Structure and composition
The structure of clay minerals is characterized by alternation of sheets, which yield the layer structure.
The composition and configuration of these sheets are different for different clay minerals. There are, however, two
basic types of sheets, composing any given clay mineral
(Grim, 1962; Bridley and Brown, 1984; Chamley,
1989): (i) tetrahedral sheets with one silicon atom
surrounded by four oxygen atoms in a tetrahedral
configuration and (ii) octahedral sheets with an aluminum,
magnesium, or iron atom surrounded by hydroxyl groups
and oxygen in a sixfold coordinated configuration
(Figure 1a–d). Alternation of one tetrahedral and one
octahedral sheet yields the 1:1 structure, whereas
alternation of two tetrahedral sheets with one octahedral
sheet results to the 2:1 structure. In most clay minerals,
substitution in tetrahedral and octahedral sites creates a
charge deficit known as layer charge, which is balanced
by potassium, sodium, calcium, or iron and/or hydrated
cations. These cations are hosted between the layers in
the interlayer space and therefore are called interlayer cations. During hydration, the interlayer cations may cause
expansion or swelling of the interlayer space of some clay
minerals (e.g., montmorillonite; a member of the smectite
family; Figure 1b). In general, clay minerals are classified
by the differences in their layered structures, the type of
cations in the structure, the magnitude of layer charge,
and the type amounts and kinds of exchangeable ions
within the interlayers (Meunier, 2005).
Kaolinite is the most important 1:1 clay mineral
(Figure 1a), whereas illite and smectites have 2:1 structures composed of two tetrahedral and one octahedral
sheet (Figure 1b, d). Chlorite is characterized by a 2:1
structure with an interlayered sheet of brucite (hydroxide
interlayer) (Figure 1c). In addition to these main types of
clay minerals, there are clay minerals that display a
mixed-layered configuration composed of different basic
structures and are classified as mixed-layer clay minerals.
The most common mixed-layer clay mineral present in
marine sediments is illite/smectite (e.g., Fütterer, 2006).
Formation of clay minerals
Clay minerals are the product of physical and chemical
weathering of primary, rock-forming aluminum silicates
(feldspars, amphiboles, pyroxenes, etc.) and hydrothermal
alteration. Different clay minerals are formed under different climate-dependent weathering conditions. Thus, clay
minerals determined in surface sediments and sediment
cores may provide information on present and past,
respectively, weathering conditions and climate of the
source areas as well as transport processes of the terrigenous sediments (see examples below and synthesis by
Chamley, 1989). The clay minerals illite and chlorite, for
example, are common weathering products of igneous
and metamorphic rocks. When mobilized during physical
weathering, these minerals are typically found in highlatitude marine sediments. Kaolinite, on the other hand,
forms under warm and humid conditions by intensive
chemical weathering of feldspars in tropical soils. Thus,
kaolinite is often referred to as a low-latitude mineral.
Kaolinite, however, is also found in polar regions in sedimentary deposits that were formed either under past
warmer and wetter climatic conditions that currently exist
or at low latitudes and later displaced northward through
plate motion. Derived from the alteration of volcanic
rocks, smectite is a good indicator of volcanic sediment
sources.
Determination of clay minerals
Clay minerals are commonly determined both accurately
and precisely in the <2 mm clay fraction by X-ray powder
diffraction (XRD) (Bridley and Brown, 1984; Moore and
Reynolds, 1997). For qualitative and semiquantitative
estimates of clay mineral composition of marine sediments, the Biscaye method (Biscaye, 1965) is widely used
in Marine Geosciences. The method uses oriented samples
of the clay fractions prepared by usual sedimentation techniques (Moore and Reynolds, 1997) and subsequent ethylene glycol (EG) solvation and assumes that the sum of the
four clay minerals is 100 %.
For example the abundance of illite, is given by the following equation:
Illite %
ð Þ ¼
4I Â 100
S þ 4I þ 2K þ 2C
where the weighting factors to convert peak areas to relative weight fractions for smectite, illite, kaolinite, and
chlorite are 1, 4, 2, and 2, respectively (Biscaye, 1965).
For smectite the 17 Å peak, after removal of the chlorite
14 Å peak, and for illite the 10 Å peak are used in the calculations. Kaolinite is distinguished from chlorite from the
peaks at 3.57–3.58 Å (kaolinite) to 3.53–3.54 Å (chlorite)
peaks using the slow-scan XRD traces. The relative
amount of each mineral is estimated from the 7 Å
kaolinite-chlorite peak (Biscaye, 1965). Further details
about the X-ray diffraction technique and different evaluation approaches can be found in Heath and Pisias (1979),
Bridley and Brown (1984), and Moore and
Reynolds (1997).
Significance of clay minerals in marine
geosciences
In Marine Geosciences, clay minerals represent a major to
dominant proportion of fine-grained deep-sea sediments
and thus are important proxies for reconstruction of present and past environmental and depositional as well as
diagenetic conditions from studies of surface sediments
and sediment cores. Some examples are presented and
discussed below. For more details the reader is referred
to the literature listed at the end.
The distribution and abundance of clay minerals within
the different ocean basins (e.g., Griffin and Goldberg,
1963; Biscaye, 1965; Venkatarathnam and Biscaye,
1973; Kolla et al., 1976; Heath and Pisias, 1979; Naidu
88
CLAY MINERALS
The structure of clay minerals is characterized by alternation of sheets, which yield the layer structure.
The composition and configuration of these sheets are different for different clay minerals. There are, however, two
basic types of sheets, composing any given clay mineral
(Grim, 1962; Bridley and Brown, 1984; Chamley,
1989): (i) tetrahedral sheets with one silicon atom
surrounded by four oxygen atoms in a tetrahedral
configuration and (ii) octahedral sheets with an aluminum,
magnesium, or iron atom surrounded by hydroxyl groups
and oxygen in a sixfold coordinated configuration
(Figure 1a–d). Alternation of one tetrahedral and one
octahedral sheet yields the 1:1 structure, whereas
alternation of two tetrahedral sheets with one octahedral
sheet results to the 2:1 structure. In most clay minerals,
substitution in tetrahedral and octahedral sites creates a
charge deficit known as layer charge, which is balanced
by potassium, sodium, calcium, or iron and/or hydrated
cations. These cations are hosted between the layers in
the interlayer space and therefore are called interlayer cations. During hydration, the interlayer cations may cause
expansion or swelling of the interlayer space of some clay
minerals (e.g., montmorillonite; a member of the smectite
family; Figure 1b). In general, clay minerals are classified
by the differences in their layered structures, the type of
cations in the structure, the magnitude of layer charge,
and the type amounts and kinds of exchangeable ions
within the interlayers (Meunier, 2005).
Kaolinite is the most important 1:1 clay mineral
(Figure 1a), whereas illite and smectites have 2:1 structures composed of two tetrahedral and one octahedral
sheet (Figure 1b, d). Chlorite is characterized by a 2:1
structure with an interlayered sheet of brucite (hydroxide
interlayer) (Figure 1c). In addition to these main types of
clay minerals, there are clay minerals that display a
mixed-layered configuration composed of different basic
structures and are classified as mixed-layer clay minerals.
The most common mixed-layer clay mineral present in
marine sediments is illite/smectite (e.g., Fütterer, 2006).
Formation of clay minerals
Clay minerals are the product of physical and chemical
weathering of primary, rock-forming aluminum silicates
(feldspars, amphiboles, pyroxenes, etc.) and hydrothermal
alteration. Different clay minerals are formed under different climate-dependent weathering conditions. Thus, clay
minerals determined in surface sediments and sediment
cores may provide information on present and past,
respectively, weathering conditions and climate of the
source areas as well as transport processes of the terrigenous sediments (see examples below and synthesis by
Chamley, 1989). The clay minerals illite and chlorite, for
example, are common weathering products of igneous
and metamorphic rocks. When mobilized during physical
weathering, these minerals are typically found in highlatitude marine sediments. Kaolinite, on the other hand,
forms under warm and humid conditions by intensive
chemical weathering of feldspars in tropical soils. Thus,
kaolinite is often referred to as a low-latitude mineral.
Kaolinite, however, is also found in polar regions in sedimentary deposits that were formed either under past
warmer and wetter climatic conditions that currently exist
or at low latitudes and later displaced northward through
plate motion. Derived from the alteration of volcanic
rocks, smectite is a good indicator of volcanic sediment
sources.
Determination of clay minerals
Clay minerals are commonly determined both accurately
and precisely in the <2 mm clay fraction by X-ray powder
diffraction (XRD) (Bridley and Brown, 1984; Moore and
Reynolds, 1997). For qualitative and semiquantitative
estimates of clay mineral composition of marine sediments, the Biscaye method (Biscaye, 1965) is widely used
in Marine Geosciences. The method uses oriented samples
of the clay fractions prepared by usual sedimentation techniques (Moore and Reynolds, 1997) and subsequent ethylene glycol (EG) solvation and assumes that the sum of the
four clay minerals is 100 %.
For example the abundance of illite, is given by the following equation:
Illite %
ð Þ ¼
4I Â 100
S þ 4I þ 2K þ 2C
where the weighting factors to convert peak areas to relative weight fractions for smectite, illite, kaolinite, and
chlorite are 1, 4, 2, and 2, respectively (Biscaye, 1965).
For smectite the 17 Å peak, after removal of the chlorite
14 Å peak, and for illite the 10 Å peak are used in the calculations. Kaolinite is distinguished from chlorite from the
peaks at 3.57–3.58 Å (kaolinite) to 3.53–3.54 Å (chlorite)
peaks using the slow-scan XRD traces. The relative
amount of each mineral is estimated from the 7 Å
kaolinite-chlorite peak (Biscaye, 1965). Further details
about the X-ray diffraction technique and different evaluation approaches can be found in Heath and Pisias (1979),
Bridley and Brown (1984), and Moore and
Reynolds (1997).
Significance of clay minerals in marine
geosciences
In Marine Geosciences, clay minerals represent a major to
dominant proportion of fine-grained deep-sea sediments
and thus are important proxies for reconstruction of present and past environmental and depositional as well as
diagenetic conditions from studies of surface sediments
and sediment cores. Some examples are presented and
discussed below. For more details the reader is referred
to the literature listed at the end.
The distribution and abundance of clay minerals within
the different ocean basins (e.g., Griffin and Goldberg,
1963; Biscaye, 1965; Venkatarathnam and Biscaye,
1973; Kolla et al., 1976; Heath and Pisias, 1979; Naidu
88
CLAY MINERALS