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The
Formation
of Minerals
THE
FORMATION
OF
MINERALS
Vivien Gornitz
New York City
March 1999
Copyright © 1999 New York Mineralogical Club, Inc.
Dr. Vivien Gornitz, left, author of “The Formation of Minerals”, being
presented with the 1998 First Place EFMLS “Each One Teach One Award”.
November 7, 1998, Stamford Connecticut. EFMLS Convention Awards
Banquet. Photo by Anna Schumate.
ii
PREFACE
This booklet forms a companion to A n Introduction to Minerals, V ol. I. It is an outgrowth of topics
discussed in a series of Study Group lectures of the New York Mineralogical Club, Inc., held at the
American Museum of Natural History, between 1997 and 1998. Basic geological concepts are
introduced to explain the processes by which minerals form in nature. This background information
gives mineral collectors and rockhounds a better understanding of why minerals are found in particular
geological settings and why certain groups of minerals tend to occur together. Factors which promote
the development of large, nearly perfect crystals are also examined. This knowledge enables mineral
enthusiasts to appreciate the special, even rare circumstances that produce well-crystallized, or gemmy,
specimens. The material in this booklet is intended to bridge the gap between colorfully-illustrated, but
sparsely-documented field guides and the more technical literature.
Vivien Gornitz, Ph.D.
New York City
March, 1999
iii
ACKNOWLEDGMENTS
The author expresses her sincerest appreciation to Mitchell Portnoy for his tireless efforts,
professionalism, and originality in arranging the layout, graphic design, and printing this booklet and
for loaning several specimens. Thanks are also extended to Karen Rice for loaning some gem
crystals and for her skill and patience in photographing the mineral specimens.
Illustration Credits
Mineral specimens are from the author's collection, except for Figures 2.5, 2.7(tourmaline), 3.3 (gold and
diamonds), 6.2, 6.11, and 9.1 (Karen Rice), Figures 2.4 (olivine), 2.7 (columbite), 3.3 (sapphires), 3.4, 6.6
(Mitch Portnoy), and Figures 4.4, 6.3 (Will Heiermann).
Figures courtesy of the Department of Library Services, American Museum of Natural History:
3.5 (AMNH trans. #2276), 5A (AMNH trans. #3628), 8.4 (AMNH trans. #526), and 9.2 (AMNH trans.
#425).
Figures 1.1, 1.5, 2.1, 2.10, 4.9, 6.1, and 6.5 from G. Robinson, Mine rals: An I llustrate d E xp lo ratio n o f th e
Dynam ic Wo rld o f Mine rals and T h e ir Pro p e rtie s , 1994. Simon & Schuster, Inc. ©1994. J. Scovil.
Reprinted with permission.
Figures 1.6, 10.9, 10.10, 10.11, 10.12, 10.13 and 10.14, N ASA.
Figure 2.6 reprinted with permission from M.B. Kirkley, J.J. Gurney, and A.A. Levinson, “Age, Origin, and
Emplacement of Diamonds: Scientific Advances in the Last Decade”, in Ge m s and Ge m o lo g y, Vol. 27, No.
1, 1991, p. 22.
Figures 2.8 and 5.3, reprinted from J. Sinkankas, F ie ld Co lle c ting Ge m s and Mine rals , 1988, with
permission of the publisher, GeoScience Press, Inc., Tucson, AZ © 1961 and 1970 by John Sinkankas..
Figures 2.11, 3.1, 3.2, 3.6, 6.4, and Table 6.1 from J.W., Barnes, O re s and Mine rals: Intro d uc ing
E c o no m ic Ge o lo g y, 1988, © John Wiley & Sons, Limited. Reproduced with permission.
Figure 5.5 reprinted with permission from E. Fritsch and G.R. Rossman, “An Update on Color in Gems.
Part 3”, Ge m s and Ge m o lo g y, Summer, 1988, © Gemological Institute of America.
Figures 6.9, 6.15, 7.6, and 9.3, reprinted with permission from N ature , © 1993, 1994, 1997, and 1998,
Macmillan Magazines, Ltd.
Figures 6.12 and 6.13 from S. Graeser, in R o c k s and Mine rals , Jan.-Feb., 1998.
Figure 6.16 from A. Audetat et al., in Sc ie nc e , v., 279, 1998.
Figures 7.3 and 7.5 reprinted with permission from R.M. Haymon and K.C. Macdonald, “The Geology of
Deep-Sea Hot Springs”, in Am e ric an Sc ie ntist, v. 73, p. 441-445, 1985.
Figure 7.4 from P. Rona, Mineral Deposits from Sea-Floor Hot Springs, in Sc ie ntific Am e ric an ,
Jan. 1986, p.84 - 92.
Figure 8.3 from W. Cummings in R o c k s and Mine rals , v., 60, 1985.
Figures 9A and 9B, 9.4, and 9.5 from Ge o m ic ro b io lo g y: Inte rac tio ns b e twe e n Mic ro b e s and Mine rals ,
eds. J.F. Banfield and K.H.Nealson, © 1998, Mineralogical Society of America.
Figure 10.3 from J. Wood, Me te o rite s and th e O rig in o f Plane ts , ©1968, McGraw-Hill Inc. Reproduced
with permission.
Figure 10.7 from J. Wood, in J.K. Beatty and A. Chaikin, eds., 1990. T h e N e w So lar Syste m , Sky Publishing
Corp. and Cambridge University Press.
Figure 10.15 from U. Ott, in N ature , © 1993, Macmillan Magazines, Ltd. Reprinted with permission.
iv
TABLE OF CONTENTS
Chapter 1.
Minerals, the Rock Cycle, and Beyond.. . . . . . . . . . . . . . . . . . . . . . . . . . 1
Chapter 2.
Minerals from the Melt. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9
Chapter 3.
Minerals from Sedimentary Deposits. . . . . . . . . . . . . . . . . . . . . . . . . . . 19
Chapter 4.
Metamorphic Minerals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25
Chapter 5.
Minerals from Cool Solutions.. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 31
Chapter 6.
Hydrothermal Minerals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37
Chapter 7.
Black Smokers — Minerals from the Deep Sea. . . . . . . . . . . . . . . . . . 49
Chapter 8.
Mineral Succession. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59
Chapter 9.
Minerals and Organisms. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 65
Chapter 10. Minerals Beyond the Earth. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 77
Appendix
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 93
About the New York Mineralogical Club, Inc.. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 95
About the Author. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 96
v
List of Figures
Chapter 1
Figure 1.1.
Figure 1.2.
Figure 1.3.
Figure 1.4.
Figure 1.5.
Figure 1.6.
The geologic rock cycle
Mineral composition of common igneous rock types
Diagram showing the main elements of plate tectonics
Diagram illustrating two types of plate collisions
Sources of hydrothermal solutions
Mars—view of surface from Pathfinder
Chapter 2
Figure 2.1.
Figure 2.2.
Figure 2.3.
Figure 2.4.
Figure 2.5.
Figure 2.6.
Figure 2.7.
Figure 2.8.
Figure 2.9.
Figure 2.10.
Figure 2.11.
Igneous processes
Crystallization sequence of minerals from a silicate melt
Topaz from the Thomas Range, Utah
Olivine crystal, northwest Pakistan
Diamond crystals
Idealized diagram of a kimberlite pipe
Pegmatite minerals: tourmaline, aquamarine, and columbite
Diagram of a gem-bearing crystal pocket from the Himalaya (San Diego) pegmatite
Amazonite from the Pikes Peak district, Colorado
Serandite and analcime, Mont Saint-Hilaire, Quebec
Schematic drawing of a porphyry copper deposit
Chapter 3
Figure 3.1.
Figure 3.2.
Figure 3.3.
Figure 3.4.
Figure 3.5.
Figure 3.6.
Cross-section through an alluvial deposit
Concentration of heavy minerals in a river at points where the stream velocity drops
Placer gems: diamond, sapphire, and gold
Bauxite—a major ore of aluminum
Halite, Krakow Saltworks, Poland
Worldwide distribution of banded iron formations (BIFs)
Chapter 4
Figure 4.1.
Figure 4.2.
Figure 4.3.
Figure 4.4.
Figure 4.5.
Figure 4.6.
Figure 4.7.
Figure 4.8.
Figure 4.9.
Approximate temperature and pressure ranges of various types of metamorphic
rocks and minerals
Kyanite, Brazil
The Kunz Garnet, almandine crystal found in New York City in 1885
Ruby from metamorphic marble, Hunza, Pakistan
Lazurite on calcite, Afghanistan
Nephrite
Jadeite
Emerald from schist, Brazil
Schematic illustration of contact metamorphism
vi
Chapter 5
Figure 5.1.
Figure 5.2.
Figure 5.3.
Figure 5.4.
Figure 5.5.
Figure 5.6.
Figure 5A
Figure 5B
Diagram illustrating formation of a cave
Botryoidal malachite, Congo (formerly Zaire)
Secondary alteration of a hydrothermal vein, showing various zones
Turquoise from Nevada and Arizona
Diffraction of light in precious opal
Boulder opal, Australia
Azurite and malachite, Bisbee, Arizona
Chrysocolla, Arizona
Chapter 6
Figure 6.1.
Figure 6.2.
Figure 6.3.
Figure 6.4.
Figure 6.5.
Figure 6.6.
Figure 6.7.
Figure 6.8.
Figure 6.9.
Figure 6.10
Figure 6.11.
Figure 6.12.
Figure 6.13.
Figure 6.14.
Figure 6.15.
Figure 6.16.
Crystallization within open cavities in lava flows
Natrolite, Paterson, New Jersey
Native copper from the Keweenaw Peninsula, Michigan
Migration of solutions from a basin into carbonate rocks
Deposition of minerals in a MVT deposit
Marcasite from the Viburnum Trend, Missouri
Fluorite and calcite, Cave-in-Rock, Illinois
Schematic diagram showing how Colombian emeralds were formed
Cross-section of an emerald vein at Muzo, Colombia
Emerald crystals with pyrite and calcite, Muzo, Colombia
Three-phase inclusion in emerald from Colombia
Geologic cross-section of the Binntal area, Switzerland, showing stacking of nappes
in the central Alps
Formation of Alpine clefts due to stresses on rocks
Gwindel quartz, Swiss Alps
The relation of hydrothermal veins to igneous activity
Longitudinal section through a quartz crystal from Yankee Lode tin deposit,
Australia
Chapter 7
Figure 7.1.
Figure 7.2.
Figure 7.3.
Figure 7.4.
Figure 7.5.
Figure 7.6.
Figure 7.7.
Figure 7.8.
Black smokers discharging high-temperature fluids and metal sulfides at vents on the
Juan de Fuca Ridge
Creatures dwelling at the hydrothermal vents
Schematic diagram of mid-ocean ridge where magma upwells from the mantle and
pushes the earth's plates apart
Submarine hydrothermal circulation
Mineral zoning at a black smoker chimney
Cross-sectional views of mineralized zones at the Juan de Fuca hydrothermal deposit
Side (a) and (b) back views of Rimicaris exoculata—the “eyeless” shrimp
The Pompeii worm, Alvinella pompejana, 6 cm (2.4 in) long
vii
Chapter 8
Figure 8.1.
Figure 8.2.
Figure 8.3.
Figure 8.4.
Figure 8.5.
Figure 8.6.
Figure 8.7.
Figure 8.8.
Figure 8.9.
Quartz on tetrahedrite over pyrite, Pachapaqui, Peru
Calcite on prehnite, Millington quarry, New Jersey
Paragenetic sequence at Millington quarry, New Jersey
Porphyry
Zoning in tourmaline (liddicoatite, Madagascar)
Pseudomorph of turquoise after apatite, Bacuachic, Sonora
Pseudomorph of native copper after aragonite, Coro Coro, La Paz Dept., Bolivia
Petrified wood, Arizona
Pyritized ammonite
Chapter 9
Figure 9.1.
Figure 9.2.
Figure 9.3.
Figure 9.4.
Figure 9.5.
Figure 9.6.
Figure 9A.
Figure 9B.
Fossiliferous limestone
Radiolarian
“Stack of coins” architecture of pearls
Emiliania huxleyi—one of the most abundant coccolithoporids in the world's oceans
Magnetosomes of magnetite
Cross-section of a salt dome
The five kingdoms of life—traditional classification
The three major domains of life—RNA-based classification
Chapter 10
Figure 10.1.
Figure 10.2.
Figure 10.3.
Figure 10.4.
Figure 10.5.
Figure 10.6.
Figure 10.7.
Figure 10.8.
Figure 10.9.
Figure 10.10.
Figure 10.11.
Figure 10.12.
Figure 10.13.
Figure 10.14.
Figure 10.15.
Meteor Crater, Arizona
Carbonaceous chondrite with chondrules (Allende)
Mineral content of a typical chondrite
Pallasite meteorite, Esquel, Argentina
The Widmanstätten structure in iron meteorites
Orientation of kamacite bands in taenite
Comparison between reflectance spectra of meteorites and asteroids
The Moon
Anorthosite from the lunar highlands
Mare basalt
Olympus Mons, giant shield volcano on Mars
Winding channel probably carved by running water
Pathfinder rover next to Yogi
Worm-like features found on the Martian meteorite ALH84001
Grains of moissanite extracted from a meteorite.
viii
CHAPTER 1
MINERALS, THE ROCK CYCLE,
AND BEYOND
Introduction
The splendid variety of colorful crystals and minerals found in nature has developed as a result of
geological processes that continuously create and destroy rocks over hundreds of millions, even
billions, of years. The bulk of the earth's crust and upper mantle is composed of around only 20 to
30 minerals, which consist predominantly of silicates and oxides. The soils and rocks under our feet,
in the mountains, the oceans, deep within the earth and the inner planets are physical aggregates of
one or more of these mineral species. The diversity of minerals on earth has increased vastly through
the actions of circulating groundwater, hydrothermal solutions, and mountain-building.
We begin our examination of how minerals form in nature by reviewing the relationships between
the rock cycle, the role of water, and the formation of minerals (Fig. 1.1). These interconnections are
described in greater detail in the subsequent chapters
Figure 1.1. The geologic rock cycle (after Robinson, 1994).
1
Minerals from Igneous Activity
The rock cycle begins with the solidification of magma, or molten silicate material, either inside the
earth or at the earth's surface. Igneous rocks are formed by the crystallization of magma. These
rocks make up approximately 95% of the upper 10 miles of the earth's crust. However, their true
abundance is concealed by a thin, widespread veneer of sedimentary and metamorphic rocks near
the earth's surface. Intrusive rocks are igneous rocks that crystallize deep inside the earth. Examples
of intrusive rocks include granite (which contains the minerals quartz, orthoclase, microcline, albite,
mica), diorite (plagioclase and hornblende), gabbro (calcic plagioclase and pyroxene), and peridotite
(olivine, pyroxene) (Fig. 1.2).
Figure 1.2. Mineral composition of common igneous rock types.
Towards the final stages of crystallization of a magma, a water and volatile-rich fluid phase separates
out from the melt, particularly from magmas of granitic composition. Minerals crystallize slowly
from this heated, fluid-rich phase, often into open cavities or pockets, where they can grow into the
large, well-terminated crystals of pegmatites that are a collector's delight! Typical pegmatite
minerals include large crystals of quartz, microcline, albite, muscovite, beryl, topaz, tourmaline,
kunzite, aquamarine, and less commonly uranium, thorium, and rare-earth minerals.
Magmas that erupt on the earth's surface are called lavas. Very hot and fluid lavas which erupt from
fissures flow over vast distances and build up large plateaus, such as the Columbia River flood
basalts of Washington State, or the Deccan Plateau, India. Lava flows that accumulate over a central
vent pile up into volcanoes, such as Vesuvius, Italy, Mount Saint Helens, Washington, or Mauna
Loa, Hawaii. The explosive escape of gases dissolved in lavas creates volcanic ash. Volcanic activity
produces a variety of volcanic , or extrusive, rocks (e.g. rhyolite, andesite, basalt; Fig. 1.2). Because
they cool more rapidly at the earth's surface, the extrusive rocks are much finer-grained than their
chemically and mineralogically equivalent intrusive counterparts.
2
The distribution of volcanoes around the earth can be explained by the theory of plate tectonics.
The earth's surface is covered by around 6 to 7 major plates, or thick slabs of crust and uppermost
mantle, which slowly but inexorably drift apart or come together like giant Arctic ice floes (Fig. 1.3).
Molten lava of basaltic composition is extruded from the mantle where the plates pull apart—along
the undersea 47,000-mile-long mid-oceanic rift system, which encircles the earth like the seams on a
baseball. Volcanoes are clustered along the crests of the mid-ocean ridges. There, hydrothermal
vents and hot springs emerge at the seafloor, heated by the interaction between submarine
volcanism and seawater. As these solutions cool, sulfides of iron, copper, zinc, and lead are
precipitated at “black smoker” chimneys.
Figure 1.3. Schematic diagram showing the main elements of plate tectonics: plates spreading apart
at the mid-ocean ridges and colliding at the ocean continent or ocean-island arc boundaries.
Volcanoes also occur in mountain ranges along the “Circum-Pacific Ring of Fire”. When an oceanic
plate is subducted or shoved under a continent, oceanic rocks are dragged to great depths within the
mantle where they begin to melt. This magma, being less dense and hence more buoyant, ascends to
the surface in explosive volcanic eruptions (Fig. 1.4). Because this magma incorporates some
continental crust that is higher in silica and aluminum than oceanic basalt, it forms an intermediate
igneous rock called andesite (Fig. 1.2). Volcanic eruptions carry certain gem crystals, such as
diamonds, peridots, and some rubies and sapphires, which have crystallized in the upper mantle, all
the way to the earth's surface.
Minerals in Sedimentary Rocks
Mountains formed by plate collisions (Fig. 1.4) wear down almost as rapidly as they rise. The detritus
shed by the rising mountains is transported toward the sea by streams and rivers. Sediments
accumulate in lakes, basins, and especially offshore, along the continental shelves and slopes.
3
Figure 1.4. Schematic diagram illustrating two types of plate collisions. (Top) Subduction of oceanic
lithosphere beneath a continent gives rise to andesitic volcanism. (Bottom) Collision of two
continents produces tectonic deformation, mountain building, and metamorphism.
Rocks freshly exposed at the earth's surface immediately begin to break down by weathering and
erosion. Weathering refers to the physical attack by the daily and seasonal cycles of temperature and
rainwater, and also chemical corrosion by dissolved carbon dioxide, and organic acids from plants.
Erosion refers to the removal of broken-down rock particles by landslides, rivers, wind, waves, and
glaciers. The eroded materials accumulate in low-lying basins, eventually compacting under the
weight of overlying sediments to form sedimentary rocks.
Minerals present in sedimentary rocks, such as sandstone or siltstone, are the worn-down remains of
pre-existing rocks (quartz, feldspar, resistant heavy minerals like garnet, magnetite, or rutile). They
have also been derived from the chemical decomposition of the original minerals in igneous and
metamorphic rocks (e.g., the clay minerals in shales or mudstones are the product of weathering of
4
feldspars). Rivers carry and concentrate hard or tough minerals along with gravels and pebbles. Such
minerals accumulate in placer deposits of economic importance (e.g., cassiterite, rutile, diamond,
gem corundums, gold, zircon, and garnet. In addition, salty solutions washed down into basins may
evaporate completely in hot, arid climates, forming evaporite minerals (e.g., halite, gypsum, borates).
Deep weathering of soils and rocks, in tropical to semi-tropical climates that vary seasonally from
wet to dry, leaches out most of the more soluble constitutents. Left behind is an insoluble residue
consisting largely of hydrated aluminum and iron oxides (e.g., bauxite, limonite, goethite). These
minerals are the main components of the lateritic soils in tropical to semi-tropical climates. Bauxite
is a major ore of aluminum.
Minerals and Metamorphic Rocks
Sedimentary rocks at the ocean floors are eventually carried back toward the continents by the
moving tectonic plates. As these sedimentary rocks approach the continent, some are carried
downward into the deep-sea trenches, where the oceanic plates plunge beneath the continent (e.g.,
off the coast of Alaska or Chile; Fig. 1.4). This underthrusting (subduction) of the oceanic
lithosphere results in partial melting of the downgoing slab. Andesitic magma ascends to the surface,
creating chains of volcanoes along the rims of the Pacific Ocean (see above). Sedimentary rocks
lying on the continental shelves and slopes, off the coast, are caught in a vise between the
continental plate and the subducting oceanic plate. As a result, they are severely crumpled, folded,
and broken. The head-on collision of two continents (for example, the Indian and the Asian Plates)
also forces sedimentary and other rocks downward where they are strongly deformed (Fig. 1.4).
Tectonic pressures ultimately push the rocks upward, forming lofty mountains, such as the
Himalayas and the Tibetan Plateau. In either case, the high temperatures and pressures to which the
rocks have been subjected cause them to recrystallize into metamorphic rocks.
The minerals of metamorphic rocks recrystallized from pre-existing minerals without remelting, in
the solid state, under a wide range of temperatures and pressures. The type of metamorphic rock
which forms depends not only on the temperature or pressure, but also the starting composition of
the rock. For example, a pure quartz sandstone and limestone will become quartzite and marble,
respectively. On the other hand, impure sedimentary rocks will produce a range of different rocks.
For example, shale--a mixture of clay minerals and fine-grained quartz—recrystallizes to a schist,
consisting of a mixture of micas and quartz, with minor amounts of almandine garnet, staurolite,
actinolite, tremolite, kyanite, etc. At even higher temperatures, the quartz, feldspar, and biotite
separate into alternating light and dark bands, forming a gneiss. At temperatures close to the melting
point (a process akin to annealing, in metallurgy), the minerals are reassembled into a granular,
almost sugary mass, reminiscent of an igneous rock. Metamorphic rocks reaching this stage are
called granulites. Finally, going one step further, increasing heat may melt the rock altogether,
forming magma anew. Tectonic forces eventually raise the metamorphic (and other) rocks from the
earth's depths back to the surface, where they are again exposed to weathering and erosion,
continuing the rock cycle.
5
Minerals and Circulating Water
Water is indispensible to the generation of an even greater variety of minerals. Minerals can
precipitate from solutions derived from rainwater that has seeped into the earth. These, fairly cool,
meteoric solutions contain carbon dioxide, which dissolves limestone along fractures and cracks in
the rock, eventually enlarging these fissures into caverns (Fig. 1.5a). Above the water table some of
the water evaporates, causing calcium carbonate to become supersaturated and to deposit from
solution. The continual dripping of water and deposition of dissolved minerals builds up stalagtites
(hanging) and stalagmites (mounds) of calcite or aragonite.
a.
b.
c.
Figure 1.5. Sources of hydrothermal solutions. (a) Meteoric water (rainwater) warms by descending
to great depths within the earth and/or passing close to buried magma chambers. (b) Superheated
water and other volatile constituents released by crystallizing magmas mix with descending meteoric
water. (c) Water in sediments is heated by the high temperatures and pressures associated with
tectonic deformation and/or metamorphism (after Robinson, 1994).
Occasionally, the migrating groundwater picks up other elements--iron, manganese, and less
commonly, copper or zinc. Minerals which precipitate from cool, near-surface solutions, typically
exhibit a botryoidal or grape-like appearance. When broken open, the botryoidal aggregates display a
radiating and concentrically-banded texture. Examples of such minerals include goethite, hematite,
rhodochrosite, malachite, azurite, smithsonite, also, chalcedony, banded agates, wavellite.
Solutions that have migrated through deep basins become hot, acidic, and more corrosive.
Descending solutions also warm up as they approach magmatic bodies at depth (Fig. 1.5b). When
sedimentary rocks are tectonically deformed or metamorphosed, any enclosed water is also heated
up (Fig. 1.5c). These hot solutions are referred to as hydrothermal. Hydrothermal solutions travel
along faults and fissures in the earth's crust. Because they are hot and corrosive, they are capable of
dissolving many minerals. But as they cool, they become very concentrated and precipitate new
minerals in veins or vugs. Typical hydrothermal minerals include galena, chalcopyrite, other
base-metal sulfides, gold, silver, together with associated gangue minerals, such as barite, fluorite,
and calcite.
6
Minerals and Organisms
Biological processes are involved in the formation of many minerals in sedimentary environments,
e.g. calcite, aragonite, pyrite, magnetite, hydroxyapatite, opal, to name a few. Entire sedimentary
deposits may be made from the remains of living organisms. Limestone is a calcium carbonate rock
that consists of the fragments of shells, corals, and marine microorganisms. Cherts are composed of
silica derived from marine microorganisms, such as diatoms or radiolarians. Phosphorite--a
phosphate-rich rock--contains fish bones and other fossil material. Coal, oil, and natural gas are the
debris of biogenic matter.
The oxygen we breathe was released during photosynthesis, by plants and algae. Microbes accelerate
the weathering of rocks and soils, oxidize sulfide minerals to sulfates, and contribute to the creation
of some metal deposits. Microbial activity may even influence the Earth's major geochemical cycles.
Thus, life acts as a geological force and plays an important role in the origin of minerals.
Minerals Beyond the Earth
Meteorites were our only samples of extraterrestrial matter until the astronauts landed on the Moon
in 1969. Dating back to 4.6 billion years ago, meteorites give us a window on the earliest days of the
Solar System. The minerals of meteorites, for example, nickel-iron, olivine, pyroxene, troilite, reveal
conditions that existed during the formation of the planets--information that has long since been
erased on Earth due to weathering, erosion, volcanism, and tectonic activity. Our reconstruction of
the earth's interior into crust, mantle, and iron core, has been largely based on the study of
meteorites. Certain anomalous inclusions in meteorites may even have originated beond the Solar
System. They may be our only actual samples of starstuff.
The lunar rocks brought back to Earth by the Apollo astronauts have enabled us to decipher the
geologic history of our nearest neighbor in space. The Moon's heavily-cratered surface records an
early, violent period of intense meteoritic bombardment, followed by a long period of volcanism.
However, being much smaller than the Earth, the Moon cooled a lot faster and thus by 3 billion
years ago, volcanic activity had ceased.
In contrast to the earth, the Moon shows no signs of plate tectonics, and lacks and atmosphere and
water. As a consequence, the number of minerals on the Moon is quite limited as compared to the
Earth. The major minerals present—pyroxenes, olivine, plagioclase, and ilmenite--are of igneous
origin, both deep-seated and surficial. They have been modified by impacting processes.
Within the last 25 years, unmanned spacecraft have explored the Solar System. Among the planets,
Mars is the most like Earth. Although still bearing scars of the early meteorite bombardment, Mars
has remained more geologically active than the Moon. Its surface is covered by extensive lava flows,
towering volcanoes, deep canyons, winding river-like valleys, desert dunes, and polar ice caps. The
composition of the Martian surface has now been probed directly by surface landers in several
locations and indirectly by instruments on board orbiting spacecraft and from earth-based
telescopes. Mars is a rusty planet--iron in its rocks has been weathered to iron oxides (Fig. 1.6). But
spectroscopic evidence of pyroxene points to an extensive volcanic history.
7
Figure 1.6. Mars—view of the surface from Pathfinder.
In summary minerals are the products of physical, chemical, and geological processes occurring
from the Earth's outer core to its skin, the bottom of the seas to the loftiest mountains, and well
into space—in other words, wherever circumstances allow atoms to assemble into crystals of the
solid state.
Recommended Reading
Minerals: An Illustrated E xploration of the Dynamic World of Minerals and Their
Properties, G.W. Robinson, Simon and Schuster, New York, 1994.
Gemstones and Their Origins, P.C. Keller, Van Nostrand and Reinhold, 1990.
8
CHAPTER 2
MINERALS FROM THE MELT
Introduction
Beneath a thin, yet widespread cover of sedimentary and metamorphic rocks, the bulk of the earth's
crust and mantle are made of igneous rocks. These rocks have crystallized from a magma, or
molten silicate mass (Fig. 2.1). Heat from the earth's interior melts rocks deep in the upper mantle or
lower crust. As molten magma slowly cools, randomly-arranged atoms or ions gradually crystallize
into the orderly, symmetrical structures of the major igneous silicate minerals (see Vol. I). The first
minerals to freeze out are those with the highest melting points, such as olivine and anorthite
(calcium-rich plagioclase feldspar, Fig. 2.2). These are succeeded by minerals with progressively
lower melting points, such as pyroxene, hornblende, biotite, and more sodium-rich plagioclase
(andesine-albite). The last minerals to solidify from a magma are quartz, alkali feldspars, and
muscovite—the main constituents of granite.
Figure 2.1. Igneous processes. Molten magma chamber at depth. Magma is brought to the surface
along deep-seated fractures or fissures. Lava (and ash) erupts from volcanoes and fissures at the
earth's surface.
Figure 2.2. Crystallization sequence of minerals from a silicate melt (Bowen's reaction series).
9
Gems from Volcanic Rocks
Topaz in Rhyolite
Magmas of granitic composition which reach the earth's surface erupt as rhyolite lava flows. At high
temperatures (between 850-600EC), lava with abundant amounts of silica, fluorine, and rare metals
will crystallize topaz and less commonly, red beryl (bixbite). These minerals are found in the Thomas
Range and Wah Wah Mountains, Utah, and also the topaz rhyolites of central Mexico.
The Thomas Range, Utah, consists of a number of rhyolite lava flows and domes which have
erupted some 6 to 8 million years ago. The lavas contained dissolved fluids rich in fluorine,
beryllium, and lithium. Bands of elliptical hollows, or lithophysae, were created by the escape of
gases through zones of weakness in the flows. Continued diffusion of fluids through the voids
promoted crystal growth. Clear sherry brown to pink topaz crystals line cavities and vugs in the
rhyolite. The topaz occurs as elongate prisms, modified by dipyramids, in clusters, or loose,
sometimes heavily included with sandy microcrystalline quartz (Fig. 2.3). The sherry color in topaz
results from exposure to natural radiation (e.g., traces of uranium) in the rhyolite. Unfortunately, the
color fades in strong sunlight. Pink topaz, on the other hand, may be colored in part by Mn+3. It is
light-stable.
The Thomas Range is also noted for small, raspberry-red, tabular,
hexagonal crystals or rosettes of red beryl (bixbite). Associated minerals
include lustrous black cubes of bixbyite (manganese iron oxide), dark redbrown spessartine-almandine, black, prismatic to acicular pseudobrookite
(iron titanium oxide), and durangite, a rare orange-red sodium aluminum
arsenate fluoride.
Figure 2.3. Topaz crystal from the Thomas Range, Utah
Peridot, Rubies, Sapphires
The magnesium-rich form of olivine, or forsterite (Mg2SiO4) is one of the
first minerals to crystallize from the melt (Figs. 2.2; 2.4). It occurs in mantle
rocks, such as dunite (a nearly pure olivine rock), or peridotite (a rock
consisting of olivine with lesser amounts of pyroxene). It is also a common,
although minor constituent of gabbro (plagioclase and pyroxene) and its
volcanic equivalent, basalt (Fig. 1.2).
Figure 2.4. Olivine (peridot), northwest Pakistan.
Peridot, the gem variety of forsterite olivine, has been mined in Zabargad Island (St. John's Island)
in the Red Sea since ancient times. Crystals of peridot from Basham, Gilgit district, NW Pakistan
resemble those from Zabargad. Peridot is also found in Myanmar (Burma). However, most peridot
used in jewelry today comes from San Carlos, Arizona, where it occurs as elliptical “bombs” in a
basalt lava flow. Kilbourne Hole, southern New Mexico, is another, less-well known, source of
10
potentially gem-quality peridot. It is an approximately 180,000 years old volcanic crater (actually a
maar—where the volcano erupted into a small lake or into water-rich sediments). The peridot is
recovered from potato-sized “bombs”, or nodules of olivine in the basalt. The olivine probably
crystallized in the upper mantle and was carried to the surface in the basaltic magma, which then
erupted explosively at the surface.
Rubies, sapphires, zircons, some spinels and garnets, are also minerals that probably crystallized in
the upper mantle, but were subsequently carried to the surface by alkali basalts. The rubies from
Chanthaburi-Trat, Thailand, probably originated in this manner, although extensive weathering of
the basalts has led to the formation of secondary placer and alluvial deposits (see Chapter 3). Other
localities furnishing gem-quality rubies or sapphires in alkali basalt include Barrington, Australia,
Pailin, Cambodia, and Dak Nong, southern Vietnam.
Profs. Levinson and Cook, University of Calgary, Canada, have proposed a two-stage process for
the origin of corundum in alkali basalt. In the first stage, aluminum-rich shales, which contain clay
minerals like kaolinite or illite, or micas, are carried to great depths in the earth's mantle by plate
collisions or subduction (see Fig. 1.4). (Bauxite or lateritic soil, rich in hydrated aluminum oxide,
could also serve a starting material, in some cases). Under the high pressures and temperatures that
prevail at such depths, the hydrous clays and micas break down into feldspar, corundum (aluminum
oxide), and water. The pressures required for the crystallization of corundum exist between depths
of 24 km to 48 km, near the boundary between the earth's lower crust and upper mantle.
In the second stage, alkali basalt magmas pick up corundum crystals on their way to the surface.
Alkali basalts do not have enough silica to react with corundum to produce aluminum silicates, such
as kyanite. During the upward journey, the rocks originally enclosing the corundum (and zircon,
spinel, etc.) were dissolved in the magma, releasing the corundum crystals. The rounded, etched, and
pitted condition of many crystals is evidence of such corrosion. Corundum, according to this
scenario, is created independently of its host rock—alkali basalt. This situation is analogous to the
origin of diamonds (see below).
Diamonds
Diamond is a form of carbon that crystallizes under
high temperatures (900-1300EC) and depths of 100
to 300 kms in the earth's upper mantle (Fig. 2.5).
Minute mineral inclusions in diamonds can now be
dated, using sophisticated, modern analytical
techniques. These studies have shown that many
Figure 2.5. Diamond crystals.
diamonds are very old—often over one billion
years old (the Earth is around 4.6 billion years old). However, the diamond crystals are usually much
older than the kimberlites in which they are enclosed. Kimberlite has merely acted as an “elevator”
that has picked up “passengers” (i.e., the diamonds) on its ascent to the earth's surface.
Kimberlite is a brecciated (fragmented) mantle rock consisting of olivine, phlogopite mica, pyrope
garnet, diopside (often chromium-rich), calcite, chromite, ilmenite, spinel, and hopefully, diamond.
It may also contain fragments of mantle rocks, such as peridotite, eclogite, and any of the other rock
types encountered on the way up. Lamproite is another diamond host rock consisting of olivine,
diopside, phlogopite, amphibole, leucite, and sanidine.
11
Studies of inclusions in diamonds establish that diamonds did not originally crystallize in kimberlite
or lamproite, but actually came from two other types of rocks: peridotite and eclogite. Peridotite is
a coarse-grained assemblage of olivine, pyroxene, pyrope garnet, and spinel. E clogite, on the other
hand, consists of a jade-diopside (omphacite) and almandine-pyrope mixture, with minor amounts
of rutile, kyanite, and corundum. Chromite and pyrrhotite are present in both peridotite and
eclogite, and thus are not very diagnostic of the source. However, garnet inclusions from these two
source rocks can be separated by means of their color and chemical composition. Pyrope inclusions
derived from peridotite contain chromium and tend to be purplish-red, whereas almandine-pyrope
from eclogite tends toward an orange-red color. Other diagnostic inclusions of peridotites are
olivine, chromite, and green chrome-diopside. Diamonds from eclogitic sources may contain
inclusions of omphacite or kyanite.
The very same minerals that occur as inclusions in diamonds also serve as “pathfinders” or indicator
minerals in the exploration for diamonds. Since kimberlite pipes are rather small in outcrop area and
not all kimberlites contain diamonds, diamonds are tracked by looking for more abundant associated
minerals that are also hard, heavy, and resistant, that can survive and concentrate in stream gravels.
The usual indicators include pyrope garnet, chrome diopside, ilmenite, and chromite. This
exploration strategy has paid off handsomely with the discovery of new diamond deposits in
northern Canada!
Kimberlite magmas, containing dissolved gases, ascended
rapidly to the earth's surface along deep-seated fractures.
Diamonds rose from depths of around 100 km in as little as
4-15 hours! This conclusion is based on the fact that the
diamonds survived the journey without reverting to
graphite—the more stable form of carbon at shallower
depths. Most kimberlite pipes are located in continental
cratons or shields (ancient and stable portions of the earth's
crust). The classic kimberlite pipe is a carrot-shaped
structure that tapers at depth and forms a broad, shallow
crater at the surface (Fig. 2.6).
Figure 2.6. Idealized diagram of a kimberlite pipe. The
names on the right side indicate the erosion levels at the
various diamond-bearing pipes in South Africa.
Historically, diamonds came from India, and then later, from Brazil. Since the late 19th century, they
have been mined in South Africa. During the 20th century, additional deposits were found in
Angola, Zaire, Ghana, Tanzania, Russia, and Australia. A new diamond mine is being constructed at
Point Lake, the Northwest Territories, Canada, in late 1998. It is expected to produce 5 million
carats within 5 years! In the United States, diamonds have been recovered from a lamproite pipe at
Murfreesboro, Arkansas. The only recently active U.S. mine was the Kelsey Mine, along the
Colorado-Wyoming state line, which opened in June, 1996 and closed in September, 1997. It
produced around 12,000 carats of rough. A 28.3 carat diamond from the mine was sold recently.
12
Pegmatites
In general, rocks that cool slowly at depth grow large crystals. Coarsely-grained, intrusive igneous
rocks are called pegmatites. Most pegmatites are of granitic composition—consisting of large
crystals of quartz, microcline, orthoclase, albite, and muscovite. Pegmatites have been exploited
economically for feldspar and muscovite mica. They are also a source of beryllium (from beryl),
lithium, boron, niobium, tantalum, rare earth metals (elements such as lanthanum, cerium,
samarium, europium, etc.), uranium, and thorium. Other elements often present include manganese,
phosphorus, and fluorine. However, it is the presence of well-formed, gemmy crystals that
particularly appeals to the collector. These include such species as tourmaline, topaz, beryl (var.
aquamarine, heliodor, morganite), kunzite, spessartine garnet, lepidolite mica, chrysoberyl, amazonite
(bluish-green microcline), smoky quartz, amethyst, rose quartz, and apatite.
Pegmatites typically form toward the late stages of crystallization of a fluid-rich, granitic magma.
Crystallization starts at temperatures as high as 750EC to as low as 300EC. Pegmatite bodies have
massive, compact shapes or intrude into surrounding rocks as sills or dikes. (Sills lie parallel to the
layers of rocks they intrude whereas dikes cut across them). Pegmatites from greater depths tend to
be unzoned and consist mainly of quartz, feldspar, and mica, whereas shallow pegmatites (from less
than 4-5 km) are often zoned and contain pockets of large crystals, sometimes containing gems and
other rare minerals, such as cassiterite, pollucite (a cesium-aluminum silicate-hydrate), columbitetantalite, samarskite, uraninite, and thorite (Fig. 2.7). These crystal pockets are called miarolitic
cavities. Some crystals can grow extremely large, for example 15 meter-long spodumene crystals
from the Etta Pegmatite, in the Black Hills, South Dakota, or the 10 meter-long beryl from the
Bumpus quarry near Bethel, Maine. Such giant crystals grow at extremely slow cooling rates in the
presence of water and other volatile constituents such as fluorine, phosphorus, and boron.
Figure 2.7. Pegmatite minerals: tourmaline from Brazil (left), aquamarine from N. Pakistan (middle),
and columbite from Maine (right).
Some well-known gem-bearing pegmatite localities include Minas Gerais, Brazil, San Diego,
Riverside Counties, California, Newry, Mt. Mica, and Mt. Apatite, Maine, the Pikes Peak area,
Colorado, Gilgit district and the Nanga Parbat region, northwest Pakistan, Ural Mountains., Russia.
The Himalaya Mine, San Diego County, California, has produced attractive gem crystals of
tourmaline since the beginning of this century. Gem-bearing pegmatite dikes have intruded into
Cretaceous-age rocks around 98 to 93 million years ago. At the upper contact between the dike and
country rocks, radiating clusters of schorl lying nearly perpendicular to the contact project
downward toward the pocket zone (Fig. 2.8). Masses of fine-grained purple lepidolite are often
13
located near gem-producing pockets and therefore act as useful guides. Surrounding the pockets are
large, blocky, tan microcline, bladed cleavelandite (a variety of albite characteristic of pegmatites),
colorless to smoky quartz, books of silvery muscovite, and occasionally beryl crystals. The Himalaya
tourmalines are bi- and multi-colored, in pink and shades of green, usually capped by green
terminations.
Figure 2.8. Schematic diagram of a gem-bearing crystal pocket from the Himalaya pegmatite, San
Diego Co., California (from J. Sinkankas, Field Collecting Gems and Minerals, with permission,
GeoScience Press, Inc., Tucson, AZ ©1961, 1970 J. Sinkankas).
In recent years, the Shingus-Dusso area near Gilgit, Pakistan has furnished highly collectible gemmy
crystals of tourmaline, aquamarine, topaz, and spessartine. The Shingus-Dusso district lies within
one of the most geologically dynamic environments in the world, at the boundary between three
major crustal plates—the Asian Plate, the Kohistan Island Arc, and the Indian Plate. The nearby
Nanga Parbat massif (elev. 26,660 feet), around which the Indus River bends in a sharp hairpin
curve, is rising at rates of up to 10 mm/yr (0.4 in/yr). Here, the Indus has carved one of the deepest
gorges in the world. The age of the gem-bearing pegmatites may be as geologically recent as 5
million years, making these among the youngest such deposits in the world!
14
The Dusso pegmatite yields lovely, pale-blue, clear, well-formed aquamarine crystals typically
encased in white albite “jackets”. The multi-colored elbaite tourmalines of nearby Stak Nala also
frequently wear cleavelandite jackets. The tourmalines are often dark green to nearly black, grading
into clearer apple-green, and colorless to pale pink zones toward the upper termination. Other
highly-sought minerals include clear colorless to sherry crystals of topaz, with multiple complex
forms, and small bright reddish-brown spessartine crystals, up to 5 cm (2 inches) across.
The pegmatites of the Pikes Peak district, Colorado furnish deep bluish-green to greenish-blue
amazonite (a variety of microcline, Fig. 2.9). Crystal pockets are enclosed by graphic granite (a
hieroglyphic-like intergrowth of quartz and feldspar). A unique feature of
the Pikes Peak amazonites is the occurrence of white overgrowths of
potassium or sodium feldspar on microcline. The potassium-feldspar
overgrowth (“white caps”) develop preferentially on alternating prism and
dome faces. Blue-green microcline adjacent to dark brown smoky quartz
crystals make strikingly handsome specimens.
Figure 2.9. Amazonite from Pikes Peak district, Colorado.
The bluish-green coloration is believed to be caused by the presence of lead. (Galena occurs in
minor quantities in the pegmatite). Lead in the form of Pb+2 ions may substitute for K+ in the
microcline, with the electrical charges being balanced by the replacement of Si+4 for Al+3 elsewhere
in the crystal lattice (see A n Introduction to Minerals, Chap. 1). Fe+3 can also substitute for Al+3. Many
samples of the Pikes Peak feldspar fluoresce a distinct crimson red, orange, and yellow under
shortwave ultraviolet light. Fluorescence enhances the detection of color zoning in the specimen.
A number of pegmatite deposits popular with New York City collectors occur in nearby
Connecticut. Among these are: the Case Quarry, Portland, CT (blue beryl), Timm's Hill prospect,
Haddam, CT (schorl in albite, iolite—gem cordierite), Swanson's Quarry, Haddam Neck, CT
(lepidolite, green-yellow beryl, North and South Hollister pegmatite quarries, South Glastonbury, CT
(beryl, lepidolite, schorl), Simpson Quarry, South Glastonbury, CT (aquamarine, schorl, lepidolite).
Agpaitic Pegmatites
An unusual type of pegmatite occurs in agpaitic rocks, which are feldspar-rich igneous rocks higher in
sodium, but lower in silica and aluminum than ordinary granitic rocks. Minerals found in these rocks
contain sodium, for example: aegirine (pyroxene), arfvedsonite (amphibole), nepheline, sodalite,
analcime, and cancrinite. The pegmatitic stage is further enriched in elements such as beryllium,
barium, titanium, zirconium, and rare earth elements, which because of their large atomic sizes do
not readily fit into the crystal lattices of the earlier-formed silicates. These kinds of igneous rocks are
very rare worldwide, but some well-known localities include the Narssarssuk complex in southern
Greenland, the Langesundfjord area of Norway, Lovozero and Khibiny massifs in the Kola
Peninsula, and Mont Saint-Hilaire, Quebec.
Mont Saint-Hilaire is familiar to collectors as a source of over 200 mineral species, many of which
are exceedingly rare and unique to this locality. The igneous intrusive complex consists of three
separate phases. The oldest intrusion is composed of gabbros and pyroxenites. The second forms a
ring dike—a circular structure with nepheline gabbros, diorites, and monzonites. The
last—occupying the eastern half of the complex—consists of nepheline and sodalite syenites, also
15
breccias with fragments of the other units and basement rocks. The intrusion is surrounded by a
corona of hornfels—a dark, fine-grained rock formed by the interaction of the igneous rocks with
the enclosing sedimentary rocks.
The pegmatites are the residual part of the magma and the last to have
crystallized. They occur as dikes, pipes, and lenses. Occasionally, dikes or lenses
open into pockets with well-terminated crystals of microcline, aegirine, and
locally analcime, serandite, and catapleiite, the latter two species highly-regarded
collectors' items (Fig. 2.10). Other species which produce large crystals include
acicular to prismatic black aegirine, bladed, tabular, bronze-brown astrophyllite,
acicular to prismatic white, tan, or gray elpidite, reddish eudialyte, and leifite
(white fibrous to acicular, hexagonal prisms, and occasionally in spherical
aggregates).
Figure 2.10. Serandite (orange) and analcime (white), Mont Saint-Hilaire,
Quebec.
Porphyry Copper Deposits
Porphyry copper deposits originate in shallow magmatic intrusions that formed at the base of
volcanoes, within several kilometers of the earth's surface. The host rocks include such silicic to
intermediate igneous rocks as granite, monzonite, and diorite (Fig. 1.2). These rocks typically have
larger crystals (porphyry) set in a finer-grained matrix. The ore minerals are deposited from hot
fluids, largely derived from the cooling magma. Ore minerals include copper sulfides (such as
chalcopyrite, bornite, and enargite), also molybdenite (molybdenum sulfide), and sometimes, gold.
The ore minerals occur as finely-disseminated sulfide grains, veins, and stockworks. The highly
cracked texture of the latter results from shattering of the rock by boiling fluids within the cooling
intrusion as pressures lessened (Fig. 2.11). Chemical interactions between the magmatic fluids and
the enclosing host rocks have hydrothermally altered the surrounding host rocks to zoned
assemblages of biotite, potassium feldspar, fine-grained micas, clays, chlorite, and quartz.
Figure 2.11. Schematic drawing of porphyry copper deposit
16
Until recently, it would not have been economically-feasible to exploit the low-grade copper ore
(typically as low as 0.5-0.7% Cu) if not for supergene (secondary) enrichment, which has
concentrated the primary ores at depth (see Chapter 5). But at present, the ore is extracted from
giant open-pit mines and concentrated by acid leaching and electrolytic deposition.
Major porphyry copper deposits occur in Arizona and New Mexico, Mexico, Peru, Chile, and Papua
New Guinea. Morenci, Arizona is at present the largest open pit mine in the U.S. The Pima district
(including the Twin Buttes, Sierrita, and Esperanza mines) south of Tucson produced 30% of the
U.S. and nearly 17% of the world's copper in the 1970s. Other important Arizona copper mines
include Inspiration, San Manuel, and Bagdad.
17
Recommended Reading
Betts, J., 1998. The quarries and minerals of South Glastonbury, Connecticut. Rock s & Minerals,
Nov./Dec. 1998.
Bideaux, R.A. and Wallace, T.C., 1997. Arizona copper. Rock s & Minerals, 72(1), 10-27.
Fisher, J., Foord, E.E., and Bricker, G.A., 1998. The geology, mineralogy, and history of the
Himalaya Mine, Mesa Grande, San Diego County, California. Rock s & Minerals, 73 (3), 156180.
Foord, E.E., Chirnside, W., Lichte, F.E., and Briggs, 1995. Pink topaz from the Thomas Range,
Juab Co., Utah. Mineralogical Record, 26 (1), 57-60.
Foord, E.E. and Martin, R.F., 1979. Amazonite from the Pikes Peak batholith. Mineralogical
Record, 10 (6), 373-384.
Harlow, G.E., ed., 1998. The Nature of Diamonds. Cambridge University Press, American
Museum of Natural History, New York, 278p.
Holfert, J., Mroch, W., and Fuller, J., 1996. A field guide to Topaz and Associated Minerals of
the Thomas Range, Utah, (Topaz Mountain) , vol. 1, HM Publ., 103p.
Horvath, L. and Gault, R.A., 1990. The mineralogy of Mont Saint-Hilaire, Quebec . Mineralogical
Record, 2(4), 284-359.
Kazmi, A.H., Peters, J.J., and Obodda, H.P., 1985. Gem pegmatites of the Shingus-Dusso area,
Gilgit, Pakistan. Mineralogical Record, 16 (5), 393-411.
Kirkley, M.B., Gurney, J.J., and Levinson, A.A., 1991. Age, origin, and emplacement of
diamonds: scientific advances in the last decade. Gems and Gemology, Spring, 1991, 2-25.
Levinson, A.A. and Cook, F.A., 1994. Gem corundum in alkali basalt: origin and occurrence.
Gems & Gemology, Winter 1994, 253-262.
Menzies, M.A., 1995. The mineralogy, geology, and occurrence of topaz. Mineralogical Record, 26
(1), 5-53.
18
CHAPTER 3
MINERALS FROM SEDIMENTARY
DEPOSITS
Background
Sedimentary rocks cover approximately 80% of the earth's surface, yet they only constitute around
5% of the upper 10 miles of the earth's crust. Thus, sedimentary rocks form a thin layer over a crust
made up predominantly of igneous and metamorphic rocks. Rainwater, ice, and extreme
temperature fluctuations cause the minerals present in igneous or other rocks to break down
chemically or physically. The silicate minerals, such as augite or plagioclase feldspar, which solidified
at high temperatures from the melt are less stable at the low temperatures and pressures that prevail
near the earth's surface (see Chapter 2). Therefore they react chemically with water, carbon dioxide,
and organic acids in the soil to generate a variety of clay minerals, silica, and dissolved sodium,
potassium, calcium, and magnesium ions. Rocks also break down mechanically by continual
exposure to water and wind.
The decomposed bits of rock are carried by rivers, glaciers, or wind to low-lying basins, where they
accumulate into layered deposits. Under the cumulative weight of overlying layers and heat inside
the earth, the sediments are gradually compacted, cemented, and converted into sedimentary rocks.
In addition, dissolved constituents carried to lakes or the sea by rivers may evaporate in shallow
basins, under extremely arid conditions.
Placer Deposits
Placer deposits are accumulations of minerals that occur together with gravels and pebbles carried
by rivers or shore currents (Fig. 3.1). Minerals found in placer deposits are typically hard (i.e.,
resistant to abrasion), lacking cleavage, insoluble, tough, not
easily oxidized, and relatively dense. Examples of minerals
found in placers include cassiterite, rutile, diamond, rubies
and sapphires, gold, zircon, and garnet. While gold is fairly
soft, it is also dense, highly malleable, and chemically
resistant, making it a good placer mineral. Even although it
is continuously abraded by pebbles in the streambed, gold is
shaped into rounded nuggets or thin flakes.
Figure 3.1. Cross-section through an alluvial deposit. Heavy minerals concentrate in between
cobbles and boulders and near the river bed.
19
Placer minerals are concentrated where the coarsest gravels occur, i.e. wherever stream velocities
decrease: at sand bars, where currents slow down at a decrease in slope, the inside bend of a
streambed, or where streams widen, join, enter lakes, or the sea (Fig. 3.2). At the shore, waves and
longshore currents move sediments and further sort minerals by density and size. Denser minerals
accumulate in gravel layers or lenses on beaches during major storms. Gems from placer deposits
are usually of higher quality than in their parent deposits,
because the abrasion and tumbling along the streambed have
disintegrated any flawed or fractured stones (Fig. 3.3). Gembearing placer deposits include the diamond-bearing beach
sands of Namibia and South Africa, and rubies and sapphires
from Mogok, Burma, the Umba River, Tanzania, and
especially Sri Lanka (Ceylon).
Figure 3.2 Concentration of heavy minerals (black) in a river
at points where the stream velocity drops (after Barnes,
1988).
Figure 3.3. Placer gems: diamond (left), sapphire (middle), and gold (right).
A fortune in glittering diamonds has been recovered from the beaches of Namaqualand, South
Africa, and neighboring Namibia, but now after nearly 70 years of stripping the beaches, the surface
deposits are nearly mined out. Therefore mining activity is shifting offshore, where, conservatively,
over a billion carats of gem-quality stones still await. The primary sources of these diamonds are 80
to 120 million-year old kimberlite pipes on the craton of the continental interior. A 100 million years
of erosion around Kimberley and other pipes has removed over 1,000 meters of the original rock
cover, releasing diamonds to be carried seaward by the Orange River and its tributaries. Only the
unflawed diamonds have survived the long journey to the sea and the subsequent pounding by
waves. As a result, over 90% of the stones are of gem quality. They average around one carat in
weight near the mouth of the Orange River. Longshore currents have spread the diamonds all along
the coast. Over the last 100 million years, large fluctuations in sea level have resulted in the
formation of diamond-bearing beach terraces at various levels, both above present sea level on
shore, and at depths ranges from 20 to 120 meters beneath the ocean surface. Diamonds are being
dredged from shallow depths along the South African and Namibian coasts. De Beers is operating
several mining vessels capable of extracting diamonds from down to 200 meters depth, off the coast
of Namibia.
Sri Lanka is one of the world's major sources of placer gemstones. Although most noted for
sapphires, the gem gravels of Sri Lanka have also yielded cat's-eye and alexandrite varieties of
20
chrysoberyl, ruby, spinel, zircon, peridot, tourmaline, beryl, moonstone, and garnet. Famous
gemstones of Sri Lankan origin include the 104-carat Stuart sapphire and the Timur “ruby”—a
361-carat red spinel—from the British royal crown jewels. The Smithsonian Institution in
Washington D.C. has on exhibit the spectacular 138.7 carat Rosser Reeves star ruby and the 423carat Logan sapphire. The 362-carat Star of India star sapphire, on view at the American Museum of
Natural History, is in fact from Sri Lanka.
Large-scale mining is concentrated around Ratnapura, 97 km southeast of the capital city of
Colombo, and the Elahera area, in the central part of the island. Gemstones are usually recovered
from the lowest gravel and sand beds in river and lake deposits, some up to 10-15 meters thick.
Mining methods are still relatively primitive. Workers dig out gem-bearing gravels using pick and
shovel from open pits or simple shafts, hoisting the ore to the surface in baskets. The gravels are
then washed in special round baskets, using a swirling motion, much like Western miners panning
for gold. The gem concentrates are then sorted manually. Gravels from modern riverbeds are
dredged.
Many of the famous historic gold strikes, such as those in California, Colorado, the Yukon, and
Australia, initially came from placers. The world's largest fossil placer deposit is in the Witwatersrand
basin in South Africa. It has produced nearly 40% of all the gold ever mined! The Witwatersrand
formed around 3 billion years ago, as wind and rain eroded the rising mountains and shed debris
into fast-flowing rivers. Over time, the land gradually sank, and the rivers kept adding more detritus,
until a nearly 7 kilometer-thick layer of sand, pebbles, gravel, and gold had accumulated. Uranium is
also found in the Witwatersrand, but is recovered only as a by-product of the gold-mining.
Recently, some geologists have questioned the placer theory for the Witwatersrand. They claim that
the gold was deposited several hundred million years later than the sediments, when heated solutions
containing dissolved gold traveled through pores and fractures in the rock and reacted chemically
with iron and carbon in the pebble layers, precipitating the gold there. They cite evidence that the
ore minerals replaced the original sedimentary grains. Similar fluid inclusions were found in fractures
from both the ore-bearing rocks and in rocks believed to have been the original source of the gold.
This was taken to mean that the gold had been carried in solution and had subsequently deposited in
both types of rock. The controversy over the origin of this deposit is by no means settled.
Residual Deposits
Intense subtropical to tropical weathering leaches away the more soluble components of rocks,
leaving behind an insoluble residue. The residual materials are enriched in iron and aluminum oxides
that typically form a hard laterite crust on the soil. In economic terms,
the most important lateritic deposits are bauxites, or ores of
aluminum (Fig. 3.4). Bauxite is a rock consisting of a mixture of
diaspore [AlO(OH)], boehmite [AlO(OH)], gibbsite [Al(OH)3], with
lesser amounts of hydrated ferric oxides, such as goethite [FeO(OH)],
lepidocrocite [FeO(OH)], limonite [Fe2O3 • nH2O], hematite (Fe2O3),
and also some minor kaolinite and silica. It is derived from the
weathering of aluminum-rich rocks, which are also poor in iron and
silica. Bauxite is found in Australia, Jamaica, Guyana, Surinam, China
and a number of other localities.
Figure 3.4. Bauxite—a
major aluminum ore.
21
Nickel is another element that is concentrated in laterite deposits, which constitute a major potential
economic reserve of this metal. Nickel-bearing laterites result from the weathering of Ni-rich mafic
to ultra-mafic rocks (i.e., basalts, peridotites, serpentinites) under tropical to semi-tropical climates.
An important nickel ore mineral is garnierite (hydrous nickel magnesium silicate). It is a bright green,
lamellar or felted, microcrystalline aggregate (H 2-4). Nickel laterites are mined in New Caledonia
and Cuba.
Evaporites
In hot, arid regions, evaporation of water concentrates dissolved salts, which then precipitate out
from saturated solutions. The least soluble minerals, such as gypsum, anhydrite, or calcite,
precipitate first. The most soluble minerals, such as halite or carnallite, are the last to deposit.
Evaporites accumulate in playas, salt lakes, or in constricted embayments that are cut off from the
ocean.
Among the more common minerals that typically form in evaporite deposits are: halite (rock salt,
NaCl), gypsum [CaSO4 • 2H2O], sylvite (KCl), carnellite (potassium magnesium chloride hydrate),
anhydrite (CaSO4), borax (sodium borate hydrate), kernite (hydrated sodium borate), colemanite
(calcium borate hydrate), ulexite (hydrated sodium calcium borate), saltpeter (niter, KNO3), trona
(hydrated sodium carbonate), epsomite (magnesium sulfate hydrate), and glauberite (sodium calcium
sulfate) (Fig. 3.5).
Figure 3.5. Halite, Krakow Saltworks, Poland (courtesy American Museum of Natural History).
Notable sources of evaporites include the Great Salt Lake, Utah, the Dead Sea, Israel, Death Valley,
California, parts of the Sahara, the Chiricahua Desert, Mexico, the Atacama Desert, Chile. Massive
rock salt deposits are found in a belt extending from New York through Michigan, Saskatchewan,
Canada, and also from salt domes in the Gulf Coast, Germany, Romania, and Iran. The sparkling
white dunes of the White Sands National Monument, New Mexico are made of tiny gypsum
crystals—wind-blown remnants of a former lake-bed. Well-crystallized sylvite occurs associated with
the salt deposits of Stassfurt, Germany. It is also found near Carlsbad, New Mexico, western Texas,
and Saskatchewan, Canada.
22
Banded Iron Formations (BIFs)
The BIFs are iron ore deposits consisting of alternating bands of silica-rich cherts and iron-rich
layers. Individual layers are only millimeters to centimeters thick, although the entire formation can
be hundreds of meters thick. These formations, also known as banded ironstones or taconites, are
usually of Lower Proterozoic age (i.e., over 2 billion years old) and are distributed in the cratons or
stable shield areas of all continents (Fig. 3.6).
Figure 3.6. Worldwide distribution of banded iron formations (BIFs) (after Barnes, 1988).
The major minerals of BIFs include siderite, hematite, magnetite, chert, with lesser amounts of
calcite, ankerite, and iron silicate. A widely-held theory is that soluble ferrous iron released by
weathering under an oxygen-poor atmosphere over 2 billion years ago, precipitated out in shallow
seas as iron silicate, or as the ferric oxides-hematite and magnetite. The banding resulted from
seasonal variations in dissolved iron. Bacteria or other microorganisms may have been partially
involved in the deposition and formation of these minerals. Disagreement remains over the original
source of the large quantities of iron. Recent studies suggest that iron-rich solutions came from
oceanic hydrothermal springs and volcanic activity. Wherever the iron is highly concentrated, the
BIFs can be mined as economic ore deposits. Some important deposits include the Lake Superior
region, USA and Canada, Hamersley, Australia, the Labrador trough, Canada, Transvaal, South
Africa, and Brazil.
23
Recommended Reading
Barnes, J.W., 1988. Ores and Minerals: Introducing E conomic Geology. Open University Press,
Milton Keynes, U.K., Chapter 3, p. 24-42.
Davidson, G., 1995. After the gold rush. New Scientist, April 15, 1995, p. 26-31.
Gurney, J.J., Levinson, and Smith, H.S., 1991. Marine mining of diamonds off the west coast of
southern Africa. Gems and Gemology, Winter 1991, p. 206-216.
Keller, P.C., 1990. Gemstones and their Origins. Van Nostrand Reinhold, New York, Chap. 1,
p. 5-17.
24
CHAPTER 4
METAMORPHIC MINERALS
Introduction
Minerals can be transformed into new minerals by heat, pressure, or chemical reactions within the
earth's crust. At high temperatures and pressures, minerals will recrystallize in the solid state, without
melting. This recrystallization, or metamorphism, takes place at depths of 2-3 km down to 35-40
km, and at temperatures over 100EC. As rocks are subjected to these conditions, both their mineral
makeup and their physical appearance are altered. Metamorphism can take place over wide regions
(regional metamorphism) or directly adjacent to igneous instrusive bodies (contact metamorphism).
Regional Metamorphism
The increasing grades of metamorphism are revealed by a sequence of “indicator” minerals. Going
from mild to severe levels of metamorphism, these indicators are: chlorite-biotite-almandine garnetstaurolite-kyanite-sillimanite. Other minerals are diagnostic of very high pressures. For example,
kyanite (Al2SiO5) has the same chemical composition as andalusite and sillimanite, but it forms at
much higher pressures than its other two polymorphs. Jadeite, a sodium aluminum pyroxene, and
glaucophane, a purplish-blue sodium amphibole, are also indicators of high pressure, but relatively
low temperature conditions.
The minerals which actually form depend on both the initial composition and the metamorphic
regime to which the rock has been subjected. A nearly pure quartz sandstone will recrystallize to
quartzite; a pure calcite limestone will become marble. These metamorphic rocks are more coarsely
crystalline than their sedimentary precursors, but have essentially the same mineral content.
Figure 4.1 Schematic illustration of
approximate temperature and pressure
ranges of various types of metamorphic
rocks and associated minerals
More chemically-diverse rocks recrystallize to a variety
of new rocks, depending on the temperatures and
pressures they have experienced. For example, the
original mixture of clay minerals and fine-grained quartz
in shale gradually recrystallizes under progressively
rising temperatures and pressures to: 1) phyllite—a
mixture of sericite, chlorite, and quartz, 2) schist—micas
(muscovite, biotite), quartz, garnet, andalusite, kyanite,
sillimanite, staurolite, hornblende, feldspar, chlorite, 3)
gneiss—micas, feldspar, quartz, garnet, hornblende , 4)
granulite—highly recrystallized feldspar, quartz, garnet,
kyanite or sillimanite, and ultimately 5) a nearly-melted
migmatite (Fig. 4.1).
25
These mineralogical changes are accompanied by changes in the texture of the rock. The very-fine
grained clays and quartz in shale are compacted into slate. Application of pressure and heat
transforms slate into a fine-grained silvery or greenish-gray foliated (layered) phyllite, and then a
more coarsely-grained, foliated schist, in which the individual mineral grains are clearly visible. In
gneiss, the minerals are further segregated into dark and light layered bands. The granulite is
coarsely-grained but only slightly layered. The foliated or layered appearance of most metamorphic
rocks is due to the nearly parallel alignment of platy minerals such as chlorite or mica, or to the lineup of prismatic or bladed minerals such as hornblende, actinolite, or kyanite (Fig. 4.2). These
minerals are lined up because they have recrystallized under pressure at right angles to the main
direction of compression within the earth.
Figure 4.2. Kyanite, Brazil.
Almandine, the most common species of garnet, is characteristic of mica schists and other
metamorphic rocks derived from clayey sediments. It occurs in fine brownish-red crystals from
Zillertal and Oetztal, Austria. Large red crystals over an inch in diameter—dodecahedrons modified
by trapezohedrons—are embedded in mica schists at Wrangall, Alaska. In the northeast U.S., giant
almandine-pyrope crystals in amphibole schist from the Barton Mine, Gore Mountain, New York
are crushed for use in abrasives and sandpaper. The Green's Farm Garnet Mine, Roxbury
Connecticut is an old classic locality popular with regional collectors. Large dark wine red to black
almandine crystals, up to 1.5 in in diameter are associated with staurolite in the schist. Small
almandine crystals are scattered throughout the biotite schist that underlies much of Manhattan. The
largest almandine crystal ever collected in New York City weighs over 9 lbs and is 6 inches in
diameter (Fig. 4.3).
Figure 4.3. The Kunz Garnet. Almandine crystal, about 6" across, found in 1885 at West 35th Street
between Broadway and Seventh Avenue, New York City. The crystal is currently on display at the
American Museum of Natural History.
26
Impure limestones, starting as mixtures of calcite, dolomite, quartz, and clay are progressively
transformed into marble containing talc, then tremolite (a calcium magnesium amphibole), diopside
(calcium magnesium pyroxene), and ultimately forsterite (magnesium olivine).
The famous gem locality of Mogok, Myanmar (Burma) is an example of a regionallymetamorphosed deposit. The rubies occur in a coarsely crystallized marble, together with spinel,
diopside, phlogopite (mica), forsterite, chondrodite, sphene, and garnet. (Because of the nearby
presence of granitic intrusions, some geologists consider the Mogok deposit to be an example of
contact metamorphism—see below). Since 1992, Mong Hsu, a locality 250 km (150 miles) southeast
of Mogok, has become a major source of Burmese rubies. The overall suite of associated minerals,
including andalusite, almandine, and tremolite, suggests that the Mong Hsu marble crystallized at
lower temperatures than did the marble at Mogok. Other metamorphic deposits of ruby in marble
are found in Hunza Valley, Azad Kashmir, NW Pakistan, Jegdalek, Afganistan, and northern Viet
Nam (Fig. 4.4).
Figure 4.4. Ruby from metamorphic marble,
Hunza, Pakistan.
Figure 4.5. Lazurite on calcite,
Afganistan.
In rare cases, sulfur may be present in the limestone (possibly originating from evaporite deposits
containing gypsum). Metamorphism alters this to lapis lazuli—a rock consisting of deep-blue lazurite
(sodium calcium aluminum silicate sulfate), calcite, and pyrite (Fig. 4.5). Lapis has been mined in
Afganistan since antiquity. More recently, Chile has supplied lower-grade lapis in quantity.
Serpentinites, or rocks consisting largely of serpentine, are produced by the addition of water to
olivine-rich igneous rocks (e.g., dunites—all olivine, or peridotites—mostly olivine with some
pyroxene). They form under “retrograde” metamorphism, so-called because temperatures and
pressures are lower than when the original igneous rock crystallized. Serpentinites occur in Cornwall,
Quebec, Zermatt, Switzerland, Norway, New Zealand, Vermont, several western U.S. states, and
closer to home, on Staten Island, New York.
Both forms of jade—nephrite and jadeite—are products of regional metamorphism. Nephrite is a
variety of the amphibole series tremolite-actinolite. Pure tremolite is a nearly white mineral,
consisting of calcium magnesium silicate. With the addition of increasing amounts of iron, grading
into actinolite, the color ranges from pale grayish-green, shades of brown, yellow, dark spinachgreen, to nearly black (Fig. 4.6). In nephrite, the fibrous, almost felted, amphibole crystals are
compactly intergrown, forming an extremely tough and durable material suitable for carving, in spite
of the moderate Mohs hardness (H 6½).
27
Figure 4.6. Nephrite carving, China.
Figure 4.7. Jadeite carving, China
(The jadeite is originally from Burma).
Nephrite is typically associated with regionally metamorphosed rocks, such as schists, gneisses, and
serpentinites. Because of its toughness, nephrite is resistant to erosion and accumulates in riverbeds
as smoothed and rounded pebbles and boulders, some weighing up to several tons. A large polished
slab of dark green nephrite several feet long is on display at the American Museum of Natural
History.
Nephrite is fairly abundant throughout the world. The nephrite used in ancient Chinese jade came
from the region near Khotan, Yarkand, and the Kunlun Mts., in eastern Turkestan. Other, more
modern Asian sources include Taiwan and Korea. Major deposits also occur in New Zealand and
Siberia. Enormous quantities of spinach-green nephrite are found along the Fraser River, British
Columbia and the Kobuk River, Alaska. Much of this jade is shipped to Hong Kong or Taiwan for
carving into trinkets, which are then re-exported back to the U.S. Nephrite deposits occur elsewhere
in the U.S. in various localities in California and Wyoming.
Jadeite is more limited in its geographic distribution. Historically, the major source has been from
Tawmaw and Hpakon, northern Myanmar (Burma). In the Western Hemisphere, jadeite comes
from the Motagua River, Guatemala.
Jadeite is a member of the pyroxene group, composed of sodium aluminum silicate. In pure form, it
is white, but with traces of iron and other impurities it takes on various shades of green, yellow, red,
lavender, bluish-gray, and black (Fig. 4.7). The most highly-prized form of jadeite is a translucent,
vivid emerald-green, known as “imperial jade”, colored by traces of chromium. Jadeite differs from
nephrite in its greater hardness (H 7), higher specific gravity (~3.5), and a more granular texture.
Jadeite is produced by a form of high-pressure, low-temperature regional metamorphism that occurs
at depths ranging from 10 to 30 km and temperatures between 150E to 300EC. These conditions can
be achieved in zones of compression associated with plate subduction. At Tawmaw, Burma, jadeite
comes from dikes within a serpentinized peridotite. Most of the jadeite within the district, however,
is recovered as rounded boulders from recent alluvial deposits in the Uru River. The Guatamalan
jadeite has weathered out of serpentinite rock along the Motagua River fault zone—part of a major
fault system which defines the boundary between the North American and Caribbean plates. As in
Burma, most of the jadeite is found as large boulders in the streambed. Much of the recentlyrecovered jadeite is grayish green, in contrast to the prehistoric blue-green jades used by the Olmec
and Costa Rican cultures and the emerald-green jades favored by the Mayas.
28
Mica-schist beryl deposits are formed under a special set of circumstances. Beryllium-rich fluids
from a granitic (or pegmatitic) magma intrude into silica-poor metamorphic rocks, such as
amphibolite or serpentinite, producing beryl crystals. Under certain rare conditions, where
chromium or vanadium are also present, the beryl may incorporate these elements and become
emerald (Fig. 4.8). Such emerald deposits are found in Brazil, the Ural Mts., Russia, the ancient Red
Sea deposits, Egypt, Manyara, Tanzania, Swat Valley, Pakistan, and North Carolina. (The famous
Muzo emerald deposit, Colombia was formed by hydrothermal solutions interacting with
sedimentary rocks—see Chapter 6). The schist type of deposit is also host to the alexandrite variety
of chrysoberyl. However, because of the metamorphic conditions, internal cracks and inclusions are
abundant; thus, most of the material is usually not of gem quality.
Figure 4.8. Emerald from schist, Brazil.
Contact Metamorphism
The intrusion of a hot magma, or molten rock, into cooler enclosing rocks literally bakes the
margins of the country rocks at their contact zone (Fig. 4.9). Fluids containing volatile constituents
that emanate from the magma into the country rocks also alter these chemically. New minerals result
from the interaction of heat and fluids. Intrusion of a magma into a limestone produces a
skarn.This contact metamorphic rock often contains economically-useful minerals, such as
magnetite, cassiterite, scheelite, sphalerite, chalcopyrite, or attractive crystals of ilmenite, spinel, uvite
(tourmaline), diopside, vesuvianite, epidote, grossular, andradite, uvarovite, and scapolite.
Figure 4.9. Schematic illustration of contact metamorphism (after Robinson, 1994, p.166).
29
Recommended Reading
Betts, J., 1997. Accounts of the discovery of the Kunz garnet. Bulletin of the New York Mineralogical
Club, Dec. 1997, p. 9-10.
Desautels, P.E., 1986. The Jade Kingdom, Van Nostrand Reinhold, NY, 118p.
Harlow, G.E., 1991. Hard rock. Natural History, Aug. 1991, p. 4-10.
Kane, R.E. and Kammerling, R.C., 1992. Status of ruby and sapphire mining in the Mogok
Stone Tract. Gems & Gemology, Fall 1992, 152-174.
Keller, P.C., 1990. Gemstones and their Origins. Van Nostrand Reinhold, NY, Chaps. 6 and 7.
Kunz, G.F., 1892 [1968]. Gems and Precious Stones of North America. Dover Publications, Inc.
New York, pp. 83, 346.
Peretti, A., Schmetzer, K., Bernhardt, H.J., and Mouawad, F., 1995. Rubies from Mong Hsu. Gems
& Gemology, Spring 1995, p. 2-25.
Sinkankas, J., 1989. E meralds and Other Beryls. GeoScience Press. Prescott, AZ, 665p.
Yedlin, L.N., 1947. Garnet at Roxbury and W. Redding, Conn. Rock s & Minerals, 22(9), 824-826.
30
CHAPTER 5
MINERALS FROM COOL
SOLUTIONS
Caves and Cavities
Water circulating at or near the earth's surface can give rise to a variety of different minerals.
Rainwater contains small amounts of dissolved carbon dioxide from the atmosphere. Microorganisms in the soil also generate carbon dioxide. Carbon dioxide dissolved in water produces
carbonic acid, as in a carbonated drink. This weakly acid water seeps underground towards the water
table, gradually dissolving limestone (a rock made up mostly of calcite [calcium carbonate] or
dolomite [calcium magnesium carbonate]) along subsurface fissures and fractures. Over time, these
fissures are enlarged into caves or caverns (Fig. 5.1).
Figure 5.1. Schematic drawing of a cave.
Water slowly dripping from the roof of a cave that lies above the water table evaporates, causing
calcium carbonate to deposit icicle-like stalactites. Drops that reach the floor eventually build up
into mounds or pinnacles called stalagmites. A cross-section of a stalactite reveals a pattern of
radiating and concentric bands, similar to the annual growth rings of trees. This growth pattern is
also characteristic of many minerals with botryoidal and reniform habits (such as goethite, malachite,
smithsonite, etc.), and implies that they have deposited from water solutions into open cavities, at
relatively low temperatures.
Cave accumulations usually consist of calcite and aragonite, both of which are polymorphs of
calcium carbonate. Other minerals occur in caves less commonly. Some minerals reported from
caverns include other carbonates such as dolomite, magnesite, siderite, also nitrates (niter), iron
oxides (goethite, hematite), manganese oxides, phosphates (brushite CaHPO4 • 2H2O); dahllite
(carbonate-hydroxylapatite), silica (opal, chalcedony, quartz), and sulfates (selenite).
31
Ore minerals in carbonate rocks may have formed as migrating
hydrothermal solutions entered pre-existing caves and deposited their
load of metallic minerals by cooling and dilution with groundwater.
Alternatively, the ore-bearing solutions may have reacted with
carbonate rocks, leading to the precipitation of the metal sulfides. Both
processes could have contributed to the formation of the
carbonate-hosted Mississippi valley type lead-zinc deposits (see Chap.
6). Percolating groundwater later oxidized the pyrite-rich sulfide ores,
releasing sulfuric acid, which is more corrosive than carbonic acid.
Sulfuric acid dissolves iron, copper, manganese, and zinc, among other Figure 5.2. Botryoidal
malachite, Congo (Zaire).
metals. These reactions produce many secondary carbonates such as
siderite (iron carbonate, FeCO3), malachite [Cu2CO3(OH)2] and azurite [2CuCO3 • Cu(OH)2],
rhodochrosite (MnCO3), and smithsonite (ZnCO3) (Fig. 5.2).
Oxidation and Secondary Sulfide Enrichment
The primary metal sulfide minerals of hydrothermal veins are chemically altered by descending
groundwater. As noted above, pyrite is converted to sulfuric acid and ferric sulfate. This corrosive
mixture further attacks other metal sulfides, for example transforming chalcopyrite to copper sulfate.
The chemical breakdown of pyrite and other sulfides creates a wide variety of secondary minerals
within the oxidized zone above the water table (Fig. 5.3). Buried sulfide deposits were often
discovered because of the colorful reddish-brown and yellow staining or “gossan” at the surface,
formed by the oxidation of iron sulfides to iron oxides (e.g., goethite, limonite, hematite) and
leaching of other metals. In addition to the iron oxides, other common secondary minerals from the
oxidized zone include malachite, azurite, smithsonite, chrysocolla—a hydrated copper silicate,
cerrusite—lead carbonate (PbCO3), and hemimorphite—hydrated zinc silicate (Table 5.1).
Figure 5.3 Secondary alteration of a hydrothermal vein, showing various zones — the leached zone
(i.e. “gossan”, top), to the oxidized zone (middle), the enriched zone directly below the water table,
and the primary ores (bottom) (from J. Sinkankas, Field
Collecting Gems and Minerals, GeoScience Press, Inc., ©
1961, 1970, J. Sinkankas).
Groundwater, trickling down through fissures, crevices, and
cavities in aluminum-rich igneous rocks, such as granite,
trachyte, and monzonite porphyry (Fig. 1.2) oxidizes and
decomposes enclosed grains of copper sulfides and apatite.
The weathering process releases copper, iron, aluminum, and
phosphorus, which recombine and precipitate as turquoise
[CuAl6(PO4)4(OH)8 • 5H2O] in thin veins, seams, or nodules
near the water table (Fig. 5.4). In humid regions, the abundant
groundwater often carries away the chief ingredients of
turquoise. Therefore this sky-blue mineral usually tends to
form in more arid climates, such as the American Southwest,
Iran, Mexico, and Peru. In Hubei Province, China, turquoise has formed in a similar manner, except
that the host rocks are low-grade metamorphic carbonaceous- and silica-rich black slates.
32
Figure 5.4. Turquoise from Carico Lake, Nevada (left) and Sleeping Beauty Mine, Arizona (right).
Top quality, hard, pure turquoise ranges from deep sky-blue to green-blue. The softer, chalkier,
more porous varieties lose water and become a pale, dull blue. This chalky material is often treated
(or “stabilized”) with resins and plastics, creating a salable product that is extensively used in
contem-porary southwestern jewelry. If iron replaces some of the copper in turquoise, chalcosiderite
[CuFe6(PO4)4(OH)8 • 4H2O] forms instead. The more iron in the turquoise, the greener the stone.
Arizona Copper Minerals
Arizona is well known to collectors for the abundance and incredible variety of colorful
secondary minerals from the oxidized zone of porphyry copper deposits. Over 160 different
copper minerals have been found, mostly from the oxidized zone. These minerals span the entire
color spectrum from the bright blues of azurite, linarite, and kinoite, the greens of malachite,
antlerite, and brochantite, the aquas of chrysocolla, turquoise, and aurichalcite, the reds of cuprite
(var. chalcotrichite) and native copper, to the rare purple of anthonyite, and lemon-yellow of
beaverite.
Bisbee is world-famous for its spectacular specimens of azurite and malachite (Fig. 5A). It also
produced a large quantity of crystalline shattuckite. Ajo was the original source of ajoite and
papagoite. Ray and Ajo have also yielded lovely native copper crystals. Kinoite was first recovered
from the Christmas mine. Virtually all of the mines furnish some chrysocolla and turquoise, the
latter mostly of the soft, chalky variety. Gem-quality turquoise has been produced at the Bisbee,
Morenci, Kingman, and Sleeping Beauty mines (the latter three still active). Gem chrysocolla is
mined at Inspiration and Ray (Fig. 5B).
Figure 5A.Azurite and
malachite, Bisbee,
Arizona.
Fig. 5B. Chrysocolla,
Arizona
33
Table 5.1 Common secondary minerals from the oxidized zone.
Mineral
Composition
Habit
Colors
Adamite
Zn2(AsO4)(OH)
Botryoidal
Yellow, green, colorless
Aurichalcite
(Zn,Cu)5(CO3)2(OH)6
Radiating, acicular crystals
Blue-green
Azurite
Cu3(CO3)2(OH)2
Botryoidal,often with
malachite; bladed crystals
Blue
Cerussite
PbCO3
Bladed crystals, often
twinned
White
Chrysocolla
Cu silicate hydrate
Massive
greenish-blue
Crocoite
PbCrO4
Acicular, prismatic
Orange-red
Octahedron, cube,
dodecahedron
Ruby-red
Cuprite
Cu2O
(Chalcotrichite-variety)
Fine, acicular
Erythrite
Co3(AsO4)2 • 8H2O
Acicular, radiating crusts
Pink
Goethite
α-FeOOH
Botryoidal, stalagtitic
Brown-black
Hemimorphite
Zn4Si2O7(OH)2 • H2O
Mammillary, stalagtitic,
fibrous, radiating
White, blue-green, yellow
Lepidocrocite
γ-FeOOH
Botryoidal, fibrous
aggregates
Red, black-brown
Limonite
FeO • OH • nH2O
(mixture)
Botryoidal, stalagtitic,
earthy masses
Yellow-brown
Malachite
Cu2CO3(OH)2
Botryoidal, reniform,
banded
Green
Mimetite
Pb5(AsO4)3Cl
Hexagonal prisms
Yellow, brown
Pyromorphite
Pb5(PO4)3Cl
Hexagonal prisms
Green, brown
Rhodochrosite
MnCO3
Reniform, stalagtitic,
rhombohedral crystals
Pink
Siderite
FeCO3
Botryoidal, rhombohedral
Yellow,brown
Smithsonite
ZnCO3
Botryoidal, reniform,
stalagtitic
White, blue, green, yellow, pink
Turquoise
CuAl6(PO4)4(OH)8 • 4H2O
Massive, rarely crystallized
Green-blue
Vanadinite
Pb5(VO4)3Cl
Hexagonal prisms
Red-orange
Variscite
Al(PO4) • 2H2O
Massive, nodular
Greenish
Wavellite
Al3(PO4)2(OH)3 • 5H2O
Spherulitic, radiating
crystals
Pale green, yellow
Wulfenite
PbMoO4
Tabular, square outline
xtals
Orange-red
34
Copper sulfate trickles downward and reprecipitates as copper sulfides (chalcocite—Cu2S and
covellite—CuS) in the reducing (oxygen-deficient) environment at or below the water table (Fig.
5.3). The copper solutions also react with primary ore minerals (e.g., chalcopyrite CuFeS2) to
produce bornite Cu5FeS4. These high grade secondary sulfide enrichments were once
economically-important sources of copper ore, that had concentrated at the upper portions of
porphyry copper deposits, such as the porphyry copper mines in Arizona, New Mexico, Utah, and
Mexico. Nowadays, the low-grade primary copper ore (0.5-1.0%) is extracted from huge open-pit
mines such as Morenci, or Sierrita and Esperanza, south of Tucson, or Bingham Canyon, near Salt
Lake City, Utah.
Secondary (supergene) enrichment most commonly involves copper sulfides. Secondary silver recrystallizes as the native metal or the sulfide, acanthite (Ag2S). Lead minerals are generally too
insoluble to produce secondary enrichments, whereas zinc sulfate is so soluble that it is washed
away.
Fluctuating Water Tables
The deposition of Australian opal was related to fluctuations in the water table. Most of the precious
opal deposits occur in Cretaceous (140 to 65 million years ago) marine sediments. During the late
Cretaceous and early Tertiary periods, the shallow marine sea became a desert. Intense and deep
weathering of the quartz-rich sediments dissolved silica, which was carried downward and
precipitated in cavities, fissures, open pore spaces in the sediments, and even replaced fossil
mollusks, bone, and wood. Seasonal or interannual variations in climate caused the water table level
to rise or fall. During drier seasons or periods, when the water table was lower, evaporation of
groundwater led to the precipitation of silica as small, colloidal spherical particles. Because
conditions remained relatively stable throughout the deposition process, the spheres grew to a
remarkably uniform size with an orderly arrangement, analogous to the stacking of oranges or apples
in a fruit stand, or the atoms within a crystal lattice. The spheres range in size between 0.15 to 0.3
micrometers (or 1.5-3.0 ten-thousandths of a millimeter).
The uniform stacking of the silica spheres into a three-dimensional array acts as a diffraction grating
(Fig. 5.5). The resulting diffraction of light produces the characteristic play-of-color in precious opal
(Fig. 5.6). Ordinary opal, on the other hand, has randomly ordered, variably-sized spheres.
Figure 5.5. Diffraction of light in precious opal.
Figure 5.6. Boulder opal, Australia.
35
Other occurrences of opal, as in Mexico or the western U.S., form in a similar manner, but are
found in vesicles and veins in siliceous volcanic rocks. Opal from volcanic sources tends to be more
prone to dehydration, resulting in crackling or “crazing”, which makes it less stable and less suitable
for jewelry than the Australian stones.
Recommended Reading
Bideaux, R.A. and Wallace, T.C., 1997. Arizona copper. Rock s & Minerals. 72(1), 10-27.
Hill, C.A. and Forti, P., 1986. Cave Minerals of the World. National Speleological Society,
Huntsville, Alabama, 238p.
Keller,P.C., 1990. Gemstones and their Origins. Van Nostrand Reinhold, NY, Chap. 2.
36
CHAPTER 6
HYDROTHERMAL MINERALS
Introduction
As water migrates through deep sedimentary basins, it warms from the earth's internal heat.
Descending water is also heated up as it comes in contact with deeply-buried magma chambers.
Deformation and metamorphism of sedimentary and other rocks releases any enclosed water, which
is also warmed up by tectonic processes (see Figure 1.5). Hot solutions are collectively termed
hydrothermal solutions. Hydrothermal temperatures range between 50EC and 200EC (122E392EF) for low temperature (epithermal) to over 400EC for high-temperature vein deposits
(hypothermal).
This chapter describes several characteristic types of hydrothermal deposits, progressing from
relatively low to high temperature occurrences.
Crystals Lining Cavities and Pockets
As basaltic magma cools at the earth's surface, rising gas bubbles may “freeze” in place, producing a
porous, vesicular texture, much like the holes in a Swiss cheese. These rounded, hollow vesicles or
vugs provide open spaces for migrating solutions. Groundwater warmed by moving through cooling
lava flows leaches out metals and silica from volcanic ash and glass (Fig. 6.1). Crystals precipitate in
these vugs from the cooling hydrothermal solutions, forming geodes. Zeolites usually deposit at
temperatures below 200EC. They commonly line vugs and cavities in basaltic lavas. Display-quality
zeolite crystals (e.g., natrolite, mesolite, scolecite, stilbite, chabazite, heulandite) and associated
minerals (e.g., prehnite, apophyllite, pectolite) have been collected from basalt in the Watchung
Ridges, Paterson, N.J. and the Deccan trap rocks, Poona, India (Fig. 6.2).
Figure 6.1. Crystallization within open cavities in lava flows (Robinson, 1994).
37
Figure 6.2. Natrolite,
Paterson, New Jersey.
Figure 6.3. Native copper from the
Keweenaw Peninsula, Michigan.
Although less famous than Paterson, the Millington Quarry, Somerset County, New Jersey yields
attractive specimens of zeolites and related minerals. The quarry is located in the southwest part of
the Hook Mountain Basalt, the youngest of three basaltic lava flows that constitute the Watchung
Mountains. The mineralization occurs within large amygdules (almond-shaped cavities) in the basalt.
Prehnite, apophyllite, datolite, and calcite are among the most common minerals at this locality.
Prehnite is found in yellowish-green botryoidal clusters associated with calcite and apophyllite.
Apophyllite forms rosettes of green, pink, or clear crystals. Natrolite occurs in acicular, radiating
translucent crystal clusters up to 5 cm long. Pectolite grows in small pinkish-reddish botryoidal
masses.
Amethyst is another mineral which commonly lines basaltic vugs. Large “cathedrals” (arch-shaped
cavities) from Brazil and Uruguay are covered with deep purple crystals. Native copper (and silver)
from the Keweenaw Peninsula, Michigan occurs as amygdaloidal deposits in basalt, replacing matrix
in overlying conglomerates and shales, and in fissures (Fig. 6.3).
Other common geode minerals (e.g., quartz, calcite, barite, pyrite, and marcasite) are found in
sedimentary formations. The cryptocrystalline (very fine-grained) varieties of quartz—chalcedony,
agate, chert, flint, jasper—occur as nodules in sedimentary rocks or in lava flows.
Mississippi Valley Type Deposits
Solutions, originally derived from rainwater that has seeped into the earth, may travel considerable
distances through porous sediments or open fractures in rocks. As these solutions descend to great
depths within downwarped structures, called basins, they are heated and dissolve soluble salts and
metal ions (Fig. 6.4). The solutions may cool as pressure is reduced, when they enter open cavities or
move upward. In addition, they may react chemically with other solutions or with the rocks they
traverse. Such changes in physical conditions may lead to precipitation of the dissolved minerals in
veins or in open spaces between rock fragments. Examples of such deposits are the lead and zinc
ores of the Mississippi Valley region and elsewhere.
38
Figure 6.4. Migration of solutions from a basin into carbonate rocks, (after Barnes, 1988).
The Mississippi Valley is host to major deposits of lead and zinc sulfides in limestones and
dolomites. In addition to galena and sphalerite, other typical minerals include calcite, barite, fluorite,
with lesser amounts of chalcopyrite, pyrite, and marcasite. The mineralization in Mississippi Valley
Type (MVT) deposits is usually stratabound (confined to rock strata) in nearly horizontal carbonate
sedimentary rock formations. Minerals are localized along faults, veins, and in solution collapse
structures (e.g., collapsed caves, sinkholes). The mineralization is epigenetic (i.e., has formed later than
the enclosing rocks) and is thus unrelated to the formation of the sedimentary host rocks.
MVT minerals have deposited from warm (100-140EC), saline solutions. The salt-rich brines picked
up many metallic elements from organic-rich shales. Such dark shales are often enriched in metals
such as copper, lead, zinc, and iron. At higher temperatures, lead and zinc are soluble in salt
solutions and are readily mobilized. The brines also contain oil droplets and methane. These fluids
thus closely resemble “formation waters” that transport petroleum to its reservoirs. The hot brines
eventually encounter carbonate rocks where they move along joints, cavities, and collapse breccias
(broken rock). The metals are precipitated as the sulfides, in the presence of a suitable source of
sulfur (Fig. 6.5), for example: 1) hydrogen sulfide (a poisonous gas known for its rotten egg odor)
generated by the bacterial decomposition of gypsum (CaSO4 • 2H2O) or anhydrite (CaSO4), and 2)
chemical reaction of sulfate minerals with organic compounds or with methane gas.
Figure 6.5. Deposition of minerals in a MVT deposit. Metal sulfides precipitate when the dissolved
metals encounter a sulfur-rich source (after Robinson, 1994).
39
Missouri Minerals
The historic Tri-State district of Missouri, Kansas, and Oklahoma was once a major source of lead
and zinc in the U.S. Cherty limestones dissolved by heated solutions produced breccias that were
favorable sites for mineral deposition. Geologic structures, such as faults, fractures, and grabens
(downdropped blocks between two parallel faults) also controlled the migration of mineralizing
solutions, and also deposition of the ore minerals.
The primary minerals in this area are sphalerite, galena, chalcopyrite, pyrite, marcasite, dolomite, and
calcite. Sphalerite occurs as “black jack” (dark color with a high iron content), as “ruby jack” (in
transparent reddish-amber colored crystals), and also in botryoidal and stalactitic masses. Galena
appears in large cubic and octahedral crystals. Secondary lead minerals (cerussite, anglesite, and
pyromorphite) formed by the oxidation of galena are also common. Attractive calcite crystals
displaying a large variety of habits have also been collected from this famous district.
Since World War II, lead production has shifted to the Viburnum Trend in southeastern Missouri.
The Viburnum Trend consists of about a dozen lead-zinc deposits
lined up in a nearly north-south trend, around 45 miles long, situated
on the west flank of the Ozark uplift. This ore belt is located along
“knobs” representing ancient islands and a band corresponding to
the axis of an ancient stromatolite reef. (Stromatolites are laminated
calcareous accumulations deposited by blue-green algae). The ore
minerals are concentrated within the reef complexes, in faultbounded solution collapse breccias, and in calcareous sandstones.
Collectible minerals from this district include large, lustrous cubes
and octahedrons of galena, some up to 8" across, sphalerite with
Figure 6.6. Botryoidal
chalcopyrite, calcite crystals, many twinned, and also some showing
marcasite from the
marcasite phantoms, also botryoidal, colloform, and stalagtitic pyrite
Viburnum Trend, Missouri. and marcasite (Fig. 6.6).
Illinois-Kentucky Fluorite
The Illinois-Kentucky mining district was originally worked for galena and
sphalerite, but has been a major producer of fluorite since 1900. However,
the last mine closed in 1995, unable to compete against new sources of
fluorite from China, with vast reserves of this mineral and cheap labor costs.
The fluorite from the Illinois-Kentucky district comes from sedimentary
rocks, largely limestones of middle to late Mississippian age (325-345 million Figure 6.7. Fluorite
years old). Around 50 narrow, highly altered dikes and sills of Permian age and calcite, Cave-in(c. 250-290 million years old), cut across the strata. The fluorite deposits
Rock, Illinois.
from this district occur as veins, replacements, and breccia pipes along
faults. The fluorite forms large, well-developed cubes of many colors (Fig. 6.7). Crystals are
frequently zoned and between 5-9 stages of crystallization have been described. The earliest crystals
are yellow, followed by white, blue, and purple in multiple zones. Green and pink fluorite are rarer.
Bitumen (an asphaltic or tarry material) and petroleum inclusions are common
40
Beautiful specimens collected from this district include barite and golden yellow calcite on purple
fluorite, white calcite on golden yellow and blue fluorite, calcite crystals exhibiting various habits,
galena associated with purple fluorite, yellow to red and black sphalerite crystals. It also occurs
banded together with fluorite and with barite. Other primary minerals are chalcopyrite, marcasite,
and pyrite. Some important secondary minerals are celestite, smithsonite, and witherite.
Closer to home, the Lime Crest Quarry, in Sparta, N.J., a popular New York Mineralogical Club
collecting destination, may be related to MVT deposits. W. Cummings describes veins of dolomite
and calcite, with pyrite, minor galena, sphalerite, barite, and fluorite—a typical MVT mineral
assemblage. In common with other MVT deposits, the minerals at Lime Crest probably deposited
from hot solutions moving along faults and in open cavities in carbonate rocks.
Colombian Emerald Deposits
Most emeralds are found in biotite schists (e.g., Miko, Zambia; Swat Valley, Pakistan; Ural
Mountains, Russia; Brazil; Egypt; Austria). These deposits formed during periods of regional
metamorphism (see Chapter 4). Beryllium, originally derived from granitic pegmatites, was localized
at the boundary with silica-poor rocks such as amphibolites or serpentinites (which were
metamorphosed from gabbros and peridotites). The latter two rock types supplied chromium and
vanadium, two elements essential for the development of the green color in emerald (beryl,
Be3Al2(Si6O18)).
By contrast, the beautiful, intense green emeralds of Colombia were deposited by hot, salty brines,
traveling through organic-rich black shales. These briny solutions resemble those that formed the
Mississippi Valley-type (MVT) deposits. Studies of fluid inclusions show that the brines were at least
330EC. These hot, salty solutions reacted chemically with the black shales and limestones, leading to
the formations of emeralds (Fig. 6.8).
Figure 6.8. Schematic diagram showing how Colombian emeralds were formed.
41
Black shales are high in carbonaceous matter (i.e., pyrobitumen). Organic material is often enriched
in heavy metallic elements such as chromium, vanadium, copper, nickel, cobalt, and beryllium. In
particular, the elements chromium, vanadium, and beryllium, together with silica and aluminum, are
the key ingredients for making emeralds. The sulfate carried in solution by the hot brines was
originally derived from evaporites. Upon encountering the organic matter in the black shale, the
sulfate was reduced chemically to sulfur (S) and pyrite (FeS2, Fig. 6.8). In turn the organic matter was
oxidized to carbon dioxide (CO2), which then reacted with calcium ions to precipitate calcite
(CaCO3). The oxidized zones, depleted in the dark organic matter, have been bleached to a greywhite color, locally called cenicero (Sp. ash). The emeralds are concentrated in veins that radiate from
the cenicero zones, but are usually not found within the cenicero zones (Fig. 6.9). Typically they are
associated with pyrite and veins of calcite and dolomite (Fig. 6.10). Accessory minerals include
quartz, parisite, albite, apatite, fluorite, and barite.
Figure 6.9. Cross-section of an emerald vein at Muzo, Colombia. Albite, crystallizing first, lines the
vein at the contacts. It was succeeded by tabular or fibrous calcite, and finally well-crystallized
emerald and blocky calcite. Pyrite is distributed throughout the vein (after Ottaway et al., 1994).
Figure 6.10. Emerald crystals with pyrite
and calcite from Muzo, Colombia
42
Figure 6.11. Three-phase inclusion
in emerald from Colombia.
Colombian emeralds owe their intense blue-green color to the removal of iron as pyrite (Figs. 6.8,
6.10). Chromium and vanadium not only color emeralds green, but also produce a red fluorescence
that intensifies the color. However, iron impurities incorporated into the beryl crystal structure,
common in emeralds from most other localities, tend to suppress this fluorescence. Thus,
Colombian emeralds, free of iron, unlike their counterparts from biotite schists, are intensely bright
green.
Emerald crystals from Colombia reveal evidence of their geologic origins. They typically contain
two- and three-phase inclusions, consisting of the brine, a CO2-rich gas phase, and in the case of the
3-phase inclusion, crystals of halite (Fig. 6.11). Solid inclusions typical of Muzo include pyrite,
calcite, and parisite. At Chivor, characteristic solid inclusions are pyrite, quartz, albite, and goethite.
Emeralds are often clouded by numerous fractures and tiny fluid inclusions (“jardin” or garden
[Fr.]). Trapiche emeralds contain black carbonaceous inclusions radiating from the center in a sixspoked pattern, when cut perpendicular to the c-axis. Some trapiche crystals have clear emerald
spokes within an opaque albite-emerald mixture.
Alpine Cleft Minerals
The Alps were formed, some 60-30 million years ago, as the result of the collision of two continental
plates: the Eurasian and the African Plates. This crunching of plates led to the overthrusting and
stacking of several thick piles of rocks called nappes (Fr. tablecloth, sheet) (Fig. 6.12). At least four
distinct nappes are recognized in the central Alps. This tectonic activity recrystallized and
metamorphosed any pre-existing rocks. The deformation also opened up clefts or fissures in the
rocks (Fig. 6.13). These fissures and faults provided pathways for the migration of hydrothermal
solutions, that were enriched in salts and carbon dioxide. Temperatures were around 400-500EC and
pressures equivalent to 4,000-5,000 times the pressure of the atmosphere at the earth's surface. As
the Alps were uplifted, the solutions cooled, pressures were reduced, and eventually minerals such as
quartz were precipitated out. Toward the end of the crystallization process, a very-fine- grained
dusting of chlorite coated the earlier-formed minerals.
Figure 6.12. Geologic cross-section of Binntal area, Switzerland, showing stacking of nappes in the
central Alps (after Graeser, 1998).
43
Figure 6.13. Formation of Alpine clefts, due to stresses on rocks. The vertical arrows indicate the
direction of greatest compression; the smaller horizontal arrows show the direction of least
compression (or extension). This results in opening of gashes, as rocks are drawn into sausage-like
sections (left), or are broken (faulted; center). As two slabs of rock are dragged past each other, “en
echelon”, or staggered, tension gashes open up (right) (after Graeser, 1998).
Clefts in the granite and gneiss massifs of the central Alps have
yielded spectacular clusters of quartz crystals, both as clear and
smoky varieties, and occasionally, amethyst. Associated with quartz
are adularia (var. orthoclase feldspar), albite, fluorite (usually as pink
octahedrons), also rare beryllium minerals (milarite, phenacite,
bavenite) and hematite (in the form of “eisenrose” or iron rosettes),
titanite (sphene), anatase, brookite, purple apatite. Habits of Alpine
quartz crystals are quite distinctive. They include: gwindel—crystals
that are bent and elongated on the a-axis, fenster—skeletal quartz,
also quartz with chlorite phantoms, rutilated quartz, scepter quartz,
and morion—almost black smoky quartz (Fig. 6.14).
Figure 6.14. Gwindel quartz,
Swiss Alps.
Rocks from the Penninic area of southern Switzerland and northern
Italy were especially severely folded and deformed. These rocks
consist of upthrusted portions of the oceanic floor, the lower crust,
and the upper mantle. These mafic source rocks are higher in iron
and magnesium, but lower in silica. Thus, unlike the alpine cleft
deposits, quartz is not common. Instead, these rocks have yielded
good specimens of chlorite, diopside, epidote, garnet, tremoliteactinolite, and serpentine.
A famous mineral locality from this zone is the Binn valley (Binntal), Switzerland. It has produced a
wide variety of rare and unique base metal sulfides and sulfarsenides. Some minerals unique to the
Binn area include jordanite (lead arsenic sulfide), lengenbachite (lead, silver, copper arsenic sulfide),
sartorite (lead arsenic sulfide), cafarsite, asbecasite, and graeserite.
Ore Veins
Many important metal deposits occur in veins, which are directly or indirectly related to igneous
activity. Hot fluids and some metals may be released directly from the cooling magma. The heat
44
given off by the magmatic intrusions is the energy source which drives hydrothermal circulation
within the earth's upper crust. Magmatic fluids may mix with deeply-circulating groundwater, which
then scavenges metals from the host rocks and redeposits them elsewhere in veins or lodes at some
distance from the igneous intrusion.
The relation of vein deposits to igneous activity is illustrated schematically in Figure 6.15. A
magmatic intrusion is shown at depth (gray, shaded area). The porphyry copper (and molybdenum)
deposits, such as at Morenci and the many other copper deposits in Arizona, are formed near the
top of the intrusive or at the base of the volcano, within several kilometers of the surface. Hightemperature tin and tungsten veins also form at depth. Closer to the surface, near the volcano, are
gold-bearing copper and iron sulfides, somewhat similar to the porphyry copper deposits. Further
from the volcano, hot springs (geothermal systems) emerge at or close to the surface, having
become considerably diluted by groundwater. Some gold, silver, and mercury deposits are associated
with such hot spring activity.
Figure 6.15. The relation of hydrothermal veins to igneous activity.
Geologists still debate the relative contributions to the formation of ore minerals made by primary
magmatic fluids or descending meteoric solutions that have been heated by the cooling magmas.
Clues as to the nature of the mineralizing solutions are furnished by studying variations in the
isotopes of certain elements, such as oxygen, hydrogen, and sulfur. These isotopic fluctuations can
be used as “fingerprints” to pinpoint the sources of the solutions and the extent to which solutions
from different sources have mixed during the precipitation of the ore minerals. In addition, analysis
of fluid inclusions in vein minerals such as quartz or calcite, can tell us much about the temperature,
pressure, and salt content of the solutions from which the host crystals grew.
A vein or lode is a tabular or flat, lens-like body that cuts across the enclosing rock, usually showing
well-defined walls. Veins are typically composed of quartz, calcite, dolomite, barite, fluorite, and
metal sulfides. Topaz, tourmaline, garnet, and other rarer minerals may occur in high-temperature
veins (Table 6.1).
45
Table 6.1. Minerals of vein deposits (after Barnes, 1988).
Hypothermal
E pithermal
Gold
Cassiterite
Wolframite
Bismuthinite
Gold
Chalcopyrite
Bornite
Molybdenite
Grey Coppers
Galena (with Silver)
Sphalerite
Gold
Acanthite
Ruby silvers
Cinnabar
Stibnite
Galena (with Silver)
Sphalerite
Realgar & Orpiment
Pyrrhotite
Arsenopyrite
Magnetite
Hematite (Specular)
Garnet
Topaz
Tourmaline
Pyroxene
Amphibole
Quartz
Pyrrhotite
(Arsenopyrite)
Marcasite
Siderite
Rhodochrosite
Other Carbonates
Adularia
Quartz
Quartz & Chalcedony
Ore Minerals
Gangue Minerals
Mesothermal
Vein minerals fill fissures, fractures, or shear zones in rocks. They also cement breccias (broken or
shattered rock). Well-terminated crystals fill veins that have formed at shallow depths, and grow into
open cavities or “pockets”, commonly known as vugs. Veins are closely-related to geological
structures such as faults or folds. Thus, they can be vertical, horizontal, or dipping at an angle, or
follow folds within the host rocks. Veins may occur singly, as sub-parallel clusters, in overlapping or
staggered groups (en echelon), and in intersecting stockworks.
Hydrothermal vein deposits are classified into three groups, based on temperature conditions
(Table 6.1):
1. Epithermal veins formed at temperatures below 300EC (572EF).
2. Mesothermal veins formed at temperatures between 240EC and 400EC (464E-752EF).
3. Hypothermal veins formed at temperatures above 400EC (>752EF).
Epithermal Veins
Epithermal minerals have deposited from solutions, largely groundwater, warmed by intrusions at
depth, which then continue upward to the surface at silicious hot springs (Figure 6.15). As pressure
is reduced near the surface, the solutions start to boil, precipitating their load of metals. The
expansion of rock caused by the boiling may have contributed to the brecciation (shattering) often
seen in this type of deposit.
Many precious metal deposits are epithermal: the Comstock Lode, Nevada, Cripple Creek,
Colorado, and much of the Mother Lode, California. Characteristic mineral associations include:
native gold, gold-silver alloys (electrum), gold tellurides (calaverite, petzite) or selenides, ruby silvers
(proustite-silver sulfarsenide and pyrargyrite—silver sulfantimonide), also stibnite, cinnabar, and
occasionally tennantite and tetrahedrite. Mercury and antimony are also commonly epithermal.
46
Mesothermal Veins
Mesothermal veins are major sources of copper, lead, and zinc deposits, and also some gold. These
veins are often associated with granitic or dioritic intrusions (see Fig. 1.2 for mineral content of
rocks). Ore solutions have mixed with groundwater. Some examples of mesothermal vein deposits
include zinc-lead deposits and Devon and Cornwall, U.K., cobalt-silver deposits, Ontario, Canada,
and gold at Farncomb Hill, Breckenridge, Colorado.
Hypothermal Veins
Major minerals are cassiterite (SnO2) (e.g., the famous Cornwall tin deposits), wolframite ([Fe, Mn]
WO4), and native gold. Hypothermal veins are commonly found above dome-like granitic
projections. The cooling granite expels water and silica-rich fluids, rich in volatiles such as boron,
fluorine, also tin, tungsten, iron, other base metals, arsenic, and sulfur. These fluids then deposit in
fractures overlying the crystallizing mass.
One well-studied hypothermal vein deposit is the Yankee Lode, eastern Australia, which is mined
for cassiterite. Fluid inclusions in vein quartz have been studied using sophisticated laboratory
instruments (Figure 6.16). The figure shows how fluid inclusions from successive zones in the
growing quartz crystal reveal a progressive cooling and dilution of the magmatic brines with
groundwater, which has led to the precipitation of the ore minerals. The temperature of 453EC
(847EF) and a salinity of 35% at the base of the crystal has dropped to 220EC (428EF) and salinity of
only 0.1% near the top. The cassiterite, with associated ilmenite, tourmaline, and muscovite, began
to deposit near the middle of the growth sequence.
Figure 6.16A. Longitudinal section through a
quartz crystal from the Yankee Lode tin
deposit. Shown are successive growth zones
and selected fluid inclusions from various
growth stages. The inclusions record the
evolution from a magmatic fluid to a
groundwater-dominated system. Thtot is the
homogenization temperature, or temperature
when the water and gas phases merge.
Figure 6.16B. Some actual inclusion are shown.
The small inset shows scars made by the laser
used to drill open and extract fluid for analysis
by a very sensitive mass spectrometer (after
Audetat and others, 1998).
47
Recommended Reading
Audetat, A., Günther, D., and Heinrich, C.A., 1988. Formation of a magmatic-hydrothermal ore
deposit; insights with LA-ICP-MS analysis of fluid inclusions. Science, 279, 2091-2094.
Becker, M.A., 1998. Recent mineral finds at the Millington quarry, Somerset County, New
Jersey. Rock s & Minerals, Sept./Oct. 1998, 320-324.
Cummings, W., 1985. Mineralization at the Millington Quarry, New Jersey. Rock s & Minerals,
60, 213-218.
Cummings, W., 1993. A Mississippi Valley Type Lead-Zinc vein of probable Paleozoic age at
the Lime Crest Quarry, Sparta, New Jersey. The Pick ing Table, v. 34 (1), p. 14-16.
Goldstein, A., 1997. The Illinois-Kentucky Fluorite district. The Mineralogical Record, v. 28, Jan.Feb., 1997.
Graeser, S., 1998. Alpine minerals. Rocks & Minerals, Jan./Feb. 1998, p. 14-32.
Missouri Issue, Rock s & Minerals, Nov./Dec. 1997.
Ottaway, T.L. and others, 1994. Formation of the Muzo hydrothermal emerald deposit in
Colombia, Nature, June 16, 1994, p. 552-554.
Robinson, G.W., 1994. Minerals: An Illustrated E xploration of the Dynamic World of
Minerals and Their Properties, pp. 77-85.
Peters, J., 1984. Triassic traprock minerals of New Jersey. Rock s & Minerals, 59, 157-182.
Sinkankas, J., 1989. E meralds and Other Beryls. GeoScience Press, Prescott, AZ, 665p.
48
CHAPTER 7
BLACK SMOKERS—MINERALS
FROM THE DEEP SEA
A New World Revealed
The world of the deep sea is the final frontier and yields numerous surprises—many revealed for the
first time within the last 20 years. The ocean floor hosts a wealth of minerals as well as strange life
forms. Gold and other precious metal deposits will soon be mined from the ocean bottom. A
portion of a “black smoker” will form the centerpiece of the new Hall of E arth exhibit at the
American Museum of Natural History (AMNH) in New York City by the summer of 1999. This
mineralized pillar was hauled from 2000 (6,600 feet) depth by a joint U.S.-Canadian expedition to
the Juan de Fuca Ridge, off Vancouver Island in July, 1998.
In 1978, the French submersible Cyana, while exploring the sea floor 2.6 km (8,600 feet) deep along
the East Pacific Rise at 21EN latitude encountered orange- and green-stained rocks that turned out
to be deposits of copper, zinc, and iron sulfides. A year later, active hydrothermal vents were
discovered on an expedition with the Woods Hole submersible ALVIN to the central volcanic zone
of the East Pacific Rise rift valley. Great billowing clouds of black “smoke” were pouring out of the
vents at 350EC (662EF), leading one of the expedition scientists to exclaim that “It was like
Pittsburgh in 1925, with all those blast furnaces going full tilt!” (Fig. 7.1). The smoke actually
consists of tiny particles of hot, newly-precipitating metallic sulfides that build up hollow tapered
spires or chimneys, the size and shape of tree trunks. These superheated, discharging hydrothermal
vents are popularly known as “black smokers”. They are perched on top of larger dome-like
structures, which represent accumulations from countless episodes of chimney construction and
destruction.
Figure 7.1. Black smokers discharging high-temperature fluids and metal sulfides at vents on the
Juan de Fuca Ridge.
The first deep-sea vents were found at mid-oceanic rifts in the Pacific Ocean, where the oceanic
plates are moving apart relatively rapidly—at rates of several inches a year (or approximately the
same rate as fingernails grow). Therefore it was believed that such vents could only occur in the
49
Pacific Ocean where the sea floor spreading rates are up to ten times faster than in the Atlantic
Ocean. Yet, indications of hot-spring mineralization were noted in the Atlantic Ocean as early as
1972. These initial suspicions were confirmed during an expedition with ALVIN in 1985. Activelydischarging black smokers were discovered on top of a mound the size and shape of the Houston
Astrodome, composed of massive metal sulfide deposits. Hydrothermal mineral occurrences are
now known from over 100 localities in the Pacific and Atlantic Oceans. Buoyant plumes of hot,
metal-rich brines have recently been detected in the Indian Ocean along a slow-spreading ridge.
A major surprise was the discovery of unique ecosystems that thrive in oases in the abyssal darkness
at the hydrothermal vents, at so-called “barbeque pits”. The vent creatures include giant tubeworms,
up to 1 meter (3.3ft) long, large clams and mussels the size of dinner plates, crabs, shrimp, heatloving worms attached to the smoker chimneys, and some wandering fish (Fig. 7.2). Active vents are
often located by going up the “crab gradient”.
Figure 7.2. Creatures dwelling at the hydrothermal vents.
Submarine Hydrothermal Vents
Hydrothermal vents are found on the ocean bottom in areas where active volcanism provides a
source of heat and where seawater can penetrate into cracks in the lava. Most submarine volcanoes
and hot springs are concentrated along a globe-encircling chain of underwater mountains 47,000
miles long. Here, molten lava spews out of the mantle and pushes the earth's plates apart a few
inches a year. Most of this volcanic activity is confined to a rather narrow zone a few miles wide in
the central rift valley of the mid-oceanic ridge system (Fig. 7.3).
Figure 7.3. Schematic diagram of mid-ocean ridge where magma upwells from the mantle and
pushes the earth's plates apart. Hydrothermal activity is concentrated at the summit graben—a
narrow rift zone at the mid-ocean ridge. Cold seawater seeps through the heavily-fractured rocks
and is heated, before returning to the sea floor at black smoker vents (after Haymon and
Macdonald, 1985).
50
The hydrothermal circulation is driven by the heat from the
upwelling molten magma (over 1,200EC or 2190EF). The
newly-solidified lavas are strongly fractured and faulted by
the ongoing seismic activity that accompanies the volcanism.
Seawater penetrates through these fissures and cracks,
heating up as it approaches the magma chamber at depth. As
the seawater descends and warms, it reacts chemically with
the surrounding basaltic rocks, losing some constituents and
gaining others (Fig. 7.4). Further reactions occur as the now
hot and buoyant brine ascends through the severely
shattered rocks. Magnesium from the seawater combines
with silica and releases hydrogen ions, thereby increasing the
acidity of the solution. Sulfate in seawater, in turn, is reduced
to sulfur, by reacting with ferrous iron in the basalt,
producing hydrogen sulfide gas (H2S). This hot, acidic brine,
rich in chlorine (from dissolved sea salt) is very corrosive and capable of leaching out many metals,
such as copper, iron, zinc, manganese, lead, silver, and gold from the basaltic rocks. These hot fluids
discharge from hydrothermal vents on the sea floor, forming black smokers (Fig. 7.1).
Figure 7.4. Submarine hydrothermal circulation. Descending seawater carries sodium, chlorine,
magnesium, and sulfate. As it heats up, it dissolves various metals and silica from the basaltic rocks.
The hot solutions discharge at black smoker vents on the sea floor (after Rona, 1986).
Minerals from Black Smokers
As the turbulent clouds of metal-rich brines emerge from the vents,
they mix with the cold surrounding seawater and precipitate fine black
particles of copper, iron, and zinc sulfides. The tiny particles
suspended in the buoyantly rising black plumes gradually settle out
and build up hollow chimney-like pillars up to 10 meters (33 feet) tall.
These chimneys are fairly fragile—several have been knocked over
accidentally during deep sea expeditions with ALVIN. However, they
can regenerate within several years at actively-discharging vents.
The concentric zoning and banding of minerals seen in cross-sections
of black smokers chimneys are caused by the sharp changes in
temperature and brine concentrations within a few inches to feet near
the active vents. At the East Pacific Rise (21EN, 109EW) calcium
sulfate (anhydrite) concentrates in the cooler, outer walls of chimneys,
with iron and zinc sulfides (pyrite, sphalerite, wurtzite). Moving
inward toward the hotter interior, one encounters a zone of
intermixed sphalerite and pyrite. Copper sulfides (chiefly chalcopyrite) line the hottest (>350EC, >662EF) inner portion of the
chimney (Fig. 7.5).
Figure 7.5. Mineral zoning at a black smoker chimney. In stage 1, anhydrite precipitates with a
mixture of sphalerite and pyrite. In stage 2, chalcopyrite precipitates in the hotter interior of the
chimney (after Haymon and MacDonald, 1985).
51
“White smokers” form at somewhat lower temperatures (300EC or 572EF). They consist of
amorphous silica, sulfur, pyrite, and barite. They are densely populated by attached worms, crabs,
and other creatures (Fig. 7.2).
The mineral zonation differs at other hydrothermal vents. On the Juan de Fuca ridge, west of
Vancouver Island and Washington State, dendritic and colloform sphalerite and silica are the first
minerals to precipitate. These are replaced by higher-temperature pyrite and marcasite on the outer
walls. Some sulfide specimens appear to have had an even more complex growth history. A
specimen from the East Pacific Rise initially crystallized as fine-grained pyrrhotite, which was later
replaced by several stages of sulfide and opal deposition.
The mineralogy of the deep-sea vents consists of metal sulfides, sulfates, native elements, silicates,
and oxides (Table 7.1). The most abundant sulfides are octahedral sphalerite, cubes and framboidal
masses of pyrite, botryoidal chalcopyrite, and also hexagonal plates of wurtzite (a polymorph of
sphalerite). Minor constituents include marcasite, galena, bornite, digenite, cubanite, chalcocite, and
covellite.
Anhydrite is widespread at active chimneys; gypsum less common. Barite rosettes together with
amorphous silica fill the walls of fossilized tubeworms. Native sulfur is found in dead, sealed
chimneys and also together with the barite and silica in the tubeworms.
Ore Deposits in the Making
Scientists working with the Ocean Drilling Program have recently drilled through two activelyforming massive sulfide deposits, one in the middle of the Atlantic Ocean at around 26EN latitude
and the other in the northern part of the Juan de Fuca Ridge, west of Vancouver Island.
At the TAG hydrothermal field, 3650 m (12,000 feet) deep in the median valley of the Mid-Atlantic
Ridge at 26EN, episodic hydrothermal activity over a 20,000 year period has created a mound, 200 m
(660 feet) in diameter and 50 m (165 feet) high, containing an estimated 4 million metric tons of
copper, zinc, and iron sulfides. This figure lies well within the range of typical land-based, volcanichosted massive sulfide deposits, such as those found in Cyprus, Oman, and Newfoundland.
Moreover, the mineralogy and structure of the TAG mound closely resembles those of its terrestrial
counterparts. The most famous of these deposits on land are the copper deposits from the Troodos
Mountains of Cyprus, which have been exploited since antiquity. (In fact, the word “copper” derives
from the name of Cyprus).
The Juan de Fuca Ridge deposit at 48E 26'N latitude, 128E40'W longitude, is a 35 meter (116 feet)
high mound, covered by oxidized metal sulfide rubble and sediments. The major minerals present
consist of pyrrhotite with less abundant isocubanite (a copper-iron sulfide), sphalerite, and wurtzite.
Much of the deposit has recrystallized into a mixture of pyrite and magnetite. This deposit is at least
100 meters (330 feet) thick and grades downward into a “feeder zone” of intensely mineralized
veins. At around 200 m (660 ft) below the sea floor, a roughly horizontal copper-rich layer (up to
16% Cu by weight) is encountered (Fig. 7.6). This “deep copper zone”, or DCZ, formed during
periods of hydrothermal activity when copper-rich solutions flowed into permeable layers underlying
more impermeable, silicified horizons. This ore deposit is conservatively estimated at 8.8 million
metric tons, not counting the copper-rich zone or the mineralization associated with a second,
nearby mound.
52
Table 7.1. Minerals from hydrothermal vents at the East Pacific Rise, 21EN (after Haymon and
Kastner, 1981).
Mineral Group
Sulfides
Sulfates
E lements
Oxides and
Oxyhydroxides
Silicates
Mineral Species
Relative Abundance
Habits
Sphalerite
Major
Pyrite
Chalcopyrite
Major
Major
Wurtzite
Marcasite
Pyrrhotite
Minor
Minor
Minor
Galena
Bornite
Cubanite
Chalcocite
Digenite
Covellite
Trace
Trace
Trace
Trace
Trace
Trace
Anhydrite
Major (Active Chimneys)
Gypsum
Caminite**
Barite
Minor (Active Chimneys)
Minor (Active Chimneys)
Minor
Jarosite
Minor
Sulfur
Minor
Radiating and Reticulated
Goethite
Limonite
Minor
Minor
n.d.
Amorphous, Colloform
Clusters
Silica
Minor
Talc
Nontronite
Trace
Major clay in ridge crest
sediments
Globules; Irregular Platy
and Tabular
n.d.
n.d.
Cube, Dodecahedron,
Tetrahedron, Twinned,
Massive
Cube, Framboids, Massive
Botryoidal, Massive,
Tetrahedrons
Hexagonal Platelets
n.d.*
Hexagonal Plates, Radiating
Clusters
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
Tabular, Radiating,
Acicular, and
Cryptocrystalline
Radiating, Acicular, Tabular
Microcrystalline
Cross-Cutting Tabular
Crystals and Rosettes
Colloform
* n.d.= no data.
** Caminite is Mg7(SO4)5(OH)4.H20. It was first discovered at black smokers, hence the name, from the Latin for chimney.
Figure 7.6. Cross-sectional views of the mineralized zones at the Juan de
Fuca hydrothermal deposit. A copper-rich zone (the “deep copper zone”)
underlies the massive sulfide mounds. Hydrothermal fluids vented
vigorously at drill holes 1035F and 1035H after being drilled (Zierenberg et
al, 1998).
The black smoker retrieved for the AMNH exhibit came from the
Endeavor segment of the Juan de Fuca Ridge, to the south of this deposit.
It was one of four chimneys hauled from a forest-like vent field 2,250
meters (7,400 feet) below the ocean surface, around 300 km west of
Washington State and Vancouver Island (47E55'N, 129E06'W). Two of the
53
smokers were still “sizzling” hot as they were brought on deck. Segments up to 1 ½ meters (5 feet)
long were recovered. The smokers were enclosed in specially-designed cages and sawed off at their
bases. Once firmly secured to the cages by cables, they were gently hoisted to the deck of the
awaiting research vessel.
At the Endeavor site, fragments of the outer smoker crust revealed beautiful crystals of “root beercolored” sphalerite, glistening chalcopyrite, and black wurtzite, as well as amorphous silica in an iron
sulfide matrix. Rusty stains of iron oxide coated many pieces. The smokers can regrow at amazing
rates. A stump from one of the chimneys that had been sawed off grew a beehive-shaped formation
of anhydrite and sulfides over 1 meter in height in less than one week!
Deep-Sea Bonanza
Hot volcanic spring deposits, a mile deep near Papua New Guinea, hold a fortune in gold, silver,
copper, and other metals. The ore deposits, discovered in 1991 and 1993 by scientists from Australia
and Canada, may contain billions of dollars worth of metals. Ore samples assay up to 26% zinc, 15%
copper, several ounces of silver per ton, and an ounce of gold per ton—assays far richer than those
of most land-based deposits. A Papua New Guinea company (the Nautilus Minerals Corp.) has
staked a claim to 1,974 square miles of the Bismarck Sea, off Papua New Guinea, and plans to begin
mining within the new few years.
An extinct submerged volcano off Lihir Island, in Papua New Guinea (around 3E20'S, 152E40'E)
may also soon be mined for gold. The seamount rises more than 600 m (~2000 ft) above the ocean
floor to a height of 1050 m (~3500 ft) beneath mean sea level. Gold-rich precipitates, assaying up to
43 parts per million gold, cover the basaltic lavas. This deposit lies only 25 km south of the worldclass Ladolam gold deposit on Lihir Island. It appears to be very similar to epithermal vein deposits
on land.
Exploitation of these ocean-bottom metal deposits raises a slew of technical and environmental
questions. The high costs of deep-sea mining and the technological challenges of raising the ore to
the surface may be offset by the high ore grade, the potentially large areas of metal deposits, and the
absence of environmental regulations—the latter a significant concern when mining on land. On the
other hand, a rich and unique assemblage of creatures live around the active hydrothermal vents.
Any mining activity will have to proceed carefully, so as not to destroy this unique biology and
habitat.
Denizens of the Deep
Scientists, descending in submersibles to the lightless abyss for the first time, were astounded to find
the abundance of life-forms surrounding the hydrothermal vents. Thriving in the relatively warm
water of the thermal springs (at least 52EF, as compared to the usual near-freezing temperatures at
the ocean bottom) were giant clams up to a foot across, huge mussels, white and yellow crabs,
tubeworms several feet long, shrimp, starfish, sea lilies, sponges, snails and sea cucumbers, and an
occasional octopus or fish (Fig. 7.2). These abyssal oases were later called “clambakes”. The deepsea creatures, first encountered only 20 years ago, belong to an entirely different and unique
ecosystem—one existing entirely independently of sunlight as a source of energy.
54
Unlike plants, which use the energy from sunlight to manufacture food by photosynthesis, the deepsea food chain is based on chemosynthesis, or the conversion of chemical energy into organic
molecules. Hydrogen sulfide, emitted in copious quantities in the vent fluids, releases considerable
energy when it oxidizes. Bacteria harness this oxidation energy, using it to convert inorganic carbon
in the form of carbon dioxide or bicarbonate, dissolved in seawater, into organic compounds. The
sulfide, in turn, is transformed into sulfate, much of which then precipitates out as calcium sulfate,
or anhydrite. The giant tubeworms and clams depend symbiotically on these bacteria for their food.
In fact the tubeworms have no digestive systems at all and would starve if not for the bacteria!
Some of the more bizarre discoveries include eyeless shrimp that somehow “see” in the dark, worms
directly attached to hot vent chimneys, and most puzzling of all—a creature the size of a sand dollar
that resembles a Chinese checkerboard!
The eyeless shrimp (Rimicaris exoculata—or “dweller of the rift without eyes”), discovered in 1985
by Peter Rona and colleagues, at the TAG hydrothermal field in the Atlantic, display bright
reflective patches on their backs (Fig. 7.7). Investigations by biologist Cindy Lee Van Dover and
others determined that these patches contained a light-sensitive pigment—rhodopsin—which
absorbs light most strongly in the blue-green part of the spectrum. This pigment is also present in
eyes of both vertebrates and invertebrates. Somehow, these shrimp are detecting low levels of
light—but what kind of light? After all, sunlight does not reach these depths—several thousands of
meters below sea level. Could the shrimps be “seeing” the heat glow from the thermal vents? The
glow from the 350EC vent plumes is mostly reddish, whereas the shrimps are most sensitive to bluegreen light. The shrimp “eye” may be well-adapted to detect active vents at long distances. Seawater
quickly absorbs the redder, longer wavelengths from the heat glow, leaving only a weak greenishblue light which the shrimp patches can spot.
Figure 7.7 (a) side and (b) back views of Rimicaris exoculata—the “eyeless” shrimp. Shaded area
indicates the location of the unique “eyes”, visible as bright spots in the dark (after Cindy Lee Van
Dover, 1988).
The Pompeii worms (Alvinella pompejana) live directly attached to “cooler” smoker chimneys (Fig.
7.8). Temperatures at the inner end of the worm tube average around 68EC (154EF), going as high
as 81EC (178EF). The outer end of the tube is around 20EC (68EF). This extreme temperature
difference of up to 60EC (110EF) from one end to the other is a biological record. Furthermore, this
worm seems to tolerate the highest temperatures known for any multi-cellular organism.
55
Figure 7.8. The Pompeii worm, Alvinella pompejana, 6 cm (2.4 in) long.
This worm colonizes the sides of active deep-sea hydrothermal vents.
Temperatures are up to 81EC (178EF) at its inner end and around 20EC
(68EF) at its outer end.
Scientists are eagerly studying the close connections that exist between the
deposition of the mineralized chimneys and the microorganisms which
form the base of the food chain at the luxuriant hydrothermal vent oases.
The bacteria are able to survive in highly toxic environments inside the
vents—poisonous hydrogen sulfide laden with heavy metals, at extremely
high temperatures. These “extremophiles” may hold clues to the origin of life and may provide
chemicals that could help cure diseases. The microorganisms include some of the most primitive
creatures at the base of the evolutionary tree. The hydrothermal vent environment represents a
popular scenario for the origin of life. All of the necessary ingredients are present: a source of heat,
energy, chemicals, and an environment shielded from the destructive hailstorm of asteroidal impacts
that devastated the earth's surface in its infancy. Commercial ventures are already underway to
harvest enzymes from symbiotic bacteria associated with the Pompeii worms. These chemical may
provide new pharmaceuticals and biocatalysts for many industrial applications.
56
Recommended Reading
Black Smokers and Hydrothermal Vent Minerals
Broad, W., 1997. First move made to mine mineral riches of seabed. The New York Times, Dec.
21, 1997.
Broad,W., 1997. Undersea treasure, and its odd guardians. The New York Times, Dec. 30, 1997.
Haymon, R.M. and Kastner, M., 1981. Hot spring deposits on the E ast Pacific Rise at 21EN:
preliminary description of mineralogy and genesis. E arth and Planet. Science L etters, 53, 363-381.
Haymon, R.M. and Macdonald, K.C., 1985. The geology of deep-sea hot springs. A merican
Scientist, 73, 441-445.
Humphris, S.E., and others, 1995. The internal structure of an active massive sulfide deposit.
Nature, 377, 713-715.
Rona, P., 1986. Mineral deposits from sea-floor hot springs. Scientific A merican, 254, 84-92.
Rona, P., 1988. Metal factories of the deep sea. Natural History, 97, 52-57.
Rona, P. 1992. Deep-sea geysers of the Atlantic . National Geographic. Oct., 1992, p. 105-109.
Rona, P. and Scott, S.D., 1993. A special issue on sea-floor hydrothermal mineralization: new
perspectives. Preface. E conomic Geology, 88, 1935-1976.
Zierenberg, R.A. and others, 1998. The deep structure of a sea-floor hydrothermal deposit.
Nature, 392, 485-488.
Deep-Sea E cosystems
Cary, S.C., Shank, T., & Stein, J., 1998. Worms bask in extreme temperatures. Nature, 391, 545-6.
Flanagan, R., 1997. The light at the bottom of the sea. New Scientist, 42- .
Travis, J., 1993. Probing the unsolved mysteries of the deep. Science, 259, 1123-1124.
Tunnicliffe, V., 1992. Hydrothermal-vent communities of the deep sea. A merican Scientist, 80,
336-349.
Van Dover, C.L., 1988. Do ‘eyeless’ shrimp see the light of glowing deep-sea vents? Oceanus,
31, 47-52.
For a first-hand account of the retrieval of the Endeavor black smokers, visit the website:
www.ocean.washington.edu/outreach/revel/subintroductiontext.htm.
57
Notes
58
CHAPTER 8
MINERAL SUCCESSION
Sequence of Deposition
When minerals crystallize from a melt or from solution, some minerals appear earlier in the
sequence of deposition and are followed by others. The order of crystallization can be reconstructed
by examining which minerals overlie or replace others in the field, or in hand specimens, or by
viewing thin sections of rock under the microscope. The sequence of deposition, or paragenesis,
offers important clues as to the mode of origin of the deposit. Knowing the range of physical and
chemical conditions over which major mineral phases are stable, one can infer the circumstances
leading to deposition. Mineral succession is one important reason why certain minerals typically
occur together. This information is also useful in the search for new ore deposits.
The most obvious case of mineral succession is when one mineral has visibly grown on top of
another. This deposition sequence is most apparent where crystals were able to grow unimpeded
into open spaces, such as veins, vugs, or pegmatite pockets. Some common examples include cubes
of galena growing on dolomite and sphalerite from Joplin, Missouri, blades of selenite on calcite
over amethyst in basaltic vugs from Brazil, quartz over tetrahedrite on pyrite from Pachapaqui, Peru
(Fig. 8.1), or calcite on prehnite from Millington quarry, New Jersey (Fig. 8.2).
Figure 8.1. Quartz on tetrahedrite over pyrite, Pachapaqui, Peru.
Figure 8.2. Calcite on prehnite, Millington quarry, New Jersey.
59
At Millington, the mineralization is concentrated in amygdaloidal layers within basalt. The
mineralization is divided into five stages, which roughly progress from the lower to the higher
stratigraphic levels of the deposit. These stages are: 1) the saline period, 2) the quartz period, 3) the
prehnite period, 4) the zeolite period, and 5) the
calcite period (Fig. 8.3). Following solidification of
the basalt, hydrothermal solutions penetrated
relatively porous and permeable amydaloidal
horizons in the basalt. Anhydrite was one of the
earliest minerals to precipitate (period 1), but was
replaced by chalcedony and drusy quartz (period 2).
Prehnite, datolite, and apophyllite deposited over
the quartz (period 3). Zeolites such as analcime,
stilbite, natrolite, also pectolite and apophyllite grew
over prehnite and datolite (period 4). Calcite
crystallized over the earlier-formed minerals such as
prehnite (Fig. 8.2), apophyllite, and zeolites (Period
5).
Figure 8.3. Paragenetic sequence at Millington
quarry, New Jersey.
The minerals of veins or pegmatites crystallize inward from the surrounding wallrock. In granitic
pegmatites, the outer margins are usually a finer-grained assemblage of feldspar, quartz, and
muscovite. Further in, one encounters coarser-grained quartz, muscovite and/or biotite, and perthite
(an intergrowth of microcline and albite). The last minerals to solidify from the magmatic fluid have
grown unimpeded into a miarolitic cavity and formed large crystals of quartz, microcline,
cleavelandite (a platy variety of albite), and occasionally tourmaline, beryl, lepidolite, also uranium,
columbite and others (see Figs. 2.7, 2.8). Cookeite, a late-stage lithium chlorite, frequently coats
tourmaline, cleavelandite, and quartz.
In igneous rocks, large crystals (phenocrysts) generally appear earlier
in the crystallization sequence than does the finer-grained
groundmass. A rock that contains phenocrysts is called a porphyry
(Fig. 8.4). As a rule, euhedral grains bounded by crystal faces deposit
before anhedral grains, lacking crystal faces. Often, several mineral
phases may precipitate more or less simultaneously, as shown by the
interlocking of grains in some granites.
Figure 8.4. Porphyry.
On the scale of a single crystal, zoning records the growth history of the crystal, much like tree rings.
The bands of varying color or composition mark successive changes in the composition and
temperatures of the fluids during growth. Fluorite, amethyst, corundum, and tourmaline are
frequently zoned (Fig. 8.5).
60
Figure 8.5. Zoning in tourmaline (liddicoatite, Madagascar).
Fluid inclusions taken from successive zones in a quartz crystal
from a tin mine in Australia have been linked to the deposition of
ore minerals (Fig. 6.16). Fluid inclusions incorporated into the base
of the crystal record a temperature of 453ºC and a salinity of 35%,
while near the top, temperatures had dropped to 220ºC and a
salinity of only 0.1%. Cassiterite, together with ilmenite,
tourmaline, and muscovite began to deposit near the middle of the
quartz growth, as shown by the presence of solid inclusions and
changes in the chemistry of the included fluids. This progressive
lowering of temperature and salinity indicates that the precipitation of the ore minerals was caused
by cooling and dilution of magmatic fluids with colder, less saline groundwater.
Mineral Replacements
Frequently, as physical or chemical conditions change during the course of mineral deposition,
earlier-formed minerals are replaced by later ones. The replacement may appear as a coating or
veining. In the latter case, solutions preferentially alter the host phase along cracks or cleavage
planes.
The hydrothermal alteration that accompanies porphyry copper deposition illustrates the
consequences of mineral replacement (Fig. 2.11). Magmatic fluids reacting with rocks at the core of
the intrusion convert the original igneous minerals to an assemblage of potassium feldspar and
quartz. This innermost zone is surrounded by a zone of clay and sericite alteration, and further from
the intrusion, by a widespread zone of chlorite, epidote and carbonates. The latter two zones are the
product of meteoric water penetrating through the igneous intrusion, being heated, and rising by
convection.
The oxidized zone of porphyry copper and vein metal deposits furnishes additional examples of
mineral replacements (Chap. 5). Primary copper, iron, molybdenum, lead and zinc sulfides were
originally deposited from high-temperature hydrothermal solutions at depth. Tectonic uplift and
erosion have subsequently brought these deposits close to the earth's surface. Descending meteoric
water has oxidized and leached the primary ores, forming a host of secondary oxides, carbonates,
arsenates, and phosphates. (Table 5.1). The original sulfide minerals have been nearly totally
replaced. In a few instances, however, remnants of the primary ores still survive. One finds cerussite
or anglesite on galena, cuprite over native copper, or smithsonite on sphalerite. A common
alteration sequence seen in specimens from Arizona copper mines is chalcopyrite or bornite,
followed by azurite, now totally replaced by malachite. The malachite was succeeded in turn by
chrysocolla, the latter topped by a coating of drusy quartz.
Other mineral replacements include the surficial oxidation and tarnishing of native silver to
argentite, or goethite on pyrite. In the zone of secondary sulfide enrichment, the primary
chalcopyrite and bornite are replaced by secondary chalcocite and covellite.
61
Pseudomorphs
A pseudomorph is a mineral which preserves the crystal habit of a former mineral. The
pseudomorph represents an extreme stage of mineral replacement, where only the outward form of
the pre-existing mineral still remains. Some common examples of pseudomorphs include goethite
after pyrite, malachite after azurite, and hematite after magnetite. Less common pseudomorphs are
turquoise after apatite (Fig. 8.6), native copper after aragonite (Fig. 8.7), and calcite after
halite—where the original salt cube has been distorted to a rhombohedron to accommodate the
calcite. Fossils are frequently replaced by minerals: opal after calcite or aragonite in mollusks, agate
or jasper after wood (Fig. 8.8), pyrite in brachiopods and ammonites (Fig. 8.9), and less commonly,
vivianite in belemnites.
.
Figure 8.6. Pseudomorph of turquoise after
apatite, Bacuachic, Sonora, Mexico.
Figure 8.7. Pseudomorph of native copper after
aragonite, Coro Coro, La Paz Dept., Bolivia.
Pseudomorphs form in several different ways. In substitution, the original mineral is replaced by
another, without any chemical reaction between the two. Examples are silicification in petrified
wood and lithification of other types of fossils. In alteration, the new material has altered or
chemically transformed the original, as for example in the change of anhydrite (CaSO4) to gypsum
(CaSO4 • 2H2O), or galena (PbS) to anglesite (PbSO4). Traces of the former mineral may still survive
at the core of the pseudomorph. In encrustation, a crust of one mineral is deposited over another. A
common example is drusy quartz encrusting fluorite or barite. Sometimes, the first-formed mineral
has completely dissolved, but evidence of its former presence is revealed by the surviving casts.
Figure 8.8. Petrified wood, Arizona.
62
Figure 8.9. Pyritized ammonite.
Recommended Reading
Barnes, J.W., 1988. Ores and Minerals: Introducing E conomic Geology, Open University Press,
Philadephia, pp. 62-63.
Becker, M.A., 1998. Recent mineral finds at the Millington Quarry, Somerset County, New
Jersey. Rock s & Minerals, 73, 320-324.
Cummings, W., 1985. Mineralization at the Millington quarry, New Jersey. Rocks & Minerals,
60 (5), 213-218.
Robinson, G.W., 1994. Minerals: An Illustrated E xploration of the Dynamic World of
Minerals and their Properties, Simon & Schuster, New York, 207p.
63
Notes
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Appendix
Geologic Setting
Important Minerals and Gems
Igneous Activity
Intrusives
Rock-forming minerals
Pegmatites
Porphyry copper
Plagioclase, orthoclase, microcline, quartz, mica, pyroxene,
amphibole (hornblende), olivine, garnet
Quartz, feldspar, mica, tourmaline, beryl, topaz, spessartine,
chrysoberyl, kunzite, columbite, euclase, monazite, fluorapatite,
amblygonite
Chalcopyrite, bornite, molybdenite, (gold)
E xtrusives (volcanism)
Lava flows
Basalt (tholeiite)
Alkali basalt
Rhyolite
Kimberlite
Rock-forming minerals as above, but fine-grained
Plagioclase, pyroxene, quartz, +/-olivine (peridot)
Plagioclase, pyroxene, olivine, zircon, ruby, sapphire
Same rock-forming minerals as granite; topaz, red beryl
Diamond, Cr-diopside, pyrope, chromite, phlogopite, olivine, spinel,
ilmenite
Sedimentary Processes
Placer deposits
Residual deposits
Evaporites
Banded iron formations
Gold, cassiterite, diamond, ruby, sapphire, zircon, rutile,
garnet
Bauxite (aluminum oxides), iron oxides, nickel-laterite
Halite, gypsum, anhydrite, sylvite, borax, trona, epsomite, glauberite
Hematite, magnetite, siderite, chert
Metamorphism
Regional metamorphism
Contact metamorphism
(e.g., skarn)
Chlorite, biotite, almandine, staurolite, kyanite, sillimanite, andalusite,
hornblende, tremolite-actinolite (nephrite), jadeite
Magnetite, cassiterite, scheelite,
sphalerite, andradite, ilmenite, spinel uvite, diopside, vesuvianite,
scapolite, rhodonite, wollastonite
Deposition from Solution
Meteoric solutions
Caves and caverns
Oxidized zone of
ore deposits
Secondary sulfide
enrichment
Fluctuating water tables
Calcite, aragonite, siderite, goethite, rhodochrosite, smithsonite,
pyrite, marcasite
Goethite, limonite, azurite, malachite,crocoite, cuprite, erythrite,
pyromorphite, turquoise, vanadinite, variscite, wavellite, wulfenite
Chalcocite, covellite, bornite, silver,acanthite
Opal, some agate, chalcedony
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Hydrothermal solutions
Geodes and vugs
Mississippi Valley
type deposits
Alpine clefts
Colombian emeralds
Black smokers
Metallic veins
Epithermal
Mesothermal
Hypothermal
Biominerals
Amethyst, zeolites, agate, chalcedony
Fluorite, calcite, sphalerite, galena, pyrite, marcasite,barite, cerussite
Quartz (clear, smoky, amethyst), adularia (orthoclase), fluorite (pink),
hematite, phenacite, titanite, apatite, anatase, brookite
Emerald, pyrite, calcite
Chalcopyrite, sphalerite, pyrite, pyrrhotite, (galena), anhydrite, barite,
sulfur, silica
Gold, cinnabar, stibnite, acanthite, adularia, milky quartz, chalcedony
Chalcopyrite, bornite, enargite, galena, silver, gold, quartz
Cassiterite, wolframite, molybdenite, bismuthinite, pyrrhotite,
magnetite, topaz, tourmaline, garnet, quartz
Calcite, aragonite, opal, pyrite, hydroxyapatite, sulfur, goethite,
magnetite, ferrihydrite
Minerals Beyond E arth
Meteorites
Irons
Stones
Stony-irons
Moon
Mars
Stardust
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Iron-nickel (kamacite, taenite), troilite, schreibersite, cohenite
Pyroxene, olivine, feldspar, troilite, nickel-iron
Iron-nickel, troilite, olivine, pyroxene
Calcic plagioclase, pyroxene, olivine, ilmenite
Pyroxene, feldspar, iron oxides, scapolite(?), palagonite, nontronite
(?), sulfates, chlorides
Graphite, diamond, moissanite, corondum, spinel, hibonite
About the New York Mineralogical Club, Inc
On September 21, 1886, in the home of Professor Daniel S. Martin at 236 West 4th Street, the New
York Mineralogical Club was formed through the efforts of George F. Kunz, B.B. Chamberlin and
Professor Martin. Monthly meetings followed at the homes of members, with hosts presiding.
George F. Kunz was elected Secretary. At the sixth meeting in March 1887, the name “New York
Mineralogical Club” was officially adopted. A constitution and bylaws were approved at the eighth
meeting. Since no president was required by this constitution, none was elected until April 1895
when George F. Kunz became the club's first president, an office he held for many years. At the end
of the club's first year, it had a membership of forty-six, including several notable mineralogists.
The club's fine collection of more than 700 mineral specimens from New York City is housed at the
American Museum of Natural History. The collection includes specimens of beryl, chrysoberyl,
garnet, tourmaline, stilbite and xenotime, together with many other species. During the 112 years the
club has been in existence several honorary members have been appointed in recognition of
contributions in the field of minerals and mineralogy. They include: Joseph Arons, Sir William
Henry Bragg*, Russ Buckingham, Lawrence H. Conklin, Madam Marie Curie*, Edward S. Dana*,
Clifford Frondel, Victor Goldschmidt*, Richard Hauck, Paul Kerr* Carl Krotki, Alfred Lacroix*,
Charles Palache*, Frederick Pough, Waldemar T. Schaller*, Leonard J. Spencer*, Ernest Weidhaas*,
Herbert P. Whitlock*. (*Deceased)
Minerals named after former and current club members include: austinite, bementite, bideauxite,
brianite, bostwickite, cahnite, canfieldite, charlesite, cliffordite, eglestonite, englishite, frondelite,
hauckite, hiddenite=green spodumene, holdenite, kempite, keyite, kunzite=lavender spodumene,
ludlockite, montgomeryite, mosesite, nealite, overite, petersite, perloffite, poughite, roeblingite,
schallerite, segelerite, sinkankasite, sklodowskite, spencerite, stenhuggarite, strunzite, whitlockite, and
yedlinite.
Currently the club has a membership of over 250. Monthly meetings, most with a guest lecturer, are
held on the second Wednesday of each month (excepting July and August) at the American Museum
of Natural History, New York City. Meetings are open to the public. Frequent field trips are
organized to interesting mineral localities in the vicinity of New York City and extended trips to
more distant states. Each month a bulletin is published for members and guests with news,
announcements and original articles about minerals and gemstones contributed by club members.
The New York Mineralogical Club is affiliated with the Eastern Federation of Lapidary and
Mineralogical Societies and the American Federation of Mineral Societies.
The club is dedicated to increasing interest in the science of mineralogy through the collecting,
describing and displaying of minerals and associated gemstones. Anyone interested in mineralogy or
gemstones is welcome as a member. Inquiries should be addressed to: New York Mineralogical
Club, Inc., P.O. Box 77, Planetarium Station, New York, NY 10024-0077.
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About the Author
Vivien Gornitz grew up in Queens and New York City, but spent a few formative years in
Switzerland, where exposure to the Alpine scenery stimulated a life-long interest in geology and
mineralogy. She studied chemistry at Barnard College, and mineralogy/geology at Columbia
University under Profs. Paul F. Kerr and Ralph J. Holmes, both former Club members. While
current research is focused on sea levels rise and coastal impacts of climate change, she still retains a
keen interest in minerals, gems and planetary exploration. She also enjoys working with stones and
setting them in silver.
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