classification scheme of mineral deposits

Earth-Science Reviews 100 (2010) 1–420

Contents lists available at ScienceDirect

Earth-Science Reviews j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / e a r s c i r ev

The “chessboard” classification scheme of mineral deposits: Mineralogy and geology from aluminum to zirconium Harald G. Dill Institute of Geosciences, Gem-Materials Research and Economic Geology, Johannes-Gutenberg-University Mainz, D-55099 Mainz, Becherweg 21, Germany

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Article history: Received 6 July 2008 Accepted 23 October 2009 Available online 18 November 2009 Keywords: economic geology mineral deposits geology mineralogy classification scheme spreadsheet

a b s t r a c t Economic geology is a mixtum compositum of all geoscientific disciplines focused on one goal, finding new mineral depsosits and enhancing their exploitation. The keystones of this mixtum compositum are geology and mineralogy whose studies are centered around the emplacement of the ore body and the development of its minerals and rocks. In the present study, mineralogy and geology act as x- and y-coordinates of a classification chart of mineral resources called the “chessboard” (or “spreadsheet”) classification scheme. Magmatic and sedimentary lithologies together with tectonic structures (1-D/pipes, 2-D/veins) are plotted along the x-axis in the header of the spreadsheet diagram representing the columns in this chart diagram. 63 commodity groups, encompassing minerals and elements are plotted along the y-axis, forming the lines of the spreadsheet. These commodities are subjected to a tripartite subdivision into ore minerals, industrial minerals/rocks and gemstones/ornamental stones. Further information on the various types of mineral deposits, as to the major ore and gangue minerals, the current models and the mode of formation or when and in which geodynamic setting these deposits mainly formed throughout the geological past may be obtained from the text by simply using the code of each deposit in the chart. This code can be created by combining the commodity (lines) shown by numbers plus lower caps with the host rocks or structure (columns) given by capital letters. Each commodity has a small preface on the mineralogy and chemistry and ends up with an outlook into its final use and the supply situation of the raw material on a global basis, which may be updated by the user through a direct link to databases available on the internet. In this case the study has been linked to the commodity database of the US Geological Survey. The internal subdivision of each commodity section corresponds to the common host rock lithologies (magmatic, sedimentary, and metamorphic) and structures. Cross sections and images illustrate the common ore types of each commodity. Ore takes priority over the mineral. The minerals and host rocks are listed by their chemical and mineralogical compositions, respectively, separated from the text but supplemented with cross-references to the columns and lines, where they prevalently occur. A metallogenetic-geodynamic overview is given at the bottom of each column in the spreadsheet. It may be taken as the “sum” or the “ mean” of a number of geodynamic models and ideas put forward by the various researchers for all the deposits pertaining to a certain clan of lithology or structure. This classical or conservative view of metallotects related to the common plate tectonic settings is supplemented by an approach taken for the first time for such a number of deposits, using the concepts of sequence stratigraphy. This paper, so as to say, is a “launch pad” for a new mindset in metallogenesis rather than the final result. The relationship supergene–hypogene and syngenetic–epigenetic has been the topic of many studies for ages but to keep them as separate entities is often unworkable in practice, especially in the so-called epithermal or near-surface/shallow deposits. Vein-type and stratiform ore bodies are generally handled also very differently. To get these different structural elements (space) and various mineralizing processes (time) together and to allow for a forward modeling in mineral exploration, architectural elements of sequence stratigraphy are adapted to mineral resources. Deposits are geological bodies which need accommodation space created by the environment of formation and the tectonic/geodynamic setting through time. They are controlled by horizontal to subhorizontal reference planes and/or vertical structures. Prerequisites for the deposits to evolve are thermal and/or mechanical gradients. Thermal energy is for most of the settings under consideration deeply rooted in the mantle. A perspective on how this concept might work is given in the text

E-mail address: [email protected]. URL: http://www.hgeodill.de. 0012-8252/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.earscirev.2009.10.011

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H.G. Dill / Earth-Science Reviews 100 (2010) 1–420

by a pilot project on mineral deposits in Central Europe and in the spreadsheet classification scheme by providing a color-coded categorization into 1. 2. 3. 4. 5.

mineralization mainly related to planar architectural elements, e.g. sequence boundaries subaerial and unconformities mineralization mainly related to planar architectural elements, e.g. sequence boundaries submarine, transgressive surfaces and maximum flooding zones/surfaces) mineralization mainly controlled by system tracts (lowstand system tracts transgressive system tracts, highstand system tracts) mineralization of subvolcanic or intermediate level to be correlated with the architectural elements of basin evolution mineralization of deep level to be correlated with the deep-seated structural elements.

There are several squares on the chessboard left blank mainly for lack of information on sequence stratigraphy of mineral deposits. This method has not found many users yet in mineral exploration. This review is designed as an “interactive paper” open, for amendments in the electronic spreadsheet version and adjustable to the needs and wants of application, research and training in geosciences. Metamorphic host rock lithologies and commodities are addressed by different color codes in the chessboard classification scheme. © 2009 Elsevier B.V. All rights reserved.

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Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.1. Mineral deposits and economic geology . . . . . . . . . . . . . . 1.2. Classification of mineral deposits through time . . . . . . . . . . The “chessboard” (spreadsheet ) classification scheme of mineral deposits . 2.1. The principles of the “chessboard” (spreadsheet ) classification scheme 2.2. The host of mineral deposits . . . . . . . . . . . . . . . . . . . 2.2.1. Magmatic host rocks. . . . . . . . . . . . . . . . . . . 2.2.2. Ore-bearing structures . . . . . . . . . . . . . . . . . . 2.2.3. Sedimentary host rocks . . . . . . . . . . . . . . . . . 2.2.4. Organic material and special host rocks . . . . . . . . . . 2.3. Type of commodity (inorganic raw material) . . . . . . . . . . . 2.3.1. Ore minerals . . . . . . . . . . . . . . . . . . . . . . 2.3.2. Industrial minerals and rocks . . . . . . . . . . . . . . 2.3.3. Gemstones and ornamental stones . . . . . . . . . . . . 2.4. Mineralizing processes . . . . . . . . . . . . . . . . . . . . . . Chromium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.1. Chemistry and mineralogy . . . . . . . . . . . . . . . . . . . . 3.2. Magmatic chromium deposits. . . . . . . . . . . . . . . . . . . 3.3. Sedimentary chromium deposits . . . . . . . . . . . . . . . . . 3.4. Metamorphic chromium (gemstone) deposit . . . . . . . . . . . 3.5. Chromium supply and use . . . . . . . . . . . . . . . . . . . . Nickel . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1. Chemistry and mineralogy . . . . . . . . . . . . . . . . . . . . 4.2. Magmatic nickel deposits. . . . . . . . . . . . . . . . . . . . . 4.3. Structure-related nickel deposits . . . . . . . . . . . . . . . . . 4.4. Sedimentary nickel deposits . . . . . . . . . . . . . . . . . . . 4.5. Nickel supply and use . . . . . . . . . . . . . . . . . . . . . . Cobalt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.1. Chemistry and mineralogy . . . . . . . . . . . . . . . . . . . . 5.2. Cobalt deposits . . . . . . . . . . . . . . . . . . . . . . . . . 5.3. Cobalt supply and use . . . . . . . . . . . . . . . . . . . . . . Platinum group elements (PGE) . . . . . . . . . . . . . . . . . . . . . 6.1. Chemistry and mineralogy . . . . . . . . . . . . . . . . . . . . 6.2. Magmatic platinum-group-element deposits . . . . . . . . . . . . 6.3. Structure-related platinum-group-element deposits . . . . . . . . 6.4. Sedimentary platinum-group-element deposits . . . . . . . . . . 6.5. Platinum-group-element supply and use . . . . . . . . . . . . . Titanium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.1. Chemistry and mineralogy . . . . . . . . . . . . . . . . . . . . 7.2. Magmatic titanium deposits . . . . . . . . . . . . . . . . . . . 7.3. Structure-related titanium deposits . . . . . . . . . . . . . . . . 7.4. Sedimentary titanium deposits . . . . . . . . . . . . . . . . . . 7.5. Metamorphic titanium deposits . . . . . . . . . . . . . . . . . . 7.6. Titanium supply and use . . . . . . . . . . . . . . . . . . . . . Vanadium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.1. Chemistry and mineralogy . . . . . . . . . . . . . . . . . . . . 8.2. Magmatic vanadium deposits . . . . . . . . . . . . . . . . . . .

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8.3. Sedimentary vanadium deposits . . . . . . . . . . . . . . . . . . . . . . . . 8.4. Vanadium supply and use . . . . . . . . . . . . . . . . . . . . . . . . . . . Iron . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.1. Chemistry and mineralogy . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.2. Magmatic iron deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.3. Structure-related iron deposits . . . . . . . . . . . . . . . . . . . . . . . . . 9.4. Sedimentary iron deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.5. Metamorphic iron deposits. . . . . . . . . . . . . . . . . . . . . . . . . . . 9.6. Iron supply and use . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Manganese . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.1. Chemistry and mineralogy . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.2. Magmatic manganese deposits . . . . . . . . . . . . . . . . . . . . . . . . . 10.3. Structure-related manganese deposits . . . . . . . . . . . . . . . . . . . . . 10.4. Sedimentary manganese deposits. . . . . . . . . . . . . . . . . . . . . . . . 10.5. Metamorphic manganese deposits . . . . . . . . . . . . . . . . . . . . . . . 10.6. Manganese supply and use. . . . . . . . . . . . . . . . . . . . . . . . . . . Copper . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11.1. Chemistry and mineralogy . . . . . . . . . . . . . . . . . . . . . . . . . . . 11.2. Magmatic copper deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . 11.3. Structure-bound copper deposit . . . . . . . . . . . . . . . . . . . . . . . . 11.4. Sedimentary copper deposits . . . . . . . . . . . . . . . . . . . . . . . . . . 11.5. Copper supply and use. . . . . . . . . . . . . . . . . . . . . . . . . . . . . Selenium and tellurium. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12.1. Chemistry, mineralogy and economic geology of selenium . . . . . . . . . . . . 12.2. Chemistry, mineralogy and economic geology of tellurium . . . . . . . . . . . . Molybdenum and rhenium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13.1. Chemistry and mineralogy of molybdenum . . . . . . . . . . . . . . . . . . . 13.2. Magmatic molybdenum deposits . . . . . . . . . . . . . . . . . . . . . . . . 13.3. Sedimentary molybdenum and rhenium deposits . . . . . . . . . . . . . . . . 13.4. Molybdenum and rhenium supply and use . . . . . . . . . . . . . . . . . . . Tin and tungsten . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14.1. Chemistry and mineralogy . . . . . . . . . . . . . . . . . . . . . . . . . . . 14.2. Magmatic tin and tungsten deposits . . . . . . . . . . . . . . . . . . . . . . 14.3. Structure-bound tungsten deposits . . . . . . . . . . . . . . . . . . . . . . . 14.4. Sedimentary tin and tungsten deposits . . . . . . . . . . . . . . . . . . . . . 14.5. Supply and use of tin and tungsten . . . . . . . . . . . . . . . . . . . . . . . Niobium-, tantalum- and scandium . . . . . . . . . . . . . . . . . . . . . . . . . . 15.1. Chemistry and mineralogy . . . . . . . . . . . . . . . . . . . . . . . . . . . 15.2. Magmatic niobium-, tantalum- and scandium deposits. . . . . . . . . . . . . . 15.3. Sedimentary niobium- and tantalum deposits . . . . . . . . . . . . . . . . . . 15.4. Supply and use of niobium and tantalum . . . . . . . . . . . . . . . . . . . . Beryllium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16.1. Beryllium chemistry and mineralogy . . . . . . . . . . . . . . . . . . . . . . 16.2. Magmatic beryllium deposits. . . . . . . . . . . . . . . . . . . . . . . . . . 16.3. Structure-bound beryllium deposits . . . . . . . . . . . . . . . . . . . . . . 16.4. Sedimentary beryllium deposits . . . . . . . . . . . . . . . . . . . . . . . . 16.5. Metamorphic beryllium deposits . . . . . . . . . . . . . . . . . . . . . . . . 16.6. Supply and use of beryllium . . . . . . . . . . . . . . . . . . . . . . . . . . Cesium, lithium and rubidium . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17.1. Chemistry and mineralogy of cesium and lithium . . . . . . . . . . . . . . . . 17.2. Magmatic cesium-, lithium and rubidium deposits. . . . . . . . . . . . . . . . 17.3. Cesium and lithium sedimentary deposits. . . . . . . . . . . . . . . . . . . . 17.4. Cesium and lithium supply. . . . . . . . . . . . . . . . . . . . . . . . . . . Lead, zinc, germanium, indium, cadmium and silver . . . . . . . . . . . . . . . . . . 18.1. Chemistry and mineralogy of lead, zinc, germanium, indium, cadmium, and silver. 18.2. Magmatic lead-zinc deposits . . . . . . . . . . . . . . . . . . . . . . . . . . 18.3. Structure-bound lead-zinc deposits . . . . . . . . . . . . . . . . . . . . . . . 18.4. Sedimentary lead–zinc deposits . . . . . . . . . . . . . . . . . . . . . . . . 18.5. Metamorphic lead–zinc deposits . . . . . . . . . . . . . . . . . . . . . . . . 18.6. Silver deposits sensu stricto . . . . . . . . . . . . . . . . . . . . . . . . . . 18.7. Lead, zinc, germanium, indium, cadmium, and silver supply . . . . . . . . . . . 18.7.1. Past and present of mining and metallurgy of lead . . . . . . . . . . . 18.7.2. Past and present of mining and metallurgy of zinc . . . . . . . . . . . 18.7.3. Final use of germanium, indium and cadmium . . . . . . . . . . . . . 18.7.4. Final use of silver . . . . . . . . . . . . . . . . . . . . . . . . . . . Bismuth . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.1. Chemistry and mineralogy of bismuth . . . . . . . . . . . . . . . . . . . . . 19.2. Deposits of bismuth . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.3. Final use of bismuth . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Gold. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20.1. Chemistry and mineralogy of gold . . . . . . . . . . . . . . . . . . . . . . . 20.2. Magmatic gold deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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66 68 68 68 69 74 75 85 88 89 89 89 93 93 94 97 97 97 98 107 107 118 118 118 118 119 119 119 120 120 121 121 122 128 128 129 129 129 130 134 136 136 136 137 140 141 142 143 143 143 143 146 148 148 148 150 153 154 166 168 169 169 172 172 174 174 174 174 175 175 175 180

4


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27.

28.

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30.

20.3. Structure-bound gold deposits . . . . . . . . . . . . . . . 20.4. Sedimentary gold deposits . . . . . . . . . . . . . . . . . 20.5. Metamorphic gold deposits . . . . . . . . . . . . . . . . . 20.6. Gold supply and use . . . . . . . . . . . . . . . . . . . . Antimony, arsenic and thallium . . . . . . . . . . . . . . . . . . 21.1. Chemistry and mineralogy of antimony, arsenic and thallium . 21.2. Antimony deposits . . . . . . . . . . . . . . . . . . . . . 21.2.1. Magmatic antimony deposits. . . . . . . . . . . . 21.2.2. Structure-bound antimony deposits . . . . . . . . 21.2.3. Sedimentary antimony deposits . . . . . . . . . . 21.3. Arsenic deposits . . . . . . . . . . . . . . . . . . . . . . 21.4. Thallium deposits . . . . . . . . . . . . . . . . . . . . . 21.5. Antimony, arsenic, thallium supply and use . . . . . . . . . 21.5.1. Antimony . . . . . . . . . . . . . . . . . . . . . 21.5.2. Arsenic . . . . . . . . . . . . . . . . . . . . . . 21.5.3. Thallium . . . . . . . . . . . . . . . . . . . . . Mercury . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22.1. Chemistry and mineralogy of mercury. . . . . . . . . . . . 22.2. Magmatic mercury deposits . . . . . . . . . . . . . . . . 22.3. Structure-bound mercury deposits . . . . . . . . . . . . . 22.4. Sedimentary mercury deposits . . . . . . . . . . . . . . . 22.5. Mercury supply and use . . . . . . . . . . . . . . . . . . Rare Earth Elements (REE) . . . . . . . . . . . . . . . . . . . . 23.1. Chemistry and mineralogy of rare earth elements . . . . . . 23.2. Magmatic rare earth element deposits . . . . . . . . . . . 23.3. Structure-bound rare earth element deposits . . . . . . . . 23.4. Sedimentary rare earth element deposits . . . . . . . . . . 23.5. Rare earth element supply and use . . . . . . . . . . . . . Uranium, thorium and radium. . . . . . . . . . . . . . . . . . . 24.1. Chemistry and mineralogy of uranium and thorium . . . . . 24.2. Uranium deposits . . . . . . . . . . . . . . . . . . . . . 24.2.1. Magmatic uranium deposits . . . . . . . . . . . . 24.2.2. Structure-bound uranium deposits . . . . . . . . . 24.2.3. Sedimentary uranium deposits . . . . . . . . . . . 24.3. Thorium deposits . . . . . . . . . . . . . . . . . . . . . 24.3.1. Magmatic thorium deposits . . . . . . . . . . . . 24.3.2. Structure-bound thorium deposits . . . . . . . . . 24.3.3. Sedimentary thorium deposits . . . . . . . . . . . 24.3.4. Uranium and thorium supply and use . . . . . . . Aluminum and gallium . . . . . . . . . . . . . . . . . . . . . . 25.1. Chemistry and mineralogy of aluminum and gallium . . . . . 25.2. Magmatic aluminum deposits. . . . . . . . . . . . . . . . 25.3. Sedimentary aluminum deposits . . . . . . . . . . . . . . 25.4. Gallium deposits . . . . . . . . . . . . . . . . . . . . . . 25.5. Aluminum and gallium supply and use . . . . . . . . . . . Magnesium . . . . . . . . . . . . . . . . . . . . . . . . . . . 26.1. Chemistry and mineralogy of magnesium . . . . . . . . . . 26.2. Magmatic magnesium deposits . . . . . . . . . . . . . . . 26.3. Structure-bound magnesium deposits . . . . . . . . . . . . 26.4. Sedimentary magnesium deposits . . . . . . . . . . . . . . 26.5. Magnesium supply and use . . . . . . . . . . . . . . . . . Calcium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27.1. Chemistry and mineralogy of calcium . . . . . . . . . . . . 27.2. Magmatic calcium deposits . . . . . . . . . . . . . . . . . 27.3. Structure-bound calcium deposits . . . . . . . . . . . . . . 27.4. Sedimentary calcium deposits . . . . . . . . . . . . . . . 27.5. Metamorphic calcium deposits . . . . . . . . . . . . . . . 27.6. Carbonate rocks supply and use. . . . . . . . . . . . . . . Boron . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 28.1. Chemistry and mineralogy of boron . . . . . . . . . . . . . 28.2. Magmatic boron deposits . . . . . . . . . . . . . . . . . . 28.3. Sedimentary boron deposits . . . . . . . . . . . . . . . . 28.4. Metamorphic boron deposits . . . . . . . . . . . . . . . . 28.5. Boron supply and use . . . . . . . . . . . . . . . . . . . Sulfur and calcium sulfate . . . . . . . . . . . . . . . . . . . . . 29.1. Chemistry and mineralogy of sulfur . . . . . . . . . . . . . 29.2. Magmatic sulfur and sulfate deposits . . . . . . . . . . . . 29.3. Sedimentary sulfur and Ca sulfate deposits . . . . . . . . . 29.4. Sulfur and calcium sulfate supply and use . . . . . . . . . . Fluorine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 30.1. Chemistry and mineralogy of fluorine . . . . . . . . . . . . 30.2. Magmatic fluorine deposits . . . . . . . . . . . . . . . . .

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186 187 188 188 188 188 189 189 195 196 196 198 199 199 200 200 200 200 200 203 203 203 203 203 203 206 206 206 206 206 208 208 210 214 220 220 222 222 222 222 222 222 224 226 227 229 229 230 230 232 235 235 235 236 236 236 241 241 242 242 242 243 246 246 246 246 247 248 252 252 252 253


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30.3. Structure-bound fluorite and topaz deposits. . . . . . . . . . . . . . 30.4. Sedimentary fluorite and topaz deposits . . . . . . . . . . . . . . . 30.5. Fluorine supply and use . . . . . . . . . . . . . . . . . . . . . . . Barium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 31.1. Chemistry and mineralogy of barium . . . . . . . . . . . . . . . . . 31.2. Magmatic barium deposits . . . . . . . . . . . . . . . . . . . . . . 31.3. Structure-bound barium deposits . . . . . . . . . . . . . . . . . . . 31.4. Sedimentary barium deposits . . . . . . . . . . . . . . . . . . . . 31.5. Barium supply and use . . . . . . . . . . . . . . . . . . . . . . . Strontium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32.1. Chemistry and mineralogy of strontium. . . . . . . . . . . . . . . . 32.2. Magmatic strontium deposits . . . . . . . . . . . . . . . . . . . . 32.3. Structure-bound strontium deposits . . . . . . . . . . . . . . . . . 32.4. Sedimentary strontium deposits . . . . . . . . . . . . . . . . . . . 32.5. Strontium supply and use . . . . . . . . . . . . . . . . . . . . . . Potassium, sodium, chlorine and bromine . . . . . . . . . . . . . . . . . . 33.1. Chemistry and mineralogy of potassium, sodium, chlorine and bromine . . . 33.2. Sedimentary potassium, sodium, chlorine and bromine deposits . . . . 33.3. Sodium, potassium, chlorine and bromine supply and use . . . . . . . Nitrogen and iodine . . . . . . . . . . . . . . . . . . . . . . . . . . . . 34.1. Chemistry and mineralogy of nitrogen and iodine . . . . . . . . . . . 34.2. Sedimentary nitrogen and iodine deposits . . . . . . . . . . . . . . 34.3. Economic consideration of nitrogen and iodine deposits . . . . . . . . Sodium carbonate and –sulfate . . . . . . . . . . . . . . . . . . . . . . . 35.1. Chemistry and mineralogy of sodium carbonate and –sulfate. . . . . . 35.2. Sedimentary sodium carbonate and –sulfate deposits . . . . . . . . . 35.3. Sodium carbonate and –sulfate supply and use . . . . . . . . . . . . Phosphorus . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36.1. Chemistry and mineralogy of phosphorus . . . . . . . . . . . . . . . 36.2. Magmatic phosphate deposits . . . . . . . . . . . . . . . . . . . . 36.3. Structure-bound phosphate deposits . . . . . . . . . . . . . . . . . 36.4. Sedimentary phosphate deposits . . . . . . . . . . . . . . . . . . . 36.5. Metamorphic phosphate deposits. . . . . . . . . . . . . . . . . . . 36.6. Phosphate supply and use . . . . . . . . . . . . . . . . . . . . . . Zirconium and hafnium. . . . . . . . . . . . . . . . . . . . . . . . . . . 37.1. Chemistry and mineralogy of zirconium and hafnium . . . . . . . . . 37.2. Magmatic, metamorphic and sedimentary zirconium (hafnium) deposits 37.3. Zirconium (hafnium) supply and use . . . . . . . . . . . . . . . . . Silica . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38.1. Chemistry and mineralogy of silica . . . . . . . . . . . . . . . . . . 38.2. Magmatic silica deposits . . . . . . . . . . . . . . . . . . . . . . . 38.3. Structure-bound silica deposits . . . . . . . . . . . . . . . . . . . . 38.4. Sedimentary silica deposits . . . . . . . . . . . . . . . . . . . . . 38.5. Metamorphic silica deposits . . . . . . . . . . . . . . . . . . . . . 38.6. Silica supply and use . . . . . . . . . . . . . . . . . . . . . . . . Feldspar . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39.1. Chemistry and mineralogy of feldspar . . . . . . . . . . . . . . . . 39.2. Magmatic feldspar deposits . . . . . . . . . . . . . . . . . . . . . 39.3. Structure-bound feldspar deposits . . . . . . . . . . . . . . . . . . 39.4. Sedimentary feldspar deposits . . . . . . . . . . . . . . . . . . . . 39.5. Metamorphic feldspar deposits . . . . . . . . . . . . . . . . . . . . 39.6. Feldspar supply and use . . . . . . . . . . . . . . . . . . . . . . . Feldspathoid . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40.1. Chemistry and mineralogy of feldspathoids . . . . . . . . . . . . . . 40.2. Magmatic feldspathoids deposits . . . . . . . . . . . . . . . . . . . 40.3. Metamorphic feldspathoids deposits . . . . . . . . . . . . . . . . . 40.4. Feldspathoids supply and use . . . . . . . . . . . . . . . . . . . . Zeolites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41.1. Chemistry and mineralogy of zeolites . . . . . . . . . . . . . . . . . 41.2. Magmatic zeolite deposits . . . . . . . . . . . . . . . . . . . . . . 41.3. Sedimentary zeolite deposits . . . . . . . . . . . . . . . . . . . . . 41.4. Metamorphic zeolite deposits . . . . . . . . . . . . . . . . . . . . 41.5. Zeolite supply and use . . . . . . . . . . . . . . . . . . . . . . . . Amphibole and asbestos . . . . . . . . . . . . . . . . . . . . . . . . . . 42.1. Chemistry and mineralogy of amphibole and asbestiform minerals . . . 42.2. Magmatic amphibole and asbestos deposits . . . . . . . . . . . . . . 42.3. Metamorphic amphibole and asbestos deposits . . . . . . . . . . . . 42.4. Amphibole and asbestos supply and use . . . . . . . . . . . . . . . Olivine and dunite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43.1. Chemistry and mineralogy of olivine minerals. . . . . . . . . . . . . 43.2. Magmatic olivine and dunite deposits . . . . . . . . . . . . . . . . 43.3. Sedimentary olivine deposits . . . . . . . . . . . . . . . . . . . . .

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5

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43.4. Metamorphic olivine and dunite deposits . . . . . . . . . . . . . 43.5. Olivine and dunite supply and use . . . . . . . . . . . . . . . . Pyroxene and inosilicates . . . . . . . . . . . . . . . . . . . . . . . . 44.1. Chemistry and mineralogy of pyroxene and pyroxenoid . . . . . . 44.2. Magmatic pyroxene and inosilicate deposits . . . . . . . . . . . . 44.3. Metamorphic pyroxene and inosilicate deposits . . . . . . . . . . 44.4. Pyroxene and inosilicate supply and use. . . . . . . . . . . . . . Garnet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 45.1. Chemistry and mineralogy of garnet-group minerals . . . . . . . . 45.2. Magmatic garnet deposits . . . . . . . . . . . . . . . . . . . . 45.3. Structure-bound garnet deposits . . . . . . . . . . . . . . . . . 45.4. Sedimentary garnet deposits . . . . . . . . . . . . . . . . . . . 45.5. Metamorphic garnet deposits . . . . . . . . . . . . . . . . . . . 45.6. Garnet supply and use . . . . . . . . . . . . . . . . . . . . . . Epidote-group minerals . . . . . . . . . . . . . . . . . . . . . . . . . 46.1. Chemistry and mineralogy of epidote-group minerals . . . . . . . 46.2. Magmatic epidote-group mineral deposits . . . . . . . . . . . . . 46.3. Metamorphic epidote-group deposits . . . . . . . . . . . . . . . 46.4. Epidote-group minerals supply and use . . . . . . . . . . . . . . Sillimanite-group minerals . . . . . . . . . . . . . . . . . . . . . . . 47.1. Chemistry and mineralogy of sillimanite-group minerals . . . . . . 47.2. Magmatic deposits of sillimanite-group minerals. . . . . . . . . . 47.3. Structure-bound deposits of sillimanite-group minerals and staurolite . 47.4. Sedimentary deposits of sillimanite-group minerals and staurolite . 47.5. Metamorphic sillimanite-group deposits. . . . . . . . . . . . . . 47.6. Sillimanite-group minerals and staurolite supply and use. . . . . . Corundum and spinel . . . . . . . . . . . . . . . . . . . . . . . . . . 48.1. Chemistry and mineralogy of corundum and spinel minerals . . . . 48.2. Magmatic deposits of corundum and spinel minerals. . . . . . . . 48.3. Sedimentary deposits of corundum and spinel minerals . . . . . . 48.4. Metamorphic deposits of corundum and spinel minerals . . . . . . 48.5. Corundum and spinel supply and use . . . . . . . . . . . . . . . Diamond . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 49.1. Chemistry and mineralogy of diamond . . . . . . . . . . . . . . 49.2. Magmatic deposits of diamonds. . . . . . . . . . . . . . . . . . 49.3. Sedimentary deposits of diamonds . . . . . . . . . . . . . . . . 49.4. Metamorphic deposits of diamonds . . . . . . . . . . . . . . . . 49.5. Diamond supply and use . . . . . . . . . . . . . . . . . . . . . Graphite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50.1. Chemistry and mineralogy of graphite. . . . . . . . . . . . . . . 50.2. Magmatic deposits of graphite . . . . . . . . . . . . . . . . . . 50.3. Structure-bound graphite deposits . . . . . . . . . . . . . . . . 50.4. Sedimentary graphite deposits . . . . . . . . . . . . . . . . . . 50.5. Metamorphic graphite deposits . . . . . . . . . . . . . . . . . . 50.6. Graphite supply and use . . . . . . . . . . . . . . . . . . . . . Clay minerals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51.1. Chemistry and mineralogy of clay minerals . . . . . . . . . . . . 51.2. Magmatic deposits of clay minerals . . . . . . . . . . . . . . . . 51.2.1. Serpentine–kaolin . . . . . . . . . . . . . . . . . . . . 51.2.2. Talc-pyrophyllite . . . . . . . . . . . . . . . . . . . . 51.2.3. Smectite . . . . . . . . . . . . . . . . . . . . . . . . 51.2.4. Vermiculite . . . . . . . . . . . . . . . . . . . . . . . 51.2.5. Mica . . . . . . . . . . . . . . . . . . . . . . . . . . 51.2.6. Chlorite . . . . . . . . . . . . . . . . . . . . . . . . . 51.2.7. Sepiolite–palygorskite (hormites) . . . . . . . . . . . . 51.3. Sedimentary deposits of clay minerals. . . . . . . . . . . . . . . 51.3.1. Serpentine–kaolin . . . . . . . . . . . . . . . . . . . . 51.3.2. Talc-pyrophyllite . . . . . . . . . . . . . . . . . . . . 51.3.3. Smectite . . . . . . . . . . . . . . . . . . . . . . . . 51.3.4. Vermiculite . . . . . . . . . . . . . . . . . . . . . . . 51.3.5. Mica . . . . . . . . . . . . . . . . . . . . . . . . . . 51.3.6. Sepiolite–palygorskite (hormites) . . . . . . . . . . . . 51.4. Metamorphic deposits of clay minerals . . . . . . . . . . . . . . 51.4.1. Mica–(chlorite) . . . . . . . . . . . . . . . . . . . . . 51.4.2. Prehnite . . . . . . . . . . . . . . . . . . . . . . . . 51.4.3. Talc. . . . . . . . . . . . . . . . . . . . . . . . . . . 51.4.4. Pyrophyllite . . . . . . . . . . . . . . . . . . . . . . . 51.5. Clay minerals and phyllosilicate supply and use . . . . . . . . . . Biological materials (“biominerals”) . . . . . . . . . . . . . . . . . . . 52.1. Chemistry and mineralogy of biological materials . . . . . . . . . 52.2. Jet and amber deposits . . . . . . . . . . . . . . . . . . . . . . 52.3. Supply and use of biological materials. . . . . . . . . . . . . . .

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53.

Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53.1. The columns: geodynamic setting and environment of deposition . . . . . . . . . . . . . . . . . 53.1.1. The ultrabasic magmatic clan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53.1.2. The basic magmatic clan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53.1.3. The intermediate and felsic magmatic clan . . . . . . . . . . . . . . . . . . . . . . . . 53.1.4. The alkaline magmatic clan and carbonatites with pegmatitic plus aplitic derivative products 53.1.5. One-dimensional structures— pipes . . . . . . . . . . . . . . . . . . . . . . . . . . . 53.1.6. Two-dimensional structures—veins. . . . . . . . . . . . . . . . . . . . . . . . . . . . 53.1.7. Duricrusts–regolith–vein like deposits . . . . . . . . . . . . . . . . . . . . . . . . . . 53.1.8. Coarse-grained clastic rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53.1.9. Fine-grained clastic rocks and massive rocks . . . . . . . . . . . . . . . . . . . . . . . 53.1.10. Limestones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53.1.11. Evaporites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53.1.12. Special sedimentary rocks and carbon-bearing hosts . . . . . . . . . . . . . . . . . . . 53.2. Metamorphogenetic mineral deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53.3. The spreadsheet correlation — bringing together the columns and the lines . . . . . . . . . . . . . 53.4. Sequence stratigraphy of epigenetic and diagenetic mineral deposits in Central Europe . . . . . . . 54. Conclusions — the importance of raw materials . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

1. Introduction 1.1. Mineral deposits and economic geology Mineral deposits resulted from physical–chemical changes in the atmosphere, hydrosphere and lithosphere (crust and the upper mantle). There are physical barriers hampering our access to mineral deposits, such as mining depth controlled among others by the geothermal gradients (5 km will hardly be crossed by exploitation methods in the foreseeable future) and the vast oceans, which cover more than 2/3 of the Earth surface. Manganese nodules are widely known as a potential source of Mn, Cu and Ni, but neither the technical nor the jurisdictional issues have so far been solved to everybody's satisfaction, so that a vast part of the sea has still to be considered as an area of scientific research of black and white smokers, rather than an area of exploration and exploitation of mineral deposits. On the other hand, who can exclude that future generations will extract all elements from seawater and consider the ocean as the only inexhaustible low-grade, large-tonnage deposit on earth? The first step has already been taken with some alkaline, earth alkaline and halogenides recovered from seawater. For the time being, a growing demand for mineral raw materials and, unlike organic raw materials, with only a few of the inorganic raw materials being renewable in a life-time, mineral resources are limited. We are in need of the knowledge and experience provided by economic geologists from academia and application (Gocht, 1978; Saager, 1984). Economic geology is not a discipline of its own, it is a mixtum compositum of various subjects of earth sciences dedicated to find new inorganic raw materials and enhance the exploitation of those already known. Geology and mineralogy are the key players in achieving these goals. Magmatic, sedimentary and metamorphic processes may concentrate some elements to such a degree that these mineral occurrences stand out not only by their mineralogical and chemical compositions but also by color, texture or structure from the unmineralized or poorly-mineralized wall and country rocks around. Hand specimens of such “abnormal” rocks, called ore, are shown in the succeeding chapters together with the sites which these rocks have been taken from, called ore bodies. The widely-known report “Limits of Growth” by the “Club of Rome” predicted in 1972 that by the year 2000 many deposits will be exhausted and many metals will no longer be available even at a high price level. The pessimistic and very constricted view of this group of persons has proved to be wrong, but not so the predictions of Hewett in 1927 who was able to demonstrate that in mining countries exploitation, export and import of metals evolve in a cyclic way (Saager, 1984). During an initial stage, countries go through a period of metal surplus. With the domestic industry on the rise, the demand

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374 374 374 374 374 375 375 375 375 379 379 379 379 381 381 381 382 382 382 384

for metals grows while the domestic metal production has already passed its climax. At the very end the domestic mining industry can no longer cater for the needs and wants of the domestic industry and metallic goods have to be imported from foreign countries. The current situation in the Asian countries China and India is another example for the rightness of this prediction made in 1927. The world production of mineral raw materials may be looked up in various publications such as the “World Mining Data”, a yearbook issued by Weber and Zsak (2007) or on the Internet (US Geological Survey http://minerals.usgs.gov/minerals/). This link is placed at the end of each section to give the reader a quick access to these data and update his/her knowledge on an annual basis. World mining data listed at the end of sections are to give a rough idea on the world raw material situation and enable the reader to critically look at the daily statistics. Some data in the paper have been derived from compilations by Huy (2007) and based on the annual reports of the Raw Materials Group (Sweden). A bridge between classical mining data and metallogenic studies has been built by the review of giant ore deposits by Laznicka (1999). 1.2. Classification of mineral deposits through time The wealth of minerals and rocks which form the building blocks of mineral deposits has very early sparked attempts among geoscientists to classify mineral deposits and refine their terminology. Lindgren (1934) and Lindberg (1922) were among the pioneers who addressed this problem in a modern way; numerous others followed suit and joined in the attempt to put in order the newly discovered mineral deposits worldwide as well as those which have already been mined for ages. Mining operations may be counted among the most long-lasting operations. Different approaches have been taken by the various students of mineral deposits to pigeonhole or classify the wide range of mineral deposits. Some provided a general overview of mineral deposits and ore-forming processes (Bateman, 1950, 1957; Schneiderhöhn, 1962; Routhier, 1963; Stanton, 1972; Hutchinson, 1983; Laznicka, 1985; Schröcke, 1986; Guilbert and Park, 1986; Carr and Herz, 1989; Pohl, 1992, 2005; Evans, 1993; Kesler, 1994; Robb, 2004; Laznicka, 2005) while others selected one group of commodity (Manning, 1995; Harben and Kužvart, 1996) or a special type of ore deposits such as those related to magmatic processes (Whitney and Naldrett, 1989). A different approach has been taken by Roberts and Sheahan (1988) and Kirkham et al. (1993) who focused on modeling ore deposits. The textbooks by Barnes (1997) on the geochemistry of hydrothermal ore deposits and Henley et al. (1984) on fluid-mineral equilibria in hydrothermal systems deviate from the norm when focusing on the classification of mineral

8


deposits. They are quoted here for two reasons. In these books, there are some groups of (hydrothermal) deposits dealt with so that the data may be used for classification schemes. First and foremost, these papers act as transition into geochemistry, a field of study which is of utmost importance when studying mineral deposits. However, this topic cannot be dealt with in any of the afore-mentioned classification schemes at length without getting the reader's attention drawn off from the gist of these classifications or the audience drowned into numbers and equations that are published elsewhere in a more appropriate way. Among the energy resources it has become common practice to treat the various commodities separately, as it was done for uranium deposits by Dahlkamp (1979), for coal deposits by Stach (1982), Diessel (1992), Moore and Shearer (2003) and Kalkreuth (2004), and for petroleum deposits by Tissot and Welte (1978), Hunt (1979), North (1985), Selley (1997) and Gluyas and Swarbrick (2004). The chessboard classification scheme presented in this paper is open for an enlargement towards coal petrography and classification schemes dealing with the accumulation of organic matter in whatever physical condition (Fig. 01.01a,b). There are special publications which try to relate metallogenesis to a particular geodynamic process such as global tectonics (Sawkins, 1990) or deal with a particular group of ore deposits such as the iron or non-sulfidic zinc deposits (Zitzmann, 1977; Boni and Large, 2003). German economic geologists have begun compiling the wealth of information gathered through times on Pb–Zn- and Fe deposits in several monographs in the “Geologische Jahrbuch” and “Beiheft zum Geologischen Jahrbuch” between 1951 and 1986 and thereby contributed to the classification of vein-type deposits, in particular. Only a few of these monographs can be quoted here. Regional-based classification schemes are numerous and in this context some studies from Central Europe may be quoted as representative of this sort of classification schemes (Pouba and Ilavsky, 1986; Dill, 1989; Osika, 1990; Walther and Dill, 1995; Dill et al., 2008a,b). Some classification schemes can still be used for the purpose which they were designed for, be it teaching or exploration, while others are still nice to read and informative on certain aspects. It has to be noted, no classification scheme is really perfect or complete. In the most recent comprehensive publication on fossil fuels, ore and industrial minerals in Central Europe two separate approaches have been taken one for the map (Dill et al., 2008a) and another for the text (Dill et al., 2008b). The map used an element-based classification scheme, avoiding any interpretation, whereas in the text a classification scheme has been adopted which makes use of various structural elements related in time and space to the Variscan and Alpine orogenies, the most significant geodynamic phases during crustal consolidation of what is called today Central Europe: (1) stratabound deposits, (2) thrust-bound metamorphogenic and/or fold-related deposits, (3) deposits controlled by collision-related granitic activity, (4) unconformity-related fault-bound hypogene and supergene deposits, (5) deposits controlled by extension-related magmatic activity along deep-seated fault zones, and (6) petroleum deposits. This classification scheme is well-established for the extraAlpine part of Central Europe (e.g. Dill, 1988a, 1989, 1994a; Dill and Nielsen, 1987; Tischendorf et al., 1995) and was extended to the Alpine realm, where similar subdivisions have been applied by Pohl (1993), Pohl and Belocky (1994, 1999) and Rodeghiero et al. (1996). 2. The “chessboard” (spreadsheet ) classification scheme of mineral deposits 2.1. The principles of the “chessboard” (spreadsheet ) classification scheme Needless to say, ideas, concepts and models in economic geology are important but they are changing rapidly. Quite often, they are dependent on the metallogenic fashion and fostered by current trends, likewise driven by industry or politics. This back and fro in economic geology can neither be the basic philosophy of exploration nor is it helpful to disclose the secrets of this mixtum composition

called economic geology to the beginner in its own camp or to people interested in it but not dealing with this matter on a permanent basis. When studying mineral deposits, nobody can afford to sideline any new chemical or physical methods, appearing on the scene but these new methods need to be handled in a careful balance combined with “noseon-rock methods” (hammer and laptop). Results obtained by means of fluid inclusion measurements, data derived from the study of stable and radioactive isotopes are sparsely used in this study. The current study has two major goals, not to cloud the reader's vision and not to distract the reader's attention from the essentials of economic geology: geology, mineralogy and chemical composition. The reader is referred to the different publications quoted for each type of deposit to get more information on those data which cannot be treated to the full extent in this paper. Ore grades reported for the various commodities in the final section of each chapter are not strict or sharp boundaries set in this case by experts from the “Gesellschaft für Metall-, Hütten- und Bergleute” in Germany (GDMB). These data have been reported in the final section of each chapter to give the reader an idea to what extent an element has to be enriched relative to the average grade in the crust, which is reported in the introduction to each chapter, to render this element feasible for exploitation. Mean values of elements may be looked up in various reference books like the “Handbook on Geochemistry” or found when visiting website likes http://www. uniterra.de/rutherford (Clark and Washington, 1924; Mason, 1958; Turekian and Wedepohl, 1961; Vinogradov, 1962; Taylor, 1964). Mineral formulae are found among others in the paper by Kretz (1983) and all minerals referred to in this study are listed in Table 01.01. Further information on minerals can be obtained by consulting encyclopedias such as Roberts et al. (1990) or textbooks such as Strunz (1970) and Ramdohr and Strunz (1978). Those who want to know what nice well-crystallized minerals in mineral galleries look like are requested to visit special websites on the internet, e.g., “Mineralogy Database” http://webmineral.com or http://www. mindat.org/. Those who make use of ore microscopy, which is still a valuable and efficient tool to understand ore genesis and not yet outdated by the electron microprobe (EMP), need to consult textbooks or determination tables published by Schouten (1962), Ramdohr (1975), Picot and Johan (1977), Craig and Vaughan (1981), Uytenbogaardt and Burke (1985) and Mücke (1989) or the “Virtual Atlas of Opaques and Ore Minerals” http://www.smenet.org/opaque-ore/. The reader will find there ore varieties on a microscopic scale for each type of ore deposit containing native elements, oxidic or sulfidic minerals. Classification schemes solely based on new concepts, models and geodynamic processes may sometimes stand on shaky grounds and may get swiftly out of fashion. In the wide range of features that may be useful to define mineral deposits, two categories underwent little change through time. It is the terminology of the host rock lithology and the ore-bearing structures on one part and the element and mineral of economic interest on the other. There may be ups and downs in demand and supply, but rarely has an element really fallen out of use, excluding uranium in modern-day Germany. Exploration does not place the main emphasis on ore types but on raw materials and it goes without saying that it is the quality and quantity of elements and the minerals that make the wheel in the head frame of a shaft go round or make load–haul-dumpers go ahead in the open pit. A chemist with industrial background who wants to know more about the basic ingredients of a certain substance in his lab uses the element for his/her search and the rockhound makes use of the mineral to find a new site where to look for something out of the ordinary. And an exploration geologist specialized on a certain type of deposit but less familiar with other commodities will enter a new frontier area through the metal's or mineral's doors rather than an ore type or a model. The first object an exploration geologist becomes aware of in the field is the lithology of the country rocks that might then turn into the wall rocks of the ore body.


9

Table 01.01 Chemical composition of minerals referred to in the text. The column on the right-hand side denotes the “line” where the mineral may be found (e.g. 12 = Sn–W). Mineral

Chemical composition

Line

Actinolite Aegerine Aeschynite Akaganeite Alabandite Albite (Na end member plagioclase) Alexandrite Allanite-(Ce) orthite Allargentum Alstonite Altaite Alunite Alunogen Amblygonite Amesite Amosite Amphibole

Ca2Fe2Mg3(Si8O22)(OH)2 NaFeSi2O6 (Ce,Ca,Fe)(Ti,Nb)2(O,OH)6 Fe7.6Ni0.4O6.4(OH)9.7Cl1.3 MnS NaAlSi3O8 BeAl2O4 (Ce,Ca,Y)2(Al,Fe)3(SiO4)3(OH) Ag0.99Sb0.01 BaCa(CO3)2 PbTe KAl3(SO4)2(OH)6 Al2(SO4)3·17(H2O) Li0.75Na0.25Al(PO4)F0.75(OH)0.25 Mg2Al2SiO5(OH)4 Fe7(Si8O22)(OH)2 XY2Z5(Si, Al, Ti)8O22(OH, F)2 X = K, Na, Y = Fe Ca, Mg, Mn, Zn, Z = Fe, Mn, Al, Ti NaAl(Si2O6)·(H2O) TiO2 SrCe(CO3)2(OH)·(H2O) Al2SiO5 PbSO4 CaSO4 Cu1.75S Ca(Mg, Fe)(CO3)2 Ni3 (AsO4)2.8H2O CaAl2Si2O8 (Na,K)AlSi3O8 Mg7(Si8O22)(OH)2 Sb2Se3 Sb Ca5 (F,Cl,OH) (PO4)3 KCa4(Si4O10)2F·8(H2O) CaCO3 Mn2.8Ca0.8Mg1.4Al4.7Fe0.4(AsO4)0.9(VO4)0.1(SiO4)2(Si3O10)(OH)6 Na3Fe2+4Fe3+(Si8O22)(OH)2 Ag2S Ag8GeS6 BaCa2Al6Si8O22·2H2O As Ca2Fe3(AsO4)3O3·3(H2O) As2O3 FeAsS Ni0.3Co0.1Ca 0.1Mn2 +1.5O1.5(OH)2·0.6(H2O) Cu2Cl(OH)3 Pd2.25Pt0.75Sn (Mg,Al)2Si4O10(OH)·4H2O (Ca,Na)(Mg,Fe,Al,Ti)(Si,Al)2O6 AuSb2 Ca(UO2)2(PO4)2·12(H2O) Ca2MgAl2(BO3)Si4O12(OH) Cu3(CO3)2(OH)2 ZrO2 (Ba,Sr)(Nb,Ti)2(O,OH)7 BaSO4 2Ca(SO4)·H2O (Ce,La)(CO3)F Be3(Sc,Al)2 Si2O18 Ca(UO2)6O4(OH)6·8(H2O) Be(OH)2 Na0.5Al2(Si3.5Al0.5)O10(OH)2·n(H2O) Cu7Hg6 BaTiSi3O9 Ag8Sb0.5As0.5Te2S3 Ca,Mg,Mn7(Ba,Sr)6(CO3)13 Be2(BO3)(OH,F)·(H2O) Fe2Al(Si,Al)O5(OH)4 Be4Si2O7(OH)2 Be3Al2Si6O8

14-19-41-44-49 44-46 24 7 8-37 7-12-13-14-19-25-27-30-41-42-43-46 14 24-48 16 32-33 10 16-19-20-38-40-49-54-55 31 13-15 32 44 44-46-48 43 5 24 49-50 16 31 9 7 2 41-46-50 41 44 20 20 5-6-7-9-12-13-24-25-32-38-41-50-51-59 43 29 47 25-26 17 16 32 21 21 21 21 2 9 4 38-57-61 5-46-50 19-20 25 30 9 39 24 33 29 24 13 25 14 57 23 5 10 32-33 14-30 7 14 14

Cu2Se Ca2Al2SiO6(OH)2 Li2Al2(Si2O6)2·2(H2O)

10 42 15

Analcime Anatase Ancylite Andalusite Anglesite Anhydrite Anilite Ankerite Annabergite Anorthite (Ca end member plagioclase) Anothoclase Anthophyllite Antimonselite Antimony native Apatite Apophyllite Aragonite Ardennite Arfvedsonite Argentite Argyrodite Armenite Arsenic native Arseniosiderite Arsenolite Arsenopyrite Asbolane Atacamite Atokite Attapulgite/palygorskite Augite Aurostibite Autunite Axinite-(Mg) Azurite Baddeleyite Bariopyrochlore Barite (baryte) Bassanite Bastnaesite Bazzite Becquerelite Behoite Beidellite Belendorffite Benitoite Benleonardite Benstonite Berborite Berthierine Bertrandite Beryl, aquamarine, heliodor, goshenite, morganite, rosterite, bixbite, emerald Berzelianite Bicchulite Bikitaite

(continued on next page)

10


Table 01.01 (continued) Mineral


Line

Biotite Birnessite Bischofite Bismite Bismuth native Bismuthinite Bixbyite Blödite Boehmite Boracite Bornite Boulangerite Bournonite Braggite Brannerite Braunite Bravoite Brazilianite Breithauptite Breunnerite Briartite Britolite-(Ce) Brochantite Brockite Bromargyrite Bromellite Bronzite Brookite Brucite Brüggenite Brunckite Brushite Bustamite Calaverite Calcite Cancrinite Canfieldite Carbocernaite Carlinite Carnallite Carnotite Carrolite Cascandite Cassiterite Cattierite Cavansite Celadonite Celestite Celsian Cerianite Ceriopyrochlore Cerite Cerussite Chabazite-(Ca) Chalcedony Chalcocite Chalcomenite Chalcophanite Chalcopyrite Chalcostibite Chamosite Charoite Chengdeite Chevkinite-(Ce) Chiavennite Chlorite Chloritoid Chloroargyrite Choloalite Chondrodite Chrisstanleyite Christite Chromite Chrysoberyl Chrysocolla Chrysoprase Cinnabarite

K(Mg,Fe)3AlSi3O10(OH,F)2 Na0.3Ca 0.1K 0.1Mn4+Mn3+O4·1.5(H2O) MgCl2·6 H2O Bi2O3 Bi Bi2S3 Mn1.5Fe0.5O3 Na2Mg(SO4)2·4(H2O) AlO(OH) Mg3B7O13Cl Cu5FeS4 Pb5Sb4S11 PbCuSbS3 Pt0.6Pd0.3Ni0.1S U0.5Ca0.3Ce0.2Ti1.5Fe0.5O6 Mn2Mn3+6SiO12 (Ni,Co,Fe)S2 NaAl3(PO4)2(OH)4 NiSb (Fe,Mg)CO3 Cu2(Fe, Zn)GeS4 (Ce,Ca,Sr)2(Ce,Ca)3(SiO4,PO4)3(O,OH,F) Cu4(SO4)(OH)6 Ca0.6Th0.3Ce0.1(PO4)·(H2O) AgBr BeO (Mg,Fe)2 (SiO3)2 TiO2 Mg(OH)2 Ca(IO3)2 ZnS Ca(HPO4)·2(H2O) (Mn,Ca)3Si3O9 AuTe2 CaCO3 Na6Ca2Al6Si6O24(CO3)2 Ag8SnS6 (Ca,Na)(Sr,Ce,Ba)(CO3)2 Tl2S KMgCl3·6 (H2O) K2(UO2)2(VO4)2·3H2O Co2CuS4 Ca (Sc,Fe) Si3O8 (OH) SnO2 CoS2 CaVSi4O11·4(H2O) K(Mg,Fe2+)(Fe 3+,Al)[Si4O10](OH)2 SrSO4 BaAl2Si2O8 (Ce,Th)O2 (Ce,Ca,Y)2(Nb,Ta)2O6(OH,F) (La,Ce,Ca)9(Mg,Fe)(SiO4)6[SiO3(OH)](OH)3 PbCO3 (Ca0.5Na,K)4[Al4Si8O24]·12H2O SiO2 Cu2S Cu(SeO3)·2(H2O) (Zn,Fe2+,Mn2+)Mn4+3O7·3(H2O) CuFeS2 CuSbS2 Fe3Mg1.5AlFe0.5Si3AlO12(OH)6 K5Ca8Si18O46(OH)·3(H2O) Ir3Fe (Ce,La,Ca,Th)4(Fe,Mg)2(Ti,Fe)3Si4O22 CaMnBe2Si5O13(OH)2·2H2O A4-6Z4O10(OH,O) A = Al, Fe, Li, Mg, Mn, Ni, Z = Al, Fe, Si (Fe,Mg,Mn)2Al4Si2O10(OH)4 AgCl CuPb(TeO3)2·(H2O) Mg3.75Fe2+1.25(SiO4)2F1.5(OH)0.5 Ag2Pd3Se4 TlHgAsS3 FeCr2O4 BeAl2O3 Cu1.75Al0.25H1.75(Si2O5)(OH)4·0.25(H2O) Ni-bearing layer silicates (willemseite) HgS

9-13-14-27-32-38-42-47-48-49-50-56-58-59 8 28-35 18 18 18 8 37 27 30-35 9 16 16 4 25 8 2 38 2 28 9-16 24 9 26 17 14 1-4 5 28 36 16 38 8 19 29 42 12-17 24 22 35 25-6 3-9 13 12 3 43 40-43-59 34 41 24 24 24 16 14-43 40 9 10 8 9 20 7 46 4 24 43 1-2-6-7-12-16-19-25-27-28-32-43-44-45-56-60 50 17 10 7 9-10 22 1 14 9 2 23


11



Line

Clausthalite Clinochlore Cr-bearing (kammererite) Clinohumite Clinoptilolite-(K) Clinozoisite Cobaltite Coffinite Colemanite Colusite Cookeite Cooperite Copper Cordierite/iolite Cordylite-(Ce) Coronadite Corundum, sapphire, ruby, padparadscha Corvusite Covellite Crandallite Crocidolite Crocoite Cryolite Cryptomelane Cubanite Cummingtonite Cuprite Cylindrite Danburite Daqingshanite-(Ce) Datolite Davidite Dawsonite Descloizite Diamond Diaspore Dietzeite Digenite Diopside Dioptase Djurleite Dolomite Donbassite Dumortierite Duranusite Dyscrasite Edingtonite Electrum Embolite Emplectite Enargite Enstatite Epidote Epistolite

PbSe Mg5(Al,Cr)2Si3O10(OH)8 Mg9Si4O16(F,OH)2 (Na,K,Ca)2–3Al3(Al,Si)2Si13O36·12(H2O) Ca2Al3(SiO4)3(OH) CoAsS U(SiO4)0.9(OH)0.4 Ca2B6O11·5 H2O Cu3(As,Sn,V,Fe,Te)S4 LiAl5Si3O10(OH)8 Pt0.6Pd0.3Ni0.S Cu Mg2Al4Si5O18 BaCe1.5La0.5(CO3)3F2 Pb(Mn4+Mn2+)8O16 α-Al2O3 Na0.6Ca0.25K0.15V8O20·4(H2O) CuS CaAl3(PO4)2(OH)5·(H2O) Na2Fe2+3Fe3+2(Si8O22)(OH)2 PbCrO4 Na3AlF6 K(Mn4+Mn2+)8O16 CuFe2S3 Mg7(Si8O22)(OH)2 Cu2O Pb3Sn4FeSb2S14 CaB2Si2O8 Sr1.2Ca0.6Ba0.2Ce0.75La0.25(PO4)(CO3)2.5(OH)0.4F0.1 CaBSiO4(OH) La0.7C 0.2Ca0.1Y0.75U0.25i15Fe5O38 NaAl(CO3)(OH)2 (Pb,Zn)Cu(OH)(VO4) C AlO(OH) Ca2(IO3)2.(CrO4) Cu9S5 CaMg(Si2O6) CuSiO2(OH)2 Cu 1.96S CaMg(CO3)2 Al5.333Si3O10(OH)8 Al 6.9(BO3)(SiO4)3O 2.5(OH) 0.5 As4S Ag3Sb BaAl2Si3O10·4(H2O) AuAg AgCl0.5Br0.5 CuBiS2 Cu3AsS4 Mg2Si2O6 Ca2Fe2.25Al0.75(SiO4)3(OH) Na 3.79Ca 0.27Mn0.04Nb1.92Ti 0.04Fe0.04 (Si2O7)2 O2 (OH)1.44 F0.56·4(H2O) MgSO4·7 H2O (Na2,K2,Ca)2[Al4Si14O36]·15(H2O) Co3(AsO4)2.8H2O CuAgSe BeAlSiO4OH Na15Ca6 (Fe, Mn)3 Zr3[Si25O73] (O,OH,H2O)3 (OH,Cl)2 Bi4(SiO4)3 Ni3MgSi6O15(OH)2·6(H2O) Cu3SbS4 Tl3AsS4 (Na2,Ca,Mg)3.5[Al7Si17O48]·32(H2O) Fe2SiO4 MnO(OH) MnWO4 La0.2Ce0.4Pr0.1Nd0.2REE0.1Nb0.9O 4 FeO(OH) Fe3+2O3·0.5(H2O) FeNb2O6 FeSe2 FeTa2O6 (Ca,Ce,Na)(Nb,Ta,Ti)2(O,OH,F)6 (FeVO4)·(H2O)

10 1 42-50 43

Epsomite Erionite-(Na) Erythrite Eucairite Euclase Eudialyte Eulytine Falcondoite (“garnierite”) Famatinite Fangite Faujasite-(Na) Fayalite Feitknechtite Ferberite (Mn wolframite) Fergusonite Feroxyhyte Ferrihydrite Ferrocolumbite (columbite-(Fe), niobite) Ferroselite Ferrotantalite Fersmite Fervanite

3 25 30 12 7-15 4 9 30-49 24 8 30-32-42-49-50-56 6 9 19-27 44 1 32 8 9 9 (and others) 9 20-12 30 24 30 25 43 16-9-6 45-46-47-49-51 27 36 9 1-7-9-14-38-42-46-49-50-51-58-59 9 9 28 7 30 8 17 43 19 17 18 9 45-46 7-9-12-17-19-41-44-47-48 13-5 28 43 3 10 14 39 18 2 9 22 43 45 8 12 24-13 7 7 13 10 13 24 6 (continued on next page)

12




Line

Florencite-(Ce) Fluocerite Fluorite Forsterite Fourmarierite Fowlerite Fraipontite Francevillite Franckeite Francolite Franklinite Freibergite Fuchsite (chromium muscovite) Gahnite Galena Galkhaite Gallite Gallobeudantite Garnet

(Ce, La)Al3(PO4)2(OH)6, (Ce,La)F3 CaF2 Mg2SiO4 Pb(UO2)4O3(OH)4·4(H2O) Zn-rich rhodonite (Zn,Al)3(Si,Al)2O5(OH)4 (Ba,P)(UO2)2(VO4)2·5H2) Pb5Sn3Sb2S14 Ca5(PO4)2.5(CO3)0.5F (Zn,Fe,Mn)Fe2O4 (Ag, Cu)12(Sb, As)4S13 K(Al,Cr)3(Si3O10(OH)2) ZnAl2O4 PbS (Cs,Tl)(Hg,Cu,Zn)6(As,Sb)4S12 CuGaS2 (PbGa3[(AsO4),(SO4)]2(OH)6) X3Y2(SiO4)3 X = Fe, Ca, Mn, Mg; Y = Fe, Al, Cr, Zr, Ti Mg5Al2(Si6Al2O22)(OH)2 Cu26Fe4Ge4S32 NiAsS PtSb1.5Bi0.5 Al(OH)3 Tl2As6Sb2S13 (K,Na)(Fe3+,Al,Mg)2(Si,Al)4O10(OH)2 Na2Mg3Al2(Si8O22)(OH)2 (Na2,Ca)Al2Si4O12·6(H2O) α FeO(OH) Au Cu12(Sb, As)4(Te, S)13 BaAl3(PO4)2(OH)5·H2O SrAl3(PO4)2(OH)5·(H2O) C Pb9As4S15 Fe2.3Fe0.5Si2.2O5 (OH)3.3 CdS Na 1.74K 0.1(Ca, Sr, Ba)0.16CO3, Fe3S4 MnO(OH) HgBi2S4 Fe7(Si8O22)(OH)2 Bi2Se3 FeSbS CaSO4·H2O Hf(SiO4) NaCl Al4Si4O6(OH)12 FeAl2(SO4)4·22(H2O) Be2(OH,F)BO3 KNa22(SO4)9(CO3)2Cl NaCa2(Fe2+4Fe3+)Si6Al2O22(OH)2 Mn2+/3+3O4 Na4Ca2Al6Si6O22S2(SO4)Cl0.5 CdS Ni3S2 Na0,3(Mg,Li)3Si4O10(FOH)2 CaFe(Si2O6) α Fe2O3 Pb10Zn(CrO4)6(SiO4)2F2 Zn4Si2O7(OH)2·H2O FeAl2O4 SnS AgTe ZnMn2O4 CoOOH (Ca,Na)2–3Al3(Al,Si)2Si13O36·12(H2O) CaV6O16·9H2O (Ca,Ce)(Al,Ti,Mg)12O19 (Pb,Sr)Al3(PO4)(SO4)(OH)6 Ag2FeSnS4 PbSnO3.n H2O (Mg,Fe)2(Al,Ti)5O10 Ba(Mn4+Mn2+)8O16 Rh0.6Pt0.3Pd0.1AsS Li2Mg3Al2(Si8O22)(OH)2

24 24 32 13-30-32-45-59 25 8 16 6-25 12 38-32 16 9-17 1-19-50 16 16 22 27 27 1-5-7-8-13-14-16-25-32-41-42-44-45-4647-49-50 30-44 16 2 4 27 22 7-8-38-59 43 43 7 19 9 13-24 24 14-20-23-38-42-46-47-50-52 16 7 16 37 7 8 23 7 10 20 28-30-31-34-37-38-40-54 39-39 35-35 27-53-54-55 31 14 37 49 8 42 16 2 15-57 9-12-16-19-20-32-46 7 1 16 49-50 12 17-10 8 3 43 6 50 19 12-17 12 49-50 8 4 13-15

Gedrite Germanite Gersdorffite Geversite Gibbsite Gillulyite Glauconite Glaucophane Gmelinite-(Na) Goethite Gold native Goldfieldite Gorceixite Goyazite Graphite Gratonite Greenalite Greenockite Gregoryite Greigite Groutite Grumiplucite Grunerite Guanajuatite Gudmundite Gypsum Hafnon Halite Halloysite Halotrichite Hambergite Hanksite Hastingsite Hausmannite Hauyne Hawleyite Heazlewoodite Hectorite Hedenbergite Hematite Hemihedrite Hemimorphite Hercynite Herzenbergite Hessite Hetaerolite Heterogenite Heulandite Hewettite Hibonite Hinsdalite Hocartite Hochschildite Högbomite Hollandite Hollingworthite Holmquistite


13



Line

Howlite Huanghoite-(Ce) Huebnerite (Fe wolframite) Humite Huntite Hutchinsonite Hyalophane Hydroboracite Hydromagnesite Hydrozincite Hypersthene Illite Ilmenite Ilvaite Indite Indium native Inyoite Iranite Isoferroplatinum Ixiolite Jacobsite Jadeite Jamesonite Jarosite Jentschite Jeremejevite Jervisite Joaquinite Johannsenite Jordanite Jordisite Juonniite Kainite Kaolinite, dickite, nacrite Kemmlitzite Kernite Kieserite Klockmannite Knopite Koesterite Kolbeckite Kolymite Kornerupine Kotoite Krennerite Kristiansenite Kutnahorite Kyanite Kylindrite Langbeinite Lanthanite Larnite Laumontite Lautarite Lawsonite Lazulite Lazurite Leonite Lepidocrocite Lepidolite Leucite Libethenite Linneite Lithiophilite Lithiophorite Livingstonite Loellingite Lomonosovite Lonsdaleite Loparite Lorandite Lorenzenite Loughlinite Ludwigite Luzonite Mackinawite

Ca2B5SiO9(OH)5 BaCe(CO3)2F FeWO4 (Mg,Fe)7(SiO4)3(F,OH)2 CaMg3(CO3)4 TlFeS2 K0.75Ba0.25Al1.75Si2.25O8 CaMgB6O8(OH)6·3(H2O) 3MgCO3 Mg(OH)2·3H2O Zn5(CO3)2(OH)6 MgFeSi2O6 (K,H3O)(Al,Mg,Fe)2 (Si,Al)4O10·H2O FeTiO3 CaFe3(SiO4)2(OH) FeInS2 In CaB3O3(OH)5·4(H2O) Pb10Cu(CrO4)6(SiO4)2F1.5(OH)0.5 Pt2.25Pd0.75Fe 0.75Cu 0.25 (Ta,Nb,Sn,Mn,Fe)O2 Mn2+0.6Fe2+0.3Mg0.1Fe3+1.5Mn3+0.5O4 NaAlSi2O6 Pb4FeSb6S14 KFe3(SO4)2(OH)6 TlPbAs2SbS6 Al6B5O15F2.5(OH)0.5 (Na,Ca,Fe)(Sc,Mg,Fe) Si2O6, NaBa2FeCe2(Ti, Nb)2(SiO3)8(OH, F)·H2O CaMn(Si2O6) Pb14(As,Sb)6S23 MoS2 CaMgSc(PO4)2(OH)·4(H2O) Mg(SO4)KCl·3(H2O) Al4Si4O10(OH)8 (Sr,Ce)Al3(AsO4)(SO4)(OH)6 Na2B4O7·4 H2O MgSO4·H2O CuSe (Ca,Ce)TiO3 Cu2ZnSnS4 Sc(PO4)·2 H2O Cu7Hg6 Mg3.5Fe0.2Al5.7(SiO4)3.7(BO4)0.3O1.2(OH) Mg3B2O6 AuTe2 Ca2ScSn(Si2O7)(Si2O6OH) CaMn0.6Mg0.3Fe0.1(CO3)2 Al2SiO5 Pb3Sn4SbS14 K2Mg2(SO4)3 (Ce,La)2(CO3)3·8(H2O) Ca2(SiO4) CaAl2Si4O12·4(H2O) Ca(IO3)2 CaAl2Si2O7(OH)2·(H2O) MgAl2(PO4)2(OH)2 Na3CaAl3Si3O12S K2Mg(SO4)2·4(H2O) γ FeO(OH) KLi2AlSi4O10F(OH) KAlSi2O6 Cu2(PO4)(OH) Co3S4 LiMn(PO4) (Al,Li)Mn3+/4+O2(OH)2 HgSb4S7 FeAs2 Na5Ti2O2Si2O7(PO4) C (Na,Ca,Ca) (Ti,Nb)O3 TlAsS2 Na2Ti2Si2O9 Na2Mg3Si6O16·8(H2O) Mg2FeBO5 Cu3AsS4 Fe0.75Ni0.25S0.9

30 24 12 44 28 22 41 30 28 16 46 19-38-43-54-55-59 5 9 (and others) 16 16 30 1 4 13 8 44-46-50 20 9 22 30 13 5 9-16-46 16 8 13 35 7-8-16-19-20-23-27-32-38-40-49-54-55-56-59 24 30 28-35 10 24 12 13 23 30 30 19 13 8 46-48-49-50 12 35 24 45 43 36 43 49 42 35 7 15 42-51 9 3 13-15-38 8 23 21 13 51 13 22 13 61 30 9 7 (continued on next page)

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Maghemite Magnesite Magnetite Magnetite/vanadiferous Malachite Malayaite Maldonite Manganite Manganotantalite Manjiroitite Marcasite Margarite Marialite scapolite Mariposite Martite Matildite Maucherite Meionite scapolite Melanterite Meneghinite Mercury native Merenskyite Metacinnabarite Metastibnite Meyerhofferite Miargyrite Microlite Millerite Mimetite Minasragreite Minnesotaite Mirabilite Molybdenite Molybdite Monazite Monetite Montmorillonite Montroseite Montroydite Mordenite Mosandrite Moschelite Moschellandsbergite Mottramite Mtorolite/mtorodite Murmanite Muscovite Nagyagite Nahcolite Natrite Natrolite Naumanite Nepheline Nepouite Neptunite Nesquehonite Nickelite/nickeline Nimite (“schuchardite”) Ningyoite Niter Nitratine Nontronite Nosean Nsutite Nyerereite Olivenite Olivine Opal Orpiment Otavite Paladium Paragonite Paralstonite Paraschachnerite Pargasite Parisite Parsettensite

γ Fe2O3 MgCO3 Fe3O4 (Fe,V)3O4 Cu2(CO3)(OH)2 CaSnSiO5 Au2Bi MnO(OH) MnNb2O6 (K,Na)(Mn4+Mn2+)8O16·n H2O FeS2 CaAl4Si2O10(OH)2 3 (NaAlSi3O8)·NaCl K(Al,Cr)2(Al,Si)4O10(OH)2 γ Fe2O3 AgBiS2 Ni3As2 3 (CaAlSi2O8).(CaCO3) Fe2+(SO4)·7(H2O) Pb13CuSbS24 Hg Pd0.9Pt0.1Te1.8Bi0.2 HgS Sb2S3 Ca2B6O6(OH)10·2(H2O) AgSbS2 Na1.5Ca0.5Ta2O6.6(OH)0.3F0.1 NiS Pb5(AsO4)3Cl VO(SO4)·5(H2O) Fe2.5Mg0.5Si4O10(OH)2 Na2(SO4)·10(H2O) MoS2 MoO3 (Ce,La,Nd,Th)PO4 Ca(HPO4) (Na,Ca)0,3(Al,Mg)2Si4O10(OH)2·n(H2O) (V,Fe)O(OH) HgO Na1.1Ca0.5K0.1Al2.2Si9.8O24·5.9(H2O) Na2Ca4(Ce,Y)(Ti,Zr)(Si2O7)2OF3 Hg2I2 Ag2Hg3 Pb (Zn,Cu)(OH)(VO4) chrom chalcedony Na3Ti3.6 Nb0.4(Si2O7)2O4·4(H2O) KAl2(Si3Al)O10(OH,F)2 AuPb(Sb,Bi)Te2–3S6 Na(HCO3) Na2(CO3) Na2Al2Si3O10·2(H2O) Ag2Se (Na,K)AlSiO4 Ni3Si2O5(OH)4 KNa2LiFe 1.5Mn 0.5Ti2Si8O24 MgCO3.3 H2O NiAs Ni2.6Mg1.7AlFe3+0.4Fe2+0.3Si3AlO10.3(OH) 7.7 (U,Ca,Ce)2(PO4)2·1–2(H2O) KNO3 NaNO3 Na0.3Fe2(Si,Al)4O10(OH)2·n(H2O) Na8Al6Si6O24(SO4)·(H2O) Mn4+0.85O1.7Mn2+0.15(OH)0.3 (Na 0.82K 0.19)2(Ca, Sr, Ba) 0.975(CO3)2 Cu2(AsO4)(OH) (Mg,Fe)2SiO4 SiO2·n(H2O) As2S3 CdCO3 Pd NaAl3Si3O10(OH)2 BaCa(CO3)2 Ag3Hg2 NaCa2(Mg,Fe)4Al(Si6Al2)O22(OH)2 Ca(Ce,La)2(CO3)3F2 (K,Ca)8(Mn,Mg)49[(OH)50/Si64Al8O168]·20H2O

7 28 7 6 9 12 19 8 13 8 7-31 32-49-50-59 5-7-38-42-47-50 1 7 16-17-18 2 5-7-38-42-47-50 31 20 23 4 23 20 30 17 13 2 16 6 7-56 35-37 11 11 24 38 15-27-30-54-57 6 23 43-57 39 23 23 16-9-6 1 5 9-12-14-32-41-43-47-59 19-10 43 43 43-52 10-17 27 2 5 28 2 2 24 36 36 54-57 42 8 37 9 1-2-7-9-19-43-45 8-9-20-28-32-40-53-55-57 21 16-38 4 43-56-59 32 23-17 50 24 8


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Pascoite Patronite Pearceite Pecoraite Pectolite Penroseite Pentlandite Periclase Perovskite (dysanalyte (Nb), knopite (Ce)) Petalite Petzite Pezzottaite Phenakite Phillipsite-(K) Phlogopite Phoenicochroite Phosphophyllite Phosphosiderite Piemontite Pimelite Pitchblende Plancheite Platinum Plumbogummite Polhemusite Polianite Pollucite Pollucite Polybasite Polydymite Polyhalite Porpezite Poudretteite Powellite Prehnite Pretulite Priceite Priderite Proustite Psilomelane Pumpellyite Pyrargyrite Pyrite Pyrochlore Pyrolusite Pyromorphite Pyrophyllite Pyroxmangite Pyrrhotite Quartz, lechatelierite, keatite, stishovite, christobalite, tridymite Radian barite Rammelsbergite Ramsdellite Rancieite Realgar Reniérite Rhabdophane Rheniite Rhodizite Rhodochrosite Rhodonite Richterite Riebeckite Rinkite Rodalquilarite Romanèchite Romeite Roquesite Roscoelite Roselite Rustenburgite Rutile Safflorite Samarskite Samsonite

Ca3V10O28·17(H2O) VS4 Ag16As2S11 Ni3Si2O5(OH)4 NaCa2Si3O8(OH) (Ni,Co,Cu)Se2 (Ni,Fe)9S8 MgO (Ca,Fe,REE)TiO3 Li0.92Al0.99Si3.99O10 Ag3AuTe2 Cs(Be2Li)Al2Si6O18 Be2SiO4 K0.8Na0.7Ca0.7Si5.2Al2.8O16·6(H2O) KMg3AlSi3O10F(OH) Pb2CrO5 Zn2(Fe,Mn)(PO4)2·4(H2O) Fe3(PO4)·2(H2O) Ca2Al1.8Mn2+0.9Fe2+0.3(SiO4)3(OH) Ni3Si4O10(OH)2·4(H2O) U3O8 Cu8Si8O22(OH)4·(H2O) Pt PbAl3(PO4)2(OH)5·(H2O) Zn0.75Hg0.25S MnO2 Cs0.6Na0.2Rb0.04Al0.9Si2.1O6·(H2O) (Cs,Na)2Al2Si4O12·(H2O) Ag16Sb2S11 Ni3S4 K2Ca2Mg(SO4)4·2(H2O) AuPd KNa2BSi12O30 CaMoO4 Ca2Al2Si3O10(OH) Sc(PO4) Ca4B10O19·7 H2O (K,Ba)(Ti,Fe)8O16 Ag3AsS3 Ba (H2O)Mn3+5O10 Ca2MgAl2(SiO4)(Si2O7)(OH)2·(H2O) Ag3SbS3 FeS2 Na1.5Ca0.5Nb2O6(OH)0.75F0.25 MnO2 Pb5(PO4)3Cl Al2Si4O10(OH)2 MnSiO3 Fe7S8 SiO2

6 6 17 2 46 10 2 28 5 15 19 14-15 14 43 44-50-51-59 1 38 49 47-48 2 25 9 4 24 23 8 15 13-15-43 17 2 35 19 30 11 43-46-47-52 13 30 51 17 8 41-43 17 7-31 13 8 16 2-23-27-30-49-55-56 8 7 40

(Ba, Ra)SO4 NiAs2 MnO2 Ca0.75Mn2+0.25 Mn4+4O9·3(H2O) As4S4 (Cu,Zn)11(Ge,As)2Fe4S16 (Ce,La)PO4·(H2O) ReS2 (K,Cs)Al4Be4(B,Be)12O28 Mn(CO3) (Mn,Fe,Mg,Ca)SiO3 Na2CaMg3Fe2(Si8O22)(OH)2 Na2Fe2+3Fe3+2(Si8O22)(OH)2 Na2.5Ca4CeTi0.75Nb 0.25(Si2O7)2O3F HFe3+2(TeO3)4Cl (Ba,H2O)2(Mn4+,Mn3+)5O10 (Ca,Fe,Mn,Na)2(Sb,Ti)2O6(O,OH,F) CuInS2 KV2(OU/)2/AlSiO3O10 Ca2(Co, Mg)(AsO4)2·2H2O Pt2.25Pd0.75Sn TiO2 CoAs2 Y0.2REE0.3Fe0.3U0.2Nb0.8Ta0.2O4 Ag3MnSb2S6

25 2 8 8 21 16 24 11 14 8 8 13 7-25-44 5-13 10 8 20 16 6-25 3 4 5 3 24-13-13-25 20-17 (continued on next page)

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Sanbornite Sanidine, orthoclase, microcline Santafeite Santanaite Saponite Sapphirine Sarabauite Sassolite Sauconite Scandiobabingtonite Scapolite Schachnerite Schapachite Scheelite Schoepite Schwatzite (mercurian tetrahedrite) Scorodite Scorzalite Seginite Seinajokite Sellaite Senarmontite Sepiolite Serendibite Serpentine, antigorite, chrysotile Siderite Siegenite Sillimanite Silver native Sinhalite Sklodowskite Skutterudite Smithsonite Smythite Sodalite Sohngeite Sperrylite Sphalerite Sphene Spherocobaltite Spinel-group minerals Spodumene-hiddenite-kunzite Stannite Staurolite Stavelotite Steenstrupin Stephanite Stevensite Stibiconite Stibiopalladinite Stibnite Stilbite-Ca Stilleite Stilpnomelane Stokesite Stolzite Strelkinite Stromeyerite Strontianite Strüverite Suanite Sugilite Sulfur native Svanbergite Sylvanite Sylvite Synchysite Szaybelyite Taaffeite Tachyhydrite Talc Tapiolite Tarkianite Teallite Tennantite Tenorite Tephroite

BaSi2O5 KAlSi3O8 (Mn,Fe,Al,Mg)2(Mn4+,Mn2+)2(Ca,Sr,Na)3(VO4,AsO4)4(OH)3·2(H2O) Pb9Pb2CrO16 (Ca,Na)0,3(Mg,Fe)3(Si,Al)4O10(OH)2·4(H2O) (Al,Mg)8(Al,Si)6O20 CaSb10O10S6 B(OH)3 (Na0.3(Zn,Mg)3 (Si,Al)4.OH2·nH2O) Ca2(Fe2+, Mn)ScSi5O14(OH) Na2Ca2Al3Si9O24Cl Ag1.1Hg0.9 AgBiS2 CaWO4 (UO2)8O2(OH)12·12(H2O) (Cu,Hg)12Sb4S13 Fe(AsO4)·2(H2O) Fe0.75Mg0.25Al2(PO4)2(OH)2 PbFe3H(AsO4)2(OH)6 FeSb2 MgF2 Sb2O3 Mg4 [(OH)2 Si6O15]·2H2O + 4H2O Ca2Mg4.5Al1.5Si3.6Al1.8BO Mg3Si2O5(OH)4 FeCO3 (Co,Ni)3S4 Al2SiO5 Ag MgAlBO4 (H3O)2Mg(UO2)2(SiO4)2·4(H2O) CoAs3 ZnCO3 Fe6.75Ni2.25S11 Na8Al6Si6O24Cl2 Ga(OH)3 PtAs2 ZnS CaTiSiO5 CoCO3 (X)(Y)2O4 X = Mg, Zn, Fe, Mn, Y = Al, Cr, Ti LiAl(Si2O6) Cu2FeSnS4 (Fe,Mg)2Al9(Si,Al)4O20(O,OH)4 (La) (La,Sc,Nd)3 Mn Cu(Mn,Fe)26 O30|(Si2O7)6 Ce (Na14Ce6Mn2+Mn3+Fe2+ 2 (Zr,Th)(Si6O18)2(PO4)7·3H2O) Ag5SbS4 (Ca0.5,Na)0.33(Mg,Fe)3Si4O10(OH)2·n(H2O) Sb3O6(OH) Pd3Sb Sb2S3 NaCa4Al8Si28O72·30(H2O) ZnSe K0.7Fe3.3Mg1.4Fe3.3Si10Al2O24(OH)3·2(H2O) CaSnSi3O9·2H2O PbWO4 Na2(UO2)2V2O8·6(H2O) CuAgS SrCO3 (Ti,Ta,Nb,Fe)2O4 Mg2B2O5 KNa2(Fe,Mn,Al)2Li3Si12O30 S8 SrAl3(PO4)(SO4)(OH)6 (Au,Ag)2Te4 KCl CaCe(CO3)2F MgB2(OH) Mg3Al8BeO16 CaMg2Cl6·12(H2O) Mg3Si4O10(OH)2 FeTa2O6 (Cu,Fe)(Re,Mo)4S8 PbSnS2 (Cu,Fe)12As4S13 CuO Mn2(SiO4)

33 6-12-13-27-30-41-45-51 25 1 43-57 30-49-50 20 30 16 13 5-7-38-42-47-50 23-17 17 12 25 23 21 38 21 20 32 20 61 30 1-2-3-7-28-32-44-45-52 7 3 30-32-48-49-50 17 30 25 3 16 7 14-42 27 4 16 5 3 1-7-30-45-47-50-51 15 12 45-49 13 39 16-17 57 20 4 20 12-43 10 7 12 12 25 17 34 13 30 8 S 49 19 35-35 24 30 14 35 44-56 13 11 12 9 9 9-45


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Tetradymite Tetraedrite Thenardite Thomsonite-(Ca) Thoreaulite Thorianite Thorite Thortveitite Thorutite Thulite Tiemannite Tinaksite Tincalconite (“borax”) Tinzenite Tirodite Tischendorfite Tlapallite Todorokite Topaz Torbernite Tourmaline, rubellite, indigolite, elbaite, dravite, schoerl, achroite Tremolite Triphylite Triplite Troillite Trona Tsumgallite Tugtupite Tulameenite Tungstenite Tungstite Turquoise Tyrrelite Tyuyamunite Ulexite Ullmannite Ulvite Umangite Uraninite Uranophane Uranospinite Uranothorite Uvarovite Valentinite Valleriite Vanadinite Vandendriesscheite Vaughanite Vauquelinite Vermiculite Vesuvianite Villiaumite Violarite Vivianite Volchonskoite Voltaite Wadeite Wickmanite Willemite Willemseite Witherite Wodginite Wollastonite Woodhouseite Woodruffite Wulfenite Wurtzite Xanthochroite Xanthophyllite/clintonite Xenotime-(Y) Zincite Zinckenite Zinnwaldite Zircon Zoisite (ortho/clinozoiste) Zunyite

Bi2Te2S (Cu,Fe)12Sb4S13 Na2(SO4) NaCa2Al5Si5O20·6(H2O) SnTa2O7 ThO2 ThSiO4 (Sc,Y)2 (Si2O7), Th0.4U0.4Ca0.2Ti2O3(OH)3 (Ca,Mn)2Al3(SiO4)(Si2O7)O HgSe K2NaCa1.75 Mn0.25Ti0.85Fe0.15Si7O19(OH) Na2B4O7.10 H2O (Ca, Mn, Fe) Al2BSi4O15(OH) Na2.5Mn0.5Mg4Fe3+(Si8O22)(OH)2 Pd8Hg3Se9 H6Ca1.5Pb0.5Cu3(SO4)(TeO3)4(TeO6) Na0.2Ca0.05K0.02Mn4+4Mn3+2 O12·3(H2O) Al2(SiO4)F1.1(OH)0.9 Cu(UO2)2(PO4)2·11(H2O) X1Y3Al6B3Si6(OH)4 X = Na, Ca , Y = Mg, Li, Al, Fe Ca2Mg5(Si8O22)(OH)2 LiFe(PO4) (Mn,Fe,Mg,Ca)2(PO4)(F,OH) FeS Na3(HCO3)(CO3)·2H2O GaO(OH) Na4AlBe(Si4O12)Cl Pt2FeCu WS2 WO3·H2O CuAl6(PO4)4(OH)8·4(H2O) Cu1.8Co0.9Ni0.3Se4 Ca(UO2)2(VO4)2·5 H2O NaCaB5O9·8 H2O NiSbS Fe2TiO4 Cu3Se2 UO2–U3O8 CaH2(SiO4)2(UO2)·5(H2O) Ca(UO2)2(AsO4)2·10(H2O) (U,Th)SiO4 Ca3Cr2(SiO4)3 Sb2O3 Fe2+2.2Cu1.8S4Mg1.7Al1.3(OH)2 Pb5(VO4)3Cl Pb(UO2)10O6(OH)11·11(H2O) TlHgSb4S7 Pb2Cu(CrO4)(PO4)(OH) (Mg,Fe,Al)3(Al,Si)4O10(OH)2·4(H2O) Ca10Mg2Al4(Si2O7)2(SiO4)5(OH)4 NaF FeNi2S4 Fe3(PO4)2·8(H2O) Ca0.Mg 0.1Cr1.2Mg0.8Fe0.3Si3.5Al0.5O10(OH)2·3.6 (H2O) K2Fe2+5Fe3+3Al(SO4)12·18(H2O) K2ZrSi3O9 MnSn(OH)6 Zn2SiO4 Ni2.25Mg0.75Si4O10(OH)2 BaCO3 Mn(Sn,Ta)(Ta,Nb)2O8 CaSiO3 CaAl3(PO4)(SO4)(OH)6 (Zn,Mn2+)2Mn4+5O12·4(H2O) PbMoO4 ZnS CdS CaMg2.2Al0.7Al2.7Si1.3O10(OH)2 YPO4 ZnO PbSb2S4 KLiFeAl2Si3O10F 1.5(OH)0.5 ZrSiO4 Ca2Al3Si3O12(OH) Al13Si5O20(OH)16F2Cl

10-18 9 35-37 43 12-13 26 26 13 26-25 48 10 46 30 8 47 4 10 8 12-13-14-30-32-42-49 25 30 44-49-50-52-58 15-38 12-15-38 7 37-43 27 14 4 12 12 9 10 6-25 30 20 5 10 25 25 25 26-25 1 20 2 16-6 25 23-22 1 38-44-50-58 9 (and others) 32 2 9 2 31 51 12 16 2 33 13 9-20-32-45-46-47-52 55 8 11 16 16 7 24 16 20 15 39 41-48 19

18


Purely descriptive parameters are used to constitute the basis for the first-order subdivision in the present classification scheme (Fig. 01.01). Both parameters, host and commodity may be coupled in a spreadsheet used in the daily routine on the PC (“spreadsheet classification scheme”) with lines (commodities) and columns (host rock and host structure) so that the relations between the element and mineral looked for on one hand and the host rock and host structure on the other are more or less self-explanatory. Some users might call this classification scheme a “chessboard classification scheme”, because playing chess has some similarities with exploration. Elements and minerals are required by the industry, rocks and structures are the targets during exploration and mining. In the x–y plot of Fig. 01.01, the variation of the host environments is shown on the x-axis and the mineralogical and chemical variation along the yaxis. Using the classification scheme in the strict sense of the word and using a spreadsheet of an .xls file (EXCEL) renders the scheme more attractive because it becomes extremely versatile. The reader can scroll up and down to see which commodity you might expect in a specific lithological realm and from left to right to get an overview where you can find a specific commodity in the magmatic, sedimentary or metamorphic realms and also find a cross over in the field of coal petrography where equivalent studies have been conducted and published in the “International Journal of Coal Geology”. A combination of the line and column codes defines the square and gives the code of type of deposit listed in the tables and used as subheadings in the text. Economic geology is a subject matter in constant change. If a new type of an ore deposit has been discovered or a mineralization has been redefined for scientific or economic reasons, no matter, key them in. The digital arena provides space for all of these arrangements and renders the spreadsheet classification scheme user-friendly. What about deposits which show up in different lines, as exemplified by the VMS deposits which are significant producers of Co, Cu, Pb, Zn, Ag and Au? The major commodity recovered from the deposit or the target mineral seems to be the most appropriate qualifier. The emplacement of an ore deposit and the enrichment of element by some orders of magnitude is rarely a monophase process. Not surprisingly, there may be recognized certain variations from the reference types listed in the “chessboard classification scheme”. The mineralization may either be the product of a complex polyphase mineralizing process, reflecting a combination of different types or may be a type so far unknown to the scientific community or to the author. Current metallogenetic models and concepts were not sidelined in this classification scheme. They are mosaic stones to paint greater pictures and referred to later on a second- or third-order level when it comes to the discussion of the individual deposits. Production figures and reserve calculations are important in economic geology but rarely presented for individual deposits rather than commodities. The Internet — see database mentioned previously — is cramped with beautiful photographs of minerals and micromounts and full-color publications presenting them are numerous whereas images showing specimens of ordinary ore and industrial minerals are scarcely found in the literature and are often hidden in company reports. Therefore in this paper ore takes priority over mineral as far as the presentation of rocks and minerals is concerned. It is mandatory, when dealing with gemstones in a paper like that, that the common minerals are presented in images to give an idea how the different varieties of gemstones differ from each other. 2.2. The host of mineral deposits Rocks are either hosts of mineral deposits or they are used themselves as mineral raw materials. The mineralogical composition of common magmatic, sedimentary and metamorphic rocks referred to in the succeeding text is listed in Table 02.01, consulting the “Glossary of Geology” (Bates and Jackson, 1987).

2.2.1. Magmatic host rocks Magmatic host rocks in the header of the diagram of Fig. 01.01 are arranged in order of increasing SiO2 content as if they had originated by fractional crystallization from an ideal parental mantle magma with the ultrabasic rocks plotted on the left-hand side and felsic rocks on the right-hand side (Carmichael et al., 1974). The most recent geological classification of magmatic rocks into plutonic, subvolcanic, dikes and extrusive/effusive rocks and their wide range of outward appearances has been published in the form of an easyto-handle field guide by Thorpe and Brown (1985). The first-order classification of principal groups of mineral deposits is based on a mineralogical and/or chemical subdivision of magmatic rocks. A closer look at the various classification schemes reveals little change in the attributes of rock types throughout the recent past so that this can be considered a rather sound basis (Iddings, 1909; Holmes, 1920; Johannes, 1931; Niggli, 1954; Hatch et al., 1972; Streckeisen, 1976, 1980; Cox et al., 1979; Ehlers and Blatt, 1982; Wimmernauer, 1985; Le Bas et al., 1986). The double-triangular plot also called Q–A–P–F diagram (quartz, alkaline feldspar, plagioclase, and foid) designed by Streckeisen (1976) allows for a subdivision of volcanic and plutonic magmatic rocks based on their mineralogical composition. The Q–A–P–F diagram combined with the pigeonhole diagram that uses the SiO2 content (wt.%) and the sum of Na2O and K2O (wt.%) yields a subdivision into ultrabasic rocks, basic, intermediate and acidic magmatic rocks suitable to describe the host rock lithologies used in this spreadsheet classification scheme. Alkaline magmatic rocks whether they are silica deficient and/or alumina deficient (ekeritic, miaskitic, and agpaitic) can be grouped also within this series. 2.2.2. Ore-bearing structures Sedimentary and magmatic host lithologies in the diagram are separated by two columns representative of structural elements. Megabreccias including fragments of various lithologies cannot be linked to a particular host rock lithology. Their striking character is the irregular shape of their fragments and structural control (Laznicka, 1988). Vein-type deposits in this classification scheme are defined as fault-related mineralizations unrelated to any magmatic or sedimentary host rock lithologies or activities, some of which are described in the succeeding paragraph. Mineralized structure zones, fissures and veinlets are ubiquitous in mineral deposits. Criss-crossing veinlets most commonly occur in acidic and intermediate plutonic and subvolcanic magmatic rocks leading into disseminated mineralization where mineral grains are scattered across the magmatic host rocks and their exocontacts. This interlaced network of irregularly shaped Cu-, Au- and Mo-bearing veins forms an integral part of what is described later in this study as hydrothermal Cu-, Au- and Mo porphyry system and, hence, dealt with in the chapters on Cu, Au and Mo. Alteration pipes and fissures mineralized with Fe and Cu sulfides are also found in another type of world-class-deposits, underneath massive sulfide base metal deposits. Vein-type Sn and W mineralizations in the immediate surroundings of acidic intrusions as well as in copulas of a granitic intrusion itself are treated as part of a granite-related mineralization rather than a vein mineralization of its own. They never occur far off the granite with which they are genetically associated. Fault-related or structure-related mineralization has been subdivided into shallow and deep veins. These mineralizing processes are veritably unrelated to any magmatic, sedimentary or metamorphic lithologies in the close vicinity. Shallow veins, as viewed in plan view in the current study, show a wide aerial extension in the basement rocks and near the basementplatform interface, named unconformity (or disconformity). Although a preferred orientation of individual veins is often seen, they do not show up in platform sediments far off the uplifted basement (Teuscher and Weinelt, 1972). These veins were emplaced in the


aftermaths of an orogeny when the magmatic and metamorphic basement rocks where uplifted and subjected to erosion or during renewed phases of uplift and erosion as a long-range effect of distant orogenic processes. In both cases a subhorizontal plane called unconformity is in control on the aerial extension of these veins which do not reach very deep into the host rock underneath (2-D shallow veins). Unconformities or disconformities may come into existence within a series of platform sediments which were only gently tilted and close to a domal structure caused by salt, mud or magmatic intrusions. Such a setting is exemplified by the Late Variscan and the Subhercynian (or Laramide) unconformities, one developed in the wake of the Variscan orogeny and the other as a long-distant effect of the Alpine orogeny (Dill et al., 2008a,b). The distribution of the Permo–Mesozoic rocks overlying the folded Variscides is controlled by eustacy and tectonic events related to the break-up of Gondwana. The question why the 2-D vein mineralization is there can be answered by giving reference to the horizontal plane and large regional faults serving as master channel ways. The question when the 2-D vein mineralization came into existence can be answered by giving reference to various thermal events. Major Permian to Jurassic tectonic events which influenced metallogenesis along the post-Variscan unconformity include (1) Stephanian to early Permian rifting accompanied by extensive magmatism, (2) Triassic to Jurassic opening of oceanic domains in the Alpine region, and (3) Mid-Triassic to early Jurassic extension in the extra-Alpine region. Isotope data increasingly show the significance of a thermal event at the Triassic–Jurassic boundary for ore formation in Central Europe (Wernicke and Lippolt, 1997). Deep veins mostly occur in uplifted basement blocks; in plan view they show the close relationship to deep-seated lineamentary fault zones, or structural elements of the folded host rock terrain which can be traced over a long distance. Rarely, they are found near the edge of basement uplifts, never are they reported from the platform sediments. Mineralization is related to thrust and shear zones acting as channel ways and loci of mineralization. At a deeper level these structures may convert into zones of mobilization along which felsic intrusions were able to penetrate the crystalline basement. The Variscan and Alpine orogenies used as an example for a metallotect hosting 2-D shallow veins, can also be used as an example for thrustbound and fold-related metamorphogenic deposits (Dill et al., 2008a, b). Activation of the continental margin resulted in the initiation of subduction in the early Late Devonian, closing of the Rhenohercynian Basin and subsequent Variscan collisional tectonics led to the accretion of shelf sediments and a magmatic activity along deepseated suture zones (Franke and Oncken, 1990). The Variscan ore shows pervasive textural distortion and strong mylonitization. It is found in cleaved psammo-pelitic series and developed along the fold axes of the Variscan anticlines. The Alpine analogues of thrustrelated or syn-orogenic veins developed during the Late Alpine deformation and were controlled by sinistral wrenching along ENEtrending faults within the Eastern Alps and the subsequent eastward lateral escape of crustal wedges towards the Pannonian realm accompanied by orogen-parallel extension and tectonic denudation of metamorphic core complexes (Ratschbacher et al., 1991; Neubauer et al., 1999). Two groups of deposits have been recognized within these syn- to late orogenic regimes (Prochaska, 1993; Pohl and Belocky, 1994). Immediately after the Cretaceous orogeny deposits were produced at high pressures by metamorphic fluids of very high salinity as a result of devolatilization of subducted South Pennic rocks. During the Paleogene orogeny under relatively low pressures, CO2-rich fluids and low to moderate salinities evolved.

19

Isotopic data indicate a deep crustal or even mantle source for CO2, while the water may have mixed sources, both surficial and metamorphic. Tectonic control of these Oligo–Miocene mineralizations is transtensional faulting, which exposed hot metamorphic rocks to fluid convection along brittle structures. These deposits conform best to the model of metamorphogenic metallogenesis by retrograde leaching- see also section on metamorphic mobilization. The color of the boxes direct the reader's thoughts to the principal realms which these mineralized structure are affiliated with (Fig. 01.01b). The 2-D vein clusters are part of large-scale basinal mineralization with closer affinities to mineralization in duricrusts and siliciclastic–calcareous shelf sediments than any deep-seated heat source. By contrast, the 1-D vein zones are deeply rooted in the basement and linked even to deeper sources of elements (mantlederived). 2.2.3. Sedimentary host rocks Sedimentary rocks result from deposition of products of chemical, physical and biological weathering and erosion of magmatic, metamorphic or pre-existing sedimentary rocks. Mechanically disintegrated material and chemical compounds are transported more or less far off the source rocks, where they are deposited and undergo lithification through time. Many classification schemes for the various rock types have been put forward and compiled in the various textbooks (Leighton and Pendexter, 1962; Füchtbauer and Müller, 1970; Selley, 1976; Flügel, 1978; Scoffin, 1987; Tucker and Wright, 1990; Friedman et al., 1992; Adams and MacKenzie, 1998; Miall, 2000; Tucker, 2001). The sedimentary host rocks are arranged in a similar way like their magmatic counterparts, reflecting a differentiation from the basin edge to the basin center or in lithological terms from autochthonous chemical residues evolving in or proximal to the provenance area via siliciclastic deposits to calcareous and evaporitic sediments with the most soluble representative on the right-hand side. 2.2.4. Organic material and special host rocks Sediments containing organic matter such as lignite and coal or hydrocarbons, which sometimes are enriched in inorganic compounds do not directly fit in this differentiation scheme and therefore they were given a column of their own. They may be classified according to suggestions made by Stach (1982), Diessel (1992), Moore and Shearer (2003) and Kalkreuth (2004), and for petroleum deposits by Tissot and Welte (1978), The missing link between organic material and mineralization has been provided by Seredin and Finkelman (2008), who gave a detailed account of metalliferous coal. 2.3. Type of commodity (inorganic raw material) 2.3.1. Ore minerals There is no yet common consensus about the meaning of the word ore. The more stringent explanation defines ore as a rock composed of ore minerals and gangue. From the standpoint of a metallurgist heavy metals (e.g. Pb) may be won at a profit from ore minerals, e.g. galena (PbS). Gangue comprises the various waste minerals for which there is no use at the current time and which after beneficiation are dumped at the mining site. Some hundred years ago, a vein mineralization was mined for Cu ore made up of chalcopyrite forming the ore mineral and siderite plus dolomite considered then as gangue minerals (Dill, 1985a). Siderite and dolomite were used by our ancestors as backfill in the mined-out

Fig. 01.01a. (see next pages). Chessboard classification scheme. Bold arrowheads lateral facies changes of host rock lithologies. Thin arrowheads mineralization emplaced at the contact of different lithologies or alternative placement of mineralization. Boxes show element composition, rocks and/or minerals of economic interest. Elements set in brackets refer to elements present in abnormally high amounts but not necessarily show up by minerals of their own, e.g. (As) in coal. Terms firmly entrenched in the literature and used to describe the ore types are given in italics. The locus typicus of ore type is written in italics and set in rounded brackets. For color coding and further symbols see legend Fig. 01.01b.

20


I Type 01 commodity

ELEMENT

Mineral ROCk

:
••c 0

N

,

~

~ ~ ~>

.s:

-~

e •

'~

~

•c "§

u

H

G

""",e

~

'.

•c ••E

~

.,

'~

2

••c ~ •E

:J

• :E0 • > w

~

J

.~

~

o

mrte--

huntile

magnesia

mag"""""

(SabJdIa EI

seawater

(Granssen)

-.}

hunUte· hydromagnesite· magnesite

magnesite-Fe

Mg-K-Br bnnescamalllle-ideseme kainite--bischoftI.e surface and

(Vettsch)

(Lefkara)

subtemmean salt deposits

magnesite-B (Bela SUms)

magnesrte-W (Tux)

magnesite

F-(U) unconformityrelated

171 limestones t a, Iraverti s C8 los

calcite

Ioe,.

salltlBS (Lamaca)

ynI- dolomite

( KunwarafS)

calcile-Pb-Zn-Ba-

Na-Mg-(epsomite)

""_ ~ 1L,;._1i'" c. ,,"os

Fig. 01.01a (continued).

pearts cora ls

30


A

B

D

C

E

29b axinite

borate dimension

danburite

stones

B

dalalile tourmaline 5.S. dravile

BO RON

Boron silicates Borates

30a

tourmaline -

poudretteite

danburiledumortierite-

carbonalite·fenile

jeremejevite-

dumortierite elbaite indigolite jeremejevite

30b

komerupine painite rubellite serendibile sinhalite poudretteite

30c 30d

komerupine (gem) granitic pegmatite

-

tourma~Cr)

axinites- gypsum! anhydrit~Sr· Mg)

sulfur

oR shale and tar sand F-{Ba)

F-(Ba-Sr)-

maar-dlatrsme

cacholong hot brines

(Latium)

F-Pb-Zn-Ba-(U)(dolomite) unconformity-

...

(Ambian Gulf)

F-Pb Zn-Ba-(Ag)

lop",

' -So residual-calcreles

MVT

pl.",,,

~

(CaVfJ.fn rock)

F·ln silicate non-sulfidic

F-Sr -Ba MVT

karsl-caJamine

(Coahuila)

rmated F-(Cu-Fe )

IOGG ( Vergenoeg)

(Ba)-F

maar-diatreme

(F)-P high energy

topaz -(emeraldmagnesite) shear veins (Ouro Preto)

Sa-F-(Sr)cach%ng

Ba-(Po-Zn) residual-karst

hot brines

So withente veins

(F)- P /ow6nergy

(F)'P IfISlJ/ar and ramptypo

Ba-FeS-Pb-Zn-Cu Ba-F..f>b-Zn-(Ag)

Ag

MVT

SEDEXISMS ( Rsmmelsberg)

(C8vs-in rock)

..

--

(Ba)-Pb-Zn-(As)

sII/cIdastJc-chen

sulfidic (SilesIan)

Pb-Zn-Ba-F-(U)-

Sa-Sr pI.ys

calcite unconformity-

M'' ""

Mn-Pb-Ba

(Imini)

(Imini)

(Tunisian)

Ba-FaS

.-..

Sr-Pb-Zn-Ba-F-

S,

(U)-calclle

residual-karst

Sr-B-(As-U) ploy. (Kirks)

(Sr)-Ba-F-

Sf-Ba

cacholOng

~r;J

S,

(IriSh)

Pb-Zn-(Sr-Ba)

Ba-Fe

unc:~o:;/Y-

Ba-Pb-Zn-Ag

(Eyre)

Mn-Pb-8a

(6 ) limniccoal

{Sr)--Ba-F-Pb-Zn(Ag)

...vr

Sr-Pb-Zn-(Ba) (Tunisian)

S, reef limestones (Cyprus)


S lignite

S oil gas

-

32


A

B

C

E

0

34d sylvite halite

Na K CI

SODIUM

Sr

POTASSIUM BROMINE CHLORINE

35a

35b 35c

N

NITROGEN IODINE

36

Sodium carbonate Sodium sulfate

37a

I soda ash sodium sulfate

37b

trona chen zeolite

37c

witherite (Ba)

sellaite-fluorile (F) nalrocarbonalite (Oldoinyo Langai)

(apatite) phosphates lazulile-scorzalile s.s.s brazilianite

p

eophosphorite

PHOSPHOROUS Apatite Fe-Mn-AI-Mg

P-apatite U-Cs(Rb-Nb-Tal brazilianite phosphate pegmatite

38a

phosphate

phosphophyllile varisei!s ivory

P-Fe-(Cu-Au-REE) (Kiruna)

38b

bones

P-Fe-TI-Nb-ZrREE-(phlogopile) (Kola)

P-Fe-REE-U ($okl!)

IOCa?

38c

38e 381 38g zircon (hyacinth)

zircon

rock crystal

quartz

citrine amethysl rose quartz chalcedony agale opal jasper obsidian onyx carnelian morion zebra rock tektite silicified wood

gravel conglomerate sand sandstone diatomile tripo/ite silex pumice scona perlile puzzolana dimension stones aggregates

Zr Hf Si

ZIRCONIUM HAFNIUM Zircon SILICA (SILICIUM) Quarb: varieties Glass

39

(zircon-sapphire) alkalibasalt

40a

agate ctJalcedony

40b 40c

ssorolite

amethysl roc k cryslal cUrine jasper hypogene amygdaloidal basaltic tuff and lava

Zr-Th-REE-P-NbTa-F-(Be) carbonalites anile

agate chalcedony opal thunder eggs

opal

opal opal geyserite

HS


•

-


F

G

I

H

K

J

33

L

M

K-Na-Mg..Br brines-camaliltekiesente-kainilebi&choffte sutfacesnd

Na-Br-(K)

N

Sr1JYPSum

..-.

seawater

subterranean san deposits Na-Mg-(Br)-

""""elEi s.::::::

Na-Mg-(K-Cs-RbU-S)

Dlava ·lake N-j.(Cr-B-CI)

N-t-(er-B-CI) playa

csJichs (Atacama)

I brines from oil and gas deposits

Ihermonacrite

,rona

saicret6S

a/kalina lakes

Na sulfatecarbonate brines

natn:m-Mg (hydro)

magnesite mirabilite Ihenardite

p/Bya lakes trona

Istralified lakes (GreenrlvsrJ

zeolile ch6rt (Lake Magadi)

apatite

P-Nb-Ti-REE

apatite

p -( e)

"'""

laterite bauxite

pia""

perennial lake

apatite remobilized

P.(U-At.S)

P-(U)

APS phoscret8s

aCfMI et1cOS8S

P apicretes

Iazu!ile metaquaffZtte

~I'" veins

P

""'-'Y

P-(N-C} (bat)guano

ivory bones

_ P

...

granulite marbleskam (7)

P

P

P

high energy

low energy

1nsuIar~ramp-

(P)-Fe Minette-(bone beds)

P-Fe-(REE)

P. F&-(REf) Rio negro

(P)-Fe pebble Fe ore

Zt (zircon I hyacinth) placer

quartz ( Pb·Cu-Zndolomite-calcite-

quartz

rock quartz

silcrotes silicified

amethyst agate

wood

,,,,,,.rock

placer

dolomite-baryte)

opal rollte quartz (Au-Ag-Cu-U ) amethyst rose quartz smoky quartz citrine

tektite impact glass

carnelian

"''''

silcretes

quartz feldspar phyilosilica

,

h"

.opal-A

quartz , chalcedony

( P)

trlpo/lte l (onenstone

dtalolJllte

mine I, ate s

'"

,d

renite andslone

ron~::;8Ie Fig. 01.01a (continued).

-

34


B

A

E

0

C

quar1z

40d

feldspar

mica smoky quartz monon

rose quartz amethyst

pegmatite

40e 401

opal-(agate-zeolite-celadonile) volcanic-hosted

soona

40 9

amazonile moonstone (adularia,

feldspar dimension stones

Feldspar

pumice dacitic

basaltic

perlite

41a

;/Y J-' % / andesite

basalt

diaba

plagioclase)

/0

a

sunstone (oligoclase)

I

41b

ornamental slones

pumice /rachyticphonolilic

40h

aggregates

Ba feldspar ' unakite'

pumice rhyolitic

41c 41d

ite

quartz granite ,yolite ·pe Idlte"

alkaline gra . e

srem

Ia

lIe askite trachyte

hosira

prophyres

feldspar kaolinite

rhyolite

feldspar

saprolite

£. g leld r iea

. ,-aDIi"

41e

moon

ne

stone

'file

411

~

41 9

~ quart,

41h

' lapis lazuli" nepheline scapolile sodalile dimension stones

Feldspalhoids

feldspar

.,.,.,.

41;

"unskrtB'

42a

scapolile

pegmatite

s

stones

42b

scapolite

,kam

Zeolites

lapis lazuli skarn

43a

zeolite Be-Na pegmatite

43b 43c 43d 43e

AI nepheline syenite nepheline syenite AI nepheHne

42d zeolites (partly syn thesized)

yalte alr/e (oido/ite

_""",./zed

42c

zeolite (showcase)

;it leldspalh

syenite askile

hauyne omamental

r

alkalin

mplex

Na zeolite pegmatite

zeolite skam zeolites geysers zeolites subaerial vokanics opal (agate zeolite celadonite) volcanic-hosted

","' " ctw" zeolite (Lake Magadi)

431

zeolites-celadonite submarine volcanic rocks

439

,oolite zBOlde-subgreenschlst facI8s



F

G

H

""",

silcretes

_-

I

K

J

quartZ

-

M

rna"""

on",

opaJ-CT

,.,.,.,. qua""'"

L

35

quartZ

quartz chalCtKlOny

""opai-CT

me" opal

graphite schist ( Tasgazkan) Sa feldspar Alpine-type veins

quartz

sa ........,

feldspar phyllosilica tes

SEDEX

heavy minerals carbonates arkoses arlr;osic arenite

paragneis _ _ _

feldspar SBDro/ite

----

""""lito

0"""",,".

rna"""y)

I_

scapoille

csksllca_

m"'"

zeolite

t,ona

me"

(Lak9 Magadi)

zeolite

zeoNla

110 b0f8.t&

bo!

Qualernary Tertiary

D _

Carbonalile Phonolile

D

Teph rile (i ncl. Lim burgile)

~ Essexile 133

134


Fig. 15.05. Alkaline magmatic rocks and Nb–Ta ore minerals. a) Calciocarbonatite (sövite) with richterite, pyroxene and uraniferous pyrochlore, Oka-Quebec, Canada. b) Pyrochlore, Vishevy Gory, Ural Mts. Russia. c) Niobite, Ampangabe, Madagascar. d) Thortveitite intergrown with orthoclase, Evje-Iveland pegmatite district, Norway.

(Cretaceous), Lueshe and Bingo, DR Congo (Cretaceous ?) (Fig. 15.04). The largest and most spectacular ones from the mineralogical point of view are subjected to an in-depth treatment in the paragraphs to come, for the Lueshe and Bingo carbonatites (see Section 26). Niobium deposits in alkaline magmatic complexes in Brazil are operated at Tapira, Araxa, Salitre I and II Catalao. Precambrian quartzites and mica schist (Precambrian Araxa-Group) were intruded by late Cretaceous carbonatites and alkaline magmatic rocks provoking a complex mineral assemblage to develop with Nb ore (pandaite: Ba pyrochlore, gorceixite, magnetite, monazite, ilmenite), phosphate, Ti mineralization (anatase), barite, REE metals and vermiculite. The Oka calciocarbonatite complex (sövite) in Quebec, Canada, is associated with alnoites, ijolites, and okaites and was intruded into Precambrian metamorphics during the mid-Early Cretaceous. The Nb deposits carbonatite complex of Oka, Canadacontains several zones of Nb ore hosted by discontinuous bands of (1) forsterite–magnetite– apatite, (2) diopside–magnetite–apatite and (3) forsterite carbonatite carrying up to 5% pyrochlore (Fig. 15.05). Niobium is partitioned in four types of pyrochlore: (1) calcite carbonatite contains pyrochlore depleted in Ta, U and Th, (2) phlogopite carbonatite is depleted in Zr and Th, (3) pyrochlore hosted in magnetite and diopside carbonatite is enriched in Ta and U and depleted in Th, (4) pyrochlore in forsterite, forsterite–magnetite–apatite and diopside–magnetite–apatite carbonatite is enriched in Th and depleted in U. Major element geochemistry indicates that fractional crystallization of calcite carbonatite enriched the magma in Fe, Mn and F led to formation of ferrocarbonatites (Eby, 1973, 1975; Treiman and Essene, 1985; Gold et al., 1986). The Oka Complex originated from a cooling magma produced by partial melting of upper mantle rocks which are inferred to be metasomatized garnet lherzolites (age: mid-Barremian Stage, mid-Early Cretaceous, 124–125 million years). 13b E: Another group of Nb deposits is related to alkaline magmatic complexes and carbonatites and Nb-perovskite-dominated laccolites.

Studies of Nb deposits in the world's largest laccolitic alkaline magmatic complex at Lovozero, Russia, have revealed an extraordinary mineral association and lithology (Fig. 15.06). During the Paleozoic, alkaline magmatic rocks intruded during a multistage process garnet–biotite– gneisses: (1) nepheline syenite, (2) rhythmites of alternating urtites, foyaites and lujavrites, (3) eudialyte lujavrite which were intruded into older rocks, (4) lamprophyre dykes. Nb ore occurs in seams of varying thickness and made up of loparite overlying pyrochlore ore. The mineral assemblage includes among others loparite (Nb–Ta–REE perovskite), murmanite (alteration product of lomonosovite), lomonosovite, eudialyte, lorenzenite and pyrochlore (Chakhmouradian and Sitnikova, 1999; Chakhmouradian and Mitchell, 2002). The Sc mineral juonniite was reported from the Kovdor carbonatite, Russia (Liferovich et al., 1997). Thortveitite and scandian columbite are known from the Fen carbonatite, Norway (Åmli, 1977). 15.3. Sedimentary niobium- and tantalum deposits 13a H: The residual Nb–P–Ti laterites and bauxite on top of the Nb–Ti– phosphate deposit Tapira/Minas Gerais, Brazil are illustrated in (Fig. 15.07). Alkaline magmatic complexes underwent intensive chemical weathering under subtropical climates resulting in many residual deposits with various beds enriched in phosphate, Nb and Ti ores. 13a I: Sn–Ta–Nb placer deposits rich in Nb and Ta in Thailand and Malaysia originated from erosion of Mesozoic Sn-bearing granites which gave rise to an accumulation of cassiterite, ilmenite, monazite, zircon, tantalite, columbite, wolframite, xenotime, gold, strüverite (Ta–Nb rutile) in the coastal zone. Eluvial and alluvial placers in W Australia derived from erosion cutting into the Sn–Ta–Nb pegmatites of the Precambrian Yilgarn Block (Pilgangoora near Wodgina tantalite placer, Moolyella alluvial COLTAN-Sn placer). Further deposits are located in Nigeria and the DR Congo. Mineral assemblages include tantalite–columbite, cassiterite, beryl, lepidolite and spodumene. The supergene mineralization of the Nigerian


km +1

135

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Fig. 15.06. Niobium deposits in the laccolitic (layered) alkaline magmatic complexes at Lovozero, Russia (Sørensen, 1970).

granites has also provoked numerous placer deposits in the vicinity of these granites (Sakoma et al., 2000). The “Older Granites” source alluvial placers enriched in tantalite over columbite (e.g. Wambe Jemaa, Egbe), while the “Younger Granites” are the source of alluvial

placers enriched in cassiterite and columbite but impoverished in tantalite. The source rocks are biotite granites of the Jos–Bukuru Complex which were emplaced within Precambrian basement rocks (MacLeod, 1956). Erosion into these Jurassic ring dyke complexes fed the residual

136


Fig. 15.07. The residual Nb–Ti phosphate deposit Tapira–Minas Gerais, Brazil. The alkaline magmatic complex underwent pervasive chemical weathering under subtropical climates. After Minerios (1979), Rio de Janeiro/Belo Horizonte).

and eluvial cassiterite placer deposits. Placer deposits occur at a distance of 6 to 10 km at maximum from the source granites and attain a thickness of between 1.5 and 6.5 m with COLTAN being concentrated in the basal lag deposits. Placer deposition may be observed to occur over a length of 0.5 to 3 km along thalweg. In the DR Congo, pervasive chemical weathering under subtropical climatic conditions of pegmatitic primary resources brought about eluvial placers deposits (max. 10 m) and kaolinic pegmatites down to a depth of as much as 80 m.

Table 15.04 World tantalum production in 2005 by country and mining company. • World mine production • Proven probable resources • Mine production by country • Australia (48.9%) • Mozambique (18.8%) • Brazil (16.7%) • Canada (4.6%) • DR Congo (2.8%) • Major producer • Sons of Gwalia (Australia 64.3%) • Cabot Corp. (USA 5%) • Highland African Mining Company (5%)

1500 t metal content 43.000 t metal content Top 5: 91.8%, Top 10: 99.4%

Top 5: 77.8%. Top 10: 88.6%

Table 15.05 World niobium production in 2004 by country and mining company. • World mine production • Proven probable resources • Mine production by country • Brazil (88.0%) • Canada (10.2%) • Australia (0.6%) • Nigeria (0.5%) • Mozambique (0.4%) • Major producer (2003) • Moreira Salles (Brasilia 38.5%) • Unocal (USA 32.2%). • Anglo American (Great Britain 10%). • Cambior (Canada 4.9%). • Sequoia Minerals (Canada 4.9%)

33,900 t metal content 4.4 m t metal content Top 5: 99.7%

Top 5: 90.5%

15.4. Supply and use of niobium and tantalum Columbite and pyrochlore need to be upgraded in excess of 0.5% Nb–Ta oxide in hard rock deposits and 200 g/m3 Nb–Ta oxide in placer deposits to render mining of niobium feasible. REE and Sc may increase the value of the ore whereas P N 0.1%, Ti, Zr, and Sn (total amount should be b4 to 8%) have a negative effect on the ore assessment. It is used for stainless steel/high-grade steel, Nb carbide, high-temperature alloys and within the nuclear fuel circle (atomic reactor). There is limited substitution of Nb by other compounds due to a lesser performance of the resulting products. Ore grade control and host minerals are identical to that for tantalum. The final use of Ta slightly deviates from that of Nb in that it is also used for catalysts, electrolyte condensers, cathodes, spinning nozzle and in medical technique (implants and instruments). Substitution of Ta is carried out by Nb, Al, ceramics, Pt, Ti, Zr for some goods, but limited recycling. The world mining productions of niobium for 2004 and of tantalum for 2005 are listed in Tables 15.04 and 15.05. Scandium finds little industrial application except in nuclear technology, the production of lighting devices and laser crystal rods. Scandium is produced in China, Kazakhstan, and Russia as a byproduct during processing of various ores or recovered from tailings and residues. Update supply and use: USGS: http://minerals.usgs.gov/minerals/pubs/commodity/niobium/ http://minerals.usgs.gov/minerals/pubs/commodity/scandium/ 16. Beryllium 16.1. Beryllium chemistry and mineralogy Beryllium contents average 3 ppm Be in the earth's crust. Due to similarities in their ionic radii, Be is capable of replacing Si during magmatic processes and more widely dispersed than accumulated in the main stage of magma differentiation. Among beryllium minerals there are some minerals such as beryl and chrysoberyl/alexandrite which have become very popular also with non-mineralogists for their aesthetic value. Ordinary chrysoberyl is yellowish-green, transparent or translucent. Alexandrite is strongly pleochroic of emerald green, red and orangeyellow colors which changes under artificial light (Fig. 16.03). Beryl has


been given different names corresponding to its color: emerald (green), aquamarine (pale blue), heliodor (yellow), goshenite (transparent colorless variety), morganite (pink), rosterite (pink) and bixbite (red) (Fig. 16.03). Chromium and vanadium give emeralds their distinctive green color. Cr2O5 contents of 0.2 to 1.5% and V2O5 contents of 0.1 to 1% with 0.2 to 1.8% FeO are required for emeralds. While beryl is relatively common as it only requires Be to develop, the geological conditions needed to bring Be into contact with Cr and/or V rather seldom occur in nature and make emerald a strongly looked-for gemstone. Other Be minerals such as behoite, bertrandite, bromellite, euclase and phenakite are known to the rockhound or mining engineers extracting Be from these minerals but seldom go over a jeweler's counter (Fig. 16.03). The stability relationships of and reactions between the various Be phases have been experimentally studied by Burt (1975) and Barton (1986), in general. Late-magmatic greisen fluids interacting with beryl-bearing pegmatites and the leucogranitic host rocks were investigated by Markl and Schumacher (1997). Franz and Morteani (1984) have carried out investigations on chrysoberyl in metamorphosed pegmatites. The following reactions are to show how some of the less well-known Be minerals originate from Be minerals widely known as gemstones or how these gem-quality minerals (in italics) may brought about by hydrothermal processes (Barton, 1986). (T = 200 °C) 2Be3 Al2 Si6 O8 þ 14H2 O⇒4BeAlSiO4 OH þ Be2 SiO4 þ 7H4 SiO4 (T = 300 °C) 4BeAlSiO4 OH þ 2SiO2 ⇒Be3 Al2 Si6 O8 þ BeAl2 O4 þ 2H2 O (T = 400 °C) 2BeAl2 O4 þ Be2 SiO4 þ 3H4 SiO4 ⇒4BeAlSiO4 OH þ 4H2 O: (T N 600 °C water saturated) Be3 Al2 Si6 O8 þ 2Al2 SiO5 ⇒3BeAl2 O4 þ 8SiO2 (dry) Be3 Al2 Si6 O8 ⇒BeAl2 O4 þ Be2 SiO4 þ 5SiO2 In the 200 °C range, phenakite and bromellite hydrate to bertrandite and behoite, respectively, and pure beryl reacts to euclase + quartz + bertrandite or phenakite (Barton, 1986). 16.2. Magmatic beryllium deposits 14a D: Beryl–emerald–euclase–hambergite-bearing granite pegmatites are known for their varied mineral assemblage (Table 16.01). In magmatic lithologies, Be minerals mainly develop ore bodies in beryl-bearing granite pegmatites where beryl attains gem-quality or forms the most-valued modification of beryl emerald. Beryllium minerals also fill veins in Be-bearing granites and in their immediate surroundings also often as gem-quality emeralds. In beryl-bearing granite pegmatites, replacement of primary or pegmatitic beryl ends up in the formation of secondary beryllium minerals. At high temperature (∼ 550 °C), gem-quality aquamarine precipitates in vugs surrounded by alteration haloes of albite, muscovite, cassiterite and fluorite (Markl and Schuhmacher, 1997). Černý (1991) gave a comprehensive overview of pegmatites hosting beryl mineralization and used beryl as the marker mineral to distinguish the so-called beryl- or LCT-type pegmatites from NYT or REE pegmatites — first order subdivision (Fig. 16.01). Columbite or phosphate acts as another discriminator for a second-order subdivision of this group into two subtypes. It has to be noted, that the classification schemes are still hotly debated and therefore in this paper, I acted with reserve as to the use of these pegmatite classification schemes. Several deposits with Be in Canada, Australia, Malawi and USA are

137

Table 16.01 Classification scheme of beryllium deposits. (1) Magmatic beryllium deposits (1) Beryl–emerald–euclase–hambergite-bearing granite pegmatites (14a D) (2) Taaffeite- and emerald-bearing skarns (14c D) (3) Replacement deposits in volcaniclastic deposits and granites (14b D) (4) Be-bearing alkaline intrusive rocks (nepheline syenite) (14a E) (5) Tugtupite-bearing alkaline intrusions (14b E) (6) Chrysoberyl within rare element pegmatites (14c E) (7) Beryl-bearing pegmatoids (14d D) (2) Structure-related beryllium deposits (1) Aquamarine veins (14a G) (3) Sedimentary beryllium deposits (1) Regolith-hosted emerald deposits (gemstone) (14a H) (2) Alluvial–fluvial chrysoberyl placer (14a I) (3) Black shale-hosted emerald deposits (gemstone) (14a J) (4) Metamorphic beryllium deposits (1) Schist-related emerald deposits with or without pegmatitic mobilizates (gemstone) (14a AB) (2) Chrysoberyl in pegmatitic mobilizates in the contact zone (1) Metaultrabasic rocks (14b A) (2) Metapelites (14b J)

encountered in zoned pegmatites with quartz-rich cores and beryl sized up to several meters in length, occurring along the margins. In unzoned pegmatites, beryl may be present but normally at a grain size too fine-grained and an amount too small to prevent beryl from being recovered economically. Beryl occurs in patches of greisen, in quartz veins and pegmatites, and as a fine-grained dissemination at the edge of zone of muscovite-rich granite (Burke et al. 1964). The greisen contains a small deposit of estimated 50 tons of beryl in the Rosses granite, NW Ireland. On the Seward Peninsula, Alaska–USA, beryllium minerals are present in veinlets and as replacements bodies. Beryl is also known from the Alto Ligonha pegmatite, Mozambique (Hutchinson and Claus, 1956) and many pegmatites in neighboring Zambia and Malawi (Holt, 1962). Beryl — aquamarine mineralization in the Mzimba district, Malawi, are obviously unrelated to granites and supposed to have been derived from metamorphic devolatilization and partial melting of basement rocks. These pegmatites are transitional between granitic metapegmatites and pegmatitic mobilizates (Dill, 2007b). Grew (1981) investigated such Be-bearing pegmatites in granulite-facies rocks and published P–T conditions at the time of pegmatite emplacement at 800 to 900 °C and 7 to 8 kbar. Similar beryl occurrences in Canada are associated with either Archaean pegmatites or with Mesozoic/Cenozoic granitoids of S-type peraluminous type (Tomascak et al., 1994). Beryl in pegmatites decomposes in course of hydration to euclase (see previous section) which may come close to gem quality as in Zimbabwe and Minas Gerais (Brazil) (O'Donoghue, 2006). In some cases Be and B two granitophile elements being enriched during the latest stages of granite intrusions combine to form hambergite as in the granite pegmatites at Anjanabanoana, Madagascar, and the Ganesh Himal area, Nepal. Phenakite is widely known from granite pegmatites at Minas Gerais, Brazil, Amelia, USA and sites near Ekaterinburg, Russia. Rhodizite is a Cs-bearing Be oxide which was found within the alkaline late stage parts of the LCT granite pegmatites at Antandrokomby and Antsongombato, Madagascar. Madagascar is the fourth largest island in the world with a geology as old as 3.5Ga and hundreds of producing mines in pegmatites, mainly in beryl-(columbite)-type or LCT pegmatites, mined for gemmy and industrial beryl and a few operating mines in NYT pegmatite deposits. (Behier, 1960; Martelat et al., 1997; Pezzotta, 2001). A subtype of the LCT pegmatite clan has been found near Lake Alaotra by Pezzotta (2001) to contain chrysoberyl gemstones. Unlike the contact pegmatites, dealt with below, mega crystals of chrysoberyl

138


Fig. 16.01. Element zonation in complex pegmatites hosting Be–Li–Cs–Nb–Ta mineralization (modified after Cerný (1991)). a) Regional zonation: schematic representation of proximal and distal facies around cogenetic granite and pegmatite with the most characteristic element on display. b) Local zonation: cross section through a complex pegmatite body with typical mineral assemblages in zonal arrangement from the core to the margin.

may be encountered as floater in miarolitic druses within the pegmatite, locally associated with tourmaline and K feldspar. According to Lewis et al. (2004) beryl deposits are bound to highlyfractionated granites (leucocratic muscovite-bearing quartz-rich granites) enriched in W, Sn, Mo and U. Fluorine used to transport beryllium in the fluid or vapor and shows up as (1) fluorite in quartz veins, (2) fluorine-enriched granites and greisens and (3) topaz mineralization. Pegmatites and aplites indicate the presence of a fluid phase. Beryl in veins tend to have a greater gem potential than beryl that stays in the melt and crystallizes within granites or in pegmatites. The criteria for the evolution of the most valuable modification of beryl, the emerald in granitic pegmatites have been compiled by Grundmann and Giuliani (2002), Schwarz et al. (2002) and Groat et al. (2005) who focused especially on the Canadian deposits. The Brazilian emeralds such as at Comisa and Capoeirana were investigated by Delaney (1996) and interpreted to have developed from the late pegmatitic down to the hydrothermal stage based on fluid inclusion analyses. (Fig. 16.03). Be- and F-enriched magmatic fluids interacted with the host rocks to scavenge Cr and V. Mafic and ultramafic schists need not necessarily be involved as might be expected by the trace elements listed (see previous sections). Emerald crystallized in vugs of alkali feldspar granites accompanied by quartz, blue topaz and aquamarine. 14d D: Beryl-bearing pegmatites rarely occur as pegmatitic mobilizates (“pegmatoids”–abyssal pegmatites which bridge the gap between metamorphic rocks/migmatites and magmatic rocks; some may also be connected with deep-seated granitic bodies) evolving along

cleavage planes of crystalline schists as at Sierra de Ancasti/Province Catamarca, Sierra Velazco/Province La Rioja, Sierra de Cordoba/Province Cordoba and Sierra de San Luis/Province San Luis, Argentine. Compared with the afore-mentioned pegmatites from Africa and North America, these pegmatites in the Andes are poor in rare-element minerals and beryl, although very widespread among the rock-forming minerals, and only attains industrial quality with no precious beryl or emerald close by. Sometimes any distinction between true magmatic and metamorphogenic Be deposits may be difficult. 14c D: Taaffeite-and emerald-bearing skarn deposits do not play an important part in terms of the recovery of Be, excluding the rare occurrences of taaffeite skarns evolving from dolostones and limestones at the contact of Be-bearing granites (O'Donoghue, 2006). Deposits are known from Sri Lanka, China and Mt. Painter, southern Australia. Vanadium-rich emerald (200–610 °C) at the Lened occurrence, Northwest Territories, Canada, is a rare example of emeralds in skarns and hosted within a fractured garnet–diopside skarn within the Cambro–Ordovician rocks. It resulted from contact metamorphism related to the emplacement of the adjacent 93 Ma Lened pluton of the Selwyn Plutonic Suite (Marshall et al., 2004). 14b D: From replacement deposits in volcaniclastic deposits and granites, no Be minerals of gem-quality have so far been reported; bertrandite developed by epithermal alteration of calcareous volcaniclastic rocks, where it is associated with adularia, fluorite, and smectite the best example of which is at Spor Mountain, Utah–USA (Petkof, 1980). 16,000 tons of beryllium are contained in this deposit. This Be amount is sufficient to last more than 100 years at current production levels.


139

Fig. 16.02. Beryllium deposit Spor Mountain—Utah, USA. a) Cross section through the host rock unit of the Spor Mountain deposit with the open pit design (after Soja and Sabin (1986)). b) Beryllium ore with fluorite containing 1% BeO from the nodular beds of the Pliocene tuffs.

Beryllium minerals are hosted by Mio-Pliocene volcanic rocks, tuffs, breccias, rhyolite and rhyodacite intrusions into Paleozoic sedimentary (Fig. 16.02). Fluorite has been discharged from chloride- and fluoridebearing fluids at 50 to 100 °C. Beryllium-bearing tuffs are also recorded from Mongolia (Kovalenko and Yarmolyuk, 1995). 14a E: Be-bearing alkaline intrusive rocks (nepheline syenite) are a target area when exploring for rare metal deposits as demonstrated by the Thor Lake zone in the NW Territories, Canada. The Thor Lake rare metals deposit was discovered in a peralkaline intrusion. Syenites, granites and their associated pegmatitic phases underwent strong

hydrothermal alteration. Elements of economic interest in this deposit are Be, Y, REE , Zr, Ta and Nb. The predominant beryllium is concentrated in phenakite with an average BeO content of 24%, double the amount of beryllium of the more common beryllium mineral beryl. Yttrium is found in xenotime that is enriched in HREE, Eu, Tb and Dy and in bastnaesite that is enriched in LREE La, Ce and Nd. Other REE-bearing mineral phases include synchysite, parasite, allanite and monazite. Tantalum and niobium are found in ferrocolumbite and fergusonite. 14b E: Tugtupite is a pinkish red mineral attaining gem quality and one of the rare Be mineral being recovered from alkaline intrusive

140


rocks (Jensen and Petersen, 1982). Its locus typicus is located in Greenland in sodalite syenites and syenites of the peralkaline Ilímaussaq intrusion discussed elsewhere. Aegirite, pyrochlore and neptunite join tugtupite in late hydrothermal veins within these alkaline magmatic rocks. 14c E: Chrysoberyl forms along with beryl in some complex rareelement pegmatites and in veins (Walton, 2004). 16.3. Structure-bound beryllium deposits 14a G: Aquamarine is common to quartz veins that fill sigmoidal tension gashes and cut a syenite of Mississippian age which was

emplaced within an extensional setting into undeformed Paleozoic sediments of the Cassiar Platform and felsic volcanic rocks of the Pelly Mountain Volcanic Belt (Turner et al., 2007). Accessory minerals in the veins include siderite, ankerite, allanite-(Ce), fluorite, and minor albite, sulfides, and Fe–Ti–Nb oxides. The fluorite that coprecipitated with beryl from several veins has been dated using Sm–Nd geochronology at 171.4 ± 4.8 Ma. Fluid temperatures were between ∼ 275 and ∼ 400 °C. Evidence gathered in this study points to a metamorphic origin for the mineralizing fluids and a local derivation of vein constituents, which distinguish the fluids at True Blue from other intrusion- related beryl-forming fluids in the northern Cordillera.

Fig. 16.03. Gem-quality beryllium mineralization and beryl minerals. a) Emerald mineralization, Comisa, Brazil. b) Emerald mineralization, Copoeirano, Brazil. c) Emerald mineralization, Habachtal, Austria. d) Emerald mineralization, Swat Mine, Pakistan. e) Emerald mineralization, Vietnam (proprietary). f) Emerald mineralization Vietnam (proprietary). g) Aquamarine crystal growing on schoerl, Engoro, Namibia. h) Euclase, Don Bosco deposit/ Minas Geraes, Brazil. i) Alexandrite triplet, Port Viktoria, Zimbabwe.


141

Fig. 16.03 (continued).

16.4. Sedimentary beryllium deposits 14a H: Regolitization under (sub)tropical climatic conditions penetrating deep into the crystalline basement may also hit metamorphic and magmatic rocks hosting emeralds as it is the case on the Brazilian Shield at Itabira (Minas Gerais). Chemical weathering is a “natural dressing plant” which liberates the emeralds from the thrash or gangue minerals by doing neither harm to the edges nor to the faces of the crystals. They may easily be collected by the miners from the kaolinitic weathering loam, where they occur as floaters.

14a I: Beryl rarely occurs in placer deposits. Beryl is hard (hardness = 7.5–8) but its specific gravity is similar to quartz (2.63– 2.68), and thus it is not recovered from heavy mineral concentrates. Emerald is commonly full of inclusions and does not survive weathering and long-distance transport. Semi-precious gem beryl and non-gem beryl are more resistant to weathering and may occur as larger, resistant minerals within alluvial and fluvial sediments. Chrysoberyl as hard as beryl has a higher specific gravity of between 3.68 and 3.78, and hence is frequently recovered from alluvial–fluvial placer deposits, e.g., in Sri Lanka and in Brazil (Walton, 2004).

Fig. 16.04. Pegmatite-related schist-hosted emerald deposits in Zambia. Source: T. Häger.

142


Table 16.02 Host rocks of emerald deposits (Guiliani et al. 2002). 1. Pegmatites without schist (e.g. Nigeria) ➢ Be- and F-enriched magmatic fluids ➢ Interaction with the host rocks to scavenge Cr and V ➢ Mafic and ultramafic schists are not involved ➢ Emerald in vugs of alkali feldspar granites ➢ Associated minerals: quartz, blue topaz, aquamarine 2. Pegmatites and greisens with a phlogopite schist (e.g. Russia, Madagascar, Zimbabwe, Brazil, Mozambique, W-Australia) ➢ Phlogopite schists at the contact zone between ultramafic rocks and rare element pegmatites ➢ Formation of dark and black schists with emerald ➢ Associated minerals: chrysoberyl 3. Schists without pegmatites (e.g. Austria, Pakistan, Brazil) ➢ Phlogopite schists at the contact zone between ultramafic rocks and rare element pegmatites regionally metamorphosed ➢ Schistose, folded and lensoid host rocks ➢ Carbonate-talc-schists and quartz lenses ➢ Emeralds are found in melanges of blue, green schist and ophiolites ➢ Interaction of ophiolites with continental anatectic fluids ➢ Stratabound mineralization of emerald in phlogopite schists and carbonate talc schists ➢ Infiltration of hydrothermal fluids during tectonic events 4. Black shales with veins and breccias (e.g. COLOMBIA) ➢ Emeralds are found in vugs with carbonates, pyrite and albite ➢ Brines react with organic matter in black shales, releasing Be, Cr and V

14a J: Emerald veins of the so-called Colombia-type intersect black pyritiferous and carbonaceous shale and slate which are associated with albitite (metasomatized black shale horizons) and tectonic breccias (“cenicero”) (Cheilletz et al., 1994; Ottaway et al., 1994; Cheilletz and Giuliani, 1996; Giuliani et al., 2000). The genetic model is described as follows: Compressional tectonics resulted in the formation of decollements that are infiltrated by alkaline fluids causing albitization, carbonatization of the shales and provoking mobilization of Be, Al, Si, Cr,

V and REE. These alkaline fluids are supposed to have originated from evaporitic layers. A drop in fluid alkalinity and pressure release were responsible for the emeralds to precipitate. Another factor controlling emerald precipitation is the presence of organic matter. 16.5. Metamorphic beryllium deposits 14a AB: Suture zone-related or pegmatite-related schist-hosted emerald deposits are known from Austria, South Africa, Pakistan, Zambia, Brazil and Madagascar (Grundmann and Morteani, 1989; Kazmi et al., 1989; Giuliani et al., 1990; Seifert et al., 2004) (Fig. 16.03). They are often closely related to the emerald deposits mentioned in the previous section on emerald deposits in magmatic rocks. A detailed account on the origin of the emerald deposits in the Kafubu Area, Zambia has been given by Seifert et al. (2004). They identified highly magnesian talc-chlorite ± actinolite ± magnetite metabasites, hosting emerald mineralization as metamorphosed komatiites (Fig. 16.04). The estimated equilibration temperatures during regional metamorphism gave a narrow interval of 590–630 °C, assuming a pressure of 400–600 MPa. Emerald-bearing phlogopite schists are confined to the contacts of quartz–tourmaline veins and quartz–feldspar pegmatites with magnesian metabasites. Pegmatites belong to the LCT family with common Be–Nb–Ta–F enrichment. The highest homogenization, i.e. minimum trapping conditions at 400–450 MPa and 360–390 °C, was recorded in the quartz of several samples. The majority of the quartz– tourmaline± emerald veins was deposited at lower pressures around 200–400 MPa. All these emerald deposits are confined to the contact zone between granitic to pegmatitic rocks and metamorphosed mafic to ultramafic rocks (epidote–chlorite–actinolite schists) and biotite– phlogopite schists. The emerald deposits of the Mingora emerald mines in the Swat Valley, northwestern Pakistan were studied by Arif et al. (1996). These deposits occur in talc-magnesite and quartz– magnesite assemblages, associated with quartz, fuchsite and tourmaline

Fig. 16.05. Model to depict the formation of emerald deposits at the contact between pegmatite and ultrabasic magmatic rocks. Modified after Walton (2004).


(Fig. 16.03). The isotopic studies suggest that the mineralization was probably caused by modified hydrothermal solutions derived from an S-type granitic magma. It is proposed that the Swat magnesites formed due to the carbonation of previously serpentinized ultramafic rocks by a CO2-bearing fluid of metamorphic origin. Emeralds, e.g. in the Habachtal, Austria, are found in mélanges of blue, green schists, carbonate–talc–schists and quartz lenses (Table 16.02). The ore mineralogy is more variegated than in the Columbian-type deposits bearing emerald, aquamarine, morganite, ±chrysoberyl and industrial grade beryl. In addition to these gemstones in the contact zone, spodumene (kunzite) may occur in related pegmatites—see (Fig. 17.04). These emerald deposits require the interaction of pegmatites and Bebearing fluids with Cr-bearing mafic/ultramafic rocks. The gemstones in the schists formed by syn- or post-tectonic regional metasomatism between felsic rocks and pre-metamorphic pegmatoids, with the adjacent Cr-bearing rocks such as schists, gneisses or serpentinites (Fig. 16.05) (Walton, 2004), or in other words the interaction of ophiolites with continental anatectic fluids. The Kianjavato/Mananjary emerald deposits, along the East coast of Madagascar evolved during such metasomatic processes at the contact between pegmatites and hornblendites during the Pan-African tectonometamorphic event (Vapnik et al., 2006). In the nearby emerald deposits, Ianapera which is of similar age, shearing was important as a mechanism to facilitate introduction of CO2rich fluid. At Irondro, Madagascar, emerald formed within the pegmatites that are surrounded by gneisses enriched in phlogopite and metabasic rocks. Even if it might be called an intra-pegmatitic emerald deposits, it is the interrelationship of meta(ultra)basic magmatic rocks and pegmatites that brought about this gemstone deposit. Emerald and Cr spodumene in quartz veins and open cavities occupy brittle tensional fractures in folded metamorphic rocks in the Rist Emerald mine, Hiddenite, North Carolina–USA (Wise and Anderson, 2006). The occurrence of Be-, Li-, Ti-, and B-bearing minerals coupled with late precipitation of pyrite, chabazite-Ca and graphite suggests that low temperatures and reducing conditions were prevalent during the waning stages of mineral precipitation. They are unrelated to nearby pegmatites. 14b A: Chrysoberyl is observed in metasomatic contact zones between Be-bearing granitic pegmatites and host ultramafics. These contact zones contain glimmerites which represent pervasive potassic alteration (Walton, 2004). Chrysoberyl may be associated with emerald and phenakite in the metasomatic contact zones. 14b J: Chrysoberyl may form in course of regional metamorphism when Be-rich pegmatites and silica-poor metapelites get in contact with each other. They are interpreted as resulting from a postpegmatitic metamorphic event at high pressure and temperature (Walton, 2004). Franz and Morteani (1984) assumed deformation and metasomatism to be crucial for the formation of alexandrite and emerald. It is concluded that the formation of Al-rich minerals like chrysoberyl in pegmatites is due to a post-pegmatitic event at high P–T conditions. The hypothesis of a desilification of a pegmatite which intruded into SiO2poor country rocks, or of the assimilation of Al2O3-rich country rocks, does not plausibly explain this mineral assemblages in pegmatites. 16.6. Supply and use of beryllium The major hosts of Be are beryl (14% BeO), chrysoberyl, Be-bearing borate (rhodizite (22% BeO)), hambergite (53% BeO), berborite (44% BeO), Be-bearing phosphate (15% BeO). At an ore grade of 0.2 to 3% BeO, Be mineralization become economic, provided there exist reasonable resources. The major field of use is an indirect use of some of its minerals in gemology and jewelry (emerald, aquamarine, heliodor, goshenite, morganite, bixbite…). Beryllium metal finds application in aerospace and defense due to its stiffness, light weight and dimensional stability over a wide temperature range (Petkof, 1985). Beryllium–copper alloys are looked for owing to their electrical and thermal conductivity, high strength and

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hardness, good corrosion, fatigue resistance and nonmagnetic properties. Beryllium oxide is a good heat conductor, has high strength and hardness and performs well as an electrical insulator. Graphite, steel, and titanium may substitute for beryllium metal, phosphor bronze for beryllium–copper alloys and aluminum nitride for beryllium oxide. It has to be noted that substitutions can result in substantial loss in performance in some applications. Update supply and use: USGS: http://minerals.usgs.gov/minerals/pubs/commodity/beryllium/ 17. Cesium, lithium and rubidium 17.1. Chemistry and mineralogy of cesium and lithium The average grade of these alkaline elements in the earth's crust stands at 20 ppm Li and 2 ppm Cs, respectively. For Li the mean values of various lithologies and aquatic regimes are given as follows: ultrabasic rocks 0.x ppm Li, basic rocks 17 ppm Li, intermediate rocks 20 ppm Li, syenite 28 ppm Li, granite 40 ppm Li, shales 66 ppm Li, sandstone15 ppm Li, carbonate rocks 5 ppm Li, river water 3 ppm Li, seawater 180 ppb Li; cesium is not at variance to Li. Lithium and cesium are both enriched in the course of fractional crystallization and concentrated in late highly differentiated granites and their pegmatitic derivatives. Lithium strongly differs from other alkaline elements by its ionic radius which does not allow substitution for most other elements in the crystal lattice of rockforming minerals and thus Li becomes enriched in the residue of highly differentiated granites as well as brines. Limited substitution for other elements is only feasible in some micas and sheet silicates of the smectite-group (see next section). Volcanic and early diagenetic processes under acidic conditions play a decisive role in the concentration of Li and the evolution of deposits of these light metals. The major Li minerals are spodumene widespread in lithium-bearing granitic pegmatites, lepidolite and zinnwaldite two Li micas which are common in cassiterite and topaz-bearing pegmatites. Further Li minerals important for the recovery of Li are petalite and amblygonite that develop in granitic pegmatites. There is one Cs mineral that deserves mentioning as ore mineral, pollucite (Table 17.01). 17.2. Magmatic cesium-, lithium and rubidium deposits 15a D: Pegmatites are important sources of Li and Cs (Bessemer City, USA, Greenbushes, Australia, Bikita, Zimbabwe). Normally these pegmatitic deposits are exploited for Nb and Ta and the light metals Li and Cs won as byproducts (Bernic Lake, Canada Manono–Kitotolo, DR Congo). Spodumene, Li mica and pollucite are extracted from these deposits to produce Cs and Li concentrates. The Li–Cs–(Rb) pegmatite, Bikita, Zimbabwe, is besides Bernic Lake (Lac-du-Bonnet), Canada, the only site where another alkaline element rubidium was found at such a high level so as to make its recovery from the ore feasible (Dixon, 1979; Anderson et al., 1998) (Figs. 17.01, 17.02). A syncline made up of Archaic ultrabasic and intermediate metavolcanic rocks in the Victoria Schist Belt gave host to the above Li-Cs-(Rb) deposit. The Greenbushes area, W Australia forms part of the Archaic Yilgarn Block and is one of the most important producers of Li,

Table 17.01 Classification scheme of lithium and cesium deposits. 1) Magmatic lithium and cesium deposits 1) Pegmatitic Li (including gem spodume) and Cs deposits (15a D) 2) Li–Cs–Rb in rhyolitic tuffs with Be and F (15b D) 2) Sedimentary lithium and cesium deposits 1) Brine deposits and salars (15a L) 1) Geothermal waters and oil-field formation waters 2) Li brines within playas in Chile 2) Clay deposits (15a J) 1) Hectorite in altered volcaniclastic rocks related to hot-spring activity

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Fig. 17.01. The Li–Cs–(Rb) pegmatite, Bikita, Zimbabwe in plan view and in a sequence of cross sections supplemented with a list of its major ore minerals (Dixon, 1979, Kippenberger et al., 1988).

Fig. 17.02. Lithium–cesium ore of the Li–Cs–(Rb) pegmatite Bikita, Zimbabwe. a) Li ore composed of bikitaite. b) Li ore composed of lepidolite. c) Li–Cs ore composed of lepidolite and pollucite.


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Fig. 17.03. Cross section through the lithium pegmatite Greenbushes, W Australia, with ore grades of Li2O (Kippenberger et al., 1988).

recovered from spodumene with tantalite, cassiterite and niobium ore mined as a byproduct (Fig. 17.03) (Hatcher and Clynick, 1990; Partington et al., 1995). Feldspar and quartz pegmatites of the Hagendorf–Pleystein pegmatite province were also mined for some years for Li which is accommodated in a wide range of Li phosphates (Dill et al., 2006b). The Hagendorf–South pegmatite, shows a welldeveloped zoning with an aplitic margin, a pegmatitic inner zone, a quartz core and a cone-shaped body with Li phosphates, Nb, Ta, U and numerous other rare element minerals (Fig. 17.05). Between 1960 and 1972 1000 tons of Li ore have been extracted mainly from triphylite as a byproduct of the then running feldspar exploitation (Bayerische Staatsministerium für Wirtschaft und Verkehr, 1978). This chain of Variscan Li-bearing pegmatites may be extended

through the Czech Republic into Poland with lepidolite pegmatite near Rožná and Dobrá Voda (Černý et al., 1995; Novák and Cerný, 2001). Gem quality kunzite, the rose colored spodumene variety formed together with cleavelandite, black tourmaline, morganite and aquamarine in many Brazilian pegmatites where the world's largest kunzite crystal weighing 37,050 carats was discovered at Conselheiro Pena (Delaney, 1996) (Fig. 17.04). Hiddenite and amblygonite are exploited at Governador Valadares. Swarms of unzoned spodumen-rich K feldspar pegmatites with montebrasite and cookeite are exploited in the Cachoeira deposit/Minas Gerais, Brazil, whose discordant and concordant of bodies were emplaced in cordierite– biotite–quartz schist without any aplitic marginal zone (Fig. 17.04) (Romeiro and Pedrosa-Soares, 2005). Spodumene-rich pegmatites

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Fig. 17.04. Varieties of spodumen. a) Pink spodumen (kunzite) at Santa Maria del Sussuahy, Minas Geraes, Brasil. b) Colorless spodumen, unknown site in Afghanistan. c) Boudinage of concordant unzoned Li-bearing K feldspar pegmatites, Cachoeira mine, Minas Gerais, Brazil. d) Fluidal orientation of lath-shaped spodumene K feldspar pegmatites, Cachoeira mine, Minas Gerais, Brazil.

were explored within amphibolites and micaschists, during the 1980s, at Koralpe, Austria, but not exploited owing to some difficulties in processing the Li ore (Göd, 1978). 15b D: Cesium, Rb, and Li are also abnormally enriched in beryllium- and fluorite-bearing tuffs, underlying rhyolite flows at Honey Comb Hills, USA (McAnulty and Levinson, 1964). The rarealkali elements are not associated with montmorillonite, but adsorbed on minerals or glass surfaces. Lithium and Cs were concentrated by hydrothermal activity and their concentrations preserved by strong aridity in this area from subsequent wash-outs.

17.3. Cesium and lithium sedimentary deposits 15a L: An ever-increasing rate of Li and Cs is produced from sedimentary Li and Cs deposits such as brine pools and salars or geothermal waters and oil-field formation waters in the USA, (Smackover–Texas, Clayton Valley Playa–Nevada, Great Salt Lake– Utah). Lithium brines are also known from playas in Chile (Salar de Atacama), Bolivia (Salar de Uyuni), Argentine and Israel (Dead Sea brines averaging 20 ppm Li) (Fig. 17.06). The brines are processed to lithium carbonate with an average grade of 40% Li 2 O. The


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Fig. 17.05. Lithium phosphate-bearing feldspar and quartz pegmatites of the Hagendorf–Pleystein pegmatite province. Main ore body at Hagendorf-Süd which formed during the waning stages of the Variscan orogeny in Central Europe. Redrawn in Dill et al. (2006).

Fig. 17.06. Lithium extraction from brines of inland playas. a) Salar stretching across the Bolivian Altiplano near Oruro. b) Salar of the Pampa De Leoncito, Argentine.

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Table 17.02 World lithium production in 2005 by country and mining company. • World mine production • Proven + probable resources • Mine production by country • Chile (45.0%) • Australia (23.9%) • Argentine (10%) • China (9.4%) • USA (6.1%) • Major producer • GEA Group AG (Germany 23.9%) • Sons of Gwalia (Australia 23.9%) • FMC Corp. (USA 10.0%) • Potash Corp. of Saskatchewan (Canada 10.0%) • Yara International ASA (Norway10.0%)

18,000 t metal content 4.1 m t metal content Top 5: 94.4%

Top 5: 77.8%. Top 10: 88.6%

concentration of Li in the brines at the well head stand at 0.135% in Chile and 0.02% Li at Silver Head (BGR, 1988). This Li content is upgraded in evapotranspiration plants. Long-distant transport and bad infrastructures are weak points in the recovery of Li from the South American salars. 15a J: An unconventional source of Li is found in clay deposits. The Li smectite hectorite is contained in altered volcaniclastic rocks and genetically related to hot-spring activity. 17.4. Cesium and lithium supply Spodumene (8.03% Li2O), lepidolite (7.7% Li2O), zinnwaldite (3.4% Li2O), petalite (4.5% Li2O) and amblygonite (7.4% Li2O) are the most looked-for minerals to recover Li from apart from “saline brines”. Concentrations exceeding 0.5% Li2O are feasible. Lithium is used for light metal alloys, as a flux in the ceramic industry and Al plants, a lubricant additive in chemical processes, for batteries, catalysts during production of rubber and in the nuclear industry and medicine. Substitution of Li by Ca, Mg and Zn is common. Its production for 2005 is shown in Table 17.02. Update supply and use: USGS: http://minerals.usgs.gov/minerals/pubs/commodity/cesium/ http://minerals.usgs.gov/minerals/pubs/commodity/lithium/ 18. Lead, zinc, germanium, indium, cadmium and silver 18.1. Chemistry and mineralogy of lead, zinc, germanium, indium, cadmium, and silver In the periodical charts of elements, lead, zinc, germanium, indium, cadmium and silver take positions far apart from each other and, not unexpectedly, behave chemically very differently. A closer look at the economic geology of these elements gives quite a different view. It reveals that these elements resemble very closely among each other with respect to the type of deposits and the mode of concentration. There is almost no Pb deposit devoid of Zn and vice versa and rarely galena is found free from any Ag or sphalerite that does not contain even traces of In, Cd or Ge. The average grade of Pb in the earth's crust is 16 ppm Pb, while Zn has a mean of 80 ppm Zn (granites: 40 ppm, basalts: 100 ppm, shales: 80 ppm). Cadmium a rather toxic element is present in the earth's crust at an average grade of 0.2 ppm Cd. Like Cd, In (average grade 0,2 ppm) and Ge (average grade 1 ppm) are recovered mainly from ZnS-bearing ores. The abundance of silver in the earth's crust is rather modest, with an average grade of 0.08 ppm Ag. This precious metal is present in a great variety of types of deposits from Au–Ag deposits to polymetallic base metal deposits at an economic grade. Pure silver deposits, however, are the exception than the rule — see section on silver deposits. In base metal deposits Ag is almost exclusively recovered from galena and to a

lesser extent from Cu sulfides enriched in Ag, although many Ag sulfides and arsenides are real eye catchers in mineralogical collections and the showcases of mineral dealers. Argentite (87.06% Ag) is often encountered in the lower part of the supergene alteration zone and proustite– pyragyrite s.s.s. (65.41%–59.75% Ag) in low-temperature vein-type deposits of the Five-element association (Ag–Bi–Co–Ni–U). Native silver is well-known for its curled aggregates that decorate many glass cabinets (Fig. 18.01). Galena as a chemical compound contains 86.6 wt.% Pb. In nature, this level of Pb is seldom reached because galena is host to a couple of minor elements such as Ag (0.01–0.3 wt.% max. 1 wt.% Ag), together with Sb (0.001–0.1 wt. Sb, max 3 wt.%) and Bi (0.001–1 wt. Bi, max. 5 wt.% Bi). These elements are not accommodated in the lattice of PbS as solitary ions but form s.s.s. in form of miargyrite or schapachite above 215 °C. The crystal habit has been found to change with temperature and the Sb contents with the cubic form seems to predominate at the lowest temperatures (Marshall and Joensuu, 1961). The Ag-bearing minerals miargyrite and matildite in galena are also of profound importance to model the genesis of Ag deposits, in that at higher temperature matildite and at lower temperature miargyrite evolve. The solubility of Ag in galena is limited by a ternary miscibility gap, which is skewed toward the high Pb portion of the system. In addition to that complex s.s.s. or intergrowth, there are numerous accessory Ag minerals that may be identified in galena under the ore microscope. Native silver may result from primary and secondary enrichment, while argentite is said to be secondary. The pyrargyrite– proustite s.s.s, polybasite, pearceite, dyscrasite and stromeyerite are considered as primary Ag sulfide inclusions. A group of minerals often observed in close intergrowth with galena is the fahlore s.s.s with tetrahedrite, tennantite, freibergite and goldfieldite. Selenium may substitute for the sulfide in galena (b18 wt.% Se) and eventually ends up as clausthalite. The Se contents in galena increase along with increasing Eh. Even Hg and Cd normally enriched in sphalerite find an uptake by galena during hydrothermal growth (Tauson et al., 2005). Antimony and lead coprecipitate from hydrothermal solutions and form a lot of chemical compounds called sulfosalts, e.g., boulangerite and bournonite. The most simple way to describe the chemical composition of Pb sulfosalts is by the equation x. PbS : y. Sb2S3 , using the end members Pb sulfide and antimony trisulfide (Ramdohr and Strunz, 1978). Of these sulfosalts, only bournonite and boulangerite attract importance as Pb ore minerals at Zacatecas, Mexico. Furthermore Pb is found in oxidizing minerals, the most common of which is cerussite (theor.: 83.5 wt.% Pb). It is the main host of Pb in the so-called calamine deposits at Brilon, Germany, Iglesiente–Monteponit, Italy, and Freihung, Germany. A common alteration product of galena, but with little preservation potential, is anglesite (theor.: 73.7 wt.% Pb) which forms by direct replacement of galena in gossans and swiftly converts into cerussite. Pyromorphite (theor.: 76.4 wt.% Pb) is isomorphous with mimetite (As) and vanadinite (V) and on account of its green tint may act as an ore guide for exploration geologists in search of Pb deposits. This is also valid for wulfenite (theor.: 60.7 wt.% Pb) which is isomorphous with stolzite (W) and abundant in gossans of sedimenthosted Pb deposits embedded in black shales and carbonaceous sandstones and limestones at Bleiberg, Austria. Sphalerite crystallizes in a cubic-hexacistetrahedral habit (theor.: 67.1 wt.% Zn). Like its companion galena, Zn sulfide bears significant amounts of Cd (0.1–0.5 wt.% Cd) which may increase in wurtzite to as much as 30 mol% CdS (Kuhlemann and Zeeh, 1995; Palero-Fernández and Martín-Izard, 2005; Di Benedetto et al., 2005). Iron is another element that take a prominent position in the ZnS crystal lattice. Ironenriched sphalerite with max. 25 wt.% Fe were given distinct names such as “marmatite” or “christophite” (e.g. Ivigut-Greenland, Trepca, Serbia). Although often present in deposits that formed under max. temperature and pressure, this element cannot be held as a geothermometer or geobarometer (Schroll, 1976). While Fe is present as FeS, chemically related Mn (max. 5 wt.% Mn) occurs as


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2 em

2 em

2 em

4cm Fig. 18.01. Non-sulfidic Pb–Zn ore (“calamine”) and associated Ag, Cd and Ge ore minerals. a) Cerrusite, Mansurabad near Yazd, Iran. b) Smithsonite (“calamine”) in Eocene limestones at Saint Eugene–Sekarna, Tunisia. c) Hemimorphite (“calamine”) in karstified limestones at Bo Ngam, Thailand. d) Curled aggregates of native silver intergrown with argentite from Kongsberg, Norway. e) Yellow–green coatings on dark sphalerite in the vein-type Pb–Zn deposit Schauinsland near Freiburg, Germany. f) Germanite stands out from the massive ore by its red cupreous color, Tsumeb, Namibia.

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alabandite. Sphalerite has variable amounts of In (b1000 ppm), Ga (b300 ppm), Ge (b300 ppm), Tl (b1000 ppm) and Hg (mainly in low T mineralization). Cu, Sn and Fe often encountered as part of mineral inclusions dissolved from the sphalerite, developing chalcopyrite, stannite, pyrrhotite and cubanite. Rather uncommon for a cubic mineral, ZnS also appears as fibrous or acicular ZnS in various high-T deposits (?) (e.g. Příbram, Czech Republic, Fenice Capanne, Italy, Porco, Bolivia) (Dill, 1979a). Collomorphous ZnS “Schalenblende” from Diepenlinchen, Germany and Gorny Slasky, Poland, also contains wurtzite at low temperature. Wurtzite is the dihexagonal–pyramidal modification (stable N1020 °C, in nature normally present only under metastable conditions) that often occurs in low-T mineralization with marcasite. It develops a sheet-like crystal structure with mixed-layer structures similar to phyllosilicates in deposits such as Carguaicollo, Bolivia (Zn–Sn veins), Balmat–N.Y., USA (high-T deposit) and Tarnowitz–Bytow, Poland (limestone-hosted Pb–Zn deposit). Zincian calamine smithsonite (b51.5 wt.% Zn), hemimorphite (b54.3 wt. Zn), and hydrozincite are observed in gossans but also in calcareous rocks undergoing contact metasomatic alteration (Aachen, Germany, Gorno, Italy, Angouran, Mediabad, Iran) (Fig. 18.01) (Gilg et al., 2003a,b). Bivalent Fe acts as a redox indicator because on redeposition of Zn carbonate under oxidizing the Fe contents used to drop. Zincian spinel-franklinite and gahnite two further oxidic Zn ore minerals (b21 wt. Zn) developed in metamorphogenic Zn deposits (e.g. Bodenmais, Germany, Franklin New Jersey, USA) (Metsger et al., 1958; Dill, 1990a; Johnson et al., 1990). They are of lesser importance than sphalerite as it is the case with willemite (strong fluorescence) an oxidic Zn minerals encountered in gossans and metamorphic deposits as at Sterrling Hill and Franklin New Jersey, USA. The same holds true for cubic-hexacistetrahedral zincite whose theoretical contents of 80.3 wt.% Zn make it the Zn mineral most abundant in Zn, but its overall presence limit its economic value. It has been described from the contact metasomatic–metamorphic deposit Sterrling Hill and Franklin New Jersey, USA (Dunn, 1979). Germanium occurs as Cu-thiogermanate minerals germanite and reniérite in the Cu deposit at Tsumeb, Namibia (see section Cu) (Fig. 18.01). In the Freiberg vein district, Germany, argyrodite (1.8 to 6.9% Ge) was found for the first time and used to extract the metal germanium (Winkler, 1886). An overview of the variation of Ge in mineral deposits was given by Höll et al. (2007). Most Ge is dispersed through silicate minerals due to the substitution of Ge4+ for the geochemically similar Si4+. Ge is unusual in that it exhibits siderophile, lithophile, chalcophile and organophile behavior in different geologic environments. There is a tendency for Ge to be enriched in silicate minerals of late magmatic differentiates (e.g., muscovite granites), rocks that crystallize in the presence of a high volatile concentration such as pegmatites and greisens and late hydrothermal fluids. At high S activities, the thiocomplex [GeS4]4− can give rise to the formation of thiogermanate minerals, e.g., argyrodite, briartite, reniérite, and germanite, which can form elevated Ge concentrations due to sorption processes in iron hydroxides and oxides. Under certain circumstances, indium occurs as native indium in greisenized ore bodies. More often it goes together with sulfur to forms minerals such as roquesite or indite. Neither mineral is of commercial value. It is the common sulfide sphalerite, where Zn is replaced by In, Cu and Fe in the sphalerite structure, according to the scheme Cu + In + Fe = 3 Zn, which provides most In. Indium-rich sphalerite is intergrown with roquesite, forming a texture interpreted as the product of roquesite exsolution. The data obtained are consistent with the existence of the pseudoternary system stannite– sphalerite–roquesite (Moura et al., 2007). The same holds true for cadmium which forms sulfides of its own like greenockite dimorphous with hawleyite and xanthochroite an amorphous form (Fig. 18.01, Table 18.01). It is common in supergene

deposits forming yellow coatings on other Cd hosts mainly sphalerite , e.g. Tsumeb, Namibia (Bortnikov et al., 1995). The Red Dog Mine in Alaska continued in 2006 to be the leading U.S. source of cadmiumbearing sphalerite concentrate, followed by the Pend Oreille Mine in Washington State (US Geological Survey). 18.2. Magmatic lead-zinc deposits 16a C: Polymetallic subvolcanic Sn–W–Ag–Zn–Pb–Cu deposits have been mined since the16 th century for Ag and still play a vital part in southern America (Fig. 18.02, Table 18.01) Ericksen and Cunningham, 1993; Cunningham et al., 1991, 1994).These deposits are characterized by veins criss-crossing altered dacitic domes. The core zone of cassiterite, bismuthinite, wolframite and arsenopyrite is surrounded by a peripheral low-T minerals assemblage of Pb-, Zn-, Ag-sulfides and sulfosalts. Both hydrothermal systems widespread in the Andes belong to a large magmatic source at depth. Emplacement of magmatic rocks, wall alteration and ore mineralization lasted for quite a long time as demonstrated by the age data for the various events: dome extrusion: 13.8 ± 0.2 Ma, sericitisation: 13.7 ± 0.1 Ma, last thermal mineralizing event: 12.1 ± 1.1 Ma, alunite veins: 8.3 –5.7 Ma. Deposits of this type are mined at El Barqueño, Mexico, (Au–Ag–Cu–Zn–Pb), in the Pachuca Real del Monte District, Mexico (Ag), Julcani, Peru, (Ag–Pb–Zn–Cu), Quiruvilca, Peru, (Ag–Au–Cu–W– Bi–Pb–Zn), Casapalca, Peru, (Pb–Cu–Zn–Ag), Korri Kollo and La Joya, Bolivia (Au–Ag–Cu–Pb–Zn) (Rye and Sawkins, 1974; Halls and Grant, 1975; Columba and Cunningham, 1993; Deen et al., 1994; Dreier, 2005; Camprubí et al., 2006). The polymetallic mineralization is also encountered in siliciclastic rocks proximal to subvolcanics (Real de Angeles, Mexico) or it occurs as replacement and skarn deposits (see succeeding section) (Pearson et al., 1988). 16a D: Some Zn–Sn–Cu veins and replacement deposits in granites, a subtype of the afore-mentioned vein-type deposits merit

Table 18.01 Classification scheme of lead and zinc deposits. (1) Magmatic lead–zinc deposits (1) Polymetallic Sn–(Ag–Pb–Zn–Cu) vein-type deposits in intermediate subvolcanics (16a C) (2) In-bearing Zn–Sn–Cu veins and replacement deposits in granites (16a D) (3) Polymetallic Ag–Zn–Pb skarn deposits (1) Skarn deposits s.s. (16d CD) (2) Skarn-vein-type deposits (mixed-type) (16e CD) (4) Volcanic massive sulfide Pb–Zn–Cu–In deposits (VMS) (1) Kuroko-type (16c D) (2) Besshi-type and Abitibi/Ural-type (16c B) (3) Cyprus-type and Atlantic-type (16d B) (4) Iberian-type/Greenstone-belt-type (16e B–16e D) (5) VMS deposits in medium to high grade metamorphic terrains of unknown type (16f BCD) (2) Structure-related lead–zinc deposits (1) Unconformity-related Pb–Zn–F–Ba deposits (16a G) (2) Thrustbound and fold-related Pb–Zn deposits (16c G) (3) Replacement deposits (sulfidic and non-sulfidic) (16b G) (3) Sedimentary lead–zinc deposits (1) SEDEX/SMS Pb–Zn–Ag–(Ba)–FeS–(Sn) deposits (16a J) (2) Sandstone-hosted Pb–(Zn) deposits (1) Sulfidic subtype (16a I) (2) Non-sulfidic subtype (16b I) (3) Carbonate-hosted Pb–Zn–(Ag) deposits (sulfidic and non-sulfidic) (1) Pb–Zn–Ba–F–(Ag) flexure- to horstbound sulfidic deposits (plus gem-quality sphalerite) (16b K) (2) Pb–Zn–(As–Ba) flexure- to horstbound sulfidic-non-sulfidic (16c K) (3) Pb–Zn–Ba–Ag–(Ni–Co) horstbound sulfidic-(non-sulfidic) (16c K) (4) Pb–Zn–Ge–Mo faciesbound sulfidic (16e K) (5) Carbonate-evaporite-associated Pb–Zn deposits (16f KL) (6) Karst-related non-sulfidic Ba–Zn–(Pb) deposits (16a–16b H) (7) Non-sulfidic Zn hypogene-(supergene) replacement deposits (16a K) (4) Ge-enriched coal seams (16a N) (4) Metamorphic lead–zinc deposits (16b J)


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Fig. 18.02. Morphological expression of the subvolcanic polymetallic Ag–Zn–Pb deposits in the Bolivian Andes. a) Subvolcanic dacitic stocks of Mina Porco, Bolivia. b) High-temperature hydrothermal fibrous ZnS (intergrowth of sphalerite and wurtzite) in subvolcanic Pb–Zn veins at Mina Porco, Bolivia. c) Dacitic stock with Pb–Zn–Sn–Ag deposits of Cerro Rico de Potosi, Bolivia.

special treatment herein for their elevated In contents. Common Pb– Zn deposits rarely see the In contents to be increased to abnormal amounts in sphalerite. Indium mineralization is closely linked to the Zn–Sn–Cu association in the base metal veins in the Erzgebirge, Germany, and interpreted to have accumulated by fluids expelled from magmas during emplacement of post-collisional lamprophyric and rhyolitic dikes. Indium concentration is especially high in Cdand Fe-rich sphalerites from Ag-base metal and Sn-polymetallic deposits (Seifert and Sandmann, 2006). In the renown In-bearing deposits at Mount Pleasant, New Brunswick–Canada, after the emplacement of the porphyry tungsten–molybdenum deposits, with negligible In contents, In-bearing veins, replacement and breccia-hosted tin-base metal deposits formed during a younger intrusional stage into both granitic rocks and associated volcanic and sedimentary rocks. Sphalerite, chalcopyrite, arsenopyrite and cassiterite are the most abundant ore minerals. Indium content correlates closely with high contents of zinc and copper, reflecting its

concentration primarily in sphalerite and, to a lesser degree, in chalcopyrite at temperatures from 400 °C to less than 200 °C. Indium and associated metals were concentrated in magmatic–hydrothermal fluids in response to cooling and dilution of the hot magmatic ore fluids, mainly as a result of mixing with meteoric fluids (Sinclair et al., 2006). 16d CD: Polymetallic Zn–Pb–Ag skarn deposits are distal representatives of the vein-type deposits within subvolcanic rocks. Skarn deposits form near magmatic rocks compositionally ranging from diorite through high-silica granite. As far as the depth of formation is concerned, they occur from deep-seated batholiths to near-surface volcanic extrusions. Zinc skarn types with garnet, pyroxene, ilvaite, pyroxenoid and amphibole are distinctive by their Mn- and Fe-rich mineralogy from other skarn types (Fig. 18.03). The pyroxene: garnet ratio and the Mn content of pyroxene systematically increase along the fluid flow path (Meinert, 1987). Skarn deposits are known at Groundhog–New Mexico, Darwin–California, USA, El Mochito,

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Fig. 18.03. Pb–Zn skarn deposits. a) Near-vertical ore beds in the oxidized part of the Pb–Zn skarn deposit of Trepca Mine, former Yugoslavia. b) Fe-enriched black sphalerite, chalcopyrite, pyrite, galena and calcite in contact metasomatic ore from Trepca Mine, former Yugoslavia. c) Replacement of preexisting acicular Mn skarn minerals by Pb–Zn sulfides at Madan Pb–Zn deposit, Bulgaria. d) Johannesite (brown) — rhodonite (red) skarn of Madan Pb–Zn deposit, Bulgaria.

Honduras, Zacatecas, Chihuahua–Fresnillo–Unidad Naica, Bismarck, Mexico, Cerro de Pasco, Uchucchacua, Peru, Kassandra District– Chalkidiki, Greece (Einaudi et al., 1981; Meinert, 1984, 1992; Grant and Ruiz, 1988; Newberry et al., 1991; Bussell et al., 1990; Gilg, 1993; Megaw et al., 1998; Ault, 2004). At Cerro de Pasco, Peru, vanadium minerals similar to those discussed for Minas Ragra, Peru, in this paper occur. They suggest mobilization of V and carbonaceous matter at depth. These deposits occur in continental settings associated with either rifting or subduction processes. A unique position is taken by the Huanzalá Pb–Zn–Ag–Cu deposit, Peru, where limestones of the Santa Formation were intruded by a quartz porphyry dike which caused the skarn formation and hydrothermal alteration on both sides of the quartz porphyry dike (Imai et al., 1985). Noteworthy is not the Pb–Zn–Cu minerals and accompanied by Ag–Sn–W minerals from Huanzalá but the pyrite Huanzalá ranks as the largest producer of pyrite specimens of showcase-quality in the world which occur in cubic, pyritohedral, and octahedral forms, with many modifications (Weibel, 1980). Precambrian Zn-rich skarn formation has been recorded from within the Gawler Craton, South Australia (Reid et al. 2009). Moderately deep distal, dike-related, Zn skarn deposits, e.g., Francisco I. Madero, Mexico, are also connected with epithermal deposits and transitional into what has been described under 16e CD (Canet et al., 2009). 16e CD: Polymetallic Ag–Zn–Pb skarn deposits and mixed-type skarn-vein deposits are grouped as a separate entity. The reason for that subdivision is put across by the Figs. 18.04, 18.05 illustrating the

structural setting of the composite vein-type and manto-like replacement deposit in calcareous rocks at Lavrion, Greece (Marinos and Petrascheck, 1956; Skarpelis, 2007; Skarpelis et al., 2008), Gyumuslug, Russia, (Evans, 1980) and Madan, Bulgaria (Bogdanov, 1979). All of them are produced by pluton-driven hydrothermal solutions migrating along a great variety of pathways from fracture zones to karst cavities and solution pipes and slowed down in their ascend by less permeable argillaceous interbeds within the carbonate host units. They are of high-temperature origin and normally formed in excess of 300 °C, rich in silver accommodated in galena, stephanite, polybasite and pyrargyrite/proustite and characterized by an extraordinary variety of minerals found in the hypogene and supergene parts of the mineralization and not to forget in the slags left over after centuries of mining and smelting (Marinos and Petrascheck, 1956; Conophagos, 1980; Bussell et al., 1990; Gelaude et al., 1996; Voudouris, 2005) (Fig. 18.06). In Lavrion (Laurion), the Greeks mined from the 7th to the 1st century BC and from 1865 to 1977 AD ore bodies at the contact of Mesozoic marble, schist and a Miocene granodioritic porphyry intrusion. Skarn-type mineralization with magnetite and pyrrhotite around the granodiorites are common, whereas calc-silicate minerals are of minor occurrence. Being part of the same metallogenic province, the base metal-bearing Fe skarn deposit at Mega Levadi on Seriphos Island, Greece, in the Aegean Sea, a granodiorite brought about andradite and hedenbergite in its close proximity (Dill et al. in press-c). The Madan Pb–Zn ore field forms part of the West Rhodope ore district in


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Fig. 18.04. Vein-type and replacement Pb–Zn ore shoots in the Madan ore field in Bulgaria (Bogdanov, 1979).

Bulgaria which is underlain by crystalline rocks of Precambrian age (Bogdanov, 1979). Favorable conditions for Pb and Zn ores existed in this region in the aftermaths of the intrusion of rhyolitic dykes during the Miocene. Almost simultaneously two types of ore bodies developed. Ore veins crosscut various lithologies of the host anticline (Fig. 18.04). Metasomatic sheet-like ore bodies with johannsenite–rhodonite skarns and quartz sulfide mineralization replaced carbonates of three marble beds. In the Pb–Zn ore, galena and sphalerite are major ore minerals accompanied by minor Cu– Fe sulfides, arsenopyrite and sulfosalts. Native silver, Ag-bearing minerals, electrum and gold are found together with rhodochrosite and manganocalcite. Skarn mineralization took place at temperatures of more than 400 °C, whereas the Pb–Zn mineralization developed around 300 °C to 210 °C. Throughout the waning stages, barite precipitated at approximately 50 °C. The ore mineralization formed at moderate depth of between 1 and 3 km 30 to 40 m.y. ago. A large amount of Pb and Zn originates from ore extracted from volcanogenic deposits: (1) Kuroko-type (16c D), (2) Besshi-type and Abitibi/Ural-type (16c B), (3) Cyprus-type and Atlantic-type (16d B) and (4) Iberian-type (16b B–16b D) deposits. The Kidd Creek Zn–Cu– Ag–Pb–Cd–Se–In-bearing volcanic-hosted massive sulphide deposit lies within the Archaean Superior Craton N of Timmins, Ontario, Canada, in the Kidd Volcanic Complex (2.710 to 2.717 Ga). The Kidd Volcanic Complex is made up of a bimodal suite of komatiitic and high silica rhyolite flows overlain by tholeiitic basalts (Walker et al., 1975). As to the volcanic host rock suite, it resembles the afore-discussed Iberian-type. The Zn–Cu VMS deposits of the Archaean Murchison greenstone belt, South Africa, represents a rifted volcanic arc and is an analogue to the Kidd Creek deposit. The Maranda J., Romotshidi and Letaba CZ mines account for an average of 190 ppm In. The copper concentrate averages 76 ppm In and the Zn concentrate 306 ppm In (Schwarz-Schampera and Herzig, 2002). These VMS-type deposits represent the largest reserves and resources of indium which are genetically linked to the highly differentiated felsic magmatic rocks. The Flin Flon Zn–Cu–Ag–Au–Pb deposits in Manitoba, Canada, are representative of a group of massive sulphide ore bodies hosted by a

Lower to Middle Proterozoic volcanogenic sequence (Bleeker, 1990). The mineralization is associated with felsic volcanics within tholeiites. The massive sulphides are emplaced within a rhyolite flows underlain by footwall basalt breccia and overlain by basalts and andesites. Many of the VMS deposits are enriched in Zn and Cu relative to Pb so that they were treated together with Cu rather than with Pb and Zn. VMS deposits in medium to high grade metamorphic terrains of unknown type are coded 16f BCD. This “pigeonhole” has to be used also for Pb and Zn deposits. 18.3. Structure-bound lead-zinc deposits 16a G: Central Europe is rich in Pb–Zn deposits which have contributed to the welfare of towns in the Harz Mining Districts (Bad Grund, Clausthal–Zellerfeld, St. Andreasberg), Rheinisches Schiefergebirge (Ramsbeck, Bensberg) Ruhr–District and the Erzgebirge (Freiberg). In the USA the Coeur d'Alene–District forms part of this type of mineralizations. Based upon the chemical and mineralogical data obtained, many deposits mined for barite fluorite and Pb and Zn may be attributed to the hypogene subtype of unconformity-related mineralization, e.g. Nabburg–Wölsendorf, Oberwolfach, Bad Grund, Germany, Hranice, Vyškov, Czech Republic, Maxonchamp , Val d' Ajol, France (Huck, 1984; Bouladon, 1989; Stedingk and Stoppel, 1993; Schneider et al., 2003; Dill et al., 2008b) (Fig. 18.07). Ascending metalliferous fluids precipitated their elements within the Variscan basement rocks and impregnated the coarse-grained porous clastic rocks of Lower Triassic age with galena, sphalerite, chalcopyrite, bravoite, pyrite, tetrahedrite, bournonite, marcasite, covellite. Based on Rb/Sr dating, a Middle Jurassic (Dogger: 170 ±4 Ma) ore age was assumed (Schneider et al., 1999). Pb isotopes of galena from sandstone-hosted deposits in the lower Triassic Buntsandstein and vein-type deposits in late Variscan greywackes and slates of the Rhenish Massif display a similar trend (Large et al., 1983; Krahn, 1988). Studies of Pb isotopes of galena from vein-type deposit of other basement uplifts yielded similar results (Höhndorf and Dill, 1986). Steep hydraulic and chemical gradients between the bedrock and the overburden are held responsible for the emplacement of hypogene ore mineralization above and below the

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unconformity by means of intrastratal solutions and/or convectivelycirculating fluids (Figs. 18.08a, 18.08b). 16c G: The so-called schistosity or cleavage-parallel veins in the Rheinisches Schiefergebirge in Germany may be attributed to Variscan thrustbound epigenetic mineralization (Walther and Dill, 1995; Wagner and Boyce, 2001) (Figs. 18.09a, 18.09b). In the Bohemian Massif the Ag-, Pb-, Zn- and Cu accumulations in the Příbram ore district of the Czech Republic were also attributed to this type of thrust-bound and foldrelated metamorphogenic deposits, although there are remarkable structural differences between the metamorphogenic veins in the Rhenohercynian and Moldanubian zones (Pouba and Ilavsky, 1986) (Fig. 18.10). Important Ag carriers besides galena are pyrargyrite, stephanite and diaphtorite. In the Freiberg mining district more than 1000 vein-type base metal mineralizations post-date the Sn-bearing mineralization. The main minerals are Pb-, Cu-, Zn sulfides and Agenriched tetrahedrite (= freibergite) and Sb sulfosalts as well as native Ag and argentite. These veins are famous for the presence of argyrodite (Baumann et al., 2000).They resulted from differentiation of magma with the development of a volatile fluid phase that escaped along faults to form these veins. 16b G: Sulphidic and non-sulphidic replacement deposits in calcareous rocks manifest that there is a continuum between

sediment-hosted deposits dealt-with in the succeeding chapter and magmatic-related skarn deposits normally discussed under magmatic ore deposits. One of the striking characteristics of the Zn–Pb–Ag–Cu– Hg deposit Rubiales–Santa Bárbara, Spain, are ore bodies of sphalerite, galena, ankerite, quartz, complex sulfides, arsenides and sulphosalts. They developed as a result of syntectonic replacements along shear zones and lithologic contacts in the time span from 307 ± 7 Ma (Rubiales) to 323 ± 7 Ma in Lower/Early Cambrian carbonates (Santa Bárbara) (Tornos et al., 1993). 18.4. Sedimentary lead–zinc deposits 16a J: Volcaniclastics- and siliciclastic hosting Pb–Zn–Ag–Cu deposits (SEDEX/SMS deposits) have been discovered in various geodynamic marine settings, in rocks undergoing very low-grade to high-grade stage regional metamorphism: Broken Hill, Mount Isa, Australia, Black Mountain (Swartberg)/Aggeneys, South Africa, New Brunswick–Bathurst, Canada, Meggen and Rammelsberg, Germany (Barnes, 1987; Sperling and Walcher, 1990; Maynard and Okita, 1991; Praekelt et al., 1997; Reid et al., 1997; Perkins, 1997; Perkins and Bell, 1998; Burton, 1998; Cartwright, 1999; Betts et al., 2003; Davis, 2004; Lentz and McCutcheon, 2006). Some are also important producers of

Fig. 18.05. The Miocene vein-type and replacement Zn–Pb deposit at Lavrion, Greece. a) idealized cross section modified after Marinos and Petrascheck (1956). b) Pb–Ag–Zn ore replacement of marble. See biro for scale (source N. Skarpelis). c) Detachment fault, which is of ore control for the Pb–Zn–Ag ore, between the footwall Upper Marble and the Blueschist unit on top. Stockwork-like subvertical veinlets are terminated by the detachment fault. Plaka, Greece. d) Fine-grained ankerite replacing the Upper Marble shows boxwork solution cavities which are lined with quartz. The mineral dissolved might be a younger-generation ankerite or ferrous dolomite of larger grain size and different morphology (“saddle dolomite” ?) near Kamariza, Greece.


Fig. 18.05 (continued).

cadmium and bismuth. Early Precambrian representatives of this group of deposits seldom allow a precise determination of the depositional environment, e.g. the metamorphosed Zn–Pb–Ag deposit Zincgruvan, Sweden, where 1.9 Ga years ago sphalerite, galena, chalcopyrite and pyrrhotite developed in volcanic siltstones, chemical sediments and limestone–dolomite skarns (Hedström et al. 1989; Sundblad, 1994, 1994; Kumpulainen et al., 1996).

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The Fe sulfide Zn(–Pb) deposit Bodenmais, Germany, in late Proterozoic rocks has a strong Ba anomaly at the edge of the mining district, which is interpreted as a metamorphosed barite rim. Even high-T metamorphism of up to 700 °C could not completely eradicate the primary zonation in this ore deposit and prevent it to be attributed to the Rammelsberg- or SHMS-type deposits (Dill, 1990a). The Devonian reference types at Meggen, Rammelsberg, some smaller massive sediment-hosted pyrite-barite deposits of the Rheinische Schiefergebirge and an newly found one near Chaudfontaine, Belgium, provide enough sedimentological data in the clastic and calcareous shelf sediments to correlate these deposits with rifting phases during the Devonian (Dejonghe, 1998 (Fig. 18.11). The pyrite- and baritebearing Zn–Pb–Cu Rammelsberg deposit shut down after more than 1000 years of mining in 1988 has become a classical example of submarine hydrothermal deposits and its name became a class term to describe stratabound deposits. The mineral association consists of sphalerite, pyrite, barite, galena, chalcopyrite, pyrrhotite, covellite, arsenopyrite and tetrahedrite (Fig. 18.12). Ore mineralization proper, followed a pre-stage of silica concentration called “Kniest”. In the post-ore stage fine-grained barite came up again. The development of stratabound base metal deposits in rift basins by convectively circulating metal-bearing fluids was depicted by Russel et al., (1981). The “Rammelsberg-type” deposits are considered to be a product of metalliferous basin-dewatering brines, conducing to SEDEX or SMS-type mineralization during the syn-rift phase of basin evolution (Maynard and Okita, 1991). The world's largest Zn– Pb–Ag deposit Broken Hill, Australia, formed in a Paleoproterozoic continental shallow water lacustrine rift environment accompanied by an invasion of high Fe–Ti tholeiitic mafic melts (Plimer, 1979, 1984, 2006). Ore fluids and contained metals were of mixed magmatic, crustal, evaporitic and possibly lacustrine origin and had a low sulfur content. Coeval multiphase high grade metamorphism and deformation has retextured, remobilized and reconcentrated parts of the Broken Hill sulfide masses that originally contained 400 Mt of sulfide rocks. The Sullivan Fe–Pb–Zn–(Sn) sulfide ore body in Canada that lies conformably near the top of the Middle Proterozoic Lower Aldridge Formation is a stratabound, massive sulfide sediment-hosted metal deposit which differs from the overall sediment-hosted massive sulfide deposits by its alteration and element content (Fig. 18.13) (Hoy, 1982; Jiang et al., 1998). The ore is composed of predominantly pyrrhotite massive sulfides, galena and sphalerite. The pyrrhotite zone also carries cassiterite, averaging around 0.4% Sn with exhalative chert, manganese and barite. Below the western apart of the main Sullivan sulfide deposit is a funnelshaped zone or tourmaline pipe rich in albite, chlorite, pyrite and carbonate which pass locally into the ore body. This rift-related ore deposit developed from sulfide-deficient brines at temperatures between 105°

Fig. 18.06. Carbonate-hosted Pb–(Zn) replacement ore. a) High-grade replacement ore with galena, pyrite and calcite from the Black Cloud Mine Leadville—Colorado, USA. b) Sphalerite, pyrite, enargite, galena and barite of the Apex Ore Body near the Sioux Ajax Fault in Paleozoic limestones of the Mammoth Mine Tintic District Utah, USA.

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Fig. 18.07. Unconformity-related vein- and sandstone-hosted Pb–Zn deposits in the Rheinisches Schiefergebirge and its overlying platform sediments. a) Galena impregnating conglomerates of the alluvial–fluvial Lower Triassic Bunter Series at Maubach, Germany. b) Pb–Zn ore with galena and marcasite from veins intersecting the Upper Carboniferous Coal measures at Christian Levin in the Ruhr District, Germany. c) Idealized cross section showing the spatial relationship between sandstone-hosted Pb–Zn deposit within the platform sediments and vein-type deposits underneath in the basement rocks.

and 200 °C under euxinic conditions in a local sink or extensional secondorder sedimentary basin on the seafloor. Earlier, deeply buried sediments devolved fluids into a deep reservoir of sandy siltstones and sandstones. Intrusion of dolerite raised the geothermal gradient and prompted overpressuring of the lower sedimentary reservoir which breached overlying sediments, forming a breccia diatreme (Fig. 18.13). There is no sharp boundary between Pb and Zn deposits bound to siliciclastic rocks and those intercalated with calcareous series. The Neoproterozoic Rosh Pinah Zn–Pb deposit, Namibia, in southwestern Namibia developed during continental rifting at ∼740 Ma within mudstones, with interbedded, dolomitic and Ba-rich carbonate lenses of the Rosh Pinah Graben (Alchin and Moore, 2005). Finely-laminated base metal ore formed in a second-order basin in anoxic sediments: Together with some further lithological features, these are indicators of a sedimentary exhalative (SEDEX) origin. The failed rift was filled mainly by a volcaniclastic rift-fill sequence. Volcanism might have provided the heat engine for contemporaneous sedimentary-exhalative and hydrothermal replacement base metal deposits, which formed within anoxic, carbonaceous mudstone horizons in the deep rift. Similar to other deposits of this type rimmed by barite a peripheral barium-rich carbonate ore body evolved at Rosh Pinah. Vertical zoning is indicated by elevated Cu contents in the footwall breccia zone, and increases in Pb and Zn towards the massive sulphide zones occurring near or at the hanging wall contact. 16a I: Sandstone-hosted sulphidic Pb–(Zn) deposits occur in Lower Cambrian siliciclastic series at Laisvall, Sweden, and in several parts along the edge of the central European basin in the so-called Germanic Triassic Bunter Series (Mechernich–Maubach, Freihung, Germany, and Largentiere, France, (Gudden, 1975; Rickard et al.,

1979; Schmid, 1981; Bjørlykke and Sangster, 1981; Dill, 1990b; Schneider et al., 1999). The sandstone-hosted Pb–Zn deposits are epigenetic or diagenetic as exemplified in the succeeding section. The Laisvall deposit has crosscutting curvilinear features like roll fronts attesting to diagenetic through epigenetic ore mineralization. Ascending metalliferous fluids did not only precipitate their elements within the Variscan basement rocks in vein-type deposits (see structure-bound unconformity-related deposits of section 20.3), but impregnated the coarse-grained porous clastic rocks, resting immediately above the Paleozoic rocks and led to the sandstonehosted Pb–Zn ore deposits at Maubach–Mechernich, Germany. The main ore minerals are galena, sphalerite, chalcopyrite, bravoite, pyrite, tetrahedrite, bournonite, marcasite and covellite. A Middle Jurassic age was assumed by Schneider et al. (1999). 16b I: The Freihung non-sulfidic deposit, however, is different in origin from the afore-described sandstone-hosted sulfidic deposits. This Pb mineralization is faciesbound to the marginal calcareoussiliciclastic beds of the Middle-Upper Triassic beds and may well be explained by a brine-mixing model under arid climatic conditions leading to a sabkha-like environment with PbCO3 which was subsequently replaced by PbS (Fig. 18.14). The isotopic composition of galena resembles isotopically average basement Pb. (Lippolt et al. 1983). An overview of Pb and Zn deposits evolving during the PermoTriassic within the Central European basin (Germanic Facies) shows one type of deposits closely linked to a sedimentary facies with Pb being concentrated near the edge of the basin in arenaceous host rocks and another one with Cu being accumulated in finer-grained clastic rocks (Figs. 18.15, 18.16). As this siliciclastic basin fill is


157

Fig. 18.08a. Cartoon illustrating the development of hypogene vein-type unconformity-related F–Ba–Pb–Zn–U deposits and fluid migration above and below this interface.

Fig. 18.08b. Ore textures of unconformity-related Pb–Zn vein-type deposits and metamorphic Pb–Zn mobilizates. a) Brecciated silicified gneiss clasts cemented by sphalerite in epi- to mesothermal veins at Roggenbach Mine-Schauinsland, Germany. b) Chalcopyrite–calcite–sphalerite massive ore from the epi- to mesothermal vein-type deposit Bad Grund, Germany. c) Calcite–sphalerite breccia ore from the epi- to mesothermal vein-type deposit Bad Grund, Germany. d) Metamorphic mobilizate of the stratabound sphalerite–calcite ore in garnet gneisses of the Ötztal Complex Schneeberg Mine, Italy.

158


Fig. 18.09a. Fold-related metamorphogenic Ag-bearing Pb–Zn vein-type deposits. a) Cross section through the Ramsbeck Pb–Zn deposit in the Rheinisches Schiefergebirge, Germany. b) Steep and flat-lying Pb–Zn veins related to the folding of the Ramsbeck Quartzite. Movement parallel the shear planes gave rise to the accommodation space for Pb–Zn mineralization (Bauer et al., 1979).

replaced by calcareous rocks, Zn becomes predominating over Pb (Fig. 18.17). (Dill, 1994a). A similar scenario has been recorded by Drovenik et al. (1980) for the U–Pb–Zn–Cu deposits that developed during the Permo-Triassic in the Slovenian part of the Alpine Mountain Range. The stratabound nonsulfide Magellan Pb deposit in the Palaeoproterozoic Yerrida Basin, Australia is a predecessor of the Freihung-type deposit.

16a–f K: Carbonate-hosted Pb–Zn–(Ag) deposits are of sulfidic and non-sulfidic type. Calcareous rocks play a significant role as host for Pb–Zn ore deposits and only the most renown out of this class of deposits may be listed here: Tri-State-District, USA, Pine Point ore field, Nanisivik, Robb Lake, Polaris, Canada, Reocin, Spain, Les Malines, France, the Upper Silesian Pb–Zn deposits, Poland, Raibl, Gorno, Sallafossa, Italy, Bleiberg-Kreuth, Lafatsch, Austria, Mežica, Topla,


159

Fig.18.09b. Ore textures of fold-related metamorphogenic and high-temperature Ag-bearing Pb–Zn vein-type deposits. a) Galena-bearing shear vein (“Bleischweif”). Stress-induced galena crystallizing (sub)parallel to the fissure walls of Braubach Rosenberg Mine, Germany. b) Pyritiferous galena–sphalerite ore with quartz in contact with slates of the wall rock (shear zone-hosted vein) at Vereinigte Bastenberg-Dörnberg Mine-Ramsbeck, Germany. c) Sphalerite–galena–freibergite vein-type mineralization of the (eb – mineral association (“Edle Braunspat” = Ag-enriched carbonate mineral association)) Brand-Erbisdorf of Freiberg mining district, Germany.

Slovenia, Olovo, Bosnia–Herzegovina, Rauschberg, Königsberg, Wiesloch, Germany, Bou Grine, Tunisia, San Vicente Peru, Navan, Lisheen, Galmoy, Ireland and Touissit–Bou Beker, Morocco, Lennard Shelf, Australia (Beales and Jackson, 1966; Brockie et al., 1968; Monseur and Pel, 1972; Sass–Gustkiewicz et al., 1982, 1999; Symons and Sangster, 1992; Disnar, 1996; Smethurst et al., 1999; Patterson and Powis, 2002). Some subtypes also contain barite and fluorite as it is the case in the English Pennines and may be refound in the section below (Dunham, 1983). Mississippi Valley-type (MVT) deposits sensu stricto are epigenetic, stratabound, carbonate-hosted Zn–Pb sulfide bodies. In places Ag, barite and fluorite may be also recovered economically. Most of the MVT deposits formed in undeformed or weakly deformed carbonate platform sediments adjacent to epicontinental sedimentary basins. Mineralization is the product of regional- or subcontinental-scale migration of warm saline aqueous solutions similar to oilfield brines that migrated through aquifers within platformcarbonate (Fig. 18.18). Calcareous host rocks are common to all of these Pb–Zn but no common model to account for the mode of ore emplacement in these calcareous rocks exist nor do they share a common element association. There are ideas invoking ore-fluid migration to compressive tectonic regimes associated with continental accretion, whereas others favor fluid migration and ore deposition in an extensional tectonic regime. As a consequence of this, these carbonate-hosted Pb–Zn deposits are treated separately under different names such as Mississippi valley-type (MVT), Silesian-type, Alpine-type and Irish-type Pb–Zn deposits. The term MVT is at times used in a wider sense independent from calcareous host rocks and used to describe a mode of ore deposition in basins with host rocks different in their lithology (siliciclastics) and the type of ore textures.

16b K: Pb–Zn–Ba–F–(Ag) flexure- to horst bound sulfidic MVT (Mississippi–Valley–Type) deposits occur in carbonate platforms and undeformed orogenic foreland rocks, where sulfides and gangue minerals occupy primary carbonate porosity provided by open-space fillings of breccias and fractures, and/or replace the host dolostone (Leach and Sangster, 1993) (Figs. 18.18, 18.19). The epigenetic, stratabound, carbonate-hosted sulfide bodies contain sphalerite and galena, in places argentiferous (10 ppm to 161 ppm Ag) in some world class deposits such as the Tri-State district of Missouri, Kansas and Oklahoma, USA. It is to the present knowledge the only type of Zn deposit which may bring about red and yellowish brown gem-quality sphalerite as demonstrated by the MVT mineralization at Áliva Mine–Picos de Europa, Spain (Gómez Fernández et al., 2000). Attribution to this type of Zn–Pb deposit in Spain to the MVT deposits is based on trace elements of galena and sphalerite. 16c K: Pb–Zn–(As–Ba) flexure- to horst bound sulfidic and nonsulfidic deposits are represented by the Polish Pb–Zn deposit. The Olkusz–Pomorzany and Trzebionka mines in the Upper Silesian ore district are among the most important producers of Zn, Pb, Ag and best studied ore deposits of the MV-type. Hosted within shallowmarine series of Middle Triassic age, ore mineralization probably continued until Middle Tertiary times (Leach et al., 1996). Uplift and karstification, fluctuation of the water table and weathering play an important part during emplacement and are also responsible for the oxidation of sulfides of the Silesian-type MVT deposits (Boni and Large, 2003). The mineralogical association consists of sphalerite, galena, pyrite, marcasite, dolomite, calcite, barite, chalcedony, quartz, Pb–As sulfides (jordanite, gratonite) (Fig. 18.20). Zinc carbonates and rare Zn silicates occur in karst pockets (Szuwarzynski, 1996; SassGustkiewicz et al., 1982; Sass-Gustkiewicz and Kwiecinska, 1999).

160


Fig. 18.10. Section across the swarm of veins within the Příbram Ag–Pb–Zn ore district (after Pouba and Ilavsky, 1985).

Cadmium, Ag, Ge, Ga, As, Tl, Sb and Ni are accommodated in the sulfides as isomorphous substitutions. Saline basinal aqueous solutions at temperatures between 75° and 200 °C (similar to oil-field brines) migrated out of sedimentary basins, through aquifers, to the basin periphery into the platform carbonate sequences (Fig. 18.21). Various models have not been put forward for the brine flow: (1) topographic or gravity-driven fluid flow model (Bethke and Marshak, 1990; Garven and Raffensperger, 1997),

(2) sedimentary and tectonic compaction model invoking to the expulsion of basinal fluids through sediment diagenesis and tectonic sediment compaction and the episodic fluid release from overpressured aquifers (Cathles and Smith, 1983; Oliver, 1986), (3) hydrothermal convection model (Morrow, 1998) involving deep convection circulation of hydrothermal brines due to buoyancy forces related to temperatures and salinity variations in long-lived flow systems that are capable of recycling subsurface solutions many times through the rock mass. Sequence stratigraphic tools may assist in correlating MVT Pb–Zn–Ba–F deposits, unconformity-related Pb–Zn– Ba–F–U vein-type deposits and supergene mineralization bound to the regolith on the sequence boundary/unconformity (Dill, 2008a). MVT Zn–Pb mineralization was emplaced in calcareous rocks of highstandsystem tracks HST (Dill, 2008a). Karstification of the calcareous host rocks of the MVT deposits was favored by an erosional surface on top of the HST equivalent to the SB. In Upper Silesia, another SB brought Paleozoic basement rocks in contact with the Middle Triassic orebearing dolomite and, in Wiesloch, Germany, the Rhein–Graben boundary fault was responsible for Muschelkalk rocks to be thrown into a position juxtaposed to Tertiary rocks of the graben fill (SassGustkiewicz and Kwiecinska, 1999). Convective and short-distance fluid flow from geopressured zones near the basin margins are essential for the emplacement of MVT deposits in the deposits in Upper Silesian deposits (Carpathian Fore Deep) as well as at Wiesloch (Rhein Graben). Long-term fluid flow lasting until the Tertiary has been proven by isotope studies for both. Isotopically very heavy S analyzed from barite (+ 93‰) is due to hydrocarbon expelled from the Tertiary beds in the adjacent Upper Rhine Graben (Gehlen, 1966). In the Upper Silesian mineral district palaeomagnetic studies allowed for fluid flow to be correlated with processes in the Carpathian Fore Deep (Leach et al., 1996). A perfect seal to the per ascensum metalliferous fluids percolating through the Lower Muschelkalk was provided by the evaporitic rocks of the Middle Muschelkalk that is representative of low system tracks (LST) and to the mineralizing fluids within the Upper Muschelkalk beds by the argillaceous rocks of the Lower Keuper which is representative of the transgressive system tracts (TST). 16c K: The stratabound and horst-bound sulfidic to non-sulfidic Irish-type Pb–Zn–Ba–Ag–(Ni–Co) deposits Navan, Lisheen, Silvermines, Tynagh and Galmoy have raised Ireland to the top level of major Zn producer in Europe (Figs. 18.22, 18.23) (Russell et al., 1981). In Lower Carboniferous carbonate rocks, ore-forming processes went on for a 50 million year time-span. In a shallow tropical sea primary sulfides (sphalerite, galena, pyrite, marcasite) syngenetically concentrated in mud in the near-shore environment. Sulfides replaced carbonate minerals, occur as sedimentary layers, disseminations, breccia/fracture infill and massive sulfides This mineralization continued into diagenesis

Fig. 18.11. Paleogeography and geotectonic setting of Rammelsberg-type deposits In the Rheinisches Schiefergebirge, Germany (Large, 1980).


161

Fig.18.12. Textures of SHMS ore deposits. a) Galena and sphalerite-enriched ore with little pyrrhotite. Bedding textures are well-preserved in the ore body of the Sullivan Mine, Canada. b) Bedded sphalerite–galena–pyrite ore at Rammelsberg Mine, Germany. c) Melier ore composed of brass-colored slumped flasers of chalcopyrite–pyrite, lensoid nodules of pyrite, gray lenses of sphalerite and dark gray argillaceous gangue minerals at Rammelsberg Mine, Germany.

Fig. 18.13. Idealized cross section through the SHMS Pb–Zn ore deposit Sullivan, Canada. Modified after Hoy (1982).

162


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179

180


Fig. 20.07. High-sulphidation gold deposits Kochbulak and Kyzilolmasai, South Uzbekistan. a) Breccia-pipe containing Au–Te minerals, Kochbulak. b) Gold-bearing stockworktextures and beresitization, Kochbulak. c) Au telluride- bearing veinlets, Kyzilolmasai. d) Alteration with disseminated fuchsite, Kyzilolmasai.

behavior in pyritic ores include very fine-grained native gold inclusions in pyrite, or the presence of gold-bearing tellurides (Table 20.01). 20.2. Magmatic gold deposits Volcanic massive sulfide deposits (VMS)-Kuroko-type (19f D), Besshi-type and Abitibi-Ural-type (19a B) Cyprus-type (19b B), Iberian-type (19e B–19e D) and VMS deposits of unknown type in medium to high grade metamorphic terrains (19i BCD) have gold concentrations of economic grade. Many of them attracted the attention of mining engineers primarily for their base metal contents such as the Pb–Zn and Cu–bearing VMS deposits and gold is won as a by-product (Fig. 18.31). Hutchinson (1990) showed in an idealized cross section through volcanic- and sediment-hosted VMS deposits that there is a continuous decrease in the Au content of ore and footwall rocks from Cyprus-type to MVT deposits (Fig. 20.01) or in other words, from VMS deposit emplaced within an attenuated oceanic crust to those evolving within thick continental crust. The silver content shows an antithetic trend to gold and is highest in deposits of crustal origin. 19c BC: The Doyon, Mouska and the LaRonde–Penna mines are located within Archaean rocks (2703 ± 2 Ma) in Quebec, Canada. All share the same magmatic rocks. The mineralized interval is found within the Blake River Group which comprise roughly comparable amounts of felsic volcaniclastics and mafic flows and volcaniclastics that were intruded by gabbro, quartz–diorite, porphyritic tonalite and trondhjemites. The ore lenses are associated with alteration zones composed of kyanite, andalusite, pyrophyllite and staurolite,

garnet (spessartite, almandine), biotite, carbonates, chlorite, epidote, phengite and quartz. The alumosilicates direct one's thought to the epithermal Au–Cu deposits which plot close to this type of deposits. In addition to the VMS deposits, these greenstone belts are also host to intrusion-related and shear zone-hosted Au–Cu vein deposits (Mercier-Langevin et al., 2007a,b). The Goldcorp and Campbell Red Lake gold mines lie within the NE part of the Red Lake Greenstone Belt in 2989 ± 3 Ma Fe-tholeiites close to the folded contact with younger chemical and clastic sediments. The deposit features regional silicification and carbonatisation that is enriched in As, Sb and Au. Penczak and Mason (1997) envisaged a metamorphosed Archaean epithermal Au–As–Sb–Zn–(Hg) vein mineralization at the Campbell Mine. Volcanogenic massive sulphides with Au, Cu, Zn and Ag are found also within the Chibougamau District of northern Quebec, Canada. A great deal of gold is found also in ductile shear zones within meta-anorthosite and gabbroic host rocks. 19a A: Gold is not only associated with other base or precious metals but also found closely associated with industrial minerals such as magnesite and talc. The prime example is known as listwaenite-type Au deposit related to ultramafic magmatic rocks, also described as a structurally controlled epigenetic deposits within or adjacent to ophiolitic ultramafic rocks. They share many features with the mesothermal gold deposits (Barr, 1980; Buisson and Leblanc, 1985). The term listwaenite is used to describe an assemblage of alteration minerals produced by carbon dioxide metasomatism of serpentinized ultramafic rocks. Ophiolite-hosted gold veins, such as Grass Valley, Alleghany, Atlin, and Cassiar, Canada,


181

Fig. 20.08. Low-sulphidation gold deposit Kremnica (epithermal Au–Ag deposit), Slovakia and Arkhaly, Kazakhstan. a) Andesitic stratovolcano with quartz stockwork-lie veinlets, Kremnica. b) Gold-bearing quartz veinlets, Kremnica. c) Adularia–gold–quartz vein, Arkhaly. c) Rhyolitic lava with fluidal texture forming the rim of the caldera, Arkhaly. e) Lahar near the caldera margin, Arkhaly.

are contained in fault-bounded, lenses of oceanic magmatic crust. Listwaenite-altered ultramafic rocks are associated with the ophiolite-hosted gold veins, but rarely host them. Gold concentrations

were also found in steatite deposits in the environs of Miassy, Russia, and Betts Cove Complex Newfoundland, Canada (Bédard et al., 2000) (see also section on talc).

Fig. 20.09. Orogenic gold deposits in the Asian Central Interior. a) Steeply-dipping mesothermal Au–Sb vein-type deposit at Akbakay, Kazakhstan (180 m level). b) Main gold-bearing quartz lode dipping gently towards the right is surrounded by stockwork-like gold-bearing quartz veinlets at Muruntau gold deposit, Uzbekistan (Wall et al., 2004). Scale bar: 1 m. The richest gold ore is confined to the stockwork mineralization.

182


Fig. 20.10. Dalradian orogenic gold-quartz veins in Northern Ireland, Great Britain. a) Horse-tailing gold-quartz veins at Curraghinalt Mine. b) Chalcopyrite–pyrite–arsenopyrite– gold mineralization at Curraghinalt Mine. c) 1st gen. arsenopyrite–pyrite mineralization with invisible gold. d) 2nd gen. galena–arsenopyrite–pyrite–quartz mineralization at Cavanacaw Mine. e) Neoproterozoic metapelitic country rock at Curraghinalt Mine.

19d D: Intrusion-related gold deposits (IRG systems) are genetically linked with relatively reduced granitoids, e.g., Pogo, Fort Knox, USA, Kidston and Timbarra, Australia (Bakke, 1995; Baker and Andrew, 1991; Mustard, 2001; Dilworth et al., 2002). The characteristic features of this type of gold deposits, which developed landward of Phanerozoic accreted terranes in the Paleozoic of eastern Australia and the Mesozoic of the northern North American Cordillera, were summarized by Thompson et al. (1999), Thompson and Newberry (2000) and Lang and Baker (2001) (Fig. 20.02). IRG systems occur outside the field of porphyry Cu–Au deposits in continental mainly carbonaceous or carbonate-bearing sedimentary units. The host rocks of IRG deposits are mainly strongly fractionated metaluminous, calcalkaline, granodiorites to granites of I-type affiliation with SiO2 contents between 50 and 76% SiO2. Deposits of Sn, W and Bi generally occur in strongly fractionated rocks with SiO2 contents between 70 and 77% SiO2. These rocks are responsible for high-crustal-level mineralizations with prevalently Bi, Mo, W, Sn, U, Au and Ag minerals that developed in a weakly oxidized to weakly reduced oxidation state according to Blevin and Chappell (1992). The mineralizations display a pronounced lateral (W–Mo, Sn, Bi, Au, As, Sb, Zn, Pb, Ag) and vertical mineral zonation (Bi increases, As, Sb decrease with depth). Their structural position relative to the faultbound Au mineralization is illustrated in (Fig. 20.03). The similarities between intrusion-related Au–W–Bi and Sn–W deposits are shown in Fig. 20.04 (Baker et al., 2005).

For the Muruntau and Charmitan gold–quartz vein deposits, Uzbekistan, Wall et al. (2004) considered the Hercynian magmatic activity played a key role by providing fluids and causing widespread fracturing (Fig. 20.09). Based on noble gas data a small input is supposed to have come from an external source (mantle and lower crust). Porphyry Cu–Mo–Au deposits including co-magmatic breccia pipes (19c DE), calc-alkaline porphyry Cu–Au–Mo deposits with Cu– (Mo) breccia pipes (19c D) and alkalic porphyry Cu–Au/diorite porphyry copper (19c E) are an important source of gold to meet the world demand. The characteristics of calc-alkaline porphyry-related Au (Cu–Mo) deposits have been listed in Table 11.04 applicable most notably to the Cretaceous and Cenozoic porphyry systems. There are also some very complex auriferous porphyry systems of Paleozoic age such as the gold deposit Oktorkoy (Taldy–Bulak Levoberezhny), Kyrgyzstan (Fig. 20.05) that does not show the entire pattern typical of porphyry-type deposits (Seltmann and Porter, 2005). It is bound to an E–W trending island arcs on a Devonian–Carboniferous active continental margin of the Central Asian paleo-ocean above a north dipping subduction of the oceanic crust. The Oktorkoy ore district is characterized by intensive thrusting and polygenic mineralization associated with multiphase granitoid complexes. The alteration pattern includes: potassium alteration, phyllitization (beresitization), propyllitization (epidote–chlorite–albite), tourmalinization and argillization. Carboniferous to Permian aged gold-bearing hydrothermal


183

Fig. 20.11. Comparison of gold fineness and grain size of placer and laterite gold (Santosh et al., 1992). a) Plots of fineness vs. grain size gold from primary gold deposits, laterite and placer gold deposits. b) Variation in grain size among different types of gold deposits in the Nilambur Valley, India.

breccia pipes bound to porphyritic intrusions were recorded from northern Queensland, Australia by Perkins and Kennedy (1998). These Au deposits bridge the gap towards the intrusion-related gold systems coded 19d D in this classification scheme. Some Au–Ag–Cu– (Pb–Zn) porphyry systems are transitional into skarn and epithermal deposits as it is the case with the Paleogene deposit at Recsk, Hungary (Seres-Hartai and Földessy, 2001). 19g CD: Gold is also recovered from many skarn deposit as a byproduct similar to many porphyry deposits during mining of base metals, tin or tungsten. In the following paragraph only pyrometasomatic Au skarn deposits are considered which deserved the term “gold skarn” in the economic sense suggested by Einaudi et al. (1981) in that they are mined exclusively or predominantly for gold and exhibit calc-silicate alteration, dominated by garnet and pyroxene genetically related to the gold mineralization. One of the first discoveries of that kind was the gold skarn deposit in the Hedley district, British Columbia, Canada (Ettlinger et al., 1992; Ettlinger and Ray, 1993). Other finds of similar Au skarn deposits were made in Andorra, Spain, at Junction Reefs, Australia, at Crown Jewel, Elkhorn, Redline USA, Nambija District, Ecuador, (Theodore and Hammarstrom, 1991; Hickey, 1992; Everson and Read, 1992; Gray et al., 1995; Fontboté et al., 2004). The Au skarn deposits were emplaced within carbonates, calcareous clastics, volcaniclastics, volcanites intrusions of calc-alkaline, less frequently of alkalic composition in orogenic belts at convergent plate margins and arc or back-arc environments. The mineralogy is highly variable and depends on the type of skarn deposit: (1) Mg skarns: gold, pyrrhotite, pyrite, magnetite, Cu–Pb sulfides, (2) pyroxene skarn: gold, pyrrhotite, arsenopyrite, Cu–Zn–Co–Bi sulfide and ±Au tellurides

with high sulfide contents and pyrrhotite: pyrite ratios, (3) garnet and (4) epidote skarns resembling pyroxene skarn with respect to mineralogy but minor sulfide and low pyrrhotite: pyrite ratios. Mueller (1997) described the Nevoria gold skarn deposit of the Archaean Yilgarn craton, Western Australia occurring in amphibolite facies greenstones between two dome-shaped granitoid batholiths. The gangue is calcic, highly reduced, with accessory arsenopyrite–loellingite, and chalcopyrite. Native gold is enclosed in hedenbergite, actinolite, almandine, and quartz and occurs together with the alloy maldonite and a suite of bismuth tellurides. This skarn formed in a deep midcrustal environment (10–15 km) during late Archaean magmatism very much distinct from the copper–gold skarns in Cenozoic continental margins and from goldrich iron-copper skarns in Cenozoic to Mesozoic island-arc terranes, which all formed at much shallower depths of 2 to 5 km. In contrast to the reduced or typical skarn deposits also reported from the Sn–W camp, that are characterized by low garnet: pyroxene ratios, hedenbergite, abundant sulfides and dominated by pyrrhotite and arsenopyrite, there are gold skarn deposits named as oxidized gold skarns by Brooks et al. (1990). Their characteristics include high garnet: pyroxene ratios, relatively Fe-poor garnet and pyroxene, low total sulfides, pyrite ≥ pyrrhotite, and minor but ubiquitous occurrences of Pb–Cu– Zn sulfides. Not surprisingly, these deposits locally are transitional to other types of shallow or near-surface gold mineralization such as epithermal deposits in which phase separation by boiling is the major process of precipitation. 19b CD: Epithermal gold deposits or shallow high- and lowsulphidation Cu–Au–Ag deposits encompass a large spectrum of mineral deposits called high sulphidation-type/quartz alunite-type Au–Ag–Cu

184



185

Fig. 20.13. Precambrian quartz–gold pebble conglomerate from the Witwatersrand Venterdorp Contact Reef at East Driefontein Gold Mine/C-Facies, South Africa.

deposits, low sulphidation-type/adularia-type Au–Ag deposits and subtypes such as hot spring Au deposits bound to subaerial siliceous sinter. The complex intertonguing of mineral assemblages in epithermal gold deposits strongly varying with depth is depicted in the cartoon designed by Buchanan (1981) (Fig. 20.06) and the comprehensive papers by Hedenquist et al. (1996) and Albinson et al. (2001). Epithermal means low temperature and refers to a heat system or hydrothermal systems emplaced at shallow depths of less than1 km in comparison to the deep-seated mesothermal or orogenic gold deposits of higher temperature. The tripartite subdivision into katathermal, mesothermal and epithermal is not new among economic geologists and has already been used by Schneiderhöhn (1962) in his short course on mineral resources although he placed more emphasis on Hg and Sb as marker elements of this type of mineralization rather than on precious metals. A brief review of these precious metal deposits reveals that through time these precious metal deposits were treated under various terms such as gold–quartz veins , andesites to gold–alunite deposits (Lindgren, 1934). There are terms like high-sulphidation, intermediate sulphidation or alunite–kaolinite type deposits (Arribas, 1995; Sillitoe, 1999; Hedenquist et al., 1998, 2003). The most widespread types of epithermal deposits are distinguished based on the sulphidation states of their primary sulfide assemblages and grouped into highsulphidation (HS) and low-sulphidation (LS) epithermal deposits with some intermediate or transitional stages (IS) between the two (Hedenquist et al., 2000). Most HS deposits are generated in extensional to neutral calc-alkaline andesitic–dacitic arcs, a few also occur in compressive arcs. Highly acidic fluids create advanced argillic lithocaps on top of porphyry systems, IS deposits although bound to a similar spectrum of host rock lithologies are not tied up with underlying porphyry systems. Many LS deposits are associated with bimodal (basalt–rhyolite) volcanic suites within extensional geodynamic settings (Sillitoe, 2002). There are also LS- non-magmatic epithermal Au– Ag deposits which were recorded form the Rhodopes mountains, Bulgaria (Marchev et al., 2004; Marton et al., 2006). Vuggy quartz surrounded by halos of advanced argillic quartz–alunite alteration and pyrophyllite are typical of HS epithermal deposits (Sillitoe, 1995). HS-type Cu–Au deposits originated from subaerial volcanic complexes above degassing magma chambers genetically related to high-level intrusions. Under acidic and strongly oxidizing conditions the magmatic hydrothermal fluids generated a varied spectrum of ore

Fig. 20.14. The Zaamar placer gold deposit, Mongolia. a) Neogene gold-bearing fluvial deposits cut into the footwall coal seam. b) Ramp leading into the open cast mine of the placer deposit runs perpendicular to the paleocurrent direction (see cars for scale).

minerals involving pyrite, enargite/luzonite, chalcocite, covellite, bornite, gold, electrum and chalcopyrite. Locally, silver sulphosalts and tellurides are also present. It is the alteration mineralogy that makes it different from the LS-type mineralization with quartz, kaolinite/dickite, alunite, barite, hematite, sericite/illite, pyrophyllite, andalusite, diaspore, corundum, topaz, zunyite, jarosite, APS minerals such as hinsdalite, woodhouseite and crandallite. A special type of alteration was found in some HS-type deposits in central Asia such as Kyzilolmasai and Kochbulak, Uzbekistan, consisting of quartz, sericite, ankerite and pyrite which was named after the Berezovskoe gold deposit, in the Middle Urals, as beresitization (Fig. 20.07). Examples of this HS-type occur at Summitville, USA, Nansatsu, Japan, El Indio, Chile, Pueblo Viejo, Dominican Republic, Rodalquilar, Spain, Kochbulak, Uzbekistan and Lepanto, Philippines. Hot spring-type gold deposits close to the paleosurfaces also occur in the western USA at Round Mountain. Porphyry-related base-metal veins and HS epithermal systems share many common features thereby demonstrating the close genetic relationship (Nelson, 2000; Einaudi et al., 2003; Plotinskaya et al., 2006).

Fig. 20.12. Fluvial and alluvial placer gold under the SEM-EDX and stereomicroscope. a) Fluvial placer gold with milled particles and wrapped at the edge of gold flakes from the Selenga River, Mongolia. b) Alluvial gold from the catchment area of the Rhein River in the Gotthard Massif, Switzerland. c) Alluvial gold–silver alloy (electrum) from the catchment area of the Rhein River, Switzerland containing 12 at.% Ag. d) Alluvial gold grain with inclusions of quartz (“rock crystal”) from the Pflaumbach drainage system near Pleystein, Germany (source: B. Weber). e) Native gold forming branched crystal aggregates, Rosia Montana, Roumania. f) Native gold intergown with Cr-bearing muscovite (fuchsite) Kerr– Addison Mine, Virginiatown, Ontario, Canada. g) Sylvanite Nagyag, Roumania.

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Epithermal systems at Creede, USA, El Bronce, Chile, Guanajuato, Tayoltita, Mexico, Baguio, Philippines, Kremnica, Slovakia, Arkhaly, Kazakhstanand Pongkor, Indonesia belong to the LS-type or adulariasericite type that formed from relatively diluted and cool solutions of almost neutral pH and under reducing conditions as magmatic and meteoric fluids are mixing. Minerals precipitated as the solutions undergo cooling and degassing by fluid mixing, boiling and decompression in stockworks and veinlets (Fig. 20.08) (Heald et al., 1987; Konečný et al., 1995; Lexa, 1999; Lexa et al., 1999; Cooke and McPhail, 2001). This processes lead to a wide range of ore minerals such as pyrite, electrum, gold, Ag minerals, chalcopyrite, sphalerite, galena, tetrahedrite, Te and Se minerals. Near-surface mineralization with Au–Ag–As–Sb–Hg may gradually change into Ag– Pb–Zn mineralization at depth. The alteration mineralogy is strikingly different from that of HS-type deposits containing silica, adularia, sericite, barite, fluorite, Ca– Mg–Mn–Fe carbonate minerals hematite and chlorite. The majority of epithermal systems evolved during the most recent geological periods. There are also older LS, gold-bearing epithermal vein system of Late Neoproterozoic age recorded from Newfoundland (Mills et al., 1999). Epithermal hot spring Au deposits bound to subaerial siliceous sinter were described among others by Marcoux et al. (2004) from the Massif Central. The lead isotopic composition in the studied deposit is not very radiogenic (206Pb/204Pb=18.20) and is similar to that of the Late Hercynian gold lodes in the Massif Central. The mineralized sinter appears to have derived from a geyser-type hot springs. Other goldbearing sinter occur at McLaughlin, USA, and within the Waiotapu geothermal system in New Zealand (Sillitoe, 1993; Tosdal et al., 1993). They form on top of various kinds of volcanic and, in places, sedimentary rocks and may be interpreted as the topmost parts of the afore-mentioned epithermal systems. In addition to gold and electrum, they contain considerable amounts of pyrite, marcasite, stibnite, sulphosalt, realgar and cinnabarite. The silicified zones rarely contain kaolinite, adularia or alunite diagnostic of the afore mentioned LS- and HS-type deposits. For Mn–Au–Ag–Fe–Ba–F low-sulphidation deposits (19a D) see Mn deposits. 19a E: Among the Au–Ag–Te deposits related to alkaline magmatic rocks, there is one example that warrants mentioning for its exclusive geological position. The Emperor gold deposit at Vatukoua of Viti Levu, Fiji, lies on the Circum-Pacific Volcanic Belt and belongs to the category of alkalic intrusion-related epithermal Au–Ag–Te veins (Muller and Groves, 1993). The host rocks are mainly olivine flow basalts, trachybasalt, pyroclastics and trachyandesite dykes of Tertiary age. Veins are dominantly composed of quartz with minor pyrite, tellurides, gold, Zn–Cu–As–Pb sulfides. The Tuvatu low-sulphidation Au–Ag telluride deposit is the second largest deposit of this type in Fiji (Scherbarth and Spry, 2006). Similar deposits are mined at Creeple Creek, USA, and Porgera in Papua New Guinea (Richards, 1997; Richards et al., 1997; Kelley and Ludington, 2002). The alteration mineralogy of these epithermal Au deposits is characterized by adularia and their ore mineralogy by Au tellurides and other Te minerals. What makes Creeple Creek, USA, different from the overall epithermal system is the association with alkalic intrusive rocks in an extensional faulting and continental rifting environment. Emperor and Porgera are situated within an oceanic arc setting. 19h CD: The disseminated Carlin type Au–Ag deposits from the locus typicus in the USA, Mesel, Indonesia, Guizhou, China, SuzdalTrend, Russia, Alsar, Macedonia and Lucky Hill, Malaysia belong to a mixed type with respect to the host rock lithology (Hon, 1981; Percival et al., 1990; Ramadorai et al., 1991; Turner et al., 1994; Dill and Horn, 1996; Groff et al., 2002; Peters et al., 2002a,b; Narseev, 2002; Leach, 2004; Emsbo et al., 2006; Volkov et al., 2006). Carbonate-bearing host rocks prevail, but non-carbonate siliciclastics as well as magmatic rocks may also act as host rock of Carlintype Au–Ag deposits. The ore bodies are highly divers and range from disseminated, stratabound at contacts between contrasting

lithologies to faultbound. They are emplaced in contrasting geodynamic setting such as passive margins and island arcs. Native gold is micron-sized and accompanied by pyrite, arsenopyrite, stibnite, realgar, orpiment, cinnabarite, fluorite, barite and Tl minerals. There is direct link to Tl, Hg, As and Sb deposits discussed later. As constraining the style of mineralization a wide spectrum of models were put forward by the above authors to explain this Au mineralization, involving epithermal models related to shallow magmatism and basin and range extension, as well as distal skarn models. Au is related to the collapse of intrusion-centered porphyry-type hydrothermal systems and mixing models of different fluids in the aftermaths of crustal extension. 20.3. Structure-bound gold deposits 19a G: Terms like mesothermal gold-quartz veins/shear-zone hosted Au deposits and turbidity-hosted gold–quartz veins have traditionally been used to describe the vein-type gold deposits (Keppie et al., 1986; Nesbitt et al., 1986; Groves et al., 1998). By definition mesothermal gold deposits formed between 200–350 °C within a series of accreted terranes. Often the criteria based on which the classification has been made are overlapping and therefore the plethora of host-rock and temperature-related classification schemes has set an end by the unifying term orogenic gold deposits by Groves et al. (1998), Hagemann and Cassidy (2000). Orogenic gold deposits have the lion share in the world's gold supply and some of the most renowned deposits such as Ashanti, Ghana, Bendigo, Ballarat, Kalgoorlie, Australia, Kumtor, Kyrgyzstan, Vasil'kovsk, Kazakhstan, Berezovsk, Russia, Mother Lode Homestake, USA, Timmins, Lamaque, Macassa, Sigma, Val d'Or camp, Canada, Salsigne, France, and Morro Velho, Brazil are attributed to this class of gold deposits. These epigenetic gold vein-type deposits in metamorphic terranes include those of the Precambrian shields, particularly the Late Archaean greenstone belts and Paleoproterozoic fold belts, and of the late Neoproterozoic and younger Cordilleran-style orogens (Goldfarb et al., 2005). The ore mineralogy is characterized by minerals containing Au, Ag, As, Sb, Te and W minerals with minor Pb–Zn–Cu mineralization. A rather atypical mineralization is reported from the Sirkka Au–Co–Ni– Co shear-zone-hosted deposit in Finland (Holma and Keinänen, 2007). They are accompanied by alteration zones rich in K mica and carbonates that surround crustal-scale faults and shear zones through siliciclastics, iron formations and carbonate rocks that generally underwent greenschist facies but may also reach amphibolite facies conditions (Figs. 20.09, 20.10). The veins are metamorphogenicdeformational by origin and formed synchronously with late phases or prior to the major deformational event of orogenies. The fluids are tectonically or seismically driven by a cycle of pressure build-up that is released by failure and pressure reduction followed by sealing and repetition of these processes (Sibson et al., 1988). Large vertical extent with subtle vertical zonation may be recognized. Most of these deposits may be subdivided into epizonal, mesozonal, and hypozonal subtypes based on pressure–temperature conditions of ore formation (Fig. 20.03) (Groves et al. 2003; Goldfarb et al. 2005). See also the structure-bound Sb (20b G) and Hg deposits (23a G) as shallow representatives of some of these gold deposits. 19a F: The Olympic Dam iron oxide breccia complex (ODBC) with its Cu–Au–Ag–U–REE deposits developed within the Mesoproterozoic (1600–1585 Ma) Roxby Downs Granite a pink to red colored, undeformed, unmetamorphosed, coarse- to medium-grained, syenogranite with A-type affinities (Reeve et al., 1990; Hitzman et al., 1992; Hopper and Correa, 2000; Belpario and Freeman, 2004; Williams and Pollard, 2003; Williams et al., 2005). The known deposits of this type are found within Early to mid-Proterozoic host rocks. According to Hitzman et al. (1992) deposits are located in areas that were cratonic or continental margin environments during the late Lower to Middle Proterozoic in close association with extensional tectonics. The Cu–


Au–Ag–U–REE mineralization occurs in a hematite-rich granite breccia complex in the Gawler Craton and is overlain by of flatlying sedimentary rocks of the Stuart Shelf geological province. The ODBC is a funnel-shaped ore body of brecciated and hydrothermally altered granite with a barren, hematite-quartz breccia “core” surrounded by mineralized hematite–granite breccia bodies. The mineral assemblage comprises chalcopyrite, bornite, uraninite (pitchblende), coffinite, brannerite, gold and silver associated with REE-bearing mineral (mainly containing La and Ce with F) such as bastnaesite and florencite. This type of U deposit counts for 66% of U to be exploited in Australia. Ore mineralization is structuralcontrolled and of hydrothermal origin and unrelated to magmatic activity. It developed in what might be called a caldera or a maar. Principal mechanisms which formed the breccia complex are (1) hydraulic fracturing, (2) tectonic faulting, (2) chemical corrosion and (4) gravity collapse. Much of the brecciation occurred in a nearsurface eruptive environment of a crater complex during eruptions caused by boiling and explosive interaction of water from lake, sea or groundwater with magma. 20.4. Sedimentary gold deposits 19a H: Laterite-saprolite Au–(PGE–Sn)–Fe–Al–Mn deposits developed on a wide variety of bedrock types especially in greenstone belts and other gold-bearing terrains, provided there is stable and prolonged weathering under tropical to subtropical conditions (Boyle, 1979; Bowles, 1986; Monti, 1987; Symons et al. 1988). Laterites, bauxites and saprolites develop in areas where geomorphology allows sufficient drainage, so that oxidation may penetrate deep and promote extensive leaching of alkaline and earth alkaline elements (Schellmann, 1971). Primary gold is partially moved in solution and re-precipitated in secondary gold particles lower in grain size and Ag contents than the primary gold ore but higher than in alluvial Au placers which may derive from such intermediate gold repositories (Santosh et al., 1992) (Fig. 20.11). Boddington in Western Australia are the most well-known lateritic gold ore deposits evolving on a hydrothermal primary ore deposit (Symons et al., 1990; Morriss, 1993; Anand and Butt, 1998). Apart from gold, electrum, heavy minerals bearing PGE and Sn may also be present within these duricrusts made up of Fe–, Al– and Mn– oxides/hydroxides and kaolinite. 19a I: Modern Au-PGE placers are a direct consequence of fluvial erosion into lateritic or gossan-related gold accumulations. Alluvial gravel and conglomerate mainly with quartz clasts are widespread and form the most preferential traps of secondary gold, e.g., Graham Island (PGE, Au), Canada, Nome (Au), USA, Westland and Nelson provinces (Au), New Zealand, San Luis Range (Au), Argentina (Emory-Moore, 1991; Youngson and Craw, 1995; Patyk-Kara, 1999; Marquez-Zavalia et al., 2004; Dill, 2008b). Gold, Pt–Fe alloys, Os–Ir alloys are found together with a wide range of other heavy minerals in alluvial, beach, and aeolian deposits (Dill et al., 2009a, b). High-energy systems such as braided-river drainage systems where gradients flatten and river velocities lessen, below rapids and falls are preferred loci of placer accumulation (Burton and Fralick, 2003). An insight into such a channel system cutting into coal seams is exposed in the Neogene Zaamar placer gold deposit, Mongolia (Fig. 20.14). The landform plays a decisive role in siting of these placer deposits (Dill and Ludwig, 2008). The winnowing action of waves causes Au concentrations in the beach environments. Gold assumes different shapes from milled gold in fluvial placer deposits attesting to the hammering effect of quartz clasts to irregularly-shaped nuggets originating from gold reprecipitation in the upper reaches of drainage systems mediated by algae and bacteria (Fig. 20.12). 19b I: The uranium-bearing conglomerates at Blind River-Elliot Lake, Canada, and Witwatersrand, South Africa occur in stratigraphic series that developed when the Earth's atmosphere was a different one from today's oxidizing conditions (Pretorius, 1981). Buried Au– U–(PGE–Th) paleoplacers or quartz pebble conglomerates formed

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from the Archaean to the Early Proterozoic (3,100–2,200 Ma), some also extending into the Middle Proterozoic such as the Tarkwa (1,900 Ma). World-class deposits are at Witwatersrand, South Africa, Elliot Lake, Canada, Sierra do Jacobina, Brazil, Tarkwa, Ghana and the Guyana Shield area in Colombia (Els and Beukes, 1987; Els, 1988; Meyer et al., 1990; Hirdes and Nunoo, 1994; Robb and Meyer, 1996; Teixeira et al., 2001; Milesi et al., 2002; Malitc and Merkle, 2004; Frimmel, 2005) (Fig. 20.13). The oligomictic mature conglomerates intercalated with sequences of less mature conglomerates and sandstones were spread across Archaean granite-greenstones. Braided streams in alluvial fans concentrated gold at the base of mature conglomerate beds laid down on an erosional surface. Carbonaceous layers akin to algal mats and interpreted as the representative of the low-energy regime deposit within the fan contain U and fine Au. Gold in the various deposits may be part of the same evolutionary event during the Precambrian at an atmosphere completely different from that of the younger geological epochs. Processes put forward to account for the variegated mineral assemblage involving quartz, gold, pyrite, uraninite, brannerite, zircon, chromite, monazite, Ru–Os–Ir–Pt–Fe alloys and sperrylite are as varied as the mineral association running from sedimentary through hydrothermal mobilization. 19a J: The banded iron formation (BIF) is not only the most important source of iron but also of utmost importance as source of gold. The origin of gold within the BIF e.g., Kraaipan Greenstone Belt, South Africa, Homestake, USA, Mt. Morgans, Western Australia, Itabira District, Brazil, Vubachikwe, Zimbabwe, Kolar District, India, Maevatanana, Madagascar range from epigenetic models to syngenetic models. Some gold deposits such as Homestake, USA, or Nevoria, Australia, may be looked at from different angles und grouped as structure-bound or BIF-related thereby demonstrating anything but a multi-stage mineralizing process, e.g. Nevoria, Australia (Sawkins and Rye, 1974). Epigenetic Au mineralization originated from metamorphogenic or magmatic hydrothermal fluids ascending from depth and discharge of elements into a favorable depositional environment during the waning stages of orogeny. Syngenetic gold mineralization is supposed to have been formed by active submarine exhalative systems within a basin with chemical sedimentary series as it was already described for the Fe deposits (Phillips et al., 1984; Saager et al., 1987; Iddaiah et al., 1994; Vielreicher et al., 1994; Hong-Rui et al., 2000; Fan et al., 2004; Hofmann et al., 2003; Hammond and Moore, 2006; Andrianjakavah et al., 2007). Native Au, pyrite, arsenopyrite, magnetite and pyrrhotite are the major ore minerals Cu-, Zn-, Pb-, Sb sulfides but tellurides may also be present at minor levels. Gangue minerals are almost the same as reported for the Fe deposits. The Serra Pelada Au-Pd–Pt deposit, Brazil, hosted by weakly metamorphosed sedimentary rocks of Late Archaean age warrants a special treatment (Cabral, 2006). Abundant coarse-grained dendrites of palladiferous gold occur together with Pd–Pt–Se, Pd– Se, Pd–Hg–Se and Pd–Bi–Se phases and native palladium. Mineralogical and geochemical features point to highly oxidizing (hematite stability field) and relatively shallow-level conditions, which are highlighted by the Se minerals. Olivo et al. (2001) recorded Au–Pd mineralizations at the Conceicao Iron Mine, Itabira District . Brazil. 19b J: With resources estimated at 1100 t Au at an average grade of 2.45 g/t, the black shale-hosted Sukhoi Log Au-PGE deposit is the largest concentration of gold in Russia (Distler et al., 2004). The Au mineralization and associated PGE mineralization are located in metamorphosed carbon-bearing Upper Proterozoic sedimentary rocks. Nearly 90 minerals occur at Sukhoi Log including native metals, intermetallic alloys, sulfosalts, phosphates, tungstates, and oxides. Native gold is the main ore mineral, but calaverite, hessite, petzite, and krennerite also occur. PGE mineralization is located on the periphery of the gold mineralization. A three-stage genetic model has been developed: (1) deposition of carbon-bearing

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terrigenous-carbonate rocks in an intraplate rift around 800 Ma, (2) regional metamorphism at 520 Ma , (3) hydrothermal activity at 320 Ma. 20.5. Metamorphic gold deposits 19i IJ: There are numerous gold deposits either structurally controlled or bound to volcanosedimentary series in metamorphic host rocks. They have been discussed in previous sections. Apart from these types of gold deposits metamorphic Au–As–Pd–Hg–Bi– Te–W mobilizates were recognized by Moravek and Lehrberger (1997) neither bound to a premetamorphic mineralization nor structurally controlled. Gold in these mineralizations often stratiform to the metamorphic structures of the country rocks originated from fluids generated by metamorphic dehydration processes. Dynamo-metamorphic processes and retrograde regional metamorphism have had a major impact on these gold mineralizations containing a variegated spectrum of gold- and non-gold minerals: gold, arsenopyrite, loellingite, scheelite, molybdenite, pyrite, pyrrhotite, stibnite, palladium-, mercury and bismuth-tellurium minerals. There seems to be no sharp boundary between these metamorphic mobilizates and those gold deposits commonly denominated as orogenic gold deposits (19a G). Some enigmatic PGE mineralizations in metamorphic basement rocks may also be correlated with processes like those referred to for gold. This is especially true for Pd-enriched mineralizations in oxidizing environments undergoing pervasive dynamometamorphism indicated by the presence of hematite and strongly sheared crystalline host rocks (Dill et al., 2009a). 20.6. Gold supply and use Native gold and electrum (N30% Ag) are the major ore minerals as far as visible gold is concerned. There are many rare tellurides which may also act as gold ore minerals, e.g., calaverite (43.56% Au) , nagyagite (18.6% Au) or sylvanite (34.4% Au). Arsenopyrite and pyrite frequently form a source of invisible gold. In unconsolidated rocks such as placers contents of greater than 0.2 ppm Au may be of interest, in hard rocks it is strongly dependant on the accompanying elements, but in some veins grades exceeding 2 ppm Au are feasible. Silver increases the value of the ore, too much Cu is detrimental. Using amalgamation for the recovery of gold, As, Sb and Cu are deleterious to the process, leaching gold from its host rocks with cyanide, Cu and carbonaceous matter worsen the output. Carbon-in-pulp is a recovery process in which a slurry of gold ore, free carbon particles and cyanide are mixed together. The solution is passed counter current through tanks containing activated carbon particles. Loaded carbon is separated from the slurry by screening and stripped in a caustic cyanide solution under heat and pressure. Subsequently gold is recovered by electrolysis or by zinc precipitation. Gold went primarily into jewelry or

Table 20.02 World gold production in 2005 by country and mining company. • World mine production • Proven + probable resources • Mine production by country • South Africa (12.1%) • Australia (10.7%) • USA (10.6%) • China (9.1%) • Peru (8.4%) • Major producer • Newmont Mining (USA 8.5%) • Anglogold Ashanti (South Africa 7.9%) • Barrick Gold (Canada. 6.9%) • Gold Fields (South Africa 5.7%) • Placer Dome (Canada. 4.9%)

2.460 t metal content 42.00 t metal content Top 5: 50.9% . Top 10: 74.2%

Top 5: 33.9% . Top 10: 47.5%

historically was used in the monetary system. Today chemical goods and dental medicine are another outlets for the final use of gold. The world mining production of gold is shown in Table 20.02. Update supply and use: USGS: http://minerals.usgs.gov/minerals/pubs/commodity/gold/ 21. Antimony, arsenic and thallium 21.1. Chemistry and mineralogy of antimony, arsenic and thallium Antimony is a very brittle element which is mainly alloyed with metals like Cu and Pb to harden the resulting compounds. Its average grade in the earth's crust is 0.65 ppm Sb. Thallium is less widespread than Sb and occurs only at a level of 0.5 ppm Tl in the earth's crust. Arsenic forms three different modifications. Brittle, gray metallic arsenic is stable under normal conditions, whereas its two other modifications, yellow and black arsenic are unstable. Exposed to oxygen at temperature greater than 180 °C, arsenic converts into As2O3 (“white arsenic”). Compared to the Sb and Tl it is rather widespread in the earth's crust with an average grade of 2 ppm As. Antimony forms a wide range of minerals from native Sb (100 wt.% Sb) in hydrothermal veins where it is primary mineral such as at Seinajoki, Finland, to complex mixtures of Sb oxides and Sb hydroxides called Sb ochre or an extremely rare carmine- red Sbsulfide-oxide named sarabauite after the Sarabau Mine in the Bau Mining District of Sarawak, Malaysia, where it was recorded from for the first time by Nakai et al. (1981). The most common ore mineral is stibnite (72 wt.% Sb) and its metaform metastibnite (72 wt.% Sb). Selenides such as antimonselite (51 wt. % Sb) are only of mineralogical interest, as it is the case with the antimonide breithauptite which occurs in hydrothermal calcite veins associated with Co–Ni–Ag ores or seinajokite replacing native antimony in an antimony deposit in Finland. Antimony oxides valentinite (83 wt. % Sb) and its polymorph senarmontite (88 wt.% Sb) can also be of economic interest. Antimony is known to form a wide range of sulfosalts that are subdivided into two basic groups, simple sulfosalts, e.g., chalcostibite, famatinite, pyrargyrite, tetrahedrite and complex sulfosalts, e.g. boulangerite group s.s.s. with zinckenite being the most simple one, jamesonite, meneghinite or samsonite. In Fig. 21.01 stibnite ore is shown on a macroscopic scale in specimen and along a shear zone at subcrop. The most common mineral association observed in antimony veins, mainly in the initial stages of Sb vein mineralization, is composed of stibnite and arsenopyrite. Arsenic, although rather seldom found in deposits of its own, is present in a great variety of ore minerals in Au, Cu and Ag deposits. Native arsenic (100% As) occurs in ore veins and some epithermal deposits such as in the Bau Mining District, Malaysia (Dill and Horn, 1996). The most common As sulfides are arsenopyrite (46% As) covering the entire field from gold veins through pegmatites to contact metasomatic deposits and realgar (70% As) and orpiment (62% As) both are common to low-temperature veins and hot springs. In a sulfur-deficient environment of deposition arsenides occur. Many of them such as nickelite (56% As) and skutterudite (72% As), both typical of the veins of the Five-element association or cobaltite (45% As) and loellingite (73% As) known from mesothermal, pegmatitic and contact metasomatic deposits have only a limited importance. Sperrylite (43% As) common to Ni–Cu–PGE deposits and ultrabasic rocks is, however, a true ore mineral of Pt and unlike other arsenides also found in stream sediments. Arsenate such as scorodite (33% As), annabergite (25% As) and erythrite (25% As) as well as As oxides such as arsenolite (76% As) form during supergene alteration of As-bearing Ni and Co deposits and as such can be used as an ore guide. Thallium sulfides, e.g. Carlinite (92.73%wt. Tl) are rare minerals. More often Tl is found together with As in various Tl ore minerals such as christite, fangite, lorandite or in arsenide–antimonides such as jentschite or galkhaite all of which are typical of epithermal deposits (Table 21.01).


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Fig. 21.01. Antimony ore minerals at outcrop, in specimens and under the ore microscope. a) Two steeply-dipping shear zones mineralized with stibnite, pyrite, pyrrhotite arsenopyrite and gold. The wall rocks are made up of phyllites and graphite schists of the Penek Formation in the Pezinok antimony mine, Lesser Carpathian Mountains, Slovakia. b) Massive stibnite ore from Lake George, Canada. c) Strained slender crystals of stibnite besides recrystallized granular aggregates from the Sb vein at Wolfersgrün, Germany.

21.2. Antimony deposits 21.2.1. Magmatic antimony deposits 20a CD: Collision (granite)-related vein-type and replacement deposits containing Pb–Cu–Zn–As–Sb mineral associations occur at Stolice, Krupanj and Zajača in Serbia which were mined for stibnite coexisting with As–Pb–Cu–Zn minerals at the contact between dacitic to andesitic subvolcanic rocks and Carboniferous limestones

Table 21.01 Classification scheme of antimony deposits. 1) Magmatic antimony deposits 1) Granite-related vein-type (and replacement) deposits (20a CD) 1) Sb-W 2) Sb–Pb–Cu–Zn–As 3) Sb–Hg–W–Ba 2) Skarn-type Sb–(Au–Hg–As) deposit (Sarawak-type) (20d CD) 3) Carbonate-hosted disseminated Sb–Au–Ag (Carlin-type) (20e CD) 4) Shallow high- and low-sulfidation Sb deposits (20b CD) 5) Near-surface low-sulfidation-type Sb–(Hg) deposits (20c CD): 6) Sb deposits related to (ultra)basic rocks and greenstone belts (20a AB) 2) Structure-related antimony deposits 1) Shearzone-hosted, mesothermal (Au-)Sb veins 1) Monotonous Sb deposits (20a G) 2) Polymetallic Sb–(Au–W–Sn–As–Zn–Pb) deposits (20b G) 3) Sedimentary antimony deposits 1) Stratabound deposit (with remobilization) 1) Monotonous sedimentary-diagenetic Sb deposit (20a J) 2) Hg–Sb–(W) deposits in carbargillites and graphite schists (20b J) 3) Polymetallic Sb mineralization in black shales (alum shales) (20d J) 2) Sb-enriched coal seams (20a N)

Fig. 21.02. Collision and (granite)-related Sb–As–Pb–Cu–Zn deposits, Stolice, former Yugoslavia/Serbia (after Janković, 1967).

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(Fig. 21.02). They are part of a larger Sb province in western Serbia dominated by granitoids of Early Oligocene (33.7–29.6 Ma) age within Carboniferous, Permian and Triassic series (Pamić, 2002). Where the granitoids got in contact with calcareous rocks magnetite–chalcopyrite skarns developed passing further outward into galena–sphalerite mineralization surrounded by a fringe of stibnite deposits. Antimony concentration went through a maximum where silicified Upper Carboniferous reactive limestones are subjacent to impermeable Upper Carboniferous schists fostering the formation of replacement ore bodies. Vein-type deposits with fault breccia whose fragments are cemented by stibnite and calcite are among the most significant Sb producers in Europe. The outermost zone of this granite-related metallogenic province is marked by fluorite deposits (Janković, 1990). A similar type of Sb deposits at the passage from reactive calcareous rocks into impervious argillaceous rocks driven by the heat of Jurassic to Cretaceous granites evolved in western Thailand from the Lampang– Phrae Province down to the south of Surat Thani Province in Peninsular Thailand (Dill et al., 2008c). Ore mineralization may be subdivided into four stages: (1) late mesothermal/wolframite–scheelite, (2) early epithermal/arsenopyrite –stibnite–fluorite–marcasite, (3) late epithermal/pyrite–Fe–W–Sb oxides, (4) hypogene to supergene/Fe–Pb–Sb oxyhydroxides. Tungsten mineralization is likely to have formed at temperatures around 300 °C (Fig. 21.03). The mineralizing fluids changed from acidic/reducing to alkaline/oxidizing. The temperature of formation spans the full range from mesothermal to epithermal, thereby demonstrating that any classification scheme using temperature-related adjectives such as epithermal or mesothermal are often unworkable in practice and sometimes may cause confusion.

At a greater distance from the felsic intrusives Sb and W were found associated with Hg. The mineral assemblage of the Gumusler deposit, Turkey, includes scheelite, barite, stibnite, gold (up to 37.3 ppm) together with cinnabarite and Sb-sulphosalts. The Sb–Hg–W and Ba–Sb ore bodies are fracture bound with disseminations and veinlets of Hg– Sb in brecciated zones. The mineralization is genetically related to postmagmatic fluids associated with the Cenomanian granitic magmatism (Akçay, 1995). Many of these Sb–Hg–W mineralizations in Turkey were the keystone for a schematic model on timebound ore deposition throughout the early Paleozoic which was applied on a worldwide scale to volcanic-sedimentary series bearing scheelite, stibnite and cinnabarite (Höll, 1966; Maucher, 1974). After a “hype” in finding timebound Sb–Hg–W mineralizations, it has today become a more silent and this element association become more silent marginalized. 20d CD: Heading south-east, the SE Asian Sb belt described in the previous section ends up in the subvolcanic-related Sb–(Au–Hg–As) skarn deposits of Sarawak, Malaysia (Sarawak-type) (Dill and Horn, 1996) (Fig. 21.04). At temperatures of greater than 400 °C, dacites and microgranodioritic porphyries of Miocene age generated wollastonite, diopside and epidote in upper Jurassic massive Bau Limestone (Hon, 1981). When temperature dropped to approximately 369 °C sarabauite and gold precipitated followed at a further drop of temperature down to 194 °C by an early generation of stibnite, senarmontite and calcite. Native arsenic, realgar and orpiment formed below 194 °C. Taken Fe and S limitation of the mineralizing fluids to the extreme produces an ore mineralizations almost exclusively consisting of native elements as exemplified by an Sb mineralization discovered in veins and pockets in limestones of the Busan Hill, Borneo (Posewitz,

Fig. 21.03. Textures of antimony–tungsten ore mineralization. a) Vein mineralization of massive stibnite with valentinite, roméite and stibiconite from Pha Had Mine, Thailand. b) Massive Sb ore of stout crystals of stibnite at Mae Ta Mine, Thailand. c) Sender crystals of stibnite in a brecciated and pervasively silicified hornfels with fluorite and quartz growing into vugs, Doi Ngom Mine, Thailand. d) Veinlets of wolframite and late-stage quartz, Doi Ngom Mine, Thailand.


191

Fig. 21.04. Epithermal and skarn mineralization at Lucky Hill–Au–Sb mineralization, Malaysia, with Fe- and S limitation. Cartoon illustrating the position of the ore body and reactions describing the evolution of the antimony ore deposit. After Dill and Horn (1996).

1889). In this mineralization native gold, native antimony, native arsenic and native copper are accompanied by some realgar, cinnabarite and stibnite. Skarn mineralizations with subeconomic As-, Sb-, Hg- and Au mineralizations similar to that in Sarawak were found within polymetamorphic complexes south of the Tauern Window at Siflitz, Stockenboi and Rabant in Austria. These Alpine mineralizations are linked to the Oligocene magmatic activity along the Periadriatic Lineament of the Alpine Mts. 20e CD: The Sb–Au mineralization of the skarn-type Sb–(Au–Hg–As) deposit passes into the disseminated Carlin- type Au–Ag deposits or sediment-hosted disseminated gold (SHDG)-type deposits (Percival et al., 1990, 1993; Arehart, 1996; Stenger et al., 1998; Bettles, 2002; Thompson, 2002) and Tl-deposits exemplified by the Alsar deposit, Macedonia— see succeeding section. Common among this style of mineralization is the submicron gold, arsenical pyrite, late stage stibnite and realgar, orpiment and cinnabarite the latter being present at the periphery of the deposit. This Sb mineralization contrasts with the monometallic–monomineralic Sb vein-type deposits elsewhere also by its dominance in native arsenic and the absence of arsenopyrite. 20b CD: The SE Asian setting is somewhat of a mirror image of the eastern Pacific rim with Sb mineralization to have formed behind the zone of late Cenozoic arc-related porphyry and epithermal Cu–Au deposits on a rather thick lithospheric crust (Dill, 1998a; Dill et al., 1997a) (Fig. 21.16). In the Bolivian Andes stibnite deposits evolved at Cosuño and Milluri in what may be called a high-sulphidation-type Sb deposit (Fig. 21.06) (Dill et al., 1997a). The Cosuño, Lipez and Milluri volcanic edifices consist of a wide spectrum of volcanic and pyroclastic rocks of dacitic composition which evolved during the Miocene and Pliocene along N–S and NNE–trending fault zones that acted as feeder channels for the Neogene magma and for the hydrothermal solutions which brought about the Sb mineralization. At Cosuño a younger sinter cap bearing opaline siliceous material covers the ore mineralization. Hydrothermal systems are still active and hot brines are discharging in the immediate surroundings of this

sinter cap. At Cosuño and Milluri stibnite, pyrite and marcasite are disseminated in aggregates of alunite, quartz and kaolinite. In both sites the Tertiary magmatic rocks are strongly silicified, alunitized and argillized, particularly in the upper parts where the alunite–stibnite mineralization developed in stockworks and small veinlets. Antimony mineralization in the Bolivian Andes cannot be discussed isolated from subvolcanic Sn and base–metal mineralization emplaced adjacent to these Sb mineralization. Evidence for this idea is furnished by the presence of complex Sn- and Pb sulfosalts within neighboring mineralization (Fig. 21.07). Chemical and mineralogical studies of the Slovak and Austrian Sb deposits demonstrate that they do not fulfill the criteria for a classification as acid sulfate-type stibnite deposits (Böhmer, 1980; Belocky et al., 1991). Alunite and kandite-group minerals are only rare constituents among the minerals of the Kremnica Au- Sb deposit. It may, thus, be denominated as low sulphidation- or quartz-adularia-type of Sb mineralization, using for classification the criteria put forward by Heald et al. (1987). 20c CD: Hot brine-related mineralization in southern Tuscany, Italy, is still going on at the edge of the geothermal fields of Lardarello, Amiata and Latera within calcareous country rocks leading to near-surface lowsulphidation-type Sb-(Hg) mineralization (Dessau et al., 1972; Latanzi, 1999). They contain pyrite, stibnite and invisible gold. The fluids span a wide range of temperature (132 to 245 °C). The onset of hydrothermal (meteoric) fluid circulation is presumed to be correlative with the Neogene volcanic activity (Fig. 21.08). Ore mineralizations in Tuscany are another example of ore deposition in a calcareous environment by Neogene subvolcanic rocks starting off with skarn-type mineralization (hedenbergite, ilvaite) and fading out in low-temperature mineralization akin to high- and low sulphidation-type mineralization (Dill, 1979b). The Sb mineralization of northern Chios, Greece, so far considered as stratabound by Höll (1966) within Upper Paleozoic sediments is fault-controlled and of hot-brine style (Skarpelis, 1999a,b). The clastic wall rock sediments are hydrothermally altered to quartzsericite–pyrite and their calcareous counterparts ankeritised. The ore

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Fig. 21.06. Diagrammatic sections through the APS-bearing volcanic-hosted stibnite deposits in central Bolivia and micrograph of stibnite mineralization. a) Acicular crystals of stibnite disseminated in a matrix of equigranular alunite. Fissures filled with white mica (subparallel streaks). b) Idealized cross section through the deposit at Milluri. c) Idealized cross section through the deposit at Cosuño.

mineral association comprises base metal sulfides, tetrahedrite, stibnite and berthierite. Precious metals Au and Ag are of low grade. The entire hot spring mineralization is related to the Mid-Miocene porphyritic rhyolites and akin to the epithermal hot-spring-type deposits encountered in western Turkey. In contrast to porphyry Cu–Au–Mo deposits, Sb deposits are distal relative to the subduction zone. They are confined to sections of the fold belt where the continental crust is thick and was subject to strong horizontal displacements. The ratio of horizontal to vertical crustal movements during the structural evolution of the fold belt and the resultant skewness of the geothermal gradient played a decisive role for the type of (sub)volcanic-related Sb deposits to evolve. 20a AB: Sb deposits related to (ultra)basic rocks and greenstone belts are quite distinct from the overall Sb ore mineralization by the low silica contents of the host rocks. Several Sb deposits occur in the Murchison range, a greenstone belt located in the Northern Province of South Africa approximately 400 km NE of Johannesburg and made this Precambrian terrain to one of the world's largest antimony producing areas (Abbot et al., 1986; Pearton and Voljoen, 1986). The Sb ore-bearing Weigel Formation is bounded on one side by the ultramafic to mafic volcanic-dominated Mulati Formation, and

on the other by the mafic to felsic volcanics of the Rubbervale Formation which is host of volcanogenic Cu–Zn mineralization. A linear shear or fault zone is defined by several deposits lined up over a length of around 50 km, (Antimony Line). The main ore minerals are stibnite, berthierite, gersdorffite, ullmannite, gersdorffite–ullmannite s.s.s., gudmundite, pyrite, arsenopyrite and scheelite. Native gold and aurostibite are found in the Sb ore. Volcanic activity was responsible for the primary ore mineralization; the final location of the mineralization along shear zones has been ascribed to polyphase remobilization and re-concentration. The close affinity of Sb to (ultra)basic magmatic lithologies and the final emplacement along shear zones has had some implications on the primary source of Sb which is held to be in the mantle from which it was remobilized along deep-seated lineamentary shear zones. When establishing the so-called Hg–Sb–W formation, Maucher (1965, 1974), who re-classified numerous cinnabarite, stibnite and scheelite mineralization as stratabound and timebound to basic volcanic rocks of early Paleozoic age, gave a direct reference to this Precambrian Sb province in South Africa. Modern oceanic/ mantle-derived fluids and their predecessors in the geological past are cast in the role of source of Sb. In the wake of this new wave created by Maucher (1965) many Sb mineralizations in almost every

Fig. 21.05. Comparison of Sb–Au deposits at the western and eastern margin of the Pacific Ocean and location of major Sb–Au and Cu–Au metallogenic provinces in SE Asia and the Andes with shearzone-hosted and acid–sulfide-type Sb mineralization (Dill et al., 2008d).

194


Fig. 21.07. Complex sulfidic and non-sulfidic vein-type Sb ore. a) Zinckenite, Mina San Francisco near Poopo, Bolivia. b) Seginite, arseniosiderite, stibiconite, senarmontite at La Angustora—Sonora, Mexico. c) Cervantite and roméite at Pha Joe Area—Kachenaburi, Thailand. d) Roméite, stibnite and cinnabarite, Pha Joe Area—Kachenaburi, Thailand.

Fig. 21.08. Epithermal polymetallic Sb–(Au–Hg) deposits. a) Sb deposit within the “Calcare cavernoso” at Cetine—Rosia, Italy. b) Cross section of the Sb deposit Manciano-Tafone, Italy (Arisi-Rota et al., 1971). c) Hot brine pool at Monterodonto Marittima, Italy. d) Adularia replaced by quartz near Fenice Capanne, Italy (Dill, 1979) low-sulphidation type. e) Alunite replaced by pyrite near Fenice Capanne, Italy (Dill, 1979) high-sulphidation type.


195

Fig. 21.09. Mineralized shear zones in the Bolivian Andes and the Western Carpathians. a) Ladder-shaped Sb-bearing ore shoots in the Churquini Sb deposit, Bolivia (modified after Ahlfeld and Schneider-Scherbina (1964)). b) Bedding-concordant Sb-bearing veins in the hinge zones of anticlinal structures (“saddle reefs”) in the Candelaria Sb deposit, Bolivia (modified after Lehrberger (1992)). c) Flat-lying mineralized Sb-bearing fault zones in the Dúbrava deposit, Slovakia (modified after Pouba and Ilavský (1986)). d) Graphite- and stibnite-bearing shearzone in the Pezinok deposit, Slovakia (modified after Chovan et al., (1992)).

country found adjacent to basic volcanic rocks was held to be stratabound and timebound (Lahusen, 1972; Höll and Maucher, 1976; Reimann and Stumpfl, 1981). Today the pendulum is swinging towards the opposite direction as exemplified by the debate on the age of the Felbertal deposit (Raith and Stein, 2006). In the near future, when the euphoria on Re–Os has died down and deposits will be looked at from different perspectives we will certainly meet somewhere in between.

21.2.2. Structure-bound antimony deposits Shear zone-hosted, mesothermal (Au-)Sb veins are subdivided into monotonous Sb deposits (20a G) and polymetallic Sb–(Au–W– Sn–As–Zn–Pb) deposits (20b G). Structure-bound antimony deposits reveal a significant difference in their mineral assemblages. One group may be denominated as monomineralic–monometallic Sb vein-type deposits. Its mineralogy is rather poor and made up of stibnite, arsenopyrite with little Pb–Cu–Zn minerals in a quartzose matrix and

196


a low Au content (Figs. 21.09, 21.10). This type of vein mineralization occurs in shear zones developing in non-metamorphosed to verylow-grade regionally metamorphosed country rocks at low temperatures. Monotonous vein-type Sb mineralization is found in a more distal position relative to the heat zone than the second group of veintype Sb mineralization (Bachmann, 1955; Dill, 1985c; Gumiel and Arribas, 1987; Aral, 1989; Berger, 1993; Fomino, 1994; Dill et al., 1995b; Neiva et al., 2006). The second group of vein-type Sb mineralization may be called polymetallic Sb–(Au) vein-type mineralization. Examples of this type were found in sheared metapelites near or in the roof rocks of granites during the collisional stage of orogenies such as Brandholz–Goldkronach, Germany, Durico–Beirao, Portugal, Brioude–Massiac–Massif Central, France, Dúbrava, Magurka, Medzibrod, Lom, Dve Vody, Slovakia, and western Thailand (Bril, 1982; Dill, 1985c, Chovan et al., 1992; Guillemette and WilliamsJones, 1993; Ortega and Vindel, 1995; Wagner and Cook, 2000). In addition to the minerals recorded from the monotonous Sb vein mineralization, these veins are also enriched in complex Pb–Sb sulfosalts. In places, these veins are abundant in wolframite and/or scheelite concentrated during an initial mineralizing phase and in visible and invisible gold sensu Cook and Chryssoulis (1990).Many of the shear-zone-hosted stibnite deposits are related to the fold dynamic of the host orogen (Figs. 21.11, 21.12). Variscan folding and shearing and anomalous crustal heat flow provided a sustainable long-lived hydrothermal activity. These veins closely resemble, lowsulfide gold-quartz orogenic or turbidite-hosted Au (Sb) lodes in late Proterozoic and Paleozoic sequences of back-arc basins, which are sediment-dominated and lack thick volcanic sequences (Hutchinson, 1987; Madu et al., 1990; Dill et al., 1997a; Dill, 1998a). Diluted CO2enriched fluids were generated by metamorphic dehydration and zones of structural weakness channeled the hydrothermal fluids during regional deformation. Recently, a model for this type of vein mineralization in Central Europe was proposed by Boiron et al. (2001), involving the mixing of metamorphic and superficial fluids in the upper crust and hydrocarbon migration. The lead isotopic data for stibnite indicate a homogeneous source of lead of crustal origin, from a dominant metasedimentary source, for the studied antimony, leadantimony and antimony–gold quartz veins in northern Portugal (Neiva et al., 2006). 21.2.3. Sedimentary antimony deposits 20a J: Stratabound sedimentary-diagenetic Sb deposits have a prime example at Xikuangshan in China. It is one of the world's biggest Sb resources emplaced by sedimentary and diagenetic processes. Weak epigenetic remobilization may have modified what is primarily a sedimentary-diagenetic Sb deposit bound to Devonian black shales and cherts (Wu, 1993; Fan et al., 2004) (Fig. 21.13). Mineralization occurs in veinlets, vugs, and in different layers of chert. Stibnite is the only ore mineral with trace amounts of pyrite, pyrrhotite, sphalerite in a primary gangue mineralization of quartz and calcite followed by secondary barite and fluorite. 20b J: Gold-bearing stibnite and cinnabarite deposits at Emirli and Halıköy, Turkey, are stratabound to graphite- and mica-schist as well as structurally controlled by veins. Part of this mineralization is attributed to the epithermal type of Sb–Au mineralization (Akçay et al. 2003, 2006). Stibnite and cinnabarite are associated with pyrite, arsenopyrite, gold, base metal sulfides, marcasite, realgar orpiment and cinnabarite which belong to a mineralization event starting at temperatures of more than 300 °C. Deposition of pyrite and arsenopyrite is probably due to cooling of the fluids and sulphidation. Sulfur in stibnite originated from sulfide minerals in the country rocks (bacterial reduction of sulfate within the pre-metamorphic sediments). Cinnabarite has heavier δ34S contents as a result of a magmatic sulfur contribution. 20c J: Polymetallic Sb mineralization in black shales or alum shales falls into the category of low-grade–large tonnage deposits with only

a few examples upgraded to a giant deposit such as at Sukhoi Log, Russia under favorable circumstances (Distler et al., 2004). Black shales such as the Paleozoic Graptolite Shales, Alum Shales and Chattanooga Shales have been known for a long time to contain a wide range of metals (Bates and Strahl, 1957). The various elements including Sb were accumulated in anoxic environments by sorption and precipitation from seawater, fossil brines or hydrothermal solutions. Carbonaceous sediments in passive margin basins contain only subeconomic concentrations of V, Mo, base metals, Au and Sb (Kříbek, 1991). Only the uraniferous black shales have attracted the attention of mining geologists and petroleum geologists alike, that is why these sediment-hosted ore types are discussed later in context with uranium (Dill, 1986a, b; Lüning and Kolonic, 2003). 20a N: Although not mined for Sb, anthracite may among other elements such as Tl, As, and Hg also contain elevated contents of Sb as shown from the late Permian anthracite at Xingren, Guizhou, southwest China (Dai et al., 2006). 21.3. Arsenic deposits Arsenic is concentrated in stratiform Cu- PGE–Ni–deposits in the basal zone of (ultra)basic complexes (21a A), porphyry Cu deposits including co-magmatic breccia pipes (21a D) and Sn Skarn deposits (21b CD) (Table 21.02). Arsenic is extracted mainly from Sn–W-, Cu- and Au ores. Porphyry copper deposits contain enargite as a major host of Cu and As (Butte, USA, Chuquicamata, Chile, Morococha, Peru, Lepanto, Philippines). Massive Cu sulfide deposits, e.g. Kidd Creek, Canada, and Cu–Ni sulfide deposits, Sudbury, Canada, and Noril'sk, Russia, also contain minerals rich in As. In gold deposits it is mainly arsenopyrite that is host to As. In many deposits, arsenic causes more than a headache where to dispose it as a waste material than to sell it as a raw material. 21a G: Only the Ambrolauri region, Georgia, warrants mentioning for its unconformity-related (?) vein-type deposits abundant in orpiment, realgar and arsenopyrite (Mindat database). 21b G: Arsenic deposits are known from the Sudetes in Poland (Fig. 21.14) (Osika, 1990). The arsenic–pyritic minerals assemblage contain besides pyrite, arsenopyrite minor amounts of stibnite, chalcocite, galena and sphalerite. The Złoty Stok As–Au deposit in Proterozoic series of schists, amphibolites, marbles and dolomitic limestones is still under discussion (Mochnacka et al., 1995) with the most recent interpretation suggesting a hydrothermal–metasomatic origin. Au is concentrated in loellingite with an average ore grade of between 3 and 5 ppm Au. Until its closure in 1961 this mine was the largest producer of primary Au in the Western Sudetes. It has many similarities with the orogenic gold deposits as far as structure and mineral assemblages are concerned. 21a N: Enrichment of As in coal is significant (Yudovich and Ketris, 2005a; Seredin and Finkelman, 2008). Arsenic is a very coalphile element. The relations are controlled by presence of As– Aspyr or Asorg. Four genetic types of As accumulation on coal are known: –Chinese type–hydrothermal As enrichment, similar to known Carlin type of As-bearing telethermal gold deposits. –Dakota type–hypogene enrichment from ground waters draining As-bearing tufa host rocks. –Bulgarian type–As enrichment resulting from As-bearing waters entered coal-forming peat bogs from sulfide deposit aureoles. –Turkish type–volcanic input of As in coal-forming peat bog as exhalations, brines and volcanic ash. 21b K: The polymetallic sulfide mineralization Lengenbach/Binntal in Switzerland is discussed as one of the strangest “deposits” on earth. It was exclusively exploited for private and public mineral collections or to investigate certain aspects of Pb–Cu–Ag–As–Tl sulphosalts that were almost exclusively encountered in this mineralization (Cannon,


197

Fig. 21.10. Vein-type Sb ore mineralization in the Bolivian Andes. a) Vein mineralization of massive stibnite, Mina Kharma. b) Veinlets of stibnite penetrating vein quartz, Mina Kharma. c) Quartzose Sb mineralization at outcrop stained with Sb ochre near Mina Caracota.

2000). The latest summary and most complete report on the mineral wealth of the Lengenbach Mine were published by Baumgärtl and Burow (2005). This “deposit” may contribute significantly to the understanding of the Alpine metallogenesis in the western Alps (Hofmann, 1994). The protore mineralization consisted of pyrite, galena and sphalerite which were deposited in Triassic dolomites. Sulfosalts containing predominantly As and Tl formed in a closed system with temperatures of up to 450 °C during regional metamorphism at around 18.5 ± 0.5 Ma. Thus, the minerals underwent significant changes and the complexity of the mineral assemblage

Fig. 21.11. Cartoon to demonstrate synmetamorphical and syntectonical Sb–(Au) mineralization in shearzones along the axial surface of an anticline. Similar rock mechanical processes were also operative during formation of some orogenic gold veins at shallow tectonic level.

Fig. 21.12. Thrust-bound metamorphogenic and/or fold-related deposits, e.g. Berga Anticline, Germany. Antimony mineralization may in places run subparallel to the bedding planes close to the fold hinge of the recumbent fold.

198


Fig. 21.13. Stratabound sedimentary-diagenetic Sb deposits Xikuangshan-Hunan, China (Fan et al., 2004).

may be used to measure the degree of Alpine remobilization. The element budget may, however, still reflects the parent material and it may be used for “minero-stratigraphic correlations” between deposits of the Alps and deposits in the extra-Alpine realm. This mineralization can be taken as the missing link between the As–Au mineralization related to the metacarbonates in the Variscides of the Polish Sudetes

and the Carlin-type deposits rich in As, Tl and Au mentioned below (Fig. 21.14). 21.4. Thallium deposits The economic situation of thallium is not at a variance with that of arsenic. Both elements are strongly toxic and their industrial use is limited. Therefore only a few mineralizations may be called Tl deposits such as the disseminated Au–As–Sb–Hg–Tl deposit Alsar, Macedonia (Percival and Radtke, 1994) (22 CD). Base–metal and silver concentrations are characteristically low, as is the case with the mineralization in the type locality in the Carlin Trend. The host rocks consist of Triassic carbonate rocks interbedded with volcanic tuff and dolomite of Tertiary age. The Au–As–Sb–Tl–Hg mineralization is Pliocene in age and its formation driven by a magmatic heat source. The chemical changes and facies variation within the mineral assemblages may be deduced from the cross section of Fig. 21.15. In addition to marcasite and pyrite, realgar, orpiment, stibnite, cinnabarite, lorandite and various Tl-bearing sulphosalt minerals precipitated. A similar type of deposit was discovered at Zalaa Uul, Mongolia (Cluer et al., 2000). The Xiangquan thallium deposit, eastern China, is exceptional, because it is the world's first thallium-only mine (Zhou et al., 2005). Thallium present as lorandite and hutchinsonite was originally precipitated from hydrothermal emanations onto the seafloor and deposited in Lower Ordovician calcareous sediments. Subsequently in the Early Cretaceous these source beds were reworked to form thallium enriched minerals in hydrothermal veins. By and large, the mechanism of discharge are not very much different from those of the more complex Alsar deposit.

Table 21.02 Classification scheme of arsenic and thallium deposits.

Fig. 21.14. Structures and textures of arsenic ore. a) Cross section through the goldbearing arsenic deposit Złoty Stok, Poland, an orogenic deep-seated deposit in the Sudetes Mountains (Osika, 1990): 1) blastomylonitic mica schists, 2) quartz plagioclase gneiss, 3) dolomitic marble, 4) metasomatic As–Au ore, 5) fault zones. b) Native arsenic (“Scherbenkobalt”) from the vein-type deposit St. Andreasberg, Harz Mountains, Germany.

(1) Magmatic arsenic and thallium deposits (1) As in stratiform Cu–PGE–Ni-deposits in the basal zone of (ultra)basic complexes (21a A) (2) As in porphyry Cu deposits including co-magmatic breccia pipes (21a D) (3) As in Sn skarn deposits (21b CD) (4) Tl in Carlin-type deposits (22 CD) (2) Structure-related arsenic deposits (1) Unconformity-related realgar and orpiment deposits (21a G) (2) Orogenic As–Au veins (21b G) (3) Sedimentary arsenic and thallium deposits (1) As in coal (21a N) (2) Tl in coal (22N) (4) Metamorphic arsenic and thallium deposits (1) As–Tl–Pb–Cu–Ag deposits in dolomitic metacarbonates (21b K)


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216


Fig. 24.10. A compilation of sediment-hosted U deposits illustrating the environment of syndiagenetic calcrete and lacustrine as well as epigenetic sandstone-hosted “roll-front”, tabular and stratal U mineralization (modified from Galloway and Hobday (1996)). Arrowheads show the direction of fluid flow.

25a I: Roll front deposits are characterized by arcuate U ore bodies that crosscut sandstone bedding, e.g., Wyoming, Texas, USA (Goldhaber et al., 1978) (Fig. 24.12). There are also some important U deposits of this type outside the USA Moinkum, Inkai and Mynkuduk, Kazakhstan, and Bukinay, Sugraly and Uchkuduk, Uzbekistan. In the unaltered sulfidized part o the rock FeS2, calcite, organic matter (OM) and Fe–Mg silicate were preserved and made visible by the gray color. In the alteration zone Fe2O3 and FeOOH formed instead of Fe sulfides, metals were leached out, OM were destroyed, Fe–Mg silicates altered and calcite removed. The rock color turns into red, brown and yellow. In the ore zone of the roll front FeS2 is abundant together with U, Mo, Se, V, Co, Cu and As, OM and calcite increased and Fe–Mg silicate are stable. Pitchblende and coffinite are commonly present as U ore minerals while V, Mo, Se, Cu and As may be recovered as by-product. Too much of calcareous gangue minerals is detrimental during ore processing. Vanadium may be accommodated within montroseite, carnotite, tyuyamunite and francevillite and molybdenum is present as jordisite and Se accommodated in the lattice of ferroselite.

25b I: Tabular subtypes of sandstone-hosted U deposits with lenticular bodies parallel to the depositional trend occur in the Grants Mineral District and New Mexico, USA, in the Westmoreland area, Australia, as well as in Niger (Akouta, Imouraren and Arlit), where paleochannels incised into underlying basement rocks. The sedimentary U deposits at the western edge of the Aïr Massif, in the Tim Mersoi sub-basin at Arlit and in the Gall basin at Agades are hosted by Upper Carboniferous units, in places, containing coal beds. Similarly, Jurassic and Cretaceous sandstones beds host U and Cu. Uranium in these units is hosted within sandstone and conglomerate (Fig. 24.13) (Bigotte and Obellianne, 1968). 25c I: Tectonic/lithologic deposits also called “stack type” occur in sandstones adjacent to a permeable fault zone like Lake Ambrosia in New Mexico, USA or Mounana, Gabon. The Franceville uranium ore bodies in Gabon are hosted by the upper sections of the Paleoproterozoic Mabinga Sandstone (Dymkov et al., 1997; Mossman et al., 2005). The host sequence comprises fluvial and deltaic sediments with poorly sorted and conglomeratic sandstones some of which containing thorite and uranothorite in heavy mineral concentrations. The most common

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