concentration equilibrium for a specific mineral (i.e., its
saturation or oversaturation), temperature, confounding
materials (such as clay particles), and/or presence of certain cations (e.g., Mg
2+ inhibits precipitation of apatite),
pH, Eh, and buffers (van Beusekom and de Jonge, 1997;
Jordan et al., 2008).
Since estuarine waters are the mixing zone of marine
water and freshwater, in many estuaries the open water
salinity in specific parts of the estuary will fluctuate
between freshwater and marine water (both extremes that
are not conducive to the precipitation of minerals by saturation or oversaturation) or may remain in the brackish
water field (also not conducive to the precipitation of minerals). The best locations for precipitation of minerals are
the hypersaline estuarine high-tidal flats where evaporation and transpiration can increase the salinity of pore
water to levels in excess of 60 ppt and into the precipitation fields of calcite, Mg calcite, or aragonite (> 60 ppt)
as carbonate minerals or gypsum (> 120 ppt) (Logan,
1974). These minerals have been recorded in hypersaline
parts of tidal flats of estuaries.
Where minerals are precipitated by pore waters, they
form in the interstices of the sediment and can cement
(i.e., lithify) the sediment. Surface crusts of cemented
sediment, cemented layers within the sediment, or nodules
(lumps) of cemented sediment can form in this manner
(Pye, 1984; Semeniuk, 2010). In other situations, precipitation of minerals result in mineral crystals scattered in the
sediment.
Some authors, through the use of hydrochemistry, have
inferred precipitation of silica in estuaries, but evidence of
the products of such direct precipitation of silica has not
yet been documented.
Under acidic conditions, phosphorus may react with
aluminum and iron to form minerals such as strengite
and variscite (though uncommon).
While framboids of pyrite are commonly considered to
be the products of bacterial activity, Wilkin and Barnes
(1997a) have documented pyrite forming in the water column. In a shallow, fjord-like estuary of the Pettaquamscutt
River (Rhode Island), there is a gradient with depth of oxic
and anoxic and sulfidic water layers in which iron
monosulfides increased with depth. Pyrite is suspended
in the water column and is present only as framboids,
whereas the sediments contain (dominantly) framboids
as well as euhedral grains and infilled framboids of pyrite.
Pyrite framboids form in the water column where there is
high supersaturation with respect to pyrite and slight
undersaturation with respect to the iron monosulfides,
suggesting that framboid nucleation and growth occurs
in the water column adjacent to, and below, the oxicanoxic interface where supply of ferrous species, sulfide
species, and suitable electron acceptors is available.
Wilkins and Barnes (1997a) conclude that there is about
70 % of pyrite burial flux to the sediments via framboids
forming at the oxic-anoxic interface and then settling.
Infilled and overgrown framboids form diagenetically
(see next section) in the sediments as a result of nucleation
and continued pyrite growth on framboids that had
originated above the sediment-water interface.
Sweeney and Kaplan (1973), in examining pyrite synthesis in the laboratory, found that the path of pyrite formation begins with reaction of hydrogen sulfide and
dissolved iron. The first product is either amorphous iron
sulfide or mackinawite with composition of FeS. In the
presence of limited oxygen, this can change to hexagonal
pyrrhotite, FeS. Further reaction of the above phases with
elemental sulfur produces greigite, Fe 3 S 4 , and, finally,
pyrite, FeS 2 . The greigite has the morphological character
of spherules. Pyrite framboids were found to form only on
spherical nuclei. Thus, chemical, crystallographic, and
textural results indicated the pathway of pyrite formation
and compared favorably with naturally occurring pyrite
in marine/coastal sediment. In these pathways, the primary product is amorphous iron sulfide or mackinawite,
and the later mineralogical alteration is part of diagenesis
(see next section).
Inorganic alteration of one mineral to another
(geological mineralization)
While inorganic precipitation of mineral matter from aqueous environments can be viewed definitively as
mineralization, geologists also view the inorganic transformation of one mineral to another by reaction, transformation, cation substitution, or crystallographic overgrowth/
modification as a variation on mineralization. In this case,
the final mineral product is the outcome of chemical and/or
crystallographic alteration of earlier-formed minerals.
The specific process of mineral alteration in shallow water
or in shallow sediment depths is one expression of
a process termed diagenesis (Larsen and Chillingar,
1979). Where such alteration is occurring contemporaneously or nearly contemporaneously with sedimentation,
it is termed synsedimentary diagenesis; where such
alteration is occurring a short time after deposition of
the sediments (e.g., 100s to 1,000s of years), it is termed
early diagenesis (as compared to later diagenesis that occurs
with deeper burial and groundwater alteration).
Common examples of diagenesis are conversion of calcite and aragonite (both CaCO 3 ) in Mg-rich environments
to dolomite (CaMgCO 3 ), the conversion of iron sulfide in
an anoxic sedimentary environment to iron oxides in an
oxic environment, the conversion of iron sulfide (pyrite)
to another polymorph iron sulfide (marcasite), crystallographic overgrowths of framboidal pyrite, and the phosphatization of carbonate minerals.
The alteration of carbonate minerals to dolomite has
been documented by Cook (1973), von der Borch
et al. (1964, 1975), and Semeniuk (2010). The phosphatization of dolomite in supratidal (hypersaline) environments has been recorded by Cook (1973). In the case of
phosphate diagenesis, while phosphate may be taken up
by organisms to form bio-matter and skeletons, if the
phosphate anion is abundant enough, a proportion may
be involved in mineral transformation in diagenesis.
442
MINERALIZATION
saturation or oversaturation), temperature, confounding
materials (such as clay particles), and/or presence of certain cations (e.g., Mg
2+ inhibits precipitation of apatite),
pH, Eh, and buffers (van Beusekom and de Jonge, 1997;
Jordan et al., 2008).
Since estuarine waters are the mixing zone of marine
water and freshwater, in many estuaries the open water
salinity in specific parts of the estuary will fluctuate
between freshwater and marine water (both extremes that
are not conducive to the precipitation of minerals by saturation or oversaturation) or may remain in the brackish
water field (also not conducive to the precipitation of minerals). The best locations for precipitation of minerals are
the hypersaline estuarine high-tidal flats where evaporation and transpiration can increase the salinity of pore
water to levels in excess of 60 ppt and into the precipitation fields of calcite, Mg calcite, or aragonite (> 60 ppt)
as carbonate minerals or gypsum (> 120 ppt) (Logan,
1974). These minerals have been recorded in hypersaline
parts of tidal flats of estuaries.
Where minerals are precipitated by pore waters, they
form in the interstices of the sediment and can cement
(i.e., lithify) the sediment. Surface crusts of cemented
sediment, cemented layers within the sediment, or nodules
(lumps) of cemented sediment can form in this manner
(Pye, 1984; Semeniuk, 2010). In other situations, precipitation of minerals result in mineral crystals scattered in the
sediment.
Some authors, through the use of hydrochemistry, have
inferred precipitation of silica in estuaries, but evidence of
the products of such direct precipitation of silica has not
yet been documented.
Under acidic conditions, phosphorus may react with
aluminum and iron to form minerals such as strengite
and variscite (though uncommon).
While framboids of pyrite are commonly considered to
be the products of bacterial activity, Wilkin and Barnes
(1997a) have documented pyrite forming in the water column. In a shallow, fjord-like estuary of the Pettaquamscutt
River (Rhode Island), there is a gradient with depth of oxic
and anoxic and sulfidic water layers in which iron
monosulfides increased with depth. Pyrite is suspended
in the water column and is present only as framboids,
whereas the sediments contain (dominantly) framboids
as well as euhedral grains and infilled framboids of pyrite.
Pyrite framboids form in the water column where there is
high supersaturation with respect to pyrite and slight
undersaturation with respect to the iron monosulfides,
suggesting that framboid nucleation and growth occurs
in the water column adjacent to, and below, the oxicanoxic interface where supply of ferrous species, sulfide
species, and suitable electron acceptors is available.
Wilkins and Barnes (1997a) conclude that there is about
70 % of pyrite burial flux to the sediments via framboids
forming at the oxic-anoxic interface and then settling.
Infilled and overgrown framboids form diagenetically
(see next section) in the sediments as a result of nucleation
and continued pyrite growth on framboids that had
originated above the sediment-water interface.
Sweeney and Kaplan (1973), in examining pyrite synthesis in the laboratory, found that the path of pyrite formation begins with reaction of hydrogen sulfide and
dissolved iron. The first product is either amorphous iron
sulfide or mackinawite with composition of FeS. In the
presence of limited oxygen, this can change to hexagonal
pyrrhotite, FeS. Further reaction of the above phases with
elemental sulfur produces greigite, Fe 3 S 4 , and, finally,
pyrite, FeS 2 . The greigite has the morphological character
of spherules. Pyrite framboids were found to form only on
spherical nuclei. Thus, chemical, crystallographic, and
textural results indicated the pathway of pyrite formation
and compared favorably with naturally occurring pyrite
in marine/coastal sediment. In these pathways, the primary product is amorphous iron sulfide or mackinawite,
and the later mineralogical alteration is part of diagenesis
(see next section).
Inorganic alteration of one mineral to another
(geological mineralization)
While inorganic precipitation of mineral matter from aqueous environments can be viewed definitively as
mineralization, geologists also view the inorganic transformation of one mineral to another by reaction, transformation, cation substitution, or crystallographic overgrowth/
modification as a variation on mineralization. In this case,
the final mineral product is the outcome of chemical and/or
crystallographic alteration of earlier-formed minerals.
The specific process of mineral alteration in shallow water
or in shallow sediment depths is one expression of
a process termed diagenesis (Larsen and Chillingar,
1979). Where such alteration is occurring contemporaneously or nearly contemporaneously with sedimentation,
it is termed synsedimentary diagenesis; where such
alteration is occurring a short time after deposition of
the sediments (e.g., 100s to 1,000s of years), it is termed
early diagenesis (as compared to later diagenesis that occurs
with deeper burial and groundwater alteration).
Common examples of diagenesis are conversion of calcite and aragonite (both CaCO 3 ) in Mg-rich environments
to dolomite (CaMgCO 3 ), the conversion of iron sulfide in
an anoxic sedimentary environment to iron oxides in an
oxic environment, the conversion of iron sulfide (pyrite)
to another polymorph iron sulfide (marcasite), crystallographic overgrowths of framboidal pyrite, and the phosphatization of carbonate minerals.
The alteration of carbonate minerals to dolomite has
been documented by Cook (1973), von der Borch
et al. (1964, 1975), and Semeniuk (2010). The phosphatization of dolomite in supratidal (hypersaline) environments has been recorded by Cook (1973). In the case of
phosphate diagenesis, while phosphate may be taken up
by organisms to form bio-matter and skeletons, if the
phosphate anion is abundant enough, a proportion may
be involved in mineral transformation in diagenesis.
442
MINERALIZATION