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Archean Greenstone Belt
1. Greenstone Belt
Greenstone belts are generally elongate, Archean to Proterozoic
terrains comprising intrusive and extrusive mafic to ultramafic
igneous rocks, felsic volcanics, and inter-flow or cover
sedimentary rocks.
Greenstone belts occur sandwiched between regions dominated by
granitoids and gneiss.
Greenstones are generally of low to moderate metamorphic grade.
The term greenstone comes from the green color of many mafic to
ultramafic constituents due to an abundance of chlorite.
A common igneous rock in greenstones is komatiite. Komatiites are
rocks with greater than 18 weight percent magnesium oxide and a
well-developed spinifex texture of inter-locking bladed or acicular
crystals of olivine or pyroxene.
Spinifex texture (named after similarities in crystal shape and pattern to
the spinifex grass that grows in South Africa and Western Australia)
implies rapid cooling or decompression of the magma.
Komatiites formed as volcanic flows and less commonly as intrusive sills.
Sedimentary sequences within greenstone belts comprise both clastic
(e.g., conglomerate, quartz arenite, shale and graywacke) and
chemically precipitated (e.g., banded iron formation and chert)
components. Greenstones may also be intruded by syn-to post-tectonic
Greenstone belts host many major mineral deposits, such as gold and
nickel. Greenstone belts were previously often thought to continue to
large depths in the crust. Reflection seismic profiles over the Norseman
Wiluna Belt of the Yilgarn Craton, Western Australia, however, indicate
that this greenstone belt has a relatively shallow (6–9 km) flat-base and
overlies a uniformly thick crust.
Contrasting models have been proposed for the origins of greenstone
belts. Some geologists believe magmatic and tectonic processes during
formation of greenstone belts in Archean times were different to presentday plate tectonics. Earth's mantle would have then been far hotter. They
cite differences between greenstone belts and Phanerozoic orogens (such
as the abundance of komatiitic lavas) and point out that there are no
modern analogues to greenstone belts. Opponents to Archean plate
tectonics contend that greenstone belts commonly represent a laterally
continuous volcano sedimentary sequence (sometimes on a granite-gneiss
basement) essentially undeformed prior to late tectonism and may not
therefore represent relics of volcanic chains. They consider that
Archean tectonics was dominated by mantle plumes and was possibly
analogous to the tectonics of Venus. Greenstone belts are interpreted as
oceanic plateaus generated by mantle plumes, similar to plume-generated
oceanic plateaus in the southern Caribbean. A mantle plume origin is also
proposed for neighboring tonalite-trondhjemite-granodiorite sequences.
The alternate view is that tectonic processes comparable to presentday plate tectonics were operative during the Late Archean, and
possibly were similar to plate tectonics since the Hadean-Archean
transition (between 4.0 and 4.2 billion years ago). In a plate tectonic
context, greenstones may have formed in volcanic arcs or inter-arc or
back-arc basins. Greenstone belts are interpreted to represent collages
of oceanic crust, island arcs, accretionary prisms, and possible
plateaus. Recent experimental work on the origin of komatiitic magmas
indicates that they were hydrous and that temperatures for their
formation do not indicate that the Archean upper mantle was
significantly hotter than today. Komatiites and similar rocks have also
been found in younger orogens. Komatiites may not therefore require
different tectonic processes or conditions for their formation, as
previously thought.
In many granitoid-greenstone terrains, greenstone belts constitute synformal
keels between circular to elliptical granitoid bodies. This outcrop pattern is
generally thought to be due to deformation resulting solely from the greater
density of greenstones compared to underlying granitoid and gneiss. Due to
gravitational instability, the underlying, less dense granitoid-gneiss basement
domed upwards and rose to form mushroom-shaped bodies called diapirs
whilst the denser greenstones sank into the basement. Shear zones were
formed along some granite-greenstone contacts due to differential vertical
displacement and upright folds developed in the greenstones. This process
of either solid-state and/or magmatic diapirism was independent to any
tectonic processes that may have acted on margins to granite-greenstone
terrains. The formation of granitoid domes in granite-greenstone terrains has
also been attributed to crustal extension (producing metamorphic core
complexes) or polyphase folding during regional shortening. The more linear
form of some greenstone belts is due to subsequent deformation, especially
the superposition of regional-scale transcurrent shear zones on earlyformed structures.
Isua greenstone belt (Greenland)
The Isua Greenstone Belt is an Archean greenstone belt in
southwestern Greenland dated at 3.8-3.7 Ga and contains the oldest
known, well preserved, metavolcanic (metamorphosed mafic
volcanic), metasedimentary and sedimentary rocks on Earth. It
consists of five tectonic domains.
Almost all the rocks are deformed and substantially altered by
metasomatism, however the transitional stages from the volcanic and
sedimentary structures to schists can clearly be seen. The different
episodes of metasomatic alterations can also be seen that produced a
diversity of metamorphic mineral assemblages from similar protoliths. The
metamorphic processes of garnet-growth events span both the early and
late Archean. New geological mapping studies are tracing the transitional
gradations between the protoliths and their diverse deformed and
metasomatised structures. These new mappings show that most of the
Isua Greenstone Belt consists of fault bounded rock assemblies
(1) derived from basalt and high-magnesium basaltic pillow lava and
pillow lava breccia,
(2) intruded by numerous sheets of tonalite,
(3) intercalated with chert-banded iron formations, and a minor
component of clastic sedimentary rocks derived from chert and basaltic
volcanic rocks.
It is thought that the recrystallized ultramafic bodies that occur in the
belt are intrusions or komatiitic flows. Studies show that these
komatiites are extremely similar to the 3.5 Ga Barberton basaltic
komatiites of South Africa, and are both
Archean equivalents of modern boninites produced by hydrous
melting in subduction zones.
The Barberton komatiites share some of the same geochemical
characteristics with modern-day boninites, including petrologic
evidence for high magmatic water content. The boninitic
geochemical signatures provide evidence that plate tectonic
processes are responsible for the creation of the belt, and that the
pillow breccias and basaltic debris indicate that liquid water existed
on the surface at the time of their formation. The most common
sedimentary rocks are the chert banded iron formations.
The 3.5 Ga Isua basalt-komatiite-chert are parental to the enclosing 2.8
Ga Amitsoq Tonalite-Trondhjemite-Granodiorite (TTG) gneisses suite.
Current arguments favor a direct mantle melting to produce the diorites
and high magnesian granodiorites found in these Archaean cratons.
According to Rapp (1999): "A full continuum of processes can be
envisioned from the generation of "pristine" TTG magmas by wet melting
of garnet amphibolite/eclogite, to hybridization of these melts by
assimilation of peridotite, and the consequent metasomatism of the subcontinental mantle. The Mg-number of Archean granitoids is perhaps the
best indicator of TTG lineage, and the extent to which the mantle was
involved in early continent formation. The evidence thus far suggests an
increasingly important role in the late-Archean, with crustal, oceanic
basaltic sources dominating in the early-Archean."
Quartz globules from undeformed pillow breccia are associated with a
complex system of quartz veins and are interpreted as remnants of a seafloor hydrothermal system, operating at the time of lava eruption and pillow
basalt formation at 3.75 Ga. The quartz globules are thought to be
interpreted as former gas vesicles filled with quartz and carbonate, and are
embedded in an altered basaltic matrix now comprising biotite, muscovite
and quartz. Silica-filling in the vesicles is thought to have been
contemporaneous with the formation of an intricate hydrothermal vein
system. During deformation, the strain is thought to have been partitioned
into the mica-rich rock matrix, when the vesicles behaved as competent
objects, eventually cutting or deforming the thinner veins. Fluid inclusions
from quartz in the vesicles resemble present-day sea-floor hydrothermal
fluids and are thought to be responsible for the alteration of the pillowed
breccia and co-precipitation of quartz and carbonate. Therefore, these
carbonates may represent the products of an early sea-floor.
In the Southwestern Isua Greenstone Belt, Kyanite has been found in
the muscovite-rich schists and was formed by metasomatic reaction.
This shows that this part of the belt was subjected to deformation in
high strain zones and high grade metamorphism during the late
Archaean. According to a study by Frei and Rosing (1999): The kyanite
are Al-rich metasomatites formed by transformation of felsic gneisses
by fluids derived from ultramafic schists. Peridotite and dunite bodies
were transformed into talc-anthophyllite-chlorite-magnesite schists
during prograde reaction with ambient fluids. During amphibolite facies
metamorphism and deformation, the ultramafic schist released fluids,
which interacted with neighboring lithologies. These fluids were
buffered at silica the talc-anthophyllite-magnesite mineral
assemblage. At this low silica activity, chlorite in the ultramafic schists
buffered (aAl+)/(aH+)3 activity ratios close to that of corundum
These fluids caused silica depletion, and high aluminum ion
activities that intersected kyanite saturation in the quartz
saturated felsic rocks. Progressive leaching in three steps
defines a Pb-Pb isochron with an age of 2847 ± 26 Ma. This age
is consistent with a Sm-Nd amphibole-plagioclase isochron of
2849 ± 116 Ma from mafic schists at Isua, and a lead (Pb) step
leaching age of 2840 ± 49 Ma (MSWD=1.43) on magnetite from
Isua Banded Iron Formation (BIF), and with regional granulite
facies metamorphism in other parts of the Greenland Archaean
Belingwe Greenstone Belt (Zimbabwe)
Samples taken from the NERCMAR drill hole in the 2.7 Ga Manjeri
Formation in the Belingwe Greenstone Belt contain oxide and sulphide
ironstones that are indicative of a complex bacteria/archaea eclogical
community. The REE compositions imply an oceanic deposition similar
to that of the late Archean. (Bickle et. al, 1999).
The Belingwe Greenstone Belt contains a 7 km succession of mafic
and ultramafic lavas and high-level intrusions which overlie a thin
sedimentary formation, itself unconformable on a granitic basement.
The lavas range in composition from andesites (4 per cent MgO) to
peridotitic komatiites (32 per cent MgO). The mineralogy and textures of
the most magnesian lavas demonstrate that they were extruded in a
completely liquid state. If the source mantle had an MgO content around
40 per cent, then partial melts in the range 35 per cent to 55 per cent
would be required to produce the most magnesian liquids observed.
Physical constrains on the origin of the mafic and ultrafic lavas
Imply a derivation from a depth of >150 km, at temperatures of
1600-2000 oC (Nisbet et al., 1977).
A Review of Greenstone Belts in the Superior Province and the Evolution
of Archean Tectonic Processes. (Development of Abitibi Greenstone Belt)
Kirsty Y. Tomlinson
Australian Centre for Astrobiology, Department of Earth and Planetary Sciences,
Macquarie University, NSW, 2109, Australia.
The origin and development of Archean greenstone belts continues to be
strongly debated, particularly with regard to the roles of subduction, plume
magmatism, rifting, diapirism and autochthonous vs allochthonous development
(e.g. de Wit, 1998; Hamilton, 2003). It is apparent from studies in the Superior
and Slave Provinces of Canada that strongly contrasting tectonic
styles may have been in operation at the same time. For example at ca. 2.7 Ga,
large diapiric batholiths and synclinal greenstone keels may suggest that
diapirism was an important tectonic process in the Slave Province (Bleeker,
2002), whereas in the Superior Province the linear distribution of belts suggests
that accretionary tectonics (i.e. plate tectonics) may have dominated
(e.g. Stott, 1997). Neither theory precludes the other, and in developing models
for Archean tectonic evolution, no one model will be equally applicable to all
Greenstone belt types
In the Superior Province, greenstone belts are typically collages that
contains more than one sequence of rocks. These sequences may
have different ages and distinct histories. There are several
common types of sequence (or stratigraphic associations) that are
found in volcanic dominated greenstone belts, and their features
indicate certain environments of formation. These features can also
be used to suggest tectono-magmatic settings but are usually open
to interpretation.
1) Flood volcanism on submerged (shallow water) continental
platforms – these sequences contain thick, laterally extensive
tholeiitic mafic flows, pillow basalts, hyaloclastite ± komatiites and
minor amounts of felsic tuff, BIF, cherty and clastic sedimentary units.
Rarely unconformities are preserved at the base of these sequence,
where thin conglomerate-quartzite-arkose-carbonate units overly
tonalitic basement. More commonly the base of the sequence is not
preserved but detrital zircon ages in sedimentary rocks correspond to
the age of nearby granitoid rocks which are inferred to represent
basement. The volcanic rocks may also contain xenocrystic zircons,
Nd isotopic evidence of older crustal involvement, and geochemical
signatures suggesting felsic crustal assimilation. These sequences
typically occur early in the development of Archean cratons and are
common at 3.0-2.9 Ga in the Superior Province (North Caribou and
Marmion terranes). In the North Caribou terrane evidence of graben
development exists in the basement suggesting extension prior to
volcanism, and plume-driven rifting has been suggested as the
environment of formation. This environment does not have a true
modern analogue as the continental flood basalts on the modern
Earth are subaerial.
2) Submarine volcanic plains – comprise massive and pillowed
tholeiitic mafic flows ± komatiite, BIF, mudstones and rare
greywackes. The lavas have juvenile Nd isotopic signatures, lack
significantly older zircon inheritance and have primitive mantle
normalised REE profiles that are flat to slightly depleted in light REE
(but less depleted than modern MORB) or slightly enriched in light
REE. The sequences often have tectonic contacts and may be
fragments of thicker sequences. They appear to have formed in
oceanic environments, and suggested tectonic settings are oceanic
plateau, back-arc basin, primitive island arc, oceanic island or
possibly mid-ocean ridge. These sequences often flank older
continental blocks and may have been accreted to continental
margins. Examples are the older (ca. 2.88 Ga?) parts of the OxfordStull Lake terrane flanking the northern margin of the North Caribou
terrane; the older (ca. 2.78 Ga) parts of the Western Wabigoon
terrane flanking the Winnipeg River and Marmion terranes; and parts
of the ca. 2.75-2.70 Ga Abitibi subprovince.
3) Diverse volcanic sequences – comprise a variety of submarine to
lesser subaerial units, dominated by basalt, with lesser andesite,
dacite and rhyolite flows, and dacite-rhyolite pyroclastic units. In some
belts komatiites are also present. Both tholeiitic and calcalkaline
signatures are often observed, with basalts ranging from slightly light
REE depleted (but less depleted than modern MORB) to moderately
light REE enriched with negative Nb anomalies (similar to modern arc
related basalts). These sequences are commonly associated with synvolcanic granitoids. Nd isotopic data suggest that some sequences are
juvenile while others have experienced minor interaction with older
enriched sources (metasomatised mantle or crustal contamination?).
They are suggested to represent arc magmatism (both island arc and
continental arc on a thin margin), arc/plume interaction, arc rifting or
back-arc magmatism. Examples are 2.83 Ga sequences in the OxfordStull Lake terrane flanking the northern margin of the North Caribou
terrane; various ca. 2.9-2.74 Ga sequences in the Uchi subprovince at
the southern margin of the North Caribou terrane; 2.75-2.71 Ga
sequences in the Western Wabigoon terrane; and parts of the ca.
2.75-2.70 Ga Abitibi subprovince
4) Continental felsic volcanic centres – comprise thick sequences
of massive calc-alkaline dacite-rhyolite flows and pyroclastics with
lesser calc-alkaline basalts and andesites and syn-volcanic plutons.
They may unconformably overly older tholeiitic to calc-alkaline
sequences but basal contacts are usually tectonised or intruded by
younger granitoids. Indirect evidence for eruption on older basement
occurs in the form of older Nd model ages, zircon inheritance, evolved
geochemical signatures and the maturity of the sequences. They are
suggested to represent continental arc magmatism and examples are
widespread at 2.73-2.71 Ga in the western Superior Province,
occurring along both the northern and southern flanks of the North
Caribou terrane (in several greenstone belts along a 2000 km long
continental margin), and in the eastern part of the Wabigoon
5) Late alkaline-shoshonitic sequences – occur locally in the
Superior Province associated with late transpressional faults (at ca.
2.71-2.68 Ga, younging from north to south), following the major N-S
shortening event. They comprise alkaline volcanic rocks with evolved
Nd isotopic signatures and geochemistry, and continent-derived
alluvial-fluvial sedimentary rocks with a large diversity of detrital zircon
ages. The sequences are thought to have formed in late pull-apart
basins and well-developed examples include 2.71-2.70 Ga sequences
along the north and south margins of the North Caribou terrane (in the
Oxford-Stull Lake terrane and the Uchi subprovince); and 2.69-2.68
Ga sequences in the Abitibi subprovince where the type locality of this
“Timiskaming” sequence occurs.
Plate tectonics in the Archean
It is clear from the Superior Province that 3.0-2.7 Ga greenstone
belts formed in both oceanic and continental environments.
Archean “oceanic” sequences may not represent true oceanic
crust generated at mid-ocean ridges but an abundance of other
oceanic environments appears to be represented by the diverse
rock record (e.g. primitive island arc, back-arc and oceanic island
sequences). Continental sequences also appear to represent
both divergent and convergent plate settings, related to rifting or
hot-spot magmatism and subduction-zone magmatism. The
diversity of greenstone belt sequences requires a diversity
of tectono-magmatic processes to generate them, and is
most consistent with the operation of plate tectonics (or
something resembling it) in the Archean.
It has been suggested from the secular evolution of Archean
granitoid rocks (e.g. Smithies et al., 2003; Martin and Moyen, 2002),
that plate tectonics may have evolved through the Archean to
become more similar to present day plate tectonics. They suggest
that prior to 3.1 Ga the angle of the down-going oceanic plate in
subduction zones may have been shallow or even flat and hence
excluded development of a mantle wedge. Tonalites were generated
by melting of the subducted oceanic plate, but the magmas did not
interact with the mantle wedge prior to their emplacement in the
crust. After 3.1 Ga a systematic increase of Cr and Ni in tonalite
suggests increasing interaction of slab melts with a mantle wedge,
indicating a gradually thicker mantle wedge and hence steeper
angles of subduction. Mantle wedge processes in subduction zones
therefore became increasingly important after 3.1 Ga which is
consistent with the increasing presence of mantle wedge-derived
rocks (such as calc-alkaline basalts, sanukitoids and Nbenriched
basalts) in the mid to late Archean record.
2. Komatiite
Komatiites are ultramafic mantle-derived volcanic rocks. They
have low SiO2, low K2O, low Al2O3, and high to extremely high
Komatiites were named for their type locality along the Komati
River in South Africa.
True komatiites are very rare and essentially restricted to rocks
of Archaean age and most are greater than two billion years old,
restricted in distribution to the Archaean shield areas.
Komatiites occur with other ultramafic and high-magnesian
mafic volcanic rocks in Archaean greenstone belts.
Komatiites are restricted to the Archaean, with few Proterozoic
and few Mesozoic or Phanerozoic komatiites known (although
high-magnesian lamprophyres are known from the Mesozoic).
This restriction in age is thought to be due to secular cooling of the
mantle, which may have been up to 500 °C hotter during the early
to middle Archaean (4.5 to 2.6 Ga). The early Earth had much
higher heat production, because of the greater abundance of
radioactive elements, as elements with a relatively short half-life,
such as the uranium isotope with mass 235, have appreciably
diminished in abundance by radioactive decay.
The youngest komatiites are from the island of Gorgona on the
Caribbean oceanic plateau.
Komatiite core
Thin section through the
coarse-bladed olivine spinifex
zone in the slowly cooled
interior of a komatiite unit.
Samples are ~0.5 m from upper
chill margin. The long, dark
stripes in the thin section outline
the shapes of olivine (now
altered to serpentine and
magnetite) blades that reach 15
cm in length in this portion of
the unit. The lighter-colored
triangles and quadrilaterals
contain tremolite, chlorite and
magnetite and represent
regions where spinel and highCa pyroxene crystallized after
the early olivine blades.
Diamonds in volcaniclastic komatiite from French Guiana
Capdevila et al. (1999) Nature
The world's main sources of non-alluvial diamonds are found in
ultrapotassic kimberlite1 or lamproite2 diatremes (pipes filled with
explosive volcanic debris), most of which have Phanerozoic ages and
are located in stable Precambrian cratons. Diamond exploration has
therefore tended to focus on such deposits. Microdiamonds are known
to occur in metamorphic rocks such as gneiss3 and eclogite4 that have
equilibrated deep in the mantle and were then tectonically transported
to the surface, but such deposits are thought to have little commercial
potential. Here we report a new type of diamond occurrence from the
Dachine region in French Guiana for which the host rock is
volcaniclastic komatiite—an unusual type of volcanic rock whose
composition and origin are quite unlike those of kimberlite and
lamproite. These komatiites form part of a Proterozoic island-arc
sequence, a tectonic setting distinct from that of all other currently
exploited diamond deposits. The discovery of diamonds in
volcaniclastic komatiite has implications not only for diamond
exploration, but also provides strong evidence that these komatiite
magmas originated at depths of 250 km or greater within the Earth.
3. Banded iron formation (BIF)
Banded iron formation is iron rich chert (cryptocrystalline silica, SiO2).
The banded colors, usually on a cm scale are due to differing amounts
and oxidation states of Fe-containing minerals: hematite, magnetite,
grunerite, siderite and sometimes pyrite. BIF is not forming today and
although it can be found in the Archean, most deposits of BIF were
formed around 2 billion years ago.
Banded iron formations are a distinctive type of rock often found in old
sedimentary rocks. The structures consist of repeated thin layers of iron
oxides, either magnetite or hematite, alternating with bands of iron-poor
shale and chert. Some of the oldest known rock formations, formed
around three thousand million years before present (3 Ga), include
banded iron layers, and the banded layers are a common feature in
sediments for much of the Earth's early history.
Banded Iron Formations are composed of alternating layers of iron-rich
material (commonly magnetite) and silica (chert). Each layer is relatively
thin, varying in thickness from a millimeter or so up to several centimeters.
BIFs are the diagenetic product of a chemical precipitate in a very ironrich-system: SiO2-FeO-Fe2O3-CaO-MgO-CO2-H2O. They occur in the
geologic record from 3.8 Ga (Isua, West Greenland) to about 1.8 Ga with
a maximal abundance at about 2.5 Ga, and a reoccurrence in
Neoproterozoic time (from about 0.8 and 0.6 Ga).
Many of the 3.8 to 1.8 Ga BIFs have very similar average chemistries and
the late diagenetic assemblages consist mainly of chert, magnetite,
hematite, stilpnomelane and carbonates (siderite, dolomite to ankerite,
calcite). Regional metamorphism results in assemblages rich in various
amphiboles, and at higher grades, various pyroxenes. Several major BIFs
in Brazil with an age of about 2.4 Ga are much richer in Fe3+ (almost all
hematite-rich) than normal, possibly as a result of deep weathering and
secondary oxidation.
The Neoproterozoic BIFs are chemically distinctly different from most
others because about 95% of their total iron is Fe3+.
3.8 Ga Isua BIF
SW Greenland
Archean Folded BIF: Ord Range, W Australia
15 cm in
vertical scale
Light: chert
jasper bands
Dark: magnetiterich bands
Stromatolite, Ord Range, W Australia
Stromatolite: An organosedimentary structure produced by sediment trapping,
binding, and/or precipitation as a result of growth and metabolic activity of microorganism principally cyanophytes (blue/green algae).
All BIFs between 3.8 and 1.8 Ga show REE patterns with pronounced
positive Eu anomalies, negative Ce anomalies and depletion in the light
REE. These patterns are the result of chemical precipitation from
solutions that represent mixtures of seawater and hydrothermal input
(of Fe and Si) from spreading centers in oceanic crust.
The Neoproterozoic BIFs (e.g., Rapitan iron-formation, Yukon, Canada)
display a lack of the Eu anomaly, and their overall REE pattern is very
similar to that of modern ocean water at 100 m. This suggests that the
hydrothermal input was highly diluted by ocean water at this late
Precambrian time. The Neoproterozoic iron-formations commonly show
a close association with glaciogenic deposits. The Rapitan BIF is
interpreted as having been deposited during a major transgressive
event with a rapid sea-level rise during an interglacial period, after
earlier buildup of ferrous iron in solution in deeper water during a glacial
Earlier banded iron formations
The conventional concept is that the banded iron layers were formed in
water as the result of oxygen released by photosynthetic cyanobacteria,
combining with dissolved iron in Earth's oceans to form insoluble iron
oxides, which precipitated out, forming a thin layer on the substrate, which
may have been anoxic mud (forming shale and chert). Each band is similar
to a varve. The banding is assumed to result from cyclic variations in
available oxygen. It is unclear whether these banded formations were
seasonal or followed some other cycle. It is assumed that initially the Earth
started out with vast amounts of iron dissolved in the world's acidic seas.
Eventually, as photosynthetic organisms generated oxygen, the available
iron in the Earth's oceans was precipitated out as iron oxides.
It is theorized that the Earth's primitive atmosphere had little or no free
oxygen. In addition, Proterozoic rocks exposed at the surface had a
high level of iron, which was released at the surface upon weathering.
Since there wasn't any oxygen to combine with it at the surface, the iron
entered the ocean as iron ions. At the same time, primitive
photosynthetic blue/green algae was beginning to proliferate in the near
surface waters. As the algae would produce O2 as a waste product of
photosynthesis, the free oxygen would combine with the iron ions to
form magnetite (Fe3O4). This cleansed the algae's environment. As the
biomass expanded beyond the capacity for the available iron to
neutralize the waste O2 the oxygen content of the sea water rose to
toxic levels. This eventually resulted in large-scale extinction of the
algae population, and led to the accumulation of an iron poor layer of
silica on the sea floor. As time passed and algae populations reestablished themselves, a new iron-rich layer began to accumulate.
Unfortunately, the algae would again proliferate beyond the capacity of
the iron ions to clean up their waste products, and the cycle would
repeat. This went on for approximately 800,000,000 years!
Later banded iron formations
Until fairly recently, it was assumed that the rare later banded iron deposits
represent unusual conditions where oxygen was depleted locally and ironrich waters could form then come into contact with oxygenated water. An
alternate explanation of these later rare deposits is undergoing much
research as part of the Snowball Earth hypothesis — wherein it is believed
that an early equatorial supercontinent (Rodinia) was totally covered in an
ice age (implying the whole planet was frozen at the surface to a depth of
several kilometers) which corresponds to evidence that the earth's free
oxygen may have been nearly or totally depleted during a severe ice age
circa 750 to 580 Ma prior to the Ediacaran wherein the earliest multicellular
lifeforms appear. Alternatively, some geochemists suggest that BIFs could
form by direct oxydation of iron by autotrophic (non-photosynthetic)
The total amount of oxygen locked up in the banded iron beds is estimated
to be perhaps twenty times the volume of oxygen present in the modern
atmosphere. Banded iron beds are an important commercial source of iron
BIF and oxygen state in the Archean
Addition of O2 to the Atmosphere
Today, the atmosphere is ~21% free oxygen. How did oxygen reach these
levels in the atmosphere? Revisit the oxygen cycle:
Oxygen Production
。Photochemical dissociation - breakup of water molecules by
Produced O2 levels approx. 1-2% current levels
At these levels O3 (Ozone) can form to shield Earth surface from
。Photosynthesis - CO2 + H2O + sunlight = organic compounds +
O2 - produced by cyanobacteria, and eventually higher plants supplied the rest of O2 to atmosphere.
Oxygen Consumers
。Chemical Weathering - through oxidation of surface materials
(early consumer)
。Animal Respiration (much later)
。Burning of Fossil Fuels (much, much later)
Throughout the Archean there was little to no free oxygen in the
atmosphere (<1% of presence levels). What little was produced by
cyanobacteria, was probably consumed by the weathering process.
Once rocks at the surface were sufficiently oxidized, more oxygen could
remain free in the atmosphere.
During the Proterozoic the amount of free O2 in the atmosphere rose
from 1 - 10 %. Most of this was released by cyanobacteria, which
increase in abundance in the fossil record 2.3 Ga. Present levels of O2
were probably not achieved until ~400 Ma.
Evidence from the Rock Record
‧Iron (Fe) is extremely reactive with oxygen. If we look at the
oxidation state of Fe in the rock record, we can infer a great deal
about atmospheric evolution.
‧Archean - Find occurrence of minerals that only form in non-oxidizing
environments in Archean sediments: Pyrite (FeS2), Uraninite (UO2).
These minerals are easily dissolved out of rocks under present
atmospheric conditions.
‧Banded Iron Formation (BIF) - Deep water deposits in which
layers of iron-rich minerals alternate with iron-poor layers, primarily
chert. Iron minerals include iron oxide, iron carbonate, iron silicate,
iron sulfide. BIF's are a major source of iron ore, they contain
magnetite (Fe3O4) which has a higher iron-to-oxygen ratio than
hematite. These are common in rocks 2.0 - 2.8 B.y. old, but do not
form today.
‧Red beds (continental siliciclastic deposits) are never found in rocks
older than 2.3 B. y., but are common during Phanerozoic time. Red
beds are red because of the highly oxidized mineral hematite
(Fe2O3), that probably forms secondarily by oxidation of other Fe
minerals that have accumulated in the sediment.
Conclusion - amount of O2 in the atmosphere has increased with
Evolution of the Hydrosphere and the Atmosphere
The atmosphere and the oceans both arose from volcanic degassing
very early in earth history. The lighter fraction made up the atmosphere
while the heavier fraction made up the oceans. Also the presence of
water served as transporter of soluble solids and gases between the
land, sea and atmosphere.
Early Degassing : Loss of Noble Gases
Helium, argon and xenon in low concentration when compared to their
cosmic abundance. It must have been lost from the earth at high rates,
early on (i.e., the first 50 million years of earth history) in order to be so
low now.
If degassing from volcanoes has always has the same composition,
the major gasses would be water vapor and CO2 while the minor
gasses would be H2S, CO, H2, N2, CH4, NH3, HF, HCl, and Ar. But
no oxygen.
The Evolution of Oxygen
Oxygen is toxic to living systems. Oxygen-mediating enzymes had
to evolve to handle the toxic properties of O2
Different segments of the solar spectrum are capable of creating O2.
1. Ultra-violet (1500 and 2100 angstroms causes the
photodissociation of H20 in the upper atmosphere. H2 drifts into
outer space, leaving O2 behind. This this mechanism accounts
for only 0.1% of the present atmospheric level i.e., PAL).
2. Visible light, used by photoautotrophs, creates far more O2 from
CO2 and H2O and photosynthetic enzymes (e.g., chlorophyll).
The carbon released by photosynthesis was held by these
organisms by the creation of organic chemicals used in growth.
Stages of Oxygen Development through Geologic Time
A. Oceanic Oxygen
1. 4.0-3.2 Billion Years Ago
a. Oxygen held in mantle and oceanic crust. Little outgassed
b. That which did was trapped in evaporite precipitation and
the oxidation of CO to CO2.
c. Fe+2 and Mn+2 accumulated in the sea from hydrothermal
vents and the leaching of volcanics and immature
sediments. These two are soluble in their +2 state. The
lack of O2 kept them reduced and soluble.
1. 3.2-2.6 Billion Years Ago Early photoautotrophic organisms release
small concentrations of O2 that oxidized Fe+2 creating insoluble
Fe+3 precipitates in Banded Iron Formations (i.e., BIF) in island arc
(i.e., Algoma type BIF) and shallow shelf environments (i.e.,
Superior type BIF) as a result of oceanic upwellings.
2. 2.6-2.0 Billion Years Ago
a. Development of extensive Superior-type BIF and manganese
deposits capturing O2 released by the photoautotrophs.
b. Stromatolites in limestones expand rapidly both providing O2 to
the ocean while capturing carbon at the same time.
c. The presence of oxygen cleared the sea of soluble iron and
manganese so that little further oxygen was needed for BIF
and manganese deposit formation.
B. Atmospheric Oxygen.
Oxygen produced by photoautotrophs was used in BIF and
manganese deposits. Once O2 production increased beyond that
needed by BIF and manganese deposit formations, could it
accumulate in the ocean and in the atmosphere.
Free oxygen began to accumulate in the atmosphere between 2.4
and 1.9 billion years ago. Evidence for this includes the following.
1. BIF deposits iron-leaching from paleosols disappear after 1.9
billion years. The presence of oxygen rendered iron and manganese
insoluble, hence they would not leach. Leaching could only occur if
iron and manganese were in the +2 state and therefore soluble and
2. Presence of detrital uraninite and pyrite 2.3 billion years ago, then
disappears. Grains of uraninite and pyrite indicate they are insoluble in the
absense of oxygen. Once oxygen is present, they dissolve and won't
occur as solid grains (i.e., detrital).
3. Development of hematite-rich paleosols by 2.2 and 2.0 billion years
indicates that there is enough oxygen in the atmosphere to oxidize iron
(i.e., rust). This indicates an increase in atmospheric oxygen of 15 fold 1.9
billion years ago.
4. Appearance of CaSO4 in evaporites after 1.9 billion years ago indicates
that there is enough atmospheric or oceanic oxygen to oxidize sulfides to
5. First appearance of redbeds (red conglomerates, sandstones and shales)
after 2.3 billion years ago
6. Accumulation of organic-rich limestones between 0.9 and 0.6 billion years
ago indicates the retention of carbon compounds from photosynthesis
which means an equal volume of oxygen had to be released into the
The atmospheric oxygen level fluctuated during the Phanerozoic.
Oxygen increased during periods when plants evolved and
expanded on the continents. Oxygen decreased during arid
periods, when sea levels drops because newly exposed coals and
organic-rich sediments would consume oxygen in their
decomposition (i.e., the opposite process of photosynthesis).
Carbon Dioxide Balance
CO2 concentration in the atmosphere was many hundred times greater
in the Archean. This was due to a.) the oxidation of CO and CH4 to form
CO2 by the ever-increasing concentrations of atmospheric and oceanic
oxygen, b.) the release of CO2 during the periods of massive volcanism
in the Archean and c.) the much lower rates of subduction due to the
lower thermal gradients in the mantle.
Higher CO2 concentrations (i.e., 80 to 600 times PAL) led to the
Greenhouse effect that kept temperatures above freezing and thereby
prevented glaciations.
Higher CO2 concentrations led to the photochemical production of
major oxidants, hydrogen peroxide (H2O2) and formaldehyde (H2CO),
that enhanced ferrous iron solution in the Archean.
As the CO2 concentration decreased through the Proterozoic,
stromatolite production decreased.
In the Phanerozoic, CO2 concentrations fluctuated for several
Uplift and mountain building due to plate tectonics induces the
outgassing of CO2, generated by metamorphism.
CO2 concentrations drops almospheric concentrations due to loss of
CO2 in the creations of deep sea and shallow water limestones
(i.e., CaCO3).
Burial of coal and carbonates by silicate sediments reduces CO2
concentration by burying it below the surface.
So uplift increases CO2 concentration, while their erosion to low relief
decreases CO2 concentration.
Atmospheric Evolution and the Development of Life
Reduced carbon and carbohydrate chemicals found in rocks 3.8 billion
years old.
First, prokarotic cell structures found around 3.5 billion years ago.
Bacterial communities (stromatolites) found in rocks 3.2 billion years
First eukarotic cells found around 1.6 billion years ago. Atmospheric
oxygen levels reached about 10% PAL. Respiration, a far more
efficient metabolic set up is now possible. Eukarotes mark the
beginning of sexual reproduction that, in turn, means greater
possibilities of genetic development.
First metazoa found about 700 to 600 million years ago (0.6-0.7 billion
years ago). Colligen produced and leads to the development of
hardparts 570 million years ago (.57 billion years).
Camrian period marks the resting removal of atmospheric CO2 from
prokarotic tissue growth to metazoan eukarotic CaCO3 shell formation.
Types of BIF
Archean Earth Banded Iron Formation (BIF) has been
suggested as a possible terrestrial analog for Early Mars
(Calvin, 1998).
Two types of BIF in the United States and Canada have been
differentiated based on their respective origins. The Algoma type
deposits in Ontario, Canada are in close proximity to ancient volcanic
centers suggesting a sub-aqueous hydrothermal origin similar to
modern day sea-floor spreading centers (Gross, 1983). The Lake
Superior type BIF deposits in the upper peninsula of Michigan are
not associated with extrusive volcanic materials and are therefore
interpreted as chemical precipitates of iron-rich waters in a shallow
sea (James, 1954). The Thermal Emission Spectrometer (TES)
discovery of crystalline, gray hematite in sedimentary basin type
deposits on Mars supports the use of Lake Superior type BIF as a
terrestrial analog.
The Sinus Meridiani and Aram Chaos hematite sites are not in close
proximity to a volcanic center, and do not exhibit any lava flow
features (Christensen, et al., 2001). The Sinus Meridiani hematite
occupies a smooth unit with abrupt boundaries suggesting that it
exists within a stratigraphic layer. The Aram Chaos hematite appears
to be within a closed basin around which outflow channels are
common suggesting an aqueous origin. In both sites, the hematite
appears to be part of layered, sedimentary rock units that suggest
aqueous environments (Christensen, et al., 2001).
The Lake Superior type BIF occurs in four principal facies: sulfide,
carbonate, silicate, and oxide (James, 1954). These facies grade
into each other in the field reflecting changes in the oxidation state
of the water and occur as thin laminae alternating with chert layers.
The mm scale laminations of these rocks will not be evident in
large- scale (3km x 6km) TES spectra. The iron-rich minerals
present in each facies are possible auxiliary minerals for the low
albedo hematite regions on Mars. These minerals are: pyrite in the
sulfide facies, siderite in the carbonate facies, minnesotaite and
stilpnomelane in the silicate facies, and magnetite and hematite in
the oxide facies. A field trip to the Lake Superior type deposits in
the Marquette and Gogebic iron districts of Michigan has provided
a thorough rock sampling of the different facies, including several
types of crystalline, gray hematite. Micaceous, specular hematite
with a schistose texture is highly metamorphosed and is probably
not seen on the surface of Mars.
Bulk, gray crystalline hematite occurs in relatively
unmetamorphosed BIF and retains its sedimentary layer nature. It
also displays a microplaty texture in some samples that is most
likely the result of low-grade burial metamorphism.
Some of the bulk, gray crystalline hematite displays magnetic
properties suggesting some mixture of magnetite and hematite.
The spectra of these bulk samples may be better analogs for
Mars than pure mineral phases. The spectra of these samples will
be presented and compared to what TES has observed.
Algoma-type banded-iron formation deposit, or Algoma-type BIF
Algoma type was formed over a much wider time range than the Lake
Superior type (from 3.8 billion to a few hundred million years ago).
Algoma-type BIFs are also finely layered intercalations of silica and an
iron mineral, generally hematite or magnetite, but the individual layers
lack the lateral continuity of Lake Superior-type BIFs.
Algoma type BIF
SYNONYMS: Taconite, itabirite, banded iron-formation.
EXAMPLES: Falcon, Lady A, McLeod, Sherman, Adams, Griffith
(Canada), Woodstock, Austin Brook (New Brunswick, Canada),
Kudremuk (India), Cerro Bolivar (Venezuela), Carajas (Brazil), part of
Krivoy Rog (Russia).
CAPSULE DESCRIPTION: Iron ore deposits in Algoma-type ironformations consist mainly of oxide and carbonate lithofacies that
contain 20 to 40 % Fe as alternating layers and beds of micro- to
macro-banded chert or quartz, magnetite, hematite, pyrite, pyrrhotite,
iron carbonates, iron silicates and manganese oxide and carbonate
minerals. The deposits are interbedded with volcanic rocks, greywacke,
turbidite and pelitic sediments; the sequences are commonly
TECTONIC SETTINGS: Algoma-type iron-formations are deposited
in volcanic arcs and at spreading ridges.
AGE OF MINERALIZATION: They range in age from 3.2 Ga to
modern protolithic facies on the seafloor and are most widely
distributed and achieve the greatest thickness in Archean terranes
(2.9 to 2.5 Ga).
formed both near and distal from extrusive centres along volcanic
belts, deep fault systems and rift zones and may be present at any
stage in a volcanic succession. The proportions of volcanic and
clastic sedimentary rocks vary and are rarely mutually exclusive.
HOST/ASSOCIATED ROCKS: Rocks associated with Algoma-type
iron-formations vary greatly in composition, even within local basins,
and range from felsic to mafic and ultramafic volcanic rocks, and from
greywacke, black shale, argillite, and chert interlayered with pyroclastic
and other volcaniclastic beds or their metamorphic equivalents.
Algoma-type iron-formations and associated stratafer sediments
commonly show a prolific development of different facies types within a
single stratigraphic sequence. Oxide lithofacies are usually the thickest
and most widely distributed units of iron-formation in a region and serve
as excellent metallogenetic markers.
DEPOSIT FORM: Iron ore deposits are sedimentary sequences
commonly from 30 to 100 m thick, and several kilometres in strike
length. In most economic deposits, isoclinal folding or thrust faulting
have produced thickened sequences of iron-formation.
STRUCTURE/TEXTURE: Micro-banding, bedding and
penecontemporaneous deformation features of the hydroplastic
sediment, such as slump folds and faults, are common, and can be
recognized in many cases in strongly metamorphosed oxide lithofacies.
Ore mineral distribution closely reflects primary sedimentary facies. The
quality of oxide facies crude ore is greatly enhanced by metamorphism
which leads to the development of coarse granular textures and discrete
grain enlargement.
ORE MINERALOGY: Oxide lithofacies are composed of magnetite and
hematite. Some deposits consist of siderite interbedded with pyrite and
GANGUE MINERALOGY (Principal and subordinate): Quartz,
siderite or ferruginous ankerite and dolomite, manganoan siderite
and silicate minerals. Silicate lithofacies are characterized by iron
silicate minerals including grunerite, minnesotaite, hypersthene,
reibeckite and stilpnomelane, associated with chlorite, sericite,
amphibole, and garnet.
WEATHERING: Minor oxidation of metal oxide minerals and
leaching of silica, silicate and carbonate gangue. Algoma-type ironformations are protore for high-grade, direct shipping types of
residual-enriched iron ore deposits.
GENETIC MODEL: Algoma-type iron deposits were formed by the
deposition of iron and silica in colloidal size particles by chemical and
biogenic precipitation processes. Their main constituents evidently
came from hydrothermal-effusive sources and were deposited in
euxinic to oxidizing basin environments, in association with clastic
and pelagic sediment, tuff, volcanic rocks and a variety of clay
minerals. The variety of metal constituents consistently present as
minor or trace elements evidently were derived from the
hydrothermal plumes and basin water and adsorbed by amorphous
iron and manganese oxides and smectite clay components in the
protolithic sediment.
Their development and distribution along volcanic belts and deepseated faults and rift systems was controlled mainly by tectonic
rather than by biogenic or atmospheric factors. Sulphide facies were
deposited close to the higher temperature effusive centres; iron oxide
and silicate facies were intermediate, and manganese-iron facies
were deposited from cooler hydrothermal vents and in areas distal
from active hydrothermal discharge. Overlapping and lateral
transitions of one kind of lithofacies to another appear to be common
and are to be expected. ORE CONTROLS: The primary control is
favourable iron-rich stratigraphic horizons with little clastic
sedimentation, often near volcanic centres. Some Algoma-type ironformations contain ore deposits due to metamorphic enhancement of
grain size or structural thickening of the mineralized horizon.
ASSOCIATED DEPOSITS: Algoma-type iron-formations can be
protore for residual-enriched iron ore deposits. Transitions from
Lake Superior to Algoma-type iron-formations occur in areas
where sediments extend from continental shelf to deep-water
environments along craton margins as reported in the Krivoy Rog
iron ranges. Oxide lithofacies of iron- formation grade laterally
and vertically into manganese-rich lithofacies, and iron sulphide,
polymetallic volcanic-hosted and sedex massive sulphide.
GRADE AND TONNAGE: Ore bodies range in size from about 1000
to less than 100 Mt with grades ganging from 15 to 45% Fe,
averaging 25% Fe. Precambrian deposits usually contain less than
2% Mn, but many Paleozoic iron-formations, such as those near
Woodstock, New Brunswick, contain 10 to 40 % Mn and have Fe/Mn
ratios of 40:1 to 1:50. The largest B.C. deposit, the Falcon, contains
inferred reserves of 5.28 Mt grading 37.8% Fe.
ECONOMIC LIMITATIONS: Usually large-tonnage open pit
operations. Granular, medium to coarse- grained textures with well
defined, sharp grain boundaries are desirable for the concentration
and beneficiation of the crude ore. Strongly metamorphosed ironformation and magnetite lithofacies are usually preferred. Oxide
facies iron-formation normally has a low content of minor elements,
especially Na, K, S and As, which have deleterious effects in the
processing of the ore and quality of steel produced from it.
IMPORTANCE: In Canada, Algoma-type iron-formations are the
second most important source of iron ore after the taconite and
enriched deposits in Lake Superior-type iron-formations.
Algoma-type iron-formations are widely distributed and may
provide a convenient local source of iron ore.
Lake Superior-type BIF deposit
Sedimentary rocks deposited in the shallow waters of continental
shelves or in ancient sedimentary basins. These deposits are
typified by the vast BIFs around Lake Superior and are called Lake
Superior-type deposits. Their individual sediment layers can be as
thin as 0.5 millimetre (0.02 inch) or as thick as 2.5 centimetres (1
inch), but the alternation of a siliceous band.
Spectral Properties of Lake Superior Banded Iron Formation:
Application to Martian Hematite Deposits
Several locations have been identified on Mars that expose bulk,
coarsely crystalline gray hematite. These deposits have been interpreted
as being sedimentary and formed in aqueous environments.
Lake Superior Type (LST) banded iron formation (BIF) was investigated
as a spectral and possible process analog to these deposits. In northern
Michigan, LST BIF formed in a sedimentary, continental shelf or shallow
basin environment under stable tectonic conditions, and the oxide facies
contains gray, crystalline hematite.
These deposits are Proterozoic in age and contain microfossils
associated with the early diversification of life on Earth. Samples
of the hematite-bearing oxide facies, as well as the carbonate
facies, were collected and analyzed for their spectral and
geochemical characteristics. Sample spectra were measured in
the visible, near-infrared, and thermal infrared for comparison with
remote and in situ spectra obtained at Mars. Thin section analysis,
as well as X-ray diffraction and scaning electron microscopy
measurements, were performed to determine detailed
geochemistry. There is no evidence for BIF at Opportunity's
Meridiani landing site, and the results of this work will provide
useful data for determining whether BIFs exist elsewhere on Mars
and are, thus, relevant to current and future Mars exploration
Kiruna type BIF
The Fennoscandian Shield, one of the major base metal provinces in
Europe, is composed of an Archean nucleus, largely unmineralized,
in the northeastern part of the Shield. This nucleus is bordered to the
southwest by Paleoproterozoic rocks. At c. 2.5-2.3 Ga sedimentary
and volcanic rocks were deposited on the Archean basement during
an extensional event. Further rifting of the continent at c. 2.1 Ga gave
rise to tholeiitic and komatiitic lavas and dikes. At the end of this
extensional event MORB-type pillow lavas were erupted. At c. 1.9 the
tectonic regime shifted to compressive and subduction related
volcanic and sedimentary rocks were deposited in a terrestrial to
shallow water environment. Southwest of these intracratonic
complexes, 1.95-1.87 Ga old volcanic arcs were accreted towards
the craton during the Svecokarelian orogeny. This orogeny involved
voluminous early calc-alkaline magmatism and ended with
migmatization, S-type magmatism and large batholithic intrusions of
A- to I type granitoids.
Mineralization related to these Proterozoic early extensional and
later comressional tectonic regimes include VMS (including
Outokumpu Cu-Zn-Co±Ni type) to epithermal VMS, sedimenthosted Zn-Pb, porphyry style Cu, gold lode style deposits, BIF´s,
mafic and ultramafic Ni±Cu±PGE deposits as well as Kiruna type
apatite-Fe deposits, epigenetic Cu-Au deposits and syngenetic
Cu deposits. The latter three types of economic deposits are
included in the diverse group of Fe-oxide-Cu-Au style
mineralizations: The Kiruna type apatite iron ores are hosted by
1.88 Ga felsic alkaline porphyries emplaced during compressional
tectonics. The epigenetic Cu-Au deposits is a diverse group of
mineralizations including vein style structurally controlled Cu-Au,
probably both 1.87 Ga and 1.77 Ga in age, and intrusive hosted,
porphyry style Cu-Au±Fe, related to both calc-alkaline and alkaline
magmatism in a compressional regime between 1.9-1.8 Ga. The
syngenetic Cu±Zn deposits restricted to the c. 2.1 Ga greenstones,
formed during extensional tectonics in intracratonic rift basins.
Kiruna SWEDEN (Province: Norrbotten)
Long (E): 20.112, Lat (N): 67.496
Type: Apatite Fe-ore
Morphology: Concordant sheet
Age of mineralization: c. 1.88 Ga
Ore minerals: Magnetite
Alteration: albitization, actinolite, biotite-chlorite
Age of host rocks: 1880±3 Ma (U-Pb), cutting dyke: 1876±9 Ma (U-Pb tit)
Nature of host rocks: trachyandesite lava, felsic volcanics, intermediatemafic volcanics
Cumulative past production and reserves: 2 000 Mt @ 60-68 % Fe
Ore age Palaeoproterozoic 1880 Ma
Ore mineralogy Hydrothermal alteration
Magnetite: Albitization and Biotitization
Deposit type
Magnetite-apatite deposits (tabular and pipe-like bodies, dykes) (Kiruna): Fe, P
Ore shape
Concordant to subconcordant mass, lens or pod of massive to submassive ore
Host rock age
Palaeoproterozoic 1880 ± 3 Ma U/Pb
Host rock mineralogy
Actinolite, Fe-Mg mica, Apatite
Host rock lithology
Acidic volcanic rock, Basic volcanic rock
Host rock formation names
Kiruna Porphyries
Fe Iron (metal)
Past Production average Reserve: 1 200 000 000 t
Average grade: 60%