Proterozoic Stromatolites
Nastopoka Sound
Carbonates et al
Francis, 2014
Bio-Chemical Sediments
Cations that are leached during weathering and transported by solution in water
include:
K+, Na+, Ca++, Mg++
Other cations are also present in solution, but at very low concentrations:
Si4+, Mn2+, Fe2+
River Input versus Sea Water Sink
Species
Av. Rivers (ppm)
HCO3SO4-Cl-
50
10
8
140
2,715
18,980
Ca++
Mg++
Na+
K+
15
4.1
6.3
2.3
413
1,290
10,800
399
0.04
6.5
0.00006
2.8
Fe
Si
Ocean (ppm)
River Input versus Sea Water Sink
Species
Av. Rivers (ppm)
Ocean (ppm)
HCO3SO4-Cl-
50
10
8
140
2,715
18,980
Ca++
Mg++
Na+
K+
15
4.1
6.3
2.3
413
1,290
10,800
399
0.04
6.5
0.00006
2.8
Fe
Si
Solubility of Minerals in Sea Water at 25oC
Mineral
Formula
Solubility (ppm)
Quartz
Opal
SiO2
SiO2.nH2O
Calcite
Aragonite
Dolomite
CaCO3
CaCO3
CaMg(CO3)2
14
15
32
Gypsum
Anhydrite
CaSO4.2H2O
CaSO4
2,400
2,500
1,400
Halite
Sylvite
Carnalite
NaCl
KCl
KMgCl3.6H2O
360,000
350,000
650,000
30,000
10
Conc (ppm)
1
120
120
Calcite Group
32/m
Rhombohedral
Although calcite (s.g. 2.75) is the stable form of CaCO3 at room
pressure and temperature and is supersaturated in sea water, it has
never been successfully precipitated from seawater in experiment.
Some organism (eg. echinoderms) secrete metastable high-Mg calcite
with 4 to 30% dissolved MgCO3, but this is converted to calcite during
diagensis. In past epochs, high-Mg calcite has been the dominant
biochemical precipitate in the oceans, rather than aragonite, which is
the dominate carbonate precipitated in the oceans today.
Dolomite Group
3
Carbonates:
Three groups of carbonate minerals:
Rhombohedral
Although dolomite (s.g. = 2.85) is a common rock in the
stratigraphic record, dolomitic sediments are rare in the
modern environment, and, despite the fact that seawater is
supersaturated in dolomite, it too has never been directly
precipitated from seawater in experiment.
Aragonite Group
2/m2/m2/m
Orthorhombic
Aragonite (s.g. 3.94) is the high pressure polymorph of CaCO3 and is slightly more soluble in sea water than
calcite. Aragonite, however, is the dominant form of CaCO3 that precipitates directly from seawater today, and is
the dominant skeletal material of most carbonate-secreting organisms. Recrystallization of aragonitic sediments
in contact with fresh water, converts aragonite to calcite, in much the same way that amorphous silica is
converted to quartz.
Controls on the Solubility of Calcite:
CaCO3 + CO2 + H2O = Ca2+ + 2HCO3K=
[Ca2+] [HCO3-]2
[CaCO3] [CO2] [ H2O]
=
CO2 + H2O =
[Ca2+] [HCO3-]2
[CO2] [ H2O]
in presence of calcite
H2CO3 (carbonic acid)
H2CO3 + CaCO3 = Ca2+ + 2HCO3- (bicarbonate ion)
The concentration of Ca2+ in water in equilibrium with calcite and air (PCO2 = 3 × 10-6
atm) is approximately 20 ppm, which corresponds to a calcite solubility of approximately
50 ppm.
the Solubility of Calcite:
CaCO3 + CO2 + H2O = Ca2+ + 2HCO3K=
[Ca2+] [HCO3-]2
[CaCO3] [CO2] [ H2O]
Inspection of the above equilibria indicates that the
solubility of calcite is directly related to the amount of
dissolved CO2, which is an inverse function of temperature.
Thus the higher the temperature, the lower the solubility of
calcite. On the other hand, increased pressure puts more
CO2 into a solution and increases the solubility of CaCO3.
The combined effects of increasing pressure and decreasing
temperature in the oceans are such that stable carbonate
sediments are restricted to water depths of less than 3500
metres. Below this “carbonate compensation depth” there
is a dramatic increase in the solubility of CaCO3, such that
it becomes undersaturated in seawater, and carbonate
sediments are rare because they dissolve. Note that the
greater solubility of aragonite compared to calcite means
that aragonite becomes undersaturated in seawater at much
shallower depths. This results in the conversion of
aragonite to calcite during diagenesis.
The above relationships are complicated by the fact that the solubility of a given mineral
is sensitive to the other species in solution, which compete and/or interfere with the
species of interest. Effectively, the presence of other ions reduces the activity coefficient
of the ions of interest in the equation for the equilibrium constant, thus requiring higher
concentrations to achieve saturation. The magnitude of this effect is proportional to the
ionic strength of the solution:
I = ½ ×  Mi × Zi2
where Mi are the molar concentrations of the species and Zi are their
ionic charges. Note: neutral species are assumed to have no effect.
Furthermore, the kinetics of nucleation play as large a role as solubility and concentration in
determining the point at which a mineral will precipitate from solution. For example, although
calcite will precipitate from fresh water, it will not from seawater, despite the fact that seawater is
supersaturated in CaCO3. This may reflect the fact that the high concentration of Mg(OH)2 in sea
water poisons the nucleation and growth of calcite.
In general, many dissolution reactions seem to be one way, that is it is much easier to dissolve a
mineral than to precipitate it. And then there is life, which plays havoc with everything.
Ostwald’s Step Rule
In general, the least stable polymorph of a chemical compound tends to
be the first to precipitate from a solution. Thus aragonite precipitates
before calcite, opal precipitates before quartz.
In modern oceans, the relatively unstable polymorph aragonite is precipitated by
organisms rather than calcite, despite the fact that it has a higher solubility. Some
have argued than in fact CaCO3 precipitates first as a amorphous gel that is converted
rapidly to aragonite, and then later to calcite during diagenesis.
Allochthonous Limestones
Most limestones were originally clastic rocks, and in many respects can be
classified according to the same grain size criteria as the siliciclastic rocks:
mudstones, shale
lutite, micrite
sandstones
grainstone
conglomerates & breccia
rudite
These types of limestone are referred to as allochthonous, in that their constituents
have been transported from elsewhere.
The major differences between
allochthonous limestones and siliciclastic rocks are in the nature of the clastic
fragments, the degree of sorting, and the matrix or cement that holds these rocks
together.
Allochthonous versus Autochthonous
Carbonates
Reef complexes are essentially living buckets, in which the walls of the
bucket are composed of boundstone of living corals and related species,
the shallow inner lagoon is filled with fine grainstones and mudstone,
while the steep outer flanks are composed of coarser grainstones and
rudestone, formed from talus off the reef front.
Boundstones
Stromatolites and Algal Mats
Mats of blue green algae that
grow in the intertidal to
supratidal zone. These organic
mats trap carbonate mud when
flooded, eventually forming a
finely laminated micrite, in
which the original organic seams
have been replaced by porous
septa or sparite cement. One of
the few clastic sedimentary
rocks in which the size of
porosity is larger than the grains.
Dominant carbonate rock in the
Precambrian,
before
the
development
of
reefs.
Stromatolites represent localized
buildups or cabbage-shaped
heads formed by a similar
mechanism.
Distribution of Carbonate
Sediments
Depth:
Active growing reefs are restricted to the
photic zone in the oceans, which extends
only to depths of 20 to 150 m, depending of
the turbidity of the water. Clastic carbonates
sediments can be transported to much deeper
environments, but become unstable at depths
greater than 3500 meters because of the high
solubility of CaCO3 in cold water at high
pressure. All this means that carbonate
sedimentation is largely restricted to
continental shelfs environments. Reefs will
not grow under turbid conditions where there
is a high terrigenous clastic input, such as at
the mouths of large rivers like the Amazon
and the Mississippi.
Distribution of Carbonate Sediments
Latitude:
Reefs grow only within  30o of
the equator, restricted to waters
with
minimum
annual
temperatures
of
18oC.
Carbonate clastic sediments,
however, can reach  60o on
either side of the equator.
Apparently,
cool
water
carbonates do form on some
high latitude continental shelves.
Longitude:
Distribution of Carbonate Sediments
Carbonate reefs are better developed on the western sides of the Earth's oceans because of the
direction of the prevalent wind and ocean currents within 30o of the equator. Deep ocean water
tends to rise on the eastern sides of the oceans and sink on the western sides. Because cold
water dissolves more CaCO3, the deep cold waters are nutrient-rich, becoming supersaturated as
they warm rising to the ocean surface, promoting high planktonic productivity, which in turn
results in turbid water conditions. Corals, on the other hand, thrive best in clear lowproductivity ocean water and as a result grow best on the western sides of ocean basins;
reaching latitudes of  30o on the western sides of oceans, but only about  5-10o latitude on the
east.
Carbonate Reef
Distribution in Time
Precambrian limestones are dominated by
stromatolite mats and mounds that develop
by the repeated coating of tidal-zone algal
mats by micrite.
The first true coral reefs developed in the
mid-Ordovician and were dominated by HiMg calcite secreting organisms that reached
their maximum development at the end of
the Devonian. Carbonate development at
this point retreated to build ups and mounds
formed of aragonite secreting organisms, and
then was arrested completely by the mass
extinction the end of the Permian.
Aragonite reefs became widespread again in
the Triassic and Jurassic, but then decreased
in the early Cretaceous, with carbonate
growth returning to mounds of calcite
secreting organisms. Extensive aragonite
reefs only returned in the Tertiary and have
continued until today, although there now
appears to be a reef die-off occurring,
possibly as a result of global warming.
Dolomite
Although dolomite is a common rock in the stratigraphic record, particularly in the Paleozoic where it is the dominate
carbonate rock, the mineral dolomite has never been successfully precipitated from sea water solutions at temperatures
less than 100oC, possibly because of the strong hydrolysis of Mg (Mg(OH)2 ) in water solutions. Furthermore, modern
day dolomitic sediments are relatively rare and largely restricted to sabkha environments (salt-encrusted supra-tidal
flats) where they appear to form by the alteration of initial aragonite precipitates. Most dolomites are thought to have
formed either early by the penecontemporaneous alteration of unconsolidated calcareous sediments or somewhat later
during the diagenesis of carbonate rocks.
Four models for penecontemporaneous dolomite:
Sabkha or hypersaline model
Rapid evaporation in hot dry climates,
such as that of the Persian Gulf, increases
the salinity of sea water in supra-tidal
flats to the point at which gypsum
precipitates, which depletes the remaining
brines in Ca, raising the Mg/Ca ratio from
5:1 to 10:1. These Mg-rich brines are
denser than seawater and sink into the
underlying calcareous sediments, reacting
with them to convert them to dolomite.
An analogous process on a larger scale
may be responsible for the stratigraphic
association of dolomite with evaporite
deposits
Mixing zone model
It has been proposed that dolomite might form in the mixing zone between ocean water and
fresh water. Although the mixing of two fluids can result in the supersaturation of mineral
phases, this has yet to be substantiated for the mineral dolomite, and many conditions of
mixing actually result in undersaturation. These results reflect the fact that while
concentrations of species are linear functions of mixing, the solubility products of minerals
and carbonate solubility as a function of PCO2 are non linear.
Low-sulfate model
The presence of SO42- is known experimentally to inhibit the nucleation of dolomite
and dolomite has been precipitated successfully in the absence of SO42-. Thus
processes that lower SO42- content have been proposed for the direct precipitation of
dolomite from sea water solutions in unconsolidated sediments. Such process
might include the bacterial conversion of SO42- to sulfide in anoxic environments,
as well as the precipitation of gypsum (CaSO4.2H2O) in the sabkha environment
described above.
Large through-put model
It has been proposed that a large through-put of sea water will, at high
water/sediment ratios, convert unconsolidated sediments from calcite, or high-Mg
calcite, to dolomite by ion exchange of Mg2+ for Ca2+. Such models require high
water/rock ratios. For example, ~ 800 m3 of seawater is required to convert 1m3 of
calcite to dolomite. This number drops to ~ 45 m3, if the solution is a brine which
has been evaporated to the point of halite saturation.
Diagenetic Models for the origin of Dolomite:
High Magnesian Calcite
Todays carbonate organisms are dominantly aragonitic, but it appears that in the past there have
been episodes during which high-Mg calcite was favoured, for reasons which remain somewhat
unclear, but likely reflect low Mg/Ca ratios. Hi-Mg calcite appears to be thermodynamically
unstable, so that during diagenesis it would revert to calcite or aragonite, giving Mg2+ to
solutions that then go on to “dolomitize” other carbonates along the fluid pathways.
Evaporite Association
Dolomite is commonly associated with evaporite deposits. Following the precipitation of
gypsum in evaporite sequences, the Mg/Ca ratios of the residual brines rise rapidly. These
residual brines are denser and sink into underlying carbonate, gradually converting it to dolomite.
the Solubility of Calcite:
CaCO3 + CO2 + H2O = Ca2+ + 2HCO3K=
[Ca2+] [HCO3-]2
[CaCO3] [CO2] [ H2O]
Temperature and PCO2
Aragonite
Calcite
aragonite
Silica:
The solubility of amorphous silica is controlled by a number of possible reactions. At low pH, the solubility is controlled by:
SiO2 + 2H2O
H4SiO4 (silicic acid) (1)
K =
[H4SiO4]
[SiO2][H2O]2
This reaction is insensitive to pH, however, at high pH ( pH > 8), H 4SiO4, a weak acid begins to dissociate and a significant
proportion of dissolved silica is present in the form H 3SiO4H4SiO4
H+ + H3SiO4-
K =
[H]+ [H3SiO4-]
[H4SiO4]
K =
[H]+ [H3SiO4-]
[SiO2][H2O]2
Combining the above two reactions, we get:
SiO2 + 2H2O
H+ + H3SiO4-
(2)
=
Under these conditions, the solubility of silica a
strong function of pH and temperature. Thus,
unlike CaCO3, SiO2 solubility is enhanced by
higher temperatures and pH’s, but relatively
insensitive to pressure.
Note:
The higher solubility of amorphous silica
versus that of crystalline quartz results in
the conversion of primary opal to quartz
during diagenesis. A similar phenomena
results in the conversion of aragonite to
calcite during diagenesis.
[H]+ [H3SiO4-] in the presence
[H2O]2 of amorphous SiO2
Chert
River waters are saturated in SiO2 and yet the
ocean is undersaturated. Clearly there must be a
mechanism for removing silica from ocean water.
In the modern environment diatoms (single cell
aquatic algae with exterior silica shells) and
radiolaria (single cell planktonic protozoan with
internal silica tests) construct their skeletons of
opaline silica. When these floating pelagic plants
and animals die, their skeletons rain to the
bottom where they accumulate as a siliceous
ooze. Modern oozes are dominated by diatoms,
but radiolaria dominated from the early Paleozoic
until the Cretaceous.
diatoms
Without their organic cover, these skeletons start
to dissolve forming the silica-saturated interstitial
water of the ooze, which then aids the
recrystallization of the opal into crytocrystalline
quartz during diagenesis. Some cherts exhibit
ripple cross lamination indicative of current
reworking, but most lack internal structures and
commonly all trace of the former organic
skeletons are erased during diagenesis.
Radiolaria
Chert
Bedded:
Characterized by layers of chert several centimeters
in thickness separated by thin seams of siliceous
shale. Form in areas where there is little other
siliciclastic or carbonate sedimentation, such as
deep water marine environments, below the
carbonate compensation level.
Commonly
associated with pillowed volcanic rocks, turbidite
sandstones, and other deep water sediments.
Nodular:
Chert nodules commonly form in shelf-type
carbonate rocks during diagenesis due to the
remobilization of silica from the skeletons of
diatoms and/or radiolarian, as well as sponge
spicules.
Abyssal Basins (> 2000 m):
The deep ocean basins are characterized by water depths of 2000 - 5000 m. The sediments in such environments are
typically clays and oozes formed by the accumulation of the skeletons of pelagic organisms that have settled through
the water column. Distal turbidites may also be present.
Clay and pelagic material also settle out in shallow water environments, but are swamped by the input of clastic
sediment. Sedimentation rates in the deep oceans, however, are very low (on the order of 1 cm/ 1000yrs).
Red Terrigenous Clay:
The most widespread sediment on the abyssal floor is red clay.
This clay probably has a complex origin, including wind blown
dust, meteoritic dust, volcanic ash, and possible clay sized
terrigenous material. The red colour of this sediment probably
reflects the relatively oxidizing nature of the cold abyssal
waters, which are poor in organic material.
Calcareous Oozes:
At water depths less than 3500 meters, calcareous oozes are
common. These represent the accumulated skeletons of the
floating foraminifera, such as "Globigerina", one of the few
modern organisms that secrete calcite tests. Such sediments
are not found in deeper water, below the carbonate
compensation depth, because CaCO3 is undersaturated in these
colder deep waters, and the skeletons dissolve.
Siliceous Oozes:
Below 3500 meters, siliceous oozes become common, formed
by the settling of the silica skeletons of diatoms (aquatic algae)
and radiolarian (planktonic protozoan).
Diatoms occur
preferentially in the colder waters of high latitudes.
Iron Formation
Iron formations are Fe - rich
deposits characterized by the
cyclic alternation of Fe-rich
and silica-rich bands on the
scale of centimeters to meters.
Although a few Fe-formations
are Paleozoic, the majority are
Precambrian in age. This age
distribution might reflect a
lower oxygen fugacity for the
Precambrian atmosphere that
kept Fe in its more soluble
Fe2+ state.
Iron Formation
Algoma-type – typically Archean - pre 2600 Ma
Exhalative deposits typically associated with active volcanic vents in deep marine
environments. Associated with mafic and felsic subaqueous volcanics, greywackes,
and black shales. Typically relatively restricted in area (a few kms) and thickness (0.1 10 m). Algoma-type Fe-formations have important connections with volcanogenic
massive sulfide deposits and commonly provide very useful stratigraphic markers in
otherwise monotonous sequences of submarine lavas.
Modern “Black Smoker”
Porpoise Cove Fe-Formation – 4.0 Ga
Iron Formation
Consist of finely laminated alternating iron and silica-rich bands. The silica is
usually in the form of grey chert or jasper, although in metamorphosed deposits it is
commonly converted to more coarsely crystalline quartz.
Four distinct facies according to the type of Fe-bearing minerals:
sulfide
pyrite
FeS2
carbonate
siderite
ankerite
FeCO3 &
CaFe(CO3)2
silicate
greenalite
minnesotaite
stilpnomelane
chamosite
Fe3Si2O5(OH)4
Fe3Si4O10(OH)2
Fe-chlorite
Fe-chlorite
oxide
hematite
magnetite
Fe2O3 &
Fe3O4
Fe formations are commonly zoned, with central sulfide and/or carbonate facies closest to volcanic
vents and marginal silicate and/or oxide facies furthest from the vent. Oxide facies dominates, but
all facies are commonly found in close spatial association. Algoma-type Fe formations are thought
to have formed in anoxic hot-spring environments, perhaps by the aid of bacteria. Deposits
associated with Black Smokers found along mid-ocean ridges are thought to be modern equivalents
Rise in Atmospheric Oxygen
Phanerozoic
Archean
Hadean
Proterozoic
Superior-type - Proterozoic – 1800 - 2400
Ma
Classic banded iron formation (BIF) consists of alternating Ferich and silica-rich bands, with the oxide facies dominating and
the sulfide facies rare. Formed on shallow shelves of stable
continental margins with little obvious connection to
volcanism, for example in the Labrador Trough. BIF are
typically one of the first sediments to be deposited in
Proterozoic basins, and are commonly associated with coarse
clastics and limestones.
The aerial extent and thickness
(several metres to 1000 m) are typically much greater than
Algoma type.
Superior-type Fe formations are thought to be formed by the
rise of reduced Fe-rich waters from the deep ocean basins onto
shallow marine shelves, whose relatively oxidizing
environment caused the precipitation of Fe as Fe3+.
Phanerozoic Ironstones
Oolitic or coated-grain Phanerozoic clastic rocks composed of goethite, hematite, and chamosite coating
detrital quartz grains. Typically relatively restricted in terms of area and thickness (1 to 10’s of metres).
Formed on shallow marine shelves, principally during the lower Paleozoic and the Jurassic-Cretaceous.
The absence of true Fe-formation in the Phanerozoic may reflect the low solubility of Fe3+ in relatively
oxidized marine waters.
Evaporites
Evaporites form in hot dry climates in which
evaporation is strong. Modern evaporite deposits are
restricted to latitudes around 30o N and 30o S, as are
modern deserts, where atmospheric convection
patterns result in the downward movement of cold air,
which becomes undersaturated in water as it is heated
and compressed.
30oN
30oS
Marine Evaporites
All marine evaporite successions exhibit the
same sequence of mineral precipitation
because the composition of sea water is
relatively constant, and appears to have
relatively little changed since the beginning
of the Paleozoic. The sequence is:
Mineral
% total solute
% Res
ρ Brine
aragonite
gypsum/anhydrite
0.3 %
4.5 %
50 %
20 %
1.1
1.126
halite
sylvite, carnalite,
Mg & K salts
77.4 %
17.7 %
10 %
<5%
1.214
1.29
Anhydrite forms by the diagenetic
dehydration of gypsum with a 38 % volume
decrease. There are many back-reactions
with the evolving brines, and numerous
diagenetic reactions that produce a complex
variety of final Mg and K salts.
Formation of Marine Evaporites
1000 metres of sea water will
produce approximately 16 metres of
precipitated evaporite sediment.
Clearly the thick ( 1 km) evaporite
successions of the interglacial
periods of the Cambrian, Devonian,
Permian, Cretaceous, and Miocene
require
mechanism
for
the
processing of enormous volumes of
water.
Formation of Marine Evaporites
Castile
Formation
Texas
Furthermore,
many
evaporite
sequences consist of only one or two
dominate facies, which alternate or
constitute thick successions, for
example:
calcite
and
gypsum/anhydrite, anhydrite and
NaCl, or NaCl and KCl.
Some
recharge mechanism is needed to
buffer brine compositions within
limited salinity ranges for long periods
of time.
alternating varves
of calcite and gypsum
Restricted Basin Recharge
Caspian
Sea
Kora-Bogaz-Gol
(Black Throat Lake)
Caspian Sea,
Turkmenistan
1995
1985
Caspian
Sea
1.2%
salinity
Caspian
Sea
Kara-Bogaz
35%
salinity
Three Models for Development of Marine Layered Evaporite Sequences
In Restricted Basins:
Deep Water - Deep Basin
Paleo-Mediterranean Sea (pre-Eocene)
Shallow Water - Shallow Basin
Persian Gulf Sabkhas
Shallow Water - Deep Basin
Dead Sea
Mediterranean
Sea
Mediterranean
Sea
Persian
Gulf
sabkha
Gibraltor
Dead
Sea
Salt Domes off Brazil
salt
domes
salt
dome
salt
dome
Continental Evaporites
The mineralogy of evaporite
sequences
developed
in
continental playas are more
varied because the composition
of the brine solutions reflects the
local
environment.
The
sequence
of
minerals
precipitated becomes critically
dependent on the ratios of anion
and cation species because of the
branching that occurs as each
evaporite mineral precipitates.
alluvial fans
Death Valley Playa
Death Valley Salts
The solubility of a mineral, and thus the point at which it precipitates
from an evaporating brine solution is determined by the product of the
activities of its constituent ions in solution:
Solubility Product K =
(aNa+)2 × (aSO4=)
(aNa2SO4)
=
(XNa+)2 × (XSO4=)
1
Once the solubility product of a mineral is reached, its precipitation will
quantitatively remove the lessor ion from solution. The concentration of
the more abundant ion, however, will continue to increase with further
evaporation. Each saturation point thus represents a junction at which
small differences in concentration lead to completing different
precipitation sequences.
The buffered composition of sea water leads to a constant evaporation
sequence in marine exaporites. Small differences in water composition in
continental environments, however can lead to distinctly different
evaporation sequences, which we take advantage of economically.