Silicate Structures

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 1.3 The structure of silicates

Since Si invariably occurs in tetrahedral coordination the fundamental unit of the silicate structure is the Si-O tetrahedra. The different types of silicate structure arise from the ways in which these tetrahedra are arranged: they may exist as seperate unlinked entitites, as linked finite arrays, as infinite 1-dimensional chains, as infinite 2-dimensional sheets or as infinite 3-dimensional frameworks. These possibilities give rise to the six primary structural types of silicates, each with a characteristic Si : O ratio (more strictly this should be the ratio of tetrahedral cations to oxygen, the reasons for which are discusssed later):

Orthosilicates (or neosilicates): independent Si-O tetrahedra, Si : 0 = 1 : 4, for example the olivine group.

Sorosilicates : two linked Si-O tetrahedra sharing one oxygen, Si : 0 = 2 : 7

Cyclosilicates : closed rings of linked Si-O tetrahedra sharing two oxygens, Si : 0 = 1 : 3, for example beryl.

Inosilicates : continuous chains of Si-O tetrahedra, sharing two oxygens (single chains, Si : 0 = 1 : 3, for example the pyroxene group) or alternately sharing two and three oxygens (double chains, Si : 0 = 4 : 11, for example the amphibole group)

Phyllosilicates: Continuous sheets of Si-O tetrahedra sharing three oxygens, Si : O = 2 : 5, for example the mica group.

Tektosilicates: Continuous framework of Si-O tetrahedra sharing all four oxygens, Si : O = 1 : 2, for example the feldspar group.

The requirement for charge balance or electronic neutrality in these different structural types is maintained by the dispersal of other cations in 6-fold (octahedral), 8-fold (cubic) or 12-fold (icosahedral or close packed) coordintaion between the individual tetrahedra or arrays of tetrahedra in the silicate structure. For example, in single chained inosilicates (Si : O = 1 : 3) there is a net excess of two negative charges per tetrahedra.


where n is the number of tetrahedra in the chain, and 2n- represents the charge excess of the chain forming elements. Theoretically the charge excess could be alleviated in a number of different ways, for example by adding one bivalent cation or two univalent cations per 3 oxygens. However, the location of these cations, which must reside in spaces (termed sites) between the individual chains of Si-O tetrahedra, must be such that they simultaneously satisfy the requirement for electronic neutrality of all oxygens in the structure. Fortunately, this constraint severely limits the range of compositions and structures found in the inosilicates (as indeed it does with all other silicates).

If we look at the detail of the single chain structure we find that in each tetrahdron there are two linking O atoms that are each bonded to two Si atoms and two peripheral O atoms each with bonds to only one Si atom. Since each Si atom shares its 4+ charge with the surrounding four oxygens of the tetrahedron, the requirement for the electronic neutrality for each of the two linking O atoms is completely satisfied. In contrast, each of the peripheral O atoms have a net excess of one negative charge. In order to satisfy this each of these oxygens can be bonded with 3 neighbouring bivalent cations in octahedral coordination (as shown above) or with four bivalent cations in 8-fold coordination. In the pyroxene group both possibilities occur, each placing profound constraints on the way in which the adjacent chains are located with respect to each other.

In double chain inosilicates (Si : 0 = 4 : 11) the chain forming elements give rise to the basic formula:


where the net excess of charges per 11 oxygens is equivalent to 6 negative charges. The sheet forming elements in phyllosilicates (Si : 0 = 2 : 5) give rise to the basic formula:


Silicate Structures, Structural Formula,
Neso-, Cyclo-, and Soro- Silicates


 This document last updated on 15-Jun-2012

Silicate Structures and Structural Formula

As we discussed in a previous lecture, the relative abundance of elements in the Earth’s crust determines what minerals will form and what minerals will be common.  Because Oxygen and Silicon are the most abundant elements, the silicate minerals are the most common.  Thus, we will spend some time here discussing the structure, chemistry, and occurrence of silicate minerals.  Our systematic discussion of the common rock forming minerals will follow in the lectures throughout the remainder of the course.
Element Wt% Atomic% Volume%
O 46.60 62.55 ~94
Si 27.72 21.22   ~6
Al   8.13   6.47
Fe   5.00   1.92
Ca   3.63   1.94
Na   2.83   2.34
K   2.59   1.42
Mg   2.09   1.84
Total 98.59 100.00 100
In order to discuss the silicates and their structures it is first necessary to remember that the way atoms are packed together or coordinated by larger anions, like oxygen depends on the radius ratio of the cation to the anion, Rx/Rz.
Rx/Rz C.N. Type
1.0 12 Hexagonal or Cubic
Closest Packing
1.0 – 0.732 8 Cubic
0.732 – 0.414 6 Octahedral
0.414 – 0.225 4 Tetrahedral
0.225 – 0.155 3 Triangular
<0.155 2 Linear
Since oxygen is the most abundant element in the crust, oxygen will be the major anion that coordinates the other other cations. Thus, for the major ions that occur in the crust, we can make the following table showing the coordination and coordination polyhedra that are expected for each of the common cations.
Ion C.N.
(with Oxygen)
Coord. Polyhedron Ionic Radius,  Å
K+ 8 – 12 cubic to closest 1.51 (8) – 1.64 (12)
Na+ 8 – 6 cubic to octahedral 1.18 (8) – 1.02 (6)
Ca+2 8 – 6 1.12 (8) – 1.00 (6)
Mn+2 6 Octahedral 0.83
Fe+2 6 0.78
Mg+2 6 0.72
Fe+3 6 0.65
Ti+4 6 0.61
Al+3 6 0.54
Al+3 4 Tetrahedral 0.39
Si+4 4 0.26
C+4 3 Triangular 0.08
The radius ratio of Si+4 to O-2 requires that Si+4 be coordinated by 4 O-2 ions in tetrahedral coordination.
In order to neutralize the +4 charge on the Si cation, one negative charge from each of the Oxygen ions will reach the Si cation. Thus, each Oxygen will be left with a net charge of -1, resulting in a SiO4-4 tetrahedral group that can be bonded to other cations.  It is this SiO4-4 tetrahedron that forms the basis of the silicate minerals.  
Since Si+4 is a highly charged cation, Pauling’s rules state that it should be separated a far as possible from other Si+4 ions.  Thus, when these SiO4-4 tetrahedrons are linked together, only corner oxygens will be shared with other SiO4-4 groups.  Several possibilities exist and give rise to the different silicate groups.

Nesosilicates  (Island Silicates)

If the corner oxygens are not shared with other SiO4-4 tetrahedrons, each tetrahedron will be isolated. Thus, this group is often referred to as the island silicate group.  The basic structural unit is then SiO4-4.  In this group the oxygens are shared with octahedral groups that contain other cations like Mg+2, Fe+2, or Ca+2.  Olivine is a good example: (Mg,Fe)2SiO4

Sorosilicates  (Double Island Silicates)

If one of the corner oxygens is shared with another tetrahedron, this gives rise to the sorosilicate group.  It is often referred to as the double island group because there are two linked tetrahedrons isolated from all other tetrahedrons.  In this case, the basic structural unit is Si2O7-6.  A good example of a sorosilicate is the mineral hemimorphite – Zn4Si2O7(OH).H2O.  Some sorosilicates are a combination of single and double islands, like in epidote – Ca2(Fe+3,Al)Al2(SiO4)(Si2O7)(OH). 

Cyclosilicates (Ring Silicates)

If two of the oxygens are shared and the structure is arranged in a ring, such as that shown here, we get the basic structural unit of the cyclosilcates or ring silicates.  Shown here is a six membered ring forming the structural group Si6O18-12.  Three membered rings, Si3O9-6, four membered rings, Si4O12-8, and five membered rings Si5O15-10 are also possible.  A good example of a cyclosilicate is the mineral Beryl – Be3Al2Si6O18.      

Inosilicates (Single Chain Silicates)

If two of the oxygens are shared in a way to make long single chains of linked SiO4 tetrahedra, we get the single chain silicates or inosilicates. In this case the basic structural unit is Si2O6-4 or SiO3-2.  This group is the basis for the pyroxene group of minerals, like the orthopyroxenes (Mg,Fe)SiO3 or the clinopyroxenes Ca(Mg,Fe)Si2O6

Inosilicates (Double Chain Silicates)

If two chains are linked together so that each tetrahedral group shares 3 of its oxygens, we can from double chains, with the basic structural group being Si4O11-6.  The amphibole group of minerals are double chain silicates, for example the tremolite – ferroactinolite series – Ca2(Mg,Fe)5Si8O22(OH)2.

Phyllosilicates (Sheet Silicates)

If 3 of the oxygens from each tetrahedral group are shared such that an infinite sheet of SiO4 tetrahedra are shared we get the basis for the phyllosilicates or sheet silicates.  In this case the basic structural group is Si2O5-2. The micas, clay minerals, chlorite, talc, and serpentine minerals are all based on this structure.  A good example is biotite – K(Mg,Fe)3(AlSi3)O10(OH)2.  Note that in this structure, Al is substituting for Si in one of the tetrahedral groups. 

Tectosilicates (Framework Silicates)

If all of the corner oxygens are shared with another SiO4 tetrahedron, then a framework structure develops.  The basic structural group then becomes SiO2.  The minerals quartz, cristobalite, and tridymite all are based on this structure.  If some of the Si+4 ions are replaced by Al+3 then this produces a charge imbalance and allows for other ions to be found coordinated in different arrangements within the framework structure.  Thus, the feldspar and feldspathoid minerals are also based on the tectosilicate framework.

General Formula for SilicatesBased on these basic structural units, we can construct a general structural chemical formula for the silicates.  But one substitution in particular tends to mess things up a bit.  This is Al+3, the third most abundant element in the Earth’s crust.  Al+3 has an ionic radius that varies between 0.54 and 0.39 depending on the coordination number.  Thus, it could either fit in 6-fold coordination with oxygen or 4-fold coordination with oxygen. Because Al+3 will go into 4-fold coordination with oxygen, it sometimes substitutes for Si+4.  If such a substitution takes place, it creates a charge imbalance that must be made up elsewhere in the silicate structure.

The other common elements in the Earth’s crust that enter the silicates do so in other types of coordination.  Ions like Al+3, Mg+2, Fe+2, Fe+3, Mn+2, and Ti+4 enter into 6-fold or octahedral sites. Larger ions like Ca+2, and Na+1, are found in octahedral coordination or 8-fold, cubic coordination sites.  Very large cations like K+1, Ba+2, and sometimes Na+1 are coordinated by 12 oxygens in 12-fold coordination sites.

We can thus write a general structural formula for the silicates as follows:


where X represents an  8 to 12 fold coordination site for large cations like K+, Rb+, Ba+2, Na+, and Ca+2.

Y represents a 6-fold (octahedral) site for intermediate sized cations like Al+3, Mg+2, Fe+2, Fe+3, Mn+2, and Ti+4.

Z represents the tetrahedral site containing Si+4, and Al+3.

the ratio p:q depends on the degree of polymerization of the silica (or alumina) tetrahedrons, or the silicate structural type as discussed above.

O is oxygen,

and W is a hyrdoxyl (OH-1) site into which can substitute large anions like F-1 or Cl-1.

The subscripts m, n, and r depend on the ratio of p to q and are chosen to maintain charge balance.

This is summarized in the table shown here.  In this table note that there is very little substitution that takes place between ions that enter the X, Y, and Z sites.  The exceptions are mainly substitution of Al+3 for Si+4, which is noted in the Table, and whether the X site is large enough to accept the largest cations like K+1, Ba+2, or Rb+1.

Site C.N. Ion
Z 4 Si+4
Y 6 Al+3
X 8 Na+1
8 – 12 K+1

Nesosilicates (Island Silicates)

We now turn our discussion to a systematic look at the most common rock forming minerals, starting with the common nesosilicates.  Among these are the olivines, garnets, Al2SiO5 minerals, staurolite, and sphene (the latter two will be discussed in the last lecture on accessory minerals).

As discussed above, the nesosilicates or island silicates are based on the isolated SiO4-4 tetrahedral groups.  In the olivines, the remaining corner oxygens form octahedral groups  that coordinate Mg+2 and Fe+2 ions.

The olivines consist of a complete solid solution between  Mg2SiO4 (forsterite, Fo) and Fe2SiO4 (fayalite, Fa).  There is limited substitution of the following end members:

Ca2SiO4 – larnite

Mn2SiO4 – tephroite

CaMgSiO4 – monticellite (which is commonly found in metamorphosed dolomites)

Also found substituting in octahedral sites are Ni+2 and Cr+3, particularly in Mg-rich olivines.

The phase diagram for the common end members of the olivine solid solution series shows that pure forsterite melts at 1890oC and pure fayalite melts at 1205oC.  Thus, the olivines are sometimes seen be be zoned from Mg-rich cores to more Fe-rich rims, although such zoning is usually limited to 5 to 10% difference between the cores and the rims.  
  • OccurrencePure forsterite is limited to metamorphosed Mg-rich limestones and dolomitic metamorphic rocks.Fo90 – 95 is found in ultrabasic igneous rocks, particularly dunites (>90% by volume olivine), and peridotites (Olivine + Cpx + Opx).Fo60 – 90 is found in basic igneous rocks likes basalts and gabbros, and sometimes in andesites, where it occurs with plagioclase and pyroxene.

    Fa100 – 40 is found in Fe-rich siliceous igneous rocks like rhyolites and granites.

    Mg-rich olivines rarely occur in quartz bearing rocks and quartz rarely occurs with Mg-rich olivine because the reaction shown below runs to the right for most pressures and temperatures.

Mg2SiO4 + SiO2  <=> 2MgSiO3
Fo          Qtz                En

Note however, that Fe-rich olivines can occur with quartz.

  • StructureThe structure of the olivines is illustrated on page 439 of Klein and Dutrow.  Note that 2 different kinds of octahedral sites occur.  One is a regular octahedron, labeled M2, and the other is a distorted octahedron, labeled M1.  Fe+2 and Mg+2 have no particular preference for either site, but if Ca+2 is present it prefers the M2 site.
  • Identifying PropertiesThe olivines are orthorhombic (2/m2/m2/m) and usually green colored in hand specimen.The most characteristic property in thin section is their surface texture that kind of looks like a piece of sandpaper (see photo on the back wall of the Mineralogy lab).Because of their good {010} cleavage and common {100} parting, they show parallel extinction relative to the cleavage or parting.

    Maximum birefringence as seen in the interference colors in thin section varies between 3rd order blue (for Fo rich varieties) and 3rd order yellow (for Fa-rich varieties), but remember that this is the maximum birefringence that will only be seen for grains with a and g parallel to the microscope stage.

    Fo-rich olivines are usually clear in thin section, but Fa-rich olivines show pale yellow, greenish yellow, or yellow amber absorption colors and sometimes show pleochroism with a = g = pale yellow, b = orange, yellow, or reddish brown.

    Because optical properties vary with composition of the olivine, 2V is useful in distinguishing olivine compositions.  Look at the graph on page 11 of Deer, Howie, and Zussman.  From the graph you can see that very Fo-rich olivines(>Fo90) are optically positive with a 2V between 82 and 90o.  Between Fo90 and Fa100 the olivine is optically negative with 2Vg  between 90 and 130 (2V between 90o and  50o.  Thus, by estimating the 2V, you should be able to estimate the composition of the olivine.

    Olivines are distinguished from orthorhombic pyroxenes (opx) easily because olivines show higher maximum birefringence and do not show the characteristic {110} cleavage of the pyroxenes.  They are distinguished from the clinopyroxenes (Cpx) which show inclined extinction relative their {110} cleavage and show a biaxial positive character with a 2V of 50 to 60o.

GarnetsGarnets are isometric minerals and thus isotropic in thin section, although sometimes they are seen to be weakly birefringent (slightly anisotropic).  They are also nesosilicates, and therefore based on the SiO4 structural unit.  The general formula for garnets is:


where the A sites are cubic sites containing large divalent cations, usually Ca, Fe, Mg, or Mn, and the B sites are octahedral sites occupied by smaller trivalent cations, like Al and Fe+3.

Garnets with no Ca in the A site and Al in the B site are called the pyralspite series.  These consist of the end members:

Pyrope – Mg3Al2Si3O12

Almandine – Fe3Al2Si3O12

Spessartine – Mn3Al2Si3O12

Garnets with Ca in the A site are called the ugrandite series and consist of the end members:

Uvarovite – Ca3Cr2Si3O12

Grossularite – Ca3Al2Si3O12

Andradite – Ca3Fe+32Si3O12

Limited solid solution exists between end members of each series.

  • OccurrenceThe garnets occur mostly in metamorphic rocks where they are often seen to form euhedral (well-formed) crystals.The Mg-rich garnet, pyrope, is found in metamorphic rocks formed at high pressure and in eclogites (basalts metamorphosed at high pressure) and peridotites (ultrabasic rocks containing olivine, Opx, Cpx, and garnet).

The Fe-rich garnet, almandine, is the most common garnet and is found in metamorphic aluminous schists.

The Mn-rich variety, spessartine, is limited to Mn-rich metamorphic rocks like meta-cherts.

  • Identifying PropertiesGarnets are generally isotropic although some may be weakly birefringent.  In hand specimen they exhibit a wide range of colors and these are sometimes seen in thin section. Color is controlled by the amounts of Fe+2, Fe+3, Mg+2, and Cr+3 present.Pyrope is usually pinkish red to purplish in hand specimen and is usually clear in thin section.Almandine is usually deep red to brownish black in hand specimen and pink in thin section.

    Spessartine ranges from black to red to brown and orange and is usually pink in thin section.

    Grossularite has a color in hand specimen that reflects the amount of Fe and Mn present and thus ranges from brown to yellow to pink. If Cr is present, the color is usually green.  In thin section grossularite varies in color from clear to brown or green in Cr-rich varieties.

    Uvarovite, with high Cr concentration is usually deep green in hand specimen and green in thin section.

    Andradite ranges from yellow to dark brown, but if appreciable amounts of Ti are present, the color could be black in hand specimen and brown in thin section.

    The composition and identity of the garnets is best determined either by association with other minerals or by more sophisticated techniques such as electron microprobe or XRD.

    Garnets are easily distinguished from other minerals by their high relief, isotropic character, and common euhedral habit.

Al2SiO5 Minerals

The Al2SiO5 minerals are common in aluminous metamorphic rocks (meta-shales and meta-mudstones) and sometimes found in aluminous igneous rocks.

In metamorphic rocks the Al2SiO5 polymorphs provide rather general estimates of the pressure and temperature of metamorphism, with Kyanite indicating relatively high pressure, andalusite indicating low temperature and pressure, and sillimanite indicating high temperature.  Better estimates of pressure and temperature are provided if two of the minerals are present in the same rock.  
  • SillimaniteSillimanite is orthorhombic with a good {010} cleavage.  It generally occurs in long fibrous crystals that are length slow, with extinction parallel to the {010} cleavage.  In sections lying on {001}that show well-developed {110} forms, the cleavage is usually seen to cut across the crystal as shown here.  Maximum birefringence is generally seen to be between 2o yellow to 2o red. Sillimanite is biaxial positive with a 2V of 21 – 31o.
  • AndalusiteAndalusite is also orthorhombic , but shows a length fast character.  It generally tends to occur as euhedral blocky crystals with a maximum birefringence in thin section between 1o yellow and 1o red.  It sometimes shows weak pleochroism with a = rose-pink, b = g = greenish yellow.  Some varieties show a cross, termed the chiastolite cross, which is made up of tiny carbonaceous inclusions oriented along crystallographic directions (see illustration on page 492 of Klein & Dutrow).  Andalusite generally occurs as euhedral crystals with an almost square prism.  It is biaxial  negative with 2V = 73 – 86o.
  • KyaniteKyanite is triclinic and thus shows inclined extinction relative to its good {100}and {010}cleavages and {001} parting.  In hand specimen kyanite is commonly pale blue in color, but is clear to pale blue in thin section.  Because of its good cleavages and parting, two cleavages or partings are seen in any orientation of the crystal in thin section.  These cleavages intersect at angles other than 90o and thus look like parallelograms in two dimensions.  Because Kyanite has high relief relative to other minerals with which it commonly occurs, it stands out in thin section and sometimes appears to have a brownish color.  This color is more due to its high relief and numerous cleavages rather than due to selective absorption.Kyanite is biaxial negative with 2V = 78 -83o

Staurolite (Mg,Fe)2Al9Si4O22(OH)2

Staurolite is a common mineral in medium grade metamorphic rocks, usually metamorphosed shales.In hand specimen and in thin section it characteristically is seen to show staurolite twinning, either the right-angle cross, twinned on {031} or the oblique cross, twinned on {231}  
It is monoclinic, but its optical properties are those of an orthorhombic mineral. It has moderate {010} cleavage, which if present, will cause parallel extinction. It’s most distinguishing property is its pleochroism, with a = colorless, b = pale yellow, and g = golden yellow. Less distinctive are its positive optic sign and 2V = 82 – 90o. In many rocks Staurolite shows twinning, and commonly forms euhedral crystals with well developed {100} and {010} crystal faces. In thin section Staurolite is commonly seen to contain tiny inclusions of other minerals, usually quartz. There are very few minerals which can be confused with Staurolite.

Zircon ZrSiO4

Zircon is a common accessory mineral in nearly all kinds of rocks, particularly the more siliceous igneous rocks, like granites, granodiorites, and syenites.   Still, it is not often found in thin section because it is so hard that it gets plucked out during the grinding of the section. Zircon usually contains high amounts of radioactive elements like U and Th.  Thus, when it is found as inclusions in minerals like biotite, it produces pleochroic haloes in the biotite as seen in thin section.  Because it contains high concentrations of U and Th, it is very useful in obtaining U-Pb and Th-Pb radiometric dates on old rocks.  It is very resistant to weathering and may also survives during metamorphism, allowing for dates to be obtained on the original rock prior to metamorphism (often called the protolith).

In hand specimen Zircon usually occurs as tiny reddish colored crystals.  In thin section, it shows extremely high relief, with w = 1.923 to 1.960 and e = 1.968 to 2.015. and is uniaxial positive. Zircon has high birefringence, with interference colors in the higher orders (lots of reds, pinks and light greens). It is commonly colorless to pale brown or pinkish brown in polarized light without the analyzer. Generally it occurs as small crystals with relief higher than almost anything else in the thin section. This latter property should tip you off to its presence.

Sphene (Titanite) CaTiSiO4(OH)

Sphene is another common accessory mineral in  plutonic igneous rocks like granites, granodiorites, and syenites.  It is also found as larger crystals in metamorphic gneisses and chlorite bearing schists.

In hand specimen as an accessory mineral, it is usually seen as small wedge-shaped crystals with a resinous to adamantine luster and brown to yellow brown color.  In thin section, Sphene,  has a relief similar to that of zircon, and is usually found in small crystals with an elongated diamond shape. It is generally brownish in color, shows a well developed {110} cleavage, and high order interference colors.


Sorosilicates are the double island silicates.  Only one important mineral group, the epidote group, has this structure.

Epidote, Clinozoisite, Zoisite

The important minerals in the epidote group are epidote, clinozoisite, and zoisite.  Since the sorosilicates are based on the Si2O7 -6 group, the structural formula can be written as:


Thus, the epidote group contains both the double tetrahedra and the single tetrahedron, separated by groups of AlO6 octahedra and Ca in nine to 10 fold coordination with Oxygen or OH.

The formula can be rewritten as:


Epidote is the Fe-rich variety and has the above general formula.  Clinozoisite is the Fe-free variety with the chemical formula:


Both clinozoisite and epidote are monoclinic (2/m).  Zoisite has the same chemical formula as clinozoisite, but is orthorhombic.

Epidote is usually pistachio green in color with perfect {001} cleavage and imperfect {100} cleavage.  It is optically negative with a 2V of 64 – 90o.  It usually shows pleochroism with a – colorless to pale yellow, b – greenish yellow, and g – yellowish green, and shows high relief relative to feldspars and quartz.  It’s birefringence is high enough to show 3rd order interference colors.  It usually shows an anomalous blue extinction.

Clinozoisite shows similar relief and cleavage to epidote, but it is optically negative with a 2V of 14 to 90o, shows no pleochroism, and lower birefringence (1st to 2nd order interference colors).  Zoisite is similar to clinozoisite, except it will show parallel extinction relative to faces parallel to the crystallographic axes.

Epidote is a common mineral in low grade metamorphic rocks, particularly metamorphosed volcanic rocks and Fe-Al rich meta shales.   Both Clinozoisite and epidote occur as alteration products of  plagioclase and as veins in granitic rocks.


The cyclosilicates are based on rings of SiO4 tetrahedra, with a Si:O ratio of 1:3  The most common minerals based on this structure are Beryl, Cordierite, and Tourmaline.  
BerylBe3Al2Si6O18 is hexagonal (6/m2/m2/m) with a strong prismatic habit with the form {100} usually the only form present.  It is usually deep green to yellowish green in color.  Beryl forms different gemstones depending on color – Aquamarine when it is pale greenish-blue, Morganite if pink, and emerald if deep green and transparent.  Beryl is a common constituent of coarse grained granitic rocks and pegmatites and is found in aluminous mica schists.In thin section, Beryl shows higher relief than quartz, and is distinguished from quartz by its negative optic sign and length-fast character.  The only other mineral that it can be confused with is apatite, but apatite shows even higher relief than Beryl.


Cordierite is (Mg,Fe)2Al4Si5O18.nH2O.  It is orthorhombic (2/m2/m2/m), but shows a pseudohexagonal character due to its common cyclical twinning on {110}.  In thin section it may show a twinning that looks like albite twinning, which makes it hard to distinguish from plagioclase.  But, cordierite is usually dusted with tiny opaque inclusions.  In thick sections it shows a pale -yellow, violet, pale blue  pleochroism.  It can be distinguished from quartz by its biaxial character.

Cordierite is a common constituent of aluminous metamorphic rocks.  It is common in contact metamorphic rocks where it is commonly associated with sillimanite or andalusite, feldspars and micas.


Tourmaline – Na(Mg,Fe,Mn,Li,Al)3Al6Si6O18(BO3)3(OH)4 is hexagonal (3m) and is commonly found as well-formed prismatic crystals, with a rounded triangular cross section perpendicular to the c crystallographic axis.

Tourmaline is a common mineral in pegmatites (SiO2 – rich igneous rocks with large grain size), where it is associated with quartz and alkali feldspar. It is also found in metasomatized rocks of all types, where it is precipitated from a Boron and Silica – rich fluid phase. 

It’s most distinguishing properties are its uniaxial negative optical character and its pleochroism with w = dark green or dark blue and e = yellow or violet. Tourmaline usually forms in euhedral crystals with well developed prism faces and extinction parallel to the prism faces.Examples of questions on this material that could be asked on an exam

(note that properties that distinguish different minerals will be included in the laboratory exam)

  1. Why are the silicate minerals the most common minerals in the earth’s crust?
  2. Give an alternative name for each of the following groups of silicates (a) nesosilicates, (b) double island silicates, (c) cyclo-silicates, (d) inosilicates, (e) sheet silicates.
  3. What would be the normal zoning pattern in a zoned crystal of olivine?  Explain why this is the case.
  4. What is the difference between the pyralspite garnets and the ugrandite garnets?
  5. Draw a pressure – temperature diagram for the Al2SiO5 Minerals.
  6. State whether each of the following minerals would be most commonly found in igneous rocks, metamorphic rocks, or both igneous and metamorphic rocks: (a) kyanite, (b) olivine, (c) epidote, (d) sillimanite, (e) andalusite (f) tourmaline (g) cordierite (h) garnet.





Silicate Mineral Structural Styles

si_tetra.jpg (31470 bytes) Under the pressure temperature conditions of the crust and mantle, silicon and oxygen will usually combine to form the complex ion SiO44-, which geometrically is a tetrahedron with silica in the center and oxygen at the four corners (a tetrahedron is a geometric shape having four triangular-shaped sides).  Linking these tetrahedra on the corners, either by sharing adjacent oxygens, or by inserting positively charged metal ions between tetrahedra, allows a great many crystal structures to be constructed.

Basically, silicate minerals show six different structural styles:

  1. Isolated tetrahedra are separated by positively charged metal ions all around.  This is for example the structure of the mineral olivine.
  2. Tetrahedra form rings that are linked together by positively charged metal ions.  This structure is for example found in the mineral tourmaline.
  3. Tetrahedra share two of their oxygens with neighbors to form chains.  Between the chains positively charged metal ions are inserted.  This is the structural style found in the pyroxene family of minerals.
  4. Two chains may combine to form double chains or ribbons.  This structural style is found in the amphibole family of minerals.
  5. When adjacent tetrahedra share the three oxygens in one tetrahedral plane, we produce sheets of silica tetrahedrons. These sheets can then be linked together by positively charged ions.  This is the basic structure of micas (biotite, muscovite), and also of clay minerals.
  6. Finally, when all four oxygens are linked to other tetrahedra, either by oxygen sharing or via an inserted metal ion, we get three-dimensional framework structures. This is the structural style of the mineral quartz, and of the feldspar family of minerals.

All, except the ring structures, are found in our rock forming minerals. Despite the limited number of components, the large numbers of resulting silicate minerals have very distinct crystalline structures, and equally distinct physical and chemical properties.   Thus, with a very small number of components a vast array of different compounds and structures can be produced. In that sense the versatile silicone plays in the world of inorganic chemistry a similar role as carbon does in the world of organic chemistry. And indeed, the two elements are direct neighbors in the periodic system and thus show many parallels in their chemical behavior.  It is for that reason that some people (scientists, and in particular science fiction writers) have speculated that carbon based lifeforms (like on our planet) may not be the only way to construct living organisms, and that there may be silicon based lifeforms out there, just waiting to be discovered.



Atoms are very small (on the order of 10-9 mm or 10-12 m), and we can not see them with a microscope. Visible light has a wave length around 400 to 700 10-9 m, about 1000 times larger.  Because the theoretical optical resolution an imaging system is by physical laws limited to 1/2 the wavelength used for observation, visible light is in a way “too coarse” for the observation of atoms.  We can get to smaller wavelengths when we use an electron microscope, but even very fast electrons have an insufficiently small wavelength.  High energy X-rays have a wave length that is small enough to image atoms, but the problem is that they penetrate pretty much everything and are very difficult to manipulate optically.   However, X-rays interact with the atoms of a crystal as they pass through it.   In certain directions the X-rays will be blocked, whereas in other directions they can pass or are reflected by the crystal lattice.  This method of observing crystals is called X-ray diffraction.  If we place a photographic film around a mineral that has an X-ray beam shot at it, these directions of passage and reflection will produce a pattern of lines that will be arranged with characteristic distances for each mineral.  From the line spacings it is possible to calculate the spacing and arrangement of crystal planes (lattice planes), and finally the spacing of the atoms in a mineral.  Although invented early this century, X-ray diffraction is still a widely used method to study the internal structure of all sorts of solids.  For example, the determination of the atomic structure of DNA molecules was accomplished by using X-ray diffraction (uses crystallized DNA).  The scientists who figured out how to do this won the Noble Price in chemistry.

xrdqtz.jpg (12083 bytes) The picture at the left is a X-ray diffractogram of quartz.  The diffraction lines are not all of the same intensity, and the intensity patterns as well are characteristic (and unique) for each mineral.
tunnelem.jpg (38626 bytes) A more recently developed tool to image atoms is the Scanning Tunneling Microscope.  This kind of microscope uses a very tiny probe (needle) that is moved across a sample surface and measures the electron distribution (intensity of the electric field, or more accurately the flow of “tunneling” electrons).   Because atoms have electrons on the outside, they show up as roundish balls.   The picture at left is an image of clusters of Sulfur atoms (4 atoms each), that are attached to a Rhenium surface.  It does show the regular arrangement of the Sulfur clusters (due to the crystal structure of the Rhenium substrate), and the fourfold nature of the Sulfur clusters.  Although the Scanning Tunneling Microscope allows us to image the geometric arrangement of atoms directly, it only allows us to image surfaces.   In order to figure out the 3D arrangement of atoms in a mineral we still have to rely on X-ray diffraction.

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