52 WHAT HAPPENS IF THE TETRAHEDRAL SHEETS AND OCTAHEDRAL

52 what happens if the tetrahedral sheets and octahedral sheets don’t fit together? the serpentine group minerals with a general for

52
WHAT HAPPENS IF THE TETRAHEDRAL SHEETS AND OCTAHEDRAL SHEETS DON’T FIT
TOGETHER?
The serpentine group minerals with a general formula Mg3Si2O5(OH)4 has
three forms : lizardite, chrysotile and antigorite. In each of these
minerals there is a different way of coping with the
tetrahedral-octahedral mismatch.
In lizardite, the sheets are flat, but the tetrahedra have to rotate
so that the apices can connect to the octahedral sheet, just as in
kaolinite.

In chrysotile, the tetrahedra are tilted and the layers are curved to
accommodate the longer repeat length wuthin the octahedral layer.
Continued growth of the curved layer results in a crystal in which the
layers are rolled up as in a carpet.

In antigorite, the tilting of the tetrahedra which leads to the curved
layers is accompanied by a periodic switching of the direction of the
tetrahedral layer.

Another consequence of the misfit between the sheets is that the
crystals are usually rather small, This is the case in clay minerals.
CLAY MINERALS
Clay minerals are fine-grained (<0.002mm) sheet silicate minerals
which form as a result of weathering of other silicates. The main clay
minerals are:
Kaolinite - Al2Si2O5(OH)4

Illite ~ K0.8Al2 (Al0.8Si3.2)(OH)2
i.e. essentially fine-grained muscovite

Smectite ~ Ca0.17(Al,Mg,Fe)2(Si,Al)4O10(OH)2 . nH2O
S
mectites form a large group of TOT layer minerals with both cations
and water molecules between the layers. They may be either
dioctahedral or trioctahedral and the net charge on the TOTlayers is
the result of various substitutions in both tetrahedral and octahedral
sheets (e.g. Al3+ for Si4+ in the tetrahedral sheets, or Mg2+ (or Fe2+)
for Al3+ in the octahedral sheet.). The layer charge is not as high as
in illite, so there are fewer cations in the interlayers, and also H2O
can easily move in and out of the interlayers, causing the structure
to expand and contract. The interlayer spacings can change from 10Å
(when the clay is dry) to 15.2Å (when there are two layers on water
molecules between the TOT layers). Smectites are therefore sometimes
referred to as the swelling clays.
Vermiculite ~ (Mg,Ca)0.3(Mg,Fe2+,Fe3+,Al)3(Si,Al)4O10(OH)2
Another group of swelling clays minerals is called vermiculite. It
usually formed by the alteration of biotite by the oxidation of some
Fe2= to Fe3+, thus lowering the negative charge in the biotite TOT
layer. It has a higher layer charge than smectite.
Chlorite: (Mg,Fe,Al)3(Si,Al)4O10(OH)2.
(Mg,Fe,Al)3(OH)6
I
n clays chlorite can have a more complex composition than the larger
chlrite crystals in rocks, and can be dioctahedral or trioctahedral.
FRAMEWORK SILICATES
In the framework silicates all [SiO4] tetrahedra share their corners
with others, generally forming rather open three-dimensional networks.
If no ion substitutes for Si, the entire framework has the composition
SiO2 and all valence bonds are satisfied.
When Al substitutes for Si in the tetrahedra, interstitial cations are
required to maintain charge balance. The openness of these framework
structures results in rather large interstitial cation sites compared
to those in the chain silicate minerals. Thus Na+ and Ca2+ will be
considered as 'small' cations compared to K+, while ions such as Mg2+
are too small to play a role in these structures.
The aluminosilicate framework minerals are by far the most abundant
minerals in the Earth's crust, the feldspars making up about 65% by
volume.
In this course we will consider only the two most important groups of
framweork silicates: the silica minerals and the feldspars.
The silica minerals
S
ilica, SiO2, occurs in a number of different forms in the Earth.
Quartz, the most common crystalline polymorph is stable up to 857oC;
tridymite is the stable form from 857oC to 1470oC, and then
cristobalite from 1470oC up to the melting point at 1713oC. The high
pressure forms of silica are coesite, stable in the deep crust of the
Earth, and stishovite which is thought to be stable in the Earth's
mantle. Stishovite has the rutile structure and is one of the very few
known materials in which Si occurs in octahedral coordination with
oxygen. The stability relationships of the SiO2 polymorphs are shown
below.
The crystal structures of quartz, tridymite and cristobalite are
related by reconstructive phase transitions, in other words, to
convert one to another as the temperature increases requires breaking
Si-O bonds and creating a new structure. This is a slow process,
especially when the transitions take place on cooling, and both
cristobalite and tridymite can be preserved metastably in volcanic
rocks.
Crystal structure of quartz.
The crystal structure of quartz can be considered as being made up of
6-fold helices, spiralling around the c axis. The figure on the left
shows one such helix. Looking along the helix axis, this gives the
appearance of a hexagonal ring (right). The interlinking of such
helices gives the full three-dimensional structure (bottom).


The structure shown above is that of high quartz which is stable above
573oC. Below that temperature the structure undergoes a displacive
phase transition to low quartz. This transition is fast and is
unavoidable, no matter how fast the quartz is cooled. No breaking of
bonds is involved, just a twisting of the tetrahadra around the
‘joint’ which is the shared oxygen atom between tetrahedra.
High quartz is hexagonal. Low quartz is trigonal.

The structure of low quartz
One consequence of the displacive transition is that there are two
equivalent ways in which the structure can distort to trigonal
symmetry. If different parts of the crystal distort in different ways,
then the boundary between these regions is related by a simple
symmetry operation and is called a twin boundary. The process of the
formation of such boundaries is called transformation twinning and is
a common phenomenon in many minerals.
The structures of tridymite and cristobalite.
These structures are related to one another, but are quite different
from quartz. Both are made up of the same structural unit which is a
layer of tetrahedra, with alternate tetrahedra pointing up and down.

In both structures these layers are stacked one on the other.
In tridymite they are stacked so that there is a two-layer repeat :
ABABAB….. Tridymite is hexagonal.
In cristobalite the layers are stacked with a three-layer repeat :
ABCABC …. Cristobalite is cubic.
When these structures are cooled they should transform cristobalite 
tridymite  quartz. As mentioned above, these are reconstructive
transitions, and so very slow and difficult. If they do not occur, the
crustobalite (and tridymite) will cool to lower temperatures (~ 200oC)
and then distort by a displacive transition to low cristobalite and
low tridymite, which have lower symmetry than the high forms.
When tridymite and cristobalite are found in rocks they are always the
low forms. These do not appear on the phase diagram because they are
not thermodynamically stable. They are metastable, just as diamond is
relative to graphite at room P,T.
The feldspars
Feldspars make up approximately 70% of the Earth’s crust, yet are
structurally the most complex mineral group because of the many phase
transitions which occur on cooling.
In the feldspars some Al3+ substitutes for Si4+ in the framework, and
charge balance is achieved by cations sited in the open spaces of the
framework structure.
The most common feldspars are the alkali feldspars, with compositions
between KAlSi3O8 and NaAlSi3O8 , and the plagioclase feldspars with
compositions between NaAlSi3­O8 and CaAl2Si2O8.

At high temperatures there is a complete solid solution between the
alkali feldspar end-members. At lower temperatures the solid solution
exsolves (or unmixes) into regions that are Na-rich and regions which
are K-rich. The result is that a single crystal of alkali feldspar
contains an intergrowth at lower temperatures. This is called a
perthitic texture. The scale of this perthitic texture depends on the
cooling rate of the rock, but can usually only be seen by using a
microscope.
There is also a complete solid solution in the plagioclase feldspars
at high temperatures. This solid solution is more complex because it
involves the coupled substitution of Na for Ca and Al for Si: Na+ + Si4+
 Ca2+ + Al3+
At lower temperature this solid solution also breaks down, but the
situation is more complex than alkali feldpars and the intergrowths
can only be seen by electron microscopy.
The structure of the feldspar minerals.
The basic high temperature structure of the feldspars is shown below.

a.
The basic building unit of the feldspar structure is the
four-membered ring of tetrahedra with a pair of tetrahedra
pointing up and a pair pointing down. Al and Si occupy these
tetrahedra. (b) The four-fold rings are joined to form a layer in
which the rings are related by mirror planes parallel to (010) and
diads parallel to the b axis. Two sets of individual tetrahedra
are distinguishable in this layer, and are labelled T1 and T2. The
T1 tetrahedra are all related to one another by symmetry, as are
the T2 tetrahedra. In the third dimension these layers are joined
to each other by the apices. The K, Na and Ca cations occupy the
large oval-shaped cavities between the rings.
At lower temperatures the feldspar structures undergo:
(i) displacive transitions, (ii) exsolution processes and (iii) Al,Si
ordering processes. These processes are not independent and result in
major complexities in the structures, which are not yet fully
understood.
The tendency for Al and Si atoms to become ordered in the tetrahedral
sites at lower temperatures is a very general phenomenon.
Understanding such Al,Si ordering transitions in minerals is a major
topic in modern mineralogy because it has important consequences to
the stability (thermodynamics) of minerals and rocks. It will be
discussed in more detail in later Mineralogy courses.
CARBONATE MINERALS : CALCITE, DOLOMITE, ARAGONITE
Trigonal and orthorhombic carbonates
====================================
The geologically important carbonate minerals may be divided into two
structural groups. The trigonal structure is adopted by carbonates of
small cations while the orthorhombic structure is formed by carbonates
of larger cations. The table below lists some members of each group
with the ionic radius (in Å) of the cation.
Trigonal (Calcite group)
Orthorhombic (Aragonite group)
Calcite CaCO3 (0.99)
Aragonite CaCO3 (0.99)
Magnesite MgCO3 (0.66)
Witherite BaCO3 (1.34)
Siderite FeCO3 (0.74)
Strontianite SrCO3 (1.13)
Rhodocrosite MnCO3 (0.80)
Cerussite PbCO3 (1.20)
Smithsonite ZnCO3 (0.74)
Dolomite CaMg(CO3)
Ankerite Ca (Mg,Fe)(CO3)
The most important carbonate minerals are calcite, dolomite and
aragonite.
T
he structure of calcite CaCO3
=============================
The structure of calcite, CaCO3, can be described in terms of a
rhombohedron (essentially a cube which has been shortened along one of
the triad axes). The Ca atoms have a face-centred distribution on this
rhombohedron, and the triangular CO3 groups lie at the centres of each
edge. This results in layers of CO3 groups lying normal to the c axis
of calcite, with layers of Ca atoms lying between them .
Calcite is the main constituent of limestone rocks.
The structure of dolomite Ca,Mg(CO3)2
The structure of dolomite is very similar to calcite but layers of Ca
cations alternate with layers of Mg cations.
The structure of aragonite
The aragonite structure is preferred by carbonates when the cation
radius is larger than about 1Å. Because Ca2+ lies on this boundary, it
can fit into either structure.
The structure also contains layers of triangular CO3 groups with the
Ca atoms in between, but arrabged in a different way to that in
calcite. The layers lie perpendicular to the c axis of the
orthorhombic unit cell.

T
he calcite-aragonite phase diagram.
From this diagram we can see that aragonite is stable at high
pressures and that calcite is the stable polymorph of CaCO3 at the
earths surface. Aragonite does form in high pressure rocks, but also
forms metastably under low pressure near-surface conditions. As well
as forming the shells of some marine invertebrates, aragonite also
forms in hot-spring and cave deposits, and may also be directly
precipitated from warm sea water.
When a mineral forms outside its stability field it must mean that the
growth occurs under non-equilibrium conditions, and that there is some
kinetic reason for its formation, i.e. it is easier to form aragonite
than calcite under those particular conditions.