chapter ii definition and na1enclature...
TRANSCRIPT
CHAPTER II
THE PRECAMBRIAN BANDED IRON FORMATION : A BACKDROP
Contents Page
DEFINITION AND Na1ENCLATURE 20
GEOGRAPHICAL DISTRIBUTION 23
DISTRIBUTION IN TIME AND SPACE 26
LITHOLOGY 30
STRATIGRAPHY 34
ClASSIFICATION 42
MINERALOGY 44
CHEMICAL COMPOSITION 49
ORIGIN 67
SOURCE 69
TRANSPORT 76
DEPOSITION 84
Depositional Environment 85
Depositional Process 87
Mode of Deposition & Banding 97
Primary Deposition & Facies 102
POST DEPOSITIONAL CHANGES 107
DIAGENESIS 107
METAMORPHISM 108
WEATHERING 117
20
DEFINITION AND NOMENCLATURE
Iron-formations (IF) are thought to be chemical sedi
mentary rocks consisting essentially of silica and iron in the
form of oxides, carbonates, silicates and sulphides. A great
majority of the known IF are mostly confined to the Precambrian,
which are also well banded with alternating layers of chert/
quartz and iron minerals. Therefore, they are commonly called
as Banded Iron-Formations (BIF).
BIF exhibit diverse physico-chemical characteristics
which make any formulation of exact definition difficult. Att
empts, however, have been made in that direction. James (1954)
has defined iron formation as "a chemical sediment, typically
thin bedded or laminated, containing 15 per cent or more iron
of sedimentary origin, commonly but not necessarily containing
a layer of chert". Gross (1965) has used the term in a more
general way for all stratigraphic units of layered or laminated
rocks that contained 15 per cent or more iron, iron minerals
being interbanded with chert and the banded structure being in
conformity with the banded structure of the adjacent. supracrustal
rock or their metamorphic equivalent. The weakness or limita
tions of above definitions are as follows:
i) The sedimentary rocks defined as banded ferruginous
quartzite and banded chert although contain iron far less than
15 per cent seem to have a genetic relationship with typical
21
banded iron-formation and a gradation in properties from the ban-
ded iron-formation to the banded ferruginous quartzite to the
banded chert is well established (Beukes, 1973). Hence, 15 per
shales
for the definition although they are a part of the iron-fornation
and considered as its sulfide facies. Similarly, many non-cherty,
sulfidic and carbonaceous sediments are considered to be an inte-
gral part of the iron-formation although they do not come under
the definition (Goodwin, 1973).
Itt .;53·31'
55 1•71/'7'l iii) As far as the mineralogy and the iron content are con-
B39'> cerned there exist not much difference between the iron-forma- ~ tion and younger iron-stone and the above definition never dis-
ting11ishes one from the other \Goodwin, 1973). Considering the
wide range of rock types with diversified properties that go into
one definition, it is felt unnecessary to try for an exact one.
Roughly it can be presumed to b:= a sedimentary rock either from
a chemical or from a biochemical precipitate of Precambriaf! origin,
mostly banded or laminated, mostly a chert band alternating with
a band of iron minerals, and containing almost all elements
in trace quantities except iron and silica.
Iron-formations are called differently in different parts
22
of the world. Wide diversity in character and different degree
of metamorphism have made the task to evolve a common scheme of
nomenclature more difficult. An adhoc committee on nomenclature
was formed at Kiev Symposium which made a good attempt in that
direction (Brandt et al, 1973). Iron-formation (IF) or Banded
iron-formation (BIF) is the most common name used in USA, Canada,
Australia and South America, and is mostly used in a lithologic
sense. Wnen the words are capitalized and nonhyphenated
like Iron Formation, it implies the corresponding stratigraphic
unit (Brandt et al, 1973; Kimberley, 1978).
The oxide facies of the iron-formations are known as
Jaspilite in USA and USSR, as banded iron stone in South Africa,
as Itabirite in South America and Western Africa, and Banded
hematite quartzite (BHQ) and Banded magnetic quartzite (BMQ) in
India.
The term banded ferruginous quartzite (BFQ) is used for
iron-formation when iron percentage is less than 15 per cent
but its use in USSR is generally for oxide and silicate facies.
The term banded chert is used when iron per cent is negligibly
small (Beukes, 1973).
In addition to these, other local names in which the
iron-formation is also known are ferruginous chert, banded hema
tite jasper, Jasper bar, calico rock, and zebra rock (Eichler, 1976).
23
GEOGRAPHICAL DISTRIBUTION
Iron-formations are widely distributed and occur in
almost all continents of the world. They are mostly confined
to the Precambrian Shield Areas and occupy a prominent place
in their stratigraphy:
In North America, it is found both in the Canadian
Shield and in USA. Abitibi Volcanic segment, Swayze Area, f
Gogma Area of Michipicoten, Lake Superior Area and Kirkland
Lake Area are prominent fields of iron-formation on the Cana-
dian Shield (Goodwin, 1973). The major iron-formation of the
United States are confined to the Lake Superior region, which
is the southern most margin of the exposed Canadian Shield, in
the districts of Minnesota, Wisconsin, Michigan, Menominee,
Marquette, Gogebic,Mesabi and Gunflint. It is also found in
the northern Rocky mountains where the Precambrian rocks form
the core of the ranges in Hontana, Wyoming and South Dakota.
Few deposits of younger age are also found in the south western
states of Arizona, Colorado and New Mexico (Bayley and James,
1973).
Most widely known iron formations in South America
occur in the states of Minas Gerais, Brazil. It is also found
as a great deposit in Serra dos Carajas of Brazil, and Morro do
Urucum and Serrania de Mutun of Brazil-Bolivia border. Other
iron-formations of South America are the Imataca complex of
Venezuela, and few deposits of Chile and Uruguay (Dorr, 1973).
Most of the iron-formations of Africa are confined to
South Africa and classified into four tectono-sedimentary units
such as (i) the greenstone belts of the Kaap~al and Rhodesian
cratons, (ii) the Limpopo metamorphic belt, (iii) the cratonic
basin of Pongola, Witwatersrand and Trans~al Supergroups, and
(iv) the Darnara Mobile belt (Beukes, 1973). It is also found as
Ijil Group, Mauritania and in Liberian Shield of West Africa
(James, 1973).
Major iron-formations of Australia are: the iron-forma
tions of Yilgarn Block and Pilbara Block of Western Australia,
the iron-formation of the fiamersley Basin of Western Australia,
iron-formations of the Cleve Metamorphics of South Australia,
hematite rich sediments of the Yarnpi Sound area of Western Aus
tralia, the Roper Bar and Constance Range iron-forrr~tions of the
Northern Territory and Queensland and the Holowilena iron-forma
tions and Braernar iron-formations of South Australia (Trendall,
1973).
Iron-formations of Europe and Asia are mostly confined
to USSR and India. In USSR, it occurs mainly along a trend which
can be followed between the coastal regions of the Azou Sea in
the south and the Kola Peninsula in the north. The deposits of
Krivoy Rog, Krurnenchung within Ukrainian Shield and those of the
Kursk Magnetic Anomalies (KMA) are well known iron-formations.
25
The ~Lecambrian iron-fot~ations also occur in Urals and the Asiatic
parts of the Soviet Union (Alexandrov, 1973).
Iron-formations of India occur in several states of India
and major types are classified into two main groups as (i) the
iron-formation of the Dharwar Group, and (ii) the iron-ore series
of Bihar, Orissa and Madhya Pradesh. The BHQ and BMQ of Adilabad,
Guntur and Nellore districts of Andhra Pradesh; the BMQ of Salam,
Tiruchira Palli and Nilgiris districts of Tamil Nadu; the BHQ of
Ratnagiri district of Maharashtra, the BHQ of Goa, the BHQ
~nd BMQ OF Bellary, Chitradurga, North Kanara, Shimoga and
Mysore of Karnataka are well known iron-formations of the Dharwar
Group. Similarly, the BHQ of south Singhbhum of Bihar, Keonjhar
and Mayurbhanja districts of Orissa and Bailadila range of Bastar
district of Madhya Pradesh makes the iron-ore series (Picharnuthu,
1974; Krishnan, 1973).
Most of the iron-formations lie close to the border of
the Cratonic masses now surrounded by younger fold belts and
platform sediments as in between South America and Africa, and
between Australia and India. Therefore, perhaps, the segmenta
tion and drifting of Gondwanaland mass has followed a Precambrian
fault belt which was perhaps the basin of the deposition of
iron-formation (Gross, 1973; Eichler, 1976).
26
DISTRIBUTION IN TIME AND SPACE
The distribution of iron-forrnatior!s in the geological
time scale involves the ag~ determination which forms a major
problem in the case of banded iron-formations because of the
negligible concentration of elements present in it which are
essential for the radiometttdating. However, the probable
age range is determined by dating the associated metasedimen
tary and metavolcanic rocks which includes the underlying base
ment rock, rocks lying just above the iron-formation and the
igneous intrusions. Direct dating by K-Ar, Rb-Sr and U-Pb
techniques have been tried by many. The age determination by
K-Ar and Rb-Sr generally proves to be the age of post-deposi
tional metamorphism whereas dating by U-Pb claims to give
values close to that cf sedimentation (Goldich, 1973; James, 198~.
Despite the weakness in basic data the average age
range of the iron-formation~: have been established and can be
classified into four age groups as:
i) the middle Archean (3500 M.Y. - 3000 M.Y.),
ii) the late Archean (3000 M.Y. - 2600 M.Y.),
iii) the early Proterozoic (2600 M.Y. - 1900 M.Y.), and
iv) the late Proterozoic to early Phanerozoic (750 M.Y. -
450 M.Y.).
27
i) Deposits of early_ and middle Archea~ ~e i 3000 !E..=.Y.)
The oldest iron-formation is considered to be Isua iron
fol~ation, Greenland with an age of atleast 3750 m.y. (James,
1983). TI1ese iron-forrrBtions are interbedded with quartzite
strata that represent the oldest rocks of the supracrustal ori
gin. The other important iron-formations belonging to this
group are Irnataca complex, Venezuela (3400 rn.y. - 3000 m.y.);
the Pilbara and Yilgarn blocks of Hestern Australia (close to
3000 m.y.); the Ukranian Shield of Greater Krivoi Rog (3500 m.y.-
3100 rn.y.) and the iron-formation of India both belonging to
the Dharwar Group and the Iron-Ore series of Bihar and Orissa
with a probable age range (3200 m.y. - 2700 m.y.) (Goldich,
1973; James, 1983).
ii) I:e~sits ~!_ late .Archean ~e (3000 ~.: 2600 ~_:.l
The well known iron-formations falling under this age
group are: the iron-formation of the Superior Province of the
Canadian Shield, (2750- 2700 m.y.), ·.the Sekakvian, Bulavayan
and Sharnvayan system of ·arnbabwe (2750- 2700 m.y.), the Witwater
sand Supergroup of South Africa (2650 rn.y.), and the eastern
Kera1ia on the Baltic Shield (2800 rn.y.) (Goldich, 1973; James,
1983). The Indian iron-formations of both Dharwar group and the
Iron-Ore series of Bihar and Orissa, and the iron-formation of
Yilgarn Block of Western Australia are sometimes included in this
28
group instead of middle Archean because of the uncertainty in
their age (Goldich, 1973).
Most of the iron-fonnations of this group fall within
the brief interval of. 2750-2700 m.y. indicating an epoch fav
ourable for iron-formation.
iii) ~sits of early Proterozoic age (2600.: 1900 m.y.)
The early Proterozoic was the major epoch in the iron
sedimentation and stands as a milestone in the history of the
earth. The six most important iron-formations of the world
fall into this group. They are: the vast iron-formation of
the Lake Superior Region Canada-USA (2200-2000 m.y.); the
Labrador Trough and its extension of Canada (2200-2000 m.y.);
Krivoy Rog and Krush Magnetic Anomaly (KMA) of USSR (2200 -
1900 m.y.); Transvaal system of South Africa Minas Gerais of
South America and the Hamersley area of Western Australia (2200 -
1900 m.y.). Most of these major iron-formations fall within
an age range of 2200-1900 m.y. representing a very favourable
time period for the deposition of iron-formation (Goldich, 1973;
James, 1983).
iv) Deposits of late Proterozoic and early Phanerozoic age
These iron-formations are very rare. A few deposits are
there in Nepal and Brazil. They are also found in the Maly Khinghan
and Uola areas of far eastern USSR. Some appear to be Archean
29
types associated with contemporaneous volcanism and otters bear
an ill defined relation to the late Proterozoic glaciation. The
most notable example of the iron-formation related to the depositE
of possible glaciogenic origin is that of the Rapitan Group,
Canada, (James, 1983).
The spatial distribution or initial tonnage of iron
formations can not be precisely determined. The margin of
errors in these estimations are certainly very large although
there are reports on these here and there in the literature.
James (1983) has also given a good account of these for some
well known iron-formations. Problems associated in the estima
tion of spatial distribution are deformtions and erosion over
periods of time in billions of years, incomplete geological
mapping and falliability of geological assessment in connection
with the geometry.
James (1983) has also given a graphical picture of their
spatial abundance to indicate that the 90 per cent of the total
deposit belong to early Proterozoic which we call as the Superior
type. Jn terms of number of stratigraphically distinct bodies
the minor atundance peak of late Archean is most notable. These
deposits number in thousands but very minor in their initial
tonnage and are called ~he Algoma type. Similarly, the phane
rozoic are also not that abundant (James, 1983).
30
Banding is the most distinct lithological feature of
the Precambrian banded iron-formation although its origin is
yet controversial. Bands rich in iron minerals c-;lternates with
bands rich in chert. Bands are quite variable in their continuity,
thickness and mineralogical make up, giving a heterogeneous
character to the rock formation. Band thickness varies from
less than one millimeter to severe 1 centimeters and continuity
varying from few meters to several kilometers. A long contin-
uity of nearly 300 km is well marked in iron-formation of
Harnersley Group and at places of South Africa and India. Meta-
morphism and orogenesis has blurred many of its primary features
and tectonic deformation has generated many structural features
like folding.
Trendall (1973) has recognised three scales cf banding.
These in order of decreasing scales are macro banding, rnescbanding
and microbanding. Large scale macrobands are due to the alter-
nation between thin layers of shales and thick layers of banded
cherty iron-formation, the thickness ranging between 0.6 - 15
meters. Mesobands represent the alternative bands of sharply
defined iron rich and iron poor layers, thickness varying from
few millimeters to several centimeters. Microbands represents
thin banding of iron minerals and chert within a mesoband, thick-
31
ness varying between 0.2 - 2 millimeters. They are also known
as varves (Trendall, 1973).
The rnicrobands within a chert mesoband consists of
iron minerals and in a bedding planes that forms the adjacent
mesobands. Red micro bands are due to very fine disseminated
hematite, bluish grey rnicrobands are for specular hematite,
grey and black microbands are due to magnetite, while and yellow
micro bands are for carbonates, green micro bands are due to
chlorites and dark grey microbands for carbonaceous materials
(Alexandrov, 1973).
Iron-formations are sometimes classified as one com
ponent, two component or three component system depending upon
the number of rock forming minerals (except chert). They con
tain iron mineral in one component system which is either hema
tite or magnetite or carbonate with associate carbonaceous
materials. In case of two component system, microbands consists
of any two above minerals forming laminae alternating with each
other. In case of three component system, one mineral remains
as a subordinate to the other two and never forms a rnicroband
of its own (Alexandrov, 1973).
Iron rich shales also show rhythmic banding of barren
quartz or jasper layers alternating with shaly (or silicate).
layers,rnicrobandings are quite conspicuous in barren chert
(quartz) band and appear as thin laminae of siderite or magnetite.
32
As the iron rich shales grade into iron-formations, carbonate
quartz layers or shaly layers are replaced by ore layers (Melnik,
1982).
The bandings are usually parallel. However, angular
unconformities sometimes exist due to erosion during deposition.
Certain iron-formation exhibit cross bedding and ripple marks
to indicate their deposition under a shallow water condition.
Other sedimentary features like slumping, sedimentary breccia
and microfaulting are well marked in many iron-formations.
Sedimentary structures like scour and fill structure and pre~
lithification slump structures like shrinkage and cracks are
also marked at places. Development of above structures are due
to compaction, desiccation and diagenetic alternation of
amorphous precipitates (Gross, 1972).
Algal structures have been observed in Biwabic iron
formation, USA (Bayley and James, 1973) and in Kuruman iron
formation, Transvaal group South Africa (Beukes, 1973). A
remarkable feature of the Hamersley iron-formation is the
occurrance of lenticular, biconvex chert lenses along the
bedding planes of the evenly banded iron-formation (Eichler, 1976).
Iron-formations having oolitic and granular structures
are mostly restricted to the Proterozoic types (Beukes, 1973;
Gross, 1972). Granuiar textures are common in slightly metamor
phosed banded iron-formation of the Lake Superior region (Bayley
33
' and James, 1973; French, 1973), and in Kururnan and Penga iron-
formation of South Africa (Beukes, 1973). The granular textures
are most uncommon in BIF of USSR · (Alexaiidrov, 1973). The gra-
nules are either spherical or ellipsoidal and consists of iron
silicates (Greenalite, minnesotite and stilpnomelane), chert
and magnetite in variable proportion. In most of the cases,
rocks containing granules grades gradually into oolitic rocks
in which oolites are rimmed with hematite (Bayley and James,
1973; Eichler, 1976).
Indian iron-formations are well banded and are either
banded hematite quartzite (BHQ) or banded magnetite quart
site (BMQ). BHQ consists of alternating layers of hematites
and chert or Jasper with layers varying in thickness from 1
to 20 rnm (Melnik, 1982). BMQ are very irregularly banded
and cherts are medium to coarse crystalline. Breciations
have been marked in the iron-formations of Bababudan
area and in Iron-Ore Series of Orissa, Quartz veins
containing quartz of different texture with coarsely
crystalline hematite cutting across the bedding planes are
common features in many iron-formations (Picharnutu, 1974;
Krishnan, 197 3).
34
STRATIGRAPHY
Stratigraphy of iron-formations differs from one deposits
to the other in their lateral continuities, thicknesses and stra
tigraphic sequences. Even the stratigraphic sequences in a parti
cular deposit is not completely uniform throughout and have diff
erent stratigraphic columns in different parts of the same basin.
Inspite of these difficulties, the generalization in the strati
graphy of iron-formation within the deposits and among the deposits
has been tried by many investigators.
Major iron-formations belonging to the Proterozoic are
quite extensive. Laterally they extend from some hundreds to
thousand of kilometers and thickness varying from few hundreds
to thousands of meters. Whereas most of the Archean iron-formations
seem to be very much limited in their lateral continuities and
thicknesses because of metamorphism, tectonic deformations and
segmentations. Their lateral continuities at present stand at
several kilometers and thicknesses in range of 10 tc 100 meters
(Gole and Klein, 1981).
The Archean iron-formations are associated with various
volcanic rocks which include pillowed andesites, tuffs, pyroclas
tic rocks or rhyolitic flows and greywackee Iron formations
usually lie over the felsic volcanics and are in turn covered
by basic volcanics. Thin beds of graphitic schist and black car-
35
bon rich mud stones containing appreciable amount c.·f lead, zinc
and copper are interbedded with iron-formations. Tney are presumed
to have been derived from tuff and volcanic ash and collected in
depressions in the depositional basins (Gross, 1972).
Goodwin (1973) has reconstructed the stratigraphy of the
Hichipicoten basin (Fig. 2. 1). In th.e eastern and central part
of the basin, the lowermost mafic volcanic is overlain by pyro-
clastic rocks followed by the unit of iron-formation which is
origin overlain by mafic volcanics. This is followed by a
discontinuous unit of clastic sediments overlain by a younger
mafic volcanics. In the western section of the basin, the lower-
most mafic volcanics are overlain by a thick clastic sediment,
the unit of iron-formation re~ining just inside the clastic
sediment which is again overlain by younger mafic volcanics. So
the iron-formation in the central and eastern part is in between
mafic and felsic volcanics whereas in western part it is inside
clastics. In the western part:, the iron-formation represent its '
oxide facies and is typically composed of interbedded chert
magnetite and jasper (hematite chert) enclosed in greywacke
mudstones. The central IF representing the carbonate facies
consists of in descending order a thick band of tended chert,
thin sulphidic and thick carbonate layer. It gradually changes
into a thin or no band of banded chert, thick band of sulphide
and a thin band of car~nate towards the eastern part representing
West
KABENUNG SECTION 0
HELEN-MAGPIE SECTION 0
36
Eost
GOUDREAU SECTION 0
IRON FACIES. Oxide ----=-G--'ro..:.d.c..ot_io:._n--'-a_l -- Carbon ate Grodalionol Sulphide
B Iron f1Jrmot1on.
j:<:·~ :::::1 Dore sediments
E_:: 2 ::-~Felsic pyroclastics.
~Mafic volcanics.
Transitional from sediments to volcanics
... ..... .. u VI
..... ... u ;: a:
s,ooo 5 10 MILES
~L_ ___ _.~. ____ ......__
0 HORIZONTAL SCAlE
6
trupr,on Ond tectonic subsidence
co~E
Fig. 2.1 Reconstructed stratigraphic section of the Michipicotan
basin (from Goodwin, 1973)
1\ABENUNG SC::CTION (Oxide Fac.es J
--- -· - - - _-_-_-_ -_ ----~~,~~~ ~·~·\·-:--~.~-
'.'
-_ -_-__ -_-_-_-_-
D Andes1t. flows
t=-=--:$3 Shaltt-greywack•
~ lntttrb&ddtte! ~ chert-magnetite
HELEN SECTION C CarbOnate Foc1ttS J
•• ,, lo.i,l
L<2J Graph1t1c chel"t
r ~:: J Granular ch~t
~ Bandect chel"t
GOUDREAU SECTION C Sulphide Fac1u J
37
.jJJiB!BB~ "6-:LV~V;,.'\:;~Y· ~: ~- ...
c --~·-·' -~
0 r"l ~
Carbonate S = S1dttr1te C = h me stone
Rhyollt•- dacite tull brecCia, flOws
Fig. 2.2 Diagrammatic cross sections of typical iron-formation
showing stratigraphic arrangements and relationship
to adjacent volcanic rocks, H- Helen range,
B- Codon range, C- Goudreau range
38
the sulphide facies of IF (Fig. 2.2).
The stratigraphy of few greenstone belts of the Kaapvaal
and Rhodesian cratons indicates a cyclical nature between vol
canism and sedimentation. In ascending order of the stratigraphy~
each volcanic cycle consists of ultramafic to mafic to felsic
rocks terminating with sedimentation consisting of banded chert,
banded ferruginous quartzite and banded iron-formation. At
places a thin unit cf ultra mafic rocks directly overlie these
rocks constituting the base of another mafic to felsic cycle.
In ascending order of stratigraphy, volcanism gradually becomes
rnir.or and minor with equivalent increase in sedimentation, and
iron-formation gradually chan&ing from dominantly bar:ded chert
to dominantly banded iron fo1~tion through the bru:ded ferru
ginous, quartzite. It is also well marked on the surface along
the strike (Beukes, 1973).
A few iron-formations of late Archean like that of
Witwatersrand, South Africa lie within normal clastic rocks
(Beukes, 1973), and similarly in the nc~rthern rocky mountc..in
of Montana, USA, highly metamorphosed iron-formation is found
interbedded witt quartzite schist and dolomite marble (Bayley
and James, 1973).
The iron-formation of the Dharwar Group of India are
closely associated with volcanics. Iron-formation of the Shimoga
and Chitradurga of the Dharwar Schist Belt are assc~iated with
39
lime stones whereas the iron-formation of the Bababudan's are
not (Pichar.Dthu, 1974). The Iron Ore Series of Bihar, Orissa
and Madhya Pradesh in their stratigraphic column includes upper
shales and volcanics, banded hematite quartzite, lower shales
and purple sandstones with some lime stones, sandy and conglo
merate beds followed by phyllitic shales, tuffs and basic lavas
(Krishnan, 1973).
The Proterozoic iron-formations are not directly asso
ciated with volcanic rocks but volcanic rocks may be present
in some parts of the stratigraphic column. The stratigraphic
sequence in case of most of the Proterozoic irpn-forn~tion are
dolomite, quartzite, red and black ferruginous shale, iron
fotmation, black shale and argillites in order front bottom to
top, although there may be slight variations here and there.
The close association of iron-formation with qt¥trtzite carbon
aceous shale, conglomerate dolomite and argillitE.s are reco
gnised throughout the world.
The stratigraphy of major proterozoic iron-formations like
Hamersley Group, Transvaal Super Group, Lake Superior Region and
Labrador Trough are found to be more or less identical (Fig. 2.3).
The stratigraphic sequence in the ascending order bears a regional
unconformity separating lower Proterozoic sediments from the Arc
hean rocks. The Proterozoic sequence starts, with a clastic unit
accompanied by a great thickness of volcanics in Hamersley Basin
40
Hamersley 2350-2000 m.y.
Transvaal 2350-2100 m.y.
Lake Superior 1900-1700 m.y.
Labrador Trough 1900-1700 m.y.
Fig. 2. 3
Dolomite IF Shale Dolomite Shale Sandstone
Pillow Lavas and Basic Tuffs
sandstone
Dolomite
IF
Dolomite
Shale
Basic Volcanics
Basic Volcanics
Sills tone and Shale
IF
Shale
Dolomite
Shale
Dolomite
Red Sandstone and
• : Conglomerate
Generalized stratigraphic columns for four major
iron-formation deposits (a) Hamersley, (b) Transvaal,
·(c) Lake Superior, (d) Labra dor Trough (from Maynard, 1983)
41
and small thickness of volcanics in Transvaal group. This
unit in the Lake Superior region and Labrador Trough consists
of thick dolomite section followed by shale. This unit is
followed by another unconform~ty ~xcept in Harnersley Basin.
Next comes a transgression marked by thin clean sand stone .
Then comes the unit of chemical sediments composed of dolomite,
iron-·formation, lime stone and carbonaceous shale in varying pro··
portion in various relative positions. In the Transvaal the
sequence starts with thick dolomite that passes gradually into
iron-formation. The Harnersley Group contains thinner dolomitEs
but more shale. But in the Lake Superior and Labrador Trough
dolomites are absent and iron-formations lie directly on the
basal sand. This chemical unit is followed by a regional uncon
formity except in the Labor~dor Trough dolomites are absent and
iron-formations lie directly on the basal sand. This chemical
unit is followed by a regional unconfirmity except in the
Laborador Trough then followed by a thick cl8.stic volcanic
sequenc:e (Maynm·d, 1983).
The stratigraphy of ircn-fmmations belonging to late
Proterozoic and early Phar:erozoic c:•re of two types.. one group
found in Malykhinghan and Ude area of far eastern USSR are asso
cit:ted with contemporaneous volcanism and are mostly of the
Archean type. The oth€~r group is associated with sediments with
glacial features. The known iron-formations of this group are
Raptan Group of north CanaC:a, Jicadigo South Series, South
America, Damara Supergroup, South Africa and Urnberatane Group
of Australia. They are mostly laminated hematite and chert,
and are found inside clastic strata of glacial on glacial
marine origin (James, 1983).
ClASSIFICATION
42
The Precambriar: banded iron-formations have been class
ified into two major groups based on their lithologiE.s, rock
associations and depositional environments (Gross, 1965, 19i'3,
1980). The iron-formations belonging to the Archean age are
associated with volcanic rocks and/or greywackes, and Fere
mostly probably deposited in intercratonic basins of E.ugeo
synclinal types. These iron-formations are associated with
greenstone belts and are known as Algoma types. On the other
tand, the iron-formations c:f the early Proterozoic time that are
associated with quartzites, lime roch: and black shales and
an: devoid of any direct volco.nic ~)ssociation are known as
the Suped or type or the Animikie type. These were perhaps
formed in intracratonic continer:tal shelf environments of
miogeosynclinal types. These superior types are quite thick
and sufficiently extensive in their aeral eYtent compan~d to
the Algoma tn-e (Garrels et al, 1973; Lepp, 1975). A compari-
43
son between these two major types of iron formation hc.ts been made
by Gross (19i'3) and Eichler (1976).
This two fold clcSssification of the Precambric_n banded
iron-forrr•ation has its own limitc..tions (Trendall, 1968; Gale
and Klein, 1981).
i) There are many iron-forrr:ations younger than 1800 m.y.
that never come under this classificc:,tion. Therefore, Garrel
et al, (197~~.) has classified the frecambrian iron-formations .ss :;
(a) Archean type, (b) Animikie t~'e, and (c) Post Animikie type;
(ii) The present scheme of claE:sification has not recognized the
similarities amor!g iron-formations but largely ba~:ed on differ
ences in their stratigraphic Dnd tectonic settings (Kimberley,
1978; Gole c.:.nd KlE:.in, 1981); (iii) The Lake Supetior t~>es are
usuEJlly preEmmed to be a relEtively near shore deposit but there
are mar:y deposj ts belonging to this group that were usuc.dly
fotmed in a deep water offshore conditions as assumed in c8se of
the iron-formation of the Hamersley Pas in (Gross, 198 0) ;
(iv) Some Superior type iron-formations like that of Hamersley
Group art: also associated with volcanic roc.ks; (v) Iron-forma
tions confined to high grade granulite tE·rrE_ins also form a
distinct class. These are highly folded !:ended magnetite quart
zites metamorphosed under granulite faciEs cor!di tion, and occur
in association with quartzitEs, mica schist,marble and metavol
canics engulfed in tonalitic gneiss (Radhatrishna et al, 1986).
44
Therefore, it is quite naturc•l that the iron-formations
with their divergent char·acters cannot be exactly classified
although a broE1d classificE1tion to abcve two major type~ are
quite preva]ent.
MINERALCX;Y
Banded iron-formations are characterjzed by alter
native bar:ds of irc;n-minerals and cheL-t. Iron because of its
variable oxidation state is very sensitive to envircnmental
coEditions and, therefore, can form different minerals in
res~lnse to the available environmental conditions like Eh,
pH and concentration of ciff£rent active species like co2, S,
Sio2 and others. Post depositional chcu:ges like di~genesis
metamorphism and, of course, weathering, changes the initially
formE.d precipi ta tE·s to primary and secondary minerals. So the
present mineralogy of the irc•n-formation could be a product of
initial C(mditiom; and subsequent post depositional changE-~s
that it is bound to ur1dergo.
James (1954,1966) has claf:sified the iron-formation
based on its mineralogy into four different facies such as
oxides, carbonates, silicates and sulphides, depending upon
the dominant iroE mineral it carry witt. These facies are
extreme cases and are gradational giving rise to mixed facies.
1 0
.c F~ .. Oq ........ w
' 0·8
' . 0 ' < ly<·;;-
' 06
' ' ' ........ 04
Hematite Fe 2 0 3 0 2
0 .......
' -0 2 FeC0 3
-0 4
pH
Fig.2.4(a) Fh-pH stability field relationship among iron oxides,
carbonates and sulphides and showing how the field
45
area is a function of iron concentration "(from Stanton, 1972)
1·0
' -" 08 ' ' ·,.~ 10·,.
06 ' ' ' ' ' 0·-' '
0·2
Fe co3
-0·-'
-0·6
' -0·8
-I·OL_--~2L---~"L_--~6~--~8~--~~--~~--~1-' pH
Fig.2.4(b) Eh-pH stability field relations.among iron oxides,
carbonates, sulphides·and silicates at 25°C and 0 . total pressure, total co2= 10 m total sulfur
(from Garrel and Christ, 1965)
46
47
Stanton ( 1972) has pointed out that thE~ facies concept is a
simple and convenient division of the different mineralogical
types. Stabilities of these minerals and their interrelations
with respect to environmental conditions like Eh, pH & activities
of active species has been well discussed by Garrel and Christ '! ., I· ;I . '.-, r.J
(1968); Huber (1959); Curtis and Spears (1968) (Fig. 2. 4,8 and 2.4b).
Oxide Facies: The main oxide minerals that make the oxide facies
are mostly hematite (Fe2o3) and magnetite (Fe3o4). The oxide
facies can, therefore, be divided into two subtypes - banded
hematite quartzite and banded magnetite quartzite. The banded
hematite rocks consists of alternating bands of hematite and
chert. Hematites are well cryst~lline and the degree of crys-
tallinity depends upon the degree of metamorphism it has ·under-
gone. Sometimes hematites show oolitic structure, indicating
their deposition in shallow water conditions. The banded mag-
netite rods not only forms alternate bands of magnetite and
chert but also contains layers varying proportion of iron sili-
cEtes, carbonates and chert. The magnetitE sub facies,. there-
fore, go with slightly higher proportior! of MgO, CaO, MnO, C02 and FeO/Fe2o3 than its hematite counterpart. In an Eh & pH
stability diagram, hematite occupy a vider field of high Eh
and high pH whereas magnetite is confined to a narrow field of
low Eh and high pH conditions and overlaps with the field of
silicates. Its formation and its area of stability is subjected
48
to the activity of co2, S, Si02 and, thereforetalways associated
with carbonates, silicates and sulphides.
Carbonate Facies: The dominant iron mineral which makes this
facies is iron carbonate, siderite (Feco3). It also contains
other minerals like ferrodolomite, ankerite and calcite as
accessories. It consists of alternating layers of carbonates
and chert and looks like a carbonate counterpart of oxide
facies. They do not show oolitic structure, therefore, must
have deposited below the level of wave action. These carbonate
facies are usually associated with carbonaceous materials and
pyrites and are, therefore, seem to be slightly rich in P2o5
and ~riO. They mostly grades into either oxide or sulphide
facies. Although the gradation of carbonate to sulphide facies
'' is common, the gradation from carbonate to oxide is rare. The
stability field of carbonate facies, in the Eh and pH stability
diagram, remains in between the stability field of oxide and
sulphide facies which explains the above facts.
Silicate Facies: Silicate facies c:f iron-formation consists
of different iron silicates depending upon the degree of
metamorphism the rock has undergone. The iron silicates are
mostly greenalite, minnesotaite, chlorite, iron amphiboles
of greenali te - commingtonite se.ries, orthopyroxene and iron
olivine called fayalite. Greenalite is believed to be th~
49
primary silicate which has given rise to other silicates in the
course of metamorphism. Their formation also depend on the
availability of other oxides like Al2o3,Tio2, CaO, Na2o, K2o
and MgO. The silicate facies can be grouped as granular and
non-granular. In the Eh and pH stability field diagram, the
silicate field overlaps the stability fields of magnetite,
carbonates, and sulphides and, hence, closely associated with_
them. Pyrite is also sometimes found as an accessory mineral.
It is not quite clear, the conditions that would permit the
precipitation of iron and silica as one phase rather than
as separate phases (James, 1954; Melnik, 1982).
Sulphide Facies: The main mineral in the sulphide facies is
disseminated pyrite with small amount of pyrrhotite and minor
siderites in black carbonaceous shales. They may contain chert
layers here and there. All sulphide facies contain abundant
carbon and clastic material indicating their formation in deep
water condition where oxygen was not sufficient to decompose
organic matter (Goodwin, 1973). Sulphide facies are mostly
restricted to the Archean iron formation (Eichler, 1976).
CHEMICAL COMPOSITION
The Precambrian Banded iron-formation being a chemica
sedimentary rock, its chemical composition may act as a reflec
tor of the overall geochemical conditions that led to its forma-
so
tion. Assuming most of the post depositional changes like dia-
genesis and metamorphi~m to be mostly isochemical excepting, of
course, dehydration and decarbonation, chemical composition stand~
as a primary feature to understand the rock formation. Hence
the objective of chemical composition cmd the range of varia-
tior: in composition of its component parts are for three main
purpo~:es (I) general petrographic and chemical characterization;
(II) deterrninaticn of ore potential, and (III) understanding of
the origin and evolutionary developments (Davy, 1983). ftny
attempt to establish the overall composition of this E~xtremely
heterogeneous rock type with all scales of banding is first met
with the problem of sampling. In addition to this, a gradual
change in the facies and alterations due to metamorphism and
weathering, however, insignificant they may be und however
fresh a sample may- be, has made the overall composition inexact.
It is nearly impossible to establish the bulk composition Clf a ; '
whole rock formation with all scales of banding from the analy-
sis of a hand specimen of a short length core. Still then
attE.!mpts have been made in this important. field to establish
the range of compositions of major oxides, trace elements,
rare earth elements and isotoper:,and to study their utility
for a better understarrling of the rock formation.
(i) Major Oxides: Major oxide chemistry of iron-fottnations
have been well reviewed by James (1966). He has studied them
51
in groups according to their facies and compared them with the
major oxide chemistry of iron stones. Lepp and Goldich (1964)
have succeeded in establishing an apprm.imate composition range
with their average values fc•r few iton-fo11nations from Canada
ar:d United States and have made a good compHrison with that of
iron stones (Table 2.1). They have found an inverse relation
ship between Si02 and co2 and have used the CaO-MgO ratio to
distinguish ircm-formations from younger deposits. On an
average, iron oxic'e ancl Si02 arE.: the only two major oxicles in
thE.~ Precambd an iror!-fc•rmatj on and other oxides occur on trace
amounts. Mn seems to be more differentiated from thE~ Fe in
the Precambrian iron-fonnation than in the younger deposits.
Trendt!l and Block (1970), Gole (1970> and Trendall (1977)
have established the major oxide composition of few iron··
formation of the Hamersley Group, Western Australia. These
values have been well documented by Davy (1983) to demonstrate
bow the values are vad able because of tbe facies gradation
and different scale of b?nding invohErl in the iron-formation.
Compositional variaticn in meso and macro bands have also been
well discussed by Davy (1983). Eichler (1976) has indicated
the average major oxide composition of different facies of
iron-formation to show how they are facies dependent (Table 2.2).
Triangular diagrams representing Fe, Si02 and CaO + MgO (Lepp
52
Table 2.1: Comparison of the chen!ical composition of the Precambrian and the post Precamt.·rian iron-formations
Range Average Range Average
Total Fe 17.1-44.2 27.8 15.2-49.9 29.0
("'. 0 ._:.]_ 2 7.3-64.2 42,8 4.5-55.7 12.9
Al2o3 0.03-13.93 1.6 0.24-16.8 6.1
CaO 0.01-10.48 1.5 0.1-33.0 14.3
MgO 0.04-11.22 2.8 0.45-7.84 2.9
MnO 0.01-5.06 1.0 0.02-1.8 0.34
P205 0.03-4.02 0.26 0.14-2.2 0.86
Ti02 0.02-0.52 0.15 (1.17-2. 2l~. 0.45
c 0.01-3.05 0.4 0.58-2.55 1.11
co2 0.1-31.56 8.1 1. 5-30.52 17.8
H20 0.25-9.29 L'.5 0.26-15.1 4.7
CaO/MgO ratio of iron-formations c:f different ages
Precambdan
Paleozoic
Mesozoic
Range
0.008-2.06
0.03-:!.9.1
0.62-47.0
(from Lepp and Goldich, 1964)
Average
0.59
8.0
8.8
53
Table 2.2: Average chemical cor:position of different sedi-mentary facies of iron-formations (compiled from James, 1966; Eichler, 1968, 1970)
Oxide Silicate Carbonate Sulphide facies facies facies facies
Fe 37.8 26.5 21.23 20.00
FeO 2.1 28.9 22.22 2.35
Fe2o3 61.63 5.6 5.75
Fes2 38.7
SiO · 2 42.89 50.7 48.72 36.67
Al2o3 0.42 0.4 0.15 6.9
Mn 0.3 0.4 0.5 0.001
p 0.003 0.07 0.09
CaO 0.1 0.1 4.6 0.13
MgO + 4.2 C.84 0.65
K20 + + 1.81
Na2o + 0.01 0.26
Tio2 + + + 0.39
co2 5.1 14.1
s + 2.76
so3s 2.6
c + ++ 7.6
H2o 0.43 5.2 2.67 1.25
(from Eichler, 1976)
and Goldich, 1964) and representing Fe2o3, Sio2 and Al2o3
(Govett, 1966) have marked the area for the iron-formations
which distinguishes them from other rock formations. Mean
chemical composition for the Algoma and the superior types
54
for different facies have been given by Gross (1986) (Table 2.3).
(ii) Trace Elements: The study of trace elements in the
banded iron-formation is very much limited. The inclusion of
trace elements in a particular rock formation seems to be a
product of geochemical environment in which it is formed and
subsequent post depositional changes to form particular min
erals to accommodate the trace elements. Therefore,the con
centration of trace elements in a particular rock formation
may be a reliable basis to interpret the physichemical condi
tion of deposition and probable source for the constituent
materials.
The Precambrian banded iron-formations seem to be very
much deficient in their trace element composition. Landergren
(1948) was perhaps the first to make an attempt to study the
abundance and statistical distribution of ferrides (Ti, V, Cr,
Mn, Co & Ni) in iron rich sedimentary rocks of Sweden. A few
trace element values in iron-formations here and there like
that of Alexandrov (1973); Plaksenko et al (1973) indicate
their paucity in some Trace Elements (TE) like Ni, Cu, Cr and
55
Table 2.3: Mean chemical composition of the Algoma and the Superior type iron-formation for the different facies
Major oxides
Si02 FeO
Fe2o3
Oxide Facies Algo- Super-rna ior
47.83 47.7
12.7 7.8
30.33 35.6
Fe2o3(T) 44.19 44.27
CaO 1.66 1.6
MgO 1.58 1.24
Na2o -0.33 O.ll
H2o 0. 71 0.14
Al2o3 2.65 1.27
P205 0.21 0.05
co2 0.88 2. 72
H2o 0. 77 1.13
(from Gibbs, 1986)
Silicate Facies Carbonate Algo- Super- Algo- Super-
rna ior rna ior
64.2 58.83 44.5 36.91
10.67 16.73 13.39 21.53
10.91 8.98 2.7 6.89
21.99 27.58 17.81 30~51
4.94 2.26 4.66 4.93
3.57 2.84 6.18 4.45
0.22 0.18 1.23 0.14
0.24 0.55 0.89 0.14
2.52 2.18 6.67 1.31
0.07 0.1 0.14 0.14
2.19 4.29 15.54 20.92
1.25 2.48 1.65 1.37
Sulphide facies Algoma
40.94
15.96
13.15
29.57
2.45
2.29
0.86
0.94
6.64
0.11
2.05
3.25
Zn. Plaksenko et al, (1973), has indicated that the average
concentrations of trace elements like Ba, Ti, Cu, Ni and V
in oxide and silicate facies of the iron-formation seem to
remain within the range of 28-110 ppm. It also marked that
the iron-formation of the volcanic association are more in
Mn, Mg, Cu, Ca, Co, S, B, P & Ni and less in Ge, V and Sr
than the iron-formation of terrigenous association.
56
Gole (1~1) has reported a few trace element values
like Ni, Cu, Cr and Zn for some iron-formations and Fe rich
shales, of the Yilgarn Block, Western Australia and have
compared their average values and compositional ranges with
that of Hamersley Group iron-formations. Fe rich shales seem
to be richer in Ni, Cr & Zn and a little poor in Cr compared
to banded iron-formation.
Majumdar et al, (1982) have analysed a few iron-forma
tion samples of Orissa, India and indicated their abnormally
low values of B, Ba, Sr, Ni, Cr, Ti, Zn and Co/Ni and have
suggested a low temperature sedimentary process in a shallow
water condition for their formation.
Both Algoma and Lake Superior type iron-formation show
a parallel trend in most of their average trace element compo
'sition compared with that of earth's crust, except of Mn,
which is abnormally high in Lake Superior type and Ni, Cu and
Zn which are strongly depleted (Maynard, 1981).
Table 2.4 Mean Trace Element Composition in ppm of the
Trace Ele-ment
sc
Ti
v
Cr
Mn
Co
Ni
Cu
Zn
Ba
Sr
Zr
Algoma and the Superior type iron-formation for their dominant facies
Oxide Facies Silicate Facies Carbonate Facies Algo- Super- Alog- Super- Algo- Super-rna ior rna ior rna ior
8 11 22 10 16
750 170 1910 1600 2090 340
60 30 110 100 170 60
78 110 140 100 390 80
1120 5130 2190 4000 1720 7390
38 30 40 30 30 30
80 30 340 - 50 220 40
50 10 20 40 80 10
60 30 50 40 90 100
210 170 130 150 240 40
70 30 90 20 210 30
40 60 60 170 80 70
, (from Gross, 1980, 1986)
57
Sulphide facies
Algoma
~10
2620
90
160
2330
80
130
630
3430
150
100
120
Trace elements analysis of the Archean iron-formation
hosting gold mineralization, Zimbabwe (Fripp, 1976) indicates
a high base metal concentration which is presumed to be due to
the gold association in sulfide facies.
Davy (1983) has summarized trace element values of
some important iron-formation to conclude that all of them are
present in trace amount and variations are due either to wrong
sampling or to the association of ore bodies. There appears
to exist not much difference in trace element content between
oxide, silicate and carbonate facies but a higher amount of
trace element in sulphide facies either additional source
material or a different depositional environment from other
facies. The mean trace element composition of the Algoma and
the Superior types has been given by Gross (1986) (Table 2.4).
Rare Earth Elements: Rare earth elements (REE) because of
their peculiar electronic configuration (n-2)fx (n-1), d1 ns2
and a uniform decrease in ionic radii, from 1.03A for La to
0.86 A for Lu, mostly show coherent behaviour with uniform
regular change in their geochemical characteristics. All of
them show an oxidation state + 3 except Eu (+2, +3)?nd Ce(+3,+4).Their
relative abundance in a system is a product of geological
processes and stands as a clue to understand the geological
history (Hanson, 1980).
Rare earth elements are characterized by their low
solubilities and mobilities in sea water. As they have a
very short residence time (50 to 6000 years) compared to the
mixing time of the ocean ( 1000 years), majority of them go
immediately into sedimentation. Of course, the heavy REE
(Gd-Lu) go with a slightly longer .residence time period
than the light REE (La-Eu) because of their better ability
to form stable complexes. Marine sediments, therefore,
usually show a depletion in heavy REE and enrichment in
light REE. In addition to this, Ce (+3, +4) and Eu (+2,
+3) because of their variable oxidation states show differ
ent mobilities under different degree of oxidationand redu
ction conditions. Hence Ce & Eu exhibit anomalous behaviour
of either enrichment or depletion compared to their neigh
bouring elements (Fryer, 1983).
The study of REE in the iron-fonnation is really
very scanty. Fryer (1977a, 1977b, 1983), Gtaf (1977), Majum
dar et al, (1984) have done some work in the field of REE in
the iron-formation to give us some understanding. Most of
59
the Archean iron-formations are characterized by low absolute
REE contents compared to the crustal REE abundances. They also
show a relative enrichment of Eu compared to their contempora··
neous clastic sediments and/or volcanic rocks (Fryer, 1977).
The low absolute REE abundance could be due to the precipi
tation of iron minerals from the sea water which has extremely
lLJ I-
20
10
ii' 2 0 z 0 r u 10
~ ---..so
lLJ ...J a.
'" <I
"' 20
10
5
2
0" 0 .
--.....___0
~ 0
6 Fonland
o WyamonQ
X Mary River
• TemOQOmo
o Michipicalen
e Akilio Stllcate - Q.,de o lsuo Sulfode
X•{0...._0_,o~ o,
~ ·-·- 0 ·--·--· 0 .~ • --. ·-·/
~~~~-L-Jl ____ ~I_L I
La Ce Pr Nd Sm Eu Gd Tb Oy Ha Er Tm Yb Lu
Fig.2.5(a) Chondrite normalized REE abundances in the Archean
iron formations (from Fryer 1983)
60
wiO ..... cr a z 0 I u
~ ....____ w _J
a. ~ <I V'l
10
61
'----'---'-----'- --1...._---'-----L-L-_L___l_ __.__.)._ - ' t ' t
La Ce Pr Nd Sm Eu Gd Tb Dy '-to !:r •m Yb Lu
Fig.2.5(b) C'.hondrite normalized REE abundances in the Proterozoic
iron-formations (from Fryer, 1983)
low REE abundances (Fryer, 1983). The positive Eu anomaly of
the Archean IF could be attributed either to an anomaly in
1 f . f E +2 the sea water resu ting rom greater transportat1on o u
in Archean conditions or due to the nature of minerals pre-
cipitated from sea water and their kds. No significant ano-
mal6us behaviour for Ce has been observed for the Archean
iron-formation (Fryer, 1983) (Fig. 2.5a). Their REE pattern
show consistently middle REE depletion, i.e. concave downward.
The Proterozoic IF have generally higher abundance
of REE, more so on the light REE related to the Algoma IF
(Fryer, 1983) (Fig. 2.5b). The-REE patterns are well fra-
ctionated with LREE enrichment. Unlike the Algoma IF they
show no consistent Eu anomaly. In fact there seems to be
no significant Eu anomaly at all.
Ce shows a slight depletion anomaly in the Proterozoic
iron formation which becomes substantial in the Phanerozoic
iron-formation. It perhaps get separated from the rest of
REE by oxidation to +4 state Eu still behaving partially
as Eu (+2).
Fryer (1983) feels that the formation of Mn nodules
in the late Proterozoic and in the Phanerozoic sea helps the
oxidation of Ce to +4 state to separate it from the rest of
the REE, Mn acting as a catalyst.
62
REE abundances and patterns exhibit a small but signi
ficant facies dependance. The carbonate and the silicate fac
ies show a constant REE pattern and trace element distribution
because of their initial origin as a crystalline precipitate
in equilibrium with sea water. But the oxide facies shows
a variation in REE patterns and trace element distribution
because of their initial precipitation as hydroxide and
subsequent diagenetic change to oxide which might have aff
ected the REE pattern and trace element distribution. Dia
genetic change of REE are very much complex and poorly under
stood (Fryer, 1983).
Graf (1977) has accounted for the REE pattern in the
iron-formation as a product of REE pattern of the source mat
erial probably a hydrothermal solution that brought iron, the
REE pattern of sea water to which the hydrothermal solution
was discharged and the degree of m~ing that took place
between two. He has explained the Eu anomaly in the Archean
IF because of their enrichment in the hydrothermal solution
due to its interaction with felsic volcanic rocks. Fryer (1983)
has criticised this on the ground of selective interaction of
hydrothermal solution only with the felsic rocks.
Isotopes
Isotopic studies of carbon, oxygen and sulphur in the
Precambrian iron-formation though seem to be very much limited
64
but they are certainly very much useful to understand this rock
formation. As a convention, the isotopic variations, for example,
oxygen are presented as in per milli where
-1 J X 1000
Standard being the Standard Mean Ocean Water isotopic ratio
(SMOW) and the difference in isotopic composition between
chemical phases A & B are expressed as
fj AB ·= 1000 ln( r,18o;$-6o) A
( s18o;~6o ) B
~A~f;B
(Perry, 1983)
Oxygen isotopic study on the iron-formation helps one
to understand the nature of the Precambrian hydrosphere and
acts as a tag to metamorphic reaction and hence to the origin
of iron minerals. It is presumed that chemical sediments are
enriched on &18o compared to the water of the basin and this
of course dependent of equilibrium temporarily. There also
exist a functional relationship between the b18o of silica
and the time of deposition. Based on this it was concluded that
the Precambrian hydrospheres were depleted in the ~180 with
a 5180 value for the Archean ocean about - 18 per milliand
the Proterozoic form -10 to -12 permilli. !here is, of course,
a big difference of opinion why the Precambrian hydrosphere
were poor in 0180 but increased with time (Melnik, 1982).
While some feel that the variation was due to the mass depo-
sition of chemogenic carbonate silicate with a variation in
temperature, others argue in favour of extensive circulation
65
of water through mantle. There are, of course, uncertainities
in the above isotopic values. While siderite of Krivoy Rog
gives a value of &18o in the range of +12.3 to 12.9 per milli
which correspond to ~18o of the then hydrosphere to be -14
to -15 per mil~ the carbonates of Hamsley area gives 5180
value + 20.12 to 21.20 percent indicating the s18o for
ocean water to be in the range -11 per mil·ll to -3.5 per milti
(Melnik, 1982).
Isotope fractionation of oxygen seems to•be facies dep-
endent and provides some insight to estimate the temperature
at which the rock equilibrated. Therefore, it could tell the
grade and temperature of metamorphism that the rock has
undergone. It also could indicate the origin of iron minerals
like magnetite whether it is primary or a secondary mineral, if
secondary then whether it is after hematite or after siderite.
The ~180 of hematite and magnetite are similar and vary
'within +1.3 to 6.8 per milli but siderite is enriched in ~18o
taking a s18o value upto + 12.9 per miU:i.. A slight
66
enrichment of s18o in magnetite indicates that the magnetite
possibly was formed at the expense of hematite. But a magnetite
substantially enriched with ~}8o , could have been derived
from siderite (Melnik, 1982; Perry, 1983).
Carbon isotopic study seems to be useful in under
standing the evolution of the biosphere and its direct or
indirect role in the precipitation of iron-formations. It
has been established that average isotopic composition of the
two forms of carbon-carbonate ( s13c = 0~0/~ and organic
( s13c = -27/,J was constant throughout the Precambrian.
The study of 013c indicates a very early evolution of
life, photosynthetic microbiote, and their participation in
the formation of BIF. It indicates that the organic process
in the formation of BIF were quite prominent in the protero
zoic but was more limited in the Archean (Becker and Claytin,
1972; Melnik, 1982). The shallow water chert facies shows a
negative s13c value ( -25 to 30/o) whereas deep water chert
carbonates show (-15 to 20/~ indicating that shallow water
precipitations were more biogenic where deep water depositions
were mostly abiogenic (Melnik, 1982).
PAlAEOBIOLOGY
It is a fact that the Precambrian iron formations were
formed before the evolution ~f any megascopic form of life.
The palaeontological study of the Precambrian iron-formation
has indicated the presence of microfossils and sedimentary
structures produced by benthic mats of microbes (bacteria,
cyanobacterja, and algal protist) called stromatolites, giving
the evidence of microbial activity, during their formation. J
These have been well summarised by La Berge (192)), and Walter
and Hofmann (1983). Archean iron-formation contain organic -
matters called kerogen which could be the degradational pro-
duct of the original cell and its isotopic composition seems
to be the source of palaeobiological information. The lack of
stromatolites in the Archean iron-formation indicates that they
were most probably deposited in a deep water condition below
the photic · zone.
Early Proterozoic iron-formation on the other hand
exhibit the presence of wind fossils both spheroidal
and filamentous which are between 1 to 10 urn in cell
diameter. They also contain speroidal sedimentary stru-
ctures called stromatolites. Mickiie and Tate proterozoic
stromatolites seems to be deprived of these features (Walter
and Hofman, 1983).
ORIGIN
The origin of the Precambrian banded iron-formation is
the most controversial but interesting aspect of it has
68
kept the investigations busy for the last hundred years. Although
sufficient field and constitutional informations are available
now the nature of the exact origin is yet unclear. Rocks are
so much diversified in their general and local features that each
of them proposes a separate model for their genesis and a uni
versal model of general acceptance free from all criticism is
yet to evolve.
Winchell and Winchell (1981) after considering diff
erent theories on iron-formation of the Lake Superior Region
concluded that no thinking person could ever attempt to explain
all the deposits by any one particular theory (Lepp, 1975).
Similar views were extended by Eugster and Chou (1973) that
banded iron-formations with their diversified characteristics
inspite of their common peculiarities cannot fit into a single
depositional model. Still then attempts are on to understand
the common thread which holds all together.
Any discussion on the origin has to be based on the
following points: (1) The source for iron and silica, (2) The
transport of iron and silica; and (3) The deposition of iron
and silica.
Any model based on these should be in a position to
explain all their peculiar features like banding and cons
titutional make up.
69
SOURCE
The source for iron and silica for the formation of the
banded iron-formation is the most debatable question. Three
main sources proposed for iron and silica are: (i) the volcanic
exhalations, (ii) the continental erosion and (iii) the oceanic
upwelling. Trendall (1968) has called these main possible
sources as "from below", "from above" and "from within". Alex-
androv (1973) has also mentioned few other sources like meta-
somatic, magmatic and cosmic origin.
(i) Volcanic Exhalation: As pointed out by Trendall (1968)
the main proponents of the volcanic exhalation source for iron
and silica are Van Hise and Leith (1911), Gruner (1972), Tanton
(1950), Guild (1953, 1957), Goodwin (1956, 1961, 1962), Oftedahl
(1958), Harder (1963), Trendal1 (1965) and La Berge (1966). There
are also a number of Russian supporters of this theory (Alexan-
drov, 1973; Melnik, 1982). All 'of them point out some direct
or indirect relationship between volcanism and iron-formation.
Guild (1953) pointed out that volcanic activities that
were dominating the Precambrian earth might have brought enough '
iron and silica to the depositional basin for precipitation.
These volcanic activities might also have lowered the pH of basi-
nal water to prevent the precipitation of other carbonates.
70
Goodwin (1956) had detected pyroclastics and lavas along
certain horizons of iron-formations of the Lake Superior Region
and thought of some cyclic coordination between volcanism and
iron-formation. From the example of Santorian volcano in the
Aegean sea which recorded the out bursts in 1650, 1707 and 1869,
he concluded that the exhalative volcanic activities depositing
ash, tuffs and pyroclastics followed by a long period of chemi~
cal volcanic exhalation which were characterised by sustained
discharge of ferrous carbonate bearing solutions or gases to
the depositional basins.
Trendall (1968) assumed that first stage of volcanic
activities formed depositional basins by a depression and com-
pensating elevation of the surrounding area. The sub-marine
volcanic activities brought enough iron and silica in the middle
stage of the basin development for precipitation. Subsequent
basin development was the erosion of basin edges to fill the
centre.
Beukes (1973) from the stratigraphic study of iron-forma-
tions associated with the greenstone belts of the South Africa
proposed that perhaps the silica for the banded iron-formation
came from an acid volcanism whereas the main source of iron for \
the iron-formation was continental weathering.
Some are also very critical about volcanic exhalation as
a source for iron and silica. They argue that volcanism was no
71
doubt a common event on the Precambrian earth and a main respon
sible factor for basin developments but it is not necessary to
think that rocks associated with volcanics must themselves be
of yolcanic origin (Trendall and Blackley, 1970; Holland, 1973).
Holland (1973) has argued that Ebeko volcano in the
Kurila island annually produces some 35-40 tons of iron. At
this rate of discharge some 1000 volcanoes would have been
necessary to deliver the total iron present in the Harnersley
basin in nearly six million years. Although this time period
never loo~dunreasonable in geological time scale but such a
large number of volcanoes along the periphery of the basin,
each hardly a few miles apart makes the volcanic source very
difficult to imagine.
Holland (1973) has also disproved the volcanic source
by comparing the iron rich sediments near active volcanic ridges
to that of usual iron-formation.
Lastly, since majority of the iron-formations, the
Superior type are completely devoid of any volcanic associa
tion, volcanism can not be considered as the major source. Still
then Goodwin (1973) and many others are completely convinced of
a volcanic source atleast for Algoma type iron formation.
(ii) Continental Erosion: As reported by Trendall (1968) there
are also many proponents to the theory of continental erosion
72
as a source for the banded iron-formation. They are: Gruner,
(1922), Gill (1927), Moore and Maynard (1929), Tayler (1949),
Sakamoto (1950), James (1954), White (1954)_, Alexandrov (1955),
Hough (1958), Huber (1959), Lepp and Goldich (1964) and Govell
(1966) and Lepp (1975). All of them suggest that the weathering
of low lying landmass under humid tropical conditions were the
main source for iron and silica.
Gruner (1922) was convinced that continental erosion
was an adequate source of iron and silica for the iron-forma
tion. He argued that Amazon River could carry the amount of iron
. equal to Biwabik iron-formation in a short time of 176000 years
which was not much unreasonable in the geological time scale
(Lepp, 1975). But it is a fact that there is neither any depo
sit of iron in Amazon off shore area not in any other river
basin to support the theory (Govett, 1966).
Lepp and Go1dich (1964) have developed a lateritic
weathering model which successfully explained the chemical diff
erentiation in the iron-formation. But this weathering model
has been well criticised by Gross (1965) and Trendall (1965).
James (1966) also supported the theory of continental
erosion as a source for iron and silica but he pointed out that
perhaps the fractionation of iron and silica from other elements
took place during the process of precipitation.
Hough (1954) and Govett (1966) have strongly believed
the idea of continental erosion as a main source for iron and
silica and taking use of the theory they have succeeded
to explain the nature of the rhythmic banding.
Beukes (1973) thought the continental erosion to be
the main source for the iron which carne periodically to the
depositional basin \vhere silica from an acid volcanic source
was precipitating continuously to give bandings.
At most of the Proterozoic iron-formations are devoid
of any volcanic association, many considered continental
weathering as a major source of Fe & silica. However, it is
very difficult to believe that iron formation ·derived from
the continental source could be so low in Al2o3 and Tio2 con
tent. Some argue that iron and silica carne in the form of
solution in pure dissolved state and were deposited s~ightly
away from the site of deposition of clay and other terrigenous
matters (Melnik, 1983).
Arguments, therefore, both in favour and against of
the continental erosion theory continues without any solution.
73
(iii) Oceanic Upwelling: Those who have looked into the ocean
as a major source for iron and silica for the iron-formation are
Borchert (1960), Holland (1973) and Dever (1974). This has been
described as "from within" by Trendall (1968). Although Borchert's
model was in connection with younger iron deposits but it provi
ded a mechanism for the concentration of iron within the sea.
Holland (1973) has strongly advocated that deep sea
water saturated with siderite containing about 3 ppm of Fe
could be oxidized during upwelling of sea water to get pre
cipitated as Fe(OH) 3 in shallow marine conditions. Similarly,
the upwelling sea water quite saturated with amorphous silica,
74
precipitated silica as their solubility decreased with decrease
of water pressure.
This oceanic upwelling theory though looks good in
some .respects but seems insufficient to explain some of the
peculiar properties like banding and chemical composition.
Secondly, assuming that the volume of the Precambrian ocean
was same as that of the present ocean, its total iron con
tent with a saturation concentration of 20 mgL-l would have
been 25.6 x 1018 gm. It is just half of the iron present in
Dates Gorge and Joffra Member combined together. It is also ..
unreasonable to think that all the irons of the ocean precipi-
tated in a localized area like Hamersley basin (Ewers, 1983).
In addition to the above three major sources for iron
and silica in the iron-formation, few other hypothesis of minor
importance has also been discussed by Alexandrov (1973).
(iv) Metasomatic Hypothesis: According to this hypothesis the
iron for the banded iron-formation was derived from the hydro-
thermal alteration of basic and ultrabasic rocks mostly
amphibolites and schists during the process of granitization.
'-'
It is based on the fact that at places a grading of iron-formation
into pyroxenites and the presence of relicts of basic and ultra
basic rocks in the iron formation has been marked (Alexandrov,
1973).
(v) Magmatic Hypothesis: Some believe that the iron ore
bands in the iron-formation are epigenetic having been formed
by pneumatolytic injection of hematite between the laminae of
quartzite. Sometimes iron-formations have been considered as
a differentiation of the basic magma (Alexandrov, 1973).
(vi) Cosmic Hypothesis: Some people are with the feeling
that when earth in the solar system moved within the galaxy,
it passed through a zone of cosmic dust and the magnetic field
of the earth increased the fall out of thes~ iron particles
over the Precambrian continents. These were washed down to the
depositional basins. But the absence of substantial amount of
Ni in the banded iron-formation greatly weakens the hypothesis
(Alexandrov, 1973).
Recently, Carey (1986) has suggested that in the absence
of vegetation on the earth surface, before the middle paleozoic,
the land surface was completely barren, and the wind was the
main carrier on the land surface. He, therefore, suggested that
dust storms were responsible for transport of iron and silica as
fine particles and their deposition on a broad epeiric basin.
Carey is a strong advocate of expanding earth theory and feels
that the Proterozoic earth was only half of the present dia-,
meter, and, therefore, no great ocean existed before the mezo-
zoic to act as a depositional basin for the iron-formation.
According to him, the bandtng was due to cyclic interruption
of normal clastic sedimentation by carbonate dust storms.
TRANSPORT
After the problem of source for iron and silica, the
next question that puzzles the investigators is their
transport from the source to the site of deposition. It is
quite undoubtful that the iron-formations are chemical sedi-
ments as they show a facies pattern and contain very little
alumina in their composition to call them terrigenous
sediments. So it is believed that whole of the iron and silica
are transported in dissolved state either in ionic form or in
colloidal form. But, since typical banded iron-formations are
characteristic only of the Precambrian time which has not been
repeated in same nature and proportion, the migration of iron
and silica in a huge quantity demands a special Precam-
.brian condition. I
The transport of iron and silica for a source like val-
· canic exhalation never poses much problem. It is assumed that
77
the sub-marine volcanic activities reduces the Eh and pH of
h f . ' 1 . ' h F +2 t t e sea water suf 1c1ent y to transport 1ron 1n t e e sta ·e.
If the volcanic orifice is on the land surface, the acid
. f ' f h . h' h F +2 . water com1ng out o 1t orms a ot spr1ng w 1c carry e 1n
it to precipitate them by coming in contact with normal sea
water (Borchert, 1960; Goodwin, 1958). Similarly, the trans-
portation is not a problem for a source like oceanic upwelling
as the upwelling current of the ocean water could carry iron
and silica for precipitation.
In contrast to the above, the transportation of iron
in case of a source like continental erosion seems to be a
major problem as iron forms a stable compound like Fe(OH) 3 at
normal Eh and pH conditions of the surface water.
Moore and Maynard (1929) believed that iron was trans-
ported as iron hydroxide hydrosol and the silica as the
coloidal silica. They were stabilized by organic matter,
which kept them from mutually precipitating one another till
they come in contact with sea water. The above process was
questioned on the gtound that as iron-formations show a facies
pattern,,they must have been transported in the form of ionic
solution (Lepp, 1975).
Mac Gregor (1927) was first to point out that the
restriction of the iron-formation to the Precambrian time could
78
be because of their formation under an atmosphere with higher
co2 and lower o2 partial pressures. The higher I{:02 in the
atmosphere perhaps accelerated the weathering of the rocks and
kept the iron in the ferrous state for easy transportation.
Others those who have supported to this idea are Tyler and
Twenhofel (1952); White (1954); Lepp and,GOldich (1964); Cloud
(1968, 1973) and Garrel et al, (1973) (Lepp, 1975).
There are many like Revella and Fairbridge (1957) and
Huber (195.9) who have challenged the above idea of an oxygen
poor Precambrian atmosphere. Firstly, they believed in the
existance of a delicate interaction between atmosphere and
hydrosphere for which the buffering capacity of the vast
ocean was capable of maintaining a constant composition of the
atmosphere throughout the geological time. Secondly, there are
also many post-Precambrian iron deposits that were most probably
formed under an oxygenated atmosphere. Hence, an anoxygenous
atmosphere could not be the sole criteria for the transport of
iron and silica (James, 1966).
An interesting attempt, therefore,was made by Carroll
(1958) to explain the transport of iron under an oxygenated
\atmosphere. He showed that iron could be transported as a
constituent of clay either as a substituting impurity in the
clay lattice or as an iron oxide coating on the surface of the
79
clay mineral. This theory is objected on the ground that the
alumina which is a major constituent of claymineral is in a
significantly low concentration in the iron-formation to prove
this model.
Beck (1972) has pointed out a high concentration of iron
and silica in some stream water in low relief and high rainfall
country of south eastern USA. He has also marked that Fe to Sio2
ratio of these waters were very much similar to that of the
Precambrian banded iron-formation. Therefore, if it is also
assumed that the Precambrian atmosphere, though not completely
anoxygenous, was certainly slightly higher in its co2 content
with a partial pressure of co2 around 0.03 · atmosphere against
a present value o·~ 0.0003 atmosphere the pH of the surface
water could be reduced from 8.17 pH at present to 6.1 pH at
the Precambrian time. Again if it is assumed that the total
volume of the surface water in the Precambrian was less than
the present, the pH of the surface water in the Precambrian
could be still lower to transport iron in ferrous state to the
site of deposition (Govett, 1966).
Melnik (1982) has made a detail physico-chemical study
of the migrational character of both iron and silica. From the
observation of their geochemical behaviour in present day
hydrosphere he has succeeded in extrapolating their nature into
the Precambrian conditions.
80
In aqueous solution, iron occurs in two valency states,
Fe+2 and F+3, and in dissolved state could be transported
either in the ionic form or in the colloidal form. In a Eh-pH
diagram, for the system of Fe-H2o, (Melnik, 1982) (Fig. 2.6),
the stability field of Fe+J is restricted to highly acidic
condition (pH less than 3) and it forms only hydroxide ions
+2 + of type Fe (OH) and Fe (OH) 2 in highly acidic conditions
and Fe (OH)z in highly alkaline state.
Although Fe+J is capable of forming complex ions in
the presence of inorganic ions most of them are only stable '
in highly acidic conditions. Such high acidic conditions for
the transport of iron either in the form of Fe+J or in the
form of its inorganic complex ion seems to be very much unlikely
on the earth surface. This is because such a highly acidic
water would immediately react with rocks to decrease the acidity
d h d . 1 . f . 1· k Ca+ 2 Mg+ 2 N ·+ K+ ue to t e ~sso ut~on o cat~ons ~ e , , a , etc.
Even if the Precambrian atmosphere is assumed.to be full with
co2 with pC02=1, the pH of surface water at this condition can
take a maximum value 3.9 which is reasonably high for the trans
port of Fe+3 and its complex ions. Similarly, although Fe+J can
form stable organic complexes but their formation in the Precarn-
brian with very little of organic matter is suspected. The
colloidal form of transport of Fe+J as a Fe(OH)3 micell, m
Fe(OH) 3 n Fe 0+ (n-x) Ci v Cl-, although possible but they would
Fig. 2.6
81
f=e2•
00 ' ' ' -0 2 .....
' ..... ..... -04
-06
-08
-10 0 2 4 6 1C 12 14 0 2 12 14
DH
D
04
02
00 ' '
-0 2 .....
' ' ' -0 4
-0 6
-0.8 ~.
- 1 0 f=et0Hl2 .. ·
c 2 4 10 12 14 0 2 4 6 8 10 12 14 oH
Relationships between iron compounds in various primary
sediments for activity of aFe= l0-2g ion/1; A= oxide,
B = silicate, C • carbonate; D= sulphide (as= 10-3 g
io~/1 (from Melnik, 1982)
82
require organic compounds like fulvic acid for a better stability.
But they completely precipitate in the presence of electrolytes
at a minimum concentration of 100 mgl-l which is nearly l-300th
part of their concentration in present day sea water (Ewer, 1983) •.
. ft.t the same time, the availability of organic compounds in the
Precambrian to form a stable colloidal solution is also a big
question.
Another method of colloidal form of transport of Fe+3
could be their stabilization by colloidal silica in the absence
of any organic compound. This mixed solution though very much
resistant to coagulation by majority of electrolytes but a
-2 trace amount of SO 4 affect their stability and at the same
time their formation is only possible with a super saturated
solution of silica nearly more than 200 mgl-l which appears
very much improbable.
Therefore, the main form of iron in dissolved state for
t be F +2 . . . f . f ransport seems to e 1n 10n1c orm as 1t never orms any
stable colloidal solution or complex ion. Its equilibrium con-
centration could be veryhigh in a reducing environment like that
of the Precambrian time devoid of free oxygen (Ewer, 1983;
Melnik, 1982). \
Silica usually occurs in monomeric form like Si(OH) 4
or H4sio4 over a wide range of pH (0-10) and changes to sili-
- -2 -4 cates of the type H3Si0 4, H2Si0 4 and SiO 4 at a higher pH
Fig. 2. 7
log a 5,
- 1
-2
-3
-4
Si02 Caml only
Solubility field of ions and precipitates of amorphous silica (from Melnik, 1981)
83
(Fig. 2.7). The solubility of amorphous silica at pH from 0
to 10 seems to be constant and amounts to 120 mgl-l with
increase of pH it increases' to 1120 mgl- due to the forma-
tion of charged ions. The major form of dissolved silica is,
therefore, H4sio4 which sometimes polymerize and condense to \
form colloidal particles. As the stability of H4sio4 depends
84
hardly on the physico-chemical parameters like Eh, pH or con~
centration of electrolytes, the mode of transport of Si02 in
the Precambrian could be same as that of the present. But
it is believed that the Precambrian hydrosphere was completely
saturated with silica or close to it for their transport and
precipitation whereas present hydrosphere is very much under
saturated, 1 mgl-l at surface and 10 mgl at depth, which could
be because of their intake by organism.
DEPOSITION
Deposition of both iron and silica under a peculiar
Precambrian condition to form banded iron-formation is the most
interesting part of the study of its origin. Although there is
no complete agreement on the nature of the depositional environ-
ment, and prqcess and mode of deposition but any depositional
model for ~he Precambrian banded iron formation has to explain
all its peculiar lithological and geochemical characteristics.
Depositional Environment
Depositional environment, in a limited sense, here means
the type of the depositional basins and their physico-chemical
conditions available for the precipitation of iron and silica.
The type of depositional basin is undoubtedly a result of the
then tectonics and the physico-chemical conditions and is cer
tainly an equilibrium product of atmosphere, hydrosphere, litho
sphere and biosphere.
Two major type of the Precambrian iron-formation, the
Algoma type and the Superior type, were perhaps deposited in
two different depositional environments.
It is presumed that the Algoma types associated with
greenstone belts and formed in Archean times were deposited in
small Archean intercratonic basins known as eugeosynclinal basins.
These basins were the major site for volcanism and iron-formation.
Goodwin (1973) has identified ten major Archean basins on the
Canadian shield and similar basins have been identified in other
parts of the world. These basins were thought to represent the
remants of the original quasi-circular structures which were bet
ween 800 to 1100 kms in diameter and at present they are hardly
350-700 kms long because of tectonic deformation. It has been
suggested that volcanoes along the margin of the subsiding basin,
which were exhuding calc alkaline and felsic lavas and pyro
clastics, supplied large quantity of chemical components to form
86
iron-formation. The precipitation of the$e chemical components
were largely dependent on the physico-chemicai conditions which
were probably determined to some extent on the depth of the basin,
and the distance between site of deposition and volcanic centres.
These basins were the ancient counterpart of the modern oceans
(Goodwin, 1973; Eichler, 1976).
The superior type iron-formations of the Proterozoic
time were believed-to have been deposited in miogeosynclinal
basins. These basins were platform environments in the sedi
mentary troughs between older cratons whereas eugeosynclinal
basins were epicontinental basin which were presumed to be the
younger depressions on the Archean cratons. These iron-forma
tions are considered to have been formed in shallow-water near
shore condition far away from volcanic centres. But there are
deposits that were probably formed in deep water far offshore
conditions (Gross, 1980). The Lake Superior and Labrador Trough
formations with their dolomitic association indicate a near
shore continental shelf deposition whereas the Krivoy Rog iron
formation has sedimentary features of a deep water, far offshore
deposition. Similarly, the Hamersley Group of iron-formation
seem to be a still deep water deposits under a stable tectonic
condition (Gross, 1980).
There is yet a debate whether these basins were insepa
rable parts of the oceans or restricted basins connected over bar
or completely separated basins to have a lacustrain environment.
\
87
Therefore, while White (1954), Holland (1973) and Dover (1974) have
favoured a marine environment, James (1954), Leep and Goldich
(1964), Cloud (1973) have supported a restricted marine basin
connected over ba~whereas Hough (1958) and Govett (1966) have
argued for a lacustrian environment for the deposition of iron~
formations.
Similarly, when Rubey (1951) argued that the Precambrian
was saline to hypersaline because of the presence of glauconite
and dolomite others like Cloud (1968) believed in nonsaline
nature of the Precambrian hydrosphere on the ground that the
microbiotes found in the Precambrian rocks are morphologically
similar to the microbiotas living in the fresh water.
Depositional Process
There is no controversy on the fact that iron-formation
being a chemical sediment, its constituents iron and silica reach
the depositional basins in dissolved state for a precipitation.
Precipitation of iron and silica takes place when physico-chemi-
cal conditions, like Eh and pH, of the solution carrying them
changes in the depositional basins affecting their solubility. \
But how these changes are brought about is a difficult question
to answer.
(A) Deposition of Iron: There are many who advocate for a
direct chemical precipitation of iron in the form of one of its
primary compounds when physico-chemical conditions are changed
in the depositional basin either due to the presence of electro
lyte or due to increase of pH and Eh. Some also propose a parti
cipation of biological organisms in the change of environmental
condition leading to a biochemical precipitation of iron. Some
also feel the chemical and biochemical process of precipitation
to go side by side or one changing over to other with time. -
Moore and Maynard (1929) propdse that the iron preci
pitated out when their colloidal solution stabilized by organic
matter come in contact with the electrolytes of sea water
(Lepp, 1975).
Goodwin (1956, 1960) was of the v1ew that acid solution
from the volcanic source containing iron in the dissolved state
when reached the depositional basin, increase? __ its pH and
gradually got concentrated to precipitate iron from its
supersaturated solution.
James (1966) also contributed the same view of initial
concentration in a restricted basin to attain a saturation level
for direct precipitation although he believed in a continental
erosion source.
Eugster and Chou (1972) have argued that the Precam
brian iron-formations were formed in an evaporative setting of
Playa Lake complex type. The precipitation of iron most probably
took place through the change of pH with the fluctuation ini
tiated by evaporation on the one hand and the flooding with
the fresh water on the other.
Holland (1973) and Drever (1974) while proposing an
oceanic upwelling model suggested that iron precipitated as
Fe(OH)3 in shallow marine water when upwelling sea water
brought them up to be oxidized by atmospheric oxygen.
Some investigators strongly advocate in favour of
the biochemical control of physico-chemical conditions for
the precipitation of iron minerals. As suggested by Stanton
(1972), there may be three possible ways for the biochemical
precipitation:
1. The organism modifying the physico-chemical nature
of its surroundings as a result of its own body function thus
creating a suitable condition for precipitation.
2. Organisms sometimes use iron to form an encrustation
over their body and deposit them when they die.
3. Sometimes organisms use iron as a source of electron
for their body function by its conversion from ferrous to
ferric state, (Fe++ Fe+++ +e).
Harder and Chamberlin (1915) has suggested the preci
pitation of iron oxide by bacteria in the Itabirite of Minas
Gerais, Brazil. Subsequently, Harder (1919) has demonstrated
89
' 90
by series of experiment, the ability of bacteria to precipitate
iron under certain conditions (Lepp, 1975).
Carroll (1958) while explaining the role of clay min-
erals for the transport of iron, has stressed the importance
of bacteria for the precipitation of ferric hydroxide or iron
sulphide.
Lepp and Goldich (1964) although supported the idea
of direct chemical precipitation of iron by pre-saturation also
believed that a low form of life such as bacteria or algae might
have played an indirect role in the chemical precipitation of
iron.
Becker and Clayton (1972),and Perry and Tan (1973) from
a radioisotopic study of c13 have concluded that iron and silica
precipitated in isolated basins but close to the proximity of
oceari. Iron was oxidized in photosynthetic zone by primitive
++ organisms using the Fe -H+ --~ Fe +e reaction.
Garrel et al. (1973) has proposed that initially,atmos-
phere was devoid of oxygen,and iron moved in ferrous form to
precipitate mainly as carbonate and silicates. But with time
procaryotic organisms started producing oxygen by photosynthesis
depositing iron as iron hydroxide biochemically. In deep water,
where organic materials and argillaceous materials got accumu-
lated, sulphate reducing bacteria produc~d H2s precipitating
91
FeS initially which later converted to Fes2 by decomposition.
Cloud (1973) has stated that blue green algae started
splitting water to release oxygen. But oxygen was sufficiently
lethal to these primitive lives. Therefore~organisms only
survived by transferring that oxygen to a nonbiological oxygen
++ acceptor such as H2o, alcohol or Fe . Therefore,1depositional
basins acted as oxygen pockets to precipitate iron as ferric
oxide.
Melnik (1982) has concluded that initially,under an
.oxygen free atmosphere,all the iron in the form of Fe+2, either
from a volcanic or from a continental source, were transported
to a depositional basin and accumulatedupto an equilibrium
concentration to precipitate chemically. The precipitation was
also accelerated by the change of pH of the solution that
carried iron to the depositional basin due to dilution and neu-
tralization. With the appearance of low form of life on the
earth's surface, oxidation reduction environments of depositional
&asins gradually changed. Oxidation of Fe+2 to Fe+J became
possible only after most of the methane in the atmosphere and
free carbon on the surface of the earth oxidized to co2 and
sulphur and hydrogen sulphide of hydrosphere to sulphate ion.
There were also simple organisms that react directly with Fe+2
causing its oxidation and deposition. The free oxygen did not
92
appear in the atmosphere till that time to create problem for
h . . F +2 f t e1r transport 1n e orm. The main iron precipitate at this
stage were siderite, greenalite and hydromagnetite. The next
important phase of the evolution was the appearance of phytopla-
nkton in the form of blue green algae which was the main produeer
of oxygen for a biogeochemical production of oxygen. These blue
green algae developed in localized basins at certain depth of
the water which was responsible to cut off the harmful ultra-
violet ray. Perhaps this stage was responsible for the vast
deposits of the Superior type iron-formation.
Ewers (1983) has mentioned about the photochemical
oxidation of Fe++ by ultraviolet light in the wavelength range
between 200-300 nm. Although it is a complex reaction, the
+2 overall reaction can be represented as Fe + 3H2o >
Fe(OH)3 + 2H+ + 1/2 H2. This took place under a reducing atmos
phere when there was no oxygen or ozone to prevent the ultra-
violet ray, and the evolving hydrogen to some extent helped in
maintaining the reducing atmosphere •. But this photochemical
method of oxidation was a complementary method to the method of
oxidation of iron as a by-product of photosynthesis as discussed
before,and,perhaps,went on till there was enough oxygen in the
atmosphere to cut off the ultraviolet light. He has indicated that
the precipitation of iron whether by oxidation of Fe+2 to Fe+3 and
93
subsequent hydrolysis to Fe(OH) 3 or due to the formation of carbo
nates and silicates are always associated with the release of H+
as shown below;
---?>> 4Fe( OH) 3 + 8H +
The hydrogen ion liberated got neutralized due to buffer action
between water and silicates. He has also suggested that the
initial precipitate could be only hydroxy oxide, which mighthave given
:rise• to carbonates, silicates, sulphides and magnetite due to
its reduction with free carbon from excess organic debris incor-
porated in primary sedimenttand subsequent reaction with active
reagents. He has, therefore, suggested the reaction of type
+2 Fe produced in this way could also react with other active
reagent in the immediate vicinity to form carbonates or sulphides.
Some object to the above photochemical oxidation of
Fe+2 on the ground that it is restricted to a low pH value, less
than pH 3. But at the photo oxidation of ferrocyanide ion
(Fe(CN)-~) at pH 11.8 in the presence of a suitable electron
scavenger has already been reported. So the confusion yet
exists (Ewers, 1983).
, (B) Deposition of Silica: Silica precipitation either as a
chemical precipitate or as a biochemical precipitate or as
both going on side by side has been well argued. But con
sensus is yet to reach.
94
Moore and Maynard (1929) has suggested that colloidal
silica stabilized by organic matter when reached the depositional
basin with high concentration of electrolytes, silica was
thrown out.
Goodwin (1956, 1960) was of the view that the acid
solution from a volcanic source carried enough dissolved silica
with it to an isolated basin to get sufficiently concentrated
for a direect precipitation .. ;..
James (1966) considered the similar method for silica
precipitation although his source was continental erosion.
Eugster and Chou (1973) have supported the idea of
chemical precipitation for silica and have argued that magadiite
or a sodium silicate gel was the precursor for the precipitated
silica. According to them, the precipitation took place in a
Playa Lake complex through the changes in pH because of fluctua-
95
tion initiated by evaporation and flooding with the fresh water
alternatively.
Holland (1973) ~hile proposing the oceanic upwelling
model throught of the precipitation of silica from the upwelling
water due to its decrease in solubility withthe decrease in
water pressure.
Some strongly oppose to the direct chemical precipi-
tation of silica on the ground that its solubility is little
affected by the change in Eh, pH between 1 to 9, and the com
centration of electrolytes in sea water (Krauskopf, 1956). The
amorphous silica in the dissolved state remains in the form of
monomeric H4Sio4 and can attain the saturation level of 100 to
150 ppm (Govett, 1966). Its precipitation seems to be a slow
process due to aGy change in the environmental condition and
never looks probable for a huge deposit of iron-formation (White
et al. 1956). Therefore, Bein et al. (1958), James (1966) and
Govett (1968) have proposed that the silica uptake by diatoms
could be a responsible factor. But the biochemical precipitation
of silica looked completely unlikely to Cloud (1973). He was of
-the opinion that there was no silica secreting procaryotes that
induced the precipitation in the Precambrian time and no eucaryotes
were known before 1.3 aeons. He, therefore, favoured the preci
pitation of H4sio4 to form Sio2 by polymerization due to decrease
of acidity to alkaline or neutral state which never looked
improbable.
96
After reviewing most of the Russian literature, Melnik
(1982) has indicated that the silica precipitation could be both
chemogenic and biochemogenic. He has proposed that, if the
source is volcanic, the hot acid water containing enough dissolved
silica caning in contact with sea water, converted part of
its dissolved silica to colloidal silica. This colloidal silica
coagulated for precipitation due to the increase of pH when acid
water got diluted and neutralized and also because of the presence
of electrolytes. The rest of the ionic silica either from the
volcanic source or from the continental source, no doubt got
accumulated in depositional basin to attain equilibrium concen
tration but their direct chemical precipitation is hardly possible
as their solubility is independent of Eh, pH ,over a wide range
from 3 to lO,and the presence of electrolytes. Therefore, he has
proposed that the possibility of the biogenic precipitation of
ionie form of silica, H4sio4?could not be ruled out,and perhaps
started from the Precambrian time for a huge deposit of iron
formation although no direct evidence yet exist (Melnik, 1982).
Ewers (1983) has discussed a few alternative methods for
the deposition of silica. One of the possible method could be the
evaporation of water to cause precipitation. He has indicated that
97
with a silica concentration of 120 mgl-1, for a deposit of iron
formation type, require a net evaporation of 3.3 meters of water
per year which is close to the upper limit of the present day
evaporation.
Another alternative method could be the freezing of a
surface layer of water causing precipitation of silica. But the
depth of ice formed and remitted annually seems to be insuffi
cient for a huge deposit.
Another interesting mechanism suggested involves the
co-precipitation of silica with Fe(OH) 3• It is attractive
because a single process links the precipitation of two major
constituents.
Thus none of the above mechanisms appear free from
objections and the guess work continues.
Mode of Deposition and Banding
Three different scales of banding with alternating bands
of iron minerals and chert are most interesting sedimentary
feature of the Precambrian banded iron-formation. The origin
of these bands are definitely a result of the mode of deposition
of iron and silica which demands a proper explanation.
VanHise and Leith (1911), and Goodwin (1956, 1962 and
1964) are the proponents of the volcanogenic theory, who have
considered each pair of iron and silica band to represent a single
98
contributary episode. A pulsating volcanic activity resulting
in the periodic entry of iron and silica into the depositional
environment gave rise to banding. Perhaps,iron because of its
high density precipitated quickly and settled first whereas
silica precipitated less rapidly and settled down after iron
due to its low density to give separate bands (Stanton, 1972).
Melnik (1982) has mentioned about a sinusoidal variation
of pH as an explanation of banding in a volcanogenic source.
According to that the deposition of iron minerals is accompanied
by liberation of protons and corresponding acidification of solu-
tion.
'3Fe +2 + 2H4Sio4 + H2o > Fe3si2o5(0H)4+6H+
Fe+2 + HC0-3
+ or > FeC03 + H
+2 + ~02 Fe2o2
+6H+ or 3Fe + 3H2o
Dropping of the pH to a threshold value ceases the precipitation
of iron minerals. The system waits for a neutralization to occur
to take the pH value above the threshold value for another batch
of precipitation to occur and the cycle is repeated for banding.
Some consider the microbanding as a product of some
periodic or seasonal change in the depositional environment.
This change could be either due to regular fluctuation in supply
99
of iron and silica to the depositional site or due to a periodic
fluctuation in the biological activity leading to a regular change
in the Eh & pH conditions for a rhythmic precipitation.
Sakamoto (1950) considered that ground water was acidic
in wet season to dissolve and transport iron to the depositional
sitejwhereas the ground water became alkaline in dry season to
carry silica. This alternative discharge of iron and silica
gave rise to fine microbanding (Stanton, 1972).
On the'other hand, Krauskoupf (1956), Huber (1959)
and James (1966) have suggested that the deposition of iron was
continuous throughout the year while silica deposition was
seasonal depending upon the explosive growth of the silica
accreting organism. It could explain the minor quantity of
iron in chert layer although some strongly object to the bio
chemical precipitation of chert.
Hough (1958) arid Govett (1966) have developed a lacus
trian theory to explain the origin of banding. They assumed
that lake water attained a density stratification or a ther
mal stratification seasonally. In summer, the surface water,
epilimnion, became hot and overturn took place in winter and
spring. In summer, iron came to the depositional basins when
the ground water was acidic and remained in hypolimnion but
the autumn overturn brought the iron to the epilimnion for an
100
oxidation and precipitation. Similarly silica came to the depo-
sitional basin in winter when the ground water was alkaline and
got precipitated in spring (Stanton, 1972).
Cloud (1973) stated that microbial photosynthesizers
flourished as long as ferrous iron and nutrients were handily
available to thempbut the organism died when there was a tem
porary depletion in Fe++ and nutrients. So the development of·
iron laminae was due to priodic flourish of microbial organism
depending on the availability of Fe++ and nutrients. Cloud
(1973) also believed that banding in iron-formation could be
better explained if it is assumed that there was an extensive
continental glaciation in the Precambrian time. The melting
of the ice could flood the continental margin where episodic
bloom of phytoplanktonic microbiatas due to plentiful supply
++ of Fe and nutrients could precipitate iron to form iron-
formation. This seasonal upwelling or variation in rate of
photosynthesis due to nutrients available, temperature and
light could account for the microbanding.
Drever (1974) in his depositional model which was
in line with Holland (1973), explained the banding due to
the variation in intensity of upwelling sea water which could
be due to seasonal effect and individual storms. According
to this model, upwelling water oxidized rapidly forming iron
layers and prolonged evaporation led to continuous and gradual
deposition of silica (Melnik, 1982).
Melnik (1982) has pointed out that as the precipi
tation of iron was dependent of the environmental conditions
like its redox potential and pH;whereas the precipitation of
silica was independent of above factors, thereforee, the depo
sition of silica was continuous whereas iron deposition was
periodic depending upon the variation in environmental con
ditions to give microbandings.
Cullen (1963) considered the mesobands that contain
101
the finer microbands as a product of tectonic movements (Stanton,
1972).Alexandrov (1973)and Trendall (1973) have indicated that
most common number of microbands in a mesobands was calculated
to be 19-21-23 which are close to 11 pairs corresponding to
a 11 year cycle of appearance of sun spots. Melnik (1982) has
considered the mesobands to represent a broad change in climatic
condition whereas each microband in a mesoband stands for
seasonal or annual changes.
Eichler (1976) has attributed the macrobanding to the
repeated change to the transgressive and regressive phases
which interrupted a constant deposition of iron-formation. A few
also consider this as a result of some tectonic changes.
102
Primary Deposition and Facies
The exact nature of the primary sediment in the case of
Precambrian iron-formation is difficult to say as the system has
undergone post-depositional changes. Therefore, it is almost
impossible to obtain first hand information on the primary phases
that were originally precipitated on the depositional basin.
Only through reference based on laboratory evidence of precipi
tation of chert and iron compounds at low temperature and pressure
and on theoretical calculation one can attempt to deduce what
could be the original compound (Klein, 1983).
James (1954) has considered the banded iron-formation
as a primary sedimentary rock and its oxide, silicate carbonate
and sulphide facies as deposits of different depth of the basin
where different oxidation-reduction conditions operated. But it
is now strongly believed that the present mineralogy is a product
of post depositional changes.
Melnik (1982) has quoted that Plaksenko considered the
iron hydroxide to be the only primary sediment. These iron hydro
xide precipitated down with unoxidized organic matter to get reduced
to Fe++ form and again to react with active compounds to form diff
erent facies. But the distribution of residual carbon in different
facies and the study of the isotopic composition has not supported
the above hypothesis.
Melnik (1982) has considered that the formation of
primary iron compounds had taken place depending upon the con-
centration of iron and active form of silica, carbon and sul-
phur. The deposition of these amorphous compounds on the
floor of the basin usually occurs sometimes after they were ·
103
formed in the water layer. Melnik has considered the following
as the primary precipitates that were initially formed: amor-
phous hydroxide of trivalent and divalent iron1Fe(OH) 3 & Fe(OH) 2;·
amorphous silica, Si02(a); finally dispersed crystalline mag
netite-Fe3o4(d); siderite Feco3;greenalite,Fe3si2o5 (OH) 4(d);
pyrite and pyrrhotite,Fes2(d) and FeS. Some do not consider
pyrite Fes2 as a primary deposit, rather a secondary product
from some monosulphides like FeS (Klein, 1983). These primary
sediments can be subdivided provisionally ~nto four groups:
oxide, silicate, carbonate and sulphide formally corresponding
to the James facies of iron-formation (Fig. 2.6).
Oxide Facies: The primary oxide facies is believed to be
mainly amorphous hydroxide of Fe+3 , Fe(OH) 3; in some cases mag
netite and in a very highly reducing condition, for instance
ln the sediments of swamps and peat bogs, amorphous hydroxide
'Of Fe+2. These are formed depending upon the Eh & pH condi-
tions of the depositional basin at a low concentration of active
form of silicic acid, carbonic acid and sulphur. On the Eh-pH
diagram (Fig. 2.6 ), it is clear that Fe(OH) 3 is deposited in
an oxidizing environment whereas magnetite is only stable in a
104
reducing environment. But Fe(OH) 3 is confined to a stability
field of highly reducing conditions. The evolution of environ-
mental conditions of the depositional basin like its Eh is a
result of the general development of atmosphere and biosphere.
In the initial stage of tbe evolution of atmosphere,,
when the conditions were reducing, perhaps the main component
of the oxide sediment was metastable dispersed magnetite
Eh = + 1.229 -0.236 pH - 0.088 log a Fe
With tne increase of Eh due to flare up of life activity, perhaps
unstable Fe(OH) 3 were formed to be converted irreversibly to
magnetite.
2+ + Fe + 3H2o = Fe(OH) 3 + 3H +e
Eh = 0.941 -0.177 pH - 0.059 log a Fe
+ 3 Fe(OH) 3 + H + e
Eh = 0.366 - 0.059
1In a later stage of atmosphere development when Eh was suffici-
ently high the stable oxide facies appeared as Fe(OH) 3
105
++ + Fe + 3H20 = Fe(OH) 3 + 3H +e
0.941 - 0.177 pH - 0.59 log a Fe
There are many who question the primary origin of magnetite. When
some consider it is a diagenetic reaction product of either hydro-
magnetite Fe 3o 4 n H2o ; or a mixture of hydroxides,1 Fe ( OH) 3 & Fe ( OH) 2 ;
others consider it as a secondary mineral formed either from the
oxidation of silicates and carbonates or due to the reduction of '
hematite (French, 1973; Klein, 1973).
Silicate Facies: As Melnik (1982) has pointed out the deposi-
tion of amorphous silica occur only in the case of fairly high
concentration of dissolved Sio2 over a wide pH range (0-10) in
the presence of active form of silicic acid, in oxidized envir-
onment an association of Fe(OH)3 + Sio2 is formed, and in reducing
environment a finely dispersed crystalline silicate appear.
Greenali te is believed to be the primary silicate which is
converted to other silicate under post deposition transformations.
From the Eh-pH diagram, it is obvious that silicates are formed
only in reducing environment in the absence of free oxygen. Repla-
cement of the field of magnetite Fe3o4 by the field of silicate
Fe3si2o5(0H) 4 theoretically explains their association. \ '
Carbonate Facies: Melnik (1982) believe that carbonates were
formed when iron was precipitated in the presence of dissolved
106
carbonic acid or as an interaction of primary hydroxide with or
ganic carbon. In Eh-pH diagram (Fig. 2.6), it is confined to a
low Eh and approximately neutral pH condition. Therefore, the
deposition of siderite could occur only in oxygen free environment
in early stage of development of atmosphere that was in the Archean
or beginning of the Proterozoic with increase of Eh, probably
in the Proterozoic, led to an 2ssociation of Feco3 and Fe(OH)3:
A lower Pco2
and corresponding high pH perhaps led to the associa
tion of, siderite +magnetite.
Sulphide Facies: Melnik (1982) has indicated that chemogenic
depositior: of iron sulphide occur as a result of interaction of
dissolved iron with active sulphur. It is yet not certain whether
pyrite or pyrrhotite is primary or the diagenetic product of some
monosulphides. On the Eh-pH diagram (Fig. 2.6), it is understood
that the sulphides can be deposited in presence of a high con
centration of active sulphur at a wide pH. value mainly in reducing
conditions. It replaces the magnetite field to some extent.
The deposition of oxide-carbonate-silicate-,and sulphide
iron sediment depend on the joint interaction of iron compound
with active form of silica acid, carbonic acid and sulphur at a
particular Eh and pH value. So it gives rise to different facies
association depending or. the available environmental conditions.
107
POST DEPOSITIONAL CHANGES
i. DIAGENESIS
Melnik (1973) has defined the term diagenesis as those
processes which derive their energy essentially from the thermy
dynamically spontaneous transitions of the metastable materials
or assemblages toward stable equilibrium phase (Ewers, 1983).
It has also been explained that freshly formed sediment repre-
sent a physico-chemical system that is out of equilibrium. The
process of alteration of sediments leading to the equilibrium under
the thermodynamic conditions of the earth's crust before conver
sion into ~ock a~e called diagenesis (Melnik, 1982).
The process occur close to the sediment water inter
face and the mechanism involves compaction, expulsion of water
from the original sediment, crystalization of original gel, clay
and colloidal particle and the precipitation of mineral at a
low temperature below 100°C (French, 1973).
Like primary deposition, any first hand information on
the process of diagenesis is not available and conclusions are
mostly drawn based on the experimental observations and theoreti
cal calculations. The mechanism involved is very much complex
and yet very much unclear.
Melnik (1982), on the basis of experimental therno-
dynamic analysis has concluded that diagenesis of hydroxide
(Fe (OH) 3 ) sediments in an oxidizing environment results in
the crystallization c·f amorphous and dispersed hydroxide with
the formation of stable gpethite or metastable dispersed hema-
tite, whereas the reducing condition leads to the formation of
megnetite. Diagenesis completely converts the hydroxide of
108
divalent iron, Fe(OH) 2, into magnetite. It changes the freshly
precipitated greenalite to its dense crystalline form. Siderite
also undergoes the similar change to a well crystalline form and
finely dispersed pyrite into its crystalline r,;odification. All
these changes lead to better stabilization and, hence, are asso
ciated with liberation of energy. These phase transformations
are irreversible and therefore spontanec:us with liberation of
energy.
Similarl~silic2 which is initially deposited as a gel,
under diagenetic conditions is converted to crystalline quartz.
Melnik (1982) from his experimental study has established the
following chc..nge of transformation
Opaline ::;:.. Silica
ii. METAMORPHISM
SiO -X-SiO -X 2 2 Crystobalite quartz
We can define metamorphism as a post depositional change
109
after diagenesis which involves the participation of intergranular
liquid under the change of temperature and pressure. So metamor
phism is a product of hydrostatic pressure, which is equal to the
fluid pressure plus the directed pressure in the solid phase.
Fluid pressure is mostly due to the partial pressure of the water
and the partial pressure of co2 dissolved on it. Change of physi
co-chemical parameters brings the equivalent changes in partial
pressures or activities of H2o and co2 in the intergranular
fluid leading to different metamorphic reactions. There is no
sharp boundary line between diagenesis and metamorphism to dis
tinguish,one from the other and the changes are mostly gradational
(French, 1973). Klein (1973) has, therefore, classified the post
depositional changes in to three stages as (1) Diagenesis to
very low grade metamorphism, (2) Medium grade metamorphism, and
(3) High grade metamorphism. This metamorphism either leads to
recrystallization or to the formation of new minerals. Develop
ment of amphiboles indicates the meoium grade metamorphism
whereas the appearance of fayalite indicates the high metamorphic
conditions.
Metamorphism of oxide facies: Metamorphism of oxide facies
~fter diagenesis leads to the formation of either finely banded
quartz hematite or quartz-magnetite or quartz hematite-magnetite
110
with some hydrous iron oxide. With increase of metamorphlc grad~~
the recrystallization of chert and iron oxide takes place with
increase of their grain size. Therefore, sometimes the grain
size of the chert or quartz in a chert on quartz richhorizon,
acts as a good indicator of the metamorphic grade the rock has
undergone. In low grade metamorph~sm, the average grain size
of quartz is less than 0.1 mm, in medium grade metamorphism
it ranges from 0.1- to 0.2 mm 1and in high grade it becomes
greater than 0.2 mm (James, 1955; Klein, 1973).
Magnetites under a low metamorphic grade seems to be
well crystallized and coarse grained than co-existing chert,
hematite and iron silicate. The origin of magnetite is yet
controversial, when some consider it as a diagenetic product
of either hydromagnetite Fe3o4 n H2o or mixture of iron
hydroxides Fe(OH) 3 and Fe(OH)z Others consider it as a
secondary mineral formed either by the oxidation of carbonate
or silicate. The reduction of primary ferric hydroxide to
produce magnetite is also presumed by many from the petrographic
textures. The increase of metamorphic grade only re~rystallizes
and increases the size of magnetite grains (French, 1973;
Klein, 1973).
Hematite 1n low grade metamorphic condition is very fine
grained and the precursor for it is assumed to be Fe(OH) 3 which
Ill
get converted to goethite FeO(OH) under diagenetic condition.
Increase of metamorphism simply recrystallizes and changes its
grain size.
In the absence of the reducing agents, metamorphism
only leads to recrystallization and making the minerals like
hematite, magnetite and quartz more pure. The associations
quartz-hematite, quartz-magnetite and quartz-hematite-magnetite
are well preserved in different grades of met&~orphism. How-
ever, the iron formation towards the contact aureoles in the
. Duluth Gabbro complex shows a replacement of hematite by
In the presence of reducing agents like organic com-
pounds or free carbon, a sequence of metamorphic reaction has·
been proposed based on thermodynamic and kinetic factor,H einatite
-->> Ma .gneti te --....::::>~ Siderite -->> Magnetite Fayalite
In the case of reduction by hydrogen, the following course of
reaction has been suggested.I·Jematite -->.;:.Magnetite -----,>;;.-
f'\ M innesotai te -->> Granuri te -~> Fayali te. It has also
been proposed that hematite is not stable at granulite facies
conditions. So it changes to magnetite (Melnik, 1982).
Metamorphism of carbonate facies; The carbonate facies of
iron-formation consists of chert or quartz, carbonates (sider-
ites, members of dolomite-ankerite series and calcite) with a
112
small amount of magnetite and iron silicate. They are also asso-
ciated with graphite and pyritic material (Klein, 1973).
Original precipitate is taken to be a finely crystall~ne
precipitate which recrystallized under metamorphic conditions.
Under medium to ·high grade metamorphic conditions, carbonates
generally give way to other mineral assemblages although origi-
nal carbonate is still preserved at places with recrystalliza-
tion and increase of their grain size. Retention of carbonates
under high metamorphic conditions is possible if the C02 par
tial prelssure in the interstellar fluid is high to prevent
the break down of carbonate. If pco2 is low, carbonates reacts
with quartz to form new silicates as follows.
Ca(Fe, Mg) (co3) 2 + 2Si02 ~ Ca(Fe,Mg) Si2o6 + 2C02 Ferro dolomite Clinopyroxene
(Fe, Mg) C03+ Si02 ~
Siderite
(Fe, Mg) sio3 + co3 Orthopyroxene
If the chemical potential of water is high while the chemical
potential of co2 is low, Ca(Fe, Mg) ~co3 ) 2 + 8Si02+ H2o~ Ferrodolom1te
( +2 ) ' Fe , Mg 7 Si8 22 (OH) 2 Grunerite
' 8(Fe, Mg) C03+
Siderite
9Si02+ H20 ~ (Fe,Mg)? Si8o22 (OH) 2 Grunerite + (Fe,Mg) Si03+8C02
Orthopyroxene
113
The petrographic evidence for the above reactions have
· been well established (Klein, 1973). ~
The petrographic evidence of oxidation of siderite to
magnetite in iron-formation of Superior type is well marked
(French, 1973)
>
Hematite and siderite association appear to be rare as
they react to form~Magnetite, Siderite + Hematite ~ Magnetite
(Melnik, 1982).
As the facies are mostly gradational, a mixed facies of
carbonate-silicate type undergoes a different sequence of reac-
tion depending on the composition of the original sediment, the
partial pressure of co2 & H2o ~n the fluid and other reducing
agents.
The possible sequence of reactions are follows ·
Greenalite + &derite --~>~ Minnessotaite + Siderite
Grunerite + siderite
Grunerite + Magnetite
Grunerite
-----=>~ Gruneri te + Fayali'te >
Fayalite. The end product is mostly fayalite (Melnik, 1982).
' Metamorphism of silicate facies: Silicate facies in a low grade
metamorphic iron-formation consists of iron minerals such as
greenalite, stilpnomelane, minnesotaite, less amount of riebeckite,
114
ferriannite, ripidolite and ferrantate. They are also associated
with carbonates (siderite, members of dolomite-ankerite series
and calcite) and iron oxide mostly magnetite (Klein, 1973; Klein, '
1982). Among all the above m1nerals, greenalite with approximate
formula (Fe,Mg) 6 Si4o10 (OH) 6 with certain amount of aluminium
and alkali metals as impurities exhibit most primary texture to
be considered as a primary or diagenetic precursor. Many greena-
lites are criss crossed and transected by two other silicates,
stilpnomelane and minnesotaite which are presumed to be secondary
minerals formed from greenalite in diagenetic or low grade meta-
morphic stage. Stilpnomelate could be from a precursor greenalite
which contained some amount of K+ and Al2o3•
Petrographic evidence indicates that minnesotaite could
be a reaction product of either greenalite or stilpnomelane or
quartz + carbonate as per the following reaction (Klein, 1982)
Fe6si4o10 (OH) 8 + 4Si02 > 2Fe3si4 o10(0H) 2 + 2H2o
Greenalite Minnesotaite
Greenalite Minnesotaite
FeC03+ Sio2 + H2o
Siderite
Fe2. 7(si, Al)4(o, OH) 12
Stilpnomelane
>
X
>
Fe3si4o10(0H) 2 +3Co2 Minnesotaite
H20 + 0.33 Fe +2
Fe3si4o10(0H) 2+H20+Al,
Minnesotaite
Na, K
us
Other silicates like chamosite and rapidolite also occur
along with greenalite and their formation could be because of high
content of Al2o3 in the original sediment (Melnik, 1982). Simi
larly, ferriannite, riebeckite or its fibrous variety crocidolite
are rare minerals in diagenetic or low grade metamorphiSm whose
origin are not very clear.
The above silicates in a low grade metamorphism are
frequently interbedded, with quartz-magnetite-carbonate horizon
and leads to a complex reaction when proceed to medium grade
metamorphism giving rise to amphiboles of commingtonite-grunerite
series, pyroxene rich schist and gneisses.
If the silicates are poor in alumina like minnesotaite,
they metamorphosed to give amphiboles belonging to grunerite
commingtonite series. Carbonates, if present, are also metamor
phosed to above amphiboles. The types of reaction that give
rise to these amphiboles are as follows (Klein, 1973).
Minnesotaite Grunerite
7Ca(Fe, Mg) (co3)2+8Si02+Hl) > (Fe,Mg)8si8o22 (0H) 2
Ferrodolomite Grunerite +7CaC03+7C02 8(Fe, Mg) C03+8Si02+H2o > (Fe,Mg) 7sr8o22(0H) 2+7C02
Siderite Grunerite
116
If shales are present which are rich in Al2o3 or silicates
of type iron-magnesium chlorites or mica of stilpnomelane are
there in low grade metamorphism, they give rise to garnets of
almandine type or sometimes to hornblende when passes into medium
grade metamorphism.
Similarly a local reduction of pH2o or increase of pco2
in the interstellar fluid converts the low grade iron silicates
to pyroxene group minerals and they are sporadically present in
medium grade metamorphism.
The aluminium free silicates under high metamorphic
conditions give rise to fayalite which are sometimes associated
with ferrohypersthene, garnet (almandine) and magnetite. In
isolated cases the amount of magnetites increases and becomes a
major mineral. If the iron minerals are rich in MgO they get
converted to magnesium-iron pyroxene instead of fayalite.
Metamorphism of sulphide facies: These are finely banded black
rock with variable amount of carbon, pyrite and chert, in which
pyrite mostly occur in thin layers. At places they are associa
ted with siderite, stilpnomelane and chlorite. Some has presumed
mackinawite (Fel+xS) as the sedimentary precursor to pyrite.
The increase of metamorphism increases the original size
of pyrites and under high grade metamorphic conditions pyrite is
converted to pyrrhotite. Carbonates and silicates most of themrich
117
in Al20jcJ metamorphosed to give orthopyroxene, clinopyroxene
and fayalite rich assemblages. The amorphous carbon metamorphosed
to give graphite (Klein, 1973, 1982).·
iii. WEATHERING
Weathering usually implies decomposition of rock at or
near the earth surface by physical and chemical processes. It
is a part of supergene alteration wnich lead to iron enrichment,
and even marked to a depth of few kilometers. Morris (1983)
has made a good discussion on the weathering aspects of BIF.
Physical weathering: According to Morris (1983) physical
weathering opens the door for chemical weathering. Mineral
or root swelling, animal burrowing and'fire are local phenomena
resulting in physical weathering. Frost fracturing plays a pro
minent role in cold climates and fatigue resulting from repe
titive temperature cycling combined with hydrolysis also/result
in rock failure.
Chemical weathering: As Morris (1983) has pointed out, chemical
weathering involves the process like ionization, hydrolysis, oxi-
dation and complex formation leading to leaching and enrichment.
Leaching leads to depletion of Mg, Ca and co2 from carbonates,
K, Mg and Si from silicates, and Ca and P from apatite.
Morris (1983) has suggested a mecr2nism for a deep
seated weathering and enrichment. According to him anodic
oxidation at depth can be represented as Fe+2 ~~~~ Fe+3 +e
and the electron is conducted by magnetite layers to out crop,
where oxygen is reduced at the surface method: o2+ 4e + 2H2o
---~> 40H~ while at depth the ferric ion hydrolyses and pre-
. . F +3 3H 0 c1p1tates, e + 2
reacts with ferrous carbonates and silicates to continue the
process. Chemical weathering is mainly governed by ·the pH,
Eh conditions and volume of water involved in hydrolysis and
118
has different effect on different minerals. C1emical weathering
leads to dissolution and leaching of Sio2 although the exact
reason is not very much.clear. Fe+2 in iron carbonates and
silicates get oxidized to Fe+3 which is precipitated as goethite
or as Fe(OH) 3 with release of H+ to accelerate the further
weathering. Similarly magnetites also seem to be affected
by weathering and gets converted to hematite through some
intermediate stages. Among the minerals hematite seems to be
totally resistant to weathering (Morris, 1983).