LECTURE NOTES 1ST SEMESTER UNIT 4
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 Silicates
Based 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:
XmYn(ZpOq)Wr
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
Al+3
Y 6
Al+3
Fe+3
Fe+2
Mg+2
Mn+2
Ti+4
X
8 Na+1
Ca+2
8 - 12
K+1
Ba+2
Rb+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.
Olivines
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.
Occurrence Pure 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.
Structure The 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 Properties The 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 and 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 = = pale yellow, = 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 2V 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.
Garnets
Garnets 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:
A3B2(Si3O12)
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.
Occurrence The 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 Properties
Garnets 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.
Sillimanite Sillimanite 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.
Andalusite
Andalusite 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 = rose-pink, = = 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.
Kyanite Kyanite 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 = colorless, = pale yellow, and = 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.
Sorosilicates
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:
Ca2(Al,Fe+3)Al2O(SiO4)(Si2O7)(OH)
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:
Ca2(Al,Fe+3)Al2Si3O12(OH)
Epidote is the Fe-rich variety and has the above general formula.
Clinozoisite is the Fe-free variety with the chemical formula:
Ca2Al3Si3O12(OH)
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 - colorless to pale yellow, - greenish yellow, and - 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.
Cyclosilicates
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.
Beryl
Be3Al2Si6O18 is hexagonal (6/m2/m2/m) with a strong prismatic habit with the form {10 0} 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
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
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 = dark green or dark blue and = yellow or violet. Tourmaline usually forms in euhedral crystals with well developed prism faces and extinction parallel to the prism faces.
COMPILED BY GDC HANDWARA
LECTURE NOTES
1ST
SEMESTER
UNIT 4
Inosilicates (Single Chain Silicates)
The single chain silicates have a basic structural unit consisting
of linked SiO4 tetrahedra that each share 2 of their oxygens in
such a way as to build long chains of SiO4. The basic structural
group is thus Si2O6 with an Si:O ratio of 1:3. The most
important inosilicates are the pyroxenes. These have a general
structural formula of:
XYZ2O6
where X = Na+, Ca
+2, Mn
+2, Fe
+2, or Mg
+2 filling octahedral sites called M2
Y = Mn+2
, Fe+2
, Mg+2
, Al+3
, Cr+3
, or Ti+4
filling smaller octahedral sites called
M1
Z = Si+4
or Al+3
in tetrahedral coordination.
The pyroxenes can be divided into several groups based on chemistry and
crystallography:
Orthorhombic Pyroxenes (Orthopyroxenes - Opx)
These consist of a range of compositions between enstatite - MgSiO3 and
ferrosilite - FeSiO3
Monoclinic Pyroxenes (Clinopyroxenes - Cpx)
The Diopside- Hedenbergite series - Diopside (CaMgSi2O6) -
Ferrohedenbergite (CaFeSi2O6)
The Sodic Pyroxenes - Jadeite (NaAlSi2O6) and Aegerine (NaFe+3
Si2O6)
Augite is closely related to the diopside - Hedenbergite series with addition
of Al and minor Na substitution - (Ca,Na)(Mg,Fe,Al)(Si,Al)2O6
Pigeonite is also a monoclinic pyroxene with a composition similar to the
orthopyroxenes with more Ca substituting for Fe, and Mg.
The compositional range of the Ca-
rich, Al-free pyroxenes in shown in
the triangular composition diagram
here. Note that there is complete
Mg-Fe substitution and small
amounts of Ca substitution into the
Orthopyroxene solid solution series.
Mg-rich varieties of orthopyroxene
are called hypersthene, whereas Fe-
rich varieties are called Ferrosilite.
There is also complete Mg-Fe solid
solution between Diopside and
Ferrohedenbergite, with some
depletion in Ca. CaSiO3 is the
chemical formula for wollastonite,
but wollastonite does not have a
pyroxene structure.
There is complete Mg-Fe solid solution between the pyroxenes, and as with most
Mg-Fe solid solutions, the Mg-rich end members crystallize at higher
temperatures than the Fe-rich end members.
Solid immiscibility is present between the Diopside - Hedenbergite series and the
Orthopyroxene series. This is seen in the phase diagram below which shows a
hypothetical phase diagram running from the orthopyroxenes to the
clinopyroxenes. Note the solvi. Pigeonite is only stable at higher temperatures
and inverts to orthopyroxene if cooled slowly to lower temperatures. Thus,
pigeonite is only found in volcanic and shallow intrusive igneous rocks, or as
exsolution lamellae in a host augite or opx (more commonly in augite).
When pigeonite or augite exsolve they may form exsolution lamellae that form
parallel to the (001) plane. At lower temperature the exsolution of Opx or augite
result in exsolution lamellae that are parallel to the (100) plane.
All pyroxenes show perfect {110} cleavage.
When viewed looking down the c-
crystallographic axis, the cleavages intersect at
near 90o angles (the angles are actually 92 - 93
o
and 87 - 880). This 90 degree cleavage angle is
most useful in distinguishing pyroxenes from
amphiboles (in amphiboles the cleavages are at
56o and 124
o.
Distinguishing Opx from Cpx in
thin section is accomplished by
noting that in all orthorhombic
pyroxenes the prismatic {110}
cleavage will show parallel
extinction. If looking down the c-
axis the extinction will be
symmetrical relative to the two
cleavage traces.
In Cpx, however, one would see
inclined extinction on all faces
except {100}. Thus, one should
check several grains for extinction
before concluding that the mineral
is Opx, since there is always a
slight chance that one is looking at
a {100} face. Note that in Cpx,
the maximum extinction angle will
only be observed if one is looking
at a {010} face.
Occurrence and Distinction of the Pyroxenes
Augite - is commonly found in both plutonic and volcanic igneous rocks,
as well as high grade meta-igneous rocks like gneisses and granulites. It is
easily distinguished from amphiboles by the nearly 90ocleavage angles, and
is distinguished from Opx by inclined extinction relative to the {110}
cleavage, as discussed above. Augite also has higher maximum
birefringence than Opx, and shows 2nd
to 3rd
order interference colors.
Augite is optically positive with a 2V of about 60o. It shows high relief,
relative to quartz and feldspars and is commonly colorless to brown or
green in thin section, showing no pleochroism.
Hypersthene - is commonly found in both plutonic and volcanic igneous
rocks and in meta-igneous rocks as well. It is distinguished from augite by
its lower interference colors and lack of inclined extinction relative to
{110}. Hypersthene is sometimes pleochroic, showing light pink to light
green colors. The chemical composition of hypersthene can be estimated
using 2V (see p. 163 of DHZ). Compositions close to Enstatite are
optically positive with a 2V of 60 to 90o, whereas intermediate
compositions are optically negative with a 2V of 50 to 90o.
Pigeonite - is generally only found in volcanic igneous rocks, although, as
mentioned above, it can occur as exsolution lamellae in augites of more
slowly cooled igneous rocks. Pigeonite is distinguished from augite by its
lower 2V of 0 to 30o, and is distinguished from hypersthene by its lack of
pleochroism, lower 2V and inclined extinction relative to the {110}
cleavage.
Aegerine (acmite) - Aegerine Augite - are sodic pyroxenes and thus are
found in alkalic igneous rocks associated with sodic amphiboles, alkali
feldspars, and nepheline. The mineral is common in alkali granites, quartz
syenites, and nepheline syenites (all alkalic plutonic rocks), and are also
found in sodic volcanic rocks like peralkaline rhyolites.
Aegerine is distinguished from other clinopyroxenes by a low extinction
angle relative to the {110} cleavage (0 -10o, with augite having an
extinction angle of 35 - 48o), and by the green brown pleochroism present
in aegerine. Aegerine is also optically negative with a 2V of 60 to 70o,
whereas Aegerine-augite has a higher 2V and can be optically positive or
negative. It is distinguished from the pleochroic sodic amphiboles by its
nearly 90o pyroxene cleavage angle.
Jadeite - is a sodium aluminum
pyroxene that is characterized by its
presence in metamorphic rocks
formed at relatively high pressure. It
can form by a reaction of Albite to
produce :
NaAlSi3O8 = NaAlSi2O6 + SiO2
Albite Jadeite Quartz
Jadeite has a lower refractive index
than all other pyroxenes, and has low
birefringence, showing low order 1st
and 2nd
order interference colors.
It is monoclinic with an extinction angle of 33 to 40o, and can thus be
easily distinguished form hypersthene. It is usually colorless in thin
section, helping to distinguish it from augite and aegerine, and has lower
birefringence than augite and aegerine.
Inosilicates (Double Chain Silicates) - The Amphiboles
The amphibole group of minerals is based on the double-chain silicate structure
as shown here. The basic structural unit is (Si4O11)-6
. The structural formula can
be written as:
W0-1X2Y5Z8022(OH,F)2
where W = Na+1
or K+1
in the A site with 10 to 12
fold coordination.
X = Ca+2
, Na+1
, Mn+2
, Fe+2
, Mg+2
, Fe+3
, in an M4 site
with 6 to 8 fold coordination.
Y = Mn+2
, Fe+2
, Mg+2
, Fe+3
, Al+3
. or Ti+4
in an M1 octahedral coordination site.
Z = Si+4
and Al+3
in the tetrahedral site.
There is complete solid solution between Na and Ca end members and among Mg
and Fe end members, with partial substitution of Al+3
for Si+4
in the tetrahedral
site, and partial substitution of F for OH in the hydroxyl site.
The composition of the
common (non-sodic)
amphiboles are shown in the
diagram here. Note the
similarity to the pyroxene
compositional diagram,
above. Actinolite is the solid
solution between Tremolite
[Ca2Mg5Si8O22(OH)2] and
Ferroactinolite
[Ca2Fe5Si8O22(OH)2.]
Cummingtonite - Grunerite is
a solid solution between
Anthophyllite
[Mg7Si8O22(OH)2] and
Grunerite [Fe7Si8O22(OH)2].
Hornblende is the most common amphibole and has more in common with the
Tremolite - Ferroactinolite series, with Al substituting into the Y sites and the
tetrahedral site. It thus has the complicated formula:
(Ca,Na)2-3(Mg,Fe,Al)5Si6(Si,Al)2O22(OH,F)2
The sodic amphiboles have the following formulae:
Glaucophane - Na2Mg3Al2Si8O22(OH)2
Riebeckite - Na2Fe3+2
Fe2+3
Si8O22(OH)2
Arfvedsonite - NaNa2Fe4+2
Fe+3
Si8O22(OH)2
All of the amphiboles except Anthophyllite are
monoclinic, and all show the excellent prismatic
cleavage on {110}. The angles between the
cleavages, however are 56o and 124
o making all
amphiboles easy to distinguish from the
pyroxenes. Looking at faces that show only a
single cleavage trace would show inclined
extinction, except in Anthophyllite.
Occurrence and Distinction of the Amphiboles
Tremolite - Occurs almost exclusively in low grade metamorphic rocks,
particularly those with a high Ca concentration, such as meta-dolomites,
meta-ultrabasic rocks. Tremolite in hand specimen is white in color and
shows a fibrous habit and the characteristic amphibole cleavage. In thin
section it is distinguished from wollastonite and diopside by its amphibole
cleavage. In thin section it is clear with no pleochroism, which
distinguishes it from other amphiboles. It shows high relief, inclined
extinction, and is optically negative with a 2V of about 85o.
Actinolite - Also occurs almost exclusively in low grade metamorphic
rocks, particularly in meta-basalts and meta-gabbros where it is commonly
associated with chlorite. It is green in hand specimen and shows the
characteristic amphibole cleavage, usually showing an elongated habit. In
thin section it shows a characteristic pale yellow to green pleochroism, has
high relief, and is optically negative with a 2V of 60 to 85o.
Hornblende - is a common mineral in both igneous and metamorphic
rocks. In igneous rocks it is found in andesites, dacites, and rhyolites, as
well as in gabbros, diorites, and granites. In metamorphic rocks it is a
common constituent of meta-basalts that have been metamorphosed to
intermediate grades of regional metamorphism (amphibolites). It is also
found in some ultrabasic rocks. In hand specimen it is dark brown to black
in color and shows the characteristic amphibole cleavage. In thin section,
it shows high relief with a characteristic green - brown - yellow
pleochroism. Optic sign and 2V angle cover a wide range and not very
useful in the distinction of hornblende.
Basaltic Hornblende (also called Oxy-hornblende)- is a dark brown to
reddish brown variety of hornblende that results from oxidation during
crystallization of basalts, andesites, dacites, and rhyolites. It usually has a
dark reaction rim that consists of opaque oxide, and is characteristically
pleochroic in yellow to brown to reddish brown colors.
Anthophyllite - does not occur in igneous rocks, but is a constituent of
metamorphic rocks. It is the only orthorhombic amphibole so it is easily
characterized by its parallel extinction relative to the {110} cleavage.
Cummingtonite - Grunerite - is more common in metamorphosed
igneous rocks where members of the series occur with hornblende. It has
been found in siliceous volcanic rocks as well. Cummingtonite is optically
positive, while grunerite is optically negative. Members of this series can
be distinguished from orthorhombic Anthophyllite by the inclined
extinction of the monoclinic Cummingtonite-Grunerite series, and can be
distinguished from tremolite and actinolite by the higher refractive indices
and higher birefringence of the Cummingtonite Grunerite series.
Glaucophane - Riebeckite - Glaucophane is a common mineral in
blueschist facies metamorphic rocks that result from low temperature, high
pressure metamorphism along ancient subduction zones. Riebeckite is
found in alkali granites, syenites, and peralkaline rhyolites. Glaucophane is
easily distinguished from the other amphiboles by its characteristic blue-
lavender pleochroism. Glaucophane is length slow, whereas Riebeckite is
length fast.
Arfvedsonite - occurs most commonly in peralkaline volcanic rocks and
alkaline plutonic igneous rocks, where it typically occurs with the sodic
pyroxene aegerine. Its blue green to yellow green pleochroism distinguish
it from the other amphiboles.
The chart below, also found in your lab assignments, summarizes the properties
used to distinguish the amphiboles.
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Phyllosilicates (Sheet Silicates)
The phyllosilicates, or sheet silicates, are an important group of minerals that includes the
micas, chlorite, serpentine, talc, and the clay minerals. Because of the special importance of
the clay minerals as one of the primary products of chemical weathering and one of the more
abundant constituents of sedimentary rocks, they will be discussed in more detail in the next
lecture.
The basic structure of the phyllosilicates is based on
interconnected six member rings of SiO4-4
tetrahedra that
extend outward in infinite sheets. Three out of the 4
oxygens from each tetrahedra are shared with other
tetrahedra. This leads to a basic structural unit of Si2O5-2
.
Most phyllosilicates contain hydroxyl ion, OH-, with the OH
located at the center of the 6 membered rings, as shown here.
Thus, the group becomes Si2O5(OH)-3
. When other cations are
bonded to the SiO4 sheets, they share the apical oxygens and
the (OH) ions which bond to the other cations in octahedral
coordination. This forms a layer of cations, usually Fe+2
,
Mg+2
, or Al+3
, that occur in octahedral coordination with the O
and OH ions of the tetrahedral layer. As shown, here, the
triangles become the faces of the octahedral groups that can
bind to the tetrahedral layers.
The octahedral layers take on the structure
of either Brucite [Mg(OH)3], if the cations
are +2 ions like Mg+2
or Fe+2
, or Gibbsite
[Al(OH)3], if the cations are +3 like Al+3
.
In the brucite structure, all octahedral sites
are occupied and all anions are OH-1
. In
the Gibbsite structure every 3rd
cation site
is unoccupied and all anions are OH-1
.
This gives rise to 2 groups of sheet silicates:
1. The trioctahedral sheet silicates where each O or OH ion is surrounded by 3 divalent
cations, like Mg+2
or Fe+2
.
2. The dioctahedral sheet silicates where each O or OH ion is surrounded by 2 trivalent
cations, usually Al+3
.
We can build the structures of the various sheet silicates by starting with the octahedral layers
similar to the structures of brucite or gibbsite, as shown below.
The trioctahedral phyllosilicates are based on the
structure where the octahedral layers are similar to
brucite, where Mg+2
occupies the cation position.
The dioctahedral phyllosilicates are based on the
structure where the octahedral layers are similar to
gibbsite, where Al+3
occupies the cation position.
The octahedral sheets in both cases are held together by weak Van der Waals bonds.
If we start with the brucite and
gibbsite structures shown above,
and replace 2 of the OH ions with
O, where the Oxygens are now the
apical Oxygens of the tetrahedral
sheets, then we get the structure of
the serpentine mineral, Lizardite,
if the octahedral layer is
trioctahedral, containing Mg+2
. If
the octahedral layer is
dioctahedral, containing Al+3
, the
structure of the clay mineral
Kaolinite, is obtained.
This leads to a tetrahedral - octahedral (T-O) structure, where each T-O layer is bonded to the
top (or bottom) of another T-O layer by Van der Waals bonds.
If 2 more of the OH ions in the octahedral
layer are replaced by O, and these O
become the apical Oxygens for another
tetrahedral layer, the this builds the
trioctahedral phyllosilicate talc or the
dioctahedral pyrophyllite. This becomes a
T-O-T layer that can bond to other T-O-T
layers by weak Van der Waals bonds.
If an Al+3
is substituted for every 4th
Si+4
in the tetrahedral layer, this causes an excess -1
charge in each T-O-T layer. To satisfy the charge, K+1
or Na+1
can be bonded between 2 T-O-T
sheets in 12-fold coordination.
For the trioctahedral sheet silicates this becomes Phlogopite (Mg-biotite), and for the
dioctahedral sheet silicates this becomes Muscovite. This makes a T-O-T - T-O-T layer that,
again can bind to another T-O-T - T-O-T layer by weak Van der Waals bonds. It is along these
layers of weak bonding that the prominent {001} cleavage in the sheet silicates occurs.
Replacing 2 more Si+4
ions with Al+3
ions in the tetrahedral layer results in an excess -2 charge
on a T-O-T layer, which is satisfied by replacing the K+1
with Ca+2
.
This results in the trioctahedral sheet silicate - Clintonite and the dioctahedral sheet silicate -
Margarite.
Because of the differences in charge balance between the trioctahedral and dioctahedral sheet
silicates, there is little solid solution between the two groups. However, within the
trioctahedral sheet silicates there is complete substitution of Fe+2
for Mg+2
and limited
substitution of Mn+2
into the octahedral sites. Within the dioctahedral sheet silicates there is
limited substitution of Fe+3
for Al+3
in octahedral sites. In addition, F- or Cl
- can substitute for
(OH)- in the hydroxyl site. As previously discussed, substitution of F
-1 stabilizes the mineral to
higher pressures and temperatures.
Another group of phyllosilicates that is more of mixture of structural types is the chlorite
group. Although chlorite is complex in that the amount of Al that can substitute Mg and Si is
variable, one way of looking at the chlorite structure is shown below.
Here, the chlorite structure is depicted as
consisting of a brucite-like layer (with some
Al) sandwiched between tetrahedral layers
that are similar to phlogopite.
Another important sheet silicate structure is that of vermiculite. This is similar to the talc
structure, discussed above, with layers of water molecules occurring between each T-O-T
layer.
Similarly, insertion of layers of water molecules between the T-O-T sheets of pyrophyllite
produces the structure of smectite clays. The vermiculite and smectite groups are therefore
expanding type sheet silicates and as the water is incorporated into the structure the mineral
increases its volume.
Although we have shown that the octahedral layers fit perfectly between the tetrahedral layers,
this is an oversimplification. If the tetrahedral layers were stacked perfectly so that apical
oxygens were to occur vertically aligned, then the structure would have hexagonal symmetry.
But, because this is not the case, most of the phyllosilicates are monoclinic.
Serpentine Group
The serpentine group of minerals has the formula - Mg3Si2O5(OH)4. Three varieties of
serpentine are known. Antigorite and Lizardite are usually massive and fine grained, while
Chrisotile is fibrous. As discussed above, the imperfect fit of the octahedral layers and the
tetrahedral layers causes the crystal structure to have to bend.
In Antigorite the bending
of the sheets is not
continuous, but occurs in
sets, similar to
corrugations, as shown
here.
In Chrisotile, the bending of the sheets
is more continuous, resulting in
continuous tubes that give the mineral
it's fibrous habit. The Chrisotile variety
is commonly referred to as asbestos.
Occurrence - Serpentine is found as an alteration product of Mg-rich silicates like pyroxene
and olivine. It results due to hydration. For example:
2Mg2SiO4 + 3H2O <=> Mg3Si2O5(OH)4 + Mg(OH)2
Olivine water Serpentine Brucite
Thus, serpentine is commonly found pseudomorphed after olivines and pyroxenes in altered
basic and ultrabasic igneous rocks, like altered peridotites, dunites, and sometimes basalts and
gabbros. It is commonly associated with minerals like magnesite (MgCO3), chromite, and
magnetite. If the rock is made up almost entirely of serpentine, it is called a serpentinite.
Properties - Because the serpentines usually occur either as fine-grained aggregates or fibrous
crystals, optical properties are difficult to determine. Most of the time, serpentine can be
distinguished by its characteristic pseudomorphing of other crystals like olivines and
pyroxenes. In hand specimen it generally tends to have a dark green color with a greasy luster.
In thin section it is clear to pale green to pale yellow, but does not show pleochroism, shows a
generally low relief compared to minerals like olivine and pyroxene with which it is associated,
and show very low order interference colors due to its low birefringence.
Talc
Talc has the chemical formula - Mg3Si4O10(OH)2. It is probably best know for its low
hardness. Although it has a micaceous structure, it is so easily deformed, that crystals are
rarely seen.
Occurrence - Like serpentine, talc requires an environment rich in Mg. It is therefore found in
low grade metamorphic rocks that originated as ultrabasic to basic igneous rocks. Rocks
composed almost entirely of talc have a greasy feel and are referred to as soapstone.
Properties - Talc is most easily distinguished in hand specimen by its low hardness, greasy
feel, and association with other Mg-bearing minerals. When crystals are present they show the
characteristic micaceous cleavage on {001}. In thin section, talc is colorless, biaxial negative
with a 2V of 0 to 30o. Like other sheet silicates, it shows the well developed {001} cleavage.
Maximum interference colors, consistent with a birefringence of 0.05 is 3o yellow. Muscovite
has a higher birefringence and higher 2V, properties which easily distinguish the 2 minerals.
Mica Group
The micas can be divided into the dioctahedral micas and the trioctahedral micas, as discussed
above. Muscovite, Paragonite, and Margarite are the white micas, and represent the
dioctahedral group, and Biotite and Clintonite (Xanthophyllite) the black or brown mica,
represents the trioctahedral group. Muscovite and Biotite are the most common micas, but the
Lithium- rich, pink mica, Lepidolite, K(Li,Al)2AlSi3O10(OH)2 is also common, being found
mostly in pegmatites.
Muscovite
Muscovite, KAl3Si3O10(OH)2, and Paragonite, NaAl3Si3O10(OH)2, are two potential end
members of the solid solution series involving K and Na. But, there is a large miscibility gap
between the two end members with Muscovite being between 65% and 100% of K-rich end
member, and Paragonite showing compositions between about 80% and 100% of the Na-rich
end member.
Occurrence - Muscovite is common constituent of Al-rich medium grade metamorphic rocks
where is found in Al-rich schists and contributes to the schistose foliation found in these rocks.
Muscovite is also found in siliceous, Al-rich plutonic igneous rocks (muscovite granites), but
has not been found as a constituent of volcanic rocks. In these rocks it is commonly found in
association with alkali feldspar, quartz, and sometimes biotite, garnet, andalusite, sillimanite, or
kyanite.
Properties - Muscovite is easily identified in hand specimen by its white to sometimes light
brownish color and its perfect {001} cleavage. In thin section, the {001} cleavage is easily
seen and it's high birefringence is exhibited by the large change in relief on rotation of the stage
and it's 2nd to 4th order interference colors. It is clear and shows no pleochroism (which
distinguishes it from Biotite), and it is biaxial negative with a 2V between 28 and 50o. One of
the most diagnostic properties of the micas, including muscovite, is the mottled or birds-eye
extinction exhibited by these minerals.
Biotite
Biotite is a solid solution between the end members Phlogopite KMg3AlSi3O10(OH)2 and
Annite KFe3AlSi3O10(OH)2, although pure Annite does not occur in nature. In addition, small
amounts of Na, Rb, Cs, and Ba may substitute for K, and like in other minerals, F can substitute
for OH and increase the stability of Biotite to higher temperatures and pressures.
Occurrence - Nearly pure phlogopite is found in hydrous ultrabasic rocks like kimberlite, and
is also found in metamorphosed dolomites. Biotite, with more Fe-rich compositions is
common in dacitic, rhyolitic, and trachytic volcanic rocks, granitic plutonic rocks, and a wide
variety of metamorphic rocks. In metamorphic rocks, biotite usually shows a preferred
orientation with its {001} forms parallel to the schistose foliation.
Properties - In hand specimen, Biotite is brown to black and shows the perfect {001}
micaceous cleavage. In thin section, it shows the perfect cleavage and mottled extinction
typical of all micas. It's most characteristic property is its pleochroism, showing yellow to
brown to green colors. Hornblende shows similar pleochroic colors, but is distinguished from
biotite by the differences in cleavage of the 2 minerals. Biotite is biaxial negative with a low
2V of 0o to 25
o.
Chlorite Group
As discussed above, the Chlorite group has a structure that consists of phlogopite T-O-T layers
sandwiching brucite-like octahedral layer. There is substantial substitution of Mg for Fe, and
Al can substitute for (Mg, Fe) in both the octahedral sites, as well as for Si in the tetrahedral
sites. Thus, chlorite can have a rather complicated formula - (Mg,Fe,Al)3(Si,Al)4O10(OH)6.
Occurrence- Chlorite is a common mineral in low grade metamorphic rocks, where it occurs in
association with minerals like actinolite, epidote, and biotite. It also forms as an alteration
product of pyroxenes, amphiboles, biotite, and garnet in igneous as well a metamorphic rocks.
Properties - In hand specimen, chlorite is recognized by its green color, micaceous habit and
cleavage, and association with other minerals like actinolite and epidote. In thin section,
Chlorite shows low relief and low birefringence, with a characteristic midnight blue to black
anomalous interference color. It shows some pleochroism in the range of green to pale yellow.
It is easily distinguished from biotite by its lower relief and anomalous interference color.
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Tectosilicates (Framework Silicates)
The tectosilicates or framework silicates have a structure wherein all of
the 4 oxygens of SiO4-4
tetrahedra are shared with other tetrahedra. The
ratios of Si to O is thus 1:2.
Since the Si - O bonds are strong covalent bonds and since the structure is
interlocking, the tectosilicate minerals tend to have a high hardness.
SiO2 Minerals
There are nine known polymorphs of SiO2, one of which does not occur naturally. These are:
Name Crystal System Density
(g/cm3)
Refractive
Index
(mean)
Stishovite Tetragonal 4.35 1.81
Coesite Monoclinic 3.01 1.59
Low () Quartz Hexagonal 2.65 1.55
High () Quartz Hexagonal 2.53 1.54
Kaetite (synthetic) Tetragonal 2.50 1.52
Low () Tridymite Monoclinic or Orthorhombic 2.26 1.47
High () Tridymite Hexagonal 2.22 1.47
Low () Cristobalite Tetragonal 2.32 1.48
High () Cristobalite Isometric 2.20 1.48
Stishovite and Coesite are high pressure forms
of SiO2, and thus have much higher densities
and refractive indices than the other
polymorphs. Stishovite is the only polymorph
where the Si occurs in 6 fold (octahedral)
coordination with Oxygen, and this occurs due
to the high pressure under which the mineral
forms. Both Stishovite and Coesite have been
found associated with meteorite impact
structures.
At low pressure with decreasing temperature, SiO2 polymorphs change from high Cristobalite -
Low Cristobalite - High Tridymite - Low Tridymite - High Quartz - Low Quartz. The high to
low transformations are all displacive transformations. Since displacive transformations
require little rearrangement of the crystal structure and no change in energy, the high ()
polymorphs do not exist at the surface of the earth, as they will invert to the low ()
polymorphs as temperature is lowered.
Transformations between Cristobalite, Tridymite, and Quartz, however, as well as
between the high pressure polymorphs and Quartz, are reconstructive transformations. Since
reconstructive transformations require significant structural rearrangement and significant
changes in energy, they occur slowly, and the high temperature and high pressure polymorphs
can occur as metastable minerals at the Earth's surface.
Quartz
Quartz is hexagonal and commonly occurs as crystals ranging in size form microscopic to
crystals weighing several tons. Where it crystallizes unhindered by other crystals, such as in
cavities in rock or in a liquid containing few other crystals, it shows well-developed hexagonal
prisms and sometimes showing apparent hexagonal pyramids or dipyramid. When it
crystallizes in an environment where growth is inhibited by the surroundings, it rarely show
crystal faces. It is also found as microcrystalline masses, such as in the rock chert, and as
fibrous masses, such as in chalcedony.
As visible crystals, Quartz is one of the more common rock forming minerals. It occurs in
siliceous igneous rocks such as volcanic rhyolite and plutonic granitic rocks. It is common in
metamorphic rocks at all grades of metamorphism, and is the chief constituent of sand.
Because it is highly resistant to chemical weathering, it is found in a wide variety of
sedimentary rocks.
Several varieties of Quartz can be found, but these are usually only distinguishable in hand
specimen.
Rock Crystal - clear Quartz in distinct crystals - usually found growing in open cavities
in rock.
Amethyst - violet colored Quartz, with the color resulting from trace amounts of Fe in
the crystal.
Rose Quartz - a pink colored variety, that usually does not show crystal faces, the color
resulting from trace amounts of Ti+4
.
Smokey Quartz - a dark colored variety that may be almost black, usually forming well-
formed crystals. The color appears to result from trace amounts of Al+3
in the structure.
Citrine - a yellow colored variety.
Milky Quartz - a white colored variety with the color being due to fluid inclusions.
Milky Quartz is common in hydrothermal veins and pegmatites.
A fibrous variety of Quartz is called Chalcedony. It is usually brown to gray to translucent
with a waxy luster. It is found lining or filling cavities in rock where it was apparently
precipitated from an aqueous solution. When it shows bands of color, it is commonly called by
the following names:
Carnelian - red colored Chalcedony
Chrysoprase - apple-green colored as a result of coloration from NiO.
Agate - alternating curving layers of Chalcedony with different colors or different
porosities.
Onyx - alternating layers of Chalcedony of different colors or porosities arranged in
parallel planes.
Bloodstone - green Chalcedony containing red spots of jasper (see below)
Very fined grained aggregates of cryptocrystalline quartz makes up rock like Flint and Chert.
Flint occurs as nodules in limestone, whereas chert is a layered rock deposited on the ocean
floor. The red variety of flint is called Jasper, where the color results from inclusions of
hematite.
Optical Properties
Quartz is uniaxial positive with a low relief and low birefringence, thus exhibited only 1o gray
to 1o white interference colors. In thin section it is almost always colorless when viewed
without the analyzer inserted. One of its most distinguishing properties in thin section is that it
usually has a smooth, almost polished-like surface texture. Quartz is easily distinguished from
the Feldspars by the biaxial nature of feldspars, and from Nepheline which is uniaxial
negative. Apatite, has similar birefringence to quartz, but is uniaxial negative and has a very
high relief.
In Chalcedony, the fibers are usually elongated perpendicular to the c-crystallographic axis and
thus are length fast. Normal quartz, when it show an elongated habit, is elongated parallel to
the c axis, and is thus length slow.
Tridymite
Tridymite is the high temperature polymorph of SiO2. Thus, it is only commonly found in
igneous rocks that have been cooled rapidly to surface temperatures, preventing the slow
transformation to quartz, the stable form of SiO2 at surface temperatures. Because of this, we
only expect to find Tridymite in siliceous volcanic rocks like rhyolites, where it commonly
occurs as wedge shaped crystals in cavities in the rock. In volcanic rocks, Tridymite is
commonly associated with Cristobalite and Sanidine.
Optical Properties
Tridymite usually occurs as orthorhombic or monoclinic wedge shaped crystals with a positive
2V between 40 and 90o. The wedge shape of the crystals is the result of twinning on {110},
and usually as 2 to 3 twinned individuals. Although it has similar birefringence to quartz and
feldspar, it has lower refractive indices, and thus shows negative relief compared to quartz and
feldspars.
Cristobalite
Cristobalite is also a high temperature SiO2 polymorph, and thus has a similar occurrence to
Tridymite. It also occurs in thermally metamorphosed sandstones. In volcanic rocks it can
occur both as a lining in open cavities, and as fine grained crystals in the groundmass of the
rock.
Optical Properties
Cristobalite is tetragonal and thus uniaxial. It has a negative optic sign and shows lower relief
than quartz, but has similar birefringence.
Opal
Opal is amorphous, and thus a mineraloid, with a formula - SiO2.nH2O.
Feldspars
The feldspars are the most common minerals in the Earth's crust. They consist of three end-
members:
KAlSi3O8 - Orthoclase (or), NaAlSi3O8 - Albite (ab), and CaAl2Si2O8 -Anorthite (an)
KAlSi3O8 and NaAlSi3O8 form a complete solid solution series, known as the alkali feldspars
and NaAlSi3O8 and CaAl2Si2O8 form a complete solid solution series known as the plagioclase
feldspars.
The feldspars have a framework structure, consisting of SiO4 tetrahedra sharing all of the
corner oxygens. However, in the alkali feldspars 1/4 of the Si+4
ions are replaced by Al+3
and
in the plagioclase feldspars 1/4 to 1/2 of the Si+4
ions are replaced by Al+3
. This allows for the
cations K+, Na
+, and Ca
+2 to be substituted into void spaces to maintain charge balance.
Compositions of natural feldspars are shown in the diagram below based on the 3 components -
NaAlSi3O8, - Albite (ab), KAlSi3O8 - Orthoclase (or) and CaAl2Si2O8. The Alkali Feldspars
form a complete solid solution between ab and or, with up to 5% of the an component. The
high temperature more K-rich variety is called Sanidine and the more Na-rich variety is called
anorthoclase.
The plagioclase feldspars are a complete solid solution series between ab and an, and can
contain small amounts of the or component. Names are given to the various ranges of
composition, as shown here in the diagram are:
Albite - ab90 to ab100
Oligoclase - ab70 to ab90
Andesine - ab50 to ab70
Labradorite - ab30 to ab50
Bytownite - ab10 - ab30
Anorthite - ab0 to an10
Plagioclase Feldspars
Plagioclase is the most common feldspar. It forms initially by crystallization from magma. The
plagioclase solid solution series is coupled solid solution where the substitution is:
Na+1
Si+4
<=> Ca+2
Al+3
Thus, the general chemical formula for plagioclase can be written as:
CaxNa1-xAl1+xSi3-xO8
where x is between 0 and 1.
The phase diagram for
the plagiocalse series is
shown here, and shows
that the Anorthite
component has a higher
melting temperature the
than the Albite
component. Thus, on
crystallization, higher
temperatures will favor
more An-rich plagioclase
which will react with the
liquid to produce more
Ab-rich plagioclase on
cooling.
Plagioclase occurs in basalts, andesites, dacites, rhyolites, gabbros, diorites, granodiorites, and
granites. In most of these igneous rocks, it always shows the characteristic albite twinning.
Plagioclase also occurs in a wide variety of metamorphic rocks, where it is usually not twinned.
In such rocks where the plagioclase is not twinned, it is difficult to distinguish from the alkali
feldspars. Plagioclase can be a component of clastic sedimentary rocks, although it is less
stable near the Earth's surface than alkali feldspar and quartz, and usually breaks down to clay
minerals during weathering.
Properties
In hand specimen, plagioclase is most commonly white colored and
shows perfect {100} and good {010} cleavage. It is most easily
identified and distinguished from quartz, sanidine, orthoclase, and
microcline, by its common polysynthetic twinning on {010}. If this
twinning is not present, plagioclase can still be distinguished from
quartz by its cleavage, but cannot easily be distinguished from the
alkali feldspars. If both plagioclase and alkali feldspar occur in the
same rock, the two can usually be distinguished by differences in
color or differences in the extent of weathering.
In thin section, plagioclase commonly shows the characteristic albite
polysynthetic twinning. This twinning is the most characteristic
identifying feature of plagioclase, and makes its identification easy
when present. Although some cross-hatched twinning may also
occur in plagioclase, it is always very simple with only one or two
cross twins per grain. Thus, be careful not to identify plagioclase as
microcline. The cross-hatched twinning in microcline is always
much more complex.
Plagioclase often shows zoning. This is exhibited by the extinction position changing from the
rim to the core of the crystal. Remember that zoning is caused by incomplete reaction of
crystals with liquid during cooling of a solid solution. Often the zoning is very complex, and is
sometimes oscillatory. Normal zoning would show Ca - rich cores and Na - rich rims, but
reverse zoning is possible under certain conditions.
In metamorphic rocks plagioclase may not show twinning making it difficult to distinguish
from orthoclase. The two can be distinguished by staining the thin section with stains that
make the K-feldspars one color and the more Ca-rich feldspars another color. In this class, we
will not have time to look at these staining techniques. You should, however, be aware, that
such staining techniques exist, so that if you need them in the future, you can use them.
The optical properties of the plagioclase series vary widely as a function of composition of the
plagioclase. In general, all plagioclases show low order interference colors, and thus, low
birefringence. Optic sign and 2V vary widely, and are thus, not very distinguishing features of
plagioclase. Although, as you have seen in lab, it is possible to estimate the composition of
plagioclase from a combination of extinction angle and twinning.
Alkali Feldspars (K,Na)AlSi3O8
As an alkali feldspar cools from high temperature to lower temperature, the crystal structure
changes from that of sanidine, which is monoclinic, through orthoclase, also monoclinic, but
with a different crystal structure than sanidine, to microcline, which is triclinic. These
transformations are order-disorder transformations, and thus require large amounts of time.
Furthermore, if the feldspar is allowed to cool very slowly, then exsolution will occur, and the
solid solution will separate into a Na-rich phase and a K-rich phase. Thus, one expects to find
sanidine in rocks that were cooled very rapidly from high temperature, i.e. volcanic rocks.
Orthoclase and microcline will be found in plutonic igneous rocks (cooled slowly at depth in
the earth) and in metamorphic rocks. In addition, in the plutonic rock types if the cooling takes
place slowly enough, then perthitic exsolution lamellae may also form.
All of the alkali feldspars have low relief and low birefringence. Thus the interference colors
may range up to 1o white. Since this is the same interference color we expect for quartz, care
must be taken to avoid confusing feldspars and quartz.
Sanidine
Sanidine generally occurs with an equant habit (almost square) and shows perfect {001} and
{010} cleavages, which readily distinguish it from quartz. Rarely does sanidine show twinning,
but when it does, it is usually simple twinning. Optic axis figures will only be found on sections
showing both cleavages. Sanidine is optically negative with a 2V of 20 - 50o. This distinguishes
it from quartz, which is uniaxial positive, and from the other alkali feldspars which show larger
values of 2V.
Orthoclase
Orthoclase is a common alkali feldspar in granitic rocks and K - Al rich metamorphic rocks. It
often shows perfect {001} and {010} cleavages which will distinguish it from quartz. Also,
quartz usually shows a smooth surface texture, while orthoclase appears much rougher.
Orthoclase is also biaxial, which further distinguishes it from quartz. The 2V of orthoclase
varies from 60 to 105o, and thus it may be either positive or negative. The 2V angle
distinguishes orthoclase from sanidine, but is otherwise not very useful because of the its wide
range.
Microcline
Microcline is the lowest temperature form of alkali feldspar. Upon cooling, orthoclase must
rearrange its structure from monoclinic to triclinic. When this happens, twinning usually
results. The twinning characteristic of microcline is a combination of albite twinning and
pericline twinning. This results in a cross-hatched pattern (often called tartan twinning) that is
the most distinguishing characteristic of microcline.
Anorthoclase
Anorthoclase is a Na - rich feldspar with approximately equal amounts of the Anorthite (Ca)
and orthoclase (K) components. Generally anorthoclase occurs in Na - rich volcanic rocks. Like
the other alkali feldspars, it has perfect {001} and {010} cleavages. Sections showing both of
the cleavages are best for determining the optic sign and 2V. Anorthoclase sometimes shows
twinning, but generally not the multiple twinning seen in the plagioclase feldspars, but a cross-
hatched twinning similar to that seen in microcline, but on a very fine scale. Anorthoclase, like
sanidine shows a low 2V of 5 to 20o, and is optically negative. Anorthoclase can sometimes be
distinguished from sanidine by the fact that anorthoclase usually forms crystals with a tabular,
elongated habit, while sanidine forms crystals with a more equant habit.
Feldspathoids
The feldspathoid group of minerals are SiO2 poor, alkali rich minerals that occur in low SiO2,
high Na2O - K2O igneous rocks. In general, these minerals are not compatible with quartz, and
therefore, are rarely, if ever, seen in rocks that contain quartz. They do, however, often occur
with feldspars. Because of the alkalic nature of the rocks that contain feldspathoids, associated
pyroxenes and amphiboles are of the sodic variety, i.e. aegerine or riebeckite.
The main feldspathoids are Nepheline (Na,K)AlSiO4, Kalsilite KAlSi2O6, and Leucite
KAlSi2O6. At high temperature there is complete solid solution between Nepheline and
Kalsilite, but at low temperature Nepheline can contain only about 12 wt% K2O.
Other similar members of the feldspathoid group are:
Sodalite 3NaAlSiO4.NaCl
Nosean 3NaAlSiO4.NaSO4
Haüyne 3NaAlSiO4.Ca(Cl,SO4)
Nepheline
Nepheline occurs in both volcanic and plutonic alkaline igneous rocks. In hand specimen,
Nepheline is difficult to distinguish from the feldspars, and thus must usually be identified by
its association with other alkalic minerals. Nepheline has a yellowish colored alteration
product, called cancrinite. Nepheline is hexagonal, and thus uniaxial, making it easy to
distinguish from the feldspars. Furthermore, it is optically negative, making it distinguishable
from quartz. It usually shows no cleavage, has low birefringence, and low relief (refractive
indices are smaller than the feldspars). The only other common mineral with which nepheline
could be confused is apatite, which is also uniaxial negative. Apatite, however, shows much
higher relief than does nepheline.
Sodalite
Sodalite occurs predominantly in alkali-rich plutonic igneous rocks, like syenites, but can also
be found in volcanic rocks. It is essentially 3 nepheline molecules with an added NaCl
molecule. It is a clear colored isometric mineral with low relief. Thus, the only thing sodalite
might be confused with is a hole in the thin section. The blue color of sodalite in hand
specimen and its association with other alkali-rich minerals is usually necessary to detect its
presence in a rock.
Leucite
Leucite is found in alkalic volcanic rocks, and is rarely found in plutonic rocks. It is a
tetragonal mineral, however, its refractive indices and are so close together that it almost
always appears isometric. It usually occurs as small, slightly rounded, low relief grains that go
extinct upon insertion of the analyzer. Commonly, leucite contains tiny inclusions within the
mineral, and sometimes shows a slight twinning, barely visible with the analyzer inserted.
Oxides
The oxide minerals are very common and usually occur as accessory minerals in all kinds of
rocks. The most common oxide minerals are the following:
Corundum - Al2O3
Corundum is hexagonal and optically negative. It occurs in Al-rich igneous and metamorphic
rocks. If transparent blue, it is the gemstone sapphire, if transparent red, it is the gemstone
ruby. When it occurs as an accessory mineral it usually shows its hexagon shaped outline when
looking down the c-axis. It has high refractive indices, thus shows very high relief in thin
section. But it has low birefringence and commonly shows lamellar twinning.
Spinel - MgAl2O4
Spinel is an isometric mineral that occurs ultrabasic rocks like peridotite, and in many low
silica ignoeous rocks like basalts, where it contains high concentrations of Cr. It is also found
in Al-rich contact metamorphic rocks. It shows a wide variety of colors depending on trace
amounts of other ions substituting for both Mg and Al. Because of the isometric nature, Spinel
is difficult to distinguish from garnet, although spinel tends to occur as much smaller crystals.
Chromite - Fe+2
Cr2O4
Chromite is a major ore of Cr. It is found in in low silica, Mg-rich igneous rocks, usually
associated with Olivine. Often it is seen as small inclusions in Olivine, indicating that it is an
early crystallizing phase in basaltic and gabbroic magmas. Chromite is isometric, and usually
opaque in thin section. Electron Microprobe analysis is usually necessary to distinguish it from
other opaque oxide minerals.
Magnetite - Fe3O4
Magnetite is one of the most common oxide minerals. It is a major ore of Fe, and is found as
an accessory mineral in all rock types. It is isometric and commonly crystallizes with an
octahedral habit. In hand specimen it is most easily identified by its strongly magnetic nature,
black color, and hardness of 6. In thin section it is opaque and thus difficult to distinguish from
the other opaque oxide minerals. As discussed below, it forms a solid solution with Ulvospinel
- Fe2TiO4.
Ilmenite - FeTiO3
Ilmenite is a major ore of Ti. It is found as a common accessory mineral in a wide range of
igneous volcanic and plutonic rocks, as well as metamorphic and clastic sedimentary rocks. It
forms a solid solution series with Hematite, as will be discussed below, and commonly occurs
along with Magnetite. Ilmenite is hexagonal, but is usually opaque which makes its distinction
from other oxide minerals difficult. Ilmenite, however, often shows an elongated or acicular
habit, whereas Magnetite usually crystallizes as more equant crystals with an octahedral habit.
Hematite - Fe2O3
Hematite is one of the most important ores of Fe. It is more oxidized than Magnetite, and thus
forms as an alteration product of magnetite as well as other Fe bearing minerals. In most
unaltered igneous rocks, hematite occurs as a component of Ilmenite in solid solution.
Hematite is hexagonal, but rarely occurs in crystals where its symmetry can be determined. It
is found in a variety of forms, ranging from oolitic spherules, to massive fine grained
aggregates, to botryoidal masses. It is most easily distinguished by its black to dark red color
and reddish brown streak. In thin section it is not easily distinguished from other opaque oxide
minerals.
Iron-Titanium Oxide Geothermometer
Under magmatic conditions, Ilmenite and Hematite form a complete solid solution series, often
called the rhombohedral series since both minerals crystallize in the hexagonal system.
Similarly Magnetite and Ulvospinel form a complete solid solution series, called the spinel
series.
The possible ranges of solid solution are
shown in the diagram to the right.
Coexisting compositions (as illustrated
by the tie line) depend on temperature
and the fugacity (similar to partial
pressure) of Oxygen. If magma is
rapidly cooled so as to preserve the
compositions of the high temperature
solid solutions, it is possible to calculate
the temperature and fugacity of Oxygen
that were present just before eruption of
the magma. Minerals that allow for
determination of the temperature of
formation of minerals are referred to as a
geothermometer. The example
illustrated here is an important one,
called the Iron-Titanium Oxide
Geothermometer.
Carbonates
The carbonates are an important group of minerals near the Earth's surface. Carbonate minerals
make up the bulk of limestones and dolostones. Are found as cementing agents in clastic
sedimentary rocks, and make up the shells of many organisms. The carbonates are based on
the CO3-2
structural unit, which has carbon surrounded by 3 oxygens in triangular
coordination. Thus each Oxygen has a residual charge of -2/3. In the carbonate structure, no
two triangles share the corner oxygens and the C-O bonds are highly covalent.
There are three structural types of carbonates:
Calcite Group Aragonite Group Dolomite Group
Calcite CaCO3 Aragonite CaCO3 Dolomite CaMg(CO3)2
Magnesite MgCO3 Witherite BaCO3 Ankerite CaFe(CO3)2
Siderite FeCO3 Strontianite SrCO3
Rhodochrosite MnCO3 Cerussite PbCO3
Smithsonite ZnCO3
In addition, there are the hydroxyl Cu carbonates - Malachite, Cu2CO3(OH)2 and Azurite
Cu3(CO3)2(OH)2.
The Calcite Group
The calcite group minerals are all hexagonal. They have Ca, Mg, Fe, Mn, or Zn divalent
cations in 6-fold coordination with the CO3-2
groups, in a structure that is similar to that of
NaCl. All members of this group show rhombohedral cleavage {01 2}, thus breaking into
rhomb-shaped cleavage blocks.
Calcite CaCO3- The most common carbonate mineral is calcite. It is the principal constituent
of limestone and its metamorphic equivalent - marble. Deposits of fine grained calcite in
powder form are referred to as chalk. It forms the cementing agent in many sandstones, and is
one of the more common minerals precipitated by living organisms to form their skeletal
structures.
Calcite is also precipitated from groundwater where it form veins, or in open cavities like caves
and caverns can form the cave decorations - like stalactites and stalagmites, and encrustations.
It is also precipitated from hot springs where it is called travertine.
Calcite does occur in rare igneous rocks called carbonatites. These form from carbonate
magmas. Calcite is also precipitated from hydrothermal fluids to form veins associated with
sulfide bearing ores.
Properties
In hand specimen, calcite is distinguished by its rhombohedral cleavage, its hardness of 3, and
by its effervescence in dilute HCl. It can range in color from white, to slightly pink, to clear,
but dark colored crystals can also occur. In thin section it is most readily distinguished by its
high birefringence, showing high order white interference colors, by its rhombohedral cleavage
and its uniaxial negative character. Because of its high birefringence, it shows a large change
in relief on rotation of the stage. Furthermore, its refractive index direction (low RI
direction) when parallel to the polarizer shows a negative relief when compared to the
mounting medium of the thin section. Calcite can be distinguished from Aragonite by the lack
of rhombohedral cleavage and biaxial nature of Aragonite.
Magnesite MgCO3
Magnesite is a common alteration product of Mg-rich minerals on altered igneous and
metamorphic rocks. Like calcite, it shows perfect rhombohedral cleavage, but unlike calcite, it
does not readily effervesce in dilute HCl. It does, however, effervesce in hot HCl. These
properties and its association with Mg-rich minerals and rocks make it distinguishable from
Calcite.
Siderite FeCO3
Siderite forms complete solid solution series with Magnesite, although the environment in
which the two minerals occur usually determines that either Mg-rich Magnesite or Fe-rich
Siderite will form, and one rarely sees intermediate end members. In hand specimen, siderite is
usually brown colored and effervesces only in hot HCl. In thin section it resembles Calcite, but
has a much higher refractive index than Calcite and is commonly pale yellow to yellow
brown in color without the analyzer inserted.
Rhodochrosite MnCO3
Rhodochrosite is the Mn bearing carbonate, and is thus found only in environments where there
is an abundance of Manganese. It is relatively rare and occurs as hydrothermal veins and as an
alteration product of Mn rich deposits. In hand specimen it show a distinctive pink color along
with the rhombohedral cleavage common to the Calcite group minerals. Hot HCl is required to
make the mineral effervesce.
The Aragonite Group
The Aragonite group of minerals are all orthorhombic, and can thus be distinguished from
minerals of the calcite group by their lack of rhombohedral cleavage. Aragonite (CaCO3) is the
most common mineral in this group.
Aragonite is the higher pressure form of CaCO3 but,
nevertheless occurs and forms at surface temperatures and
pressures. When found in metamorphic rocks it is a good
indicator of the low temperature, high pressure conditions of
metamorphism, and is thus commonly found in Blueschist
Facies metamorphic rocks along with Glaucophane. Water
containing high concentrations of Ca and carbonate can
precipitate Aragonite. Warm water favors Aragonite, while
cold water favors calcite, thus Aragonite is commonly found
as a deposit of hot springs. Aragonite can also form by
biological precipitation, and the pearly shells of many
organisms are composed of Aragonite. Fine needle-like
crystals of Aragonite are produced by carbonate secreting
algae.
Properties
In hand specimen, Aragonite, like calcite effervesces in cold HCl. But, unlike Calcite,
Aragonite does not show a rhombohedral cleavage. Instead it has single good {010} cleavage.
It is usually transparent to white in color and forms in long bladed crystals. Twinning is
common on {110}, and this can produce both cyclical twins, which, when present, make it look
pseudohexagonal, and single twins. In thin section Aragonite is distinguished by its high
birefringence, showing high order white interference colors, its biaxial character with a 2V of
about 18o, and extinction parallel to the {010} cleavage.
The Dolomite Group
Dolomite - CaMg(CO3)2 and Ankerite - CaFe(CO3)2 form a complete solid solution series,
although because Mg-rich environments are much more common than Fe-rich environments,
Mg-rich dolomites are much more common than Ankerites. Ankerite is common mineral in
Pre-Cambrian iron formations. Dolomite is a common constituent of older limestones, probably
the result of secondary replacement of original calcite. It is also found as dolomitic marbles,
and in hydrothermal veins..
Dolomite is a unique chemical composition, as
can be seen in the Magnesite - Calcite phase
diagram shown here. Two solvi exist at low
temperatures. Thus, any high Mg-calcite -
dolomite solid solutions that might exist at high
temperatures would form nearly pure calcite
and pure dolomite at surface temperatures, and
similarly, any Magnesite - Dolomite solid
solutions that might exist at high temperatures
would form nearly pure Magnesite and pure
Dolomite at low temperatures. Thus,
Magnesite and Dolomite commonly occur
together, as do Calcite and Dolomite.
PropertiesDolomite, and therefore rocks containing large amounts of dolomite, like dolostones,
is easily distinguished by the fact that dolomite only fizzes in cold dilute HCl if broken down to
a fine powder. Also, dolostones tend to weather to a brownish color rock, whereas limestones
tend to weather to a white or gray colored rock. The brown color of dolostones is due to the
fact that Fe occurs in small amounts replacing some of the Mg in dolomite.
In thin section it is more difficult to distinguish from calcite, unless it is twined. In order to
facilitate its identification in thin section, the sections are often stained with alizarin red S.
This turns calcite pink, but leaves the dolomite unstained.
If calcite and dolomite
are twinned, they are
easily distinguishable
from one
another. Calcite shows
twin lamellae that are
parallel to the
rhombohedral cleavage
traces and parallel to
the long direction of
the cleavage rhombs.
Thus, the lamellae bisect the acute angle between the cleavages. Dolomite also has twins
parallel to the cleavage faces and parallel to the long direction of the rhombs, but also has twin
lamellae that are parallel to the short dimension of the rhomb. Thus, dolomite would also show
twin lamellae that would bisect the obtuse angle between the cleavage traces.
Accessory Minerals
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 = 1.923 to 1.960 and = 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.
Apatite Ca5(PO4)3(OH,F)
Apatite is another very common and almost ubiquitous (always present) accessory mineral in
igneous rocks and many metamorphic rocks. If the rock contains any phosphorous it is usually
found in apatite. Apatite is hexagonal, hence uniaxial with a negative optic sign. Its refractive
indices = 1.624 to 1.666 and = 1.629 to 1.667 are higher than both quartz and nepheline,
giving apatite a higher relief than these minerals. Its birefringence, expressed as 1o gray
interference colors is similar to that of quartz and nepheline. Quartz, however, is optically
positive. Nepheline, while optically negative, shows much lower relief than does apatite. The
crystal form of apatite is usually distinctive. If cut parallel to {0001}, it usually has a
hexagonal outline. If cut parallel to the C axis, it appears as doubly terminated prisms.
COMPILED BY
GDC HANDWARA
LECTURE NOTES
1ST
SEMESTER
UNIT 4
Pleochroism
With the upper polar removed, many coloured anisotropic minerals display a
change in colour - this is pleochroism or diachroism.It Produced because the two
rays of light are absorbed differently as they pass through the coloured mineral and
therefore the mineral displays different colours. Pleochroism is not related to the
interference colours.
UNIAXIAL OPTICS
Uniaxial minerals have only one optic axis, and belong to the hexagonal and
tetragonal systems.
Minerals in this group include:
nepheline NaAlSiO4
apatite Ca5(PO4)3(F,Cl,OH)
calcite CaCO3
dolomite (Ca,Mg)CO3
quartz SiO2
zircon ZrSiO4
tourmaline - borosilicate
On rotating the calcite rhomb one dot remained stationary but the other dot rotated
with the calcite about the stationary dot.
The ray corresponding to the image which moved is called the Extraordinary
Ray - epsilon.
The ray corresponding to the stationary image, which behaves as though it
were in an isotropic mineral is called the Ordinary Ray - omega.
The vibration direction of the ordinary ray lies in the {0001} plane of the calcite
and is at right angles to the c-axis.
The extraordinary ray vibrates perpendicular to the ordinary ray vibration direction
in the plane which contains the c-axis of the calcite.
If instead of using a calcite rhomb we had used a slab of calcite which had been cut
in a random orientation and placed that on the dots, two images would still appear.
If the random cuts were such that they were perpendicular to the c-axis, then light
travelling through the calcite, along the c-axis would produce only one image
andwould not become polarized.
The c-axis coincides with the optic axis, which is the direction through the mineral
along which light propogates without being split into two rays.
For calcite,
1. The index of refraction for the ordinary ray is uniform omega = 1.658,
regardless of the direction through the grain that the light follows.
2. The index of refraction for the extraordinary ray, epsilon, is variable ranging
from 1.486 to 1.658. The index is dependant on the direction that the light
travels through the mineral.
o If light travels perpendicular to c-axis, epsilon = 1.486.
o If the light travels along the the c-axis, epsilon = 1.658.
o For intermediate directions through the grain epsilon will fall between
the two extremes.
Calcite is used as an example of the formation of the two rays because of the large
difference between the refractive indices (birefringence (delta)).
for calcite, delta = 0.172.
For minerals with a lower birefringence, e.g. quartz, delta = 0.009, the two images
are still produced but show very little separation. The quartz would have to be 20-
25X as thick as the calcite to see the same separation of the dots.
UNIAXIAL OPTIC SIGN
LIGHT PATHS THROUGH UNIAXIAL MINERALS
nepsilon refers to the maximum or minimum index of refraction for the extraordinary
ray,.
nepsilon' refers to an index of refraction for the extraordinary ray which is between
nomega and nepsilon.
For uniaxial minerals any orientation will provide nw, but only one orientation, cut
parallel to the c-axis will yield nepsilon maximum. This orientation is the one which
exhibits the highest interference colour as delta (birefringence), is greatest, and
therefore DELTA (retardation) is greatest
(DELTA = d(ns-nf))
In Calcite omega > epsilon, 1.658 versus 1.485. In other minerals, e.g. quartz, omega < epsilon, 1.544 versus 1.553.
This difference in this refractive index relationship provides the basis for defining the optic sign of uniaxial minerals.
Optically positive uniaxial minerals omega < epsilon
Optically negative uniaxial minerals omega > epsilon
Alternatively,
if extrordinary ray is the slow ray, then the mineral is optically positive.
if extraordinary ray is the fast ray, then the mineral is optically negative.
epsilon refers to the maximum or minimum index of refraction for the extraordinary ray, the value recorded in the mineral
descriptions in the text.
epsilon' refers to an index of refraction for the extraordinary ray which is between omega and epsilon.
For uniaxial minerals any orientation will provide omega, but only one orientation, cut parallel to the c-axis will yield
epsilon maximum. This orientation is the one which exhibits the highest interference colour as delta (birefringence), is
greatest, and therefore retardation (DELTA) is greatest.
(DELTA = d(ns-nf))
Hexagonal and tetragonal systems are characterized by a high degree of symmetry
about the c-axis. Within the 001 or 0001 plane, at 90° to the c-axis, uniform
chemical bonding in all directions is encountered.
Light Paths Through a Mineral
Light travelling along the c-axis is able to vibrate freely in any direction within the
001 or 0001 plane.
No preferred vibration direction allows light to pass through the mineral as if it
were isotropic, this orientation has the lowest interference colour - black to dark
grey.
If the light passes at some angle to the c-axis, it encounters a different electronic
configuration and is split into two rays of different velocities.
The vibration vector of the ordinary ray is parallel to the 001 or 0001 plane, i.e.
perpendicular to the c-axis. The extraordinary ray vibrates across these planes,
parallel to the c-axis.
The ordinary ray has the same velocity regardless of the path it takes, because it
always vibrates in the same electronic environment.
The extraordinary ray velocity varies depending on the direction. If the light travels
nearly parallel to the c-axis, the extraordinary ray vibrates ~ parallel to 001 or
0001, so that nepsilon'~nomega.
If the light travels at right angles to the c-axis, the extraordinary ray vibrates across
the 001 or 0001 plane and nepsilon is most different from nomega.
For intermediate angles to the c-axis:
nomega > nepsilon'
and, nepsilon' > nepsilon.
Whether the extraordinary ray has a higher or lower RI than the ordiniary ray
depends on the chemical bonding and the crystal structure.
In the lab you will determine the indices of refraction for a uniaxial mineral using
grain mounts and the immersion method.
UNIAXIAL INDICATRIX
The indicatrix is a geometric figure, constructed so that the indices of refraction are
plotted as radii that are parallel to the vibration direction of light.
In isotropic minerals the indicatrix was a sphere, because the refractive index was
the same in all directions.
In uniaxial minerals, because nomega and nepsilon are not equal, the indicatrix is an
ellipsoid, the shape of which is dependant on its orientation with respect to the
optic axis. In positive uniaxial minerals, the Z indicatrix axis is parallel to the c-
crystallographic axis and the indicatrix is a prolate ellipsoid, i.e. it is stretched out
along the optic axis.
All light travelling along the Z axis (optic axis), has an index of refraction of nomega,
whether it vibrates parallel to the X or Y axis, or any direction in the XY plane.
The XZ and the YZ planes through the indicatrix are identical ellipses with
nomega and nepsilon as their axes, with the radii of the ellipses equal to the magnitude
of the RI for the ray.
Plotting the indices of light travelling in all directions produces the prolate
ellipsoid, whose axis of revolution is the optic axis, for uniaxial positive minerals;
nomega < nepsilon.
For optically negative minerals the X indicatrix axis corresponds to the optic axis
and the indicatrix is an oblate ellipsoid, i.e. flattened along the optic axis, and
nomega > nepsilon
In each case, for positive and negative minerals the circular section through the
indicatrix is perpendicular to the optic axis and has a radius = nomega.
The radius of the indicatrix along the optic axis is always nepsilon.
Any section through the indicatrix which includes the optic axis is called a
principal section, and produces an ellipse with axes nomega and nepsilon.
A section through the indicatrix perpendicular to the optic axis produces a circular
section with radius nomega.
A random section through the indicatrix will produce an ellipse with axes
nomega and nepsilon
The indicatrix is oriented so that the optic axis is parallel to the c crystallographic
axis.
Random Section Vibration Directions
Random section through the uniaxial indicatrix will give nomega and nepsilon'.
Light travelling from the origin of the indicatrix outwards, construct a wave normal
to the wave front.
A slice through the centre of the indicatrix, perpendicular to the wave normal
forms an ellipse with axes of nomega and nepsilon.
omega vibrates at 90° to the optic axis = short axis of the ellipse
epsilon' vibrates parallel to the optic axis = long axis of the ellipse.
The magnitude of the axes = nomega and nepsilon
BIREFRINGENCE AND INTERFERENCE COLOURS
Birefringence, difference between the index of refraction of the slow and fast rays
and the interference colours for uniaxial minerals is dependant on the direction that
light travels through the mineral.
1. In a sample which has been cut perpendicular to the optic axis, the bottom
and top surfaces will be parallel. The angle of incidence for the light
entering the crystal = 0° and the wave front are not refracted at the interface
and remain parallel to the mineral surface.
o A cut through the indicatrix, parallel to the bottom of the mineral, will
yield the indices and vibration directions of the light. A slice through
the indicatrix is a circular section, with radius nomega.
o No preferred vibration direction, so light passes along the optic axis as
an ordinary ray and retains whatever vibration direction it had
originally.
o Between crossed polars the light passing through the mineral is
completely absorbed by the upper polar and will remain black on
rotation of the stage, The birefringence = 0.
2. Cutting the sample such that the optic axis is parallel to the surface of the
section the following is observed.
o The indicatrix section is a principle section, as it contains the optic
axis. The indicatrix forms an ellipse with axes = nomega and nepsilon,
with the incident light being split into two rays such that:
the ordinary ray vibrates perpendicular to the optic axis,
the extraordinary ray vibrates parallel to the optic axis.
o The birefringence is at a maximum, and in thin section this grain
orientation will display the highest interference colour.
3. A mineral cut in a random orientation, with normally incident light;
o The ordinary ray produced has an index, nomega and vibrates
perpendicular to the optic axis.
o The extraordinary ray has an index nepsilon' and vibrates in the plane
containing the optic axis.
o nepsilon < nomega maximum or minimum, the birefringence is
intermediate between the two extremes.
EXTINCTION IN UNIAXIAL MINERALS
Uniaxial minerals will exhibit all four types of extinction discussed earlier.
The type is dependent on:
1. the orientation that the mineral is cut
2. the presence of cleavage(s) in the grain
Tetragonal minerals
1. Zircon ZrSi04- poor prismatic
2. Rutile Ti02 - good prismatic
o are prismatic and either elongate or stubby II to c axis.
o display prismatic (parallel to c)
o or pinacoidal (perpendicular to c) cleavage.
Depending on how the crystal is cut, and how its indicatrix is cut, dictates what
will be seen in thin section.
Hexagonal Minerals
1. Quartz - SiO2 - no cleavage
2. Apatite - Ca5(PO4)3(F,C1,OH) - rare pinacoidal, prism
3. Calcite - CaC03 - 1 of two cleavages rhombohedral
4. Nepheline - NaAlSiO4 - no cleavage
Hexagonal minerals will exhibit the following forms prisms, pinacoids, pyramids
and rhombohedrons which will exhibit prismatic, pinaciodal and rhombohedral
cleavages.
The birefringence, interference colours and any cleavage displayed by hexagonal
minerals is a function of how the grain has been cut.
PLEOCHROISM IN UNIAXIAL MINERALS
Pleochroism is defined as the change in colour of a mineral, in plane light, on
rotating the stage. It occurs when the wavelengths of the ordinary &
extraordinary rays are absorbed differently on passing through a mineral,
resulting in different wavelengths of light passing the mineral.
Coloured minerals, whether uniaxial or biaxial, are generally pleochroic.
To describe the pleochroism for uniaxial minerals must specify the colour which
corresponds to the ordinary and extraordinary rays.
e.g. Tourmaline, Hexagonal mineral
o omega = dark green
o epsilon = pale green
If the colour change is quite distinct the pleochroism is said to be strong.
If the colour change is minor = weak pleochroism.
For coloured uniaxial minerals, sections cut perpendicular to the c axis will show a
single colour, corresponding to ordinary ray.
Sections parallel to the c crystallographic axis will exhibit the widest colour
variation as both omega and epsilon are present.
BIAXIAL MINERALS
Include orthorhombic, monoclinic and triclinic systems, all exhibit less symmetry
than uniaxial and isotropic minerals.
Minerals in these crystal systems exhibit variable crystal structure, resulting in
variable chemical bonding.
The crystallographic properties of orthorhombic, monoclinic and triclinic minerals
are specified by means of the unit cell measured along the three crystallographic
axes.
It is also necessary to specify 3 different indices of refraction for biaxial minerals:
alpha, beta, gamma are used in text.
where alpha < beta < gamma
A variety of other conventions have been used or suggested, make sure that you are
aware of the convention used in the text you are using, if it is not Nesse.
The maximum birefringence of a biaxial mineral is defined by
(gamma - alpha)
Clarification
1) It takes 3 indices of refraction to describe optical properties
of biaxial minerals, however, light that enters biaxial minerals
is broken into two rays - FAST and SLOW.
2) Ordinary - extraordinary terminology is not used. Both
rays behave as the extraordinary ray did in uniaxial minerals.
The rays are both extraordinary and are referred to as SLOW
RAY and FAST RAY.
o slow = gamma' , between beta and gamma (higher RI)
gamma > gamma' > beta
o fast = alpha' , between alpha and beta (lower RI)
alpha < alpha' < beta
BIAXIAL INDICATRIX
The biaxial indicatrix is similar to the uniaxial indicatrix, except now there are
three principal indices of refraction instead of two. The biaxial indicatrix is
constructed by plotting the principal indices along 3 mutually perpendicular axes.
nalpha plotted along X
nbeta plotted along Y
ngamma plotted along Z
again, nalpha < nbeta < ngamma
So that the length of X<Y<Z.
Indicatrix is a triaxial ellipsoid elongated along the Z axis, and flattened along the
X axis.
Indicatrix has 3 principal sections, all ellipses:
X - Y axes = nalpha & nbeta
X - Z axes = nalpha & ngamma
Y - Z axes = nbeta & ngamma
Random sections through the indicatrix also form ellipses.
The uniaxial indicatrix exhibited a single circular section, a biaxial indicatrix
exhibits two circular sections with radius = nbeta; the circular sections intersect
along the Y indicatrix axis, which also has a radius of nbeta.
Look at the X - Z plane in the above image.
The axes of the ellipse are = nalpha & ngamma.
The radii vary from nalpha through nbeta to ngamma.
Remember that nalpha < nbeta < ngamma, so a radii = nbeta must be present on the X - Z
plane.
The length of indicatrix along the Y axis is also nbeta, so the Y axis and radii nbeta in
X - Z plane defines a circular section, with radius nbeta.
In the biaxial indicatrix the directions perpendicular to the circular sections define
the OPTIC AXES of the biaxial mineral. Optic axes lie within the X - Z plane,
and this plane is the OPTIC PLANE.
The acute angle between the optic axes is the optic or 2V angle.
The indicatrix axis, either X or Z, which bisects the 2V angle is the ACUTE
BISECTRIX or Bxa.
The indicatrix axis, either X or Z, which bisects the obtuse angle between the optic
axes is the OBTUSE BISECTRIX or Bxo.
The Y axis is perpendicular to the optic plane and forms the OPTIC NORMAL.
OPTIC SIGN
For biaxial minerals optic sign is dependant on whether the X or Z indicatrix axis
is the acute bisectrix.
if Bxa is X, mineral is -ve
if Bxa is Z, mineral is +ve
In the special case where 2V = 90°, mineral is optically neutral.
Another convention used is to identify the angle between the optic axes bisected by
the X axis as the 2VX angle; and the Z axis as 2VZ angle.
These two angles can vary from 0 to 180°, such that the following relationship
holds:
2VX + 2VZ = 180°
Using this convention the optic sign is determined by the following:
if 2VZ < 90°, the mineral is +ve.
if 2VZ > 90°, the mineral is -ve.
Light travelling through biaxial minerals is split into two rays -
FAST and SLOW rays which vibrate at 90° to each other.
The vibration directions of the FAST and SLOW rays are defined, or determined,
by the axes of the ellipse or section through the indicatrix, which is oriented at 90°
to the wave normal.
The Refractive Index corresponding to the FAST ray will be between nalpha and
nbeta, and is referred to as nalpha'.
The Refractive Index corresponding to the SLOW ray will be between nbeta and &
ngamma, and is referred to as ngamma'.
With this convention the following relationship will be true for all biaxial minerals:
1. X - will always correspond to the fast ray and will have the lowest RI.
o RI = nalpha, always fast
2. Y - will be either the fast or the slow ray depending on which other
indicatrix axis it is withand its refractive index will be between the lowest
and highest RI for the mineral.
o RI = nbeta, either fast or slow
3. Z - will always correspond to the slow ray and will have the highest RI.
o RI = ngamma, always slow.
COMPILED BY
GDC HANDWARA
LECTURE NOTES
1ST
SEMESTER
UNIT 4
REFLECTION AND REFRACTION
At the interface between the two materials, e.g. air and water, light may be reflected at the interface
or refracted (bent) into the new medium.
For Reflection the angle of incidence = angle of reflection.
For Refraction the light is bent when passing from one material to another, at an angle other than
perpendicular.
A measure of how effective a material is in bending light is called the Index of Refraction (n),
where:
Index of Refraction in Vacuum = 1 and for all other materials n > 1.0.
Most minerals have n values in the range 1.4 to 2.0.
A high Refractive Index indicates a low velocity for light travelling through that particular medium.
Snell's Law
Snell's law can be used to calculate how much the light will bend on travelling into the new medium.
If the interface between the two materials represents the boundary between air (n ~ 1) and
water (n = 1.33) and if angle of incidence = 45°, using Snell's Law the angle of refraction =
32°.
The equation holds whether light travels from air to water, or water to air.
In general, the light is refracted towards the normal to the boundary on entering the material
with a higher refractive index and is refracted away from the normal on entering the material
with lower refractive index.
In labs, you will be examining refraction and actually determine the refractive index of
various materials.
ISOTROPIC INDICATRIX
To examine how light travels through a mineral, either isotropic or anisotropic, an indicatrix
is used.
INDICATRIX - a 3 dimensional geometric figure on which the index of
refraction for the mineral and the vibration direction for light travelling
through the mineral are related.
Isotropic Indicatrix
Indicatrix is constructed such that the indices of refraction are plotted on lines from the origin that are
parallel to the vibration directions.
It is possible to determine the index of a refraction for a light wave of random orientation travelling
in any direction through the indicatrix.
1. a wave normal, is constructed through the centre of the indicatrix
2. a slice through the indicatrix perpendicular to the wave normal is taken.
3. the wave normal for isotropic minerals is parallel to the direction of propagation of light ray.
4. index of refraction of this light ray is the radius of this slice that is parallel to the vibration
direction of the light.
For isotropic minerals the indicatrix is not needed to tell that the index of refraction is the same in all
directions.
Anisotropic minerals differ from isotropic minerals because:
1. the velocity of light varies depending on direction through the mineral;
2. they show double refraction.
When light enters an anisotropic mineral it is split into two rays of different velocity which vibrate at
right angles to each other.
In anisotropic minerals there are one or two directions, through the mineral, along which
light behaves as though the mineral were isotropic. This direction or these directions are
referred to as the optic axis.
Hexagonal and tetragonal minerals have one optic axis and are optically UNIAXIAL.
Orthorhombic, monoclinic and triclinic minerals have two optic axes and are
optically BIAXIAL.
Calcite Rhomb Displaying Double Refraction
Light travelling through the calcite rhomb is split into two rays which vibrate at right angles
to each other. The two rays and the corresponding images produced by the two rays are
apparent in the above image. The two rays are:
1. Ordinary Ray, labelled omega w, nw = 1.658
2. Extraordinary Ray, labelled epsilon e, ne = 1.486.
Vibration Directions of the Two Rays
The vibration directions for the ordinary and extraordinary rays, the two rays which exit the
calcite rhomb, can be determined using a piece of polarized film. The polarized film has a
single vibration direction and as such only allows light, which has the same vibration
direction as the filter, to pass through the filter to be detected by your eye.
1. Preferred Vibration Direction NS
With the polaroid filter in this orientation only one row of dots is visible within the
area of the calcite rhomb covered by the filter. This row of dots corresponds to the
light ray which has a vibration direction parallel to the filter's preferred or permitted
vibration direction and as such it passes through the filter. The other light ray
represented by the other row of dots, clearly visible on the left, in the calcite rhomb is
completely absorbed by the filter.
2. Preferred Vibration Direction EW
With the polaroid filter in this orientation again only one row of dots is visible, within
the area of the calcite coverd by the filter. This is the other row of dots thatn that
observed in the previous image. The light corresponding to this row has a vibration
direction parallel to the filter's preferred vibration direction.
It is possible to measure the index of refraction for the two rays using the immersion oils, and
one index will be higher than the other.
1. The ray with the lower index is called the fast ray
o recall that n = Vvac/Vmedium
If nFast Ray = 1.486, then VFast Ray = 2.02X1010 m/sec
2. The ray with the higher index is the slow ray
o If nSlow Ray = 1.658, then VSlow Ray = 1.8 1x1010 m/sec
Remember the difference between:
vibration direction - side to side oscillation of the electric vector of the plane light
and
propagation direction - the direction light is travelling.
Electromagnetic theory can be used to explain why light velocity varies with the direction it
travels through an anisotropic mineral.
1. Strength of chemical bonds and atom density are different in different directions for
anisotropic minerals.
2. A light ray will "see" a different electronic arrangement depending on the direction it
takes through the mineral.
3. The electron clouds around each atom vibrate with different resonant frequencies in
different directions.
Velocity of light travelling though an anisotropic mineral is dependant on the interaction
between the vibration direction of the electric vector of the light and the resonant frequency
of the electron clouds. Resulting in the variation in velocity with direction.
Can also use electromagnetic theory to explain why light entering an anisotropic mineral is
split into two rays (fast and slow rays) which vibrate at right angles to each other.
PACKING
As was discussed in the previous section we can use the electromagnetic theory for light to
explain how a light ray is split into two rays (FAST and SLOW) which vibrate at right
angles to each other.
The above image shows a hypothetical anisotropic mineral in which the atoms of the
mineral are:
1. closely packed along the X axis
2. moderately packed along Y axis
3. widely packed along Z axis
The strength of the electric field produced by the electrons around each atom must
therefore be a maximum, intermediate and minimum value along X, Y and Z axes
respectively, as shown in the following image.
With a random wavefront the strength of the electric field, generated by the mineral, must
have a minimum in one direction and a maximum at right angles to that.
Result is that the electronic field strengths within the plane of the wavefront define an ellipse
whose axes are;
1. at 90° to each other,
2. represent maximum and minimum field strengths, and
3. correspond to the vibration directions of the two resulting rays.
The two rays encounter different electric configurations therefore their velocities and indices
of refraction must be different.
There will always be one or two planes through any anisotropic material which show
uniform electron configurations, resulting in the electric field strengths plotting as a circle
rather than an ellipse.
Lines at right angles to this plane or planes are the optic axis (axes) representing the direction
through the mineral along which light propagates without being split, i.e., the anisotropic
mineral behaves as if it were an isotropic mineral.
INTERFERENCE PHENOMENA
the colours for an anisotropic mineral observed in thin section, between crossed polars
are called interference colours and are produced as a consequence of splitting the light
into two rays on passing through the mineral.
RETARDATION
Monochromatic ray, of plane polarized light, upon entering an anisotropic mineral is
split into two rays, the FAST and SLOW rays, which vibrate at right angles to each
other.
Development of Retardation
Due to differences in velocity the slow ray lags behind the fast ray, and the distance
represented by this lagging after both rays have exited the crystal is the retardation - D
The magnitude of the retardation is dependant on the thickness (d) of the mineral and
the differences in the velocity of the slow (Vs) and fast (Vf) rays.
The time it takes the slow ray to pass through the mineral is given by:
during this same interval of time the fast ray has already passed through the mineral
and has travelled an additional distance = retardation.
substituting 1 in 2, yields
rearranging
The relationship (ns - nf) is called birefringence, given Greek symbol lower case d
(delta), represents the difference in the indices of refraction for the slow and fast rays.
In anisotropic minerals one path, along the optic axis, exhibits zero birefringence,
others show maximum birefringence, but most show an intermediate value.
The maximum birefringence is characteristic for each mineral.
Birefringence may also vary depending on the wavelength of the incident light.
INTERFERENCE AT THE UPPER POLAR
Now look at the interference of the fast and slow rays after they have exited the anisotropic
mineral.Fast ray is ahead of the slow ray by some amount = D.Interference phenomena are
produced when the two rays are resolved into the vibration direction of the upper polar.
Interference at the Upper Polar - Case 1
1. Light passing through lower polar, plane polarized, encounters sample and is split into
fast and slow rays.
2. If the retardation of the slow ray = 1 whole wavelength, the two waves are IN PHASE.
3. When the light reaches the upper polar, a component of each ray is resolved into the
vibration direction of the upper polar.
4. Because the two rays are in phase, and at right angles to each other, the resolved
components are in opposite directions and destructively interfere and cancel each other.
5. Result is no light passes the upper polar and the grain appears black.
Interference at the Upper Polar - Case 2
1. If retardation of the slow ray behind the fast ray = ½ a wavelength, the two rays are OUT
OF PHASE, and can be resolved into the vibration direction of the upper polar.
2. Both components are in the same direction, so the light constructively interferes and
passes the upper polar.
MONOCHROMATIC LIGHT
If our sample is wedged shaped, as shown above, instead of flat, the thickness of the
sample and the corresponding retardation will vary along the length of the wedge.
Examination of the wedge under crossed polars, gives an image as shown below, and
reveals:
1. dark areas where retardation is a whole number of wavelengths.
2. light areas where the two rays are out of phase,
3. brightest illumination where the retardation of the two rays is such that they are
exactly ½, 1½, 2½ wavelengths and are out of phase.
The percentage of light transmitted through the upper polarizer is a function of the
wavelength of the incident light and retardation.
If a mineral is placed at 45° to the vibration directions of the polarizers the mineral
yields its brightest illumination and percent transmission (T).
POLYCHROMATIC LIGHT
Polychromatic or White Light consists of light of a variety of wavelengths, with the
corresponding retardation the same for all wavelengths.
Due to different wavelengths, some reach the upper polar in phase and are cancelled, others
are out of phase and are transmitted through the upper polar.
The combination of wavelengths which pass the upper polar produces the interference
colours, which are dependant on the retardation between the fast and slow rays.
Examining the quartz wedge between crossed polars in polychromatic light produces a range
of colours. This colour chart is referred to as the Michel Levy Chart and may be found as
Plate I in Nesse.
At the thin edge of the wedge the thickness and retardation are ~ 0, all of the wavelengths of
light are cancelled at the upper polarizer resulting in a black colour.
With increasing thickness, corresponding to increasing retardation, the interference colour
changes from black to grey to white to yellow to red and then a repeating sequence of
colours from blue to green to yellow to red. The colours get paler, more washed out with
each repetition.
In the above image, the repeating sequence of colours changes from red to blue at
retardations of 550, 1100, and 1650 nm. These boundaries separate the colour sequence into
first, second and third order colours.
Above fourth order, retardation > 2200 nm, the colours are washed out and become creamy
white.
The interference colour produced is dependant on the wavelengths of light which pass the
upper polar and the wavelengths which are cancelled.
The birefringence for a mineral in a thin section can also be determined using the equation
for retardation, which relates thickness and birefringence.
Retardation can be determined by examining the interference colour for the mineral and
recording the wavelength of the retardation corresponding to that colour by reading it
directly off the bottom of Plate I. The thickness of the thin section is ~ 30 µm. With this the
birefringence for the mineral can be determined, using the equation:
See the example below.
This same technique can be used by the thin section technician when she makes a thin
section. By looking at the interference colour she can judge the thickness of the thin
section.
The recognition of the order of the interference colour displayed by a mineral comes
with practice and familiarity with various minerals. In the labs you should become
familar with recognizing interference colours.
EXTINCTION
Now we want to examine other properties of minerals which are useful in the
identification of unknown minerals.
Anisotropic minerals go extinct between crossed polars every 90° of rotation. Extinction
occurs when one vibration direction of a mineral is parallel with the lower polarizer. As a
result no component of the incident light can be resolved into the vibration direction of
the upper polarizer, so all the light which passes through the mineral is absorbed at the
upper polarizer, and the mineral is black.
Upon rotating the stage to the 45° position, a maximum component of both the slow and
fast ray is available to be resolved into the vibration direction of the upper polarizer.
Allowing a maximum amount of light to pass and the mineral appears brightest.
The only change in the interference colours is that they get brighter or dimmer
with rotation, the actual colours do not change.
Many minerals generally form elongate grains and have an easily recognizable cleavage
direction, e.g. biotite, hornblende, plagioclase.
The extinction angle is the angle between the length or cleavage of a mineral and the
minerals vibration directions.
The extinction angles when measured on several grains of the same mineral, in the same thin
section, will be variable. The angle varies because of the orientation of the grains. The maximum
extinction angle recorded is diagnostic for the mineral.
Types of Extinction
1. Parallel Extinction
The mineral grain is extinct when the cleavage or length is aligned with one of the
crosshairs.
The extinction angle (EA) = 0°
e.g.
o orthopyroxene
o biotite
2. Inclined Extinction
The mineral is extinct when the cleavage is at an angle to the crosshairs.
EA > 0°
e.g.
o clinopyroxene
o hornblende
3. Symmetrical Extinction
The mineral grain displays two cleavages or two distinct crystal faces. It is possible to
measure two extinction angles between each cleavage or face and the vibration
directions. If the two angles are equal then Symmetrical extinction exists. EA1 = EA2
e.g.
o amphibole
o calcite
4. No Cleavage
Minerals which are not elongated or do not exhibit a prominent cleavage will still go
extinct every 90° of rotation, but there is no cleavage or elongation direction from which
to measure the extinction angle.
e.g.
o quartz
o olivine
Exceptions to Normal Extinction Patterns
Different portions of the same grain may go extinct at different times, i.e. they have
different extinction angles. This may be caused by chemical zonation or strain.
Chemical zonation
The optical properties of a mineral vary with the chemical composition resulting in
varying extinction directions for a mineral. Such minerals are said to be zoned.
e.g. plagioclase, olivine
LECTURE NOTES
1ST
SEMESTER
UNIT 4
RELIEF
Refractometry involves the determination of the refractive index of minerals, using
the immersion method. This method relys on having immersion oils of known
refractive index and comparing the unknown mineral to the oil.
If the indices of refraction on the oil and mineral are the same light passes through
the oil-mineral boundary un-refracted and the mineral grains do not appear to stand
out.
If noil <> nmineral then the light travelling though the oil-mineral boundary is
refracted and the mineral grain appears to stand out.
RELIEF - the degree to which a mineral grain or grains
appear to stand out from the mounting material, whether it is
an immersion oil, Canada balsam or another mineral.
When examining minerals you can have:
1. Strong relief o mineral stands out strongly from the mounting medium,
o whether the medium is oil, in grain mounts, or other minerals in thin
section,
o for strong relief the indices of the mineral and surrounding medium
differ by greater than 0.12 RI units.
2. Moderate relief o mineral does not strongly stand out, but is still visible,
o indices differ by 0.04 to 0.12 RI units.
3. Low relief o mineral does not stand out from the mounting medium,
o indices differ by or are within 0.04 RI units of each other.
A mineral may exhibit positive or negative relief:
+ve relief - index of refraction for the material is greater than the index of
the oil.
- e.g. garnet 1.76
-ve relief nmin < noil
- e.g. fluorite 1.433
It is useful to know whether the index of the mineral is higher or lower that the oil.
This will be covered in the second lab section - Becke Line and Refractive Index
Determination.
BECKE LINE
In order to determine whether the idex of refraction of a mineral is greater than or
less than the mounting material the Becke Line Method is used
.
BECKE LINE - a band or rim of light visible along the grain
boundary in plane light when the grain mount is slightly out
of focus.
Becke line may lie inside or outside the mineral grain depending on how the
microscope is focused.
To observe the Becke line:
1. use medium or high power,
2. close aperture diagram,
3. for high power flip auxiliary condenser into place.
Increasing the focus by lowering the stage, i.e. increase the distance between the
sample and the objective, the Becke line appears to move into the material with the
higher index of refraction.
The Becke lines observed are interpreted to be produced as a result of the lens
effect and/or internal reflection effect.
LENS EFFECT
Most mineral grains are thinner at their edges than in the middle, i.e. they have a
lens shape and as such they act as a lens.
If nmin > noil the grain acts as a converging lens, concentrating light at the centre of
the grain.
If nmin < noil, grain is a diverging lens, light concentrated in oil.
INTERNAL REFLECTION
This hypothesis to explain why Becke Lines form requires that grain edges be
vertical, which in a normal thin section most grain edges are believed to be more or
less vertical.
With the converging light hitting the vertical grain boundary, the light is either
refracted or internally reflected, depending on angles of incidence and indices of
refraction.
Result of refraction and internal reflection concentrates light into a thin band in the
material of higher refractive index.
If nmin > noil the band of light is concentrated within the grain.
If nmin < noil the band of light is concentrated within the oil.
BECKE LINE MOVEMENT
The direction of movement of the Becke Line is determined by lowering the stage
with the Becke Line always moving into the material with the higher
refractive index. The Becke Line can be considered to form from a cone of light
that extends upwards from the edge of the mineral grain.
Becke line can be considered to represent a cone of light propagating up from the
edges of the mineral.
If nmin < noil, the cone converges above the mineral.
If nmin > noil, the cone diverges above the mineral.
By changing focus the movement of the Becke line can be observed.
If focus is sharp, such that the grain boundaries are clear the Becke line will
coincide with the grain boundary.
Increasing the distance between the sample and objective, i.e. lower stage, light at
the top of the sample is in focus, the Becke line appears:
in the mineral if nmin >noil
or in the oil if nmin << noil
Becke line will always move towards the material of higher RI upon
lowering the stage.
A series of three photographs showing a grain of orthoclase:
1. Photo 1 – The grain is in focus, with the Becke line lying at the grain
boundary.
2. Photo 2 – The stage is raised up, such that the grain boundary is out of
focus, but the Becke line is visible inside the grain.
3. Photo 3 – The stage is lowered, the grain boundary is out of focus, and
the Becke line is visible outside the grain.
When the RI of the mineral and the RI of the mounting material are equal,
the Becke line splits into two lines, a blue line and an orange line. In order to see
the Becke line the microscope is slightly out of focus, the grain appears fuzzy, and
the two Becke lines are visible. The blue line lies outside the grain and the orange
line lies inside the grain. As the stage is raised or lowered the two lines will shift
through the grain boundary to lie inside and outside the grain, respectively.
Index of Refraction in Thin Section
It is not possible to get an accurate determination of the refractive index of a
mineral in thin section, but the RI can be bracket the index for an unknown mineral
by comparison or the unknown mineral with a mineral whoseRI is known.
Comparisons can be made with:
1. epoxy or balsam, material (glue) which holds the sample to the slide n =
1.540
2. Quartz
o nw = 1.544
o ne = 1.553
Becke lines form at mineral-epoxy, mineral-mineral boundaries and are interpreted
just as with grain mounts, they always move into higher RI material when the stage
is lowered.
OPTICS
In Isotropic Materials - the velocity of light is the same in all directions. The
chemical bonds holding the material together are the same in all directions, so that
light passing through the material sees the same electronic environment in all
directions regardless of the direction the light takes through the material.
Isotropic materials of interest include the following isometric minerals:
1. Halite - NaCl
2. Fluorite - Ca F2
3. Garnet X3Y2(SiO4)3, where:
o X = Mg, Mn, Fe2+
, Ca
o Y = Al, Fe3+, Cr
4. Periclase - MgO
If an isometric mineral is deformed or strained then the chemical bonds
holding the mineral together will be effected, some will be stretched, others
will be compressed. The result is that the mineral may appear to be
anisotropic.
COMPILED BY
GDC HANDWARA
LECTURE NOTES
1ST
SEMESTER
UNIT 4
PHYSICAL PROPERTIES OF MINERALS
1. Introduction
The physical characteristics of minerals include traits which are used to identify and describe
mineral species. These traits include color, streak, luster, density, hardness, cleavage, fracture,
tenacity, and crystal habit.
Certain wavelengths of light are reflected by the atoms of a mineral's crystal lattice while
others are absorbed. Those wavelengths of light which are reflected are perceived by the viewer
to possess the property of color. Some minerals derive their color from the presence of a
particular element within the crystal lattice. The presence of such an element can determine
which wavelengths of light are reflected and which are absorbed. This type of coloration in
minerals is termed idiochromatism; different samples of an idiochromatic mineral species will all
display the same color. Other minerals are colored by the presence of certain elements in
mixture. Different samples of such a species may exhibit a range of similar colors. Still other
mineral species may usually be colorless, but may display several different and startling colors
when trace amounts of impurities, or elements which are not an integral part of the crystalline
lattice, are present. Coloration which is caused by the presence of an element foreign to the
crystal lattice, whether in mixture or in trace amounts, is termed allochromatism. Certain
elements are strong pigmenting agents and may lend vivid colors to specimens when they are
present, whether as a part of the crystal lattice, in mixture, or as an impurity. These elements are
termed the chromophores.
Streak is the color which a mineral displays when it has been ground to a fine powder. Trace
amounts of impurities do not tend to affect the streak of a mineral, so this characteristic is usually
more predictable than color. Two different specimens of the same species may be expected to
possess the same streak, whereas they may display different colors.
Minerals are either opaque or transparent. A thin section of an opaque mineral such as a
metal will not transmit light, whereas a thin section of a transparent mineral will. Typically those
minerals which possess metallic bonding are opaque whereas those where ionic bonding is
prevalent are transparent. Relative differences in opacity and transparency are described as
luster. The characteristic of luster provides a qualitative measure of the amount and quality of
light which is reflected from a mineral's exterior surfaces. Luster thus describes how much the
mineral surface 'sparkles'.
The property of density is defined as mass per unit volume. Certain trends exist with respect
to density which may sometimes aid in mineral identification. Native elements are relatively
dense. Minerals whose chemical composition contains heavy metals, or atoms possessing an
atomic number greater than iron (Fe, atomic number 26), are relatively dense. Species which
form at high pressures deep within the earth's crust are in general more dense than minerals
which form at lower pressures and shallower depths. Dark-colored minerals are typically fairly
dense whereas light-colored ones tend to be less dense.
Hardness is defined as the level of difficulty with which a smooth surface of a mineral
specimen may be scratched. Hardness has historically been measured according to the Mohs
scale. Mohs' method relies upon a scratch test to relate the hardness of a mineral specimen to the
hardness of one of a set of reference minerals. Hardness may also be measured according to the
more quantitative but less accessible diamond indentation method.
Cleavage refers to the splitting of a crystal along a smooth plane. A cleavage plane is a plane
of structural weakness along which a mineral is likely to split. The quality of a mineral's
cleavage refers both to the ease with which the mineral cleaves and to the character of the
exposed surface. Not every mineral exhibits cleavage.
Fracture takes place when a mineral sample is split in a direction which does not serve as a
plane of perfect or distinct cleavage. A mineral fractures when it is broken or crushed. Fracture
does not result in the emergence of clearly demarcated planar surfaces; minerals may fracture in
any possible direction.
The characteristic of tenacity describes the physical behavior of a mineral under stress or
deformation. Most minerals are brittle; metals, in contrast, are malleable, ductile, and sectile.
The term crystal habit describes the favored growth pattern of the crystals of a mineral
species. The crystals of particular mineral species sometimes form very distinctive, characteristic
shapes. Crystal habit is also greatly determined by the environmental conditions under which a
crystal develops.
2. Color:
When different wavelengths of visible light are incident upon the eye they are perceived as
being of different colors. Three different varieties of color receptors in the eye correspond to
light possessing wavelengths of approximately 660 nm (red), 500 nm (green), and 420 nm (blue-
violet). The eye then interprets the color of incident light according to which color receptors have
been stimulated. For example, if monochromatic light which stimulated the red and green color
receptors equally and did not affect the blue-violet receptors was detected, then the eye would
interpret this light as possessing a wavelength halfway between those of red and green light. The
eye would therefore register an incident light wave with a wavelength of approximately 580 nm
and the viewer would percieve the incoming light as yellow. Incident polychromatic light which
stimulated the red and green color receptors equally and did not affect the blue-violet ones would
also be interpreted as yellow light, regardless whether or not the incoming light actually
contained a component with a wavelength close to 580 nm. The incident polychromatic light
might possess only a red and a green component of equal intensity; it would nevertheless be
interpreted by the eye as yellow light. The phenomenon called color is thus a description of the
differentiation by the eye between various wavelengths and combinations of wavelengths of
visible light.
When light is incident upon a mineral specimen, some wavelengths are absorbed by the
atoms of the crystal lattice while others are reflected. Those wavelengths which were not
absorbed are reflected off of the mineral's surfaces and enter the eye of the viewer. The color
which is perceived by the viewer depends on the wavelengths of light which are reflected rather
than absorbed by the mineral. The property of color in minerals is thus due to the absorption of
particular wavelengths of light and the reflection of others by the atoms of the crystal lattice.
The color exhibited by certain mineral species may depend upon which crystallographic axis
is transmitting the light. Such species may demonstrate several different colors as light is
transmitted along various different axes. This phenomena of directionally selective absorption is
termed pleochroism.
Idiochromatism and the Chromophores
The color of many mineral species is derived directly from the presence of one or more of the
elements which constitute the crystal lattice. The color of such minerals is a fundamental
property directly related to the chemical composition of the species. Minerals which exhibit this
type of coloration are called idiochromatic minerals. Idiochromatic coloration is a property
possessed by a mineral species as a whole. In such species color can successfully be utilized as a
means of identification.
Ions of certain elements are highly absorptive of selected wavelengths of light. Such
elements are called chromophores; they possess strong pigmenting capabilities. The elements
vanadium (V), chromium (Cr), manganese (Mn), iron (Fe), cobalt (Co), nickel (Ni), and copper
(Cu) are chromophores. A mineral whose chemical formula stipulates the presence of one or
more of these elements may possess a vivid and distinctive color.
Examples of idiochromatic minerals abound. For instance, the copper carbonate malachite is
consistently green; the copper carbonate azurite and the copper silicate chrysocolla are each a
distinctive and predictable blue. Rhodochrosite is always red or pink; samples of sulphur are a
bright, recognizable yellow. Each of these distinctive colors is due to the fact that the chemical
composition which defines the mineral species specifies inclusion of one of the chromophores
within the lattice structure.
Allochromatism
Most minerals which are composed entirely of elements other than the chromophores are
nearly colorless. However, certain specimens are sometimes observed to possess vivid
coloration. Color in such instances is due to the presence of an impurity. If one of the
chromophores is present within a mineral whose chemical formula does not include it, then the
foreign element constitutes an impurity or a defect in the lattice structure. Coloration in minerals
which is due to the presence of a foreign element is termed allochromatism. In such cases the
color of the mineral may differ radically from the nearly colorless shade expected of the species.
Some minerals demonstrate a range of colors due to the presence in mixture of one of the
chromophores. For example, the substitution of a quantity of iron for zinc atoms within the
crystal lattice of sphalerite (ZnS) implements a change from white to yellow in the color of the
mineral. Proportionally larger inclusions of iron will progressively result in a brown and
eventually a black mineral specimen. In such cases the color of the sample is directly
proportional to the amount of the pigmenting element which is present in the crystal lattice.
Not all allochromatism in minerals is due to presence of substantial amounts of a
chromophore in mixture, however. The property of color may sometimes be highly dependent on
the inclusion of trace amounts of impurities. The presence of even a minute quantity of a
chromophore within the crystal lattice can cause a mineral specimen to exhibit vivid color. For
example, trace inclusions of chromium (Cr) in beryl are responsible for the deep green of
emerald, while the purple of amethyst is due to trace amounts of iron (Fe) in quartz and the pink
of rose quartz is due to trace inclusions of titanium (Ti). Samples of the mineral corundum which
include tiny amounts of chromium are deep red, and the gem is then called a ruby, while samples
containing iron or titanium impurities produce blue gems termed sapphire.
Trace amounts of an impurity do not affect the basic chemical composition or the chemical
formula of a mineral, and thus do not affect its classification as a species. Trace amounts of the
various chromophores, however, can cause several samples of a single species to differ radically
in color. (Beryl, corundum, and quartz provide examples of this possibility.) Because it varies so
widely, color is a property which is sometimes of little use in identification. However, the
idiochromatic minerals are consistently of distinctive color. The green of malachite, the blue of
azurite, the pink of rhodocrosite, and the yellow of sulphur are easily recognized and are
therefore quite useful in the identification of these species.
3. Streak
Streak is the color of a mineral substance when it has been ground to a fine powder.
Typically an edge of the sample will be rubbed across a porcelain plate, leaving behind a 'streak'
of finely ground material. The material in a streak sample thus consists of a powder composed of
randomly oriented microscopic crystals rather than a lattice structure containing the uniformly
oriented unit cells which compose a macroscopic crystal.
Although color is a property which may vary widely between two different specimens of the
same mineral, streak generally varies little from sample to sample. The presence of trace
amounts of an impurity may radically affect the property of color in a macroscopic crystal
because each unit cell is aligned within the crystal structure, thereby forming a diffraction
grating. Minute amounts of a strongly absorptive impurity within the structure may highly affect
which wavelengths of light are reflected from this diffraction grating. This change may greatly
modify the absorption of certain wavelengths of incoming light, altering the percieved color of
the specimen. In a streak sample, however, each of the microscopic crystal grains of the sample
is randomly oriented and the presence of an impurity does not greatly affect the absorption of
incoming light. Because it is not typically affected by the presence of an impurity, streak is a
more reliable identification property than is color.
4. Luster
Minerals may be categorized according to whether they are opaque or transparent. A thin
section of an opaque mineral such as a metal will not transmit light, whereas a thin section of a
transparent mineral will. The absorption index of an opaque mineral is high. Light which is
incident upon an opaque mineral such as a metal is unable to propagate through the mineral due
to this high rate of absorption, and will thus be reflected. Opaque minerals typically reflect
between 20% to 50% or more of the light incident upon them. In contrast, most of the light
which is incident upon a transparent mineral passes into and through the mineral; transparent
minerals may reflect as little as 5% of the incident light and as much as 20%. Typically those
minerals which possess metallic bonding are opaque whereas those where ionic bonding is
prevalent are transparent.
Relative differences in opacity and transparency are described as luster. The term luster
refers to the quantity and quality of the light which is reflected from a mineral's exterior surfaces.
Luster provides an assessment of how much the mineral surface 'sparkles'. This quality is
determined by the type of atomic bonds present within the substance. It is related to the indices
of absorption and refraction of the material and the amount of dispersion from the crystal lattice,
as well as the texture of the exposed mineral surface.
Minerals are primarily divided into the two categories of metallic and nonmetallic luster.
Minerals possessing metallic luster are opaque and very reflective, possessing a high absorptive
index. This type of luster indicates the presence of metallic bonding within the crystal lattice of
the material. Examples of minerals which exhibit metallic luster are native copper, gold, and
silver, galena, pyrite, and chalcopyrite. The luster of a mineral which does not quite possess a
metallic luster is termed submetallic; hematite provides an example of submetallic luster.
The property of streak can aid in distinguishing whether a specimen has a metallic or a
nonmetallic luster. Metals tend to be soft, implying that more powdered material may be
obtained from the streak sample of a metal than a nonmetal. Metals are also opaque, transmitting
no light. Minerals which possess a metallic luster therefore tend to exhibit a thick, dense, dark
streak whereas those which possess a nonmetallic luster tend to produce a thinner, less dense
streak which is also lighter in color.
Adjectives such as "vitreous', 'dull', 'pearly', 'greasy', 'silky' or 'adamantine' are frequently
used to describe various types of nonmetallic luster.
Dull or Earthy :
Minerals of dull or earthy luster reflect light very poorly and do not shine. This type of luster is
often seen in minerals which are composed of an aggregate of tiny grains.
Resinous
A surface of resinous luster possesses a sheen resembling that of resin. Such materials have a
refractive index greater than 2.0. Sphalerite (ZnS) demonstrates a resinous luster.
Pearly
Pearly luster appears iridescent, opalescent, or pearly. This is typically exhibited by mineral
surfaces which are parallel to planes of perfect cleavage. Layer silicates such as talc often
demonstrate a pearly luster on cleavage surfaces.
Greasy
A surface which possesses greasy luster appears to be covered with a thin layer of oil. A light-
scattering surface which is slightly rough, such as that of nepheline, may exhibit greasy luster.
Silky
Silky luster occurs when light is reflected off of an aggregate of fine parallel fibers; malachite
and serpentine may both exhibit silky luster.
Vitreous
Vitreous luster occurs in minerals with predominant ionic bonding and resembles the reflective
quality of broken glass. The refractive index of such minerals is 1.5 to 2.0. Many silicates
possess this type of luster; quartz and tourmeline both demonstrate vitreous luster.
Adamantine or brilliant
A brilliant luster such as the sparkling reflection of diamond is known as adamantine. Minerals
of adamantine luster have high refractive indices (1.9-2.6) and are highly dispersive and
translucent. Covalent bonding or the presence of heavy metal atoms or transition elements may
result in adamantine luster.
Metallic lustre
Metallic (or splendant) minerals have the lustre of polished metal, and with ideal surfaces will
work as a reflective surface. Examples include galena,6[6] pyrite[7] and magnetite
Submetallic lustre
Submetallic minerals have similar lustre to metal, but are duller and less reflective. A
submetallic lustre often occurs in near-opaque minerals with very high refractive indices,2such
as sphalerite, cinnabar and cuprite.
Waxy lustre
Jade
Waxy minerals have a lustre resembling wax. Examples include jade[11] and chalcedony.[12]
Optical phenomena
Asterism
Sapphire cabochon
Asterism is the display of a star-shaped luminous area. It is seen in some sapphires and rubies,
where it is caused by impurities of rutile.[12][13] It can also occur in garnet, diopside and spinel.
Aventurescence
Aventurine
Aventurescence (or aventurization) is a reflectance effect like that of glitter. It arises from
minute, preferentially oriented mineral platelets within the material. These platelets are so
numerous that they also influence the material's body colour. In aventurine quartz, chrome-
bearing fuchsite makes for a green stone and various iron oxides make for a red stone.
Chatoyancy
Tiger's eye
Chatoyant minerals display luminous bands, which appear to move as the specimen is rotated.
Such minerals are composed of parallel fibers (or contain fibrous voids or inclusions), which
reflect light into a direction perpendicular to their orientation, thus forming narrow bands of
light. The most famous examples are tiger's eye and cymophane, but the effect may also occur in
other minerals such as aquamarine, moonstone and tourmaline.
Color change .
.Alexandrite
Color change is most commonly found in Alexandrite, a variety of chrysoberyl gemstones.
Other gems also occur in color-change varieties, including (but not limited to) sapphire, garnet,
spinel. Alexandrite displays a color change dependent upon light, along with strong pleochroism.
The gem results from small scale replacement of aluminium by chromium oxide, which is
responsible for alexandrite's characteristic green to red color change. Alexandrite from the Ural
Mountains in Russia is green by daylight and red by incandescent light. Other varieties of
alexandrite may be yellowish or pink in daylight and a columbine or raspberry red by
incandescent light. The optimum or "ideal" color change would be fine emerald green to fine
purplish red, but this is exceedingly rare.
Schiller
Labradorite
Schiller, from German for "twinkle", is the metallic iridescence originating from below the
surface of a stone, that occurs when light is reflected between layers of minerals. It is seen in
moonstone and labradorite and is very similar to adularescence and aventurescence.[14]
5. Density
The property of density is defined as mass per unit volume:
µ = m/V
The geometric structure of the unit cell of a mineral determines the volume which it occupies.
The masses of the atoms which compose the unit cell decree the mass of each cell. The identity
of the atoms which compose the unit cell is specified by the chemical formula of the mineral.
Density is therefore directly related to both the physical structure of the unit cell and the
chemical composition of each species of mineral.
One method of measuring the density of a sample entails the use of one dense liquid and
another miscible liquid of lower density. A solution of the two substances is created in which a
crystal of the mineral in question remains suspended and neither sinks nor floats. The weight of a
known volume of the solution is then measured, and the density of the solution and thus the
density of the crystal are calculated from this information. Bromoform (CHBr3, density 2.9
g/cm3), soluble in acetone; di-iodomethane (CH2I2, density 3.3 g/cm3), soluble in chloroform,
CHCl3; and Clerici's solution (a solution of thallium formate and thallium malonate; density 4.4
g/cm3), soluble in water, are some heavy liquids and their solvents which are commonly used in
this process.
Density has historically been equated by mineralogists with the concept of specific gravity.
Specific gravity is a unitless quantity which is defined as the ratio of the weight of a substance to
the weight of an equal volume of water at a temperature of 4° Celsius. This ratio is equal to the
ratio of the density of the substance to the density of water at 4° Celsius.
G = µ / µwater
Specific gravity has therefore classically been measured by weighing a mineral specimen on a
balance scale while it is submerged first in air and then in water. The difference between the two
measurements is the weight of the volume of water which was displaced by the sample. The
specific gravity of the mineral specimen is thus:
G = mair / [mair - mwater]
Because the density of water at 4° Celsius is 1.00 g/cm3, the density of a mineral in units of
grams per centimeter cubed (g/cm3) is equal to its (unitless) specific gravity.
The field geologist sometimes uses a very rough estimation of the density of a hand-held
sample as a lue to identification. Certain rough trends relating mineral density to various other
factors are sometimes useful. Native elements, which contain only one type of atom and whose
molecular structure is that of cubic or hexagonal closest packing, are relatively dense. Minerals
whose chemical composition contains heavy metals - atoms of greater atomic number then iron
(Fe, atomic number 26) - are more dense than atoms whose chemical composition does not
include such elements. Minerals which formed at the high pressures deep within the earth's crust
are in general more dense than minerals which formed at lower pressures and shallower depths.
A general trend relating color to density is also prevalent; this trend states that dark-colored
minerals are often fairly heavy whereas light-colored ones are frequently relatively light. A
geologist is thus given cause to remark upon a sample which seems to reverse this trend. For
example, graphite is dark colored but of low density (C; 2.23 g/cm3) while barite is light in color
but unexpectedly heavy (BaSo4; 4.5 g/cm3). The noted oddity of unexpectedly high or low
density with respect to color provides the field geologist with a clue as to the identification of
such atypical materials.
6. Hardness
Hardness has traditionally been defined as the level of difficulty with which a smooth surface
of a mineral specimen may be scratched. The hardness of a mineral species is dependent upon
the strength of the bonds which compose its crystal structure. Hardness is a property
characteristic to each mineral species and can be very useful in identification.
Certain trends exist in hardness with respect to mineral class. (For a description of the
various classes of minerals, please refer to the discussion on mineral classification contained in
Section 4.) Native elements are typically soft, although iron (Fe) and platinum (Pt) are relatively
hard and diamond (C) is exceptionally hard. Compounds of heavy metals are soft. Sulphides and
sulpho-salts, with the exception of pyrite, are relatively soft; halides are soft; carbonates and
sulphates are usually soft. Oxides are typically hard while hydroxides are softer. Anhydrous
silicates tend to be hard, while hydrous silicates are softer.
The Mohs Scale
The property of hardness has historically been measured according to the Mohs scale, which
was created in 1824 by the Austrian mineralogist Friedrich Mohs. Mohs based his system for
measuring and describing the hardness of a sample upon the definition of hardness as resistance
to scratching. Mohs' method thus relies upon a scratch test in order to relate the hardness of a
mineral specimen to a number from the Mohs scale.
In order to define his scale, Mohs assembled a set of common reference minerals of varying
hardnesses and labled these in order of increasing hardness from 1 to 10. The reference minerals
of the Mohs scale are as follows:
Talc
Gypsum
Calcite
Fluorite
Apatite
Orthoclase
Quartz
Topaz
Corundum
Diamond
Each reference mineral will scratch a test specimen with a Mohs hardness less than or equal to its
own. Each reference mineral can be scratched by a specimen with a hardness equal to or greater
than its own. If a reference mineral both scratches and can be scratched by a certain test
specimen, then the specimen is assumed to possess a hardness equal to that of the reference
mineral in question.
The set of reference minerals of the Mohs' scale can be supplemented by a few common
household items. A fingernail has a Mohs hardness of 21/2; a copper penny 3, window glass
51/2, and a knife blade approximately 6.
The hardness of an unknown sample can be determined to within 1/2 increment by using the
scratch test. Mineral hardnesses determined by the scratch test should never be given in decimal
form, because the Mohs scale does not provide measurements of such precision.
The hardness of a mineral may vary with direction and crystallographic plane. This effect is
usually small. However, species exist in which the variance in the hardness along different axes
is notable. For example, the mineral kyanite (Al2OSiO4) typically forms elongated crystals. The
Mohs hardness parallel to the length of a kyanite crystal is 5, whereas the Mohs hardness
perpendicular to the length of such a crystal is 7. A second example is provided by the mineral
halite, which is softer parallel to its cleavage planes than it is at a 45° angle to the cleavage
planes.
The Diamond Indentation Method
Investigations more recent than those completed by Mohs have used the diamond indentation
method to quantitatively determine hardness. According to this method, a diamond point is
pushed into a planar mineral surface under the weight of a known load. The diameter of the
indentation thereby produced is then measured under a microscope. The diamond indentation
hardness of a sample is equal to the mass of the load applied divided by the surface area of the
indentation produced. The units in which diamond indentation hardness is recorded are therefore
kilograms per millimeter squared (kg/mm2).
Tests utilizing the diamond indentation method have shown that in order for a point
fashioned from a certain material to scratch a surface the hardness of its constituent material
must be 1.2 times that of the surface. Thus on an ideal hardness scale, each subsequent reference
material would have a hardness of approximately 1.2 times that of the material preceeding it. It
must be noted that the intervals between reference points on the Mohs scale are not, in fact,
equal. The interval between subsequent reference points on the scale increases as the hardness of
the reference materials increases. The skill with which Mohs chose his reference materials
becomes apparent when one notes that each of his samples is approximately 1.6 times the
hardness of the last.
The Mohs scale provides a means of testing hardness which is far more readily available to
amateur geologists than the diamond indentation method. It has therefore remained the standard
scale by which hardness is measured.
7. Cleavage
A cleavage plane is a plane of structural weakness along which a mineral is likely to split
smoothly. Cleavage thus refers to the splitting of a crystal between two parallel atomic planes.
Cleavage is the result of weaker bond strengths or greater lattice spacing across the plane in
question than in other directions within the crystal. Greater lattice spacing tends to accompany
weaker bond strength across a plane, because such bonds are unable to maintain a close
interatomic spacing.
Both the positioning of crystal faces in a mineral and the property of cleavage are derived
from the crystalline structure of the species. However, despite the fact that every mineral belongs
to a specified crystal system, not every mineral exhibits cleavage. A mineral such as quartz may
demonstrate beautiful, well-developed crystals and yet possess no distinct planes of cleavage.
Cleavage planes, if they exist, are always parallel to a potential crystal face. However, such
planes are not necessarily parallel to the faces which the crystal actually displays. Fluorite, for
example, has octahedral cleavage yet forms cubic crystals. Nonetheless, the property of cleavage,
if it is present, can offer important information about the symmetry and inner structure of a
crystal.
The quality of a mineral's cleavage refers to both the ease with which the mineral cleaves and
to the character of the exposed cleavage surface. The quality of a sample's cleavage is typically
described by terms such as 'eminent,' 'perfect,' 'distinct,' 'difficult,' 'imperfect,' or 'indistinct.'
'Eminent' cleavage describes the case in which cleavage always occurs readily and is in fact
difficult to prevent from occurring. The mineral mica, for example, cleaves readily into thin, flat
sheets. A mineral which demonstrates 'perfect' cleavage breaks easily, exposing continuous, flat
surfaces which reflect light. Fluorite, calcite, and barite are minerals whose cleavage is perfect.
'Distinct' cleavage implies that cleavage surfaces are present although they may be marred by
fractures or imperfections. 'Difficult' or 'indistinct' cleavage produces surfaces which are neither
smooth nor regular; samples possessing such cleavage tend to fracture rather than split.
Cleavage may be determined by the examination of surfaces which have actually broken. It
may also be determined by inspection of the interlacing systems of cracks which permeate the
structure of certain specimens. These systems of cracks are beautifully apparent within
transparent crystals such as fluorite or calcite.
8. Fracture
A mineral fractures when it is broken or crushed. Fracture takes place when a mineral sample
is split in a direction which does not serve as a plane of perfect or distinct cleavage. In other
words, fracture takes place along a plane possessing difficult, indistinct, or nonexistant cleavage.
The difference between fracture and indistinct cleavage is not clearly delineated.
Unlike perfect or distinct cleavage, fracture does not result in the emergence of clearly
demarcated planar surfaces which run parallel to possible crystal faces. Fracture is
nondirectional: minerals which do not possess distinct cleavage may fracture in any possible
direction.
Fractured surfaces may in some minerals possess a characteristic appearance which can aid
in identification. Examples of distinctive types of fracture are 'conchoidal,' 'irregular,' and
'hackly' fracture.
Conchoidal
Conchoidal fracture results in a series of smoothly curved concentric rings about the stressed
point, generating a shell-like appearance. The familiar ripples of a broken glass bottle
demonstrate this type of fracture. Quartz and olivine are two mineral species which possess
conchoidal fracture.
Irregular
Irregular or uneven fracture results in a rough, rugged surface.
Hackly
The term 'hackly' describes a fractured surface with multiple small, sharp and jagged
irregularities.
9. Tenacity
The property of tenacity describes the behavior of a mineral under deformation. It describes
the physical reaction of a mineral to externally applied stresses such as crushing, cutting,
bending, and striking forces. Adjectives used to characterize various types of mineral tenacity
include 'brittle,' 'flexible,' 'elastic,' 'malleable,' 'ductile,' and 'sectile'.
Brittle
Most mineral species are brittle, and will crumble or fracture under pressure or upon the
application of a blow. Such materials break or powder easily.
Flexible
A mineral which is flexible rather than brittle will flex as opposed to breaking under the
application of stress. However, a mineral which is merely flexible and not also elastic will be
unable to return to its original shape when the stress is removed. Flakes of molybdenite and
scales of talc are two substances which are flexible but inelastic.
Elastic
An elastic mineral will deform under external stress but will resume its original shape after the
stress is removed. If it is bent, it will flex, but will return to its previous position when the stress
dissappears. The mineral called mica is both flexible and elastic.
Malleable
Native metals such as copper, silver, and gold are easily flattened with a hammer. This type of
tenacity is termed malleable. Metallic-bonded minerals tend to be malleable, and may be
pounded out into thin, flat sheets.
Ductile
Some malleable materials are also ductile, and may be drawn out into a thin wire without
crumbling.
Sectile
Some minerals may be sliced into smooth sheets with a knife, although these may possibly still
crumble under a blow from a hammer. Materials possessing this rare type of tenacity are called
sectile minerals. The species chlorargyrite (AgCl) offers an example of a sectile mineral.
10. Crystal Habit
The term crystal habit describes the favored growth pattern of the crystals of a mineral
species, whether individually or in aggregate. It may bear little relation to the form of a single,
perfect crystal of the same mineral, which would be classified according to crystal system. Subtle
evidence of the crystal system to which a mineral species belongs is, however, frequently
observed in the habit of the crystals which a specimen displays.
The terminology used to describe crystal habit is not intended to replace the precise
nomenclature of crystallography. Instead, it is intended as a supplement to this system.
Discussions of crystal habit are more descriptive than precise; for this reason the terminology is
suited to the discussion of mineral samples discovered in the field. Naturally formed specimens
are rarely quantitatively perfect.
The crystals of particular minerals species sometimes form very distinctive, characteristic
shapes. Crystal habit is thus often useful in identification.
Although each mineral species typically forms according to a few preferred shapes, crystal
habit is largely determined by the environmental conditions under which a crystal develops. For
example, aqueous solutions near or surrounding a crystal contain the elemental substances which
it needs to continue growth. The direction from which a growing crystal may obtain such
solutions is a factor which will affect its eventual shape. Higher environmental temperatures
during formation increase ion mobility and aid in crystal formation; the rate at which the
environment cools determines how much time a mineral is allowed to form large crystals. The
amount of space available for a crystal to fill affects its final shape and size. Surface energy
relations are also quite important to the direction of crystal growth; this process is not yet fully
understood.
Adjectives used to describe the habit of individual crystals are 'equant,' 'prismatic,' and
'tabular.' Aggregates of crystals may also be termed equant or prismatic, while aggregates of thin,
flat, tabular crystals may be 'bladed.' Thin sheets, flakes or scales are termed 'foliated,'
'micaceous,' and, if feathery or delicate, 'lamellar' or 'plumose.' Crystal aggregates resembling
long, slender needles, hair, or thread are termed 'acicular,' filiform,' 'capillary,' or 'fibrous.' An
aggregate of crystals forming a network or lattice is 'reticulated;' one composed of branches
which radiate starlike from a central point is 'stellated' while a branching and treelike mineral
growth is 'dendritic.' 'Colloform' crystal habits termed 'botryoidal,' 'mamillary,' and 'reniform'
display spherical, bulbous or globular lumps. Smaller spherical forms are of 'pisolitic' or 'oolitic'
habit; ovoid clusters or formations are 'amygdaloidal.' Tapered, column-like formations are
'stalactitic' or 'columnar' while concentrically banded formations are of 'concretionary' habit.
Minerals whose flat crystal faces are covered with shallow, parallel grooves are 'striated;' a fine
furry layer of crystals growing over a massive lump constitutes a formation of 'drusy' habit.
Following is a list of descriptive terms which are applied when discussing crystal habit.
Equant
A crystal which is equant or equidimensional possesses approximately the same side length in
every direction. Crystals of garnet are often of equant habit.
Prismatic
A prismatic crystal is elongated in one direction like a prism. The mineral tourmaline often forms
crystals of such habit.
Tabular
Tabular crystals appear tabular or platelike in shape.
Bladed
A specimen displaying bladed habit possesses a collection of elongated, flat crystals suggestive
of knife blades. Gypsum often displays crystals of bladed habit.
Foliated
Crystals of foliated habit are separable into leafy structures or display leaflike projections. The
word 'foliated' is derived from the Latin term folium, meaning 'leaf.'
Micaceous
Minerals of micaceous habit form as thin, flat sheets or flakes which are easily peeled or split off
the larger mass. Muscovite provides an example of micaceous habit.
Lamellar or lamelliform
Crystals of lamellar habit form thin scales or plates which may resemble gills or lamellae. The
term is derived from the Latin word lamina, meaning 'thin plate.'
Plumose
A mineral specimen of plumose habit displays fine, feathery scales resembling plumes. 'Plumose'
is derived from the Latin term pluma, or 'feather.'
Acicular
The adjective 'acicular' means needlelike in shape. An acicular aggregate of crystals contains
many long, slender crystals which may radiate out like needles or bristles from a common base.
Acicular crystals are typically long and narrow like a pine leaf and seem to possess a sharp point.
The mineral natrolite often exhibits acicular crystals.
Capillary
An aggregate of crystals of capillary habit resembles an intricate network of tubules. Capillary
crystals appear long, slender, and fine, like delicate hairs. The term 'capillary' is derived from the
Latin word capillus, 'hair.'
Filiform
A mineral possessing crystals of filliform habit exhibits many hairlike or threadlike filaments.
"Filiform' is derived from the Latin word filum, 'thread.'
Fibrous
Specimens possessing fibrous habit exhibit clumps of sinewy, stringy, or hairlike fibers.
Reticulated
A mineral specimen of reticulated habit seems to display a lattice, net, or network of small
crystals. The word 'reticulated' is derived from the Latin term rete, or 'net.'
Stellated
A mineral of stellated habit possesses several branches which radiate outwards from the center in
a pattern resembling a star. The word 'stellated' stems from the Latin term stella, or 'star.'
Dendritic
Dendritic crystals form a divergent branching structure reminiscent of an arborescent, organic
growth such as a tree or a dendrite. Native copper sometimes exhibits this habit.
Colloform
Specimens of colloform habit exhibit spherical, rounded, or bulbous shapes. Botryoidal,
reniform, and mammillary habits are subsets of this category.
Botryoidal
The word 'botryoidal' means 'resembling a bunch of grapes,' or globular. Specimens of malachite
frequently provide examples of botryoidal crystals. The Greek word botrus, 'bunch of grapes,'
provides the linguistic root of botryoidal.
Mammillary
Samples possessing mammillary crystal habit display soft, rounded curves.
Reniform
Reniform crystal habit displays the shape of a kidney. The mineral species hematite provides
samples which exemplify both mammillary and reniform habit. 'Reniform' is derived from the
Latin renes, 'kidney.'
Oolitic
Crystals of oolitic habit form small spheres or grains which resemble fish roe. Oolites are often
found in limestones.
Pisolitic
A mineral of pisolitic habit develops round, pea-shaped forms. These are larger and slightly more
uneven than an oolite and are usually composed of calcium carbonate. The word 'pisolitic' is
derived from the Greek term pisos, 'pea.'
Amygdaloidal
A mineral of amygdaloidal crystal habit demonstrates small almond-shaped nodules called
amygdules. The term stems from the Latin word amygdala, or 'almond.'
Stalactitic
Stalactitic or columnar crystal habit refers to the tall, tapered, columlike appearance of an icicle
or a limestone stalactite. Such formations are built up by the dripping of mineral-laden solution.
The minerals calcite and aragonite (CaC03) typically form stalactites. The term is derived from
the Greek word stalaktos, 'dripping.'
Concretionary
A concretion develops when mineral matter is concentrically deposited around a nucleus and
colored and banded layers are build up. Malachite often exhibits such formations.
Striated
Minerals whose crystals are of striated habit display shallow parallel grooves or lines along flat
crystal faces. Pyrite often demonstrates square, striated crystals.
Drusy
A sample exhibiting drusy habit displays a surface covered with a fine furry layer of tiny
crystals.
Massive
Massive or earthy habit describes a large, lumpy mass which has no apparent crystal form. In
such a sample the crystals are too tiny to be observable by the eye and are interlocked and
mingled; the specimen lacks visible crystals.
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LECTURE NOTES
1ST
SEMESTER
UNIT 4
POLARIZATION OF LIGHT
All of this introductory material on light and its behavior brings us to the most critical
aspect of optical mineralogy - that of Polarization of Light.
Light emanating from some source, sun, or a light bulb, vibrates in all directions at
right angles to the direction of propagation and is unpolarized.
In optical mineralogy we need to produce light which vibrates in a single direction
and we need to know the vibration direction of the light ray. These two requirements
can be easily met but polarizing the light coming from the light source, by means of a
polarizing filter.
Three types of polarization are possible.
1. Plane Polarization
2. Circular Polarization
3. Elliptical Polarization
Three Types of Polarized Light
In the petrographic microscope plane polarized light is used. For plane polarized light
the electric vector of the light ray is allowed to vibrate in a single plane, producing a
simple sine wave with a vibration direction lying in the plane of polarization - this is
termed plane light or plane polarized light.
Plane ploarized light may be produced by reflection, selective absorption, double
refraction and scattering.
When light travels through any other medium it is slowed down, to maintain
constant frequency the wavelength of light in the new medium must al
1. Reflection
Unpolarized light strikes a smooth surface, such as a pane of glass, tabletop,
and the reflected light is polarized such that its vibration direction is parallel to
the reflecting surface.
The reflected light is completely polarized only when the angle between the
reflected and the refracted ray = 90°.
2. Selective Absorption
This method is used to produce plane polarized light in microscopes, using
polarized filters.
Some anisotropic materials have the ability to strongly absorb light vibrating in
one direction and transmitting light vibrating at right angles more easily. The
ability to selectively transmit and absorb light is termed pleochroism, seen in
minerals such as tourmaline, biotite, hornblende, (most amphiboles), some
pyroxenes.
Upon entering an anisotropic material, unpolarized light is split into two plane
polarized rays whose vibratioin directions are perpendicular to each other, with
each ray having about half the total light energy.
If anisotropic material is thick enough and strongly pleochroic, one ray is
completely absorbed, the other ray passes through the material to emerge and
retain its polarization.
This method of producing plane polarized light was employed prior to selective
absorption in microscopes. The most common method used was the Nicol
Prism. See page 14 and Figure 1.14 in
3. Scattering
Polarization by scattering, not relevant to optical mineralogy, is responsible for
the blue colour of the sky and the colours observed at sunset.
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