Archean Greenstone Belt

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Archean Greenstone Belt
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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.
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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 granitoids. 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 (69 km)
flat-base and overlies a uniformly thick crust.
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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 present-day 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.
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The alternate view is that tectonic processes
comparable to present-day 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.
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Isua greenstone belt (Greenland) The Isua
Greenstone Belt is an Archean greenstone belt in
SW 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. 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. New geological
mapping studies are tracing the transitional
gradations between the protoliths and their
diverse deformed and metasomatized structures.
These new mappings show that most of the Isua
Greenstone Belt consists of fault bounded rock
assemblies (1) derived from basalt and high-Mg
basaltic pillow lava, (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.
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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.
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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 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 gt150 km, at temperatures of
1600-2000 oC (Nisbet et al., 1977).
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Greenstone Belts in the Superior Province and the
Evolution of Archean Tectonic Processes
(Development of Abitibi Greenstone Belt) 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 areas.
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Greenstone belt types
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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.
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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 Oxford-Stull 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.
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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
syn-volcanic 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 Oxford-Stull 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
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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 subprovince.
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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.
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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.
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2. Komatiite Komatiites are ultramafic
mantle-derived volcanic rocks. They have low
SiO2, low K2O, low Al2O3, and high to extremely
high MgO. 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 Archean 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.
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Komatiites are restricted to the Archean, 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 Archean (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.
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Komatiite core
21
Komatiite
22
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
high-Ca pyroxene crystallized after the early
olivine blades.
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Diamonds in volcaniclastic komatiite from French
Guiana The world's main sources of non-alluvial
diamonds are found in ultrapotassic kimberlite or
lamproite 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
gneiss and eclogite 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. A
new type of diamond occurrence from the Dachine
region in French Guiana has been found. The host
rock is volcaniclastic komatiite 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.
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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.
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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.
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BANDED IRON-FORMATIONS (BIFS) WORLDWIDE. BIFs
are the diagenetic product of a chemical
precipitate in a very iron-rich-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.
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3.8 Ga Isua BIF SW Greenland
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Archean Folded BIF Ord Range, W Australia
Dimension 15 cm in vertical scale Light
chert jasper bands Dark magnetite- rich bands
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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 micro- organism principally
cyanophytes (blue/green algae).
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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 period.
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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.
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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 re-established
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!
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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 iron-rich
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
superconitent (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) microbes. 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 ore.
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BIF and oxygen state in the Archean
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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 ultraviolet Produced O2 levels
approx. 1-2 current levels At these levels O3
(Ozone) can form to shield Earth surface from UV
?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)
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Cyanobacteria Throughout the Archean there was
little to no free oxygen in the atmosphere (lt1
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.
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?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
time.
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Supplement-1
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.
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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
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• 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. Fe2 and Mn2 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.

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• 3.2-2.6 Billion Years Ago Early photoautotrophic
  organisms release small concentrations of O2 that
  oxidized Fe2 creating insoluble Fe3
  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.

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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 mobile.
43
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 sulfates.
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 atmosphere.
44
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).
45
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.
46
In the Phanerozoic, CO2 concentrations fluctuated
for several reasons. 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.
47
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
old. 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.
48
Supplements-2 Types of BIF Archean Earth
Banded Iron Formation (BIF) has been suggested as
a possible terrestrial analog for Early Mars
(Calvin, 1998).
49
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).
50
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.
51
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52
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.
53
Algoma-type banded-iron formation
deposit, or Algoma-type BIF deposit characteristi
cs 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.
54
Algoma type BIF
55
ALGOMA-TYPE IRON FORMATION SYNONYMS Taconite,
itabirite, banded iron-formation. COMMODITIES
(BYPRODUCTS) Fe (Mn). 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).
GEOLOGICAL CHARACTERISTICS CAPSULE DESCRIPTION
Iron ore deposits in Algoma-type iron-formations
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 metamorphosed.
56
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). DEPOSITIONAL
ENVIRONMENT / GEOLOGICAL SETTING They 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.
57
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.
58
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 pyrrhotite.
59
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 iron-formations are protore
for high-grade, direct shipping types of
residual-enriched iron ore deposits.
60
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.
61
Their development and distribution along volcanic
belts and deep-seated 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
iron-formations contain ore deposits due to
metamorphic enhancement of grain size or
structural thickening of the mineralized horizon.
62
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.
63
ECONOMIC FACTORS 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
401 to 150. 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 iron-formation 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.
64
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.
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66
 Lake Superior-type BIF deposit characteristics
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.
67
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.
68
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 missions.
69
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.
70
Mineralization related to these Proterozoic early
extensional and later comressional tectonic
regimes include VMS (including Outokumpu
Cu-Zn-CoNi type) to epithermal VMS,
sediment-hosted Zn-Pb, porphyry style Cu, gold
lode style deposits, BIFs, mafic and ultramafic
NiCuPGE 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-AuFe, related to both
calc-alkaline and alkaline magmatism in a
compressional regime between 1.9-1.8 Ga. The
syngenetic CuZn deposits restricted to the c.
2.1 Ga greenstones, formed during extensional
tectonics in intracratonic rift basins.
71
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 18803 Ma
(U-Pb), cutting dyke 18769 Ma (U-Pb
tit) Nature of host rocks trachyandesite lava,
felsic volcanics, intermediate-mafic
volcanics Cumulative past production and
reserves 2 000 Mt _at_ 60-68 Fe
72
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73
Mineralization Ore age Palaeoproterozoic 1880
Ma Ore mineralogy Hydrothermal alteration
Magnetite Albitization and Biotitization
74
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 Economy Fe Iron (metal) Past
Production average Reserve 1 200 000 000
t Average grade 60