superior type and a volcanogenic cherty iron-formation of Algoma type which includes the keratophyre, metabasic, and ultrabasic formations.

Geologic Setting and Depositional Environments of Banded Iron-Formations:

An examination of the main areas of occurrences Precambrian BIF of different formational types has been summarized in the following keeping the general regularities of distribution of the iron-form and to their age, spatial, and genetic relationships.

Despite many common features such as banding, mineral assemblages, sedimentary fades and chemical composition, iron-formations occur in quite different sedimentary environments.

Most of the Algoma-type iron-formations are found in Archean eugeosynclinal basins, whereas most of the Middle Precambrian Superior-type iron-formations are found most commonly on the margins of stable continental shelves in miogeosynclinal environments of intercratonic and cratonic basins.

Archean Basins:

The most complete description of Archean tectonic basins of the Canadian Shield is given by Goodwin and Shklanka (1967) and Goodwin (1973). Ten major Archean basins have been recognized in the Superior and Churchill provinces and, to a more limited extent, also in the Slave Province of the Canadian Shield, mainly identified on the basis of iron-formations, which are typically associated with greenstone belts within the Archean basins.

The elliptical basins are, in their present structurally deformed state, about 350-700 km long and are thought to represent remnants of originally quasi-circular structures in the tectonically mobile Archean crust, with diameters of between 800 and 1100 km.

Similar large basins have been recognized on shield areas elsewhere in the world, in Australia and in the Rhodesian and Kaapvaal Cratons of South Africa.

Earlier Kloostennan (1973) has described several giant “ring volcanoes”, with diameters of between 300 and 900 km, from the Amazon part of the Guiana Shield, which, however, have a younger age and were emplaced between 1,900 and 1,800 m.y. ago. The relationship of some known iron-formation deposits with these structures is not yet clear.

It was proposed by Goodwin (1973) that Archean basins of the Canadian Shield represent centres of crustal spreading in areas where vertical thermal streams (“thermal plumes” or “hot spots”) have been particularly active in the thin Archean crust. A meteorite impact scar theory is considered unlikely.

The wide distribution of basins of this type in outcropping and un-eroded Archean shield areas suggests that Archean basins have to be considered as first-order features of the crustal evolution in Early Precambrian time and perhaps as the ancient counterparts of modern oceans.

The development of the present oceans could well have originated from zones of fragility in the Precambrian crust, indicated by the distribution of many of the major iron-formation bearing old geosynclines, now broken apart by “continental drift” in the course of younger plate-tectonic movements.

The Michipicoten area of north-central Ontario was quoted by Goodwin (1973) as a typical example of such volcano-tectonic Archean basins. It encompasses several greenstone belts and intervening batholithic complexes. Arc-type volcanics with effusions and felsic pyroclastics, coarse-grained clastics, including graywackes and thick conglomerates, and iron-formations from oxide to sulfide stains are the most important lithologic associations of the marginal parts of such basins.

The distribution of iron-formations within the ambient volcanic sequences in the greenstone belts follows the normal depositional pattern, with oxide-, carbonate- and sulfide-facies in that order from the marginal to the inner parts of a basin. Oxide-facies is by far the most common facies.

Carbonate- and sulfide-facies iron-formations are equally widely developed but, in general, form only thin and discontinuous lenticular bodies of lesser significance. Commonly, iron-formations of this type are closely related to pyroclastic members of mafic to felsic volcanic sequences and shale-graywacke successions. Thus, there is common agreement to attribute the derivation of these iron-formations to volcanic activities.

Figures 2.1 and 2.2 taken from Goodwin (1973), demonstrate with a stratigraphic section of the Michipicoten basin, Ontario, the facies transitions from the shelf to the inner parts in a typical Archean basin.

Contrary to the Archean iron-formations, which are closely associated with volcanic successions in giant interarc basins of the Archean basement, the bulk of the major Middle to Early Proterozoic iron-formations was deposited in miogeosynclinal or epicontinental environments, together with clastic sequences, transgressive on the older basement. Oxide facies is the most abundant type and it is commonly not associated with volcanic rocks. These iron-formations of the Superior-type are enormous sheet deposits with thicknesses of up to several hundred metres.

They are persistent over many hundreds of kilometres to more than a thousand kilometres in extent and several hundreds of kilometres in width. The major iron-formations with an age bracketed between 2,600 and 1,800 m.y., many of them probably coeval at about 2,000 m.y., are prominent rock units of the shelf parts of elongated, large miogeosynclinal or intercratonic basins along the Archean shield areas, which they commonly overlie with profound angular and erosional unconformity.

The dimensions of these Proterozoic troughs, with extents of not much less than 1,000 km to more than 1,000 km, widths in the order of 100 km to several hundreds of kilometres and sedimentary thicknesses of several thousand metres, are comparable worldwide as Pflug (1967) has pointed out in his comparison of Precambrian miogeosynclines. Thus, it appears that this-kind of depositional environment for iron-formations represent a characteristic feature of the Precambrian evolution about 2,100-2,000 m.y. ago.

Prominent examples of the wide lateral distribution of continuous iron- formations are the Labrador Trough and the occurrences of the Lake Superior region, which correspond in age and other features and which can be followed over a distance of about 3,200 km along an old shoreline bordering the Canadian Shield.

The sedimentary basin of the itabirites in Minas Gerais, Brazil, had a minimum extent of 500 km and a minimum width of 160 km.

Iron-formations of the Superior-type are, however, not restricted to miogeosynclinal or platform environments in sedimentary troughs between older cratons. They are also frequently found in younger depressions on Archean cratons (intracratonic basins) such as the Witwatersrand, Transvaal and other iron- formations as described by Beukes (1973) from South Africa. Many of these basins, generally, but not always, are lacking contemporaneous volcanism.

They reflect conditions of restricted basins with deposition of iron-formation in shallow and quiet water, sometimes under highly saline conditions. This is indicated, for example, in the Kuruman and Penge Iron-Formations by locally high sodium concentrations.

The deposition of the iron-formation in the Hamersley basin, Western Australia, an originally ovoid basin of 500 by 250 km, is believed by Trendall (1973) to have taken place under essentially quiescent sedimentary conditions in water depths of 50-200 m.

The remarkable continuity of micro-banding over almost 300 km for single bands resembles strongly younger evaporite deposits and suggests deposition in a restricted basin free from currents and wave activity, similar to evaporitic formations.

In detailed studies of iron-formations of the Labrador trough, Quebec, Dimroth (1972) recognizes two main depositional facies: a lagoonal platform facies of thin-bedded oolitic, granular and conglomerate alternations with chert, deposited in water probably less than 10-15 m deep and a basin facies deposited at less than 100 in (perhaps less than 50 m) depth. Presumable pseudomorphs of chalcedony after gypsum suggest local evaporitic conditions and deposition in a probably warm and dry climate.

In opposition to this, some other Late Proterozoic iron-formations appear to have been deposited under cold-climate conditions. In the Northern Hemisphere the iron-formation of the Rapitan Group in the Mackenzie District of northwest Canada, deposited between 800-600 m.y. ago, alternates with clastics of probably glacial origin.

In the Southern Hemisphere, similarly, Martin (1965) described glacio-marine associations with iron-formations in both miogeosynclinal and eugeosynclinal environments of the Dainara Super-group of South West Africa, the age of which is bracketed between 620-1,000 m.y. ago.

It is of interest that these two roughly contemporaneous iron-formations on opposite sides of the globe reflect similar glacial depositional conditions. (The conclusion arises that they have been located in the pen-glacial outskirts of extensive continental glaciers or perhaps of the then polar ice caps of the ancient globe).

However, cold-climate sedimentary conditions for deposition of iron-formations in general cannot be considered as a crucial factor, since most of the major Superior- type iron-formations, grouped in the interval between 2,600-1,800 m.y. ago, have been deposited contemporaneously around the world, including the equatorial zones. Hence, different climatic factors appear to be of basically little significance for the deposition of iron-formation.

Iron-formations in Proterozoic miogeosynclines and post-orogenic intracratonic basins are major and extensive rock units. They represent a characteristic facies laid down during a certain phase in the evolution of Proterozoic geosynclines under relatively shallow-water epicontinental or shelf conditions.

The typical lithofacies associations of these miogeosynclinal rock sequences are essentially thick coarse-to fine-grained quartzites with conglomerates and pelitic interruptions in the marginal and bottom parts of the transgressive sequence and predominantly pelitic rocks, sometimes graphitic shales, with local intercalations of quartaites, limestones or dolomitic rocks in the higher parts.

Iron-formation, frequently in several horizons interrupted by shales and/or quartzites, was normally laid down in the lower and middle parts of the section. With thicknesses of several hundreds of metres, it may make up as much as 10 per cent of the total basin filling.

In general, these clastic-chemical sequences are devoid of any volcanic associations, in opposition to Archean banded iron-formations. However, volcanic activity contemporaneous with deposition of iron-formation may also be observed, presumably merely parallelling the clastic-chemical deposition in the course of the structural development of the basin and without any direct connection with the origin of the iron-formation.

In many major iron areas, for example, the Lake Superior region or in the “Quadrilatero Ferrifero” of Brazil, it can be observed that the main iron-formation of the miogeosynclinal sequence was laid down in the lower part, underlain and overlain by non-volcanic clastic rocks.

Pflug (1965) who studied the marginal parts in the Diamantina area, Dorr II (1973) in the central area of the basin investigations in the southeastern part of the “Iron Ore Quadrangle”, where several features indicate an offshore uplift zone, which can be traced further to the north, parallel to the old “So Francisco Craton”, as an island-arc zone restricting the miogeosynclinal basin.

The thickness of the Caue Itabirite varies from about 50 m to more than 600 m with an average of about 250 m. The principal iron-formation covers an area of about 160 km lateral width and can be traced over an extent of at least 500 km in a NNE – SSW direction.

The diagrammatic section shows a transgressive overlap of lower Minas Series quartzites and phyllites on the pre-Minas basement (Rio das Velhas Series; plus 2,700 m.y.) and the successive deposition of the Caua-Itabirite in an at least partly barred and continuously subsiding basin.

Increasing subsidence stopped deposition of itabirite and led subsequently to eugeosynclinal sedimentary conditions with thick sequences of shales (phyllites) and graywackes with local quartzite wedges and dolomite lenses. Conglomerates and clastic hematite quartzites, probably reworked itabirites, are locally inter-fingering with the itabirites.

Thin and insignificant iron-formations of commonly lenticular shape may locally follow in the higher parts of the sections, embedded in eugeosynclinal sediments with or without adjacent rocks of volcanic origin. It is here suggested that these minor successive horizons represent stragglers of the vanishing stages of iron-formation deposition incidentally conflicting with volcanic material when volcanic activity started in the progressively subsiding basin.

It is likely that major depositional troughs have been divided into several separate shallow basins. This may be supposed when lithologic associations of single correlative iron-formations differ considerably from one to another in the same district as it is the case in the Menominee, Marquette, Gogebic, Mesabi and Gunflint ranges of the Lake Superior District (Figure 2.4 and Figure 2.5).

Summary of depositional types iron-formations. Relative to the principal environments of deposition, a grouping into two main types of banded iron-formation has been explicitly established by Gross (1965): the Algorna-type and the Lake Superior type. A comparison of the characteristic features of these two types is shown in Table 2.2.

The late Proterozoic iron-formations with an age of less than 1,600 m.y. show many features common with the bulk of the older iron-formations. However, there are some notable differences in their general geological setting. First of all, their regional distribution is much less abundant in time and in space and they are also less prominent units in the associated rock sequences.

The mineralogical and chemical composition seems to be more variable. Chert is normally observed but may also be lacking. In the silicate-facies, high-aluminous chamosite is more common than low- aluminous greenalite. In some deposits essentially higher contents of iron, far exceeding the normal range in banded iron-formations is present.

Thus, Gabrielse (1972) mentioned from the Snake River iron-formation, at the base of the coarse clastic Rapitan Group in northwest Canada, an average content of 46 per cent iron in the hematite-jasper. Dorr II (1973) describes from the Mato Grosso, Brazil the jaspilitic banded hematite deposit of Urucum, which has an average iron content of 56.9 per cent (range 48.7-62.1 per cent Fe) and is one of the highest grade sedimentary iron-formations in the world. The deposit is inter-bedded with detrital clastic material and, abnormally, with horizons of high-grade sedimentary manganese oxides.

Occurrences of Palaeozoic banded hematite jaspers with an average iron content of 56.5 per cent are known in the Himalayas of Nepal. The origin of these deposits, as well as those of most of the younger Phanerozoic banded iron-formations summarized by O’Rourke (1961), is in many respects not comparable with the older iron-formations and remains enigmatic.

Sedimentary Facies of Iron-Formations:

The development of different facies of iron-formation dependent on depositional environments is similar to that of Phanerozoic marine-sedimentary iron deposits of the Minette-type explicitly defined by Borchert (1960). As is commonly known, iron occurs in nature in different oxidation states, depending mainly on the environment of deposition.

The kind of chemically precipitated iron minerals is dependent principally on the pH and Eh conditions as the chief controlling factors of kind and amount of the precipitate, as has been demonstrated by Krumbein and Garrels (1952). Concentrations as well as changes of P/T conditions have no significant influence on the kind of the precipitates.

As stressed by Krumbein and Garrels (1952), the pH and Eh values in normal marine environments show only slight variations between pH 8.4/Eh + 0.4 near the surface and pH 7.5/Eh + 0.1 near the bottom.

For iron- formation facies reflecting reducing depositional conditions, James (1954) consequently suggested that iron-formations containing sulfide-facies have been deposited in basins which were partly cut off from the open sea by bars, thus allowing the formation of restricted conditions with strongly negative Eh for the deposition of iron-sulfides.

James (1954) and later Gross (1965) stressed the close relationship between composition, mineralogy and the oxidation state of chemically precipitated sediments and depositional environment, which allows a classification of distinctive facies types related to controlling physico-chemical factors.

Thus, James (1954) separated four distinctive facies of iron-formation, namely oxide, silicate, carbonate and sulfide facies, which were deposited in that order from shallow to deep water within the basin, in chemical environments with relatively high Eh (oxide facies) to strongly negative Eh (sulfide-facies) and intermediate Eh values in the carbonate and silicate facies.

This concept can be applied worldwide to all types of banded iron- formation, to the Algoma-type as well as to the Superior-type. Figure 2.7 illustrates the relative position of iron-formation facies in a restricted sedimentary basin.

It has to be noted that the facies are commonly inter-gradational and a complete successive pattern is, not developed in nature. Associations of carbonate-sulfide facies have already been described from the Michipicoten basin. Sulfide facies is practically absent in all major iron-formations and apparently confined to the Archean type. Dorr II (1973) mentions only a thin one in Gabon, Africa.

There is a remarkable difference between the great iron-formations of the Northern and the Southern Hemispheres as to their homogeneity and facies association as Dorr II (1973) has pointed out. Contrary to many major deposits in North America and Russia, in the iron-formations of the Southern Hemisphere, the carbonate and silicate facies are rare exceptions and oxide facies is by far the dominant type.

Transitions between hematite – magnetite and silicate – carbonate facies have been described by Gastil and Knowles (1960) from the Labrador trough. From the Gunflint Iron Formation, Minnesota, Goodwin (1956) describes lateral transitions between magnetite – silicate and carbonate facies.

World Distribution of Banded Iron Formation:

I. Europe:

On the East European platform, BIF occur in the western part, in the Ukrainian, Baltic, and Voronezh shields, and in the eastern part of the platform in Kazakhstan.

i. Iron-Formations of the Ukrainian Shield:

In the Ukrainian shield, BIF are developed in a number of synclinorium zones in different structural levels of the Precambrian. They are spatially unconnected, which makes it difficult to correlate them and establish their common genetic regularities. Usually several zones of development of BIF are distinguished.

The Konka – Belozerka synclinorium, which was stabilized in the first Precambrian cycle 2,700 m.y. ago, is considered to be the first and oldest zone. Rocks with an absolute age of 3,500 m.y. have been established in here. The cherty iron- metabasite and cherty iron-ultrabasic formations of the Lower Konka group were developed essentially at that time.

In the final stages of development of the zone, basic volcanism was superseded by acid volcanism and the rocks of the cherty iron- keratophyre formation of the Upper Konka group were deposited. According to Semenenko (1973) the rocks of this zone were metamorphosed to schists and hornstones and reach the rank of gneiss only in the marginal parts of the zone at the contact with granitoids.

The second zone in the Bazaviuk synclinorium, which began to developed in the period of 2,700-2,800 m.y. and stabilized 2,300 m.y. ago. In this zone the cherty iron deposits consist of a cherty iron-metabasite formation in the Lower Bazaviuk group and a cherty iron-keratophyre formation in the Upper Bazaviuk group. The rocks of this zone are diverse in degree of metamorphism: from phyllite to pyroxene-gneiss.

The third zone is confined to the Orekhov – Pavlograd synclinorium developed 2,300-1,800 m.y. ago. Semenenko (1973) correlates the Korsak and Mangush zones of the Pri-Azo area with this zone. Both cherty iron metabasite and cherty iron-gneiss formations are present in the Orekhov- Paviograd synclinorium, the former being limited stratigraphically to the lower part of the section and the latter of it. A feature of the iron cherts of this zone is more intensive metamorphism under amphibolite and granulite facies conditions. Less metamorphosed rocks occur only in individual synclines, for instance the Gulyaypole.

The fourth and main zone is the Krivoy Rog – Kremenchug, developed 1,800 – 2,000 m.y. ago. The accumulation of rocks began here, as in the other zones, with the eruption of meta-basic and ultrabasic rocks, which then probably were locally succeeded by rocks of the keratophyre series. Most of the iron cherts, however, are not related to the initial but rather to the middle stage of development of the zone, when the sedimentary cherty iron-formation of the Middle suite of the Krivoy Rog group was formed. Maximum development of iron cherts is characteristic of the Saksagan and Kremenchug basins.

No meta-igneous rocks have seen definitely established in the sections of the Middle suite, a fact which makes one to attribute the cherty iron strata to the remote type of formation or even to believe that the original material of the sediments was the product of sub-aerial weathering. Depending on the typical paragenetic associations of the rocks, Semenenko (1973) distinguishes these types of cherty iron-formations on the Ukrainian shield: cherty iron-schist (sedimentary), cherty iron-keratophyre (orjaspilite-leptite), cherty iron-metabasite and cherty iron- ultrabasite.

The cherty iron (sedimentary,) formation (CIS) is the most uniform in facies. The cherty iron beds are traced for 10-20 km; their thickness reaches 100 – 300 m and their iron content reaches 30-35 per cent. According to Semenenko et al (1959) these rocks were formed under conditions of prolonged hydrothermal activity of submarine volcanoes, with limited introduction of ashy material and widespread development of chemogenic deposits. On the Ukrainian shield the Krivoy Rog – Kremenchug cherty iron strata of the Middle suite and also the rocks of the Upper suite of the Upper Konka group of the Belozerka syncline are of this type.

The cherty iron-keratophyre formation (CIK) is characterized by the paragenetic association of acid dacite-rhyolite lavas, tuff-keratophyres, and cherty iron deposits. The formation is more uniform in facies. It is believed that the cherty iron sediments were formed near volcanic centers at times when volcanism was ending.

The cherty iron-metabasite (CIM) formation was formed in regions of submarine basaltic volcanism. The iron cherts in the formation are not uniform in facies and amount to 10-30 per cent of the total thickness of the sequence. The formation is extensively developed in the lower metabasic series of the Konka. Belozerka, Verkhovtsevo, and Sura synclines.

The cherty iron-ultrabasite formation (CIU) is not widespread; occurs in the Konka syncline and in the Kudashev sector of the Verkhovtsevo syncline are low- grade banded cherty iron intercalations of no great thickness inter-bedded with ultrabasic schists.

Other views are concerning the genetic classification of the cherty iron-formations of the Ukrainian shield and their spatial and temporal relationships. Kalyayev (1965) distinguishes three types of cherty iron-formation: sedimentary, volcanogenic- sedimentary, and the jaspilite situation within the spilite-diabase formation.

The sedimentary jaspilite formation, analogue which is the Krivoy Rog iron-ore sequence, is characterized by an association with jespilitic and psammitic sedimentary rocks whose sources were products of sub-aerial weathering. The sediments of this formation were deposited in the Krivoy Rog- Kremenchug marginal depression and also in individual parageosynclinal zones.

The volcanogenic-sedimentary jaspilite formation was formed in intra- geosynclinal synclinorium zones; its equivalent is the cherty iron-formation of the Belozerka syncline.

The jaspilite sub-formation developed in the spilite-diabase formation is known both in intrageosynclinal synclinorium zones and in the inner zones of the Krivoy Rog-Kremenchug marginal depression.

The interrelationships of these types of jaspilite formations are represented on the composite diagram and in the scheme of development of the Greater Krivoy Rog geosyncline, which suggest that the sedimentary-volcanogenic and sedimentary jaspilite formations are temporal analogs but were formed in different structural zones of the Greater Krivoy geosyncline: the former in the inner zones (Verkhovtsevo, Sura, Konka) and the latter in the outer zones of the marginal depressions.

The sources of the material in these types of formations also are different in the former the products of volcanism predominate, in the latter, the products of sub-aerial weathering, forming a Precambrian geosynclinal flysch. Thus all the cherty iron-formations belong the same geosynclinal cycle, to its pre-orogenic half, and the sedimentary and sedimentary- volcanogenic jaspilite formations were formed after deposition of the Jaspilite sub-formation of the diabase-spilite formation.

ii. Iron-Formations of the Voronezh Shield:

BIF occur in stratigraphic sequences of Voronezh shield. Layers of highly metamorphosed iron-rich rocks have been established in the old gneisses and migmatites of the Oboyan group, which is assign to the Archean (Figure 2.6).

In the Mikhaylonka group, which is correlated with the Verkhovtsevo series of the Ukrainian shield, iron cherts are inter-bedded with keratophyres and their tuffs and orthoamphibolites.

BIF are very extensively developed in the deposits of the Lower Proterozoic Kursk group, mainly in the middle suite. The iron formations are associated with metasediments-phyllites.

The iron cherts of the well-studied Mikhaylovka area of constitute several discontinuous layers forming two belts: a northeastern one 450- 500 km long and a northwestern one 600 km long. These belts skirt the arch uplift of the crystalline basement, as it were. The total thickness of the iron cherts of the Kursk group is of the order of 100-1200m. Iron cherts consisting of heavily landed and concretionary martite-hematite quartzites are also known among the rocks of the upper suite.

Plaksenko (1969) distinguishes these cherty iron-formations: cherty iron gneiss (Oboyan group), cherty iron-metabasite (Mikhaylovka group and lower suite of the Kursk group), cherty iron-schist (middle suite of the Kursk group and cherty iron- clastogenic (upper suite of the Kursk group). A volcanic source is assumed for the material of the cherty iron-metabasite formation; the formation of the cherty iron- schist formation is not connected with volcanism, but is explained by the arrival of material from the weathered layer on Archean iron-bearing rocks.

iii. Iron-Formations of the Baltic Shield:

On the Baltic shield, BIF are known in three major structural-facies zones:

1. The Karelian zone of the Karelides,

2. The Kola-Norwegian zone of the Karelides, and

3. The vast zone of the Svecofennides in the southwestern part of the shield.

The first two zones are separated by the Belomorides, which most investigators consider to be older than the Karelides. The question of the age relationships of the Svecofennides of Sweden and Finland to the Karelides is not settled yet.

In the Karelian zone of the Karelides, BIF are known in northern Finland in the spilite-diabase formation and are associated with basic volcanic rocks – pyroclastics, basaltic lavas and agglomerates – in northern Norway: graphite schists and limestones are occasionally encountered in the sections.

In the Kola – Norwegian zone of the Karelides, situated in the central part of the Kola peninsula and extending into northern Norway, iron cherts are associated with gneisses, schists of basic composition, amphibolites, and paracharnockites of the Kola group. In this sequence, which can be assigned to the volcanogenic-sedimentary formation of the spilite-diabase series, both meta-sedimentary and meta-igneous rocks occur.

In the Lake Imandra area the cherty iron-formations of the Kola group are associated with volcanic rocks of more acid composition, similar to the rocks of the porphyry-leptite formation of central Sweden. In the West Karelian zone iron cherts are restricted mainly to the Gimoly group of Lower Proterozoic age: 2,000-2,600 m.y.

The Gimoly group lies on the Archean basement of the Karelides and is overlain by the Parandovo and Bol’shoye Ozero group, where iron cherts also are known. The largest deposits of iron-formation-the Kosv and Mezhozero – are restricted to large synclinal structures preserved in the Proterozoic granites and migmatites.

The iron cherts of the upper sedimentary cycle are associated with the products of acid volcanism: lavas, tuffs. In the Syecofennides, BIF are known in central and northern Sweden (Kiruna area) and in southern Finland, also associated with volcanogenic rocks.

Formational group of the same type, similar to the Proterozoic formations of Karelia, are found in all these areas. Acid volcanic rocks of the porphyry-leptite formation, and sometimes carbonate rocks, are noted in paragenesis with iron cherts; there are no typical primary clastic formations.

In the Kiruna area the largest deposits of iron-formation and apatite- bearing iron ores occur in paragneisses with volcanic rocks of syenitic and rhyolitic composition and belong to the cherty iron-leptite porphyry formation correlate this formation at Kiruna with sequence of basic volcanics and iron cherts of the Bol’ Ozero group, which lies with angular unconformity on the leptite-porphyry chertv iron- formation of the Ginioly group and constitutes the upper part of the Lower Proterozoic section of Karelia.

Despite the diversity of the BIF and the complexity of their interrelationships with the enclosing rocks, some general features have been established. In a detailed formational analysis Chernov showed that the cherty iron- formations constitute the lower parts of the sections both in the Karelides and in the Svecofennides, and that their formation reflects the initial stages of geosynclinal development of the Baltic shield.

In many areas the BIF are underlain by thick piles of conglomerates which rest on an old basement more than 2,600 m.y. old. The BIF are overlain by volcanic rocks of the spilite-diabase series or by clastic flyschoid formations formed in the final stages of geosynclinal development.

iv. Iron-Formations of Central Kazakhstan:

Precambrian iron occupies large structures of the area – the Karsakpay and Betpak-Dala belts, the west limb of the Maytyube anticlinorium, etc.

In the Karsakpay synclinorium the BIF are restricted to the Karsakpay group of Early Proterozoic age. The rocks of the Karsakpay group constitute a narrow north- south belt extending almost uninterruptedly for 200 km. Four macrorhythms, united into individual suites, are distinguished in this series.

These macrorhythms reflect cyclicity in the manifestation of basaltic volcanism: in each rhythm the lower part of the suite consists of porphyritoids after tuffs and lavas of basaltic, less often of andesitic, composition, and the upper part consists of quartz-sericite schists, marbles, and iron-rich rocks.

In the Maytyube anticlinorium iron cherts are known among the deposits of the Satan suite, which consists mainly of clastic formations: mica schists, phyllites, conglomerates. The Satan suite lies stratigraphically above the Karsakpay series and belongs to the Upper Riphean or Lower Vendian.

Among the iron cherts of the Karsakpay synclinorium, three types of sequences and correspondingly three types of formations are identified.

The association of BIF with basic volcanic rocks is typical of the first type; it can be correlated with the jaspilite formation of the Algoma series in the Lake Superior district, which was formed in the greenstone formation. The rocks of the Verkhovtsevo group of the Ukrainian shield are also considered analogs.

The second type is distinguished by the association BIF with meta-sedimentary rocks-the analog of the jaspilite formation of the Saksagan group of the Krivoy Rog and the group of Canada; this formation follows the greenstone and jaspilite formations of Keewatin type in time.

The third type is characterized by the association of thin and rapidly-pinching-out iron cherts with clastic sediments, carbonate rocks and carbonaceous shales. These deposits belong to the remote jaspilite formation, developed either at the same time as the jaspilite formation of Krivoy Rog type or after it (analog of the rocks of the upper suite of the Saksagan district).

II. North America:

The main regions where Precambrian cherty iron-formations occur are on the Canadian shield, chiefly in the southwestern part of it.

The oldest iron cherts (2,300-2,700 m.y., Kenoran orogeny) are spatially and genetically related to the volcanogenic greenstones of the Algoma group. On this basis Gross (1973) distinguished the volcanogenic cherty iron- formation of Algorna type. The rocks of this formation are extensively developed in the Michipicoten region (Churchill, Helen, and other districts).

In the Churchill district iron cherts are associated with metadiorites, andesites, dacite breccias and tuffs, carbonate rocks, and graywackes. Also typical of the Helen district is the relationship of the cherty iron sediments to volcanism, and a relationship has been established between the composition of the sediments, their location in the sedimentary basin, and the amount of volcanic products (Figure 2.2).

The facies details shown in the sketch reflect the regularities of deep deposition of volcanic derivatives on the slopes of the original sedimentary basin. The BIF which occur in the volcanic rocks are a combination of three facies-banded chert, a sulfide member, and a carbonate member (Figure 2.7).

The banded chert member consists of alternating bands of cherty material and iron minerals, chiefly siderite with magnetite, pyrite, and pyrrhotite. It gradually grades into lenticular sulfide members which in turn also are gradually replaced by the underlying carbonate member.

In the western part of the district a sedimentary association is developed, consisting of alternating bands of chert, magnetite, and jasper, which occur in gray wackes and argillites. Thus the thick oxide facies of the western part of the district, rich in cherty material, gradually changes eastward into a thick carbonate facies containing cherty material in the center, and then into a thin sulfide facies poor in cherty material.

An interesting BIF of Algoma type is found in northwestern Canada on the Yukon-Mackenzie divide. Traces of mud flows, tuffs, ash, conglomerates, and syngenetic breccias i.e., products testifying to deposition in a tectonically active basin with explosive volcanic activity have been established in these cherty iron- formations.

The iron cherts of the Animikie group, extensively developed in the Lake Superior district (Figures 2.7a & b) in the states of Minnesota (Cuyuna, Mesabi, and Gunflint belts), Michigan and Wisconsin (Marquette, Gogebic, Menominee, and other belts), are younger (Huronian orogeny, consolidation 1,700-1,900 m.y. ago).

The rocks of the Animikie group lie on rocks of the Algoma group and are overlain by quartzites and greenstones of the Keweenawan group (see Figure 2.7a & b). Within the Animikie group itself, iron-rich and barren quartzites, paraschists. tuffs, breccias, and dolomites are interbedded and succeed one another in facies.

Besides the Lake Superior district, analogous rock associations occur in the Labrador trough, in the eastern part of the Canadian Shield. Gross assigned these iron cherts to the Superior type, characterized by the development of unstratified cherty and iron-rich deposits with granular or oolitic texture.

Sediments of Superior type were formed in the late Precambrian in miogeosynclinal conditions of a continental shelf, succeeding a volcanogenic cherty iron-formation of Algoma type, deposited in a eugeosynclinal setting.

III. South America:

Deposits of Precambrian BIF, rich in iron ores of Cerro Bolivar in Venezuela and in the states of Para and Minas Gerais in Brazil, are known. In the Sierra das Carajas area in Para, iron cherts were found not long ago in the rather inaccessible forests of the Amazon basin, and are not well known.

In Minas Gerais there is a large deposit; the iron cherts form a 500 km belt on the edge of the Brazilian shield. The Rio das Velhas group, the age of which is older than 2,800-3,000 m.y., and the Minas Gerais group, 1,800-1,300 m.y. old and lying uncomfortably on the gneisses of the Rio das Veihas group, are distinguished.

In the Rio das Veihas group iron cherts form thin lenticular bodies in Archean gneisses. The association with meta-volcanic rocks and graywackes is typical of them.

The Minas Gerais group is divided into the lower and upper clastic strata, and also the middle chemogenic Itabira strata. The BIF are traced for a distance of about 200 km and fill a deep symmetrical syncii which plunges to the northeast. The iron-formations consist of unbanded ferruginous varieties and thin-banded hematite-carbonate itabirites.

IV. Indian Platform:

On the Indian platform, iron cherts constitute part of the Dharwar system, 2000- 2700 m.y. old, which forms extensive synclinal zones and belts on the periphery of the Indian shield.

The iron cherts are of two main types:

1. Thin-banded hematitic jaspilites and

2. Banded magnetitic cherts.

Jaspilites constitute the thick Singhbhum strata (Jharkhand and Orissa), the Drug, Bastar (Chhattisgarh) and Goa, Karnataka deposits. In the Singhbhum deposit the banded hematite rocks are underlain by shales, calcareous sandstones, conglomerates, and phyllites, beneath which lie tuffs and basic lavas.

The total thickness of the formation is 2,000-2,500m, up to 1,000m of which is jaspilite. A very low degree of metamorphism is typical of the ore minerals, hematite predominates, magnetite is rare, and silica is sometimes represented by opal.

The age and genetic relationships of the hematite and magnetite rocks found in Goa, Chhattisgarh, and other areas are unknown. Krishnan (1982) considers the magnetite rocks to be the more metamorphosed analogs of the hematite rocks and unites them both into the Dharwar formation. However, ideas exist that the magnetitic rocks are older. Volcanism presumably is the source of the iron of the Precambrian BIF of India.

V. Africa:

On the African platform, iron cherts are known in several structural positions which, however, are not unequivocally distinguished. The oldest iron-formations (2900-3400 m.y.) are restricted to greenstone belts and consist of banded chert, banded ferruginous chert, and BIF proper.

These rocks are intimately associated with cycles of volcanic activity and are constituent of them. In ultrabasic rocks of the greenstone belt, the BIF may inter-finger along the strike with various types of rocks in one and the same cycle. Under these conditions the cherts and BIF consist of lenticular bodies, and several lenses may be restricted to one stratigraphic zone. Analogous structure is observed in basic and acid rocks, but the cycles are more completely developed.

The younger Kuruman and Penge iron-formations are component parts of a more prolonged (between 1,950 and 2,300 m) and more important period of chemical deposition in the basin of the Transvaal system. The BIF are underlain by dolomites in the transition zone there are inter-bedded carbonaceous limestones, banded ferruginous cherts, and carbonaceous- argillaceous shales with pyrite, that is, typical carbonate shelf deposits.

VI. Australia:

Iron cherts are known in three structural positions on the Australian platform. The oldest are in Western Australia in the Pilbara and Yilgarn Blocks, which were stabilized 2,700-3,000 m.y. ago. The iron cherts are not spatially persistent and not very thick.

The second group (1,700-2,000 m.y.) constitutes the Hamersley BIF in Western Australia, which occur in a syncline between old blocks. The thick strata of the formation (about 2700m) consist of iron-rich rocks inter-bedded with clay shales, dolomites, lavas, and tuffs.

The iron-formation is underlain by conglomerates, metabasites, pyroclastics, and pillow lavas, and is overlain by quartzites, conglomerates, graywackes, and dolomites. Slight metamorphism, the temperature of which corresponds to the normal geothermal gradient, is typical of the cherty iron-formations.

The third group consists of the BIF of the Cleve Metamorphics (South Australia), with an age of about 1,780 m.y. A direct relationship to volcanic processes has not been established for the cherty iron-formations of Australia, but Trendall (1973) suggests that the source of the ore material was rhythmic volcanism.