Chapter 10 - Geology and the Gem Minerals

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CHAPTER 10 – GEOLOGY AND THE GEM MINERALS
This chapter is a simple outline of those geological environments which favor the
formation of gem minerals. The earth is a dynamic planetary body continuously changing
through the activities of two cyclic phenomena.
The weather cycle is driven by energy from the sun. It involves circulation of water
within the earth’s atmosphere and hydrosphere and alters rocks at the earth’s surface by the
process of weathering. Sedimentary rocks are formed by the weather cycle. Near surface
weathering of feldspar and other primary minerals yields clay with a byproduct of colloidal silica
which is the source of opal, agate and the chalcedony varieties.
The tectonic cycle is driven by energy from the earth’s interior. It is a sluggish
circulation, over geologic time, within the earth’ mantle which brings heat and light elements to
the earth’s surface. Igneous and metamorphic rocks are formed by the earth’s tectonic cycle and
are the host rock for the formation of most gem minerals.
Structure of the Earth
Like a hard-boiled egg, the earth has a yolk (core), a white (mantle) and a shell (crust)
(Figure 10-1).
The core of the earth is thought to be largely metallic iron and nickel. The outer core
seems to have the low rigidity characteristic of molten metal, and currents within it may be the
source of the earth’s magnetic field.
The mantle of the earth surrounds the core like the white of the egg and represents about
80% of the earth’s volume. It appears to be very heavy rock, rich in magnesium and iron and
relatively poor in silicon, aluminum and the alkali elements. The major mineral of this rock is
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probably olivine (peridot) like that found in stony meteorites and the mantle xenoliths in basalt
flows. Rock within the mantle is not totally rigid but moves slowly by plastic flow, thereby
circulating heat from the core to the crust like the convection currents in boiling water.
The crust of the earth, like the shell of the egg, is thin and rigid and may be broken into
rigid plates which slip about on the earth’s surface in response to the circulating mantle rock
beneath them. Eight large plates and numerous smaller ones make up the earth’s crust (Figure
10-2). Some plates are entirely ocean crust and other plates have floating masses of low-density
continental rock which ride high on the ocean crust. Plate boundaries may divide ocean basins or
split continents but tend to parallel continental margins.
The ocean crust is thin, heavy, mafic and geologically young. It is composed almost
entirely of basalt which is a dark, fine-grained volcanic rock, extruded largely in the ocean
basins where crustal plates separate. Mafic (Mg-Fe) rocks are relatively silica poor (about 50%
SiO2) and tend to be rich in magnesium, iron and calcium. Continental crust is thick, light
weight, felsic and may be youthful to very old. Felsic (feldspar-silica) rocks are silica rich (about
70% SiO2) and tend to be rich in aluminum and the alkali elements sodium and potassium.
Continental crust forms where crustal plates converge and is represented by a vast variety of rock
types, which, summed together, have the composition of granite.
The earth’s crust rests in isostatic equilibrium (Gr. isos = equal, Gr. statos = standing)
on the underlying plastic mantle (Figure 10-3). Where weight is placed on the earth’s crust, it
subsides, and where weight is removed, the crust rises. Where mountains stand elevated, the
continental crust is thick, and light-weight crustal rock is pressed down into the plastic rocks of
the mantle, like an iceberg floating in the sea. As the mountains wear down, their light-weight
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roots float upward causing the mountains to rise, and the processes repeats until the mountains
are eroded completely away, exposing the igneous and metamorphic core of the mountains
(basement rock), and the roots of the mountains are raised to the level of the continental base.
Plate Tectonics:
The shell fragments of our egg paradigm can interface in only three ways. 1- They may
pull apart, exposing the white of the egg. 2- One fragment may slip underneath another fragment.
3- Fragments, with relatively straight edges, may slip past one another without separation or
overlap. If we were to trace the outer edge of any specific shell fragment, we would note that
everywhere the edge is one of the three types, i.e., one might see a “pull apart” edge on one side
of the fragment and a “slip underneath” edge on another side.
Transform plate boundaries form where plates of the earth’s crust slip past one another,
offsetting topographic features, without separation or overlap and are represented by the San
Andreas Fault, in southern California, or the North Anatolian Fault, in northern Turkey. These
large strike-slip faults are characterized by shallow earthquakes but no important gem-making
processes.
Spreading centers form where the earth’s crustal plates pull apart (Figure 10-4). Plate
separation releases pressure on the underlying mantle rock which partially melts* to form the
hot, fluid lavas of the basalt flows which make up the ocean floor. Ocean crust is
Created at spreading centers and, where a spreading center divides a continental mass (e.g., East
African rift valley), a long narrow sea (e.g., Red Sea) heralds the beginning of a broad, basaltic
ocean basin (e.g., Atlantic Ocean) Characteristic of spreading centers are: shallow earthquakes,
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mild fissure-eruption flood basalts (Figure 10-13), and rift valleys (e.g., Jordan Valley) flanked
by highlands (e.g., Judean Wilderness and Jordanian Highlands), which are the result of thermal
expansion of the crust due to heat escape from the mantle. The only gem mineral directly
associated with the spreading centers is peridot, which may be brought up from the mantle rock
as phenocrysts (Gr. phaino = apparent, Gr. krystallos = crystal) in basalt (Figure 10-15B).
*Partial melting implies that some minerals of the rock will melt while others do not. Every crystalline mineral has a
specific melting point, and the more silica rich the mineral the lower its melting temperature. Therefore, the silica
rich minerals melt first, yielding a molten phase (magma) which is more felsic than the original rock and a solid
residue that is more mafic. By this process, felsic elements (Si, Al, Na, K) are brought to the earth’s surface as
continental crust, and the mafic elements (Mg, Fe, Ca) descend into the ultramafic rocks of the earth’s mantle.
Subduction zones form where one crustal plate under-rides another (Figure 10-5). The
under-riding plate is always ocean crust, but the over-riding rock mass may be either ocean crust
or continental crust. The under-riding plate plunges into the mantle where it is partially absorbed
into the mantle rock and partially melts to form the low-temperature, viscous magma of granitic
batholiths (Gr. bathos = deep, Gr. lithos = stone) and violent volcanoes.
Converging crustal plates generate strong compressional forces. If the subduction zone
lies along a continental boundary, these compressional forces fold the off-shore sediments
forming a mountain belt of sedimentary rock parallel to the continental margin (e.g., Andes Mts.
of South America). In the deep core of the mountain system, these sedimentary rocks are
subjected to the high temperature and pressure of deep burial (i.e., metamorphism) and to the
stress forces of compression and the temperatures of magmas generated by the partial melting of
the subducting ocean crust. These felsic, viscous magmas tend to rise through the
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metamorphosed sediments, due to the lower specific gravity of the magma, and may burst
through in the tops of the overlying mountains as violent stratovolcanoes (e.g., volcanoes of the
Cascade Range of Washington and Oregon) (Figure 10-14). More often, the viscous, felsic
magmas cool at depth beneath the mountain range as large granite batholiths, surrounded by
high grade metamorphosed sediments. This complex of granite batholiths and high grade
metamorphic rocks, commonly called “basement complex”, is exposed at the surface by the
uplift and erosion described above, and herein we find many of the minerals we call gems.
Chemistry of the Earth’s Crust:
Ninety-two elements occur naturally in the rocks of the earth’s crust, however, only eight
occur in amounts greater than one weight percent (see Table 1-2). All common minerals are
some combination of these eight elements and all other elements are considered trace elements.
Since oxygen makes up half of the earth’s crust, by weight, most ionic mineral structures are an
arrangement of large oxygen anions with small, interstitial cations holding the oxygen anions
together to form the crystal structure. And, since silicon makes up half of the remaining 50%,
most minerals are silicates. The most common minerals in the earth’s crust are shown in Table
10-1.
We are familiar with quartz as sand grains or pebbles, but the other common minerals are
less familiar to us, because they are not stable at the earth’s surface where they are exposed to
the water and gasses of the earth’s atmosphere (i.e., weathering). Feldspars and the
ferromagnesian minerals (pyroxenes and amphiboles) decompose to clay or mud, with
byproducts which enrich the salinity of the oceans and cement the sand and pebbles to produce
sandstones, conglomerates and other detrital sedimentary rocks. Ferric hydroxide and colloidal
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silica byproducts color the sediments with rust and form veins of chalcedony and opal in the thin
blanket of sedimentary rock which covers 80% of the land mass. Besides the colloidal silicas
(e.g., opal and agate), a few phosphates (e.g., turquoise and variscite) and carbonates (e.g.,
malachite and calcite) form in this sedimentary blanket and qualify as gem minerals.
The great majority of familiar gems are formed in the granites and metamorphic rocks of
the underlying “basement”, exposed in the heart of modern mountain systems or on the low
rolling hills or plains of continental shields, where the deeply eroded roots of ancient mountains
are exposed (Figure 10-5).
Table 10-1 Mineral Abundance in the Earth’s Crust
Feldspar Mineral Group
60%
Alkali Feldspar (20%)
Plagioclase Feldspar (40%)
Amphibole and Pyroxene
17%
Quartz
12%
Micas
4%
Olivine
1%
Magnetite and Ilmenite
1%
95%
The above figures represent the approximate mineral abundance in the earth’s crust but
do not represent the mineral abundance at the earth surface. In the above mineral list, only quartz
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is truly stable at the earth’s surface. Olivine, pyroxene, amphibole and feldspar break down on
exposure to the earth’s atmosphere to yield various varieties of clay. The micas and magnetite,
although moderately stable, will decompose with time to clay and rust.
Igneous Rocks - Plutonic (Intrusive)
Granite and Granite-like Rocks (Figure 10-6) - Most intrusive igneous rocks are
granite formed as a huge, shapeless mass (batholith) of cooling magma at depths of several
kilometers in subduction zones (Figure 10-5). A granite magma is a chemical system of elements
resulting from the melting of select minerals (i.e., those minerals with the lowest melting points)
in the ocean crust being subducted. Granite rock is coarse-grained, because the magma cools
slowly at depth, and consists largely of feldspar varieties (K-feldspar and plagioclase) and
quartz. The granite-like rocks (e.g., granodiorite, quartz monzonite, syenite, etc.) differ only in
the relative proportions of quartz and the feldspar varieties.
A granite magma is a chemical system at a temperature above the melting point of its
minerals. Each mineral crystallizes from the system when the magma cools to its crystallization
temperature (i.e., melting point temperature), and each mineral takes from the magma those
elements required by its composition. The earliest (i.e., highest temperature) minerals to
crystallize are as mafic as the magma composition can allow and are more mafic than the magma
they crystallized from. From a granitic magma, the first major minerals to form might be an
amphibole [e.g., hornblende Ca2(Mg,Fe+2)4(Al,Fe+3)(Si7Al)O22(OH)2] or dark mica, e.g., biotite
[K2(Mg,Fe+2)5Al(Si5Al3)O20(OH)2], which will probably be the only dark (i.e., mafic) mineral in
the granite. Contemporaneous with the dark minerals will be a Ca-rich feldspar (plagioclase)
which will become more Na-rich as the crystallization proceeds. The final minerals in the main
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crystallization sequence of a granitic magma will be orthoclase feldspar [K(Si3Al)O8] and quartz
[SiO2], yielding a solid rock of quartz, two feldspars (K-rich orthoclase and Na-rich plagioclase)
and minor dark minerals. Quartz is the left-over silica, after the other major elements are
exhausted from the magma, and it crystallizes last (i.e., lowest temperature) or melts first from a
granite rock.
Dark “granites”, like gabbro or diorite (Figure 10-7), are much less common than the
light granites described above. They are coarse-grained, like granite, but contain major
percentages of iron-magnesium-rich silicates (mafic) which provide the overall dark color. They
may contain neither quartz nor alkali-feldspar (i.e., K-rich feldspar or Na-rich feldspar) and
consist largely of Ca-rich plagioclase feldspar [Ca(Si2Al2)O8] and pyroxene
[Ca(Mg,Fe+2)(SiO3)2], sometimes with olivine [(Mg, Fe+2)2SiO4]. For the more mafic minerals
to form, we would have to begin with a magma more rich in iron, magnesium and calcium than
the common granitic magmas, and such magmas are less common in subduction zones.
Pegmatites and Hydrothermal Veins - Although pegmatites (Figure 10-8) and
hydrothermal veins are almost an afterthought in the crystallization of granite magma, we
consider them here because they are the host for a great variety of gem minerals.
Only the big eight elements (O, Si, Al, Fe, Ca, Mg, Na, K) are likely to be represented in
the major minerals of a common granite. A typical granitic magma may, however, contain five or
more weight-percent of dissolved water, only a small percentage of which may be forced into the
amphiboles and micas as the hydroxide anion (OH)−. The final magma phase is very silica-rich
and very volatile-rich (i.e., H2O, CO2, SO2, H2S, etc.). Steam and carbon-dioxide are odorless,
but the sulfur gasses we expect to smell in thermal spring areas, where they escape from cooling
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magmas below (Figure 10-9). The final magma fluids are relatively low temperature but highly
fluid, due to their high volatile content, and they contain all the trace elements, incompatible in
the structures of the common minerals of a granite. Burial pressure on the underground magma
chamber may squeeze the interstitial fluids from the cooling granite body, forcing them into the
enclosing country rock as distinct pegmatite dikes or hydrothermal veins (Figure 10-8), or these
volatile-rich fluids may merely saturate the surrounding rock along mineral grain boundaries
(i.e., contact metamorphism Figure 10-25).
A granite pegmatite dike is a potential jewel box made up largely of coarse crystals of
the final minerals to crystallize from a normal granite magma, i.e., alkali feldspar
[(K,Na)(Si3Al)O8], white mica [KAl2(Si3Al)O10(OH)2] and quartz [SiO2]. The mineral crystals
may be very large, due to the high volatile (gas) content, which makes this residual magma
highly fluid. Injected into pre-existing fractures in the enclosing rocks, this highly fluid magma is
frozen against the cold wall rock as a fine-grained border zone of quartz, feldspar and mica. The
intermediate and core zones cool more slowly and may form very large crystals in the presence
of an aqueous fluid which separates from the remaining silicate magma. The intermediate zone
may contain intergrowths of K-feldspar and quartz, in the proportion 72% K-feldspar and 28%
quartz, which resembles cuneiform writing, hence the name “graphic granite”(Figure 10-10).
Volatiles in the fluid may force open cavities in the core zones which are the “jewel boxes” of
the pegmatite.
The vast majority of pegmatites are simple pegmatites containing little more than the
three basic pegmatite minerals, and most pegmatites form at great depths, associated with highgrade metamorphic rocks, where pressures are too high to allow volatiles to expand to form
voids. Gem pegmatites are extremely rare and are complex pegmatites which contain the trace
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elements essential to gem formation and which form at relatively shallow depths where highvolatile fluids separate from final-stage magmas to form “pockets” where gem minerals may
grow, relatively unobstructed, into an open cavity.
The most common trace elements in this phase are the light metals of the first period of
the periodic table (i.e., lithium, beryllium and boron) and the volatile fluorine. Spodumene
[LiAl(SiO3)2], beryl [Be3Al2(SiO3)6], tourmaline [Na(Li,Al)3Al6(SiO3)6(BO3)3(OH,F)4] and
topaz [Al2(SiO4)(F,OH)2] are the most common resulting gem minerals, although many other
pegmatite gem minerals are possible, e.g., chrysoberyl [BeAl2O4], phenakite [Be2SiO4],
euclase [BeAlSiO4(OH)], amblygonite [LiAl(PO4)F], beryllonite [NaBePO4], petalite
[Li(Si4Al)O10], etc.. We may expect pegmatite gem minerals to contain light trace elements in
some combination with the felsic elements Si, Al, Na, K and we might anticipate crystals that
could be faceted to yield quite large gemstones.
A hydrothermal vein (Gr. hydro = water and Gr. therme = heat) forms from the aqueous
solutions which remain after the last magma has crystallized and may grade away from a
pegmatite dike when the alkali feldspar and white mica (muscovite) have mostly crystallized and
little remains but silica and hot water, not greatly above its surface boiling point. The silica may
crystallize as massive quartz, milky with tiny liquid inclusions, or as cryptocrystalline (Gr.
krypto = hidden) quartz (i.e., chalcedony). Fluorine gas asserts itself in fluorite [CaF2] and
cryolite [Na3AlF6] and sulfur gases may deposit sulfur crystals at fumerole vents (Figure 10-9) or
form sulfates like gypsum [CaSO4·2H2O] or barite [BaSO4]. Milky quartz is a common host for
native gold (Figure 7-6d), and a great variety of heavy metallic elements are concentrated in the
final solutions and combine with sulfur to form sulfide ore deposits of mercury, arsenic,
antimony, tin, tungsten, molybdenum, and the common base metals copper, lead and zinc.
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Igneous Rocks - Volcanic (Extrusive)
Vulcanism is a continuous spectrum from gentle basalt flows that flood the landscape to
the violent explosions that blast away mountains, and most existing volcanoes fall near the ends
of the spectrum.
Mild volcanoes (Figure 10-11) are characteristic of plate spreading where basaltic lavas
ooze onto the ocean floor through opening fissures, making vast beds of “pillow” basalt (Figure
10-12), as new ocean crust (e.g., Mid-Atlantic Ridge). Similar vulcanism results over mantle
“hot spots” (e.g., Hawaiian Islands) where basaltic lavas originate in the upper mantle and rise
through the crust to form island chains of giant shield volcanoes, as the ocean plate moves over
the stationary hot spot in the mantle. Where crustal spreading divides a continent, fields of
fissure eruption basalts (Figure 10-13) flood vast continental areas (e.g., Columbia River
Plateau). Mild vulcanism is characterized by:
Hot and fluid basaltic lavas
Nearly 100% lava flows (i.e., little solid ash)
Fissure eruption
Gentle slopes of broad shield volcanoes
Almost continuous eruption
Molten lava in the throat of the volcano and lava lakes in the crater.
Basalt - Basaltic magmas represent the high-temperature melting of iron-magnesium rich
(mafic) silicates, usually where mantle rock is exposed at spreading centers. At spreading
centers, highly fluid mafic magmas usually reach the surface as lavas which crystallize as a dark,
fine-grained rock of olivine [(Mg, Fe+2)2SiO4], pyroxene [Ca(Mg,Fe+2)(SiO3)2], and Ca-rich
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plagioclase [Ca(Si2Al2)O8], i.e., basalt (Figure 10-15). Olivine crystallizes first and may grow to
significant size before being surrounded by fine-grained plagioclase (labradorite) and pyroxene
in the final rock. Olivine (gem name “peridot”) in basalt is usually too small to be of practical
value and forms the green sands of Hawaiian beaches.
The violent volcanoes (Figure 10-14) are the consequence of plate subduction. They are
the islands of oceanic island arcs (e.g., South Pacific islands) or appear in the youthful and
growing mountain ranges along continental margins, inland of deep ocean troughs (e.g., Andes).
Their violence may exceed the energy of many atomic bombs and some have altered the course
of history and world climate (e.g., Vesuvius, Santorini, Krakatoa). Violent vulcanism is
characterized by:
“Cool” and highly viscous rhyolitic lavas
Nearly 100% pyroclastics and ash flows
Central eruption from giant, symmetrical volcanic cones
Steep slopes of layered stratovolcanoes
Intermittent eruption with long periods between eruptions
Solid lava in the throat of the volcano, requiring huge pressure build up for eruption
Andesite (Figure 10-16) is basalt, with more sodium-rich plagioclase (andesine) and no
olivine. It is more felsic and less dark than basalt and commonly green or purple with visible
plagioclase phenocrysts (Gr. pheno = visible and Gr. krystallos = crystal). Andesite is
characteristic of huge stratovolcanoes like those in the Andes Mountains, from which it derives
its name, and is associated with crustal plate subduction
Rhyolite (Figure 10-17) is the volcanic equivalent of granite with the same chemical and
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mineral composition. Rapid cooling at the earth’s surface, however, produces a light-colored,
fine-grained volcanic rock, usually pink or pale lavender, where no individual crystals are
visible, without magnification. Very rapid cooling of the viscous, rhyolitic lavas may yield
uncrystallized volcanic glass, i.e., obsidian (Figure 10-18), which is usually black, owing to very
tiny magnetite (Fe0.Fe2O3) crystals scattered throughout the transparent glass.
As with granite, the final stages in the cooling of rhyolitic lava may yield highly mobile,
low-temperature, pegmatite solutions of water and gasses rich in silica. These mineralizing
solutions and gasses may force fractures and voids within the rhyolite or associated ash beds and,
if endowed with unique pegmatite trace elements, they may form crystals, in fissures and
cavities, of topaz, beryl or other minerals characteristic of pegmatites.
Sedimentary rocks
Sedimentary rocks (Figure 10-19) are stratified, secondary rocks, made from the physical
break down or chemical decay of any pre-existing rock. Although igneous rocks make up most
of the earth’s crust, rocks at the earth’s surface are largely sedimentary and form a thin blanket
over the igneous-metamorphic basement. Although feldspar is by far the most abundant mineral
group in the earth’s crust, it is largely absent in the sedimentary blanket, because feldspars are
chemically unstable under surface conditions. Feldspars react with oxygen, water and carbondioxide in the air to decay to clays (i.e., mud) with byproducts of colloidal silica and soluble
salts.
2Na(AlSi3)O8 + 2H2O + O + CO2  (Al2Si2)O5 (OH)4 + 4SiO2 + Na2CO3
feldspar
kaolinite clay
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colloidal silica
salt
Ferromagnesian minerals (i.e., olivine, pyroxenes, amphiboles) also decompose to other varieties
of clay with similar byproducts plus colloidal ferric-hydroxide. The soluble salt byproducts are
the salts of the oceans and lakes and are largely salts of Na, K, Ca and Mg, which combine with
the gasses CO2, SO2 and Cl2 in the sea to form carbonates, sulfates and chlorides.
Colloidal silica is transported by groundwater and deposited as a gel in the fractures and
pore spaces of clastic sediments as a cementing agent for sand and gravel (Figure 10-20) or
deposited as opal or chalcedony in veins, voids or cavities in any rocktype. Colloidal ferrichydroxide is also carried and deposited by groundwater as limonite or goethite (Fe2O3·n H2O),
which cements and colors detrital sediments and forms “iron” nodules.
Quartz is remarkably stable over a wide range of temperatures and pressures and is stable
at the earth’s surface forming sand or pebbles of quartz which are glued, or cemented, together to
form sandstone or quartz conglomerate. Sand grains are most commonly quartz, and the
cements are most commonly calcite, chalcedony or limonite, but an infinite variety of other
minerals and rocks may act as pebbles, sand grains or glue. Although most gems may be
recovered from modern river gravels or conglomerates (Figure 10-21) or ancient
metaconglomerates, the gems pebbles did not originate in these detrital sediments but were freed,
by weathering, from the source rocks where they did form. Most gem minerals tend to be
chemically stable and to resist the attack of weather.
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Shale is very fine-grained sedimentary rock and is composed of clay minerals, which are
the products of weathering. Shale is normally thinly layered due to the natural orientation of clay
particles which are tiny, mica-like sheets. A wide variety of clay minerals, have a wide range of
composition, however, all clays are essentially aluminum hydroxides. Unbalanced electrical
charges at the edges of tiny clay layers attract available cations of many other elements becoming
the raw materials of metamorphic gem minerals. Metamorphosed shale, always aluminum rich,
may contain a greater variety of minerals than any other geological environment, and many gem
minerals form in metamorphosed shale, e.g., jade, garnet, iolite, spinel, etc.
Chemical sediments are deposited chemically or organically from oceans, lakes, streams or
groundwater. Calcium-carbonate (CaCO3), as calcite or aragonite, is by far the most common
chemical sediment and has been deposited in huge amounts in shallow seas throughout geologic
time. Calcite is the principal constituent of all limestones and marbles which occur in an infinite
variety of colors and textures, due to a wide range of impurities. In addition to the iron oxides (i.e.,
rust) which provide the red, yellow and brown in most rock varieties, limestones may contain large
amounts of graphite from decayed marine life, and gray to black limestone is very common. High
temperature metamorphism removes graphite as carbon-dioxide (CO2), and the resulting marble
tends to be white masses of coarse crystalline calcite. Residual graphite may remain as black or
gray streaks or patterns in the marble.
Most marine organisms take calcium-carbonate from seawater to form protective shells or
armor, and many limestones are composed almost entirely of calcium-carbonate shells or the
discarded habitats or skeletons of a wide variety of animals that lived in the sea. Coral is the
calcium-carbonate “apartment house” for myriads of tiny marine organisms called polyps. Oysters,
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mussels, abalone and other shell-fish, which line their shells with calcium-carbonate nacre, yield
pearls which are alternating layers of calcium-carbonate and a natural plastic, called conchiolin.
Calcium-carbonate is not very soluble and is precipitated from ocean waters, as fast as it is
produced by the weathering of plagioclase feldspars, pyroxenes or amphiboles and carried to the
sea by rivers and streams. Although the ocean waters are saturated with calcium-carbonate, it is a
minor constituent of the soluble salts in sea water. Chalk is fine-grained, rather pure calciumcarbonate deposited as a fine calcite mud or ooze.
Calcium-carbonate is also the principal load of groundwater. It is carried in a soluble form
as calcium-bicarbonate [CaH2(CO3)2] which, on exposure to the air, releases CO2 gas and water to
become insoluble calcium-carbonate (CaCO3). It is calcium-carbonate that makes culinary water
“hard” and deposits in pipes and teakettles and white rings on windows, “crystal” goblets and
recently washed cars.
Calcium-carbonate deposited from ground waters is called travertine (Figure 10-22). It
may be deposited from hot spring waters as a highly porous, spongy and fragile rock called tufa, or
sinter, or it may be deposited by cold, descending waters in tiny fractures or large veins or in caves
as solid, layered, onyx-like rock. This solid, cave travertine is almost always colored yellow to
brown by minor iron impurities and is often misnamed “alabaster” when used to make vases, table
tops or fireplaces.
Hot spring water is often groundwater, heated by hydrothermal water or gasses (e.g., CO3),
which may deposit calcite, usually as fine crystals, in association with various ore minerals. In such
an environment, a whole family of metal carbonates may form, e.g., siderite (FeCO3),
rhodochrosite (MnCO3), smithsonite (ZnCO3), magnesite (MgCO3) and dolomite [CaMg(CO3)2].
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This is an isomorphous mineral family ( i.e., same trigonal crystal structure, but different
compositions) with similar properties. Rhodochrosite, containing manganese as an essential
constituent, is always rose pink and smithsonite, which usually contains enough copper ions to
color it pale blue-green, are both sometimes cut and polished as ornamental stone. Aragonite is an
orthorhombic polymorph of calcite (i.e., same composition but different crystal structure) and will
revert slowly to calcite under normal surface conditions of temperature and pressure. With slightly
higher pressure or lower temperature, aragonite becomes the stable, polymorph, in some cave
deposits and the calcium-carbonate layers in pearls or shells. Strontianite (SrCO3), witherite
(BaCO3) and cerussite (PbCO3) have large cations and the same orthorhombic structure as
aragonite. i.e., they are isomorphs of aragonite.
Metamorphic Rocks
Metamorphic rocks (Figure 10-23) are secondary rocks derived from any previously
existing rock which is subjected to elevated temperatures and/or pressures sufficient to alter its
texture, fabric, or mineral combinations but insufficient to bring about melting and its return to a
magma. Metamorphism is normally associated with subduction zones (Figure 10-5) where surface
rocks of all kinds are carried into the mantle as part of the subducting plate. In the zone of mountain
building, where sedimentary formations are deformed by compressional forces, temperature
increases with depth, and pressure is directional (sheer pressure) near the surface grading to
uniform pressure, or uniform compression, at depth.
Any chemical system (i.e., combination of elements) will form one combination of
minerals, under one specific range of temperature and pressure, and perhaps another combination
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of minerals, under another set of temperature-pressure conditions. This is metamorphism, and many
gem minerals are formed in this way. Several types of metamorphism are defined by a range of
temperature and a range of pressure, and metamorphic “facies” are defined by the combination of
minerals which the temperature, pressure and chemistry produce.
At shallow depth, where the under riding plate, carrying ocean sediment, goes beneath the
overriding plate, temperature is quite low but shearing pressure is high (blueschist
metamorphism). Under these conditions, ocean salts rich in sodium, may be added to the mix of
ocean clay sediments and “shear” minerals are formed, like jadeite [NaAl(SiO3)2] and glaucophane
[Na2Mg3Al2(Si4O11)2(OH)2].
In the main mountain building system, where sedimentary rocks are highly deformed and
masses of granitic magmas are generated, temperatures and pressures are low near the surface
becoming greater at depth (dynamothermal metamorphism). At moderate depths, shearing
pressures and moderate temperatures may alter a stable mineral combination in ocean sediments
(e.g., sand, clay, iron oxides) to form a different set of minerals (e.g., mica, chlorite, garnet) and the
oriented texture of a schist (Figure 10-24A) without changing the total chemical composition of the
overall rock. At greater depth (i.e., higher temperature and more uniform pressure), the mica and
chlorite may become unstable and change to feldspar and amphibole or pyroxene, and the oriented,
flaky schistose fabric of the rock changes to the granular, deformed layering of a gneiss (Figure 1024B) or, at even greater depth, the coarse granular texture of a granite-like rock. Almandite garnet
[Fe3Al2(SiO4)3], staurolite [Fe2Al9O6(SiO4)4(OH)2], iolite [Mg2Al3(Si5Al)O18] and kyanite
[Al2SiO5] are common gem minerals of schists and gneisses.
Where molten rock intrudes solid rock at shallow depth (contact metamorphism), the
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temperature is high and pressure low, heat from the molten rock, conducted to the solid wallrock,
may alter its texture and recombine its elements to new mineral combinations, without altering its
total chemistry (Figure 10-25). Gem minerals like grossular garnet [Ca3Al2(SiO4)3], andalusite
[Al2SiO5} and idocrase [Ca10(Mg,Fe)2Al4(Si2O7)2(SiO5)5(OH)4] are metamorphic contact minerals.
Notice that all the mineral formulas for metamorphic minerals in the paragraphs above contain only
the eight most abundant elements, plus hydrogen from water. These metamorphic minerals reflect
the composition of the original sedimentary or igneous rocks from which they formed, without
addition or subtraction of foreign elements.
When volatile-rich molten rock, perhaps a pegmatite magma, intrudes a solid rock,
however, the “country rock” may receive more than heat from the magma. It may be infiltrated
with gasses and fluids, from the magma, containing trace elements to be added to minerals of the
new mineral combinations. This form of contact metamorphism is called metasomatism and may
yield gem minerals like tourmaline [Na(Mg, Fe, Li, Al)3Al6(SiO3)6(BO3)3(OH,F)4] and lazurite
[(Na,Ca)8Al6Si6O24(SO4,S,Cl)2]. In metasomatic minerals we commonly see the addition of gasses
(SO2, S, Cl, F) and trace elements (Li, B), from cooling magmas, and residual salt brines (NaCl)
from marine sediments.
Metamorphism of shale, composed mostly of clays [e.g., Al2Si2O5)(OH)4 ], yields
metamorphic minerals rich in aluminum. Carbonate sediments, composed mostly of calcite and
dolomite [i.e., CaCO3 and CaMg(CO3)2], yield metamorphic minerals rich in calcium and
magnesium. Basaltic rocks, rich in mafic minerals [e.g., pyroxene CaMg(SiO3)2 and amphibole
Ca5(Mg,Fe,Al)5(Si4O11)2(OH)2], yield metamorphic minerals rich in magnesium, iron and calcium.
Limestone (CaCO3) containing an aluminum-hydroxide (e.g., bauxite) may yield corundum
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(Al2O3), as a metamorphic mineral, and if the limestone is magnesian calcite, spinel (MgAl2O4)
may result. Corundum and spinel must form in a silica-poor environment
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