Geochemical and Tectonic Evolution of Arc-Backarc Hydrothermal Systems: Implication for the Origin of Kuroko and Epithermal Vein-Type Mineralizations and ... Cycle (Developments in Geochemistry)

Developments in Geochemistry 8 Geochemical and Tectonic Evolution of lrc-Backarc Hydrothermal Systems Implication for ...

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Developments in Geochemistry 8

Geochemical and Tectonic Evolution of lrc-Backarc Hydrothermal Systems Implication for the Origin of Huroko and fpithermal Vein-Type lllineralizations and the Global Geochemical Cycle

Developments in Geochemistrg 1. W.S. Fyfe, N.J. Price and A.B. Thompson FLUIDS IN THE EARTH’S CRUST

2. P. Henderson (Editor) RARE EARTH ELEMENT GEOCHEMISTRY

3. B.A. Mamyrin and I.N. Tolstikhin HELIUM ISOTOPES IN NATURE 4. B.O. Mysen STRUCTURE AND PROPERTIES OF SILICATE MELTS

5. H.A. Das, A. Faanhof and H.A. van der Sloot RADIOANALYSIS IN GEOCHEMISTRY 6. J. Berthelin DIVERSITY OF ENVlRONMENTAL BIOGEOCHEMISTRY

7. L.W. Lake, S.L. Bryant and A.N. Araque-Martinez GEOCHEMISTRY AND FLUID FLOW

Developments in Geochemistry 8

Geochemical and Tectonic Evolution of flrc-Backarc Hydrothermal Systems Implication for the Origin of Huroko and fpithermal Vein-Type mineralizations and the Global Geochemical Cycle BY

Naotatsu Shikazono

2003 Elsevier

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Preface In and near the Japanese Islands many Neogene hydrothermal ore deposits have been formed from the middle Miocene to the present time, and many subaerial active geothermal systems occur. Some of them are associated with base-metal (Cu, Zn, Pb, Fe, Mn) and precious-metal (Au, Ag) mineralizations. Representative types of Neogene deposits include Kuroko and epithermal vein-type deposits. Kuroko deposits, which are strata-bound massive sulfide-sulfate deposits, are wellknown because (1) many studies have been done and many papers (more than 1,000) have been published since the work by Ohashi (1919), (2) original ore textures are preserved due to the absence of metamorphism, and (3) geological and physicochemical environments of ore deposition were well-elucidated. Summaries of previous studies on Kuroko deposits have been published in the 1970s and early 1980s (Ishihara, 1974; Ohmoto and Skinner, 1983). However, no summary written in English after the early 1980s has been published, although considerable works on ore deposits have been carried out. Epithermal vein-type deposits in Japan have also been well-studied. More than 1,000 papers (mostly in Japanese) have been published. However, a general overview of the ore deposits is not yet available in English. Previous work on Kuroko and epithermal vein-type deposits in Japan will be summarized in Chapter 1. The descriptions of individual vein-type and Kuroko deposits are not covered in this book; they can be found in the references listed at the end of each chapter. By integrating geological (e.g., distribution of ore deposits, age of ore formation, host and country rocks, associated volcanic activity, tectonics, paleogeography), mineralogical (opaque, gangue, and hydrothermal alteration minerals), and geochemical (fluid inclusions, stable and radiogenic isotopes, minor elements of ore and country rocks, thermochemical calculations) data on the two types of deposits, the genesis, depositional mechanism and origin of ore deposits are described and discussed in Chapter 1. Temporal and spatial relationships between the two types of deposits and the evolution of tectonics and hydrothermal systems associated with the mineralization during the Neogene age in and around the Japanese Islands are considered. During the last three decades, subaerial geothermal areas in the Japanese Islands have been explored considerably and geothermal energy plants were developed. It was recognized that some active geothermal systems are accompanied by present-day basemetal and precious-metal mineralizations. In 1990s, hydrothermal venting and mineralization were discovered on the sea floor of the back-arc basin, back-arc rift, and island arc surrounding the Japanese Islands as well as other western Pacific regions.

vi

Preface

In Chapter 2, a geochemical, geological and mineralogical summary of active subaerial and submarine back-arc basin hydrothermal systems and mineralizations is given. The characteristic features of above-fossil and active subaerial and submarine hydrothermal systems are compared with fossil hydrothermal systems (epithermal veintype and Kuroko deposits), and the causes for the differences in the characteristic features are considered. Characteristic features of Paleozoic-Mesozoic volcanogenic stratiform Cu deposits (Besshi-type deposits) are compared with those of midoceanic ridge deposits and Kuroko deposits. In Chapter 3, hydrothermal and volcanic gas fluxes from submarine back-arc basins and island arc are estimated. These fluxes are compared with midoceanic ridge hydrothermal fluxes. Particularly, hydrothermal flux of CO2 is considered and the influences of this flux on global long-term carbon cycle and climate change in TertiaryQuaternary ages are discussed in Chapter 4.

Acknowledgements Several acknowledgements are in order. I am very much indebted to the late Professors Emeriti T. Tatsumi of the University of Tokyo, advisor of my Ph.D. thesis, and T. Watanabe of the University of Tokyo for their valuable advice to study epithermal vein-type and Kuroko deposits in Japan. I learned from them an importance of the integration of geochemistry, geology and mineralogy for studying the genesis of hydrothermal ore deposits. I acknowledge the late Professors T. Fujii of Tsukuba University and A. Tsusue of Kumamoto University who taught me applications of thermodynamics, kinetics and hydrodynamics to ore genesis. Professors T. Nakamura, K. Nagasawa and S. Takenouchi's papers on the detailed mineralogical and fluid inclusion studies of vein-type deposits in Japan were especially valid and useful to writing this book. Discussions with Professors Dick Holland, Ulrich Petersen, Ei Horikoshi, Hiroshi Ohmoto, Clif Farrel and Udo Fehn on the genesis of Kuroko deposits during my stay at Harvard University as a Post-Doctoral Fellow (1979-1981) contributed very much to my research on Kuroko deposits. I appreciated Dick Holland's hospitality while I stayed at Harvard University during my 1997-1998 sabbatical year (during which parts of the draft of this book were written) from Keio University. Keio University provided a grant for my one-year stay at Harvard. This volume is indebted a great deal to many people of the Geology Department of the University of Tokyo, the Applied Chemistry Department of Keio University, and Geology Department of Tokyo Gakugei University. I particularly would like to mention Drs. J.T. Iiyama, E. Hirokoshi, M. Utada, T. Sato, H. Shimazaki, Y. Kajiwara, M. Watanabe, J. Date, K. Kase, S. Doi, T. Urabe, T. Mizuta, K. Hattori, M. Aoki, R. Kouda, M. Shimizu, K. Takeuchi, E. Uchida, Y. Shibue, N. Takeno, Y. Morishita, M. Tamura, S. Nakashima, H. Kawahata, K. Fujimoto, Y. Kato, A. Imai, T. Nagayama, O. Ishizuka, M. Hoshino, S. Kimura, Y. Ishikawa, H. Kashiwagi, Y. Ogawa, H. Honma and M. Nakata. I very much appreciate Miss M. Aizawa for her skillful and patient word processing.

Preface

vii

Finally, I would like to dedicate the b o o k to m y wife, Midori Shikazono, daughters, Chikako and Hisako Shikazono, and parents, N a o h a r u and Yoshiko Shikazono, for their moral support of m y academic research. Naotatsu Shikazono Keio University

References Ishihara, S. (ed.) (1974) Geology of the Kuroko Deposits. Mining Geology Special Issue, 6, 437 pp. Ohashi, R. (1919) On the origin of Kuroko of the Kosaka mine. J. Geol. Soc. Japan, 26, 107-132 (in Japanese). Ohmoto, H. and Skinner, B.J. (eds.) (1983) The Kuroko and Related Volcanogenic Massive Sulfide Deposits. Econ. Geol. Mon., 5, 604 pp.

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Table of Contents Preface ......................................................................... Chapter 1 . Miocene-Pliocene Hydrothermal Ore Deposits in and around the Japanese Islands ................................ ............ 1.1. General overview of metallogeny and tectonics in the Japanese Islands ....... 1.2. General overview and classification of hydrothermal ore deposits of Neogene age ............................................... ................ 1.3. Kuroko deposits ...................... ............................ 1.3.1. Geological characteristics ........................................ 1.3.1.1. Distribution ................................................ 1.3.1.2. General geology, country rocks and tectonic setting . . . . . . . . . . . 1.3.1.3. Age of mineralization ....................................... 1.3.1.4. Metals enriched and metal ratios ............................. 1.3.2. Mineralogical characteristics .......... ............................ 1.3.2.1. Metal zoning, and ore and gan minerals .................. 1.3.2.2. Hydrothermal alteration ..................................... 1.3.3. Geochemical characteristics ......................... 1.3.3.1. Fluid inclusions ............................................. 1.3.3.2. Gas fugacities . . . . . . . . .......................... 1.3.3.3. Chemical compositions 1.3.3.4. Stable isotopes ........ ......................... 1.3.3.5. Radiogenic isotopes ............................. 1.3.4. Depositional mechanism and origin of ore fluids ...................... 1.3.4.1. Depositional mechanism ......................... ...................................... 1.3.4.2. Origin of ore flui 1.4. Epithermal vein-type deposits ........... ................................. 1.4.1. Geological characteristics ...................................... 1.4.1.1. Distribution ............................... ............. 1.4.1.2. Age of mineralization . . . . .......................... 1.4.1.3. Volcanic activity related to mineralization .................... 1.4.1.4. Metal enriched and metal ratios .............................. 1.4.2. Mineralogical characteristics ......................................... 1.4.2.1. Metal zoning ..... ................................... 1.4.2.2. Ore minerals . . . . . . ...................................... 1.4.2.3. Gangue minerals ............................................ I .4.2.4. Hydrothermal alteration zoning ..............................

V

1 1

6 15 15 15 15 19 20 23 23 30 38 39 41 48 51 54 61 61 77 83 84 84 84 87

88 88 88 88 94 98

X

Table of Contents

1.4.2.5. Spatial and geochemical relationships between propylitic alteration and advanced argillic alteration: a case study on the Seigoshi-Ugusu district. central Japan ....................... 1.4.2.6. Chemical composition of alteration minerals . . . . . . . . . . . . . . . . . 1.4.2.7. Causes for hydrothermal alteration ........................... 1.4.3. Geochemical characteristics .......................................... 1.4.3.1. Fluid inclusions ............................................. 1.4.3.2. Estimate of temperatures from the electrum-sphalerite-pyriteargentite assemblage ........................................ 1.4.3.3. Gas fugacities .............................................. 1.4.3.4. Chemical composition of ore fluids .......................... 1.4.3.5. Stable isotopes .............................................. 1.4.3.6. Lead isotopes ............................................... 1.4.3.7. Rare earth elements (REE) .................................. 1.4.4. Se- and Te-type Au-Ag deposits ..................................... 1.4.5. Depositional mechanism and origin of ore fluids ...................... 1.4.5.1. Depositional mechanism . . . . . . . . . . . . . .................... 1.4.5.2. Origin of ore fluids .......................................... 1.4.6. Hishikari deposit: an example of Japanese epithermal Au-Ag vein-type deposits ............................................................. 1.4.6.1. Geology and vein system .................................... 1.4.6.2. Hydrothermal alteration ..................................... 1.4.6.3. Mineralogy ................................................. 1.4.6.4. Geochemical features ....................................... 1.4.6.5. Interpretation of Si02 mineral zoning in terms of kineticsfluid flow-mixing model ..... ............................. 1.4.6.6. Gold precipitation due to mixing of fluids in epithermal system 1.5. Evolution of tectonics and hydrothermal system associated with epithermal and Kuroko mineralizations ................................................. 1.5.1. Paleogeography and stress field ...................................... 1S.2. Volcanic activity ..................................................... 1S.3. Tectonic influence on temporal and spatial relationships in Kuroko and vein-type deposits in southern Hokkaido, Japan ....................... 1S.4. Geochemical features of sedimentary rocks formed in the Japan Sea as a proxy for hydrothermal activity ..................................... 1S . 5 . Mode of subduction and formation of back-arc basin .................. 1.6. Other hydrothermal ore deposits ................................ 1.6.1. Polymetallic vein-type deposits ...................................... 1.6.1.1. Ashio deposit ............................................... 1.6.1.2. Tsugu deposit .............................................. 1.6.1.3. Kishu deposit ............................................... 1.6.1.4. Obira deposit ...............................................

100 113 122 124 124 124 129 141 143 158 158 159 170 170 176 183 184 186 186 187 196 199

201 202 204 206 213 225 231 231 234 240 240

Table of Contents

xi

1.6.1.5. Temperature and sulfur fugacity estimated from iron and zinc partitioning between coexisting stannite and sphalerite and coexisting stannoidite and sphalerite ......................... 1.6.2. Hg and Sb deposits .................................................. 1.6.3. Gold-quartz vein-type deposits (mesothermal-hypothermal vein-type deposits) ............................................................ 1.6.3.1. Geology, mineralogy and geochemistry ...................... 1.6.3.2. Gold-quartz vein-type deposits in Yamizo Mountains, central Japan ....................................................... I .6.4. Hot spring-type gold deposits ........................................ Chapter 2 . Present-day Mineralization and Geothermal Systems in and around the Japanese Islands . . . . . . . . . . . . . . .............................. 2.1. Subaerial geothermal system and min ............................. 2.1.1. Chemical compositions of geothermal waters controlled by hydrothermal alteration mineral assemblage .................................... 2.1.2. Na-K-Ca geothermometer ........................................... 2.1.3. Present-day mineralization in subaerial geothermal areas in Japan ...... 2.1.3.1. Nigorikawa ...... ................................... 2.1.3.2. Osorezan ................................................... 2.1.3.3. Okuaizu . . . . . . . . . ................................ .... 2.1.3.4. Sumikawa ....... ........................................ 2.1.3.5. Arima hot springs ........................................... 2.1.3.6. Beppu hot springs . . . . .............................. 2.1.3.7. Fushime ............. .............................. 2.2. Comparison of active geothermal systems with epithermal vein-type deposits . 2.2.1. Distribution . ..................................

...............................................

.................................. .................................... 2.2.3.2. Gangue and a nerals ..............................

2.2.4. Geochemical features of hydrothermal fluids .......................... 2.2.4.1. Gas fugacities ................................. 2.2.5. Geological and tectonic environment and volcanism . . . . . . 2.3. Submarine geothermal systems and associated mineralization ................ 2.3.1. Submarine metal precipitation at back-arc basins around the Japanese islands .............................................................. 2.3. I .1. Okinawa Trough ......................... ............. 2.3. I .2. Izu-Bonin Arc ........................... ............. 2.3.2. Characteristics of back-arc deposits in the Western Pacific . . . . . . . . . . . . . 2.3.2.1. Tectonic settings, geologic structure and volcanic rocks 2.3 2 . 2 . Metal contents . . . . . . . . . . . . . . . . ......................... 2.3.2.3. Mineralogy ................................................. 2.3.2.4. Chemical and isotopic compositions of hydrothermal solution .

241 247 249 249 258 261 295 295 295 302 311 31 1 312 315 320 321 323 324 324 324 327 327 328 333 333 333 334 335 336 337 337

xii

Table of Contents

2.3.3. Comparison of present-day back-arc deposits with Kuroko deposits .... 2.3.4. Spatial relationship between back-arc deposits and epithermal gold deposits ............................................................. 2.4. Comparison of back-arc deposits with midoceanic ridge deposits ............. 2.4.1. Hydrothermal solution ............................................... 2.4.1.1. Chemical compositions ..................................... 2.4.1.2. Isotope data (6"0, 8D, 634S, 613C, 87Sr/86Sr,3He/4He, 6"B, ' ....................................................... 84 L1) 2.4.2. Metal ratios and mineralogy ......................................... 2.4.3. Mechanism of formation of chimney and ore deposits ................. 2.4.3.1. Zonation and sequence of mineral precipitation ............... 2.4.3.2. Mineral composition ........................................ 2.4.3.3. Ore texture ................................................. 2.4.3.4. Grainsize .................................................. 2.4.3.5. Sulfur isotope data .......................................... 2.4.3.6. Mineral particle behavior in hydrothermal plumes ............ 2.4.3.7. Model for the formation of sulfate-sulfide chimneys and massive deposits on the seafloor ................................. 2.4.4. Hydrothermal alteration .............................................. 2.5. Besshi-type deposits in comparison with Kuroko deposits and midoceanic ridge deposits ................................... ........................ 2.5.1. General features and classification......... ........................ 2.5.2. Geological characteristics ............................................ 2.5.2.1. Distribution ................................................ 2.5.2.2. Age of formation of ore deposits ............................. 2.5.2.3. Host rocks and tectonics ......................... 2.5.3. Metamorphism and hydrothermal alteration ........................... 2.5.4. Mineralogical characteristics ......................................... 2.5.4.1. Opaque and gangue minerals ................................ 2.5.4.2. Ore texture ................................................. 2.5.5. Geochemical features ................................................ 2.5.5.1. Sulfur isotopes .............................................. 2.5.5.2. Metal ratios ................................................ 2.5.5.3, Se/S of sulfide ore .......................................... 2.5.5.4. Co and Ni of sulfide ore ..................................... 2.5.5.5. Goldinore ................................................. 2.5.5.6. Lead isotopes ............................................... 2.5.5.1. Rb/Sr and Nd/Sm isotopic compositions .................... 2.5.5.8. Geochemical environment of ore deposition ..................

350 350 354 354 354 359 361 366 361 368 368 368 369 369 370 371 373 373 375 375 375 378 379 379 382 383 383 385 390 390 391 392 393 394

Chapter 3 . Hydrothermal Flux from Back Arc Basin and Island Arc and Global 407 Geochemical Cycle .................................................. 3.1. Major element (alkali. alkali earth. silica) flux ............................... 407 413 3.2. Volatile element (COz. S. As) flux ..........................................

Table of Contents

.................... ........................... 3.2.2. Causes for high COz concentration and origin of C02 of hydrothermal solution from back-arc basins ........................................ 3.2.3. S flux .......................... ......................... 3.2.4. As flux . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.3. Other elemental flux ........................................................ 3.3.1. Hgflux ............................................................. 3.3.2. Mn flux ............................................................. 3.3.3. Ba flux ........................ ................................. 3.4. Comparison of back-arc hydrothermal flux with midoceanic ridge hydrotherma1 flux . . . . . . . . . . . . . . . . . ............................................

X ... lll

413 417 420 421 423 423 424 424 424

Chapter 4 . Influence of Hydrothermal CO2 Flux on Tertiary Climate Change . . . . . . 431 4.1. Tertiary climate change in relation to CO2 flux by volcanic. hydrothermal and metamorphic activities ..................................................... 431 4.2. Computation on global long-term carbon cycle and climate change ........... 439 Chapter 5 . Summary ............................................................

449

Subject Index ...................................................................

453

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Chapter 1 Miocene-Pliocene Hydrothermal Ore Deposits in and around the Japanese Islands

1.1. General overview of metallogeny and tectonics in the Japanese Islands Before mentioning the characteristics of Kuroko and epithermal vein-type deposits in Japan, it is worthwhile to briefly describe the metallogeny, geology, geophysics, and tectonic situations of the Japanese Islands. Japan is situated along the boundary between Eurasia and the Pacific plate (Fig. 1.1). The metallogeny, geology, geophysics and geochemistry of Japan are largely controlled by this tectonic setting. The great variety of mineral deposits of Japan reflects the complex geotectonic environments. An intimate relationship exists between igneous and hydrothermal activity, which in turn reflects the plate tectonic history of Japan. Many Japanese ore deposits have produced many different metals, and they contain almost all common and useful minerals, although many deposits are small in size. Important metallic ore deposits include Besshi (Kieslager)-type (strata-bound cupriferous pyritic deposits), strata-bound Mn-Fe-type, skarn-type, Kuroko-type and vein-type. Dominant non-metallic deposits are limestone, clay, native sulfur, zeolite, silica and gypsum deposits. The deposits are divisible into three groups, based on their ages of formation: Carboniferous-Jurassic, Cretaceous-Paleogene and Tertiary-present. Carboniferous-Jurassic deposits, closely associated with submarine volcanic rocks, are of two kinds: Besshi (Kieslager)-type, and strata-bound Mn-Fe-type deposits. Besshi-type deposits are cupriferous pyritic deposits and occur mainly in metamorphic terranes (Sanbagawa, Sangun, Abukuma and Hidaka) and in some other areas (Chichibu and Shimanto; Fig. 1.2). The geological and geochemical similarities of these deposits and modern midoceanic ridge deposits (e.g., Juan de Fuca ridge, Guaymas) suggest a similar origin. For instance, the sulfur isotopic compositions of both types of deposits are equal to or higher than mantle values (generally + 1%0 to +4%0), suggesting mantle origin, perhaps modified by seawater-basalt interactions. In the Chichibu Zone, the intimate association of abundant strata-bound Mn-Fe deposits, limestone-dolomite and silica (chert) with basic volcanic rocks suggests an ocean-ridge hydrothermal origin. Jurassic-Cretaceous Besshi-type and Mn-Fe strata-bound deposits are present in Hidaka, Hokkaido (Fig. 1.2). Geochemical data and geological evidence all point to a midoceanic ridge environment of ore formation. 334S values of Shimokawa Besshi-type

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deposits range mostly from +7%o to +10%o (Miyake and Sasaki, 1980), suggesting a contribution of seawater sulfate in addition to mantle source sulfur. In C r e t a c e o u s - P a l e o g e n e time many skarn-type and vein-type deposits formed associated with intense granitic activity. Granitic rocks are divisible into magnetite-series and ilmenite-series (Ishihara, 1977). Magnetite-series granitoids are present in North Honshu (Kitakami) and in the inner zone o f Southwest Honshu (San-in) Belt and ilmenite-series granitoids in the outer zone of Southwest Honshu (San-yo) belt. Metals associated with these two types of granitic rocks are distinct: Mo, Cu, Pb, Zn, Au and A g with the magnetite-series; Sn, W and rare earth with the ilmenite-series. Isotopic (Sr, S and O) data suggest that the ilmenite-series granitic m a g m a was influenced by

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4

Chapter 1

contamination of carbon-bearing sediments, whereas the magnetite-series granitic magma ascended from the lower crust without a significant interaction with carbonaceous matter. In the Kitakami district (North Honshu) gold vein-type (mesothermal-type) deposits and Cu-Fe skarn-type deposits occur associated with magnetite-series granitic rocks. Ore deposits associated with volcanic rocks generally exhibit polymetallic (Cu, Pb, Zn, Sn, W, Au, Ag, Mo, Bi, Sb, As and In) mineralization. Sulfur isotopic values of sulfides from these deposits are close to 0%o, suggesting a deep-seated origin of the sulfide sulfur. Clay deposits (pyrophyllite, sericite and kaolinite) are associated with both felsic volcanic rocks and ilmenite-series granitic rocks of late Cretaceous age in the San-yo Belt. Tertiary-Recent mineralization in Northeast Japan includes both epithermal veintype and Kuroko deposits (Fig. 1.3). Kuroko deposits occur only in the Green tuff region, whereas vein-type deposits occur both in the Green tuff region and in subaerial Tertiary-Quaternary volcanic region (Non-Green tuff region). K-Ar ages of formation of Kuroko and vein-type are middle Miocene (13-t- 1 Ma) and Plio-Pleistocene (3 4-2 Ma), respectively. Precious metal vein-type deposits include enrichments in Au, Ag, Hg, Te, Se, Sb, As, S, and Bi. Base metal vein-type deposits contain Pb, Zn, Mn, Ag and Cu, whereas Kuroko deposits are enriched in Cu, Pb, Au, Ag and Ba. In Southwest Japan, two styles of vein-type mineralization (Hg and Sb) formed from middle Miocene to the present. Many Hg and Sb deposits are present along the Median Tectonic Line, associated with the Setouchi andesites and ilmenite-series granitic rocks (Fig. 1.3). These different sites of hydrothermal and ore-forming activity may have resulted from the mode of subduction of the Pacific Plate. Mariana-type subduction (characterized by a steep angle of subduction and back-arc basin formation; Uyeda and Kanamori, 1979) during middle Miocene caused WNW-ESE extension, submarine hydrothermal activity, thick accumulation of bimodal (basaltic and dacitic) volcanic activity (Green tuff) and Kuroko-type formation (Shikazono and Shimizu, 1993). Plio-Pleistocene Chilean-type subduction (shallow-dipping subduction zone, E - W compression; Uyeda and Kanamori, 1979) and oblique subduction of the Pacific Plate beneath the North American Plate led to uplift and expansion of land area, subaerial hydrothermal activity accompanied by meteoric water circulation, subaerial andesitic volcanic activity and formation of vein-type deposits. Figure 1.4 shows the quantities of metals produced f¥om the metallic deposits in Northeast and Southwest Japan. This figure demonstrates that a large quantity of base metals (Cu, Pb and Zn) and precious metals (Au and Ag) was concentrated in the deposits of Northeast Japan, while they are much less abundant in the deposits of Southwest Japan. Subduction of the westward-advancing Pacific Plate under Northeast Japan was active during the Miocene, while in Southwest Japan, subduction along the Nankai Trough began more recently (5-10 Ma) and the Miocene ocean plate was inactive. These different plate motions may cause differences in quantities of sulfide ores and kinds of metals concentrated in the deposits. Lead and sulfur isotope data suggest that during subduction, pelagic sediments and altered basalt were incorporated into the magma in the island-arc trench systems of Northeast Japan and that lead and sulfur in vein-type and

Miocene-Pliocene Hydrothermal Ore Deposits

5

8, P Zn,S,~, ( £ A~) <~ Nn H£ Sb gure~ @NStS Bed~ ma~sanesee~s~s

Figure 1.3. Distribution of mineral deposits and tectonic provinces of the Neogene in Japan (Tatsumi, 1970). I, Zone of Miocene volcanism (Green Tuff region; mainly submarine); II, Zone of Miocene volcanism in the Ryukyu Arc (mainly subaeriaI); III, Zone of Neogene volcanism along the Median Tectonic Line (mainly subaerial); IV, Zone of Late Neogene folding, mainly in the Green Tuff region; V, Zone of Neogene intrusives and extrusives in the Shimanto terrane. Kuroko deposits in Northeast Japan originated from these materials. Antimony, mercury and sulfur in the H g - S b deposits in Southwest Honshu may have been derived from the shallow level of the crust under the Shimanto Group. Large epithermal gold vein-type deposits occur at major a r c - a r c junctions (Figs. 1.5 and 1.6); specifically, Chishima (Kurile)-Northeast Honshu, Northeast H o n s h u I z u - B o n i n and Southwest H o n s h u - R y u k y u . This m a y result from hydrothermal activities and mineralizations caused by intense volcanism at the a r c - a r c junctions. Hydrothermal c l a y - s i l i c a deposits (kaolinite, halloysite, sericite, montmorillonite and silica) and zeolite deposits occur in Tertiary-Quaternary volcanic regions. These deposits are distributed in areas o f epithermal gold mineralization.

Chapter 1 Miocene-Pliocene

Quaternary

Isla2d HE

i

Valcanic ' ,I

Japan

SW dapa.~n.n 0

DRI 30000

1

20000

40000

Kuroko-type deposit

6000

U/ZA

2000

Vein-type deposits

20000 100000 1 2 3

NN

Sulfur deposits

Figure 1.4. Comparison of quantities of ore deposits formed in late Cenozoic in NE and SW Japan. Weight per kilometer length of island arc (lshihara, 1978).

Quaternary sulfur deposits are distributed along the present volcanic front. Intersections of transverse faults proposed by Carr et al. (1973) and the present volcanic front coincide with the locations of clusters of the sulfur deposits (Nishiwaki and Yasui, 1974). Recently, it was found that mineralization is taking place in and around the Japanese Islands: sulfide-sulfate chimneys were discovered at back-arc depressions of the Ryukyu Arc Okinawa (Trough) and Izu-Bonin Arc (Smith Rift). The geologic settings are similar to those of the Miocene Kuroko deposits. The Ryukyu Arc belongs to the Mariana-type because its back-arc region is under extensional stress and the Okinawa Trough probably is a nascent back-arc spreading basin (Uyeda, 1991). The Izu-Bonin Arc may also be a Mariana-type, at present. However, it is likely that the arc was a Chilean type because intense epithermal gold mineralization took place at 1-3 Ma in the Izu Peninsula. Gold-rich silica precipitates at the Osorezan volcano, which is located in the most northern part of Honshu, have features very similar to epithermal Te-bearing gold vein-type deposits of the Plio-Pleistocene.

1.2. General overview and classification of hydrothermai ore deposits of Neogene age Main hydrothermal ore deposit types of Neogene age that formed in and around the Japanese Islands are Kuroko deposits and epithermal vein-type deposits. This classification is based on the form of the deposits. Kuroko deposits are strata-bound and massive in form (Fig. 1.7) and syngenetically formed on the seafloor and/or sub-seafloor environment. Vein-type deposits are fissurefilling and epigenetically formed (Fig. 1.8). Elemental association can be used to sub-classify these deposits. Major metal elements produced from Kuroko deposits are Cu, Pb, Zn, Ba, Ca, Fe, Au, and Ag. Average ore grade and tonnage are summarized in Table 1.1. Horikoshi and Shikazono (1978) classified Kuroko deposits into three sub-types: C sub-type (composite ore type),

Miocene-Pliocene Hydrothermal Ore Deposits

Figure 1.5. Three island arc junctions in the Japanese Islands (Kubota, 1994). Y sub-type (yellow ore type), and B sub-type (black ore type), according to Cu, Pb and Zn ratios (Fig. 1.9). However, the variation in the ratio is not wide, compared with epithermal vein-type deposits. Therefore, characteristic differences in each sub-type of Kuroko deposits are not discussed here. Major epithermal vein-type deposits in Japan are base-metal type and preciousmetal type which are classified based on the ratios of base metals and Au and Ag which have been produced during the past (Table 1.2). Base-metal vein-type deposits may be divided into Pb-Zn-type and Cu-type (Otsu and Harada, 1963). However, this sub-classification is not considered here for simplicity of discussion.

T A B L E 1.1 Size and composition of ore deposits in the Hokuroku basin (Tanimura et al., 1983) Deposit name

Discovery (year)

Size (max. length x width x thickness, m)

Average ore grade Cu (%)

Zn (%)

Ph

Fe

(%)

(%)

1.13 2.28 1.32 2.60 I. 15 1.50 3.02 1.00 2.32 2.82 2.70 2.39

1.4 1.3 10.0 7.8 1.3 3.0 3.0 6.5 6.5 2.4 1.5 3.6

0.4 0.2 1.8 1.0 0.2 0.5 0.5

12.8 20.4 5.2 13.4 8.9 15.8 20.0

1.2

%0

0.5 0.4 0.3 1.0

7.0 24.4 14.8 3.6

14.7 23.4 6.0 15,4 10.2 18.2 23.0 8.0 8.0 28.0 17.0 21.1

Shakanai ore group (Shakanai Mines Co., Ltd.) No. I 1962 3 0 0 x 1 5 0 x 12 No. 2 1963 ?x?x7 No. 3 1963 400x 120x 6 No. 4 1963 400 x 300 x 40 No. 5 1964 350 x 70 x 13 No. 7 1965 350 x 250 x 15 No. 8 1965 430 x 170 x 40 No. II 1967 4 0 0 x l l 0 x 10

2.3 2.3 1.1 1.7 1.9 1.3 0.7

14.6 12.3 10.0 2.9 3.4 3.2 1.0

3,2 7,6 6.2 0.7 1.0 0.9 0,2

12.2 8.8 10.4 19.1 14.8 22.6 28.7

Matsuki ore group (Mitsubishi Metal Corp.) Matsuki 1964 350 × 90 x 30 Takadate 1963 150 × 150 x 30 Takadate South 100 × 70 x 15

1.9 3,74

11.8 2.00

3.4 0.80

0.89

10.1

3.32

Hanaoka-Shakanai area Hanaoka ore group (Down Mining Co., Ltd.) Tsutsumizawa [912 150 × 80 x 1 l0 Doyashiki 1916 270 x 200 x 100 Kamiyama 1919 60 x 40 x 55 Nanatsudate 1929 100 x 40 × 50 Higashi Kannondo 1935 35 x 15 x 25 O y a m a 1 Nishi 1938 60 × 13 x 30 Kannondo 1939 40 x 35 × 45 lnarizawa 1940 40 x 20 x 15 Ochiaizawa 1941 50 x 35 x 20 Oishizawa 1941 50 x 40 x 20 Oyama 2 1941 ll0x 17x30 Malsnmine 1963 600 x 400 x 110

Ezuri-Fukazawa

s

(%)

All (ppm)

Ag (ppm)

m m

m

Tonnage (1,000 metric tons)

441 8,946 932 677 105 60 258 34 207 90 100 30,000

0.5

57

t4.0 10. l 12.0 22.0 17.0 26.0 33.0

2.0 1.7 0.8 0.3

270 260 410 25

15.7 21.2

18.0 24.38

0.6

55

660 1,200

4.1

4.7

1.2

180

3,000

540 360

3,600 430 1,000

2,800

area

Ezuri ore group (Dowa Mining Co., Ltd.) Ezuri 1975

~.

T A B L E 1.1 (continued) Deposit name

Discovery (year)

Size (max. length × width × thickness, m)

Average ore grade Ca

(%) Fukazawa ore group (Dowa Mining Co,, Ltd.) Tsunokakezawa 1 1973 500 x 300 x 5 Kanayama 1976 210 × 90 x 8 Manjyaku 1979 190 x 190 x 13

1.13 1.6 1.0

Zn

(%) 15.4 19.0 10,1

Ph

(%) 3.3 5.8 1.5

Fc

(%)

4.4 7.9 3.6

S

(%) 5.1 9,1 4.1

AH

(ppm) 0.6

Ag (ppm)

93

Tonnage (1,0130 metric tons)

e~ I

~z

3,000

,5 g-

K o s a k a area

Uwamuki ore gnmp (Dowa Mining Co., Ltd.) No. 1 I962 1 5 0 × 1 0 0 x 14 No. 2 1965 200 × 150 x 40 No. 4 1966 350 x 1 0 0 x 17

0.6 0.8 0.8

11.5 7.8 8.3

4.2 1.8 2.8

4.1 7.7 5,5

4.7 8.8 6.3

0,7 -

Uchinotai ore group (Dowa Mining Co., L t d ) West 1959 400 x 300 x 70 East 1960 400 x 300 x 40

2.8 2.0

4.0 4.4

1.1 1.5

17.2 13.6

19.8 15.6

0.8

Motoyama ore grcmp (Dowa Mining Co., Ltd.) Motoyarna 1861 300 x 700 x 50

2.2

4.5

0.8

20.6

23.7

Furutobe ore group (Mitsubishi Metal Corp.) Yunosawa Daikoknzawa 1959 250 x 1 0 0 x 15 Daikokuzawa West $960 t70 x 70 x 40 Daikokuzawa Easl 1960 100 x 60 × 20 M a g a r i y a z a w a East 1962 200 x 80 x 15 M a g a r i y a z a w a West 1962 150 x 50 x 15

1,9 1.1 2,8 1.5 1.5

4.3 1.0 6.2 0.9 0.9

0.9 0.1 1.4 0.2 0.2

17,4 15.3 24,4

20.0 17.6 28.0

A inai ore group Ytmosawa Suehiro Daikoku Benten Yokodawara Hagoromo

0,7 4.7 2,2 1.9 2.0 1.9

tr 8.3 5,1 3.1 2.9 10.3

lr 1.7 1.3 1.2 0.7 3.5

21.8 17.6 16.1 15.7 18.3 20.0

25.0 20,2 18.5 18.0 21.0 23.0

120

130

160

930 2,580 5,240 4,000 15,000

Furutobe-Ainai area

1942 1955 1956 1957 1960 1967

200 50 180 200 150 150

x x x x × x

150 x 50 40 x 30 80 × 50 60 x 20 80 x 40 80 x 10

1.3

51

3,800

t.0 0.5 0.4 9,1

260 140 130 620

220 2,000 1,500 210

10

Chapter 1

r",, 12.912.8 2.93-049( !," ~,~,.;" 4.9~2.7\ ~ ~P "'t'~',,T/'~..,, ~[~] 3[W-] 5~b] 13

50.t < lO-50t I ~lOt o 0.II I

5.8

o 501oo 200 300km 1 ' ~ 3 " J I

I

I

I

I

(4.6)

3.6

0.6 1.1,3.7

1.1 4 0~--1.0 514 , ~1.5~1.8 ~-3.7 -4.3 Figure 1.6. Distributionand temporal and spatial relationshipof late Cenozoic gold deposits in the Japanese Islands. 1: Quartz vein-typegold deposits with little to no base metals. 2: Gold silver deposits with abundant base metals. 3: Distribution boundary of gold deposits formed during the Miocene. 4: Location of PlioPleistocene gold deposits at the actual island arc junctions. 5: Location of Plio-Pleistocenegold deposits in front of the actual island arc junctions. Numbers in the figure are K-Ar ages of epithermal Au-Ag veins (Kubota, 1994). Several sub-classifications of epithermal precious-metal deposits have been proposed: mineralogy, host-rock composition and elemental association (Lindgren, 1928), gold-silver ratios of metal weights (Ferguson, 1929; Nolan, 1933), mineral paragenesis (Nishiwaki et al., 1971), and production ratios of metals (Heald-Wetlaufer et al., 1983). Recently, epithermal gold deposits were divided into several types based on gangue minerals, and physicochemical environment of ore deposition (pH, H2S concentration of ore fluids). They are hot spring-type (Silberman, 1982; Berger, 1983a; Berger and Eimon,

11

Miocene-Pliocene Hydrothermal Ore Deposits

Epidote-rich Basalt 300m Acid/stuff'/,,~, ~,..--,/ ~ / ~ " ~'~A/kj~'~'+ " +~.~. L'ff~.~

.

.

~.

'.

J

.

~

~ .

ALTERATION

~XZ2~%8~\~,// ~ # \\%//..~.~,// ~/!

~. .

~

v

~

~,~//--,~.Stock. It Sill. . . . . Ore

AlteredB

~ - ~. ..' .. ..~. ,',

~,o~,p,,~

//-~ Ferruginous Chert (Fe-Mn) Barite Ore Black Ore ~ YellowOre ~

C h l o r i t e - r i c h Basalt

O ~~" ~

+~'Y++'I:~&i ,~o ~ , o ,. .. .. .. .. .. .. . . ....... ~..:::iii

. . : . : . : ~ . ~ , ,

Clay/Mudstone Sericite/Chlorite Alteration

,~l

400m

Figure 1.7. Schematic distribution of Kuroko orebody and hydrothermally altered rocks (modified after T. Sato, 1974).

Shishimano Dacite L~

Age 0.66 l.lMa

300 , /<'/

Portal/', t

~ ~ &'~

N / \ Hishikari Lower t Andesites \

," ", - ,, ,- -

Old Hishikari'~"~"~.~• \ / / '~ ,. ,, -Yamada / ~ \

\t \ /\/'" "~ \ 1' N / \ ,

\/\ x,"/ x\,'\ ," \ ? ' 4 / - ~ v V ~ ' - - < ) , l ~ / \ , 0.95 d W / x/\ / \ ,,"%.v,.~jc.,,v~7"-ff-~ ;" ,,, z \ ~

r/=/ =I~I

100ML'~-,,.kL.." ~ ; - . ; ~ { , / , . i~, '~

x'-"--x"-/Jf" z ""

~

I

Vein

1r

(Honko)

/

/

~

~-~

v

\ " \ z \ z \"~'4 1"79Ma /

tl

(Sanjin)

0.78-1.05Ma

,,,

\,,

pre

-Neogen~

Figure 1.8. Schematic northwest-trending section across the Main and the Sanjin deposits of Hishikari mine (Ibaraki and Suzuki, 1993).

12

Chapter 1

TABLE 1.2 Estimated total productions of Au, Ag and other metals and Ag/Au total production ratio (Ag/Au, by weight ratio) from the individual vein-type and disseminated-type deposits in Japan (Shikazono, 1986). Type I-A: gold-silver-rich deposits, Type I-B: base-metal-rich deposits, Type 2: disseminated-type deposits Mine

Deposit type

Au (M.T.)

Ag (M.T.)

Ag/Au

Hokuryu Sanru Numanoue Kohnomai Kitami Tokusei Taiho Teine Oe-Inakuraishi

l-A I-A I-A 1-B I-B 1-A I-A I-B 1-B

2.9 6.7 1.1 71.4 0.2 1.2 0.12 10.2 2.5

11.3 40 81.2 1219 22.2 14.7 10.2 158 109.5

3.9 6.0 71.4 17.0 11I 12.0 85 15.5 43.8

Todoroki Toyoha

1-A 1-B

5.7 2.3

209 914

37 404

Eniwa Koryu Chitose Oogane Shizukari Yagumo

1-A I-A I-A 1-A 1-A 1-B

0.71 0.76 22.8 1.6 7.4 0.4

5. I 22.2 105 35.0 5.1 70

7.1 29.4 4.7 19.4 7 175

Jokoku

1-B

-

152

-

Furokura

I-B

1.2

45

37.5

Osarizawa

l -B

6.2

251

40.5

Ani Takanosu Matsuaka

I-B I-B 1-B

1.0 0.6 1.53

32.0 3.9 62.6

30.8 6.6 40.9

Innai Hosokura

1-A l -B

1.0 2.9

400 527

400 184

Isobekoyama Handa Yatani

1-B 1-A 1-B

3.2 1 1.7

2.3 13 64

0.73 13 39

Other metals (M.T.)

Cu: 4720 Pb: 8850

Cu: 7291 Cu: I924 Pb: 17316 Zn:48100 Mn:307840 Pb: 226410 Zn: 558478

Cu: 70

Pb: 12000 Zn: 23200 Mn: 6444 Pb: 13810 Zn: 35906 Mn: 2762 Cu: 15500 Zn: 262500 Cu: 341000 Pb: 806000 Zn: 155000 Cu: 20770 Cu: 4770 Cu: 3150 Pb: 25650 Zn: 39600 Pb: 21534 Zn: 588596 Cu: 3850 Cu: 1270 Pb: 29210 Zn: 58420

Years of production 1928-1943 1925-1974 1923-1951 1917-1974 1934-1964 1930-1942 1912-1928 1932-i971 1890-1974

1903-1974 i914-1974 1929-1943 1903-1957 1936-1974 1932-1955 1918 1962 1931-1969

1941-1978

1904-I974

1885 1931-1969 1908-1950

1871 - 1953 1898-1977 1932 1919-1966 1870-1974

Miocene-Pliocene Hydrothermal Ore Deposits

13

TABLE 1.2 (continued) Mine

Deposit type

Au (M.T.)

Ag (M.T.)

Ag/Au

Other metals (M.T.)

Years of production

Mikawa

1-B

1.7

29.4

17.4

1942-1961

Sado Takatama Takahata Nebazawa Tochigi Ashio Ohito Toi Seigoshi Mochikoshi Yugashima Rendaiji

1-A 1-A 1-A 1-A 1-A 1-B 1-A 1-A 1-A I-A 1-A 1-A

57-77 28.8 3.1 1.0 0.115 3.1 1.03 18.4 13.5 4.9 2.2 5.6

1310 279.9 2.9 65 7.6 390 2.36 214 455 104 29.8 276

Cu: Pb: Zn: Cu:

9.7 0.95 65 66 125.8 2.3 11.6 34 21 13.5 50

Nawaji Shimonomoto Kishu Okinoura Takeno Kohmori Nakase Ohmidani Ikuno

1-A 1-B I-B 1-A 1-A 1-B I-B 1-A 1-B

1.5 1.3 2.2 4.5 4.6 0.123 2.6 0.3 2.1

25 80 179 4.4 91 28.0 12.8 79 403

16.7 62.5 80 1.0 19.5 228 4.9 267 194

Tada

1-B

0.0

0.2

8000

Sakoshi-Odomari OmorI Bajo Talo Fuke Okuchi Onoyama Yamaganc lsobe-Arakawa Kushikino Akeshi Kasuga Iwato

1-A 1-B 1-A 1-A 1-A 1-A 1-A 1-A 1-A 1-A 2 2 2

1.1 1.4 13.0 36.6 1.9 21 1.3 37.1 1.5 53.6 2.4 3.3 4.4

9.7 65.7

8.8 48.1

158.6 1.1 15.6 0.6 37.1 10 488 1.4 0.69 5.7

4.3 0.6 0.74 0.43 1.0 6.7 9.1 0.6 0.21 1.3

5281 1686 7857 5400

Cu: 671795

Cu: 1000 Mn: 15840 Pb: 3680 Cu: 89436

Cu: 20910 Sb: 1921 Cu: Pb: Zn: Sn: Cu: Pb: Zn:

76076 27664 152152 1521 12 2 7

Cu: 6331

1601-1970 I429-1974 1930-1976 1942-1974 1908-1950 1877-1966 1930-1952 1916-1965 1935-1976 1929-1962 1937-1972 1914-1959 1929 1971 1956-1962 1940-1974 1925-1942 1920-1949 1928-1968 1956-1966 1914-1974 1940-1973

1940-1973

1977-1982 I889-1918 -1945 1903-1973 1937 1947 1905-I974 1934-1963 i628-1955 -1970 1914-1974 1915-1974 1929-1966 1932-1980

14

Chapter 1

Cu

0

"Y" sub-type 0

0.- ~"~"~" .~s S @@

"C" sub-type

"B" sub-type

Pb

~@

•

8

•• •

\ Zn

Figure 1.9. Available data on the Cu, Pb and Zn ratio of total ore in a single unit deposit in the HanaokaKosaka district, marked with three sub-typesof Kurokodeposits (Horikoshi and Shikazono, 1978).

1983), quartz-alunite-type (Berger, 1983b; Berger and Eimon, 1983), high sulfidation and low sulfidation-type (Hedenquist, 1987), a low and high sulfur distribution and an alkali-type (Bonham, 1984, 1986), and Te- and Se-bearing types (Shikazono et al., 1990). Most of epithermal precious-metal vein-type deposits in Japan can be classed as adularia-sericite-type, and low sulfidation-type. Very few hot spring-type deposits (quartz-alunite-type, high sulfidation-type) are found in the Japanese Islands. A summary of various characteristic features of adularia-sericite type (low sulfidation-type) is given mainly in section 1.4. A few examples of hot spring-type deposits occur in the Japanese Islands. The characteristics of this type of deposits are described briefly in section 2.7. Shikazono et al. (1990) divided epithermal precious-metal vein-type deposits into Te-bearing and Se-bearing deposits. As will be considered later, Te-bearing deposits are regarded as intermediate-type of adularia-sericite-type and hot spring-type. The distinction between these two types of deposits is discussed in section 1.4. During the Miocene age, polymetallic vein-type (xenothermal-type, subvolcanictype) and gold-quartz vein-type (mesothermal-hypothermal-type) mineralizations occurred mainly in middle to western part of Japan. They are described in section 1.6.1. In section 1.6.2, Hg and Sb vein-type deposits are described. Each deposit type is distributed in a different metallogenic province (Fig. 1.3) (Tatsumi, 1970). Epithermal vein-type deposits occur in Miocene-Pliocene volcanic terrain.

Miocene-Pliocene Hydrothermal Ore Deposits

15

Polymetallic vein-type deposits occur in middle Miocene volcanic terrain in central and western Japan.

1.3. Kuroko deposits Hirabayashi (1907) defined "Kuroko" as an ore which is a fine compact mixture of sphalerite, galena, and barite. This definition can be applied to "black ore", but not to "yellow ore" or "siliceous ore" because these minerals are not abundant in these ores. Kinoshita (1944) defined "Kuroko deposit" as a deposit genetically related to the Tertiary volcanic rocks, consisting of a combination of Kuroko (black ore), Oko (yellow ore), Keiko (siliceous ore), and/or Sekkoko (gypsum ore) (Matsukuma and Horikoshi, 1970). The deposit is generally defined as a strata-bound polymetallic sulfide-sulfate deposit genetically related to Miocene bimodal (felsic-basaltic) volcanism (T. Sato, 1974).

1.3.1. Geological characteristics 1.3.1.1. Distribution Kuroko deposits occur in the Green tuff region which is characterized by thick altered volcanic and sedimentary piles of Miocene age. Distributions of Kuroko deposits and names of the representative mines are given in Fig. 1.10 and Table 1.1. Metals produced during the past are summarized in Table 1.1. Large Kuroko deposits occur in the Hokuroku district in Northeast Honshu (Fig. 1.11). Small numbers of Kuroko deposits are found in other districts such as Southwest Hokkaido, the northern part of Honshu (Shimokita Peninsula district), the middle of Northeast Honshu (Wagaomono and Aizu districts) and western Honshu (San-in district) (Fig. 1.10). It is clear in Fig. 1.10 that the distribution of Kuroko deposits is restricted in a narrow zone in the Green tuff region which was called a "Kuroko belt" by Inoue (t 969). This belt was formed by rapid subsidence under the extensional stress regime and is thought to have been a back-arc depression zone at middle Miocene age. The relationship between tectonic setting and formation of Kuroko deposits is discussed in section 1.5. 1.3.1.2. General geology, country rocks and tectonic setting A large number of studies on the general geology and stratigraphy in the Kuroko mine areas have been done. During t960-1970 many drillings were carried out by metal mining companies and the Metal Mining Agency of Japan. These data clarified geologic structure and stratigraphy of the mine areas. Many Kuroko deposits are distributed in the Hokuroku district, Northeast Honshu (Fig. 1.9). Therefore, general geology and stratigraphy of the Hokuroku district is briefly described below mainly following T. Sato (1974, 1977), Tanimura et al. (1983) and Ishikawa (1991) (Table 1.3). The lowest rock units are basement rocks, composed of phyllites, cherts, and minor sandstone probably of Paleozoic age. The oldest Tertiary formation, which is called Ohya

16

Chapter 1 oo

HOKKAIDO 200 km

HOKUROKU DISTRICT WAGAOMONO

04

DISTRICT

AIZU

HONSHU SANIN

~.~'<~ S~~HIKOKU

~

~.~ /

JKYUSHU IG GI ( F

Q

~

~t4

DISTRIBUTION OF KUROK0-TYPE MASSIVE SULFIDEDEPOSITS IN JAPAN °.% Network, Stratiform sulfide, Gold pyrite, *

°%

~,~

Pyrite, or typical Kuroko

Kuroko-typeGypsumor Barite Green TuffBeltof Japan

f;3~!~.~'~ Clusters of Kuroko Deposits

Figure 1.10. The distribution of the Green Tuff belt of Japan and the Kuroko-type massive sulfide deposits within it. Major mining districts are labeled and ore deposit clusters outlined (Cathles, 1983a).

(Menaichizawa, Sasahata), is composed mostly of brecciated andesite lavas and andesitic hyaloclastics. This formation is conformably overlain by the formation (Hotakizawa, Sunakobuchi) composed of thick sequence of basaltic lavas and tuff breccias with minor intercalations of mudstone and felsic tuff. The formation which is mostly composed of dacite lavas, tuff breccia and mudstone (Hanaoka, Yukisawa, Uwamuki formations) conformably overlies the Hotakizawa and Sasahata formations. The thickness of these formations is 300-400 m. Kuroko ore deposits occur at the upper part of this formation. White rhyolite lava domes characterized by intense sericite alteration have a close spatial relationship with Kuroko deposits.

Miocene-Pliocene Hydrothermal Ore Deposits

N

|

~

x

x,z

I /~

/

,--'"

",,)

X~,.

/.-

Oinzan~, ~LakeTow..~ada

? /CNamanyamaI~ Furutobe ,~,. ~ ~-

/~ H~naoka! ].:/ / d~ IShakanm"/

"I

__

17

\-~'x '~Kosaka'.,~ I , _):.-~'~- ~'lt /~" " ~

-

',,/,....

r=k===W=~.-,:.\ \ 1 ~-

J

.~Somaki i" " ",~

; ( ( ./'=.... ~"

•

-4"

/ t

Osarizawa ~, ',~,,= '/~fd ]

.,.

"-,.,

O |

-'--"~ Fault ----'-1 Anticline [--'~---1 Syncline

~ ~L. ~

M. Mioc. Sed. Basin

Mioc.

Sed. Basin

i

10 km |

---] Kuroko Deposit ~

Network Vein Deposit

Principal

---1 Vein Deposit lntrusives Figure 1.11. Geologic structure and ore deposits in the Hokurokudistrict, Akita Prefecture,North Honshu (T. Sato, I974).

The formation composed of alternation of dacitic tuff and mudstone and basalt lava (Tsutsumizawa, Kagaya and Akamori formations) conformably overlies the Hanaoka, Yukisawa and Uwamuki formations. These formations are on average 150 m in thickness. The interbedded felsic tuff and mudstones (Shishigamori, Shigenai and Harukizawa formations) of middle Miocene conformably overlie the Tsutsumizawa, Kagoya and Akamori formations. The thickness of these formations varies widely, ranging from 40 m to 250 m. The younger formations (Ittori, and Tobe formations) of late Miocene to Pliocene overlie the Shishigamori, Shigenai and Harukizawa formations and are comprised mostly of mudstones, interbedded felsic tufts, and tuffaceous sandstones. The total thickness of these formations is ca. 500 m. The formations of Pleistocene unconformably overlie the

18

Chapter I

TABLE 2.3 Simplified stratigraphic column in the Hokuroku district and correlations with the Oga stratigraphy (Tanimura et al., 1983) HOKUROKU DISTRICT AGE

BLOW'S ! OGA PENINSULA M.Y ZONE

PLEISTOCENE -

o

PLIOCENE - 5-

LATE MIOCENE

N.2~ N.ZZ N.20 N.I9 N.)8

K]lauro I~ (marine) FunokowoE (marine)

N.17

IO

N.I6 N.IS

OnnogawoE (marine)

N{ NJ N.[1

Tobe E

Sqinokomi E (01akI F.)

Shokona{ F. (Okuzu E)

N.S EARLY MIOCENE ~20-

OLIGOCENE

25"

"60" /

M.M.A.j.

OOWA MININB CO., LTD.

Weslern

JI/T-W

J~IU¥1Um~

Su~okobuchl P

MO-a'

~

Central

_~;

Ul-b T2182

{

~aolln S,dti. *A-z~.

~-

TO-o

MO-b rO-b MI-a TI

Ealtlr,B TowedoVo~nlCl

T,.oo,o,~,.,J

BGztn~

Itl6ri E

k.14

MIDDLE MIOCENE

SHAKANA]MINES CO.,LTO. and AK{TA PRER GOV~

~

"-~,TIUIlU~II¢IG"F'.'.-'~'. XogayeE~-,*.-.-.- Akofaorl F,o- %~

..... , HoaooknF..~. . . . . . .

Yuk~snwa

Uwomukl

T4/04

T5/M3/B~

MonzenG.

Ohya F. (MenoLchlzawa F.}

Toy/Ts

Oa~.ement

9aslme~!

P

E ", : ;-~

~

l:

"

'

"' "

. . . .

Ho~ok~zawa~

"

"

'

"

*.

'

'

"

. . . . .

"

~

"'"~

~,,~;c:,lzewa F. : : : : : ' : : : : : : : : : ~~,:::::::::::::::::::::::::::::::::::::::::.::

L':::':::::::::'

Right column represents the average thicknesses and names of formations in three parts (western, central, and eastern) of the Hokuroku district, adopted by the Dowa Mining Company, Ltd.; second and, third columns from right are, respectively, the formation names adopted by the Metal Mining Agency of Japan and by the Shakanai Mines Company, Ltd., and the Akita Prefecture government; heavy horizontal and near-horizontal lines represent specific times, for example, the Menaichizawa and Sasahata formations are contemporaneous and correlate with parts of the Monzen Group and Daijima Formation at Oga.

a b o v e formations. T h e formations are basin-fill sediments, Towada v o l c a n i c s and terrace and a l l u v i u m deposits. T h e g e o l o g i c history in the H o k u r o k u district is divided into five phases as follows (Tanimura et al., 1983). ( l ) A n d e s i t e v o l c a n i s m and subsidence from near-shore to shallow-sea environments ( 3 0 ( ? ) - c a . 17 m.y. ago). (2) M a j o r subsidence, formation o f the H o k u r o k u basin and basaltic v o l c a n i s m ( 1 7 - 1 6 m.y. ago). (3) Felsic and basaltic v o l c a n i s m s and f o r m a t i o n o f K u r o k o ores ( 1 5 - 1 4 m.y. ago). (4) A c c u m u l a t i o n o f m u d s t o n e s and tufts in q u i e s c e n t sea ( 1 4 - 5 m.y. ago).

Miocene-Pliocene Hydrothermal Ore Deposits

19

(5) Differential uplifts, broad folding, andesitic volcanism and formation of smaller basins (5 m.y. ago-present). The most important geologic events at the time of Kuroko mineralizations are rapid subsidence just before the mineralization and bimodal volcanism (contemporaneous basic and felsic volcanism) (Konda, 1974). Several different hypotheses on the tectonic setting of the Kuroko mine area have been proposed. They include volcanic front of island arc (T. Sato, 1974; Horikoshi, 1975a), rifting of island arc (Cathles, 1983a), back-arc depression (Fujioka, 1983; Uyeda, 1983), and back-arc basin. In recent years, many hydrothermal solution venting and sulfide-sulfate precipitations have been discovered on the seafloor of back-arc basins and island arcs (e.g., Ishibashi and Urabe, 1995) (section 2.3). Therefore, it is widely accepted that the most Kuroko deposits have formed at back-arc basin, related to the rapid opening of the Japan Sea (Horikoshi, 1990). The summary of the bulk chemical compositions (major elements, minor elements, rare earth elements), 87Sr/86Sr (Farrell et al., 1978; Farrell and Holland, 1983), microscopic observation, and chemistry of spinel of unaltered basalt clarifies the tectonic setting of Kuroko deposits. Based on the geochemical data on the selected basalt samples which suffered very weak alteration, it can be pointed out that the basalt that erupted almost contemporaneously with the Kuroko mineralization was BABB (back-arc basin basalt) with geochemical features of which are intermediate between Island arc tholeiite and N-type MORB. This clearly supports the theory that Kuroko deposits formed at back-arc basin at middle Miocene age.

1.3.1.3. Age of mineralization The age of Kuroko mineralization can be estimated from (1) K - A r ages of igneous rocks associated with Kuroko deposits and (2) foraminiferal assemblages in mudstone directly overlying Kuroko deposits. Many K - A r ages of igneous rocks in Kuroko mine area in Hokuroku district yield 11-16 Ma. Ohmoto (1983) considered based on these data that the age of Kuroko formation was 11-16 Ma. However, almost all of igneous rocks associated with Kuroko deposits were hydrothermally altered after the mineralization (Shikazono et al., 1998). Therefore, it is likely that the K - A t ages of altered igneous rocks give younger ages than that of Kuroko mineralization. In contrast to K - A r ages, foraminiferal ages yield more accurate ones. Horikoshi (1987) investigated the foraminifera1 assemblage in mudstone overlying Kuroko deposits in the Shakanai area of Hokuroku district and concluded that the Kuroko mineralization age is 15-16 Ma. However, foraminiferal age data obtained by Horikoshi (1987) are scarce. Yoshida and Yamada (2001) compiled K - A r age of igneous rocks in Hokuroku district (which are considered to have formed simultaneously with Kuroko deposits) at 12.7 Ma. Considering uncertainty of foraminiferal and K - A r ages it seems reasonable that the Kuroko deposits in Hokuroku district formed in 14-12 Ma (more likely 13.612.7 Ma).

20

Chapter 1

However, several small Kuroko deposits (e.g., Yunosawa in Hokuroku district, Kuroko deposits in Hokkaido) occur in the formation younger than middle Miocene age (16-14 Ma), suggesting younger ages (12-13 Ma). 1.3.1.4. Metals enriched and metal ratios Many elements are concentrated in Kuroko deposits. Metals and minerals recovered from the ores are Cu, Pb, Zn, Fe, Au, Ag, S, BaSO4 and CaSO4. Size, average ore grade and tonnage of representative ore deposits are summarized in Table 1.1 (Tanimura et al., 1983). Total tonnage concentrated in the Hokuroku district is Cu, 1.2 x 108 ton; Zn, 2.5 x 106 ton; and Pb, 0.8 × 106 ton, respectively. Tonnage of BaSO4 and CaSO4 is given in Table 1.4. Figure 1.9 shows the proportion of Cu, Zn and Pb contents of Kuroko ore (Tatsumi and Ohshima, 1966; Horikoshi and Shikazono, 1978). Horikoshi and Shikazono (1978) divided Kuroko deposits in the Hanaoka-Kosaka area of Hokuroku district into three sub-types based on the ratio of Cu to Pb and Zn which increases in order of the B (black ore), C (composite ore), and Y (yellow ore) sub-types (Fig. 1.9). Characteristic features of these three sub-types were summarized by Horikoshi and Shikazono (1978) and are briefly decribed below. The Shakanai No. 1 deposit in the Shakanai mine is a good example of the B subtype (Kajiwara, 1970a). The ore of the B sub-type deposits consists of predominantly of galena and sphalerite with lesser amounts of chalcopyrite. Ore deposits of this sub-type are usually not directly associated with dacite lava dome. However, it is known that domeshaped dacite occurs below some of this sub-type deposit (Kajiwara, 1970a; Tanimura et al., 1974). Total ore quantity of a single unit deposit is generally small, about one million tons. C sub-type deposits are often called typical Kuroko deposits. Sato (1970) and Horikoshi (1976) published the schematic sections of Kuroko deposits referring to the general geology of this sub-type. The major Kuroko deposits belong to this sub-type. The largest Kuroko deposit is the Doyashiki deposit in the Hanaoka mine which belongs to C sub-type. The total ore quantity may be more than 10 million tons. The second largest deposit of Kuroko deposits is the Motoyama deposit of this sub-type. About 7 million tons of

TABLE 1.4 Estimated total amount of barite and sekko (gypsum + anhydrite) Shikazono, t983) BaSO4 (105 t) Kosaka

Fukazawa

Uchinotai Uwamuki Motoyama Tsunokakezawa Manjaku Kanayamazawa Matsumine

13. l 5.7 I 1.8 10.6 2.6 1.5 12

CaSO4 (105 t) Wanibuchi Fukazawa Yokota Motoyama Hamago

16 1-2 14 10

21

Miocene-Pliocene Hydrothermal Ore Deposits

ore were mined. The ore contains 2.3 million tons of black ore, 1.1 million tons of yellow ore and 3.7 million tons of siliceous ore. This means that ore deposit consists of roughly equal quantities of black, yellow and siliceous ores. A part of geology of the C sub-type is exhibited strikingly in the abandoned open-pit of the Motoyama deposit. The Uchinotainishi deposit of the C sub-type was described by Horikoshi (1969). The hydrothermal activity responsible for the mineralization of the C sub-type deposits was mostly preceded by the uplift of lava dome and the subsequent steam explosion (Horikoshi, 1969). Some of the Kuroko deposits consist predominantly of pyrite containing a small amount of chalcopyrite. The ore deposits consisting predominantly of pyrite, either with an economical value of chalcopyrite or not, are called the Y sub-type deposits, which occur above dacite lava dome or lava flow, while copper-poor deposits occur mostly in pyroclastic rocks and are associated with a large amount of gypsum. The Matsumine deposit in the Hanaoka mine is typical of the Y sub-type. The Matsuki and Takadate deposits in the Matsuki mine are also classed as this sub-type (Kuroda, 1978). Many pyrite-rich ore bodies

1oJG>. t omk2

°m/ m

~

~ U c h i n o t a i - n i U

-200m

-204~ •/

\

m 0

-100 m

-200 m

Figure 1.12. Distribution of two different sub-types of Kuroko deposits in the Kosaka district, Akita Prefecture. "Y" sub-type deposits have not yet been discovered in the area. The top pre-Tertiary basement is contoured showing some depressed structures (Horikoshi and Shikazono, 1978).

22

Chapter 1

associated with a large amount o f g y p s u m ore were mined in the Hanaoka mine. It seems likely that these ore bodies are composed o f several unit deposits o f this sub-type. Figure 1.12 shows the areal distribution o f the B and C sub-type deposits in the Kosaka district. The Y sub-type deposits have not yet been found in the district. It appears that two zones characterized by the distribution of each sub-type deposit are distributed north-southernly in the Kosaka district as well as in the Hanaoka district (Fig. 1.13). Pyroclastic rocks in the Kosaka formation, in which all deposits occur, become thicker to the east, and probably moved from the eruptive centres to the east (Horikoshi, 1969). These types of evidence may indicate that the sea at that time became deeper to the east. Figure 1.12 shows also the top of the pre-Tertiary basements. Ore deposits, either B or C sub-type, occur above the crater-like depressions o f basements. The Shinsawa deposit is the sole example o f B sub-type in the midst of the H a n a o k a - K o s a k a district, so-called Hokuroku basin (Fig. 1.13). The Tsunokakezawa deposit in the Fukazawa mine and ore deposit in the Ezuri mine are also the B sub-type.

I

.( \

\

/,Ioo..

g ; ',,.l_..~"

Tsutsumlzawo

J

N

Ok :\

X .\_,oo.

',,x

/-"

m

Bsub,ype

"C" sub-type

\ -.,oo,,

/ ~

~

"/~

0

J

G

ShokooaiNo,tf

1000 m

Figure 1.13. Distribution of three different sub-types of the Kuroko deposits in the Hanaoka district. The top of MI mudstone is also shown to visualize the structure of country rocks (Horikoshi and Shikazono, 1978).

Miocene-Pliocene HydrothermaI Ore Deposits

23

The Y, C and B sub-types roughly correspond to types 1, 2 and 3 as defined by Urabe (1974a), who classified Kuroko deposits based on hydrothermal alteration and ore mineral assemblages: type 1, kaotinite-pyrophyllite~diaspore-type; type 2, sericitechlorite-type; type 3, sericite-chlorite-carbonate-type. Hydrothermal alterations in the Kuroko mine area are described in section 1.3.2. Most large Kuroko deposits belong to type 2 (or C-subtype). Type 1 occurs mostly in Northeast Honshu and Hokkaido. Type 3 deposits are distributed in middle Honshu and Southwest Honshu (San-in district). Most of the previous studies have been carried out on the deposits in the Hokuroku district. A summary of the mineralogical and geochemical characteristic features of Kuroko deposits in this district is given below (sections 1.3.2 and 1.3.3).

1.3.2. Mineralogical characteristics

1.3.2.1. Metal zoning, and ore and gangue minerals

Typical Kuroko deposits (C sub-type according to Horikoshi and Shikazono (1978)) are usually composed of gypsum ore, siliceous ore, yellow ore, black ore, barite ore and ferruginous chert ore in stratigraphically ascending order (Fig. 1.7). The main constituent minerals in each ore are as follows: gypsum, anhydrite, Mg-chlorite (gypsum ore), quartz, pyrite, chalcopyrite (yellow ore), sphalerite, galena, pyrite, barite, chalcopyrite, tetrahedrite-tennantite, bornite, electrum (black ore), barite, quartz (barite ore), microcrystalline quartz, hematite (ferruginous chert ore). Figure 1.14 shows the distribution of minerals in each ore zone (Matsukuma and Horikoshi, 1970). The occurrence of ore minerals in Kuroko deposits was described in Shimazaki (1974), Matsukuma et al. (1974) and Urabe (1974a). Sphalerite has been studied by many workers and has been used to restrict the chemical environment of ore deposition. Iron contents of sphalerite from replacement siliceous ores and fissure-filling vein ores are much higher than those from layered ores (black ore and yellow ore) (Takahashi, 1963; Sato, 1969; Urabe, 1974b; Urabe and Sato, 1978). That is to say, iron contents of sphalerite generally decrease stratigraphically upwards in a single unit of ore deposit (Urabe and Sato, 1978). Urabe (1974b) and Urabe and Sato (1978) explained this trend by increasing of oxygen fugacity due to the mixing of hydrothermal solution with ambient cold oxygenated seawater. However, it is also likely that this trend was caused by the decreasing of temperature towards stratigraphically upwards. Generally, iron contents of sphalerite buffered by iron silicates such as chlorite decrease with decreasing of temperature in active geothermal systems (Fig. 1.15) (Hayba et al., 1985). Chlorite is common gangue and alteration minerals in Kuroko deposits. Therefore, this process seems plausible. Iron contents of sphalerite are different in layered ores in different ore deposit and different sub-types of Kuroko deposits. Iron contents of sphalerite from the B sub-type deposits (Uwamuki No. 4, Shakanai No. 1, Ezuri, and Fukazawa deposits) show wide range but generally less than 0.2 wt% (Ono and Sato, 1995). The average value, however, is probably lower than the C sub-type deposits (e.g., Uchinotai deposits). This may

24

Chapter 1

Mineral

Yellow ore Siliceousl ore Siliceous PowdeP/ Bedded PyCpSp ,to J YO YO E}O

Black ore .... GaFz Mononlne BO~nite BO raI~BO BO

Pyrite

Barite ore

Ferrugineus quartz

I

Chatcopyrite Bornits

u

Idaits

Fukuchilite Chalcocite

LII L

Covelline

,, ,

,L.

ass

Betlchtlnite Fehlerz

Bournonite Enargite

m ,.

!,

., ,

Luzonite

Argentite

m m

immm mm

INm(m mmm me m m m m l W

Stromeyerits Jalpaite Polybasite Pyrareyrite

IIIIIII

Galena Sohalerite

Hi-disulphide

i

III

m ~

mm

!P

Mfflerite

I

i iii

Bi,mlnerals iiiii

Electrum Silver

?

Sulphur Magnetite

4IlL

Hematite

i

Quartz

I

Carbonates

III

i ill

I

ii

III

Clays Barite

m

Figure 1.14. Schematic diagram showing mineralogical changes in various kinds of ores of Kuroko deposits (Matsukuma and Horikoshi, 1970).

suggest that the C sub-type deposits formed at higher temperatures and close to the volcanic centre and B sub-type deposits are distal type and formed at lower temperatures. Tetrahedrite-tennantite composition varies widely in Kuroko deposits (Yamaoka, 1969; Yamaoka and Nedachi, 1978a; Yui, 1971; Horii, 1971; Shimazaki, 1974; Kouda, 1977; Shikazono and Kouda, 1979; Ono and Sato, 1995; Ishizuka and Imai, 1998).

25

Miocene-Pliocene Hydrothermal Ore Deposits 0.5 ÷

0.4 0.3 0.2 03 14.

0.1 0 -0.1

0

,,-I

-0.2 -0.3 -0.4 -O.5 200

220

240

260

HomoganizationTemperature°C

280

300

Fig. 1.15. Diagram showing the homogenization temperature of fluid inclusions vs. the iron content of the host sphalerite growth zone for sample locality NJP-X on the OH vein. The line shows the predicted iron content of the sphalerite if the sulfur fugacity of the system had been buffered by the triple point - - Fe-chlorite (daphnite), pyrite, hematite (Hayba et al., 1985). Generally, tetrahedrite-tennantite composition from Kuroko deposits is characterized by high Zn content, low Fe content, high Cu content, and low Ag content compared with those from vein-type deposits in Japan (Fig. 1.16). Rarely, it contains Hg up to 1 wt% (Ishizuka and Imai, 1998). Tetrahedrite-tennantite composition varies widely in a single orebody. For instance, Kouda (1977) analyzed tetrahedrite-tennantite from Fukazawa-Tsunokakezawa deposit and showed that the Fe and Zn contents are in a range of 0-5.5 wt% and 4 . 5 - 1 0 wt%, respectively. Wide compositional zoning and heterogeneity in a tetrahedrite-tennantite grain are c o m m o n (Yamaoka, 1969; Yui, 1971; Yamaoka and Nedachi, 1978a). Positive correlation between Zn contents of coexisting tetrahedrite-tennantite and sphalerite exists (Shikazono and Kouda, 1979). This relation can be explained in terms of the following reaction.

[(Cu, Ag)10Zn2(As, Sb)4Sl3]tet -]- (FeS)sp = [(Cu, Ag)10Fe2(As, Sb)4Sl3]tet +

(ZnS)sp (1-1)

where sp = sphalerite, and tet = tetrahedrite-tennantite. Electrum occurs mainly in the black ore zone dominantly composed of sphalerite, bornite, galena and barite. It is common in B sub-type such as the Ezuri and Fukazawa deposits. It occurs in brown compact black ore consisting of barite, galena,

26

Chapter 1

AG WT%

FE WT%

60

8

7

50

6 40

t'~: A

~

';',

A

5 4

30

3

20

2

10

1 10

20

30

40

SB W'P/o 35

f

C

30 f..J t'. 25

1

50

CUWT%

2

3

4

5

6

7

8

9 10 11 12 ZN WT%

D ",

A

,~

20

~ C ~'D

15

is

2o

:Is

AS WT%

Figure 1.I6. Chemical composition of tetrahedrite-tennantite (Shikazono and Kouda, 1979). A: Au-Ag vein-type deposits, B: Kuroko deposits, C: Taishu Shigekuma Pb-Zn vein-type deposits, D: Skarn deposits (Kamioka). sphalerite, tetrahedrite-tennantite, chalcopyrite, bornite, and Ge-bearing minerals (argyrodite), stromeyerite, pearceite, and mckinstryite (Ono and Sato, 1995; Ishizuka and Imai, 1998). Electrum is often associated with pearceite and/or tetrahedrite-tennantite (Ishizuka and Imai, 1998). Exceptionally, electrum occurs abundantly in the siliceous ore of the Nurukawa deposit (Yamada et al., 1987). The mode of occurrence of electrum was described by Matsukuma (1985). Chemical compositions of electrum from Kuroko deposits were summarized by Shikazono (1981) and Shikazono and Shimizu (1988a). The Ag content of electrum from Kuroko deposits varies widely from 4.7 to 89.4 atomic% (Fig. 1.17). Electrum with low Ag

27

Miocene-Pliocene Hydrothermal Ore Deposits 40 u

0"

30

LL

20

10

0

20

40

60

80

100

Nag Figure 1.17. Frequency histogram for the Ag content of electrum from Kuroko deposits in Japan (Shikazono and Shimizu, 1988b).

content occurs in siliceous and yellow ores. For instance, electrum in siliceous ore from the Nurukawa deposit contains about 20 Ag atomic%. Electrum with high Ag content occurs in black ore. Electrum in the brown ore occurring in upper part of black ore contains Hg up to 11 atomic% and Ag contents are positively correlated to Hg contents (Ishizuka and Imai, 1998). Native silver is found in bornite-rich black ore (Matsukuma and Yui, 1979; Matsukuma, 1985) and it is thought to be secondary mineral. Compositional zoning in electrum grain is common (Shimazaki, 1974; Imai et al., 1981). The Ag content of rim of electrum grain is higher than that of core. Although Ag content varies widely, it is generally lower than that of epithermal vein-type deposits. Although analytical data on bornite are few, some data show high Ag contents (max. 1.45 wt%) (Matsukuma, 1985). Pyrite is the most abundant ore mineral. It occurs as euhedral, framboidal, and colloform forms. Abundance of framboidal pyrite increases stratigraphically upwards. Colloform pyrite contains appreciable amounts of As and Cu (Nakata and Shikazono, unpublished), whereas these contents of euhedral and framboidal pyrite are less than the detection limit of an electron microprobe analyzer. Ishizuka and Imai (1998) found that the As content increases toward outer rim and reaches up to 5 wt% in the rim of colloform pyrite from the Fukazawa deposit.

28

Chapter 1 1.0.

~2:~?~~2 o' 002 $2

o

oI 03

of

04

0-8-

05

0"6" o6

oO6

=,

06 I.L

06

O.Z,-

0606

06

0-2

9oo-9 08 " ~ o 8 8 o8

o9 0"8

1~0

1'.2

1J4

o707 1:6

I ~8

2:0

A~ in 4 ( A I , S i )

Figure 1.18. Variation of Fe2+/(Fe2+ + Mg) and tetrahedral AI of chlorite from hydrothermal ore deposits: Japanese Neogene Cu-Pb-Zn vein-type (open circle) and Kuroko deposits (solid circle). Localities: 1 Ashio, 2 Yatani, 3 Toyoha, 4 Kishu, 5 Sayama, 6 Mikawa, 7 Furutobe, 8 Hanaoka, 9 Wanibuchi, 10 western Bergslagen (Shikazono and Kawahata, 1987). Dominant gangue minerals in Kuroko deposits are quartz, barite, anhydrite, gypsum, chlorite, sericite, and sericite/smectite. Morphology o f quartz changes from euhedral in the centre to the irregular in the margin of the deposits (Urabe, 1978). No amorphous silica and cristobalite have been found. Quartz is abundant in siliceous ore, barite ore and tetsusekiei ore. Minor amounts o f M g - m i n e r a l s (talc, M g - c h l o r i t e ) occur in sekko ore. Chlorite occurs in sekko ore and it contains high amounts of Mg (Fig. I. 18). Kuroko deposits are characterized by large amounts of sulfate minerals (barite, anhydrite, and gypsum). Estimated total amount of barite and sekko (gypsum ÷ anhydrite) from individual deposit is shown in Table 1.4. Sr contents o f gypsum, anhydrite and barite

29

Miocene-Pliocene Hydrothermal Ore Deposits Matsumine

6O

c- 10-

<

0 ..Q

5

E Z .

.

.

.

.

500

1000

.

.

.

1500 Sr

2000

con tent

(ppm)

Shakanai

t~o

t'- 10

<

0 r.~

5

E Z

2o o

Sr c o n t e n t (ppm)

Fukazawa

cO (l) cO >,

c- 10

<

0

$ ..Q

5

Z ....

N n .F1.

500 ,

•

I000 -

,

•

.r]

~500 2000 Sr c o n t e n t (ppm) -

,

.

.

.

.

t

Figure 1.19. Strontium contents of anhydrites from the Matsumine, Shakanai, and Fukazawa deposits (Shikazono et al., 1983).

are 150-2,000 ppm (gypsum, anhydrite), and 0.3-3 wt% (barite), respectively (Shikazono et al., 1983) (Fig. 1.19). Mixed layer clay mineral (sericite/smectite) is found in Kuroko ore bodies and altered dacitic rocks underlying the ore. This mineral is thought to have formed by the

30

Chapter 1

interaction of sericite with hydrothermal solution whose pH decreased by the deposition of sulfides by the reaction of MC12(aq) + H2S -+ MS + 2H + + 2C1- (where M is metal such as Zn and Pb) (Tamura, 1982). Tamura (1982) thought of the following mechanism of formation of mixed layer clay mineral (sericite/smectite). Mg 2+ and Ba 2+ in hydrothermal solution may play an important role to form the interstratification of the minerals. From the experiment in which Li + has caused the formation of the interstratified structure, Mg 2+ with most similar ionic radius (0.65 A) to Li + (0.60 ,~) may accelerate the interstratification. Tomita and Sudo (1981) and Shimoda et al. (1974) have synthesized the mixed layer minerals from sericite reacted with LiNO3 solution. Li + gets into the unoccupied site in the octahedral layer of dioctahedral muscovite (sericite). To retain electroneutrality in muscovite structure, interlayer K + is easily released. Potassium is held strongly adjacent to interlayer regions where the potassium is replaced, i.e., the two regions will alternate. This is interstratification. Reichenbach and Rich (1968) have revealed that interlayer K could be removed almost completely from muscovite by using 0.1 N BaC12 solution at 120°C. This K - B a exchange may induce to form the interstratification. To verify the above interpretation, the analysis of Ba and Li in sericite/smectite is required. Preliminary analysis of sericite/smectite shows high Ba content (Ogawa and Shikazono, unpublished).

1.3.2.2. Hydrothermal alteration Several studies identified the hydrothermal alteration halo in the dacitic rocks surrounding Kuroko deposits (Fig. 1.20) (Shirozu, 1974; Utada et al., 1974, 1981; Utada, 1980; Ishikawa et al., 1976; Izawa et al., 1978; Date et al., 1983; Urabe et al., 1983; Ishikawa, 1988; Marumo, 1989; Inoue and Utada, 1991; Shikazono et al., 1995). For example, Date et al. (1983) recognized the following alteration zones in the Fukazawa Kuroko mine area of Hokuroku district from the centre (near the orebody) to the margin: (1) sericite-chlorite zone (zone III in Figs. 1.20-1.22) characterized by quartz + sericite -t- Mg-rich chlorite; (2) montmorillonite zone (zone II in Fig. 1.20) characterized by Mg.Ca-type montmorillonite + quartz -t- kaolinite 4- calcite 4- sericite -tFe-rich chlorite; and (3) zeolite zone (zone I in Fig. 1.20) characterized by clinoptilolite + mordenite + Mg.Na-type montmorillonite 4- cristobalite 4- calcite or analcime + Mg.Na-type montmorillonite + quartz + calcite -t- sericite -t- Fe-rich chlorite (Fig. 1.20). Kaolin minerals (kaolinite, dickite, nacrite), pyrophyllite and mica-rich mica/smectite mixed layer mineral occur as envelopes around barite-sulfide ore bodies in the footwall alteration zones of the Minamishiraoi and Inarizawa deposits, northern part of Japan (south Hokkaido) (Marumo, 1989). Marumo (1989) considered from the phase relation in A1203-SiO2-H20 system that the hydrothermal alteration minerals in these deposits formed at relatively lower temperature and farther from the heat source than larger sulfide-sulfate deposits in the Hokuroku district. Date et al. (1983) found the existence of an Na20-depleted dacite mass with a lateral dimension of 1.5 x 3.0 km immediately below the ore horizon (Figs. 1.23 and 1.24) and the mass is useful indicator of exploration of Kuroko ore deposits. This Na20 depletion is considered to be due to the destruction of plagioclase attacked by potassium-

31

Miocene-Pliocene Hydrothermal Ore Deposits D-3

I

TKIS;

/

/

I

I

N

TKI?9

f

o

TKt$3

®

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/

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/I

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LEGEND ."'''4

ORE

(3~i

DEPOSITS

...........' ORE ZONE

~

, , ~ '~" - ' ~ " ~ .

T,.~.. TKt/~

~K,o3)

~"'

ALTERED DACITE I N FOOTWALL

....

•

ZONE I

4)

ZONE

II

O

ZONE It"

•

ZONE IU

•

ZONE

IV

TK207 e

TKI&~ e

TK 1~,6 e

ZONE BOUNDARY

Figure 1.20. Zoning map of alteration minerals in unit D3 around the Fukazawa deposits (Date et aI., 1983).

rich hydrothermal solution and the formation of K-sericite in discharge zone. It is thought that potassium is added from ascending hydrothermal solution and sodium in dacite removed to hydrothermal slution. Hashiguchi et al. (1983) also showed that the variation of Na20 content in the footwall rock is particularly useful for detailed exploration based on the large number of chemical analyses of footwall rocks and statistical treatment of the geochemical data. Singer and Kouda (1988, 1992) confirmed based on the statistical analysis of the distribution of minerals and bulk compositions of altered rocks in the Hokuroku district

32

Chapter 1 N

om

S

=

o

°T 1

-:7-

°

LEGEND

•

T1, MI.T2, M2 :ROCK UNITS IN HANGINGWALL T3, D3, T4,04 :ROCK UNITS IN FOOTWALL :INTRUSIVE DACITE

:INTRUSIVE OOLERITE { ~

0

T2

0.5

i

:ZONE I

~:ZONE

U ~nd II

~:ZONE

Ill

~:ZONE

IV

: BASALT LAVA

................ :ZONES BOUNOARY

:KUROKO DEPOSITS

~FF~:Na POORPART IN FOOTWALL

1,0 km

Figure 1.21. Zoning of alteration minerals in the Fukazawa area. The location of the profile line is shown in Fig. 1.20 (Date et al., 1983).

HANGINGWALL u~ ~

ZEOLITE

I

MONTMORILLON I T E

It'

Ca-MONT.

==~ S~RICITE -~ i

11

Ill

IV

ANALCI ME MOROENITE No-MONT.

0

~u

CHLORITE

0

PLAG IOCLASE HYDROTHEIIMAL

ALTERATION DI A G E N E ~ L,~

FOOTWALL I[ u') c= ~ t~ Q:

ZEOLITE

MORDEN I T E

NtONTMORILLON1TE

Co-MONT.

~

SERICITE

~

CHLORITE

III

Mg-rlch

PLAGIOCLASE HYDI !OTHERMAL

ALTERATION DIAGENES I S )EPLETIVE ADDITIONAL

ELEMENTS ELEMENTS

CoO Na20 SiO2

CoO

No20

CoO, FeO

IMgO

!FezO3

Figure 1.22. Summary of alteration minerals and zoning around the Fukazawa deposits (Date et al., 1983). Legend: unbroken line = present or formed in considerable quantity; dashed line = present in small amounts, or uncertain formation; definition of zones by Date et al. (1983).

33

Miocene-Pliocene Hydrothermal Ore Deposits

i

~FIAe..

I

"

;<-'~

/J~

. %.

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.

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.~.'~ ~ % .

~-~r-.<~z.

~

• \

~, f

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e~

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~-,,g.,~./~%Z;

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e

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° °

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soo

,ooo~

F i g u r e 1.23. Distribution m a p o f low N a 2 0 a n o m a l i e s in the K o s a k a a r e a ( D a t e et al., 1983).

that Na20 depletion, as well as the presence of sericite, anhydrite and gypsum near the orebody, is a useful indicator for the exploration of Kuroko deposits. Figures 1.25 and 1.26 show the map of probabilities of centres of Kuroko deposits based on the data by Singer and Kouda (1988). Green et al. (1983), Urabe et al. (1983) and Matsuhisa and Utada (1993) showed that the analysis of whole-rock 6180 values is a useful method of exploration for volcanogenic massive sulfide deposits. They found that concentric zoning patterns of the whole-rock 8180 values and of alteration minerals are particularly well developed in the footwall volcanic rocks in Kuroko mine area. For instance, 8180 values of igneous and sedimentary rocks are +16.9-4-2.7%o in the zeolite zone, +11.1+2.5%o in the montmorillonite zone, and +6.7 4- 1.3%o in the sericite-chlorite zone in the Fukazawa area (Green et al., 1983) (Figs. 1.27 and 1.28). Figure 1.29 shows the 6180 variations as functions of temperature and water/rock ratio (Green et al., 1983). This calculated result indicates that the spatial variation of 8180 values of the footwall rocks and alteration

Chapter I

34

. .

.

TWNOKAKEZAWA No2

.

$? /

LEGEND Low Na, 0 Zonr (Naz0<0. 3 2 1 )

Drill hole 15 or more w r n p l * ~ taken 0

SARUMA Vrln

KUROKO 0.porit

./

0

.

/

4 /

!4 OT f.**,

sornpt..

y/q./. >.

token 0

100

500

M

1000 J

Figure 1.24. Distribution map of low Na20 anomalies in the Fukazawa area. The distribution of green dacite is also shown (Date et a]., 1983).

mineralogy can be interpreted as a result of interaction between the rocks and seawater at different temperatures (25-200°C for the montmorillonite zone and 200400°C for the sericite-chlorite zone) under water-dominated (e.g., waterlrock ratio > 1) conditions and most of the hydrogen isotope data of the whole-rock samples from the Hokuroku district (6D = -34%0 to -80%0) can be explained by this interpretation (Green et al., 1983). However, the variation of of altered rock depends not only on waterlrock ratio and temperature, but also on the degree of mixing of hydrothermal solution and cold seawater. This mechanism is considered in section 1.4.2. Ohmoto et al. (1983) found that 6 1 8 0 values alteration zones and major and base metal contents of footwall dacite along the section in the Fukazawa area correlate with each other (Fig. 1.30). Shikazono et al. (1998) found that carbonates are common alteration minerals in the Uwamuki mine area of Hokuroku district and carbonate alteration superimposed on chlorite alteration. They showed that the mode of occurrences and the Mg/(Mg

+

35

Miocene-Pliocene Hydrothermal Ore Deposits

I

N 5 km

Probability ~ ~7

P
[~

0.001<=P<=0.1

~

FURUTOBE

/

0.1
P>0.5

Kuroko Deposit Area considered ~

1 ~

i

UCHINOTAI SHAKANAI

ODATE

TOWADA MINAMI

Figure 1.25. Map of probabilities of centres of Kuroko deposits based on sodium depletion, sericite, and gypsum plus anhydrite(Singerand Kouda, 1988). Fe) ratios of magnesite and dolomite occurring in hanging wallrocks are useful in the exploration for concealed volcanogenic massive sulfide-sulfate deposits. 313C and 3180 of carbonates from the Kuroko mine area are plotted in Fig. 1.31. 313C and 3180 data lie between igneous (3~3C = -7%o, 3180 = +8%0) and marine carbonate value (313C = 0%0 and 3180 = +20%o). This indicates that magnesite and dolomite formed due to the interaction of hydrothermal solution with the biogenic marine carbonates. Dolomite and magnesite data are plotted close to marine carbonate values, suggesting that they formed in the central zone close to the ore bodies due to the interaction of hydrothermal solution with the biogenic marine carbonates (Fig. 1.31). 313C and 3180 of manganoan calcite in altered basalt directly overlying Kuroko orebody are close to igneous carbonate values, suggesting they formed from ascending hydrothermal solution at discharge zone.

36

Chapter 1

Figure 1.26. Map of probabilities of centres of Kuroko deposits based on sodium depletion in the Hokuroku district (Singer and Kouda, 1988).

Superimposed alterations are common in the Kuroko mine area (Inoue and Utada, 1991). For example, K-feldspar, kaolinite, alunite, pyrophyllite and diaspore alterations cut chlorite alteration, indicating that they formed later than chlorite alteration (Inoue and Utada, 1991). Inoue and Utada (1991 ) thought, based on detailed descriptions of the hydrothermal alterations in the Kamikita mine area, North Honshu, that hydrothermal alterations in this district started from 13 Ma and ended at 3 - 4 Ma. Pyrophyllite and diaspore alterations were reported from several Kuroko deposits, although they are not common (Urabe, 1974a). This type of hydrothermal alteration is thought to have occurred at a later stage than the hydrothermal alterations associated with Kuroko mineralization (sericite, chlorite, and zeolites) (Utada, personal communication, 1995). As well as felsic volcanic rocks, basalt occurs in the Kuroko mine area. It is also intensely and hydrothermally altered. Shikazono et al. (1995) studied the hydrothermal

37

Miocene-Pliocene Hydrothermal Ore Deposits

i/,

1"

- N-

,' e~

~'~

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~11A.8

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!/

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1~.5 •

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LEGEND

8,5 010.5

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~=='%1 lS.8 ~ ~1o.7 TAKARAKURA t/[ III • t~z.8

KANAYAMAZAWA ~

''~

-

--A'

09.5

ZONE h Zeolite Zone Ih M o n t m o r i l l o n i t e Z o n e IIh S e r i c i t e - C h l o r i t e Z o n e

0 I

z

1 km I

: Kuroko Ore Body : Vein Deposits e6.0 : 8180 .R.(%o) Value

17.8

12.0

10.3 11.9

Figure 1.27. Areal distribution of the whole-rock8180 values of footwall volcanicrocks in the Fukazawa area. The boundaries for the alteration zones are modifiedfrom Date et al. (1983) (Green et al., 1983). alteration of basalt overlying the Kuroko orebody in the Furutobe mine area of Hokuroku district. They showed that hydrothermally altered basalt can be divided into chlorite-rich rock and epidote-rich rock. Chlorite-rich rock occurs widely, whereas epidote-rich rock occupies a smaller area, close to the orebody. It was found that the CaO, Na20 and SiO2 contents of the bulk rock correlate negatively to MgO content, while FeO and NFe contents correlate positively to MgO content (Fig. 1.32) and these changes can be explained by seawater-basalt interaction at elevated temperature. The MgO/FeO ratios of

38

Chapter 1

.

Mont.

Facies [] Mudstone

Ser.+ Chl.

Basalt, Andesite

Facies

0

[] Tuff J • Dacite

5

10

8180 (%o)

15

20

25

Figure 1.28. Whole-rock 3J80 values of Miocene volcanic and sedimentary rocks from the Hokuroku district, grouped by alteration zones. Each square represents one sample. Mont. = montmorillonite, Ser. = sericite, Chl. = chlorite, av. = average (Green et al., 1983).

chlorite and actinolite and the Fe203 content of epidote from the basalt are greater than those of midoceanic ridge basalt (Figs. 1.33 and 1.34) probably owing to the differences in the FezO3/FeO and M g O / F e O ratios of the parent rocks. The lower CaO content and the higher N a 2 0 content of the bulk rock compared with altered midoceanic ridge basalt are interpreted in terms of the difference in original bulk rock composition. 3D of epidote in hydrothermally altered basalt is in a range of -36.5%~ to 43.0%o (Shikazono et al., 1995). Using -37.5%0 as an average 3D value of epidote, the fractionation factor for the H isotope exchange reaction between epidote and H 2 0 (Graham et al., 1980), and 280°C as the temperature of formation of epidote estimated from fluid inclusion study was calculated to be +2.3%~. This value is close to that of hydrothermal fluid issuing from the East Pacific Rise 21°N (+2.0%~, Bowers and Taylor, 1985). This strongly suggests that epidote formed at a discharge zone in submarine hydrothermal system (Shikazono, 1984). It is also notable that the 3D values of epidote from the basalt from Kuroko mine area are close to those of hydrothermally altered midoceanic ridge basalt in the Costa Rica Rift and Galapagos Ridge (-31%~ to -45%~) (Kawahata et al., 1987). 1.3.3. Geochemical characteristics

A large number of geochemical studies on Kuroko deposits (fluid inclusions, gas fugacities, chemical and isotopic compositions of ore fluids etc.) have been carried out. These are summarized below.

39

Miocene-Pliocene Hydrothermal Ore Deposits 25

[6m0i.÷6-'

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Water / Rock ( Atomic Oxygen) Figure 1.29. Calculated changes in the ~180 values of volcanic rocks (~I80 = +7.0%o) as a result of equiIibrium oxygen isotope exchange with waters of different initial compositions. The dotted areas represent the 3180 ranges of rocks in the zeolite and the sericite-chlorite zones (Green et al., 1983).

1.3.3.1. Fluid inclusions Fluid inclusions from Kuroko deposits were studied first by Tokunaga and Honma (1974) who showed that Kuroko deposits formed in a range of 200-260°C for the siliceous

40

Chapter 1

(N)I~3

]] Hole NO

114

180

lance (Kin} 4 .~ration Zone - . . ZeoI

,

?

161

I~8 %54~ 5

~

5~ I~68 1~1 I03

o Ore Zone 0

Mont

Mont

~,

{s}

z

20f

(t}]

(,.ol :(

1

(S)

(NI

Cu

(ppm)

MgO

(wt%)

CaO (wt.%)

:61

0

Zn:f

'

:

!

~'"

'

mt

(ppm) Ioo

:t

(wt%)

z

SF :t

{ppm)

.

I"

"

'PP" ':I'

KeO (wt.%) 4

i

OI

I

Figure 1.30. Comparisons of the ~5180values of footwall dacite, alteration zones, and other geochemical halos along the section in the Fukazawa area (Ohmoto et al., 1983).

ore, 100-240°C for barite and 140-170°C for black and yellow ores. Lu (1969) reported homogenization temperatures of the fluid inclusions in quartz from the Uchinotai-Higashi deposit of the Kosaka mine to be 200-250°C. Watanabe (1970) studied fluid inclusions from Ainai deposits and found that the homogenization temperature for the siliceous ores (190-300°C) is higher than that for the black ores (120-290°C). Marutani and Takenouchi (1978)clarified the variations in homogenization temperature and salinity of inclusion fluids in quartz from stockwork siliceous orebodies at the Kosaka mine (Fig. 1.35; Urabe, 1978). They showed that the temperature decreases stratigraphically upwards from stockwork ore zone (280-320°C) to bedded ore zone (260-310°C). Pisutha-Arnond and Ohmoto (1983) carried out fluid inclusion studies of the stockwork siliceous ores from five Kuroko deposits (Kosaka, Fukazawa, Furutobe, Shakanai, and Matsumine) and revealed that black ore minerals (sphalerite, galena, barite) and yellow ore minerals (chalcopyrite, quartz) formed at 200-330°C and 330 4-50°C, respectively, and salinities of the ore fluids remained fairly constant at about 3.5-6 equivalent wt% NaC1. They analyzed fluids extracted from sulfides and quartz; Na ---0.604-0.16 (mol/kg H20), K = 0 . 0 8 ± 0 . 0 5 , Ca = 0.064-0.05, Mg ----0 . 0 1 3 ± 0 . 0 0 8 , C1 ---- 0.82-t-0.32, C (as CO2) -- 0 . 2 0 ± 0 . 1 5 and less than 6 ppm each for Cu, Pb, Zn and Fe.

41

Miocene-Pliocene Hydrothermal Ore Deposits 0 -1

-2

0

~-3

0

¢,0

o ^~"~ //~///

.g/°--orJ..o" ,, 0///

+o0,% .~" . o

-4

/I

-5

q. 0 /

//

-6

~/

0

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0

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0

-7 I

5

f

t

i

s

I

i

10

,

~

t

I

I

15 a18 0 ( ~ )

I

I

~

I

20

~

i

,

i

I

25

,

Figure 1.31. Bivariable plot of oxygen versus carbon isotopic compositions of carbonates. Solid circle: magnesite; open circle: dolomite; open square: calcite; A: oxygen and carbon isotopic compositions of igneous carbonates; B: oxygen and carbon isotopic compositions of marine carbonates (Shikazono et al., 1995).

Using homogenization temperature and freezing temperature data (Fig. 1.36) and pressure-temperature diagram of N a C 1 - H 2 0 - C O 2 system, minimum seawater depth at the time of Kuroko mineralization is estimated to be 1,000-2,000 m, if boiling of ore fluid did not occur (Fig. 1.37, Pisutha-Arnond and Ohmoto, 1983). This estimated depth is close to that of seawater associated with present-day hydrothermal mineralizations at back-arc basins such as Okinawa Trough (section 3.3). However, Lu (1983) found that the salinity and homogenization temperature of fluid inclusions from the Uchinotai-East ore deposit varies widely and thought that this variation is due to boiling of Kuroko ore fluids. If his argument was correct, the depth could be estimated to be 1,000-1,500 m. Two hypotheses of seaftoor depth at the time of mineralization have been proposed based on foraminiferal data, ca. 3500 m (Guber and Ohmoto, 1978; Guber and Merrill, 1983) and 1500 m (Kitazato, 1979). Considering seafloor depth of present-day ore formation at back-arc basins and fluid inclusion data mentioned above, shallow seaftoor depth hypothesis (Kitazato, 1979) seems more likely. If the pressure-temperature condition of Kuroko ore fluids was close to the boiling curve, the depth could be estimated to be 1,000-1,500 m, which is similar to that for present-day back-arc mineralization such as Okinawa Trough.

1.3.3.2. Gas fugacities Sulfur fugacity (fs2). If electrum is in equilibrium with argentite, the equilibrium constant for the sulfidation reaction (K1-2), 4 Ag (electrum) + 82 = 2 Ag2S (argentite)

(1-2)

where Ag (electrum) is the silver component in electrum, can be expressed as, KI-2 =

a2 I(a 4 ~ Ag2S/I, AgJS2)

(1-3)

42

Chapter 1 30"

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2+oi

2~2~o 22 +2s 7 6o°

2s..~g ~o o$ Z~eoI o" °~ 011

030 012 o13 029

=tl

v

oa

%

017 g o o3

Ol

2

4

6 MgO

I

12

( wt.% )

14

16

,

.

2

8

.

.

10

.

12

}&

McjO ( w t . % )

16

43

Miocene-Pliocene Hydrothermal Ore Deposits A

I

0

! r~

E

D

Z

10

II

~2

)3

I~

15

m6

J7

Fe 2 0 3 Content (wt.%) of Epidote Figure 1.33. Frequency (number of analyses) histogram for Fe203 (wt%) of epidote from the Kuroko basalt. A: epidote coexisting with albite, B: epidote coexisting with chlorite, C: epidote coexisting with pyrite, D: epidote coexisting with hematite and calcite (Shikazono et al., 1995).

where aAg2S and aAg denote the activity of Ag2S in argentite and the activity of A g in electrum, respectively and fs2 is sulfur fugacity. Barton and T o u l m i n ( t 9 6 4 ) have derived a relationship b e t w e e n fs2, temperature, and the A g c o n t e n t of e l e c t r u m in e q u i l i b r i u m with argentite, u s i n g the equation of W h i t e et al. (1957) for the chemical potential of A g in electrum in c o m b i n a t i o n with the equation

Figure 1.32. The relationship between MgO concentration and other major constituents. Solid circle represents the sample of relatively fresh rock which contains original clinopyroxene. A: H20(+) vs. MgO. B: CaO vs. MgO. C: Na20 vs. MgO. D: K20 vs. MgO. E: SiO2 vs. MgO. F: FeO vs. MgO. G: Fe203 vs. MgO. H: EFe vs. MgO. I: A1203 vs. MgO. (Shikazono et al., 1995).

44

Chapter 1 15

A m

P;H J

10

m

< t-

"6

$ r~

E

5

. . . . . .

E-I . . . .

I

I' I

ill

I11 z.J . l . l . l

0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6 1.8 2.0 MgO/FeO (in wt.%) of Chlorite

15

B

>. ¢,t)

<

_c

r~

E

z

C

5

i

0

•

I¢IMI~IC c l c l c l IClClClC c l C l c l clclcl IclClClCfc c l c l c l ~ 1 , • .I¢IC1¢1~l~ c1¢1cl

0.2 0.4 0.6 0.8

. ,

1.0 1.2 1.4 1.6 1.8 2.0

MgO/FeO (in wt.%) of Chlorite Figure 1.34. Frequency histogram for MgO/FeO ratios (in wt%) of chlorite from the basalt studied (A) and MORB (B). Data sources are: Shikazono and Kawahata (1987), Humphris and Thompson (1978) (M: Mid-Atlantic Ridge) and Kawahata (1984) (C: Costa Rica Rift, Galapagos Spreading Centre). The data on chlorite from MORB are taken from typical metabasalt and not from quartz-chlorite breccia and veins which formed in a hydrothermaI upflow zone (Shikazono et al., 1987).

of Richardson and Jeffes (1952) for the standard free energy change of reaction, 4Ag (native silver) + $2 ----2Ag2 S (argentite)

(1-4)

The equation derived by them can be expressed as, log f s 2 = (1/4.576T) {- 4 1 9 8 0 + 16.52T - 18.296T log NAg

+ (4(1 -- NAg)2[5650 -- 1600(1 -- NAg) -- 1.375T]) }

(1-5)

where NAg and T denote mole fraction of Ag in electrum and absolute temperature (Kelvins), respectively.

45

Miocene-Pliocene Hydrothermal Ore Deposits Bedded orebody L-40

,

BSO ,

::.,

~e

:t ,

eoe

YSO

,

L-55

BSO

'

,,

YSO

'

|

i•

(

¢

ee eee eeeeeee

I

I

tee e.e oe ceel.¶•

I

1

i!L

Stockwork orebody L-70

BSO

'

S,

• l eeoe e

•e~eee

L- 85

.

BSO ,

|

I

YSO

I

I

'

°Oeeee~eee

I|,;

,o,~o*,e

e

I

t

:|

eee oeoe ole, •lee OQII

k-100

BSO ,

,

,

eel•o••

:;:"F:

....

,

,,"

.

YSO

'

I

i ..e

"e g :t eF,:" b eeee te.eyee.

!

200 ° 300 ° Temperature °C

Figure 1.35. Summarized results of homogenization temperature determination in quartz from Uwamuki No. 4 Orebody shown for Kuroko-type (BSO) and Oko-type (YSO), and siliceous ores and for each level (Marutani and Takenouchi, 1978).

Barton and Toulmin (1964) and Barton (1980) have derived the correction to equation (1)-(5) which is necessary due to the solubility o f Au in argentite as a function o f temperature and electrum composition. Based on this equation and NAg, we can place a limit on fs2 and temperature (Fig. 1.38). This application to ore fluids responsible for Kuroko deposits has been done by Sat• (1969), Kajiwara (1970b) and Shikazono and Shimizu (1988a). B o r n i t e - c h a l c o p y r i t e - p y r i t e assemblage also defines i s z - t e m p e r a t u r e region. Combining the FeS content of sphalerite coexisting with b o r n i t e - c h a l c o p y r i t e - p y r i t e , f s z temperature can be determined (Kouda, 1977).

46

Chapter 1 I

Oe

g

•

6

Z

5

i

o

.

.

.

I

|

.

P r i m a r y fluid inclusions

O" ¢J

.

.

.

!

JU79SK19

S e c o n d a r y fluid inclusions

0

,.-,.

.

o

(b,

4

sw

1I

e

~2

o?

D

i

Low T

0

Secondary Inclusions (16) b.

.

.

&

.

s

|

.

• o @

@

. . . .

!

150

. . . .

I

200

. . . .

I

250

Temperature

eQ

300

(°C)

Figure 1.36. Homogenization temperature and salinity of inclusion fluids (Pisutha-Arnond and Ohmoto, 1983)•

•

"

I

Pure Wot =~t-

I I

tso-

LIQUID

//,~'/

-

,,,

=

/

loo

- ,sod

/ //

~'

-.-I=o~o

.

,

I

o - . I.'~

c,O~

/

/

Depth of Seowoter

J

/

/

(meters)

. /

,~

/

,.;.'~1 /

VAPOR

'*

A

Fluid I n c l u s i o n F i l l i n g T e m p s . 100

;~00

300

Temperoture ( % )

I

350"¢

400

Figure 1.37. Estimation of minimum depth of seawater at the time of Kuroko mineralization. P-T diagram of NaC1-H20-CO2 is from Drummond (I981) (Ohmoto et al., 1983).

47

Miocene-Pliocene Hydrothermal Ore Deposits I

I

|

I

t

*

i

!

I

,

-5

-10

O4

09

"0'~

-15

0

--I

-20

-25

0

I

100

!

200

300

Temperature °C Figure 1.38. The formative temperature and f s 2 of the black ore from the Shakanai mine. The arrow mark shows an assumed trend of deposition from earlier to later stage (Kajiwara, 1970b). py: pyrite, bn: bornite, cp: chalcopyrite.

Oxygen fugacity (foz). The relationship betweeen

is 2

and

/O 2 can

be derived from the

following reaction. H z S + 1 / 2 0 2 = H 2 0 + 1/2S2

(1-6)

Equilibrium constant for this reaction (K1-6) is expressed as, 1/2 1/2

K16 = (a.2o/

2)/(aH2s/o2

)

(1-7)

where a = activity. Assuming aH20 = 1 and FHzS (activity coefficient of HzS) = 1, we obtain, log fo2 = log fs2 -- 2 logmHzS -- 2log KI-6

(1-8)

where m = molality. Using this relation, we can estimate fo2, mH2S and fs2 at a given temperature. The temperature and f s 2 c a n be estimated from fluid inclusion homogenization temperatures, and chemical compositions of sphalerite and electrum coexisting with argentite and pyrite. H2S concentration of hydrothermat solution is thought to be controlled by chlorite-pyrite assemblage. It is deduced from the H2S concentration of present-day submarine hydrothermal solution which is 1-10 m m o l / k g H 2 0 (Gamo, 1995). Using estimated fs2, temperature, and mH2S, we can place a limit of f Q .

Carbon dioxidefugacity fc02. CO2 fugacity (fc02) of ore fluids is estimated based on CO2 concentration of fluid inclusions analyzed. By using equilibrium constant of the reaction, CO2(g) + H 2 0 ----H2CO3, and assuming aH20 to he unity, fco2 can be estimated.

48

Chapter 1 ~,t,l,~,,l~[[,I,t[,I,,,tl,,~,l

~,~l,t~[ix,,~lt,~,l~Z,l,~,l

Dolomite+K-mica+~.

Magnesite+Kaolinite+Quartz <:3

0

2

o

~--2 0

0

-r

t~

-4

o

Caleite+Mg-chlorite +K-feldspar

-6 -8

0

,,,I

50

....

I,,,,I,,,,I,,,jll,,z

100

150

200 Temperature( ~ )

250

300

_o =

Mg-chlorite "6

licit

0

50

,lllillLl

100

,I;lll,lllwl~w

150

200

Temperature (*C)

250

300

Figure 1.39. Relationship between fco2, aH2co3 and temperature for silicate--carbonate equilibria. Thermochemical data used for the calculations are taken from Helgeson (1969) (Shikazonoet al., 1998). Magnesite, dolomite and calcite occur in hydrothermally altered rocks near Kuroko orebody. The following equations are used to constrain f c Q values of hydrothermal solutions (Shikazono et al., 1998) (Fig. 1.39). 5CaMg(CO3)2 + (dolomite) =

KA13Si3OIo(OH)2 + 3H20 + 3SIO2 (K-mica) (quartz)

MgsA12Si3OIo(OH)8 + KAISi308 + 5CACO3 + 5CO2 (Mg-chlorite) ( K - f e l d s p a r ) (calcite)

(1-9)

5MgCO3 -t- A12Si2Os(OH)4 -t- 7H20 + SiO2 (magnesite) (kaolinite) (quartz) =

MgsA12Si3Olo(OH)8 + 5H2CO3 (Mg-chlorite)

(1-10)

Since temperature of formation of carbonates can be estimated from homogenization temperature of fluid inclusions in carbonates, we can place a limit of CO2 from the above equilibrium relationships. The estimated CO2 range is 1-0.01 mol/kgH20. 1.3.3.3. Chemical compositions o f ore fluids Assuming chemical equilibrium between Kuroko-forming minerals and Kuroko ore fluids, the chemical composition of Kuroko ore fluids can be estimated, using thermochemical data. Calculations for the purposes of estimating the chemistry of Kuroko

49

Miocene-Pliocene HydrothermalOre Deposits

ore fluids have been carried out by several investigators (Sato, 1969; Kajiwara, 1970b; Shikazono, 1976; Ohmoto et al., 1983). For example, if sphalerite is in equilibrium with aqueous solution, the chemical equilibrium for the following reaction can be derived. ZnS = Zn 2+ + S 2-

(1-11)

Assuming that activity of ZnS is unity, solubility product of ZnS (KI-ll) is expressed as, K i-11 = a z n 2 + a s 2 -

( 1-12)

Activity of S 2- is related to total dissolved sulfur concentration (~S) and pH in the region in which H2S is predominant among aqueous sulfur species. S 2-

+ 2 H + = H2S

(1-13)

The equilibrium relation for this reaction is expressed as,

KI_13=aH2s/(a2+as 2 )

(1-14)

Combining this relation with (1-12), and assuming aH2s = NS, we obtain, 2 K 1-14)= (azn2+]ES)/(a2+Kl-13) Kl-ll=(azn2+au2s)/(aH+

(1-15)

In the Kuroko ore fluids with 1 molal C1 concentration predominant dissolved Zn species are thought to be zinc chloro complexes such as ZnC12. For simplicity, assuming that ZnC12 is predominant among dissolved Zn species (~2Zn = mzncl2=aznc12/YZnCl2, where YZnCl2 = activity coefficient of ZnC12 and it is assumed to be 1, and ~2Zn is total dissolved Zn concentration), the relation between azn2+ and 2 Z n can be derived from the following chemical reaction, ZnC12 = Zn 2+ + 2C1-

(1-16)

The equilibrium constant for (1-16) is, Kl-16 = (KI-11KI-13a2H+/~S)(a21 /NZn)

(1-17)

Therefore, we obtain, log NZn = log K - 2pH + 2 log mcl + 2 log YcI

--

log NS

(1-18)

where K = (K1-I1Kl-~3)/K1-16 and Yc1- = activity coefficient of C I - . Thus, NZn can be calculated if the values of K1-11, Kl-13, KI-16, mci-, Yct- and NS are available. NS in ore fluids is generally in a range of 10-2-10 -3 mol/kg H 2 0 based on IES in present-day geothermal waters and fluid inclusion analytical data (Shikazono, 1972a). YcI- is represented as a function of ionic strength and temperature. Ionic strength is related to salinity which can be approximated as C I - concentration. C1- concentration can be estimated from fluid inclusion study. Concentrations of other metal elements can be also estimated based on the procedure similar to that mentioned above.

50

Chapter 1

An example of the calculated results on the chemical compositions of Kuroko ore fluids is given in Tables 1.5 and 1.6. Because of uncertainties of equilibrium constants, NS, pH, temperature, fo2 and other parameters (activity coefficient, ionic strength, activity of water, pressure), the estimated values of concentrations may have uncertainties of-t-1 in logarithmic unit. However, it can be concluded from the thermochemical calculations and fluid inclusion data that the Kuroko ore fluids have the following chemical features. (1) Slightly acidic. (2) C1- concentration is similar to or higher than that of seawater. (3) Mg concentration is very low, compared with that of seawater. (4) The concentrations of base-metal elements and Ba are considerably higher than those of seawater. Several workers have intended to estimate the chemical compositions of Kuroko ore fluids based on the chemical equilibrium model (Sato, 1973; Kajiwara, 1973; Ichikuni, 1975; Shikazono, 1976; Ohmoto et al., 1983) and computer simulation of the changes in mineralogy and chemical composition of hydrothermal solution during seawater-rock interaction. Although the calculated results (Tables 1.5 and 1.6) are different, they all show that the Kuroko ore fluids have the chemical features (1)-(4) mentioned above. As will be discussed later, the experiments (Hajash, 1975; Mottl and Holland, 1978) and theoretical studies on seawater-rock interaction (Wolery, 1978; Reed, 1983) indicate that the Kuroko ore fluids characterized by (1)-(4) above are formed by seawater-crustal rock interaction at elevated temperatures.

TABLE 1.5 Chemical composition of Kuroko ore solution, estimated based on the solubility data of Kuroko forming minerals (Shikazono, 1985c)

pH Cl Na K Ca Mg Fe SO4

H2S Cu Zn Pb Au Ba Sr T (°C) fo2 (atom)

Kuroko ore solution

Seawater

4.5 :t_0.5 33000 ppm 12500 ppm 4000 ppm >500 ppm 1-10 ppm 300 ppm 10-4 m 5 × 10 3 m 5 ppm 20 ppm 20 ppm 10 3 ppm 30 ppm 1-5 ppm 250°C ~ 50°C

7.89 18800 ppm 10760 ppm 399 ppm 412 ppm I294 ppm 2 x 10 3 ppm 2712 ppm

10 - 3 5 ± I a t o m

0.3 x 10 . 3 ppm 3 x 10 3 ppm 0.03 x 10.3 ppm 0.005 x 10 ? ppm 0.33 x 10 3 ppm 7.9 ppm

51

Miocene-Pliocene Hydrothermal Ore Deposits TABLE 1.6

Concentrations of sulfur and metal species in the Kuroko ore-forming fluids of pH = 4.5 (Ohmoto et al., 1983) Species

ESO 2-

H2S £Fe ECu IBBa li3Zn EPb EAg H4SiO4

Maximum actual, or minimum value

Concentration (Iog molality)

200°C

250°C

300°C

350°C

Minimum Maximum Maximum Actual Actual Max mum Maximum Minimum Maximum Minimum Maximum Minimum Maximum Minimum Actual

-3.5 -2.0 -2.6 -4.3±0.6 -4.6±1.1 -5.7±0.2 -4.1 -3.0 -3.5±0.6 -4.5±0.6 -4.8±0.6 -5.8±0.6 -5.9±0.3 -6.9±0.3 -2.4

-3.5 -2.0 -2.4 -3.3±0.7 -4.4±1.1 -5.4±0.2 -3.9 -4.7 -3.0i0.7 -4.7±0.7 -4.3±0.7 -6.0~0.7 -4.7±0.4 -6.4i0.4 -2.2

-3.5 -2.0 -2.0 -2.4±0.7 -5.2±1.2 -5.2±0.2 -3.5 -4.2 -2.4±0.7 -4.4±0.7 -3.6±0.7 -5.6±0.7 -4.0±0.4 -6.0±0.4 -2.0

-3.5 -1.5 -0.6 -1.6±0.9 -2.6±1.3 -4.1±0.3 -4.1 -5.4 -2.2±0.9 -5.2±0.9 -2.5±0.9 -5.5±0.9 -2.7±0.5 -5.7±0.5 -1.9

1.3.3.4. Stable isotopes 6D and 6180. 3D and 3180 of the Kuroko ore fluids were estimated based on analyses of fluid inclusions, Kuroko-forming minerals and hydrothermal alteration minerals (e.g., Pisutha-Arnond and Ohmoto, 1983). Estimated 8D and 8180 of Kuroko ore fluids are plotted on 3D-3180 diagram (Fig. 1.40).

40

,

,

.

,

,

,

,

o

,

Kuroke Ore-forming Fl@ds

2O

I

SW ! ~z I Periods ii1,111& IV

Periods I%V

0 ~D -20 (%o) -40

- ~-.

B

I I I

~

ndary i n_s J

-60

"Magmatic"FLuid] ' ~

-80 -100

-12

-1o

I

-6

!

-4

I

I

-2 0 6180 (%4

|

2

I

4

I

6

i

8

10

Figure 1.40. Summary of 8D and 8180 values of the Kuroko ore-forming fluids and of low-temperature secondary inclusions (Ohmoto et al., 1983). SW: seawater.

52

Chapter I

They are -30%o to -3%~ for 5D and -2%o to +5%o for ~180. The origin of Kuroko ore fluids based on these data is discussed in section 1.3.4. (]34S ofsulj'qd¢$. A large number of ~348 data on sulfides are available. Figure 1.41 shows the summary of sulfur isotopic compositions of sulfides (chalcopyrite, sphalerite, galena, pyrite) (Shikazono, 1987b). Because pyrite exceeds 80% of total amount of sulfide sulfur in the deposit it is reasonable to assume that 534S of pyrite represents ~34S value of sulfides in Kuroko deposits. ~34S values of sulfide minerals from Kuroko deposits vary widely in a range from -6%0 to +9%o. Pyrite with low ~34S value (-6%0 to 0%o) is

<

GALENA

"5

=E

Z

1 -6-5-4-3-2

-

0

1

2

3

4

5

6

,<

7

8

9

10

~3~S(O/oo)

CHALCOPYRITE

"6

Z1 -6

-5

-4

-3

-2

-1

-0

1

2

3

4

5

6

7

8

9

10

~

1'0

9

1'0

~34S(% o)

SPHALERITE

"6

=E

Z

1

;/ 1

-6-5

4-3-2-1

-0

1

2

3

4

s

s

7

~

8~4S(O/c~)

PYRITE

,-= ICI E-I E-] -6

-5

-4

-3

-2

~

-1

.

-0

1

.

2

.

3

.

4

5

.

6

7

8

~3~S(O/c~)

Figure 1.41. Sulfur isotopic compositions of sulfide minerals from Kuroko deposits (Shikazono, 1987b).

53

Miocene-Pliocene Hydrothermal Ore Deposits

s3~s(%,o) -6

1

mudstones end turfs (ht-qtz zone) .(bar-rich zone) --

•

blaok-ores

-5

1

. i.i . . . . . . . . .

o I

+2

i

f

~= _

•--.

'~ I

i

+4 t

I

+6 t

1

+8 t

1

+10

+22

i

I

__-~_-2-=x_. . . . .

-

+ 24

f

i

~,,~ --

eo

•

bar

,,~ ,.~ ,."

r' ~ j

#

/

yellow-ores pyrite-ores veins and disseminated

~ ]

cp

~

,,, I,,PY

Miooene sea

ores

Figure 1.42. Sulfur isotopic variation and the vertical zonation of ores in the Shakanai No. 1 deposit (Kajiwara, 1971).

framboidal type (Komuro and Sasaki, 1985). The origin of framboidal pyrite is not well understood. However, it seems likely that this type of pyrite formed biogenetically and is not of hydrothermal origin. Most ~34S values of hydrothermal pyrite having euhedral morphology range from +2%0 to +7%0. The approximate ranges of variation of ~34S values of other sulfides are: chalcopyrite; +2%0 to +7%0; galena: 0%0 to +4%0; sphalerite: +2%0 to +7%0 (Sasaki, 1974). Sasaki and Kajiwara (1971) estimated average ~34S value of sulfide minerals to be +4.6%0 based on the measurements of mill concentrates from representative Kuroko deposits. In individual deposits, ~34S of sulfides generally increases stratigraphically upwards (Fig. 1.42). (Kajiwara, 1971). Based on the sulfur isotope evidence, Kajiwara (1971) deduced that the ore solutions underwent a progressive cooling and oxidation due to mixing with seawater. Sulfur isotopic data of separated pyrite as the commonest sulfide mineral (Kajiwara, 1971; Kajiwara and Date, 1971) show different ~348 values for the three sub-types of Horikoshi and Shikazono (1978). The ~34S values of pyrite in the C sub-type deposits are higher than the ~34S values of pyrite from the Y and B sub-types. The ~34S values of pyrite from the Y sub-type seem to be slightly higher than those from the B sub-type. Kajiwara and Date (1971) are of a different opinion: the 334S values from the Kosaka district are higher than those in the Hanaoka district, because all sulfur isotopic data from the C sub-type were obtained in the Kosaka district. The sulfur isotopic data on the obtained Uwamuki deposits of the B sub-type in the Hanaoka district indicate systematic decrease in ~348 passing from the yellow ore (+7%0) to the black siliceous ore (+5%o) (Bryndzia et al., 1983). Kajiwara and Date's data (1971) include three ~348 values of pyrite in the Doyashiki deposit of C sub-type in the Hanaoka district. The main Doyashiki

54

Chapter 1 5/+.9

5'

+40

SD

YU

+

AK +35

~.~.-30

TN06

CO

NK

"KS, f t '

•

rN

•.~e

I"IZ HT ll.,rr

''I +"H""",

V

+20 1<13 IN

OG

.15 *5

.10

+15

*20

so(sMow) Figure 1.43. Relationship between 3t80 and ~34S values of barite from some Kuroko and vein-type deposits. Abbreviations are: HN, Hirano; NK, Nagaki; AK, Akakura; YU, Yunosawa; SD, Sado; KY, Katsuyama; TN, Teine; OS, Osarizawa; MZ, Mitsuzawa; HT, Hata; OE, Oe; AI, Akaishi; KN, Kohinata; MT, Miyatamata; KZ, Karuizawa; FN, Funauchi; KO, Koyama; IK, Inakuraishi; and OG, Ogoya. "K.S.R." indicates "Karoko sulfate region" (Watanabe and Sakai, 1983).

deposit was, however, mostly mined out when their samples were collected in about 1969. Furthermore, the mine calls the main Doyashiki of the C sub-type and the satellite deposits of Y sub-type together with the Doyashiki deposits (Takahashi and Suga, 1974).

and 3180 of sulfates. 834S and 8180 of sulfates (anhydrite, gypsum, barite) of Kuroko deposits were reported by Sakai et al. (1970) who indicated that 834S and 8180 values of sulfates are very close to but slightly higher than Miocene seawater sulfate value (834S = +20%o to +21%o; 8180 = 0%o). After Sakai's pioneering work, Watanabe and Sakai (1983) and Kusakabe and Chiba (1983) analyzed large amounts of barite, gypsum and anhydrite for 834S and 8180 and confirmed the conclusion drawn by Sakai et al. (1970) (Fig. 1.43). t]34S

1.3.3.5. Radiogenic isotopes Lead isotopes. Sato and Sasaki (1973) concluded on the basis of a remarkable narrow range in lead isotopic composition of Kuroko ores that lead of Kuroko ore came from deep-seated source which originated from subducting pelagic sediments.

55

Miocene-Pliocene Hydrothermal Ore Deposits o

38.8"

central

38.6.

2ospb 2°4Pb 38.4,

o

Honshu.,,/~'.

o

•

i. ',

I o i

/"~--'~Hokuroku ,,4" j ores / ~ .~

382"

.p">.%~ j

Z

38.0.

...~/

' / I

•

I• o I"iI I /" 0 /" / / , ~ i

. / / I- "II /

~

I

[northern /\Honshu

/"

_.

o Kuroko • Epigenetic

t

rowth curve

207pb 15.6- !

2o4pb

~(7 o

15.5-

~

"?.'"~ "--£'o • /

•

/ ~ Honshu

15.4- centrcl| Honshu

18~

'

1~4

'

2o6pb/2O4pb

ld.6

Figure 1.44. Isotopic composition of lead in black ore (open symbols) and in yellow ore (closed symbols) in the Hokuroku district. The isotopic fields for black ore from the Fukazawa, Shakanai, and Kosaka deposits are outlined (Fehn et al., 1983).

Lead isotopic data on Kuroko deposits, vein-type deposits in Honshu and volcanic rocks are summarized and plotted in Fig. 1.44 (Fehn et al., 1983)• Although lead isotopic compositions of Kuroko ores occupy a narrow isotopic range, within a given ore deposit, black ore has a uniform isotopic composition but is significantly higher in radiogenic lead than yellow ore (Fig. 1.44, Table 1.7; Fehn et al. 1983). On the basis of the lead isotopic distribution, Fehn et al. (1983) concluded that a major part of the lead in Kuroko deposits was derived from igneous rocks, probably volcanic rocks with a significant contribution coming from the underlying preNishikurozawa formations and the yellow ore seems to have a greater igneous rock lead component than does the black ore.

Strontium isotopes. Strontium isotopic compositions (87Sr/86Sr) of anhydrite, gypsum and barite from Kuroko deposits are summarized in Fig. 1.45 (Farrell et al., 1978; Honma and Shuto, 1979; Farrell and Holland, 1983; Yoneda et al., 1993; Yoneda and Shirahata, 1995). 87Sr/86Sr values of anhydrite and gypsum are slightly lower than that of seawater, suggesting that most of the strontium was derived from seawater, but a small amount of

56

Chapter 1

TABLE 1.7 Isotopic composition of leads from Kuroko deposits (Fehn et at., 1983) Sample number

Field number

2°6pb/2°4pb

2°Tpb/2°4pb

2°8pb/2°4pb

W 1112132 K- 105 W 1112125

18.489 18.463 18.428

15.603 15.589 15.567

38.678 38.623 38.561 a

JU 78UM64 JU 78UM67

18.458 18.475

15.584 15.594

38.577 b 38.619 b

76-12-2-3 JU 7SUMI25 JU 78UM 129 CF 082713 CF 111163B CF 11 1163Y

18.469 18.464 18.470 18.473 18.491 18.434

15.594 15.5s9 15.594 15.589 15.604 i5.560

38.618 b 38.592 b 38.618 b 38.622 38.665 38.508 c

CF CF CF CF CF CF CF CF CF CF CF CF

18.550 18.580 18.568 18.565 18.566 18.558 18.560 18.575 18.561 18.546 18.542 18.518

15.621 15.622 I5.607 15.617 15.607 15.615 15.595 15.604 15.618 15.596 15.589 15.592

38.686 38.736 38.729 38.659 38.702 38.698 38.665 38.734 38.709 38.637 c 38.613 c 38.620 a

CF7 11282 CF 11286 S 11

18.527 t8.536 18.479

15.621 13.620 15.583

38.707 38.701 38.584 a

S 30

18.527

I5.617

38.624

S 35

18.539

15.598

38.683

MT9

18.451

15.567

38.517

M 8

18.448

15.585

38.378

Kosaka area

Uchinotai East 1 2 3

Uwamuki 2 4 5

Uwamuki 4 6 7 8 9 10 11 Fukazawa area

Tsunokakezawa 12 13 14 15 16 17 18 19 20 2I 22 23 Matsurnine-Sbakanai

11197 34511 l 11922 I 11810 34516 110197 082610B l 11913 111813 I1184 682610Y 111920

area

Shakanai 4 24 25 26

Shakanai 8 27

Shakanai 11 28

Matsumine 29

Matsuki 30

If not stated otherwise all samples were measured on galena from black ore. a Galena taken from yellow ore; b From Sato et al. (1981); c Yellow ore sample.

57

Miocene-Pliocene Hydrothermal Ore Deposits TABLE 1.7 (continued) Sample number Other

Field number

2°6pb/2°4pb

2°7pb/2°4pb

2°spb/2°4pb

HA 41

18.521

15.587

38.623

CF 11233

18.525

15.594

38.626

FU 17

18.556

15.603

38.662

deposits

Hanawa

31 Omaki

32 Furutobe

33

T

~10

[ ] Gypsum []

c ,< "6

Anhydrite

5

E "1 Z

1

,

.

.m..N.I

0.7070

N N N ~ ~

N N ,

0.7080

N N N

0.7090 87Sr/86Sr

Figure 1.45. Variation of the 87Sr/86Sr ratio of anhydrite and gypsum from Kuroko deposits (Farrell and HoIiand, 1983; Shikazono et aI., 1983).

strontium was from igneous rocks. 87Sr/86Sr ratios of barite are in a range from 0.706 to 0.708, suggesting smaller contribution of seawater strontium than anhydrite and gypsum. 87Sr/86Sr of barite from small Kuroko deposits (Iwami in San-in and Minamishiraoi in Hokkaido) are lower (Yoneda, et al., 1993; Honma and Shuto, 1979; Farrell and Holland, 1983) than that from large Kuroko deposits in Hokuroku district. This indicates that a large seawater circulation did not occur in a small Kuroko mine area and that seawater circulation is important for the formation of Kuroko deposits.

Rare earth elements (REE). Analytical results of REE contents of hydrothermally altered volcanic rocks in Kuroko mine area and Kuroko ores are summarized as follows (Shikazono, 1999a) (Fig. 1.46). (1) Positive Eu anomaly is observed for barite, Kuroko ores, ferruginous chert (tetsusekiei), and hydrothermally altered basaltic and dacitic rocks overlying the Kuroko ores. (2) Negative Eu anomaly is observed for hydrothermally altered dacite underlying the Kuroko ores and anhydrite in the dacitic tuff breccia.

58

Chapter I

Hydrothermally altered il dacite and anhydrite underlying the i Kuroko ores E lO0

Barite, Kuroko

i

ore ant

Hydrothermally altere basalt overlying the Kuroko ores

err i

i i

::

::

Ns.s~

i

NS-~S

::

!

i

~

sa2

NS-36

"-- NS-33 NS-39

NS.2

i La Ce

• Nd

i SmEu

i Tb

i i Yb L~

i La Ce

i SntEu

i Tb

i i Yb Lu

I La Ce

i Nd

i SmEu

i Tb

t t Yb Lu

Figure 1.46. REE patterns of the altered volcanogenic rocks and Kuroko ores. Data sources: Shikazono (1999a). (A) Hydrothermally altered dacite and anhydrite underlying the Kuroko ores. (B) Barite, Kuroko ore and ferruginous chert. (C) Hydrothermally altered basalt overlying the Kuroko ores (Shikazono, 1999a).

(3) N e g a t i v e Ce a n o m a l y is o b s e r v e d for h y d r o t h e r m a l l y altered chlorite-rich basalt o v e r l y i n g the K u r o k o ores. (4) N e g a t i v e Ce a n o m a l y and positive Eu a n o m a l y are o b s e r v e d for epidote-rich altered basalt near the orebody. (5) L i g h t rare earth e n r i c h m e n t is distinct and R E E contents are relatively high for the ferruginous chert. The R E E pattern for fresh v o l c a n i c rocks in the Kuroko m i n e area studied by Dud~is et al. (1983) is shown in Fig. 1.47 which shows no negative Ce and no positive Eu a n o m a l i e s and L R E E (Light Rare Earth E l e m e n t ) are not enriched c o m p a r e d with H R E E

1000

Sample / Chondrite

1000~ ~_S a m p l e / C h o n d r i t e

a

Basaltic Rocks

b

~

Acidic Rocks

I

lOO

~

,,,..

I0

1 La

!...... _ " : 2 _ _ _ -

10

Ce

Sm Eu

Tb

YbLu

La Ce

Pr

Nd

S m E u G d T b Dy Ho E r

Yb

Figure 1.47. REE patterns of the fresh volcanic rocks in the Kuroko mine area. Data source: Dud~s et al. (1983). (a) basalt; (b) acidic rocks (Shikazono, I999a).

Miocene-Pliocene Hydrothermal Ore Deposits

59

(Heavy Rare Earth Element). Therefore, it is considered that negative Ce and positive Eu anomalies in hydrothermally altered volcanic rocks, Kuroko ores, and ferruginous chert and LREE enrichment in the Kuroko ores have been caused by hydrothermal alteration and precipitations of minerals from hydrothermal solution responsible for sulfides-sulfate (barite) mineralization. Negative Eu anomaly is found for anhydrite sample and altered dacite underlying the Kuroko ores. One of the possible explanations for this negative anomaly is selective leaching of Eu by circulating hydrothermal solution of seawater origin. Sverjensky (1984) has shown from thermochemical calculations that Eu 2+ is more abundant than Eu 3+ under the reduced environment at high temperatures. Thus, it is considered that Eu is leached more efficiently from the rocks compared with the other rare earth elements. Alderton et al. (1980) have shown that Eu in the granitic rocks in southwest England is depleted due to the sericitization of feldspar-bearing assemblage. Negative Eu anomaly has been reported on highly silicified volcanic rocks around the volcanic rocks around the volcanogenic polymetallic massive sulfide deposit at Que River, Tasmania (Whitford et al., 1988). Thus, it is plausible that negative Eu anomaly of dacitic rocks was caused by the sericitization. This negative Eu anomaly indicates that the alteration minerals in the volcanic rocks underlying the Kuroko ores did not precipitate from ascending hydrothermal solution which interacted at relatively high temperatures (more than 250°C) and reduced condition (Eu2+/Eu 3+ is more than 1) and have positive Eu anomaly. Date et al. (1983) and Green et al. (1983) have shown based on numerous analytical data that Na, Ca and Sr were depleted from footwall dacitic rocks below the Fukazawa Kuroko deposits and this depletion was caused by the addition of K from ascending hydrothermal solution accompanied by the destruction of plagioclase, K-feldspar and volcanic glass and formation of sericite. This evidence supports the above interpretation: The selective leaching of Eu from footwall dacite was caused by the hydrothermal solution. The REE pattern for anhydrite is different from that of seawater, indicating that anhydrite did not precipitate due to the simple heating of seawater that was suggested by Sakai et al. (1970) and Sato (1973). This REE pattern could be explained in terms of the mixing of hydrothermal solution and cold seawater and low degree of seawater/hydrothermal solution mixing ratio (Shikazono et al., 1983). Negative Eu anomaly is also found in the fresh and altered dacitic rocks (Dudfis et al., 1983). Therefore this negative anomaly in anhydrite is also explained in terms of an influence of sericitization of dacite accompanied by the depletion of Eu. Sverjensky (1984) calculated the dependency of Eu2+/Eu 3+ in hydrothermal solution on f Q (oxygen fugacity), pH and temperature. According to his calculations and assuming temperature, pH and fo2 for epidote-stage alteration of basalt and Kuroko ores (Shikazono, 1976), divalent Eu is considered to be dominant in the rocks and hydrothermal solution. Thus, it is reasonable to consider that Eu in the rocks was removed to hydrothermal solution under the relatively reduced condition more easily than the other REE which are all trivalent state in hydrothermal solution. Thus, it is likely that Eu is enriched in epidote-rich altered volcanic rocks. Probably Eu was taken up by the rocks from Eu-enriched hydrothermal solution which was generated by seawater-volcanic rock interaction at relatively low water/rock ratio.

60

Chapter 1

A negative correlation between Mg content and Ca content of hydrothermally altered basalt and dacite from the Kuroko mine area exists. This correlation indicates that Ca in the rocks is removed to fluid by the exchange of Mg in seawater. Eu may behave in the manner similar to Ca during seawater-volcanic rock interaction because of the similarity of their ionic radii. Positive Eu anomaly is observed for hydrothermal solution issuing from the hydrothermal vent on the seawater at East Pacific Rise (Bence, 1983; Michard et al., 1983; Michard and Albarbde, 1986). Guichard et al. (1979) have shown that the continental hydrothermal barites have a positive Eu anomaly, indicating a relatively reduced environment. Graf (1977) has shown that massive sulfide deposits and associated rocks from the Bathurst-Newcastle district, New Brunswick have positive Eu anomalies. These data are compatible with positive Eu anomaly of altered basaltic rocks, ferruginous chert and Kuroko ores in Kuroko mine area having positive Eu anomaly and strongly support that Eu is present as divalent state in hydrothermal solution responsible for the hydrothermal alteration and Kuroko mineralization. Hydrothermally altered basalt in Kuroko mine area can be divided into chloriterich and epidote-rich one (Shikazono et al., 1995). Chlorite formed under the higher water/rock ratio and lower temperatures than epidote (Shikazono, 1984; Shikazono and Kawahata, 1987). At the stage of chlorite formation, a large amount of seawater cycled (seawater/basalt ratio, 40 by mass) (Shikazono et al., 1995). It is considered that Ce4+/Ce 3+ was high during the interaction of relatively low-temperature seawater and rocks at this stage. Ce is present as trivalent in the original rocks. With the proceeding of the reaction of oxidized fluid (modified seawater) with volcanic rocks, Ce is removed to fluid as Ce 4+ which is stable in oxidized fluid as complexes. Removed Ce 4+ may be incorporated into Mn-hydroxides and oxides (e.g., Goldberg et al., 1963; Courtois and Clauer, 1980). It is commonly observed that the weathered rocks exhibit a negative Ce anomaly (Ludden and Thompson, 1979; Menzies et al., 1979). Although these Mn minerals are not found in the Kuroko mine area, a positive Ce anomaly may be caused by the fixation of Ce 4+ in the altered rocks. The negative anomaly is interpreted in terms of the above consideration. On the contrary, the other rare earth elements (except for Eu) in the rocks are present as trivalent. Thus, it is likely that Ce 4+ is more soluble than the other rare earth elements by the reaction of oxidized seawater with rocks. Thus, the negative Ce anomaly of altered basalt could be interpreted in terms of low-temperature interaction of oxidized and relatively unreacted seawater having a negative Ce anomaly caused by the incorporation of Ce into Mn-hydroxides with rocks at late-stage of submarine hydrothermal activity. The REE data, combined with alteration minerals and concentration of major elements in hydrothermally altered rocks, could be used to reconstruct the structure and evolution of a submarine geothermal system accompanied by Kuroko mineralization (Shikazono, 1999a). At the stage of Kuroko mineralization, evolved reacted seawater enriched in Eu, Ca, and Sr formed at low seawater/rock ratio (ca. 1 by mass) and at relatively reduced condition (Eu2+/Eu 3+ greater than 1). Selective leaching of Eu, Ca and Sr occurred from the dacitic rocks underlying the Kuroko ores. The hydrothermal solution enriched

Miocene-Pliocene Hydrothermal Ore Deposits

61

in Ca, Eu and Sr issued from the discharge zone on the seafloor and sulfides and sulfates were precipitated from such hydrothermal solution due to the mixing with cold seawater. Positive Eu anomaly is strong and REE contents are low for the barite-rich ores, whereas positive Eu anomaly is weak, but REE contents are high for the ferruginous chert ore. This difference suggests that hydrothermal signal (low REE and positive Eu anomaly) was strong and was preserved at the stage of the precipitation of sulfides and barite and was weak and not preserved at the ferruginous chert stage. Probably the high REE concentrations of the ferruginous chert is due to the effect of bottom-seawater mixing causing scavenging of REE in cold seawater by Fe- and Mn-hydroxides and oxides forming the ferruginous chert. This view is consistent with the results obtained by Klinkhammer et al. (1983), Ruhlin and Owen (1986) and Olivarez and Owen (1989) who showed the scavenging of REE in seawater by Fe-oxides and Fe-hydroxides near midoceanic ridges. The effect of bottom-water mixing was discussed also for Archean iron formations and Red Sea metalliferous sediments by Barrett et al. (1988). The REE study indicates that the concentrations of REE, particularly Eu and Ce in altered rocks and ore minerals are useful indicators of oxidation state, intensity of discharging hydrothermal solution and evolutionary stage of submarine hydrothermal activity. 1.3.4. Depositional mechanism and origin of ore fluids

1.3.4.1. Depositional mechanism Some mechanisms of anhydrite deposition in Kuroko deposits. Shikazono et al. (1983) considered the depositional mechanism of anhydrite based on the mode of occurrence, texture, Sr content, nature of the contained fluid inclusions and isotopic composition of Sr, S and O in anhydrite together with the mineralogy of the sekko ore, combined with their experimental study on the patitioning of Sr between coexisting anhydrite and aqueous solution. The following is their discussion on the depositional mechanism of anhydrite. Several mechanisms could be invoked to explain the deposition of anhydrite in Kuroko deposits. These include: (1) recrystallization of gypsum and/or anhydrite of evaporite origin; (2) precipitation of anhydrite due to the cooling of hydrothermal solutions; (3) precipitation of anhydrite due to the boiling of hydrothermal solutions; (4) replacement of calcic minerals, such as feldspars, in volcanic rocks or of calcareous foraminifera in mudstones; (5) simple heating of seawater without interaction with country rocks, either above the seawater-sediment interface (Kajiwara, 1971) or beneath the seawater-sediment interface (Farrell et al., 1978; Farrell and Holland, 1983); (6) heating of seawater accompanied by interaction with country rocks; and (7) mixing of ascending hydrothermal solutions with seawater at a site either above the seawater sediment interface (Sato, 1973) or beneath the seawater-sediment interface. Several of these mechanisms can be ruled out based on the geologic environment of the Hokuroku district, because the geologic environment of the Hokuroku area during the Miocene was quite different from that of areas of evaporite formation. The precipitation of anhydrite from hydrothermal solutions has been studied extensively by various workers (e.g., Marshall et al., 1964a,b). The salinity of the inclu-

62

Chapter 1

sion fluids is less than ca. 5 wt% (~1 mol/kg H20) (e.g., Marutani and Takenouchi, 1978; Pisutha-Arnond and Ohmoto, 1983). Simple cooling of hydrothermal solutions is therefore virtually ruled out as a mechanism of anhydrite deposition. This also holds true for the proposal that anhydrite precipitation was due to boiling of the hydrothermal solutions. Studies of fluid inclusions in minerals from Kuroko deposits have yet to produce evidence of boiling in the inclusion fluids (e.g., Marutani and Takenouchi, 1978). Kumita et al. (1980) have investigated the distribution of foraminifera in the M1 and M2 mudstones that overlie the sulfide horizons of the Shakanai mines and have found that the relative abundance of calcareous foraminifera compared to arenaceous foraminifera decreases toward the sulfide orebody. Higher ratios of calcareous foraminifera to arenaceous foraminifera are found in mudstones stratigraphically higher than the thicker sekko body. This evidence suggests that some of the calcium in the anhydrites may have been derived from calcareous foraminifera in the mudstone. However, anhydrite generally occurs in the uppermost part of the tuff and tuff breccia of the T3 unit, which underlies the M2 and M I mudstones. The quantity of anhydrite in the mudstones is very small. This suggests that most of the calcium in the sekko anhydrites was not derived from the calcareous foraminifera distributed in the mudstones. Foraminifera are not found in the T3 tuff unit, and textures showing the replacement of foraminifera by anhydrite have not been observed. If seawater is simply heated, anhydrite begins to precipitate at approximately 110°C if PH20 is equal to Ptotal. The concentration of Ca 2+, Sr 2+ and SO42- in Miocene seawater was probably similar to that of present-day seawater (Graham et al., 1982). The partition coefficient of Sr between anhydrite and aqueous solution has been measured (Shikazono and Holland, 1983) as has the solubility of anhydrite in seawater (Marshall et al., 1964a,b); the variation in the Sr content of anhydrite with temperature during simple heating of seawater can therefore be calculated, using a value of 0.25 for the partition coefficient, Kdsr, experimentally determined by Shikazono and Holland (1983). Figure 1.48 shows that the Sr content of anhydrite which initially precipitated from seawater at about 110°C, should be about 1500 ppm; the Sr content of the anhydrites from Kuroko deposits is between 200 and 2,000 ppm (Fig. I. 19). The low Sr content of most of these anhydrites is difficult to explain in terms of the simple heating of seawater unless the experimentally determined values of KdSr (Shikazono and Holland, 1983) are not applicable to anhydrite deposition in Kuroko deposits. The relationship between the Sr content and the 87Sr/86Sr ratio of a number of anhydrites from Kuroko deposits is shown in Fig. 1.49. Most of the data points fall close to a trend line along which the Sr content of anhydrites increases with decreasing 87Sr/86Sr. The samples from the Fukazawa mines include sekko and paragenetically late vein anhydrites. The line marked A in Fig. 1.49 represents the position of anhydrites precipitated during the heating of Miocene seawater if KdSr had a value of 0.24 during the process of anhydrite precipitation in Kuroko deposits. The large separation between line A and the analytical data for the Kuroko anhydrites is striking. Miocene seawater was apparently not the only source of Sr in these anhydrites. Unfortunately, the processes by which the trend line of the anhydrite compositions was generated are still impossible

63

Miocene-Pliocene Hydrothermal Ore Deposits

3000

"6 E

~2000 c 0

0

1000

60

8

w

I

I

i

I

1O0 120 140 Temperature('C )

i

l

160

Figure 1.48. Change in the strontium content of anhydrite precipitated during the heating of normal seawater without any seawater-rock interaction (Shikazono et al., 1983).

87Sr/86Sr

0.709

R=0.1 0.2

0.3

0.4

0.5. .

. . .

A

B

C /

0.708 I

0.707

[l ./

lJ !a t it

1!

tl

0.706

0.705

lb!

5 0

i

1000

J

2000 S r c o n t e n t (pprn.)

Figure 1.49. Change of the strontium content and 87Sr/86Sr ratio of Kuroko anhydrite during the deposition and dissolution due to the mixing of hot ascending solution and cold solution (normal seawater) (Shikazono et aI., 1983). R mixing ratio (in weight) = S.W./(S.W.+H.S.) in which S.W. and H.S. are seawater and hydrothermal solution, respectively. Open triangle: Fukazawa deposits, Solid triangle: Hanawa deposits, Open square: Wanibuchi deposits, Solid square: Shakanai deposits. Concentration of Ca 2+, Sr 2+ and SO 2 of H.S. are assumed to be 1,000 pprn, 1 ppm, and 10 -4 mol/kg H20, respectively. Concentrations of Ca 2+, Sr 2+ and SO 2 of S.W. are taken to be 412 ppm, 8 ppm, and 2,712 ppm. Temperatures of H.S. and S.W. are assumed to be 350°C and 5°C (Shikazono et aI., 1983).

64

Chapter 1

to define uniquely. The anhydrites containing the lowest concentration of Sr have an isotopic composition close to that of Miocene seawater. This can be explained in three ways: (1) by a value of KdSr much lower than 0.24 during the precipitation of anhydrite in Kuroko deposits, (2) by the reaction of seawater with country rocks, and (3) by the mixing of seawater with one or more solutions in which the Sr/Ca ratio is much smaller than that of seawater. The first alternative is unlikely but not impossible. Past experience suggests that Kd Sr probably depends on the degree of supersaturation of the solutions with respect to calcite during calcite precipitation (Katz et al., 1972). In the experiments by Shikazono and Holland (1983), the solutions from which anhydrite was deposited were considerably supersaturated. It is therefore possible that the values of KdSr extracted from their experimental data are higher than those which controlled the incorporation of Sr in anhydrite during the formation of Kuroko deposits. Experiments at very low degrees of anhydrite supersaturation are needed to determine whether this is a possible explanation for the low Sr content of some of the Kuroko anhydrites. The reaction of seawater with country rocks is also a possible but unlikely explanation. Tertiary volcanic sediments in the vicinity of Kuroko deposits are altered and tend to have lost both Ca and Sr (Farrell and Holland, 1983). The ratio of Sr loss to Ca loss is roughly equal to the Sr/Ca ratio in seawater. If seawater was the altering medium, its St/Ca ratio was probably not strongly affected by the alteration process. The 87Sr/86Sr ratio would be intermediate between an initial value of 0.7088 and ca. 0.740 the 87Sr/86Sr ratio of unaltered Tertiary volcanics of the Hokuroku basin. It is unlikely, therefore, that this type of alteration can account for the Sr content and for the isotopic composition of Sr in the anhydrites at the upper end of the trend line in Fig. 1.49. On the other hand, mixing of seawater with solutions which have a St/Ca ratio much smaller than that of seawater could have led to the deposition of Kuroko anhydrites. Mixing with solutions containing high concentrations of Ca 2+ and very little Sr 2+ or SO] could lead to the precipitation of anhydrite whose Sr content reflects the low St/Ca ratio of the resulting mixtures. Their isotopic composition would be close to that of seawater. If the solutions that mixed with seawater contained low concentrations of Sr with an isotopic composition of ca. 0.7050, anhydrites precipitated from such mixtures would have tended to lie along a family of curves such as curve B in Fig. 1.49. The position of these curves depends on the concentration of Sr 2+, Ca 2+, and SO 2- in the mixing solution as well as on the isotopic composition of its Sr. It seems unlikely that mixing with a series of such solutions is a reasonable explanation for the trend of the data for Kuroko anhydrites in Fig. 1.49. It seems more likely that this trend was generated by the mixing of solutions from which the low Sr, high 87Sr/86Sr anhydrites were deposited with solutions characterized by a considerably higher Sr/Ca ratio and an SVSr/S6Sr ratio of ca. 0.7070. These solutions could have been the hydrothermal solutions from which the barites and presumably the sulfides in the several Kuroko mine areas were deposited. The 87Sr/S6Sr ratio of barites falls in the range 0.7069 to 07079 (Farrell and Holland, 1983): unfortunately the Sr/Ca ratio of the solutions from which the barites were deposited is poorly constrained. If two solutions containing Sr and Ca are mixed, and if no Sr or Ca containing phase is precipitated, the 87Sr/86Sr ratio of the mixtures is related to their -

-

Miocene-Pliocene Hydrothermal Ore Deposits

65

Sr/Ca ratio by the relationship:

{(87 Sr/86Sr) _ (87 Sr/86Sr)1 }/{(87 Sr/86Sr) _ (87 Sr/86Sr)2 }

,~{(mca2+/msr2+)-(mca2+/msr2+)l}/{(mca2+/msr2+)-(mca2+/msr2+)2

} (1-19)

where m denotes molality. The bar indicates the properties of the mixtures, and the subscripts 1 and 2 indicate those of the two end-member solutions. If the amount of anhydrite precipitated during mixing is sufficiently small so that the Sr/Ca ratio of the solutions is not thereby affected, the Sr content of the anhydrites will be related to the isotopic composition of the contained Sr by curves such as curve C in Fig. 1.49. If the quantity of anhydrite precipitated from the mixture is sufficiently large so that the Sr/Ca ratio of the solutions is affected significantly, the Sr content of the anhydrites of any given 87Sr/S6Sr ratio will be greater than those along the calculated mixing curve. The data for the concentration and the isotopic composition of Sr in the analyzed Kuroko anhydrites are generally consistent with such a mixing model. Unfortunately, the model is not unique, and additional chemical and/or isotopic constraints are needed to test its validity. Ohmoto et al. (1983) and Kusakabe and Chiba (1983) also reached the conclusion that the 8180 vs. 87Sr/86Sr relationship and ~34S vs. temperature relationship of barite from the Fukazawa deposit in the Hokuroku district may be explained by a mixing model with a seawater contribution of less than 20% at temperatures around 200°C. Sato (1973) and Ohmoto et al. (1983) calculated the amounts of sulfides precipitated due to the mixing of ascending hydrothermal solution and cold seawater. Their calculations showed that the calculated ratios of the amounts of minerals precipitated are generally consistent with those in Kuroko ore deposits. At early stage of mineralization, anhydrite formed by the mixing of seawater with hydrothermal solution which interacted in shallow part with volcanic rocks at seawaterdominated condition by decreasing Sr/Ca ratio of seawater but not so changing 87Sr/86Sr. In this stage Mg-chlorite formed from such hydrothermal solution dominantly originated from seawater. After this stage, sulfides, barite, and quartz precipitations took place at the sub-seafloor and on the seafloor by the rapid mixing of hydrothermal solution with cold seawater. Rapid precipitation of minerals occurred from the supersaturated solution under non-equilibrium condition. Equilibrium processes cannot explain the following two points. (1) ~34S values of coexisting sulfate and sulfides, and (2) barite/quartz ratio in orebody. If sulfur isotopic equilibrium between coexisting sulfates and sulfides was attained, using average ~34S values of sulfates and sulfides, +22%o and +5%o, respectively, we could estimate temperature using the equation by Ohmoto and Rye (1979). This temperature seems too high compared with temperature estimated from fluid inclusions and mineral assemblages (section 1.3.3). That means that sulfates and sulfides precipitated under the condition far from equilibrium. If hydrothermal solution in which H2S is dominant aqueous sulfur species and is free mixed with cold seawater in which high amounts of SO]- are contained,

SO]-

66

Chapter 1 6O 5O

• Black ore

40

• Yellow ore

30

• Siliceous ore L)

20

¢=

10 0-

•

0

#~ 10

~., 20

$ 30

=-40

, 50

. 60

70

Quartz Content [wt% ]

Figure 1.50. Quartz and barite content of Kurokoore. but no H2S, it is thermochemically predicted that both quartz and barite precipitate with increasing the cold seawater/hydrothermal solution ratio because solubility of quartz decreases with decreasing of temperature and that of barite decreases with increasing S O ] - concentration, which means decreasing of temperature. However, the barite and quartz contents of Kuroko orebody do not positively correlate with each other (Fig. 1.50).

Precipitation of barite and quartz. Barite and quartz are the most common gangue minerals in the submarine hydrothermal ore deposits such as Kuroko deposits and backarc basin deposits (e.g., Okinawa, Mariana deposits) (Halbach et al., 1989; Shikazono, 1994; Shikazono and Kusakabe, 1999). These minerals are also common in midoceanic ridge deposits. The observations of hydrothermal vents at midoceanic ridges and back-arc basins indicated that the minerals precipitate from the hydrothermal solutions which mix with cold ambient seawater. The mineralogical studies of the chimneys from these areas clarified that metastable phases (e.g., amorphous silica, wurtzite, marcasite, native sulfur) are common in the chimneys. The metastable phases are thought to have formed from the highly supersaturated solution. The degree of supersaturation (or saturation index) for barite in Kuroko deposits is estimated to be less than 100 (Shikazono, 1994). The degree of supersaturation with respect to quartz solubility is high because amorphous silica is precipitating from the solution. The above lines of evidence suggest that the precipitation of minerals in submarine hydrothermal ore deposits on the seafloor is taking place from the fluids with high flow rate at the orifices of the chimney (ca. 1-10 m/s) and with high degree of supersaturation under the non-equilibrium conditions. Gamo (1995) revealed based on the chemical and isotopic compositions of hydrothermal fluids from midocean ridges that the precipitation of minerals and interaction

Miocene-Pliocene Hydrothermal Ore Deposits

67

of fluids and sediments under the seafloor affect the chemistry of fluids discharging from the seafloor. Shikazono et al. (1983) indicated that the anhydrite in Kuroko deposits formed at subseafloor environments. Some applications of the coupled fluid flow-reaction model were carried out to the ore-forming process (e.g., Lichtner and Biino, 1992). However, a few attempts to understand quantitatively the precipitations of minerals from flowing supersaturated fluids in the submarine hydrothermal systems have been done (Wells and Ghiorso, 1991). Wells and Ghiorso (1991) discussed the silica behavior in midoceanic ridge hydrothermal system below the seafloor using a coupled fluid flow-reaction model. The behavior of silica and barite precipitation from the hydrothermal solution which mixes with cold seawater above and below the seafloor based on the thermochemical equilibrium model and coupled fluid flow-precipitation kinetics model is described below. As noted already, Kuroko deposits are characterized by the following zonal arrangement in ascending stratigraphic order: siliceous ore (quartz, chalcopyrite, pyrite), yellow ore (chalcopyrite, pyrite), black ore (sphalerite, galena, barite), barite ore (barite and quartz) and ferruginous chert ore (microcrystalline quartz, hematite). Quartz is abundant but barite is poor in the siliceous ore, though barite veinlets occur in this zone. Barite is common in the black ore and abundant in barite ore. Barite is also found in ferruginous chert ore (Kalogeropoulos and Scott, 1983). Quartz is poor in massive sulfide ore horizons (yellow and black ores). The distribution of quartz and barite in Kuroko deposits suggests that barite and quartz precipitate by different mechanisms. Figure 1.50 shows the relationship between barite and quartz contents of the Kuroko ore samples. This suggests that quartz and barite formed separately in different parts (Ohmoto et al., 1983). Amorphous silica and barite precipitate simultaneously from white smoker in midoceanic ridge hydrothermal system (Edmond et al., 1979). It is inferred that amorphous silica precipitates in the chimney at a later stage than sulfides and sulfates (anhydrite and barite) which constitute chimneys from which black smoker is emerging. It is thought that the precipitation of amorphous silica is caused by conductive cooling from the hydrothermat solution which flows laterally in the chimney (Herzig et al., 1988). Barite is abundant in back-arc basin hydrothermal system such as Okinawa, Manus and Mariana (Shikazono and Kusakabe, 1999). In these chimneys, coprecipitation of barite and amorphous silica is taking place from the solution characterized by lower temperatures and lower flow rate than the black smoker. Solubilities of quartz and amorphous silica in aqueous solutions increase with increasing of temperature (Holland and Malinin, 1979). Solubility of barite depends on salinity and temperature (Blount, 1977). The solubility of barite in hydrothermal solution having more than 1 molal NaC1 concentration increases with increasing temperature, while a solubility maximum exists in the solution with NaC1 concentration less than ca. 0.2 molal (Blount, 1977).

68

Chapter 1

==

./

<

om

0

1O0

200

300

400

500

600

Precipitated Amount of Quartz (mg/kg • H20) Figure 1.51. Relationship between precipitated amount of quartz and that of barite.

Salinity of ore fluids responsible for Kuroko deposits is in a range of 0.5-1 tool/1 which is estimated based on the fluid inclusion studies (e.g., Marutani and Takenouchi, 1978). Thus, it seems likely that the precipitation of barite takes place by decreasing of temperature and increasing of concentration of sulfate ion caused by the mixing of hydrothermal solution having 0.5-1 molal NaC1 with cold ambient seawater. Quartz precipitates also due to the mixing of hydrothermal solution with ambient cold seawater if equilibrium between solution and quartz is attained. The thermochemical equilibrium calculations on the amounts of minerals precipitated due to the mixing of ascending hydrothermal solution with cold seawater have also shown that the amounts of quartz precipitated correlate to that of barite precipitated (e.g., Ohmoto et al., 1983; Janecky and Seyfried, 1984; Bowers et al., 1985). Previous studies on the precipitations of barite and quartz calculated the saturation index with respect to barite and quartz during the mixing of hydrothermal solution and seawater. However, the amounts of barite and quartz precipitated were not obtained. Therefore, the precipitated amounts due to the mixing was calculated. The relationship between the amounts of barite and quartz precipitated is shown in Fig. 1.51, which indicates that these amounts are positively correlated. However, as ah'eady noted, the barite content in Kuroko ore inversely correlates to the quartz content and the occurrences of barite and quartz in the submarine hydrothermal ore deposits are different. The discrepancy between the results of thermochemical equilibrium calculations based on the mixing model and the mode of occurrences of barite and quartz in the submarine hydrothermal ore deposits clearly indicate that barite and quartz precipitated from supersaturated solutions under non-equilibrium conditions. Thus, it is considered that the flow rate and precipitation kinetics affect the precipitations of barite and quartz. It was attempted to derive the relationships in the precipitated amounts of barite and quartz, flow rate and precipitation rate using the coupled fluid flow-precipitation

69

Miocene-Pliocene Hydrothermal Ore Deposits

nit J

I

Mixing

Precipitation

/

q C, Figure 1.52. Uniformly mixed kinetics model, q: volume flow rate (m3/s), Ci: initial concentration (molal), l: height (m), r: radius (m), C: concentration (molal), V: volume of the system (m3). model. T w o g e o l o g i c sites o f the precipitations o f barite and quartz are taken into account: one is the site a b o v e the seafloor and the other one is that b e l o w the seafloor. The calculations w e r e m a d e based on the u n i f o r m l y m i x e d precipitation kinetics m o d e l (Fig. 1.52) for one c o m p o n e n t and one d i m e n s i o n a l system r e p r e s e n t e d by Eq. (1-20).

dC / dt = - k ( a / m ) ( c - Co) + (q / V) ( Ci -- C)

(1-20)

w h e r e k = precipitation rate constant, A = surface area on w h i c h barite and quartz precipitate, M = mass o f h y d r o t h e r m a l solution, C = c o n c e n t r a t i o n o f Ba 2+ and H4SiO4 o f the solution in the system, Ci: the c o n c e n t r a t i o n o f input solution, q = v o l u m e flow rate, and V = v o l u m e o f the system. C a l c u l a t e d results are shown in Figs. 1.53 and 1.54. A s s u m e d values o f parameters (l, r, A / M ) for the calculations are g i v e n in Table 1.8. A c o m p a r i s o n o f the calculated results (Figs. 1.53 and 1.54) with the m o d e o f o c c u r r e n c e s o f quartz and barite in the s u b m a r i n e h y d r o t h e r m a l ore deposits indicates TABLE 1.8 Parameter values used for the computations l (m)

r (m)

A/M (m2/kg)

q/V (l/s)

1

1

2 3

1 0.1

0.01 0.1 0.1

0.2 0.02 0.02

v 10 v

v

v: velocity (m/s); A: surface area (m2); M: mass of fluid (kg); q: volume flow rate (m3/s); V: volume of system (m3).

Chapter i

70

0.16

I

0.12

----

l=lm,r=0.01m l=lm,r=0.1m I I--0.1m,r=-0.1m CBai

O

<

~.~ 0.08

X

0.04

0

~"

\

\

~

\

-,,

i

2

. . . i . . . .T . . . " r '.- . .---..-t. . . . i,. . . .¢--..

3

4

5

6

7

8

9

10

Velocity(m/s) Figure 1.53. Relationship between precipitated amount of barite and velocity of fluid at 200°C. v: velocity (m/s), CBai: initial concentration of Ba. 14 12

l=Im,r=O.Olm I l=Im,r=O.Im I l=O.Im,r=O.Im Cs o2i

- - - -

10

81',` "~

4

21| 0 1

-8

\

\",,~

\\\ \ i" - ~

-7

-'-'~...,-,---

-6

- ".~. ~ - -

-5

--~

. . . . .

-4

L------.J

-3

. . . .

-2

-1

Log 1~ Figure 1.54. Relationship between precipitated amount of quartz and velocity of fluid at 200°C. Csio2i: initial concentration of SiO2; v: velocity (m/s).

Miocene-Pliocene Hydrothermal Ore Deposits

71

that the precipitations of barite and quartz take place in different sites in the submarine hydrothermal system. Quartz tends to precipitate under the conditions of high temperature, high A/M, and slow fluid flow rate, while barite under the conditions of low temperature, low A/M, and high fluid flow rate. These results seem to be in agreement with the mode of occurrence of barite and quartz. Barite in the black ore is thought to have formed from the fluids with high flow rate (1-10 m/s) on the seafloor. Barite is found in the chimney from which black smoker is venting. On the other hand, quartz formed in the siliceous ore in high A/M, replacing the dacite below the seafloor. Ferruginous chert in which abundant silica occurs formed below the seafloor by the mixing of ferruginous sediments and hydrothermal components (Kalogeropourous and Scott, 1983). Barite-silica chimney found in back-arc basin formed in the conditions similar to that of ferruginous chert and barite bed in the Kuroko deposits; temperature is relatively low (ca. 150-100°C), and flow rate of fluids may be slow. The above mentioned study on barite and quartz precipitations in Kuroko deposits is summarized as follows. (1) Quartz occurs abundantly in feeder ore (stockwork siliceous ore) in Kuroko deposits. (2) Quartz coexisting with barite also occurs in the ferruginous and barite ores in Kuroko deposits. (3) Barite is abundant in the massive strata-bound ore bodies (black and barite ores) in Kuroko deposits and occurs in the ferruginous chert ore in Kuroko deposits, and chimneys in active deposits at back-arc basins. The coupled fluid flow-precipitation kinetics model calculations indicate the following results: (1) Quartz or amorphous silica tends to precipitate from the solution having relatively high temperature and low flow rate and under high A/M condition. (2) Barite tends to precipitate from the solution with relatively high flow rate and low temperature and under low AIM. (3) These predictions are generally in agreement with the observations; homogenization temperatures of fluid inclusions in quartz from siliceous ore zone and in barite from black ore zone in the Kuroko deposits is relatively high, ranging from 350 to 250°C, and tow, ranging from 250 to 150°C, respectively. (4) The results of calculations are in agreement with the occurrences of barite and silica and chemical features of discharging fluids in the submarine hydrothermal ore deposits; namely, quartz is inferred to precipitate in subseafloor environment and barite in seabottom environment. A/M for the stockwork siliceous ore zone in Kuroko deposits is low, while that for black ore in the Kuroko deposits and chimney may be high. The above-mentioned consideration indicates that important factors controlling the precipitations of barite and silica are surface area/water mass ratio (A/M), temperature, precipitation rate constant (k) and flow rate (v), and the coupled fluid flow-precipitation models are applicable to understanding the distributions of minerals in submarine hydrothermal ore deposits.

72

Chapter I

Barite precipitation highly depends on SO ] . and Ba 2+ concentrations in the fluids. That means that the mixing ratio of hydrothermal solution and seawater is also an important factor for the precipitation of barite, together with the factors mentioned above.

Precipitation mechanism of barite. The precipitation mechanism of barite from aqueous solutions at temperatures below 100°C have been extensively studied by many workers (e.g., Elving and Leineweber, 1950; Wagner and Wuellner, 1952; Nielsen, 1955, 1957, 1958, 1959a,b, 1961; Nielsen and Tort, 1984; Collins and Leineweber, 1956; Takiyama, 1959a,b; Fabrikanos and Lieser, 1962; Lieser and Wertenbach, 1962; Nancollas and Liu, 1975; Klein and Frontal, 1964; Blount, 1974; Liu et al., 1976; Rizkalla, 1983). However, barites in back-arc basin deposits precipitated in the temperatures more than 100°C, such as 150-300°C in Kuroko deposits (e.g., Pisutha-Arnond and Ohmoto, 1983). Thus, Shikazono (1994) carried out the experiments of barite precipitation at elevated temperature (150°C). It was found from his experiments that the morphology of barite varied with the barium chloride concentration. Dendritic crystals with rod-like, spindlelike, and star-like or cross-like habits, and irregularly shaped crystals with rough surface formed from solutions of high barium chloride concentrations (BaCl2, 0.08-0.8 molal). Feather-like dendritic crystal did not form. The results obtained by Shikazono (1994) are generally in agreement with those obtained by previous investigations in their barite precipitation experiments at less than 100°C. For example, Liu et al. (1976) found that dendritic crystals of barite formed at 25°C from aqueous solution containing BaSO4 higher than 2.0 × 10 - 3 molal, which is in agreement with the result by Lieser and Wertenbach (1962). They synthesized feather-like barite from aqueous solutions containing 2.0 x 10.2 molal BaSO4. Suito and Takiyama (1954) and Takiyama (1959a) reported that spindle-shaped dendritic barite crystals having ragged edges formed from aqueous solutions having 1.0 x 10 2 to 2.0 x 10-2 molal BaSO4. Sasaki and Minato (1983) reported that barites precipitated at 60°C from an aqueous solution having the total cation (Ba(NO3)2 + Pb(NO3)2) concentration greater than 0.004 molal were very fine-grained, and complex dendritic particles crystals, while rectangular (Ba,Pb)SO4 crystals formed from a solution of lower concentration. From the values of BaCl2 concentrations and the dissociation constant of BaC1+ (Helgeson and Kirkham, 1976), the concentrations of Ba 2+ in the experimental solutions are estimated to be 0.034-0.047 molal. Assuming the above values of SO42 and Ba 2+ concentrations, the boundary condition for the formation of dendritic barite and wellformed barite is estimated to be - 6 . 2 to 6.0 for the log(mBa2+)i(mso2)i value at 150°C (Fig. 1.55). The log(mBa2+)i(mso2-)i values for the solutions of the previous investigators (T = 25-100 C) and observed morphologies of barite m their experiments) are compared as shown in Fig. 1.55. It shows that the morphology of barite crystals changes with an increase in the concentration product, (mBa2+)i(mso2)i, from well-formed (rectangular, rhombohedra) (R) through rod-hke, star-hke, and spmdle-hke dendrmc (D) to feather-hke dendritic crystals (DF). The concentration products obtained in the experiment by Shikazono (1994) can be compared with the solubility product for large well-formed polyhedral barite crystals o

4

.

•

.

.

.

.

,4

.

.

.

.

.

.

73

Miocene-Pliocene Hydrothermal Ore Deposits -2

DF

-3

DF

-4

D

DF

D D D

0-~

D

D

m 0"~-8

0 __1

-9 -10 [

~o

I

,oo

t

,5o

i

200 V (°(3)

Figure 1.55. The relationships between the concentration product, (Ba2+)i(SO42-)i,at the initiation of barite precipitation, and morphologies of barite crystals (Shikazono, 1994). The dashed line represents the boundary between dendritic barite crystals and well-formed rhombohedral, rectangular, and polyhedral barite crystals. The 150°C data are from Shikazono (1994); the others from other investigations. D: dendritic (spindle-like, rodlike, star-like, cross-like) barite; DF: feather-like dendritic barite; W: well-formed rectangular, rhombohedral, and polyhedral barite. (i): The boundary between the diffusion-controlled mechanism (Di) and the surface reaction mechanism (S) for barite precipitation at 25°C estimated by Nielsen (1958); ®: The solubility product for barite in 1 molal NaC1 solution at I50°C based on data by Helgeson (1969) and Blount (1977). A-B: The solubility product for barite in 1 molaI NaCI solution from 25 to 150°C based on data by Helgeson (1969). adopted by Helgeson (1969) and experimentally obtained by Blount (1977). The boundary determined by Shikazono (1994)'s study (log(mBa2+)i(mso2)i = - 6 . 1 ± 0 . 1 at 150°C) 4. lies above the concentration products in 1 molal NaC1 solunon by Helgeson (1969) and Blount (1977) which are plotted as two (the concentration product based on the Blount (1977) experimental results is - 7 . 3 in log unit) for 150°C) and the curve A - B from 25 to 150°C (Fig. 1.55). The above comparison indicates that the solubility of dendritic barite crystals is higher than that of well-formed polyhedral large barite crystals used by Blount (1977). The saturation index for the boundary between dendritic barite and well-formed barite can be estimated to be 10 - 6 ' I / 1 0 .7.3 which is equal to ca. 20. It is usually believed that the growth o f dendritic crystals is controlled by a bulk diffusion-controlled process which is defined as a process controlled by a transportation of solute species by diffusion from the bulk of aqueous solution to the growing crystals (e.g., Strickland-Constable, 1968; Liu et al., 1976). The appearances of feather- and star-like dendritic shapes indicate that the concentrations of pertinent species (e.g., Ba 2+, SO 2 - ) in the solution are highest at the corners of crystals. The rectangular (orthorhombic) crystal forms are generated where the concentrations o f solute species are approximately the same for all surfaces but it cannot be homogeneous when the consumption rate o f solute is faster than the supply rate by diffusion (Nielsen, 1958).

74

Chapter 1

The dependency of rate of precipitation of barite from aqueous solution on time at room temperature studied by Nielsen (1958) suggests that the precipitation of barite from solutions of high levels of supersaturation (i.e., more than 30 as saturation index (S.I.) which is defined as the ratio, (mBa2+)(mso2-)/Ksp, where Ksp is the solubility product m molahty for equlhbrlum) is controlled 15y a bulk diffusion mechanism, while the surface reaction mechanism (polynuclear growth) dominates at low S.I. (i.e., less than 30). The boundary between the bulk diffusion mechanism and the surface reaction mechanism (polynuclear growth mechanism), suggested by Nielsen (1958) at 25°C, is shown in Fig. 1.55 ((~ in the figure). His estimated boundary is consistent with that between dendritic and well-formed barite which was determined by scanning microscopic measurements by other workers (e.g., Fisher and Rhinehammer, 1953; Okada and Magari, 1955; Liu et al., 1976). Nielsen and Toft (1984) summarized the relationship between precipitation mechanism of sparingly solute electrolytes (e.g., BaSO4, AgC1, CaF2) and the degree of supersaturation, distinguishing the surface reaction-controlled and diffusion-controlled regions by plotting growth rates of barite on a PBa-PSO4 diagram (where PBa log EBa; EBa is total dissolved barium concentration; Pso4 log ESO4, a total dissolved sulfate concentration). Nancollas and his group (e.g., Nancollas and Liu, 1975) have suggested second order kinetics for barite precipitation at 25°C with a surface-controlled mechanism even at an S.I. as high as 56. These studies have demonstrated that the surface reaction mechanism dominates to the S.I. of ca. 10-100 and that above this S.I. the bulk diffusion mechanism controls the precipitation of barite. This boundary between two different precipitation mechanisms for barite determined by the experiments at 150°C (Shikazono, 1994) roughly coincides with that between dendritic crystals and well-formed crystals which has been experimentally determined at temperatures lower than 100°C. Several morphologies of dendritic barite, such as feather-like, rod-like, spindlelike, star-like, and cross-like crystals have been recognized when the (mBa2+)i(ms02-)i values were considerably higher than the equilibrium values, although no detailed studies have been made on the relationship between the various morphologies of dendritic barites and the degree of supersaturation. Barite is abundant and widespread in Kuroko deposits. However, it is concentrated especially in the upper horizons (black ore and barite ore), increasing upwards within the black ore. Such a trend is also observed frequently in the tetsusekiei, but rarely in the yellow ore. Barite occurs also as vein-fillings in the stockwork siliceous ore. There are four main types of barite occurrence. Type A is well-formed coarsegrained barite in the black and yellow ores, associated with sulfide minerals (sphalerite, galena, chalcopyrite, pyrite). Fine-grained barite (type B) is also associated with sulfides in the black and yellow ores. Fine-grained barite tends to occur in the upper parts of black ore and also in yellow ore. According to Eldridge et al. (1983), fine-grained sulfides and barite were primitive, rapidly precipitated due to the mixing of ascending hot solution with cold seawater at the seafloor (Sato, 1973; Eldridge et al., 1983). They converted to the coarser-grained barite through dissolution and recrystallization within the ore piles. Fine-grained barite tends to occur in black ore, while coarse-grained barite is found in the .

.

.

.

.

.

4

=

=

- -

- -

Miocene-Pliocene Hydrothermal Ore Deposits

75

lower horizons of black ore and also in yellow ore. Type C is fine-grained barite in the massive barite ore horizon. Under an optical microscope, type C appears as an aggregate of radial crystals, intergrowing with fine-grained quartz. Type C is more elongated in shape compared to types A and B. Most barites (types A, B, and C) are rectangular, lath-shaped, platy, and tabular. Rare type D barite crystals are rectangular and polyhedral. Polyhedral, rhombohedral, or coarse-grained barite crystals (type D) are rarely observed in vugs of black ore or in the inner parts of the chimney recovered from the Hanaoka mine (Shimazaki and Horikoshi, 1990; Shikazono, 1992). The various barite types were examined by scanning electron microscopy (SEM). No dendritic (feather-shaped, spindle-like, rod-like, star-like, or cross-like) barite was found. Coarse-grained barite (type A) and fine-grained barite (type B) have generally rectangular shape but sometimes irregular shapes with rough surfaces. The surface of type B is rougher than that of type than that of type A. Type A and B barites are comprised of very fine-grained, disk-like barite microcrystals. The grain size of an individual microcrystal ranges from ca. 0.1 to 5 Ixm. Massive barite crystals (type C) are also composed of very fine grain-sized (several txm) microcrystals and have rough surfaces. Very fine barite particles are found on outer rims of the Hanaoka Kuroko chimney, while polyhedral well-formed barite is in the inner side of the chimney (type D). Type D barite is rarely observed in black ore. These scanning electron microscopic observations suggest that barite precipitation was controlled by a surface reaction mechanism (probably surface nucleation, but not spiral growth mechanism) rather than by a bulk diffusion mechanism. There are two interpretations for the formation of fine disk-like barite particles which constitute types A, B, and C barites. One is homogeneous nucleation caused by rapid mixing of hydrothermal solution with cold seawater and the coagulation of resulting fine particles. The other is heterogeneous nucleation and growth of barite on the surfaces of pre-existing barite crystal. Calculations of settling velocities for very fine particles in hydrothermal plumes issuing from the ocean floor suggest that very fine (less than 10 g m in diameter) particles cannot settle onto the nearby ocean floor (Shikazono, 1992). However, barite microparticles with grain sizes less than 10 ~ m are common in Kuroko deposits. Therefore, heterogeneous nucleation is likely rather than homogeneous nucleation. Sulfate-sulfide chimneys have been recently discovered from the Mariana back-arc basin (e.g., Kusakabe et al., 1990). They are composed of barite, sphalerite, galena, chalcopyrite, and silica. Barite is most abundant in the chimney. Scanning electron micrographs of barites in a Kuroko chimney show that they are well-formed polyhedra. Morphologically similar barite is also found in the inner side of a chimney from the Mariana back-arc basin. Such barites appear to have formed by recrystallization of fine barites which occur in the outer side of the chimneys. The various morphological features of barites from the Kuroko and Mariana deposits, when combined with the experimental studies on barite precipitation, suggest that the surface reaction mechanism was dominant for the formation of these barites. This implies that the concentration product, (mBa2+)(mso2-), at the initiation of barite precipitation was probably less than ca. 100 times that for equilibrium. This estimate can be evaluated based on the strontium and sulfur isotopic studies

-..-1

TABLE 1.9 Chemical composition of hydrothermal solution experimentally interacted with rocks

T (°C) P (bar} W/R Day's pH Na K Ca Mo Fe Mn Si AI CO2 SO4 H2S Cu Ni Zn Pb Au Hg As Cd Sb Ba Sr

1 (ppm)

2

3

4

5

6

7

8

9

10

11

12

13

14

350

350

455 764 349 3.7 14.9 288 0.04 1475 250 <0.1 0.02 0.3 l 0.17 0.01 0.017 0.02 0.7 0.02 0.04 1.2

1561 1330 2 5.6 8.8 1154 0.05

202 1,5 10.5 8.9 184(I 4.1

4(I(I 1000 5 14 3.7 10750 2300 313 9.0 62.5 54 1275

500 1000 5 14 3.0 10250 460(1 370 5 490 76 1200

500 1000 5 ~4 1.0 11250 950 310 250 7.1 7.1 1500

500 1000 5 14 3.5 8250 1860 1530 14.5 1100 120 1300

500 1000 5 14 3.5 10300 425 1100 7.6 385 70 1100

300 700 3 236 5.35 11084 726 1856 5.6 2.0 2.7 416 <0.26 159 1.39 6.60 <0.05 <0.07 <0.15

400 700 3 ~00 3.85 l 1156 762 1565 10.5 117 53 1684 <(}.26 573 5.65 7.54 <(}.05 <0.07

262 500 50 100 4.65 10300 370 1420 351 45 110 1350 0.21

-

10 16 4.8

300 10(,10 5 t4 3.0 750(I 1150 96 345 25.5 4.6 650

300

10 124 5.2

300 1000 09 16 4.2

2.0 5.5

2.1 6.9

10 40 0.11 0.7(I 1.5 (1.05 0.052 0.06 0.6 0,09 0.08 2.1

1.5

4.1

1 2: graywacke; 3-7: rhyolite: 8: andesite; 9-13: basalt; 14: seawater.

1.1

92 5.6 10963 489 818 0.l 3.5 0,05 580 0.37 118 20 10 0,01 0.01 0.1

0,135

0.50

0.38

<0.15

7.89 10760 399 412 1294 2 x 10 3 0.2 x 10 -3 2900 x 10 3 2 x 10 - 3 145 2717 0.3 x 10 3 0.6 x 10 3 3 x 10 -3 0.03 × 10 -3 0,005 × 10-3 0.1 x 10 3 3 x 10 _3 0.1 x 10 -3 0.33 x 10 -3 20 x 10- 1 7.9 ~"

Miocene-Pliocene Hydrothermal Ore Deposits

77

of barite and strontium content of anhydrite from Kuroko and Mariana deposits by Shikazono et al. (1983), Kusakabe and Chiba (1983), and Kusakabe et al. (1990). These authors indicated that these sulfate minerals precipitated due to the mixing of the hydrothermal solution with cold seawater with a contribution of the end member hydrothermal fluid to seawater being less than 20%. Ohmoto et al. (1983) calculated S.I. of barite to be less than ca. 101'5 during the mixing and cooling of fluids. The estimated value is in agreement with that of Shikazono (1994).

1.3.4.2. Origin of ore fluids Origin of ore fluids is constrained by (1) chemical compositions of ore fluids estimated by thermochemical calculations (section 1.3.2) and by fluid inclusion analyses, (2) isotopic compositions of ore fluids estimated by the analyses of minerals and fluid inclusions (section 1.3.3), (3) seawater-rock interaction experiments, (4) computer calculations on the seawater-rock interaction, and (5) comparison of chemical features of Kuroko ore fluids with those of present-day hydrothermal solutions venting from seafloor (section 2.3). During the last two decades, many experimental studies on the seawater-rock interaction at elevated temperatures (100~00°C) have been conducted. Particularly, detailed seawater-basalt interaction experiments have been done. Several experimental studies on seawater-rhyolite interaction and seawater-sedimentary rock interaction are also available (Bischoff et al., 1981). Examples of chemical compositions of modified seawater experimentally interacted with various kinds of rocks are shown in Table 1.9. Several factors such as C1- concentration, water/rock ratio and temperature are important in controlling the chemical composition of the hydrothermal solution interacted with the rocks. For example, water/rock ratio affects the alteration mineralogy (Mottl and Holland, 1978; Seyfried and Mottl, 1982; Shikazono, 1984). For example, at low water/rock ratio, epidote is stable, while chlorite at high water/rock ratio (Shikazono, 1984; Shikazono and Kawahata, 1987). From rock-water interaction experiments (Table 1.9) and analytical data on fluid inclusions we can derive the relationship between the concentrations of alkali, alkali earth and base-metal elements and concentration of C1- ion in the hydrothermal solutions experimentally interacted with rocks, in the natural hydrothermal solutions and in fluid inclusions (Figs. 1.56-1.58). It is seen in Figs. 1.56-1.58 that the concentrations of elements increase with increasing of C1- concentration. Especially base-metal (Zn, Fe, Mn, Pb) and Ba concentrations increase rapidly with increasing of C1- concentration. This relationship strongly suggests that C1- concentration is a very important factor controlling the chemical compositions of ore solution. Detailed discussion on the relationship between C1- concentration and concentrations of alkali and alkali earth elements were carried out by Shikazono (1978a) (see section 2.1). In Fig. 1.59 the relationship between temperature and concentration of elements (Zn, Ba) at constant C1- concentration which is equal to that of seawater obtained by the experimental studies and analytical data on natural hydrothermal solution (geothermal water) are shown. It is seen that the concentrations of base-metal elements (Zn, Fe, Mn, Cu, Pb) and Ba increase with increasing of temperature. Concentrations of these

78

Chapter 1

Log(Zn) 3

OOX

2 1 x

0

x

x •

x x

x

-;

x x

-3

x

-Z

:;

o

:3

4

,5 Log(CI)

Figure 1.56. Relationship between the zinc and CI- concentration in geothermal waters and hydrothermal solution experimentally interacted with rocks (Shikazono, 1988c).

Loq(Fe)

,"

3 o

G

,"

• sS

-1

~

:2[ -2

o0

-3

0

1

,

st

,,~f

," X

o0

I,'* I

;, ,

,o . ,

X

,~I ' /s " X

1 0

:

,"

~. sesX

!

x, " "

S

•

""

•

e °

oO S.W

:;

:3

4

5

Log(Cl)

Figure 1.57. Relationship between the iron and C] concentration in geothermal waters and hydrothermaI solution experimentally interacted with rocks (Shikazono, 1988c).

79

Miocene-Pliocene Hydrothermal Ore Deposits

Log (Ba)

CI = S.W.

4 3 AR.

2

% %

gR.

1 0

~X

X

0

%

-1

%

%

%

%

-2

%

%

%

%

-3

%

SW

-4

800

1000

600

400

200

T(°C)

0

Figure 1.58. Relationship between Ba2+ concentration and temperature of geothermal waters and hydrothermal solution experimentally interacted with rocks (AR: acidic rocks, BR: basic rock) (Shikazono, 1988c).

Loq(Zn)

CI = S.W.

5 4 3 I

2

6

%

1

%

~X

0

%

,%

-1

X"

-2

-31000

;-.

x I

I

800

I

I

600

i

1

400

I

I

200

,

%

%

I

o

SW

T(°C)

Figure 1.59. Relationship between zinc concentration and temperature of geothermal waters and hydrothermal solution experimentally interacted with rocks (Shikazono, 1988c).

elements at hydrothermal conditions (200-400°C) are much higher than those of seawater. Therefore, it is evident that the effect of temperature on the concentrations of ore-forming elements (base metals, Ba) is strong. For example, the concentrations of base metal elements in modified seawater interacted with acidic volcanic rocks at 300°C are Fe, 100

80

Chapter 1

TABLE 1.10 Isotopic compositions of Kuroko ore solution (K.O.), seawater (S.W.) and magmatic water (M.W.) (Shikazono, 1978) K.O.

S.W.

3348E,o,S ~348I],r,S ~)34Stota1

+21 tO +24%o +2 to +5%o >+2 tO +5%o

+20%o

M.W.

+20%0

0 to +1%,)

~180 ~D 87Sr/86Sr

0 to +5%0 to 30 to -3%0 0.7062-0.7087

0%o 0%o 0.709

+7 to +9%° --48 to --85% 0.704

334Sz,o,S: Sulfur isotopic composition of oxidized sulfur species. ~34S~,r,S: Sulfur isotopic composition of reduced sulfur species. ~348total:Total sulfur isotopic composition of aqueous sulfur species. ppm; Ba, 10 ppm; Zn, 1 ppm; and Cu, 0.5 ppm (Table 1.9). These concentrations are roughly in agreement with those estimated by solubility calculations on Kuroko ore fluids (Tables 1.5 and 1.6). Isotopic compositions of minerals and fluid inclusions can be used to estimate those of Kuroko ore fluids. Estimated isotopic compositions of Kuroko ore fluids are given in Table 1.10. All these data indicate that the isotopic compositions lie between seawater value and igneous value. For instance, 87Sr/86Sr of ore fluids responsible for barite and anhydrite precipitations is 0.7069~).7087, and 0.7082-0.7087, respectively which are between present-day seawater value (0.7091) and igneous value (0.704~3.705). From these data, Shikazono et al. (1983), Farrell and Holland (1983) and Kusakabe and Chiba (1983) thought that barite and anhydrite precipitated by the mixing of hydrothermal solution with low 87Sff86Sr and seawater with high 87Sr/86Sr. ~34S values of sulfides are +2%o to +7%o, which lie between igneous value (0%o (typical value)) and seawater value (+20%o). This suggests that both seawater sulfur and sulfide sulfur of the igneous rocks were incorporated into the Kuroko ore fluids of modified seawater origin. Probably, the sulfide sulfur of Kuroko ore fluids were derived from partial reduction of seawater sulfur and dissolution of pyrite and anhydrite in the country rocks. It is noteworthy that 334S values of sulfides from small B sub-type are smaller than large C sub-type and Y sub-type. This difference could be explained in terms of the extent of seawater sulfate reduction to H2S. The above argument on the calculation of chemical composition of ore fluids, seawater-rock interaction experiments, and isotopic compositions of ore fluids clearly demonstrates that Kuroko ore fluids were generated by seawater-rock interaction at elevated temperatures. The chemistry of present-day hydrothermal solution venting from back-arc basins and midoceanic ridges (sections 2.3 and 2.4) also support this view. There is another opinion on the origin of Kuroko ore fluids. Sawkins (1982) thought that intrusive felsic magmas were the source of the metals and heat in Kuroko hydrothermal systems. He stressed the contributions of magmatic fluid and seawater in

Miocene-Pliocene Hydrothermal Ore Deposits

81

ore fluids from which Kuroko deposits formed based on 3D, 3180 (Hattori and Sakai, 1979; Hattori and Muhlenbachs, 1980) and salinity data on inclusion fluids. His calculated mixing ratio of magmatic fluid to seawater was 1/9. The problem on his magmatic model is that metal content of magmatic fluid mixed with cooler hydrothermal solution of seawater origin is high and metal sulfides tend to precipitate even under the small mixing ratio of magmatic fluid to hydrothermal solution probably at very high temperatures. Another problem of magmatic hypothesis which has been critically thought by Ohmoto et al. (1983) is sulfur behaviour during this type of mixing. High salinity of Kuroko ore fluids does not solely mean magmatic contribution. Instead salinity variation can be reasonably explained by subcritical boiling of fluids of seawater origin. ~180 and 3D data (Fig. 1.40) do not clearly show the mixing of seawater and magmatic water which was favoured by Ishihara and Sasaki (1978). Ohmoto et al. (1983) pointed out that the first problem of the magmatic hypothesis is the large difference in the S concentrations between Kuroko ore fluids and magmatic fluids; fluids derived from acidic magmas at temperatures >750°C are more likely to contain ~ 1 0 -1 to > 1 M ~2S (Burham and Ohmoto, 1980), whereas Kuroko ore fluids appear to have contained 10-3-10 -2 M ZS at 250 4-50°C. The second and most serious problem pointed out by Ohmoto et al. (1983) is that magmatic fluids with +5%e sulfur which was estimated based on 334S values of granitic rocks in Green tuff region by Ishihara and Sasaki (1978) are unlikely to form sulfides of ~+5%o at temperature <300°C under the chemical conditions of the Kuroko ore fluids. Even if we accept that ~D and 3180 values lie between the two end member ~ values, this feature can be also explained by low ~D and high 3180 solution being generated by low seawater/rock ratio condition. Lead isotopic data support this interpretation; namely, these data clearly indicate leaching of lead from the rocks (Fehn et al., 1983). Horikoshi and Shikazono (1978) indicated that 3D of ore fluids for B sub-type which is located at centre of Hokuroku basin is higher, suggesting large contribution of seawater, while 3D of ore fluids of Y sub-type located at the margin of Hokuroku basin is lower, suggeting meteoric water contribution. Ohmoto et al. (1983) thought that the involvement of meteoric water in Kuroko ore fluids is unlikely because they estimated the entire Hokuroku district was under more than 2,500 m of seawater depth (probably ca. 3,500 m and the district was located more than 50 km away from the nearest island based on the studies of foraminifera assemblage in mudstone overlying Kuroko deposits (Guber and Ohmoto, 1978; Guber and Merrill, 1983). However, the seawater depth is inferred to be shallower probably 1,000-2,000 m from the following reasons. One of the reasons is that the seawater depth of hydrothermal ore depositions recently discovered at back-arc basins (Okinawa Trough, Izu-Bonin, Mariana-Trough, Marius Basin, North Fiji Basin, Lau Basin) is mostly 1,000-2,000 m. Oki and Hayasaka (1978) reported an occurrence of recent arenaceous foraminifera in a shallow submarine caldera, which usually lies at a depth of more than 4,000 m on the bottom of the open ocean. The seawater in the caldera has a lower pH than normal seawater due to the inputs of volcanic gas and hydrothermal solution with low pH. Lower pH raises the carbonate compensation depth (CCD) there. Therefore, the estimate of

82

Chapter 1

E v

o)

7-]:

20

19

18

3"~ ~ 17

I ~.~

I 16

15

time

14

(Ma)

Figure 1.60. Variation of subsidence rate for syn-rift basins in the Uetsu district, northeast Honshu (Yamaji, I990). The line of boxes shows the spatially averaged subsidence rate. The rate after 15 Ma is not clear because of uncertainty in paleobathymetry. However, the rate probably decreased to the order of 10-100 m/m.y. If the rate had been of the order of 1 km/m.y, after 15 Ma, the water depth of the inner arc region at 14 Ma would have been much deeper than modem, young, back-arc basins.

seawater depth based on paleobathymetry using CCD as a reference point in the Miocene open ocean by Guber and Ohmoto (1978) might be invalid (Kitazato, 1979; Matoba, 1983; Yam@, 1990). Kitazato's (1979) estimate of 1,000-2,000 m as the seawater depth of Kuroko formation seems more likely. Akimoto and Hasegawa (1989) also estimated to be 1,000-2,500 m. Their estimated seawater depth is similar to that of present-day back-arc basins from which hydrothermal ventings occur (Okinawa Trough, Mariana). Yamaji (1990) studied paleobathymetry of the Northeast Japan Arc at 17-14 Ma (Fig. 1.60) and showed that the arc subsided rapidly from 16 to 15 Ma (Fig. 1.61). According to his paleobathymetry, it is likely that shallow sea (sublittoral ~ 1 5 0 m and upper bathyal 150-500 m) existed near the site of Kuroko formation at 15-16 Ma. Considering the discussion above, it cannot be ruled out that meteoric water was involved in the Kuroko ore fluids. In fact, ~D and 3180 values of fluid inclusions from Iwami Kuroko deposit, west Honshu (San-in district) are considerably lower than those of seawater, indicating an involvement of meteoric water. However, 3D and 3180 values of Nurukawa Kuroko ore fluids (Y sub-type) are plotted between seawater value and igneous (or magmatic) value (Ishiyama et al., 2001), suggesting an involvement of igneous (or magmatic) fluids and no contribution of meteoric water. Here, igneous fluids mean the fluids controlled by igneous rocks, which were generated under a low water/rock ratio.

83

Miocene-Pliocene Hydrothermal Ore Deposits

®

Q •

S

SS

S

/~'Oo s~ ,~ ,

100 km

N = NON-MARINE S = SUBLITTORAL (0-150 m) o = UPPER BATHYAL (150-500 m) • = M I D D L E T O L O W E R BATHYAL (500-2500 m)

Figure 1.61. Paleobathymetry of the NE Japan Are at 17-14 Ma (circled numbers) (Yamaji, 1990). The arc subsided rapidly from 16 to 15 Ma. The entire inner arc region became at middle bathyal depths until 15 Ma, except that Yuri hill remained at shallow marine depths. The hill was submergeduntil 14 Ma (Matoba, 1981). Stippled area, Aosawa basalt (Tsuchiya, 1988). Hatched area, Yuri hill. Note that submarine volcanoes are neglected in this figure because volcaniclasticrocks usually contain few fossils.

The most important conclusion derived from the isotopic studies mentioned above is that isotopic characteristics of Kuroko ore fluids were caused dominantly by seawater-volcanic rock interaction at elevated temperature and by the mixing of seawater with small portions of igneous water or the hydrothermal solution whose chemical and isotopic compositions are controlled by water-rock interaction under the rock-dominated condition and also small proportion of mixing of meteoric water. However, it cannot be decided at present which processes (degree of seawaterrock interaction or mixing ratio of seawater, igneous water and meteoric water) are important for the generation of Kuroko ore fluids solely from the isotopic studies. But experimental and theoretical considerations on seawater-volcanic rocks interaction and origin of hydrothermal solution at midoceanic ridges suggest that Kuroko ore fluids can be produced dominantly by seawater-volcanic rock interaction.

1.4. Epithermal vein-type deposits Epithermal vein-type deposits can be divided into four types based on total metal produced and metal ratio: base-metal type, precious-metal (Au, Ag) type, Sb-type and Hg-

84

Chapter 1

type based on the association of metals. Table 1.2 summarizes estimated total productions of Au and Ag, and Ag/Au. Ag/Au of base-metal type is higher (58 by weight ratio) than that of Au-Ag type (13 by weight ratio). The characteristic features of base-metal and precious-metal types are summarized below and those of Sb- and Hg-type deposits are described in section 1.7.

1.4.1. Geological characteristics 1.4.1.1. Distribution Epithermal Au-Ag vein-type deposits occur widely in the Japanese Islands (e.g., Northeast Hokkaido, Southwest Hokkaido, Northeast Honshu, Sado Island, Izu Peninsula, Kyushu) (Fig. 1.62). The deposits occur in young (late Miocene, and Pliocene), subaerial volcanic regions and in submarine-altered volcanic regions (so-called Green tuff region). The distribution area is located at arc-arc junction (Kubota, 1994) and margin of Green tuff region (Figs. 1.5 and 1.6). These deposits are the principal gold and silver producers in Japan (Table 1.2). Epithermal base-metal vein-type deposits are distributed in the Green tuff region (Southwest Hokkaido, Northeast Honshu) (Fig. 1.62). The distribution area of this type of deposits is nearly same as that of Kuroko deposits. For example, large deposits (Osarizawa Cu-(Au); Ani Cu-Au; Hosokura Pb-Zn deposits) occur in Northeast Honshu, but are more widely distributed in the Green tuff region than Kuroko deposits. The base-metal vein-type deposits in Northeast Japan occur chiefly in Oligocene and early-middle Miocene submarine strata in the members of the Monzen (60-25 Ma), the Daijima (25-15 Ma), and the Nishikurosawa (16-14 Ma), but sometimes they are found in those of the Onnagawa (13-7 Ma) and the Funakawa (7-2.5 Ma) stages of late Miocene (Fig. 1.63) (Nakamura and Hunahashi, 1970). Their distributions are mainly confined to the marginal part of the depressional sedimentary basin which formed during the Nishikurosawa-Onnagawa stage, and are characterized by the presence of rhyolite and dacite or the presence of the so-called Tertiary granite (Nakamura and Hunahashi, 1970). In Southwest Hokkaido, Cu-Pb-Zn veins and Kuroko deposits are distributed in the centre of a volcanic depression zone where submarine rocks predominate, while Au-Ag veins do not (Shikazono and Shimizu, 1988a). Many Cu-Pb-Zn vein-type deposits are hosted by organic sedimentary rocks such as shale and mudstone but almost all Au-Ag deposits occur in altered volcanic rocks. This difference in the host rocks affects the chemical features of ore fluids (fo2, fs2, fco2) (section 1.4.4). 1.4.1.2. Age of mineralization The age of formation of epithermal vein-type deposits can be estimated from K-Ar ages of K-bearing minerals (adularia, sericite) in veins and in hydrothermal alteration zones nearby the veins. A large number of K-Ar age data have been accumulated since the work by Yamaoka and Ueda (1974) who reported K-At age data on adularia from Seigoshi Au-Ag (3.7 Ma) and Takadama Au-Ag deposits (8.4 Ma). Before their publication on the K-Ar ages of these deposits it was generally accepted that epithermal

85

Miocene-Pliocene Hydrothermal Ore Deposits II

I

I

0 I00 200km '

Sanru Q

Hokuryu

.KonomQi

Todor¢

Yakumo J okok u

Imaiishizaki

Osorizawa

Jlwoto 4o °

Takachi

11nnai iYatanl IHanda Koruizowa

Mizobe

Hoshin~

Takeno

Omidani tAkc

Bojo

\

~1~Seigoshi ..----Yugashimo --Nowaji .Kawozu

35 o

Toi "~ T ~ :ai~1 L ~Sakoshiodomori

Okuchi Hishikari mgano

1J

J

2

~ushikino

lourQ

iwoto I

133 °

I

138°

Figure 1.62. Location of epithermai-type deposits in Japan (Shikazono and Shimizu, 1988a). l: Green tuff and subaeriaI volcanic region of Tertiary/Quaternary ages, 2: Main PaIeozoic/Mesozoic sedimentary terranes, 3: Main metamorphic terranes. TTL: Tanakura tectonic line, ISTL: Itoigawa-Shizuoka tectonic line, MTL: Median tectonic line. Open circle: epithermaI Au-Ag vein-type deposits, solid circle: epithermal base metal vein-type deposits, open triangle: epithermal Au disseminated-type deposits.

vein-type deposits in Tertiary volcanic zone formed in Miocene age, the same as Kuroko deposits because Miocene rocks host the veins. The K - A r age data are summarized in Figs. 1.64 and 1.65. It is obvious in these figures that (1) ages of formation of epithermal vein-type deposits vary widely from 15 to 1 Ma, but are mostly 6-1 Ma, (2) epithermal vein-type deposits have been formed

86

Chapter I

/

Figure 1.63. Distribution of vein-type deposits in a part of the inner belt of Northeast Japan. A: the sedimentary members of the Onnagawa stage, B: the sedimentary members of the Nishikurosawa stage (Nakamura and Hunahashi, 1970).

[] ° _

>,

Au-Ag Vein

[ ] Cu-Pb-Zn Vein

Kuroko

< "5 t-

E Z

15

10

5

1

Age (Ma) Figure 1.64. Ages of formation of Neogenie base-metal vein-type, Au-Ag vein-type and Kuroko deposits, estimated from K-Ar age and paleontologic data (Shikazono, 1987b).

later than Kuroko deposits (15-16 Ma), (3) base-metal veins have been formed earlier than Au-Ag veins (Sawai and Itaya, 1996) (Fig. 1.65), and (4) the frequency of ages of mineralization has three peaks: 10 Ma, 5-6 Ma, and 2-1 Ma.

87

Miocene-Pliocene Hydrothermal Ore Deposits Au-Ag vein

I

°+f +I

I

I

r

,

Cu-Pb-Zn vein

0L

F I~

Mn vein

Mn strata-bound

z , ,, Barite velnF-- ]

F-'l

Massive barite

L,

Pyrite vein

II,

,,

Kuroko

K-At age(Ma) 10

L 1=

Figure 1.65. Histogram of K-Ar ages of hydrothermal ore deposits in the Shakotan-Shikotsu district, Hokkaido (Sawai and Itaya, 1996).

1.4.1.3. Volcanic activity related to mineralization Volcanic rocks associated with this type of mineralization are mostly andesitic rocks. For example, Watanabe (1990b) found that the age of andesitic rocks (flat lava) becomes younger eastward, correlating to the age of epithermal mineralization in Southwest Hokkaido, suggesting a genetic relationship between andesitic volcanism and mineralization. Izawa and Urashima (1989) showed a good correlation of ages of andesitic volcanic rocks and A u - A g deposits in south Kyushu. Their ranges are 4 Ma-1 Ma. However, west of Izu Peninsula, this correlation is unclear and bimodal volcanism (acidic and basaltic activities) seems related to the epithermal A u - A g mineralization.

88

Chapter 1

1.4.1.4. Metal enriched and metal ratios The ore deposits can be classed into two types based on the types of associated metals: Au-Ag rich deposits (Type A) from which Au and Ag are produced as main products, and base metal (Cu, Pb, Zn, Mn, (Sn), (W), (Bi), (Mo), (Sb)) rich deposits (Type B) from which Au and Ag are recovered as byproducts. The deposits are associated with felsic and intermediate volcanic rocks but generally not with felsic plutonic rocks. In Japan Au-Ag deposits associated with granitic rocks (e.g., Au-Ag vein-type deposits in Kitakami) occur commonly. However, these plutonic-type deposits are not described here. Total tonnages of production of Au, Ag and other associated base metal elements were estimated from various records (e.g., Geological Survey of Japan, 1980), as shown in Table 1.2. The Ag/Au ratio of the Type 1-A is lower than that of the Type 1-B from which base metals other than Au and Ag have been produced. In general, the vein-type deposits which have produced large amounts of the base metal elements especially Pb and Zn produce small amounts of Au, but sometimes large amounts of Ag (e.g., Toyoha, Yagumo, Hosokura, Yatani). The vein-type deposits which are rich in Mn have also high Ag/Au (e.g., Ohe-Inakuraishi, Yagumo, Rendaiji). Some vein-type deposits which have produced Cu tend to have large Ag and Au amounts (e.g., Furokura, Osarizawa, Kishu, Ikuno, Tada), but some small deposits which have produced small amounts of Cu have low Ag/Au ratios (e.g., Chitose, Takanosu, and Isobe-Koyama). These deposits, except Chitose from which very small amount of Cu has been produced, are stockwork in form and not typical vein-type deposits. Based on the total production data on individual mines, total production of Au and Ag and Ag/Au ratio from each type of deposits were calculated (Table 1.2). Type 1-A deposits have produced the largest amount of Au and Ag. Total production of Au from the Type 1-B and 2 is small. Ag production from the Type 1-B is lower than but not so different from that of Type 1-A. Average Ag/Au ratio of Type I-A is about 13 and that of Type 1-B is 58.

1.4.2. Mineralogical characteristics 1.4.2.1. Metal zoning Orebody zoning (Park and Macdiarmid, 1963) is observed in Cu-Pb-Zn deposits. For example, in Osarizawa deposit, which is one of the largest Cu-Pb-Zn deposits in Japan, ore metal zoning from deeper to shallower parts is Cu -+ ZnPb --+ AuAg. In the Yatani deposits, Pb-Zn ore occurs in the deeper part, while Au-Ag ore in the shallower part (Figs. 1.66 and 1.67). However, it is uncertain that Au-Ag vein changes continuously to Pb-Zn vein in deeper parts. Orebody zoning is not observed in Au-Ag veins, although the Au/Ag ratio of ore changes considerably with depth. District and regional zonings (Park and Macdiarmid, 1963) are generally not found in Cu-Pb-Zn mine district nor in Au-Ag mine district. 1.4.2.2. Ore minerals Epithermal base-metal vein-type deposits are characterized by the abundant occurrence of sulfides (chalcopyrite, pyrite, sphalerite, galena), and a scarcity of Au-

89

Miocene-Pliocene Hydrothermal Ore Deposits

Au.AgVein

Fault

/" J

$1IJ

/

~o H 2

%." "'dz.., ,." , ~ ' - -

--'-¢~-'~----

A Bo

wTu o

tI

~' WH "

500 m

_~..~P K3 ,K2 - ' ~ - . .

!

B

- ' " '--Esu ~ ~ -

" eo

\

Figure 1.66. Distribution of Au-Ag veins and Zn-Pb veins of the Yatani deposits. ETI: East-Tengu No. 1, ET2: East-Tengu No. 2, TI: Tengu No. 1, T2: Tengu No. 2, T3: Tengu No. 3, KH: Kanizawa-Honpi, WT: West-Tengu, WT2: West-Tengu No. 2, WH: West-Honpi, H2: Honpi No. 2, KU: Kanizawa Uwabanhi, YH: Yatani-Honpi, W7U: W7-Uwabanhi, E8U: E8-Uwabanhi, E3U: E3-Uwabanhi, N01F: No. 1 fault, N02F: No. 2 Fault (Shikazono and Shimazaki, 1985).

minerals. Ag-minerals (e.g., argentite, pyrargyrite, polybasite) are commonly recognized. Pyrrhotite, magnetite, hematite, and marcasite are occasionally observed. Electron microprobe analyses have revealed that many varieties of Ag mineral occur in the base metal and Au-Ag deposits. The most important features on the occurrences of Au-Ag minerals can be summarized as follows. Au-Ag vein-type deposits (Type l-A): (1) Se minerals such as naumannite, aguilarite, Se-bearing argenitie, and Se-bearing polybasite are found in some deposits (e.g. Sanru, Koryu, Takadama, Kushikino). (2) Au-Ag-Te minerals are rare throughout all of the studied deposits, but sometimes found in some deposits (e.g., hessite from Fuke, Ohkuchi, Arakawa). (3) The common Ag-Au minerals are argentite, As-Se minerals and Ag sulfosalts (polybasite, pyrargyrite, pearceite). (4) Bi, Pb, Zn, and Sn-bearing Ag minerals have not been found. (5) Electrum is abundant, compared with Type 1-B deposits. Native silver is poor in amounts. Base metal vein-type deposits (Type l-B): (1) The common Ag minerals are argentite, and Ag sulfosalts (pyrargyrite, polybasite). Ag sulfosalts are abundant in the late stage of mineralization and argentite occurs in the early stage of mineralization (e.g., Ohe-Inakuraishi, Toyoha).

90

Chapter 1

NO3F ¢,NO2F Ac

~B

, ~/i/~ /

K3

K2

?: i

i

"/7 •

1

11 ::

WH

,/

, /~w7 I

I I

o I

I

200 m 1

Figure 1.67. Vertical cross section along A-B in Fig. 1.66. Abbreviations are the same as those in Fig. 1.66 (Shikazono and Shimazaki, 1985).

(2) Se and Te minerals have not been found in these deposits. (3) Bi, Pb, Zn, and Sn-bearing Ag minerals are rarely found. (4) Abundance of electrum is small, although native silver is abundant in some deposits (e.g., Toyoha, Ikuno). A large number of analytical data on chemical composition of sphalerite are available (Shikazono, 1974a; Watanabe and Soeda, 1981 ). The FeS content of sphalerite from epithermal base-metal vein-type deposits varies widely mostly from 1 to 20 mol% (Fig. 1.68). The FeS content of sphalerite from epithermal Au-Ag vein-type deposits is low, mostly less than 1 mol% (Fig. 1.68). The FeS content of sphalerite from epithermal Se-type is slightly higher than that of epithermal Te-type. The FeS content of sphalerite depends on iron minerals coexisting with sphalerite. The FeS content of sphalerite coexisting with hematite is low (0.5-3.0 mol%), while that of sphalerite coexisting with pyrrhotite is high (mostly 10-20 mol%; Shikazono, 1975). In individual deposits (Toyoha Pb-Zn, Hosokura Pb-Zn deposits), the FeS content of sphalerite coexisting with pyrite varies widely in a range of 1-15 mot%. Mn and Cd contents of sphalerite from epithermal base-metal vein-type deposits are low except Mn content of sphalerite coexisting with alabandite which contains 4.4 wt% Mn (maximum value) (Shikazono, 1975).

91

Miocene-Pliocene Hydrothermal Ore Deposits

Frequency

Frequency Kuroko-typedeposits

10

EpithermalAu-Ag vein-typedeposits

15

m

20 FeS mole %

1"0

1'5

20 FeS mole °/~

Frequency

~rL,~ EpithermalCu-Pb-Zn

•

5

10

15

20 FeS mole%

Figure 1.68. Iron content of sphalerite from Kuroko, epithermaI Au-Ag vein-typeand epithermal base metal vein-typedeposits (Shikazono, 1977a). The Ag content of electrum from epithermal Au-Ag vein-type deposits is mostly in a range of 40-70 atomic% (Fig. 1.69). The Ag content of electrum from the Se-type varies widely (Fig. 1.70). The average Ag atomic% is 50 to 55. The Ag content of electrum from the deposits associated with Te mineralization (e.g., Chitose, Fuke, Takeno) is low, ranging from 26.0 to 40.6 atomic% (Fig. 1.70). Very few data on the chemical composition of electrum from epithermal basemetal vein-type deposits are available. However, it is evident that the Ag content varies widely (Fig. 1.71). The Ag content of electrum from the Osarizawa and Okuyama Cu deposits is low (NAg (Ag atomic%) = 8.6-17.7), while the Ag content of electrum from Pb-Zn-Mn deposits (Toyoha, Oe, Inakuraishi, and Imai-Ishizaki) is high (NAg = 60-80). Motomura (1986) reported that the Ag content of electrum from these deposits is higher than that from epithermal Au-Ag vein-type deposits. The geochemical implication of the Ag content of electrum is discussed in section 1.4.4. Chemical compositions of tetrahedrite-tennantite from epithermal base-metal vein-type deposits are characterized by (1) wide compositional variations, and (2) higher Zn and Sb contents and Ag and lower Fe, As, and Cu contents, compared with Kuroko deposits (Shikazono and Kouda, 1979).

92

Chapter I

100

0 ¢OLL

I

5O

. L• 0

20

40

60

80

100

NAg

Figure 1.69. Frequency histogram for the Ag content of electrum from epithermal Au-Ag vein-type deposits in Japan (Shikazono and Shimizu, 1988a).

The differences in Zn/Fe ratio of tetrahedrite-tennantite in epithermal vein-type and Kuroko deposits and that of sphalerite in these deposits can be interpreted in terms of the following exchange reaction: ((Cu, Ag)10Zn2(As,Sb)4S 13)tet "b (FeS)sph = ((Cu, Ag)10Fe2(As,Sb)4S 13)tet Jr- (ZnS)sph (1-21) where (FeS)sph, (ZnS)sph, are FeS and ZnS components in sphaterite, and ((Cu, Ag)]0 Fe2(As, Sb)4Sl3)tet and ((Cu, Ag) ioZn2(As, Sb)4S13)tet components in tetrahedrite-tennantite, respectively.

93

Miocene-Pliocene Hydrothermal Ore Deposits Te t y p e d e p o s i t s

• n=5

Date Okuyama Kushikino Agawa Chugu Fuke Okuchi Chitose Takeno Sado Kato

.,~n=lO • n=2

--~=n=15 = =-n=69 ~n=lO *n=17 -- = •n=6 -- --= n = l

Sakoshi Hishikari Koryu Chitose Sanru Kushikino Takadama Yatani Nebazawa Omidani

• n=l

Se t y p e d e p o s i t s

, n=61

u

n=20,~ •

L

L n=22 = n=41-i

I

0

i

I

I

I

=

=

= = n=44 I

50

• n=8

=

I

NAg

I

I

I

100

Figure 1.70. Ag content (atomic fraction of Ag) of electrum from the Te-type arid Se-type deposits (Shikazono et al., 1990). n: number of analyses.

30-

>20-

o c-"

O" LL

10"

0

2"0

40

60

80

NAg

100

Figure 1.71. Frequency histogram for the Ag content of electrum from epithermai base metal vein-type deposits in Japan (Shikazono and Shimizu, 1988a).

If equilibrium is attained for the above reaction, equilibrium constant (K1-21) is expressed as, KI-2I =

(azntet/aFetet)/(aZnspJaFesph)

(1-22)

94

Chapter 1

where aZn~et, aFetet, aZnsph, and aFesph are activities of tetrahedrite-tennantite and activities of ZnS and FeS in sphalerite, respectively. This equation suggests that the Fe content of tetrahedrite-tennantite positively correlates with that of sphalerite at constant temperature and pressure, indicating Fe and Zn contends of tetrabedrite-tennantite are useful to estimate physicochemical parameters (fs2, fo2 etc.) as well as Fe content of sphalerite, although detailed study on thermochemical properties of tennantite-tetrahedrite solid solution is still needed. The chemical compositions of Ag-minerals have been obtained from several A u Ag and base-metal vein-type deposits (Shikazono, 1978b; Ohta, 1991, 1992). For instance, Shikazono (1978b) found that argentite (acanthite) from epithermal A u - A g vein-type deposits (Seigoshi, Ohmidani) contains appreciable amounts of selenium but that from epithermal base-metal vein-type deposits (Toyoha, Ikuno) does not. High selenium contents of galena and Ag sulfosalts (polybasite, pyrargyrite) from epithermal A u - A g vein-type deposits (Kushikino) were reported by Takeuchi and Shikazono (1984) (Table 1.11). The geochemical implications of selenium contents of sulfides for the physicochemical environment of epithermal ore deposition will be discussed in section 1.4.4.

1.4.2.3. Gangue minerals Quartz is the most abundant gangue mineral. It occurs commonly in Au-Ag and P b - Z n deposits but is scarce in Cu deposits. Chalcedonic quartz coexisting with Au-Ag minerals occurs abundantly in A u - A g deposits. Amethyst is generally rare and occurs as a late-stage mineral in Au-Ag and Pb-Zn deposits. Magnetite is common in P b - Z n - M n and Cu deposits but has not been reported in A u - A g deposits. It commonly coexists with other iron minerals such as hematite, pyrite, pyrrhotite, siderite, and chlorite and also occurs in both the main stage of sulfide mineralization and in the late stage of mineralization. The occurrence of hematite is generally more widespread than magnetite in all types of deposits, especially Cu-Au deposits. It tends to occur in late-stage mineralization, generally later than the sulfide mineralization. Inesite is common in A u - A g deposits, especially in the Izu peninsula (Kato, 1928). Generally it is found associated with highgrade gold-silver ores, as, for example, in the Kakehashi vein of the Kawazu mine. It also occurs rarely in Pb-Zn deposits such as Toyoha and Tatsumata. The occurrence of other M n - C a silicates such as johannsenite, bustamite, rhodonite, pyroxmangite, tephroite, and penwithite has been reported from A u - A g and Pb-Zn deposits, but these minerals are not common. They have not been reported from Cu deposits. Hydrated calcium silicate minerals such as xonotlite, truscottite, and gyrolite are rare but have been reported from several Au-Ag deposits. They do not coexist with A u Ag minerals but instead are found with quartz, carbonates, and johannsenite. However, in the Keisen No. 3-2 vein in the Hishikari Au-Ag deposits, a close association of electrum with truscottite, smectite and calcite is observed (Imai and Uto, 2001). Prehnite is found with A u - A g minerals in the A u - A g vein (Kanisawa vein) of the Yatani P b - Z n - A u - A g deposit but is not found in the Pb-Zn vein (Yatani-Honpi vein).

Miocene-Pliocene Hydrothermal Ore Deposits

95

Calcium silicates such as wairakite, epidote, prehnite, laumontite, and stilbite are common in the wall rocks of some A u - A g deposits in the Izu peninsula. Epidote occurs as a gangue mineral coexisting with sulfides and quartz in some Cu deposits, but none of the other above-mentioned Ca and Mn silicates have been reported from these deposits. Laumontite is a common mineral in propylite, which is the host rock for A u - A g deposits. Other zeolites such as mordenite and dachiardite are not generally common, but they are the main gangue minerals associated with A u - A g minerals in the Ohnoyama and Awagano A u - A g deposits. Chlorite is abundant in Cu-Pb-Zn-rich deposits but is scarce in Au-Ag-rich deposits. Fe chlorite is the most common and Fe-Mg chlorite is subordinate (Shirozu, 1969). Almost all of the chlorite is classified as orthochlorite which can be regarded as part of the clinochlore~taphnite solid solution series. In general, chlorite is intimately associated with sulfide minerals such as sphalerite, galena, pyrite, chalcopyrite, and pyrrhotite. A 7 A septechlorite was reported from the Toyoha Pb-Zn deposits (Sawai, 1980). Interstratified chlorite-smectite and vermiculite-saponite are rather common minerals in A u - A g deposits (e.g., Yoneda and Watanabe, 1981), but they have not yet been reported from other deposits. Smectite commonly occurs in A u - A g deposits, and to a lesser extent in Pb-Zn deposits, but only rarely in Cu deposits. The occurrence of beidelite in the Seigoshi and Ohkuchi A u - A g deposits has been described (Nagasawa et al., 1981), and sericite is a common gangue and alteration mineral in C u - P b - Z n rich deposits, where it occurs as a late-stage mineral (Nagasawa et al., 1976). Sericites (Watanabe et al., 1982) from A u - A g deposits occur in the wall rocks as alteration products. In addition, kaolinite commonly occurs as a late-stage mineral in some A u - A g deposits, but not together with the ore minerals. Sometimes kaolinite has been found as an alteration product of adularia after main-stage mineralization. Carbonate minerals occur in almost all the vein-type deposits. In general, they are more abundant than Ca and Mn silicates, but their abundance varies widely with different types of deposits. Large amounts of Mn carbonates (rhodochrosite and manganoan calcite) occur in P b - Z n - M n deposits, moderate amounts in Pb-Zn and Cu deposits, and small amounts in A u - A g deposits. Calcite is abundant in all types of deposits. Siderite is common in Cu deposits, especially in C u - A u deposits, but it is uncommon in A u - A g and Pb-Zn deposits. Siderite coexists with hematite, pyrite, chlorite, and rarely with magnetite; it is considered to have been precipitated after the main stage of sulfide mineralization. Other carbonates such as ankerite, dolomite, and kutnahorite are not common. Ankerite has been reported from several C u - P b - Z n deposits where it is found with sulfides and also with oxides such as magnetite and hematite without sulfides. Dolomite is found in some A u - A g deposits but is rare in other deposit types. Siderite and ankerite coexist with oxides and sulfides, but calcite generally occurs as a late-stage mineral associated with quartz and sericite. Mn carbonates are found with sulfides in P b - Z n - M n deposits such as Ohe and Inakuraishi, but this assemblage is not common. Carbonates from epithermal base-metal and A u - A g vein-type deposits are different. Mn- and Fe-carbonates are common in base-metal vein-type deposits and calcite is abundant in A u - A g vein-type deposits. Shikazono (1973) revealed that the iron content

T A B L E 1.1 l R e p r e s e n t a t i v e analytical data on ore m i n e r a l s from K u s h i k i n o A u - A g d e p o s i t (Takeuchi and S h i k a z o n o , 1984) S a m p l e No.

Period

Cu

Au

Aq

Zn

Fe

Cd

Mn

Sb

As

S

Se

Total

FeS a (tool %)

Electrum .

.

IV

-

65.0

33.0

.

76081806a, L6

lI

-

60.5

38.7

.

.

.

.

.

.

.

.

99.2

7 7 0 7 2 6 0 3 a , 1.8

lib

-

60.0

38.5

.

.

.

.

.

.

.

.

98.5

77072601a, L8

lla

-

63.0

35.0

.

76081810c, L9

lib

-

57.0

42.0

.

.

.

.

.

.

.

.

99.0

76081809f, L9

llb

-

67.0

34.0

.

.

.

.

.

.

.

.

101.0

Za, L 6

IV

2.63

-

69.5 ~

0.09

0.00

-

-

6.49

2,22

11.84

6.36

77080212b, L6

IV

3.69

-

67.57

0.(X/

0.00

-

-

8.62

0.74

11.49

6.91

99.02

78031601 c, L B

Ila

3.09

-

69.63

0.00

0.00

-

-

9.44

(t.79

14.81

1.66

99.42 99,79

.

.

.

.

.

.

98.t)

77080212e, L6

.

.

.

98.0

.

Polybasite 99.14

77072602f, L8

Ila

2.67

-

67.69

0.01

0.00

-

-

11.24

0.33

14.32

3.53

77072602i, L8

Ma

3.01

-

66.24

0.00

0.00

-

-

10.60

0.75

14.90

3.75

99.25

77072602d, L8

lla

3,t8

-

69.52

0.(X)

0.00

-

-

10.05

0,93

15.09

1.20

99,97

7 7 0 7 2 9 0 l c , 1.8

I1

2.77

-

68.32

0.11

0.00

-

-

9,98

0.55

12.31

5.57

99.61

7 7 0 7 2 9 0 B c , 1.8

llb

2.48

-

68,15

0.02

0.00

-

-

9.90

0.25

ll.90

6.71

99.41

7 7 0 7 2 7 0 1 a , 1.8

lib

2.68

-

68.91

0.(X)

0.00

-

-

9.65

0.51

12.26

5.92

99.93

7 6 0 8 1 8 0 9 c . L9

IIb

3.80

-

65.78

0.10

0.31

-

-

9.85

0.37

12.04

7.30

99,55

7 6 0 8 2 3 0 3 c , 1.9

IIb

3.84

-

70.83

0.10

0.00

-

-

9.14

0.36

12.86

2.99

100.12

76081809b, L9

IIb

4.12

-

65.09

0.25

0.63

-

-

9.84

0.54

11.50

7.57

99.52

76081806a, L9

IIb

4.13

-

71.36

0.44

0.26

-

-

7.35

0.46

10.01

5.01

99.02

76081806c, L9

Ilb

4.32

-

68.50

0.25

0.00

-

-

9.18

0.44

12.57

4.68

99.94 ¢5

T A B L E 1.11 ( c o n t i n u e d ) I

Sample No.

Period

Cu

Au

Aq

Zn

Fe

Cd

Mn

Sb

As

S

Se

Total

FeS ~

~.

(tool % ) Tetrahedrite

77080221a, L6

II

23.68

-

20.20

5.24

1.29

-

26.19

1.25

22.30

0.01

77080212e, L6

IV

23.92

-

19.16

4.81

1.71

-

-

27.79

0,02

22.60

0.134

100.05

77072901c,

18

IIb

23.01

-

20.60

4.34

2.02

-

-

26.07

1.16

23.00

I).06

100.26

77072912c, L8

lib

32.42

-

7.14

5,97

1.05

-

-

27.23

1.71

24.00

0.00

99.52

76081809b, L9

IIb

22.99

-

20.90

3.86

2.57

-

26.07

1.23

23.37

0,25

101.24

76081B06b,

1.9

IIb

28.54

-

12.10

6,48

0.57

-

-

27.95

0.73

23.54

0.02

99.93

76081809a,

1.9

IIb

22.02

-

211.39

3,87

2.64

-

-

26.33

1.34

22.88

0.00

100.47

100.16 ~-~

~z Naumanni~

77080212,

1.6

7 6 0 8 2 3 0 3 , 19

IV

-

-

74.17

-

0.64

25.69

100.50

lib

-

-

80.57

-

8.96

7.65

97.18

100.45

~"

Pyrargy~te

1, L d

Ilb

0.00

-

59.98

0.00

0.00

77072602A, LB

lla

0.00

-

59.82

0.00

0.00

77072602B, L8

IIa

0.00

-

60.53

0.00

0.00

-

-

22.46

0.37

16.45

1.19

-

22.06

0.56

17.08

0.00

99.52

-

22.42

0.15

17.12

0,26

100.48

-

Sphalerite

77072901a,

63.84

0.76

0.78

0.05

-

-

32.59

65.13

0.06

0.00

0.10

-

-

33.11

18

lib

0.10

_

m

77072909e, L8

IV

0.00

m

m

77072601e, L8

IIa

0.10

62.72

1.56

0.52

0.00

-

-

33.56

-

98.46

2.5

76081801a, L9

lib

0.24

62,93

1.38

0.42

0.111

-

-

33.86

-

98.84

2.0

76081802b, L9

IIb

0.18

7 6 0 8 1 B 10b, L 9

IIb

0.33

76081810c,

lib

1.73

19

r

98.12

1.2

99.40

0.1

63.65

1.80

0.39

0.00

-

-

33.67

-

99.69

2.8

63.29

1.98

0.83

0.00

-

-

33.08

-

99.51

3.0

62,14

2.65

0.73

0.01

-

-

32.40

-

99.66

2.1

a F e S : C o r r e c t e d v a l u e e x c l u d i n g the e f f e c t o f X - r a y s c a t t e r i n g c a u s e d b y c h a l c o p y r i t e .

"--3

98

Chapter 1

(FeCO3) of Mn-carbonates varies widely in a range of 10-3-10 -1 mole fraction. Siderite from Au-Ag vein-type deposits (Ohmori) contains appreciable amounts of zinc (0.8-5.8 wt% as ZnO) (Shikazono, 1977b). Adularia is abundant in Au-Ag deposits, where it is commonly found with AuAg minerals; only rarely does it occur in Pb-Zn and Cu deposits. Albite is very rare and is reported only from the Nebazawa Au-Ag deposits. Barite is a common gangue constituent in P b - Z n - M n deposits, especially those in the southwestern part of Hokkaido and the northern part of Honshu, where it is usually a late-stage mineral coexisting with carbonate and quartz but rarely with sulfide minerals. Other rare gangue minerals include fluorite, apatite, gypsum, bementite, rutile, and sphene, but they have not been studied. Main gangue minerals of the Se-type deposits comprise quartz, adularia, illite/ smectite interstratified mixed layer clay mineral, chlorite/smectite interstratified mixed layer clay mineral, smectite, calcite, Mn-carbonates, manganoan calcite, rhodochrosite, Mn-silicates (inesite, johannsenite) and Ca-silicates (xonotlite, truscottite). In comparison, the Te-type deposits contain fine-grained quartz, chalcedonic quartz, sericite, barite, adularia, chlorite/smectite interstratified mixed layer clay mineral and rarely anatase. Carbonates and Mn-minerals are very poor in the Te-type deposits and they do not coexist with Te-minerals. Carbonates are abundant and barite is absent in the Se-type deposits. The grain size of quartz in the Te-type deposits is very fine, while large quartz crystals are common in the Se-type deposits although they formed in a late stage and do not coexist with Au-Ag minerals. Principal gangue minerals in base-metal vein-type deposits are quartz, chlorite, Mn-carbonates, calcite, siderite and sericite (Shikazono, 1985b). Barite is sometimes found. K-feldspar, Mn-silicates, interstratified mixed layer clay minerals (chlorite/smectite, sericite/smectite) are absent. Vuggy, comb, cockade, banding and brecciated textures are commonly observed in these veins. The predominant gangue minerals vary with different types of ore deposits; quartz, chalcedonic quartz, adularia, calcite, smectite, interstratified mica/smectite, interstratified chlorite/smectite, sericite, zeolites and kaolinite in Au-Ag rich deposits; chlorite, quartz, sericite, carbonates (calcite, rhodochrosite, siderite), and rare magnetite in Pb-Zn rich deposits; chlorite, sericite, siderite, hematite, magnetite and rare epidote in Cu-rich deposits (Sudo, 1954; Nagasawa et al., 1976; Shikazono, 1985b).

1.4.2.4. Hydrothermal alteration zoning Among the epithermal vein-type deposits in Japan, four major types of hydrothermal alteration can be discriminated. They are: (1) propylitic alteration, (2) potassic alteration, (3) intermediate argillic alteration, and (4) advanced argillic alteration. The definitions of these types of alteration are mainly based on Meyer and Hemley (1967) and Rose and Burt (1979) who classified the hydrothermal alteration in terms of alteration mineral assemblages. Representative propylitic alteration minerals include epidote, albite, carbonates, quartz, chlorite, sericite, and smectite. The less common minerals are mixed-layer clay minerals such as chlorite/smectite and sericite/smectite and zeolite minerals. The term "propylite" is widely used to describe altered volcanic rocks recognized

Miocene-Pliocene Hydrothermal Ore Deposits

99

in mine areas. However, the term is somehow ambiguous, because it is defined differently by different investigators (e.g., von Richthofen, 1886; Kato, 1928). Propylite is defined here as the andesitic, dacitic and rhyolitic rocks which have undergone the alteration to form the characteristic minerals such as chlorite, epidote, albite, K-feldspar, etc. Probably there are two different origins of this type of alteration. The propylite may be formed either by metamophism or by hydrothermal alteration. However, it is extremely difficult to distinguish rigidly between these two origins. Here the studies on the propylitic alteration which is considered to be intimately related to igneous activity in the mine area are summarized. Usually lateral and vertical zonations are observed in the area of propylitic alteration. Generally epidote, actinolite and chlorite tend to occur in the central and deeper parts, while at the marginal and shallower parts zeolites, mixed-layer clay minerals and smectite are commonly observed. The following zonation is generally recognized from the central to marginal parts: chlorite --+ chlorite/smectite -+ smectite. This zoning pattern is observed in the Fuke Au-Ag district (Inome et al., 1981), Seigoshi Au-Ag district (Shikazono, 1985a), Hosokura Pb-Zn district (Suzuki et al., 1982), Toyoha PbZn district (Yoshitani, 1971; Sawai, 1984), Miyatamata, Arakawa and Nissho C u - P b Zn districts (cited in Utada, 1980). At the central parts of some mine districts listed above, intrusive plutonic bodies have been observed. The formations of high temperature alteration minerals including epidote, actinolite, prehnite and wairakite as suggested by Yoshitani (1971), Fujii (1976), and Shikazono (1985a) are considered to be attributed to these intrusive activities. Zeolite minerals (wairakite, laumontite etc.), mixed-layer clay minerals and smecite occur in the upper part of the propylitically altered rocks (e.g., Seigoshi, Fuke, Kushikino), but they are sometimes poor in amounts. Generally carbonates are more abundant in the mine area as in the Toyoha district. Temporal relationship between the formation of high temperature propylitic alteration minerals (epidote, actinolite, prehnite) and low temperature propylitic alteration minerals) (wairakite, laumontite, chlorite/smectite, smectite) in these areas (Seigoshi, Fuke, Kushikino) is uncertain. Potassic alteration has been recognized in the Numanoue (Otagaki, 1951), Konomai (Urashima, 1953), Takarakura (Utada, 1980), Seigoshi (Shikazono and Aoki, 1981; Shikazono, 1985a; Imai, 1986), Hosokura (Suzuki et al., 1982; Konno et al., 1984), Chitose (Takatori and Nohno, 1985), Karuizaiwa (Sugaki et al., 1986) and the Ikutahara district, Hokkaido (e.g., Yahagi, Ryuo) (Matsueda et al., 1992) where adularia occurs dominantly. In the Mikawa (Sudo et al., 1953, Nagasawa, 1961, 1962), Ohe (Tsukada and Uno, 1980), Toyoha (Okabe and Bamba, 1976), Rendaiji (Watanabe and Nagai, 1986) and Kushikino (Izawa et al., 1987), sericitization is dominant. The area of the potassic alteration is not wide, compared with the propylitically altered area. The width of potassic alteration zone away from the vein is generally within several tens of meters (ca. 50 m) (Shikazono and Aoki, 1981; Imai, 1986). The potassic alteration is usually found in the intermediate vicinity of the vein in the epithermal deposits in Japan. Thus it is evident that this type of alteration occurs genetically related to the ore deposition. Lateral zonation from a sericitic envelope to an intermediate argillic envelope is common in the porphyry copper deposits and vein-type deposits in granodioritic rocks

100

Chapter I

(Meyer and Hemley, 1967). However, such lateral and concentric zonation has not been reported from the epithermal vein-type deposits in Japan. Montmorillonite-rich and silicarich zones exist in the upper part of the Au-Ag veins such as the Seigoshi and Takadama (Nagasawa et al., 1981). In contrast to the hardly investigated lateral zonation around Japanese epithermal vein-type deposits, a few examples of vertical zonation are known. Potassic alteration grades upwards into intermediate argillic alteration in the wall rocks for the Toyoha (Okabe and Bamba, 1976), Ohe (Tsukada and Uno, 1980), Chitose (Hasegawa et al., 1981) and Kushikino (Imai, 1986). Advanced arigillic alteration is found at the upper horizon than the sites of potassic and intermediate argillic alterations where the Au-Ag mineralization occurs (e.g., Seigoshi, Yatani, Kushikino, Hishikari). This type of alteration takes blanketform in upper part and vein-form in lower part (Iwao, 1962; Shikazono, 1985a). The conspicuous zonation from upper to lower horizon is known at the Ugusu silica deposit, namely, silica zone, alunite zone, kaolinite zone and montmorillonite zone (Iwao, 1949, 1958, 1962). It is rather difficult to determine the sequence of each type of alteration in a mine area. However, it is widely accepted that the hydrothermal alteration proceeds as follows: propylitic alteration ~ potassic alteration and intermediate argillic alteration -~ advanced argillic alteration. The actual sequence alteration might be more complicated and superimposition of each type of alteration could be common. Usually propylitic alteration precedes the base metal and Au-Ag mineralizations. Potassic and intermediate argillic alterations are nearly contemporaneous with ore deposition. Advanced argillic alteration is also nearly contemporaneous with base-metal and Au-Ag mineralization as found in the Seigoshi and Yatani Au-Ag mine areas based on K-Ar datings (Shikazono, 1985e), although advanced argillic alteration was not caused by the hydrothermal solution responsible for the Au-Ag mineralization. Advanced argillic alteration at the Ugusu silica deposit is inferred to be caused by the mixing of volcanic gas containing SO2 and groundwater (Shikazono, 1985a). Generally, the chemical composition of rocks does not considerably change during the propylitic alteration. The components which are added to the rocks are only H20, CO2 and S (e.g., Okabe and Bamba, 1976). Considerable changes in the chemical composition of rocks occur during the advanced argillic alteration. For example, SiO2 content of highly silicified rocks of the Ugusu silica mine reaches 99% (Iwao, 1962). This silicification is caused by the considerable leaching of elements from the rocks by acid hydrothermal solution except Si and addition of Si from hydrothermal solution.

1.4.2.5. Spatial and geochemical relationships between propylitic alteration and advanced argillic alteration: a case study on the Seigoshi-Ugusu district, central Japan In the Izu Peninsula, located in the middle part of Honshu, more than 20 epithermal Au-Ag vein-type deposits have been mined. Large Au-Ag mines are located in the western part of the peninsula. The Seigoshi mine is the largest one. The country

Miocene-Pliocene Hydrothermal Ore Deposits

101

TABLE 1.12 Generalized stratigraphic succession in the Izu Peninsula and the area surveyed (Shikazono, 1985a)

~ ~e" O° om R" °

' Daruma

i

Odoi

andesite lava +500m andesite lava +500m

~

Tanaba andesite lava .~ =~" Andes te pyroclastics +500m

~'

e- O

K0shimoda andesite lava Andesite volcanic breccia +500m

I-- v

Nekko Dacite

~' "~ ~. R"

dacite lava tuff breccia

+500m I

e= !~

~)

Yagisawa dacite lava andesitic tuff, Formation tuff breccia +1000m •-dacitic tuff u~ breccia andesite lava Tom (~ andesitic tuff, breccia +1000m •= O. Formation sandstone J~ ~ dacitic tuff ~ volcanic conglomerate Ugusu andesite lava pyroclastics >" Formation conglomerate sandstone +500m

rocks in the Seigoshi district have suffered intense propylitic and advanced argillic alterations. The Izu Peninsula is mainly composed of pyroclastic and volcanic rocks of Tertiary-Quaternary age. The general geology of the peninsula has been well studied (Tayama and Niino, 1931), and thus, it is briefly described below. The generalized stratigraphic succession in the Izu Peninsula is shown in Table 1.12. The oldest rock exposed in the peninsula is the Yugashima Group of Miocene age that is composed of submarine andesitic-dacitic pyroclastic and volcanic rocks with small amounts of sandstone. This rock is metamorphosed from zeolite facies to epidoteprehnite-pumpellyite facies. The total thickness of the group is variable from place to place, but, generally exceeds 1500 m. The Shirahama Group which is mainly composed of felsic tuff and sandstone overlies the Yugashima Group conformably or unconformably in different places. The age of this group is considered to be late Miocene to Pliocene. The thickness of this group is also variable from place to place, but the average thickness is more than 1,000 m. These Tertiary rocks are overlain unconformably by subaerial Quaternary andesitic rocks. The generalized geology and schematic stratigraphic succession in the SeigoshiUgusu district is shown in Fig. 1.72. The andesitic and dacitic rocks of Pliocene age unconformably overlie these Tertiary rocks. The Ugusu Formation, corresponding to the lower horizon of the Yugashima group, is characterized by the predominance of andesitic pyroclastic rocks. The maximum thickness of the Ugusu Formation is estimated

102

Chapter 1

l~

....

v v v vVVV'V'V'V

yvVvVvV~vVvVvVvVvVv3yvVvV~vVvVvVvVv vVv ~,Vv v VvVv~v~vV~Vv~v~,,v, v yvV~ v ~WW~VW4 v v v yvV~VvVvV,,Vv~ VvVW,

~AAAA

A A A A A A

A A A A A

~l~

........... o

0

o

o

o

0

o

o

o

0

o

o"o° o

o

0

o

o

o

o

o

o

A A ~

Daruma-Odoi Volcano (andesite lava) 1 ~ Tanaba Volcano (andesite lava, pyroclastics) Koshimoda Volcano (andesite lava, volcanic breccia) Nekko Volcano (dacite lava, tuff breccia) Yagisawa formation ........ (dacitic lava, andesitic tuff, tuff breccia dacitic tuff breccia)

[~

Toi formation (andesite lava, andesitic tuff breccia, sandstone, dacitic tuff, volcanic conglomerate) ~ - ~ Ugusu formation (andesite lava, pyroclastics, sandstone, conglomerate) Intrusive rock of diorite porphyry

Figure 1.72. Generalized geological map of the Seigoshi-Ugusuarea (Shikazono, 1985a).

to be more than 500 m. The Toi Formation, corresponding to the upper horizon of the Yugashima Group, conformably overlies the Ugusu Formation. This formation is mainly composed of pyroclastic and volcanic rocks of andesitic composition. The average thickness of this formation is 1,000 m. The Yagisawa Formation, which may correspond to the Shirahama Group, unconformably overlies the Toi Formation. This formation is composed of dacitic and andesitic lavas and pyroclastic rocks. Its total thickness is ca. 1,000 m. Andesitic and dacitic rocks of Pliocene or Pleistocene unconformably overlie the Yugashima and Yagisawa formations. In the northern part of this district, Quaternary andesitic rocks are distributed. These rocks are composed of an alternation of thick lava flows of hypersthene-augite andesite and their autobrecciated parts. Many faults trending generally from north to south are developed, mainly in the Yugashima Group (Fig. 1.73). The axis of the minor foldings in the Yugashima rocks is also generally north to south. Basic and felsic intrusive and dike rocks often occur in the

103

Miocene-Pliocene Hydrothermal Ore Deposits

t

J

,,' ,'~... Advanced Argillic ~" . \ ,' p.. ,' Alteration m T "1 ,' ~Funabara T \ ,'..----0 . . . . 4. area " ~ " : " ~ o ~Seigosi~i'-. " , ~ o i ,• T o"',l ~ ~, ~ "',. . ,,mine . co ~ n e ', ,.~ ,'/ . . , . , Propylltlc 'j .." .,' ', 'J" ,,', ",<~ Alteration .. ,; ~ F9 ,,'~ ,, ~,-" :/Ugusu ,' Mochikoshi ,: m ne : \ '. mine /,"

', F / .,a-x" - 2.'. 't

, ;

Amagi mine : , , ' ~ A ' " ?~ ,'

"J'.

" 'J

",3

',.

~, < 3 ,

F

"~

.... -..2~ ', "

5 km

;',. . . . . . . . . . . . ', ~ ' ,

•t

'

'." "

['~2q Diorite-Porphyry ~ ~

Silicified Rock of Advanced Argillic Alteration Gold-SilverQuartz Vein Fault

Figure 1.73. Distribution of epithermaI Au-Ag vein-typedeposits, propylitic and advanced argillic alterations and intrusive rocks of diorite prophyry(Shikazono, 1985a). Yugashima, trending from north to south and from east to west. Occasionally the A u - A g quartz veins are hosted by these intrusive rocks and the upper horizon of the Yugashima Group. These intrusive rocks have suffered propylitic alteration. The host rocks for the advanced argillic alteration are generally younger formations such as Pliocene andesite, although the Yugashima Group rocks have also suffered the advanced alteration occurring as a vein in form. The distribution of the A u - A g vein-type deposits in this district is shown in Fig. 1.73. The propylitic alteration is intimately associoated with these deposits. The vein is composed of rhythmic banding of quartz layers and fine-grained sulfides such as argentite, acanthite, sphalerite, galena, pyrite and chalcopyrite, and electrum. The principal gangue minerals are quartz, calcite, adularia and interstratified chlorite/smectite. Minor minerals are inesite, johansenite, xonotlite and sericite. These gangue minerals except for quartz, adularia, calcite and sericite are not found in the wall rocks. The Seigoshi and Toi deposits occur in the andesitic pyroclastic rocks of the upper horizon of the Yugashima Group and basic intrusive rocks. Distributions of the wallrock alteration minerals from underground in the Seigoshi mine and on the surface near the

104

Chapter 1 West <

> East

',, • Ao~-'. . . . .

...... 2 . . . . jJ

...... \--~ii-Ne~-2-_:-_:-2-_-2yuo

_. . . . ~ , 2 - 1 ..-+

5--4 j

~ \ "

', ~-

'

-.

,

,

.:~J/-~,," ~,, opx J

A~

J'~:f-

f~f J'~

•

+-

?I,

o~

\

"

\

'

', ',

I~.

'

/

"

I t

e/

',

--

, /l

', ,

#5-, "v ,

.......

/

~

T

I

I I

SeigoshiAu-AgVein

Figure 1.74. Zonal sequence of the propylitic alteration in E-W section of the Seigoshi-Toi mine area (Yug = yugawaralite; Heu = heulandite; Stil = stilbite; Opx = orthopyroxene; Mont = montmorillonite;Mor = mordenite; Lm = laumontite;Wr = wairakite; Chl = chlorite; pr = prehnite;ep = epidote; Py = pyrite; Kf = K-feldspar; Cpx = clinopyroxene)(Shikazono, 1985a). Deep(A)

)

Shallow(B)

Mordenite Yugawaralite Heulandite Stilbite Laumontite Wairakite Montmorillonite Chlorite Epidote Prehnite K-feldspar Pyrite

Pyrrhotite Magnetite Sphene

.....

Figure 1.75. Zonal sequenceof the propylitic alteration in section A-B in Fig. 1.74 (Shikazono, 1985a). mine are shown in Fig. 1.74. This type of alteration has a vertical zonal arrangement (Fig. 1.75). The abundance of the alteration minerals also changes from the portion near the A u - A g - q u a r t z vein to the portion away from the A u - A g - q u a r t z vein. From the deep to shallow and from the portion near the A u - A g - q u a r t z vein to the portion away from the vein, the hydrothermal alteration zoning is observed: (1) epidote-prehnite-K-feldsparchlorite zone; (2) wairakite-laumontite zone; and (3) stilbite-heulandite-smectite zone. The rocks of the underground levels of the Seigoshi mine have suffered alterations (1) and (2). The boundary between these zones cuts the strata with a small angle. The boundary between each zone is gradual. In zone (1), quartz, K-feldspar, epidote, chlorite, prehnite and sphene are predominant alteration minerals. Epidote, prehnite and carbonate replace plagioclase phenocryst. Epidote often occurs as a veinlet with several millimeters wide, together with prehnite. K-feldspar, calcite and quartz tend to occur as a veinlet. Chlorite replaces pyroxene

Miocene-Pliocene Hydrothermal Ore Deposits

105

and also occurs as a veinlet often with pyrite. Orthopyroxene is completely replaced by chlorite, but clinopyroxene is sometimes preserved. Amphibole is often replaced by carbonate. The rim of the original magnetite is replaced by pyrite and sphene. In zone (2), wairakite and laumontite are commonly found. Wairakite occurs as a veinlet together with laumontite. Laumontite occurs as a veinlet and also as filling amygdule. Very small amounts of epidote are found with these zeolite minerals. Small amounts of interstratified sericite/montmorillonite are found. Veinlets of epidote, prehnite, quartz, K-feldspar and chlorite appear at the deeper part of this zone. Yugawaralite occurring as a veinlet is rarely found. In zone (3), stilbite, heulandite, and minor amounts of chabazite and mordenite are found mainly as veinlets and filling amygdule. The rocks of this zone have not been significantly altered. Original mafic minerals such as clinopyroxene and orthopyroxene are altered to smectite but sometimes they are preserved. Small amounts of carbonate minerals are found in this zone. A zonal sequence of opaque minerals is also found. Pyrite is found in zones (1) and (2). Very tiny amounts of pyrrhotite coexisting with pyrite are found in zone (1). Magnetite is common in zone (3) but this mineral is thought to be original. Fluid inclusion studies have been carried out on quartz samples from the veinlets in the Yugashima Group and diorite porphyry and from the A u - A g vein. Homogenization temperatures of inclusion fluids in the quartz coexisting with laumontite and stilbite range widely from ~240°C to 380°C as shown in Fig. 1.76. This wide range of homogenization temperatures and the coexistence of vapor- and liquid-rich fluid inclusions in the same quartz crystal suggest that boiling took place when these zeolites and quartz were precipitated. The homogenization temperatures of fluid inclusions in the quartz which is in contact with epidote and prehnite range from ~235 to 285°C. All fluid inclusions in the quartz are liquid-dominated type and vapor homogenized into liquid. The range of temperatures for each alteration zone can be estimated from the following chemical reactions and thermochemical data available for these reactions laumontite

: wairakite + H20

(1-23)

yugawaralite = laumontite + quartz

(1-24)

yugawaralite = wairakite + quartz ÷ H20

(1-25)

stilbite

= laumontite + 3 quartz + H20

(1-26)

heutandite

= laumontite + 2 quartz + 3 H20

(1-27)

stilbite

= heulandite + 3 quartz + 3 H20

(1-28)

The coexistence of laumontite and wairakite is common in zone (1). If the saturated water vapor pressure is equal to 0.3 of total pressure (Zeng and Liou, 1982), the temperature for equilibrium reaction (1-23) and saturated water vapor pressure are estimated to be ~170°C and 230 bar, respectively (Liou, 1971b). Zeng and Liou (1982) have shown that yugawaralite is stable at less than ~230°C and a total pressure of 500 bar, under the condition of quartz saturation. However, if the activity of SiO2 is not unity, the boundary for reactions (1-24) and (1-25) may shift to lower temperatures. Liou (1971a) studied the equilibrium for reaction (1-26) and showed that the equilibrium

106

Chapter I o o

0 °o.

a

~0 = <

LZ

=

2 ~ ~

o ¢:.-

v

-8 &

¢- "0 ~

o o r-

U~

r'-

"E

o

-800 • 700 -600

O cO

.~0 i~ o "~

n -500

g

-400 • 300 "200

• loo

=~

0 - 100

-200

o

N

AdvancedArgiIlic AIteratlon

;i . . . . . . ~ ~ . , [ ~ v' :, O ooo,

. .

;

oo%oi ~G e.= p' e,ooo o o 0 0>3 oooo;. :,,. ~~ eoe ee el, ee, ~,ee= q o ~ o e o poo J e 13 (/} O e ~ e e *4e e e e e . . . . .

o i

200 300 400 ( T ' C ) ~ o:: Zeolile Zone ooo"

. . . .

EpidoleZone ~

Au-AgVein ,

,

L'1~" ,

I I

.£ -6~'. ............ "0 o : 5 , . . . . . . . . . . . . .

O. 0

[~] Argillic Alteration ] Montrnorillonite Advanced

]

Zone Zeolile Zone

] Epidole Zone

[ ~ Au-Ag Vein Figure 1.76. Schematic c o l u m n at section o f the Seigoshi A u - A g mine area, central J a p a n (Shikazono, 1985a).

temperature and pressure are 170-185°C at 2 - 5 kbar respectively. However, on the basis of experimental studies (Liou, 1971b; Maruyama et al., 1983) it is safe to expect that stilbite was formed under lower temperature conditions, probably less than ~150°C. Thus, it is considered that zone (3) was formed under temperatures less than ~ 150°C. In summarizing the fluid inclusion studies and stability of zeolite minerals, the most likely temperature range o f zone (1), (2) and (3) is estimated to be ~ 2 5 0 - 2 8 0 ° C , 150-230°C and < 150°C, respectively. Boiling of fluids for zones (2) and (3) suggests that the depth for the zeolite zone is probably less than 500 m from the surface and the epidote zone is more than 500 m.

107

Miocene-Pliocene Hydrothermal Ore Deposits +2 +1 0 ~-~ 1 O 2

A The G e y s e r s Larderello \ ,, -. k \

Broadlands . . t~ \(wel~ll) . _, . - " .~

tlveragera%\, , Showashinzan -'~,~, ;*" "

'~ . ' " .~'./~huachapan .." - " El Tatio . -

3

• Mexixali (well5)

Satsuma

\. . ~,; ~ ~, . . . " Nasu Chausudake - , % Wairakei ,,o . ' , ~.-,, • . . . . c ' Showashinzan ~ 0 * {average) :~:4[~ ~"

Iwojirna

...~':i~i~'~ :

::~

\B

-4

50

200

2;0

3;0

3;0

4;0

450

T(°C)

Figure 1.77. Range of hydrogen sulfide fugacity (fH2S) and temperature for the propylitic alteration (epidote zone) and the advanced argillic alteration (silica-rich zone) and some active geothermal systems. A = propylitic alteration; B = advanced argillic alteration. Data on active geothermal systems are taken from Ellis and Mahon (1977). The calculation of fH2s of Showashinzan for saturated water vapor pressure condition was based on the analytical data on volcanic gas by Matsuo (1961). The calculation of fH2S of Satsumalwo-Jima was based on the analytical data on volcanic gas by Matsuo et al. (1974) assuming PH20 = 0.5 kbar. ThermochemicaI data necessary for estimating fH2s for the propylitic and advanced argillic alterations are taken from Bird and Helgeson (1981) and Giggenbach (1981). The curves A-A', B-B' and C-C' represent the equilibrium of epidote(xp~ =0.30) - - K-mica(aK mica=0.9) - K-feldspar(a~ feidspar=0.95) -- pyrite - calcite - chlorite(aFco_0.5) where Xpis, aK-mica, aK-feldspar, and aFeo are mole fraction of pistacite component in epidote, activity of K-mica component in mica, activity of K-feldspar component in K-feldspar and activity of FeO component in chlorite, and hematite + liquid sulfur .~- pyrite + H2S, respectively (open circle = vapor-dominated system; solid circle = hot-water-dominated system; solid triangle = volcanic gas) (Shikazono, 1985a).

Ranges of $2, 02 and H2S fugacities were estimated as shown in Figs. 1.81, 1.82 and 1.77, respectively. $2 and 02 fugacities were estimated from the alteration mineral assemblage and homogenization temperature data. H2S fugacity was estimated from the chemical compositions of epidote and chlorite following the procedure by Giggenbach (1980, 1981) as shown in Fig. 1.77. CO2 fugacity can be inferred from the following chemical equilibrium relations (Bird and Helgeson, 1981): 3K-mica + 4calcite + 6quartz = 2clinozoisite + 4K-feldspar ÷ 4CO2 + 2H20

(1-29)

Based on the analytical data of K-mica, epidote and K-feldspar and using thermochemical data on these minerals (Helgeson and Kirkham, 1974; Helgeson et al., 1978; Bird and Helgeson, 1981), the fco2 range for the propylitic alteration was estimated (Fig. 1.78). The regional distribution of advanced argillic alteration in this district is shown in Fig. 1.73. The alteration zone of this type is distributed with a given trend, in general running from north to south. The distribution area of this type of alteration is more restricted than that of the propylitic alteration (Fig. 1.73). This type of alteration is well observed in the Funabara area, underground in the Seigoshi mine and the Ugusu silica mine (Figs. 1.73 and 1.79). The original rocks for this alteration are different in these

108

Chapter 1

T.Alfina I */Bagn/re @4

Rotorua V O~.~J'~

1

0 0 11

0

0 _1

-1

Broadlands /~" Nilan/d~Krafla (~zildere (~Ngawha ~ ~ 2 p h a s e ~ ~'~'~'~~'~'r ~

"-~L~------''~2Krafl pha%e ~ \ "CerroPrieto Sa,,onSea

/ ~ ~ A '

,~,and

" Seigoshi

-2 -3

-4 I

150

[

200

I

250

I

300

T(°C)

Figure 1.78. Range of carbon dioxide fugacity (fco2) and temperature for the propylitic alteration (epidote zone) in the Seigoshi area and some active geothermal systems. Seigoshi = propylitic alteration of the Seigoshi district. Data on active geothermal systems are taken from Helgeson (1967, 1968), D'Amore and Panichi (1980), Giggenbach (1980, 1982), Bird and Norton (1981), and Arndrsson and Gunnlaugsson (1983). ThermochernicaI data necessary for estimating fco2 for the propylitic alteration are taken from Helgeson and Kirkham (1974), Helgeson et al. (1978), and Bird and Helgeson (1981). The curves A-B and A ' - B ' are equilibria for epidote(xp~=o._~o) - - K-miCa(aK,.~c~ o.9) - - K-feldspar(aKqeid~p~r-0.95) - - calcite assemblages at saturated water vapor pressure condition.

areas; Tanaba andesitic rocks (Funabara area), basic intrusive rocks, pyroclastic rocks of the Yugashima Group and Tanaba Andesite (Seigoshi area) and Koshimoda Andesite, and andesitic lava and pyroclastic rocks of the Yugashima Group (Ugusu area). The elevation level of this alteration is high, ~200-400 m higher than the top of the A u - A g vein. Forms of the alteration halo are also different in these three areas; lenticular and mushroom like (Funabara), vein (Seigoshi) and strata-bound and lenticular (Ugusu). Among these three areas, the most intense alteration has occurred in the Ugusu area. Thus the description on the alteration in this area is given below as an example of the advanced argillic alteration. The Ugusu silica mine, situated 4 km south of the Seigoshi mine (Fig. 1.73), produces 6 x 105 tons of silica ore annually, which contains more than 95 wt% SiO2. The geology and the lateral and vertical alteration zonings of this mine are shown in Figs. 1.79 and 1.80. Quartz is the most predominant phase in the central zone and the silica content of this zone is 95-99 wt%. Native sulfur is found in the pores of the highly silicified rocks together with small amounts of topaz. Alunite occurs abundantly, surrounding this zone, with quartz. Dickite, sericite and sericite/montmorillonite mixed-layer mineral occur in the more peripheral zone. Montmorillonite (tri-type, contrasting with di-type of the propylitic alteration) occurs in the most peripheral zone. Generally, the boundary between the alunite zone and surrounding clay zone is sharp, but the boundary between the silica zone and alunite zone is gradual. Pyrite occurs in the clay zone. Hematite

109

Miocene-Pliocene Hydrothermal Ore Deposits /i 400 ~ ',--/-'.~,; - 5OO

.

)\

/

,,/

...~'~'--.o~ Hakko

.

-

"x.J.

..

.....6~a.'~::

:ii:"'=ii "'':.

BB

::::: ::..:::

Silica-rich Zone Alunite-rich Zone

L

j

~

1kin

Clay-rich Zone

~2~ Unconformity

Figure 1.79. Geology and alteration zoning in the Ugusu silica mine (modified from Iwao, 1949, 1962). Toi F = Toi Formation; Koshimoda = Koshimoda andesite; Hakko = Hakko orebody; Shibayama = Shibayama orebody. Numbers indicate metres above sea level. Alteration zoning in the section of A'-B ~ is shown in Fig. 1.80 (Shikazono, 1985a)

Montmorillonite a Cristobalite

Ser./Mont

.

Sericite Kaolin Alunite Topaz Native Sulfur Quartz Pyrite Hematite Rutile

.

.

Margin(A') Lower .

.

.

Center(B') Upper

.

........

___ ___

................ .................

Figure 1.80. Zonal sequence of the advanced argitlic alteration from the central to marginal zone in section of A'-B ~ in Fig. 1.79 and from upper horizon to lower horizon (Shikazono, 1985a)

o c c u r s as d i s s e m i n a t i o n a n d v e i n l e t s w i t h q u a r t z in t h e silicified a n d a l u n i t e zones. R u t i l e c o m m o n l y o c c u r s in the a l t e r a t i o n z o n e , e s p e c i a l l y in the c e n t r a l z o n e . A n a t a s e t e n d s to o c c u r in t h e m a r g i n a l p a r t o f the a l t e r a t i o n z o n e . S m a l l a m o u n t s o f o~-cristobalite a n d t r i d y m i t e o c c u r in the m a r g i n a l z o n e . T h e fluid i n c l u s i o n s c a n b e d i v i d e d i n t o t w o types: v a p o r - a n d l i q u i d - r i c h fluid i n c l u s i o n s . T h e filling d e g r e e o f fluid i n c l u s i o n s f r o m s o m e s a m p l e s f r o m the silicified a n d a l u n i t e z o n e s is v a r i a b l e a n d h o m o g e n i z a t i o n t e m p e r a t u r e s v a r y widely. T h i s i n d i c a t e s

110

Chapter 1

that the liquid-vapor separation occurred during the hydrothermal alteration process for the silica and alunite zones. Although the filling degree is variable, the homogenization temperature of the samples from the silica-rich zone is high, being in the range of 285-430°C; that from the alunite-rich zone in the range of 240-360°C and that from the alunite-clay-rich zone in the range of 220-280°C. This indicates that a steep temperature gradient existed in the hydrothermal alteration zones probably due to the mixing of high temperature volcanic gas and low temperature groundwater. Based on the hydrothermal alteration mineral assemblages and the fluid inclusion, the probable range of gas fugacities (fs2, fOR, fH2S) and temperature can be seen in Figs. 1.81 and 1.82: these estimated fugacities are quite different from those of the propylitic alteration. The characteristic features and differences of the two types of alteration are schematically summarized in Fig. 1.76, indicating the vertical changes in geology, alteration minerals and fluid inclusion characteristics in the Seigoshi-Ugusu district. The vertical alteration zoning from the shallower portion to the deeper portion in this district is summarized as follows: a massive part of the advanced argillic alteration zone; the zeolite zone of the propylitic alteration; and the epidote zone of the propylitic alteration. The advanced argillic alteration zone is developed near the unconformity boundary between relatively permeable pyroclastic rocks and the overlying relatively impermeable pyroclastic rocks. The upper portion of this district (zeolite and advanced argillic zones) is considered to be a two-phase (vapor and liquid) separation zone. It is worth noting that the mineralogical sequence in some active geothermal areas characterized by high chloride concentrations is similar to that found in the propylitic alteration of this district. For example, the alteration zoning in Wairakei, New Zealand, is similar to that of the Seigoshi district: in the shallower zone montmorillonite, illite-montmorillonite mixed-layer mineral, and laumontite occurs, and in the lower part

s~ c'd

co

f

s ~. "" j..-z.----~_ ~3,

-15 -20 -25

150

200

250

300

T (°C)

Figure 1.8i. Range of sulfur fugacity (.Ds2) and temperature for the propylitic alteration (epidote zone) and the advanced argillic alteration (silica- and alunite-rich zones (Shikazono, 1985a)). A = propytitic alteration; B = advanced argillic alteration. Thermochemical data on the fs2-temperature boundaries for the equilibria of: liquid sulfur ~- S2gas; hematite + pyrite ~- magnetite + S2gas; and pyrite ~ pyrrhotite + S2gas were taken from Heigeson (1969) and Rau et al. (1973a,b). S(1) = liquid sulfur; S2(v) = S 2 g a s ; ht = hematite, m t = magnetite; py = pyrite; po = pyrrhotite.

tll

Miocene-Pliocene Hydrothermal Ore Deposits

•

i.i i-

-30 -35 -40 04 O

"o

-45 -50 -55 -60

I

I

I

I

150

200

250

300

T (°C)

Figure 1.82. Range of oxygen fugacity (f%) and temperature for the propyiitic alteration (epidote zone) and the advanced argillic aIteration (silica- and alunite-rich zones (Shikazono, i985a)). A = propylitic alteration; B = advanced argilIic alteration. Thermochemical data on the fo2-temperature boundaries for the equilibria of: liquid sulfur + hematite ~ pyrite + H20 for saturated water vapor pressure condition; hematite ~ magnetite + O2; and pyrite + magnetite ~ pyrrhotite + 02, are taken from Helgeson (1969) and Rau et al. (1973a,b). Abbreviations used are the same as in Fig. 1.81. epidote and wairakite are found (Steiner, 1968). Both areas (Seigoshi and Wairakei) are characterized by c o m m o n occurrences of zeolites (laumontite and wairakite) and epidote and small amounts o f carbonates. The Larderello region (Italy) is also characterized by the occurrence of wairakite, laumontite and epidote (Cavaretta et al., 1982). Environmental conditions o f gas fugacities (fHzS, fco2) and temperature at the Wairakei and Larderello regions are similar to those of the Seigoshi district as shown in Figs. 1.77 and 1.78. The estimated depth o f the part affected by the zeolite and epidote alterations in the Seigoshi district is consistent with those of Wairakei and Larderello. The formation of epidote, K-feldspar, prehnite, wairakite and calcite in the geothermal area is considered to be due to the loss o f CO2 gas and rapid precipitation from the solution supersaturated with respect to quartz (Browne, 1978). The widespread occurrence o f these minerals in the Seigoshi district seems to be consistent with the above-mentioned consideration, namely that these minerals usually occur as veinlets rather than the replacements of original minerals and filling amygdule. In particular, many veinlets of epidote, prehnite and wairakite are found near the A u - A g - q u a r t z veins. There are three possible mechanisms for generating the strongly acid solution which caused the advanced argillic alteration: (1) alteration caused by the vapordominated system as inferred by White et al. (1971); (2) alteration caused by the oxidation o f HzS near the surface; and (3) alteration by volcanic gas a n d / o r hot water condensed from a volcanic gas. A m o n g them, (3) is the most attractive mechanism given the following evidence and considerations.

112

Chapter 1

(1) The temperature of the advanced argillic alteration estimated from the fluid inclusion studies and mineral assemblages, varies widely from ~220°C to ~420°C. The homogenization temperature for the central part of the alteration zone is from ~285°C to ~430°C. This temperature range is higher than the temperature of the vapor-dominated system defined by White et al. (1971) who showed that vapor-dominated systems such as Larderello (Italy), The Geysers and Mud Volcano (Yellowstone Park, Wyoming, USA) and Matsukawa (northeast Japan) are characterized by the temperature of around 240°C. (2) The advanced argillic alteration minerals are similar to those of the vapordominated system. But, in general, the minerals indicating relatively higher temperatures, such as pyrophyllite and diaspore, are lacking in the vapor-dominated system mentioned above. However, in Matsukawa which may be a vapor-dominated system at present, these high-temperature minerals like pyrophyllite, diaspore and andalusite have been reported (Sumi, 1968a). However, these minerals are considered to have been formed during the past and are not presently forming (Sumi, 1968b). (3) The minerals containing F and C1 such as topaz and zunyite are common in the advanced argillic alteration. The HF activity of the solution was estimated to be ~0.01 based on the F content of topaz from the Ugusu mine (Shibue and Iiyama, 1984). This high F concentration is observed in Crater Lake (Mt. Ruapehu, New Zealand) (Giggenbach, 1974) and is interpreted to be due to the injection of acid fumarolic gases to the lake (Giggenbach, 1974). Gases collected from a fumarole at Showashinzan and Satsumaiwo-Jima also have a high H F / H 2 0 volume ratio which is in the range of (2-5)× 10 -4 (Matsuo et al., 1974; Mizutani and Sugiura, 1982). (4) If alunite, K-mica and kaolinite (which are common minerals in the advanced argillic alteration) are in equilibrium, the concentration of H2SO4 can be estimated based on the experimental work by Hemley et al. (1969); the concentration of H2SO4 at 200°C and 300°C is 0.002 and 0.012 M, respectively. This may suggest that it is difficult to form such a high concentration of sulfate ion only by oxidation of H2S. (5) The fugacity of H2S (fH2s) for the advanced argillic alteration is plotted in Fig. 1.77 together with those for some active geothermal systems and volcanic gas. The fH2S for the advanced argillic alteration was estimated on the basis of hematite-pyriteliquid sulfur equilibrium. Estimated fH2S for the advanced argillic alteration is lower than that for the vapor-dominated system (Larderello and The Geysers) but is similar to that of the volcanic gas collected from one of the fumaroles of the Showashinzan volcano in Japan. Only a few areas subjected to solfataric alteration have been well studied. Satsumaiwo-Jima is probably the only area where alteration caused by volcanic gas has been studied in detail. The zonal sequence from the centre to the margin and from a topographically higher level to a lower level is from a silica-rich zone through an alunite-rich zone to a clay (montmorillonite)-rich zone (Yoshida et al., 1976). This pattern is similar to that of the advanced argillic alteration in the Seigoshi-Ugusu district, although oe-cristobalite is abundant and quartz is poor in the Satsumaiwo-Jima. But in the Seigoshi oe-cristobalite is poor and quartz is abundant. In the Satsumaiwo-Jima pyrophyllite, diaspore, zunyite and topaz are not found, although these minerals occur in the Seigoshi-Ugusu district. The occurrence of alunite in the Satsumaiwo-Jima indicates that the formation of alunite is largely controlled by the existence of groundwater. The

Miocene-Pliocene Hydrothermal Ore Deposits

l 13

alunite zone in this island is strata-bound in form. In the Ugusu mine the occurrence of alunite is very similar. It is interesting to note that the advanced argillic alteration tends to occur near the unconformity boundary (Fig. 1.79). The similarity in the mode of occurrence of alunite from the Satsumaiwo-Jima and from the Ugusu mine, the existence of the unconformity and the steep temperature gradient from central to marginal part in the Ugusu mine may suggest that the mixing of groundwater with volcanic gas and/or hot water condensed from the volcanic gas is the most important mechanism for the formation of advanced argillic alteration also in the Satsumaiwo-Jima area. The coexistence of these two types of hydrothermal alteration and associated high sulfidation-type A u - A g deposits (acid alteration) and low sulfidation-type Au Ag deposits (neutral alteration) is observed not only in the Seigoshi-Ugusu district, but also in the other epithermal A u - A g vein-type mine districts in Neogene volcanic region in Japan: Izu Peninsula, Takadama, Kushikino, Hishikari, Akeshi-Kasuga, Yatani, and Harukiyama A u - A g mine (Hokkaido) (Yamada, 1995) districts. In these districts both the epithermal vein-type deposits and highly silicified and clay-rich rocks are found in close proximity in time and space. The characteristic differences between the two types of coexisting associated alterations in each area are very similar to those in the Seigoshi-Ugusu district. For instance, advanced argillic alteration zone lies at topographically higher levels and in younger formations than the epithermal vein-type deposits and associated propylitic alteration. However, some differences exist in a different area. For instance, in the Kasuga-Akeshi area, fine Au particles occur in the highly silicified rocks. Surrounding the highly silicified rocks, the following zonal sequence is observed from the centre to the margin (Tokunaga, 1955; Saito and Sato, 1978): an alunite zone, kaolinite zone and a montmorillonite zone. This sequence is similar to that observed in the Ugusu deposits, but Au has not been detected in the silicified part of the Ugusu silica deposits. Topaz and zunyite have not yet been reported from the Kasuga silicified rock, while they are found in that of the Ugusu. Homogenization temperatures for the silicifled rocks in the Kasuga, Akeshi, Iwato and Kushikino areas are generally lower than those for the Ugusu deposits (Takenouchi, 1981). Boiling phenomena have been observed in the fluid inclusions from the silicified rocks in these areas (Takenouchi, 1981). Yamada (1995) described acid and potassium alterations and high sulfidation and low sulfidation mineralizations in the Harukiyama district (Hokkaido). He indicated based on K - A r data on the alterations that the high sulfidation-type have been formed first by the residual magmatic fluid subsequent to the intrusion of quartz porphyry magma and then the low sulfidation-type could have been formed by the circulation of meteoric water generated by the heated emanation from hot magma solidified but still similar temporal relationship is observed in the other districts (Osorezan: Aoki, personal communication, 1990; Yatani: Shikazono, unpublished; Seigoshi-Ugusu: Shikazono, unpublished).

1.4.2.6. Chemical composition of alteration minerals Although a wide range of alteration minerals has been recognized in epithermal systems considered here, few of their chemical compositions have been determined. Trioctahedral chlorite occurs commonly in geothermal and hydrothermal areas, whereas the occurrence of dioctahedral chlorite is very limited. For instance, donbasite

114

Chapter I

does not occur in geothermal and hydrothermal areas, but sudoite (Al-chlorite) commonly does not occur in Kuroko mine area (e.g., Tsuzuki and Honda, 1977). Dioctahedral chlorite has not been reported from the Neogene C u - P b - Z n vein-type deposits in Japan; instead, trioctahedral chlorite is common (Shirozu et al., 1975). In Kuroko mine area, the common chlorite is trioctahedral Mg chlorite. The structural formula for trioctahedral chlorite is represented by (Mg6-x-yF@ + Alx)(Alx Si4-x)O10(OH)8. It is convenient to plot a diagram of Fe2+/(Fe 2+ + Mg) (in atomic fraction) against ~VA1/(Si + IVA1)to display compositional variations in chlorite (Fig. 1.83) (e.g., Hey, 1954; Foster, 1962, Nagasawa et al., 1976). Although a large body of analytical data on hydrothermal chlorite from geothermal and hydrothermal areas is available, there are few data on the chemical composition of original fresh volcanic rocks. The relationship between the ratio MgO/FeO in chlorite and that in the fresh host-rocks indicates that the value of MgO/FeO of chlorite generally satisfies a line of 1 : 1 slope. The correlation between MgO/FeO in the host rock and that in the chlorite implies that the MgO/FeO value of chlorite from propytitically altered rocks associated with the mine areas and from altered rocks in terrestrial and submarine geothermal areas is largely affected by MgO/FeO ratio of original fresh rocks. However, most of the data from Kuroko and Neogene C u - P b - Z n vein-type deposits deviate significantly from this line. Chlorite compositions from areas (Toyoha Pb-Zn vein, Kuroko deposits) deviate significantly from a line of 1 : 1 slope. This deviation implies that the Fe2+/Mg value of chlorite from these areas is controlled not only by the FeO/MgO value of the fresh host rocks, but also by factors such as the ratio of Fe 2+ to Mg 2+ in the fluid phase. As noted already, several investigators have shown isotopically that seawater played an important role in the formation of Kuroko deposits (e.g., Sakai et al., 1970; Kajiwara, 1971; Hattori and Sakai, 1979; Farrell and Holland, 1983). Mg-rich chlorite occurs in gypsum-anhydrite bodies in many Kuroko deposits. Farrell and Holland (1983), Shikazono et al. (1983), and Kusakabe and Chiba (1983) suggested the involvement of large amounts of seawater or seawater-dominated hydrothermal solution in the formation of the gypsum-anhydrite bodies. It is reasonable to assume from the large number of experimental studies on rock-seawater interaction at elevated temperatures that the seawater-dominated fluid phase, which interacted with volcanic rocks at a high water/rock ratio, contained appreciable amounts of Mg but very small amounts of Fe (Seyfried and Mottl, 1982). Therefore, it is likely that Mg-rich chlorite precipitated from a solution with a high proportion of Mg 2+ to Fe 2+. In contrast to the Kuroko hydrothermal system, there is no evidence for the involvement of seawater in the hydrothermal system associated with the C u - P b - Z n vein mineralization. Results of hydrogen and oxygen isotopic studies indicate that large amounts of meteoric water were incorporated into the ore fluids responsible for these vein-type deposits (Hattori and Sakai, 1979). If F e - M g chlorite is assumed to be in equilibrium with a fluid phase and pyrite, the ratio of Fe 2+ to Mg 2+ in the fluids may be related to factors such as pH, f Q , temperature and total dissolved sulfur concentration (ES). This relationship can be derived from the following chemical reactions: MgsA1Si3A1Olo(OH)8 + 5 Fe 2+ = FesA1Si3A1Olo(OH)8 + 5 Mg 2+

(1-30)

115

Miocene-Pliocene Hydrothermal Ore Deposits 1.0 oI 03 01

04

0.8

05

._.0.6

o)

o 6

oO6

4LL

o6

v

o6

u_ 0.4

°6o6

o6

0.2

9 ~ 910_10 • •8 o 8 8 o8

09

0.0

018

1:0

1:2 1:4 AI in 4(AI,Si)

o7o 7

1:6

1:8

2~0

Figure 1.83. Variation of Fe2+/(Fe 2+ -}- Mg) and tetrahedral A1 of chIorite from hydrothermal ore deposits: Japanese Neogene Cu-Pb-Zn vein-type (open circle) and Kuroko deposits (solid circle). Localities: 1: Ashio (Nakamura, 1960, 1963); 2: Yatani (Hattori, 1974); 3: Toyoha (Shikazono 1974a, Sawai, 1984); 4: Kishu (Shirozu, 1958); 5: Sayama (Shirozu, 1958); 6: Mikawa (Nagasawa, 1961); 7: Furutobe (Shirozu et al., 1975); 8: Hanaoka (Hayashi 1961, Hayashi and Oinuma, 1965; Tsuzuki and Honda, 1977; Shirozu et aI., 1975); 9: Wanibuchi (Sakamoto and Sudo 1956, Iwao and Minato 1959, Katsumoto and Shirozu, 1973); 10: western Bergslagen (Baker et al., 1983) (Shikazono and Kawahata, 1987).

FeS2 + 2 H + + H 2 0 = F e 2+ + 2H2S + 1 / 2 0 2

(1-31)

w h e r e MgsA1Si3A1OIo(OH)8 and Fe5A1Si3A1Olo(OH)8 represent M g - c h l o r i t e and Fechlorite, respectively. Generally, activities o f liquid H 2 0 and FeS2 do not deviate f r o m unity; thus these values are a s s u m e d to be unity.

116

Chapter 1 From equations (1-30) and (1-31), we obtain, for the H2S dominant region,

log(aFe-chl/aMg-chl) = log K1-30 + 5 log Kl-31 -- 5 log aMg2+ -- 10 log ES - 10 log gHzS -- (5/2)1og fo2 -- 10pH

(1-32)

where g is the activity coefficient. For the SO 2- dominant region,

1og(aFe-chl/aMg-chl) = log K ~_30+ 5 log K 1-31 - 5 log aMg2+ 10 pH -- 10 log ES - 10 log YSO2- + 35/2 log fo2 + 10 log KI-34

(1-33)

where KI_34 is the equilibrium constant for the following reaction, H2S+202=SO ] +2H +

(1-34)

Equations (1-32) and (1-33) imply that the aFe--chl/aMg-chl of chlorite depends on foz, aMgz+, ~S, pH, temperature and ionic strength. From equations (1-32) and (1-33), it can be shown that the aFe-chl[aMg-chlof chlorite in equilibrium with pyrite decreases with increasing fo2 at constant temperature, pH, aMg2+, ionic strength, and ES in the region where H2S is dominant, whereas it increases with increasing f Q in the SO ] - dominant region. Previous studies on the estimates of f Q , fs2, pH, ZS and temperature for Kuroko ore deposition have been reported, for instance, by Kajiwara (1971), Shikazono (1976) and Ohmoto et al. (1983), who showed that the fo2 of the Kuroko ore fluids lies in the region close to the SO42-/H2S boundary (in this expression, SO ] and H2S represent the concentration of the total dissolved oxidized sulfur species and total dissolved reduced sulfur species, respectively). These estimates seem consistent with the tow Fe2+/Mg value of chlorite from Kuroko deposits. Chlorite from the Toyoha Pb-Zn vein type deposits is associated with sphalerite, pyrite and, rarely, pyrrhotite (Shikazono, 1974a). The iron content of the sphalerite associated with chlorite is 1.2-2.9 wt%. The temperature of formation of chlorite in the Toyoha deposits is estimated to be 200-250°C from fluid inclusion data (Shikazono, 1974a, 1975; Yajima and Ohta, 1979), which indicates that the chlorite was formed in a relatively reducing environment in which reduced sulfur species predominate. This estimate seems consistent with the composition of chlorite from the Toyoha deposits and from the other Neogene C u - P b - Z n vein-type deposits in Japan, which contains up to 40 wt% FeO (Nakamura, 1960, 1963; Shirozu, 1969; Shikazono, 1974a; Hattori, 1974). Shikazono (1974a,c, 1977a, 1978b) and Hattori (1975) showed that these deposits formed in the environments where species of aqueous reduced sulfur predominated. This estimated range of oxidation state and ratio of concentration of aqueous reduced sulfur species to oxidized sulfur species appears to be in agreement with the chemistry of chlorite from these deposits. However, the Fe2+/Mg of chlorite also depends on the other factors, such as IES, pH and aMg2+. From equations (1-32) and (1-33), it is obvious that increasing IgS and aMg2+ also causes a lower Fe2+/Mg in chlorite. The Fe 2+ /Mg and Fe-3+ /Fe 2+ values of chlorite from Kuroko deposits and Neogene C u - P b - Z n vein-type deposits differ greatly (Fig. 1.83). Chlorite from Kuroko deposits contains lower Fe2+/Mg and higher Fe3+/Fe 2+ values than this from the Neogene vein-type deposits in Japan. The most likely explanation for these differences is that these two types of deposit formed at different states of oxidation, although other

117

Miocene-Pliocene Hydrothermal Ore Deposits

variables such as pH, temperature, aMg2+ and NS are also possible important factors, as discussed above. Generally, the Fe z + / M g value of chlorite from a given geothermal or hydrothermal area is variable. The Fe2+/Mg value of chlorite in the Toyoha mine district varies widely. Sawai (1984) has shown that the iron content of chlorite away from the P b - Z n veins. Iron content of chlorite in the Toyoha veins is very high (40 wt% FeO: Shikazono, 1974a; Sawai, 1984). Variations in iron content of chlorite from the host rocks toward the C u Pb-Zn veins have also been studied for other Neogene vein-type deposits in Japan (e.g., Ohe, Ashio deposits: Hayashi, 1979). Variations in iron content of chlorite from Japanese epithermal mine districts suggest that the iron content of chlorite in the discharge zone of a hydrothermal system is higher than that in the recharge zone. Mottl (1983) has also suggested that the FeZ+/Mg value of chlorite in the discharge zone of submarine geothermal systems is higher than that in the recharge zone. The Fe2+/Mg of chlorite from different parts of a geothermal system is higher than that in the recharge zone. The FeZ+/Mg of chlorite from different parts of a geothermal system (recharge zones versus discharge zones) is probably not constant. In order to evaluate the effect of the fluid movement at discharge versus recharge zones, the temperature dependence of Fe2+/Mg 2+ in fluids in equilibrium with the isochemically recrystallized crystal rocks will be considered below. By taking a value for Fe2+/Mg of chlorite that is equal to that of the average andesitic and basaltic rocks (0.6), and assuming that chlorite is an ideal solid solution of 14 A Fe-chlorite and 14 A Mg-chlorite, the dependence of a aFez+/aMg2+in fluids on temperature was calculated by using thermochemical data for chlorite from Walshe and Solomon (1981). Figure 1.84 shows that aFez+/aMg2+of fluids increases with increasing

-2

+-3 04

~-4 o

200

250

3()0 350 Temp.(°C)

Figure 1.84. Variation of aFe2+/aMg2+of hydrothermal solution in equilibrium with chlorite having constant Fe2+/Mg (=0.6), as a function of temperature. The significance of points A and B is discussed in the text (Shikazono and Kawahata, 1987).

118

Chapter I

of temperature. Therefore, if fluids initially in equilibrium with chlorite having the same FeO/MgO as that of average andesitic and basaltic rock at elevated temperature (for example, at point A in Fig. 1.84) ascend rapidly without interaction with the surrounding rocks, chlorite precipitating from fluids at lower temperature (for example, at point B in Fig. 1.84) could contain appreciable amounts of Fe 2+ compared with Mg. This mechanism could lead to the formation of chlorite having an unusually high content of iron. It is also noteworthy that the vein chlorite in the altered basalt from the Costa Rica Rift contains higher concentrations of iron than the chlorite that replaces mafic minerals in the rock (Kawahata, 1984). In this case it is likely that the flow rate of ascending fluids from which vein chlorite precipitated is high compared with the rate of reaction between the ascending fluids and surrounding rocks. The Fe2+/Mg value of chlorite precipitating from ascending fluids depends on the extent of deviation from equilibrium between fluids and surrounding rocks. As discussed in detail by Giggenbach (1984), a number of processes such as adiabatic and conductive cooling of fluids and mixing of fluids can cause this deviation. The above considerations suggest that chlorite occurring in the discharge zones of hydrothermal systems would contain a higher concentration of iron than that occurring in recharge zones. If a magnesium-rich solution such as seawater or a seawater-dominated hydrothermal solution was involved in the hydrothermal systems, it could be expected that chlorite, even that occurring in discharge zones, would contain a high content of Mg. It is widely accepted that ascending hydrothermal solutions mixes with cold seawater at the time of ore formation at Kuroko (e.g., Hattori and Sakai, 1979; Shikazono et al., 1983). The Mg concentration of this hydrothermal solution increased with higher degrees of mixing. Involvement of large amounts of seawater at the site of ore deposition could be one of the reasons why chlorite from Kuroko deposits contains high amounts of Mg. In addition to a large concentration of Mg in these fluids, relatively high fo2, NS and SO42-/H2S, and low pH, could lead to the large Mg content of chlorite in Kuroko deposits, as discussed above. Consequently, the composition of chlorite in the discharge zone depends largely on the chemical nature of fluids (factors such as Fe2+/Mg 2+, SO42-/H2S, pH, aMg2+) and temperature. Movement of fluids may also be an important cause for the variability in the ratio of Fe 2+ to Mg in hydrothermal chlorite. Wide compositional variations in chlorite from the hydrothermal ore deposits in Japan, including Kuroko and Neogene C u - P b - Z n vein-type deposits, are considered to reflect the variable chemical nature of ascending ore fluids and fluids that mix with ascending ore fluids at discharge zone. The variations in Fe and Mg contents of the 14 ,~ Fe-chlorite-14 ,& Mg-chlorite solid solution are considered here. However, structural formulae for chlorite are not as simple as those considered here. As mentioned by Walshe and Solomon (1981), Stoesell (1984), Cathelineau and Nieva (1985) and Walshe (1986), chlorite solid solution may be represented by six components, and accurate thermochemical data on each end-member component at the hydrothermal conditions of concern are necessary to provide a far more rigorous calculation of the equilibrium between chlorite and hydrothermal solution. However, the above argument demonstrates that the composition of chlorite is a highly useful indicator of physicochemical conditions of hydrothermal solution and extent of water-rock interaction.

Miocene-Pliocene Hydrothermal Ore Deposits

119

Epidote is one of the most common alteration minerals occurring in geothermal and mine areas. The factors controlling the chemical composition of epidote from geothermal areas have been examined from thermochemical points of view. Bird and Helgeson (1981) discussed the effects of temperature, CO2 and 02 fugacities, and activity ratios (aFe3+)/(aH+) 3, (aca2+)/(au+) 2, and (aai3+)/(aH+) 3 in geothermal waters on the variations in iron and aluminum contents of epidote. Giggenbach (1981) calculated the effects of partial pressure of CO2 gas and temperature on iron content of epidote in equilibrium with alteration minerals such as kaolinite, K-mica, K-feldspar, albite, paragonite and calcite in geothermal systems. D'Amore and Gianelli (1983) showed the dependence of iron content of epidote on 02 fugacity for the epidote-K-feldspar-albitetremolite-chlorite equilibrium assemblage. Wolery (1978) and Reed (1982, 1983) have indicated based on a computer calculation of the change in chemistry of aqueous solution and mineralogy during seawater-rock interactions that epidote is formed under the low water/rock ratio less than ca. 50 by mass. Humphris and Thompson (1978), Stakes and O'Nell (1982) and Mottl (1983) have also suggested on the basis of their chemical and oxygen isotopic data of the altered ridge basalts that epidote is formed by seawater-basalt interaction at elevated temperatures (ca. 200-350°C) under the rock-dominated conditions. If epidote can be formed preferentially under such low water/rock ratio, the composition of epidote should be influenced by compositions of the original fresh rocks. Shikazono (1984) summarized analytical data of the epidote from geothermal areas to consider the relationship between the composition of epidote and that of the original fresh rocks and to inspect the other factors controlling the compositional variations in epidote. The discussion on the epidote composition by Shikazono (1984) is described below. Chemical compositions of epidote and original rocks in several geothermal and mine areas including Seigoshi (Japan, epithermal Au-Ag vein-type mine area), Yugashima (Japan, ancient geothermal area), Furutobe (Japan, Kuroko mine area), Ohtake (Japan, active geothermal area), Mid-Atlantic ridge, Costa Rica rift, Mitsuishi (Japan, pyrophyllite deposits), Shimokawa (Japan, ancient ophiolite area associated with Besshitype deposits), Reydarfjordur (Iceland, ancient geothermal area), Kushikino (Japan, epithermal Au-Ag vein-type mine area), Hachimantai (Japan, active geothermal area), Broadlands (New Zealand, active geothermal area), Wairakei (New Zealand, active geothermal area), Larderello (Italy, active geothermal area), and Sarmiento ophiolite complex (Chile, ancient ophiolite suite). In individual districts, these rocks are generally subjected to intense hydrothermal alteration. The original rock composition was estimated from analytical data of fresh rocks in the same area. Figure 1.85 shows that the Fe203 content of epidote positively correlates to Fe203 content of original rocks, although in general Fe203 content of epidote from each geothermal area has a variation over several wt%. Therefore, it is inferred that Fe203 content of original rocks has a large influence on iron content of epidote. Original rocks in Fig. 1.85 are mainly andesitic and basaltic rocks of island arcs and ocean ridge. In Fig. 1.85 iron contents of epidote from two different geologic environments, island arc and oceanic ridge or ophiolite, are summarized. It can be seen in Fig. 1.85 that the iron content of epidote from ridge basalt and ophiolite is generally lower than

120

Chapter 1

S

M. q

~13 O

2'

u..

1-

d"

sht

F,L=--~ Y

M, o ,cT J

%)

0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 Fe20s Content of Epidote (weight

Figure 1.85. Relation between Fe203 content of epidote and that of original fresh rocks. S: Seigoshi, Y: Yugashima, N: Noya, F: Furotobe, O: Ohtake, M: Mid-Atlantic ridge, C: Costa Rica rift, Mi: Mitsuishi, Sh: Shimokawa (Shikazono, 1984)

that from the volcanic rocks of island arcs. One of the most likely explanations for this difference is the difference in Fe203 content of original rocks; Fe203 contents of oceanic basalts and ophiolite are generally lower than those of island arc volcanic rocks. The positive correlation shown in Fig. 1.85 is consistent with the considerations on seawater-volcanic rock interactions under the hydrothermal conditions by Wolery (1978), Humphris and Thompson (1978), Reed (1982, 1983), Stakes and O'Neil (1982) and Mottl (1983). They suggested that epidote is formed under a low water/rock ratio less than ca. 50 by mass. Mineralogical changes resulted from seawater-volcanic rocks interactions have been elucidated by these recent theoretical studies. The epidotes except those from the Furutobe Kuroko mine area, Mid-Atlantic ridge, Costa Rica rift, and Shimokawa are suspected to have been formed in subaerial geothermal environments. Thus it is uncertain in these cases that the epidote tends to form at a low water/rock ratio. However, it seems likely that the Fe/A1 ratio of epidote formed in subaerial geothermal areas is also reflected by that of original rocks because the concentrations of Fe 3+ and A13+ in geothermal waters interacted with the rocks under the hydrothermal condition are generally very low except for strongly acid and high chloride solutions. Iron content of epidote depends on kind of iron minerals coexisting with epidote. The iron content of epidote coexisting with hematite is relatively high. For example, the epidote coexisting with hematite from the Yugashima, Furutobe and Reydarfjordur has higher iron content than that from the other districts. This relation holds even in geothermal systems (Salton Sea: Keith et al., 1968; Furutobe: Shikazono et al., 1995). On the other hand, the epidote coexisting with pyrite and pyrrhotite contains generally smaller amounts of iron as observed in the Larderello (Cavaretta et al., 1982), Shimokawa, MidAtlantic ridge, and Costa Rica rift. This difference in the iron content of epidote coexisting with different opaque minerals could be explained by the difference in oxygen fugacity. The iron content of epidote coexisting with prehnite from the Seigoshi district is smaller than that without prehnite. It has been clarified that prehnite is stable at the low

Miocene-Pliocene Hydrothermal Ore Deposits

121

CO2 fugacity and that the iron content of epidote in equilibrium with prehnite is lower than that in equilibrium with other minerals such as K-feldspar, K-mica and calcite under such low CO2 fugacity conditions (e.g., Cavaretta et al., 1982). Therefore, it seems clear that iron content of epidote is affected also by CO2 fugacity. Although Fe203 content of epidote from a given geothermal area has roughly a positive correlation to that of original rocks, the variation in Fe203 content of epidote from a given geothermal area ranges over several weight percents. This variation is too large to be explained only in terms of the difference in Fe203 content of original rocks because the composition of epidote varies sometimes so widely even in a single grain. Compositional zoning in epidote grains from subaerial and island arc geothermal systems is commonly known; usually, iron content of epidote increases from core to margin in a crystal. This tendency is observed in the Seigoshi, Furutobe, Yugashima, and Larderello areas, even though a decrease in the iron content from core to margin is rarely recognized in a single grain from the Larderello (Cavaretta et al., 1982). Complicated compositional zonings of epidote from the Salton Sea geothermal area are also found (McDowell and McCurry, 1978). In contrast, epidote from midoceanic ridge basalts (Melson and Van Angel, 1966; Humphris and Thompson, 1978; Kawahata, 1984) shows no wide compositional variation in a single grain. Epidote from a typical midoceanic ridge basalt has a small variation range of iron (Humphris and Thompson, 1978). The variation range in epidote from the Costa Rica Rift studied by Kawahata (1984) is wider than that observed by Humphris and Thompson (1978). Kawahata (1984) considered that the alteration minerals including epidote in the DSDP Hole 504B in the Costa Rica rift were formed by the interaction of midoceanic ridge basalt and ascending hot fluids. The basalt in the Shimokawa area is probably an abyssal tholeiite (Kohsaka, 1975). In this area, strata-bound cupriferous pyritic deposits are found and general features of these deposits resemble those of the ore deposits of Cyprus type and the ore deposits found in East Pacific Rise 21°N. Therefore, the wide variation range in epidote from this area may be explained by the effects of the chemical compositions and gaseous fugacities of ascending hot fluids responsible for the ore formation together with the chemical compositions of original basalt. The compositional zoning of an epidote grain and wide compositional variations in epidote grain from each geothermal area are difficult to explain because iron content of epidote depends on many factors other than the chemical composition of original rocks. However, the differences in the variation range and compositional zoning of epidote from the subaerial and oceanic ridge geothermal systems suggest that the change of physicochemical variables such as oxygen fugacity and temperature during the hydrothermal activities in both geothermal systems is different. Few data on the chemical compositions of feldspars (albite, K-feldspar) are available. Fujii (1976) indicated that K-feldspar and albite in the propylite of west Izu Peninsula, middle Honshu are of nearly end member composition. Nagayama (1992) showed that K-feldspars in the Hishikari A u - A g vein and in the host andesitic rock have different composition; N a / K ratio of K-feldspars from the vein is lower than that from the host rocks.

122

Chapter 1

1.4.2.7. Causes for hydrothermal alteration Hydrothermal alteration is reflected by the changes in many variables (temperature, water/rock ratio, extent of water-rock interaction (reaction progress), reaction rate, flow rate of fluids etc.) (Fujimoto, 1987). Theoretical and experimental works on hydrothermal alteration were reviewed by Meyer and Hemley (1967), and Rose and Burt (1979). In the last two decades, great progress has been made in the field of hydrothemal alteration studies, mainly from computation works on water-rock interactions at elevated temperatures (e.g., Wolery, 1978; Reed, 1983, 1997; Takeno, 1989). These studies revealed the relationship between the changes in chemical composition of hydrothermal solution and the relative abundance of minerals in the rocks. There are different approaches to the study of hydrothermal alteration. For instance, Shikazono (1978a) showed the relationship between chemical composition of hydrothermal solution in equilibrium with the alteration minerals and C1- concentration in hydrothermal solution. Giggenbach (1984) calculated the effect of temperature on the chemical composition of fluids buffered by alteration minerals. The causes for the hydrothermal alteration considered below are mainly based on the works by Shikazono (1978a) and Giggenbach (1984). The effect of the extent of water-rock interaction is not taken into account. Figure 1.86 illustrates the variations in the chemical composition of chloride-rich hydrothermal solution in equilibrium with common alteration minerals with temperature. Figure 1.86 demonstrates that (1) the chemical compositions of hydrothermal solution

B E~ 0

Na

H

Ca 2.

----4. i

-1

A

~

,-

I

-2

-3

~

150

200

r

250

t

300

Temp.(°C ) Figure 1.86. Variation in chemical compositions (in molaI unit) of hydrothermat solution with temperature. Thermochemical data used for the calculations are from Helgeson (1969). Calculation method is given in Shikazono (1978a). Chloride concentration in hydrothermal solution is assumed to be i moI/kg H20. A-B: Na + concentration in solution in equilibrium with low albite and adularia, C-D: K + concentration in solution in equilibrium with low albite and adularia, E-F: H4SiO4 concentration in equilibrium with quartz, G-H: Ca 2+ concentration in equilibrium with albite and anorthite (Shikazono, 1978a, 1988b).

Miocene-Pliocene Hydrothermal Ore Deposits

123

depend on alteration minerals, temperature and C1- concentration; (2) K + and H4SiO4 concentrations increase with an increase in temperature, while Na + concentration does not largely depend on temperature. Calculations were made, assuming that the dominant anion is C1-. From these calculations, it is considered that the potassic alteration occurs when hydrothermal solution initially in equilibrium with propylitic alteration minerals ascends rapidly and interacts with country rocks at lower temperature ( I - J - K in Fig. 1.86). In this case, addition of K + to the rock takes place. K-bearing minerals such as sericite and K-feldspar precipitate from the fluid accompanied by the destruction of plagioclase in the country rocks and liberation of Ca, Sr, and Na to the fluids. Dissolutions of anorthite and albite components in a plagioclase occur by this mechanism. Thus, it is likely that Ca, Sr and Na are extracted from the rocks. It is expected that SiO2 content of the country rocks increases with progressive alteration because solubility of SiO2 decreases with a decrease in temperature ( O - P - Q in Fig. 1.86) (e.g., Holland and Malinin, 1979). If fluids initially in equilibrium with quartz ascend rapidly, some metastable minerals (amorphous silica, cristobalite, wairakite) may precipitate because of supersaturation with respect to SiO2 (e.g., Wolery, 1978; Bird and Norton, 1981). Important processes for the supersaturation and deviation from the equilibrium between fluids and rocks are adiabatic boiling, mixing of fluids and conductive cooling of fluids (Giggenbach, 1984). Formation of albite which is characteristic mineral of propylitic alteration occurs by heating of rocks and descending fluids at recharge zone in the hydrothermal system (Giggenbach, 1984; Takeno, 1989). Thus, it is considered that the propylitic alteration takes place at recharge zone in the hydrothermal system, while potassic alteration at discharge zone. Besides the effect of temperature, boiling and degassing play an important role for potassic alteration. Degassing of CO2 is accompanied by an increase in pH, causing the depositions of adularia and calcite (Browne, 1978). Probably, formations of some zeolite minerals (wairakite, yugawaralite, taumontite) are also caused by degassing and boiling of fluids and increase in pH (Shikazono, 1985a). There are two important chemical reactions which cause intermediate argillic and advanced argillic alterations H2S + 202 = 2H + + SO]

(1-35)

4SO2 + 4 H 2 0 = H2S + 3 H2SO4

(1-36)

Oxidation of H2S (reaction (1-35)) occurs under the near-surface environment. Oxygen may be supplied from oxygenated groundwater. These oxidation reactions liberate H + ion, leading to a decrease in pH. Under low pH conditions intermediate argillic alteration minerals (e.g., kaolinite, sericite) are stable. When temperatures of volcanic gases containing SO2 decrease, the reaction (1-35) proceeds to the right hand side. This reaction causes a considerable decrease in pH due to the formation of sulfuric acid. Advanced argillic alteration is formed by the interaction of volcanic gas with groundwater. The above interpretation on the alteration zoning is mainly based on thermodynamics. However, it is necessary to consider the influence of kinetics and fluids flow on

124

Chapter 1

the hydrothermal alteration processes to interpret the precipitations of metastable phases such as cristobalite. The coupled precipitation kinetics-fluid flow model was applied to the distribution of SiO2 content and K20 content of the hydrothermally altered andesite in the Hishikari A u - A g mine area, south Kyushu, Japan by Shikazono et al. (2002). This will be described in section 1.4.6. 1.4.3. Geochemical characteristics Numerous geochemical data (fluid inclusions, stable isotopes, minor elements) on the epithermal vein-type deposits in Japan are available and these data can be used to constrain geochemical environment of ore deposition (gas fugacity, temperature, chemical compositions of ore fluids, etc.) and origin of ore deposits.

1.4.3.1. Fluid inclusions Substantial amounts of homogenization temperature data on the Neogene vein-type deposits in Japan are available (e.g., Enjoji and Takenouchi, 1976; Shikazono, 1985b) and they are summarized in Fig. 1.87. Homogenization temperatures vary widely within a given deposit type and even within a single deposit. However, the range of homogenization temperatures differs according to the type of deposit: 190°C 145 250oc for Au-Ag-rich deposits, 200°C to 250°C for Pb-Zn-Mn-rich deposits, and 200 ° to 350°C for Cu-Pb-Zn-rich deposits (Fig. 1.87) (Shikazono, 1985b). Homogenization temperatures are not same as the formation temperature. Therefore, we need pressure correction to estimate the formation temperatures from homogenization temperatures. However, the homogenization temperatures are in good agreement with electrum-sphalerite temperatures (Shikazono, 1985d) (Figs. 1.88 and 1.89). Therefore, pressure corrections to homogenization temperatures of fluid inclusions necessary to obtain formation temperatures are relatively small (less than 30°C). Salinities of inclusion fluids from epithermal vein-type deposits clearly indicate that the salinities of inclusion fluids from these types of deposits are distinctly different, that is, 20-2 NaC1 equivalent wt% (base-metal vein-type deposits) and 0-3 wt% (Au-Ag vein-type deposits) (Shikazono, 1985b) (Table 1.13). Salinities of inclusion fluids from Kuroko deposits (0.5-5 wt% NaC1 equivalent concentration) are between these two types of deposits. This kind of difference is observed in epithermal deposits in other countries (Hedenquist and Henley, 1985).

1.4.3.2. Estimate of temperatures from the electrum-sphalerite-pyrite-argentite assemblage As noted already, the Ag content of electrum in equilibrium with argentite and FeS content of sphalerite in equilibrium with pyrite are expressed as a function of fs2 and temperature, we can estimate temperature from Ag content of electrum and FeS content of sphalerite.

125

Miocene-Pliocene Hydrothermal Ore Deposits Filling Temperature ('C) Au-Ag Dep.

150

Sanru Teine-Takinosawa Todoroki-Chuetsu Todoroki-Shuetsu Chitose-Daikoku Chitose-Daikoku No2 Chitose-Fukujin Yatani-Kanizawa Takatama Sado-Ohdachi Nebazawa Seigoshi Taio Fuke-Honpi Ohkuchi Hishikari Arakawa Kushikino

[

200 i

•

I

•

•

350

3OO ] ]

250 ]

I

I I

I

•

I

•

¢

:

I

•

I

•

I

Pb-Zn'Mn Dep. Ohe-Senzai Inakuraishi Toyoha-Tajima Toyoha-Hadma Toyoha-lzumo Yagumo-Ohgid

•

I •

I

6

Cu" Pb oZn Dep. Oppu Osarizawa AnMnari Hosokura-Shoko Hosokura-Ohtake Hosokura-Hakuho Ohizumi Yatani-Honpi Nanetsu Tochigi Ogoya Taishu-Okutomi Taishu-Shinotomi Taishu-Himi Taishu-Taisho Taishu-Misoge Taishu-Tsurue Taishu-Akushidani Taishu-Shirocake Taishu-Amanohara

•

I

¢

¢

•

I

I •

]

I

[

•

]

I

Figure 1.87. Summary of filling temperatures of fluid inclusions from Neogene vein-type deposits in Japan. Solid circIe represents average filling temperatures of fluid inclusions for individual deposits (Shikazono, 1985b).

126

Chapter 1 350

300

_~ 250

g

200

J1 ~

'I l

150

100

150

200 250 300 350 Electrum-Sphalerite Temperature (°(3)

400

Figure 1.88. Electrum-sphalerite temperatures vs. homogenization temperatures of fluid inclusions from epithermal Au-Ag vein-type deposits in Japan. Electrum-sphalerite temperatures were calculated from the iron content of sphalerite and the silver content of electrum (Equation 6 in Shikazono (1985a)). The average and range of homogenization temperatures and electrum-sphalerite temperatures for a given ore deposit are represented by a solid circle and line, respectively. 1 = Ohmidani-Fusei, 2 = Todoroki-Chuetsu, 3 = TodorokiShuetsu, 4 = Taio no. 9, 5 = Toyoha-Tajima, 6 = Nawaji, 7 = Seigoshi no. 2, 8 = Toi, 9 = Yugashima, 10 = Yatani-Tengu, 1l = Yatani-Kanizawa, 12 = Oh~Senzai, 13 = Ohe-Senzai, 14 = Sado, 15 = NebazawaManzai no. 3, 16 = Kamioka, 17 = Chitose-Daikoku, 18 = Yatani Honpi, 19 = Yunoura (Shikazono, 1985a). In order to obtain reliable formation temperatures based on electrum-sphalerite geothermometer, the following conditions have to be satisfied. (1) The coexisting sphalerite, pyrite, electrum, and argentite must have been at e q u i l i b r i u m at the time of their precipitation. Although it is difficult to evaluate this condition, it is c o m m o n l y observed that these minerals are in direct contact with each other without evidence o f mutual replacement texture. Therefore, it is likely that these minerals have been precipitated nearly contemporaneously. (2) The FeS content o f sphalerite and the Ag content of electrum have not c h a n g e d considerably during the post-depositional period. It is unlikely that the sphalerite c o m p o s i t i o n c h a n g e d d u r i n g the cooling stage, because this mineral is one of the most refractory sulfide minerals. In general, the Ag c o n t e n t of electrum with a large grain size increases from core to margin. Such regularity o f compositional z o n i n g observed in a large grain, wide c o m p o s i t i o n a l variation in a large grain, and the relationship between grain size and compositional range in a grain suggest that the electrum composition is retained d u r i n g the post-depositional period. The A u - A g interdiffusion coefficient for electrum has b e e n experimentally studied ( C z a m a n s k e et al., 1973). The values of the A u - A g interdiffusion coefficients, the compositional z o n i n g pattern in electrum, and the temperature range of c o n c e r n (ca. 2 0 0 - 3 0 0 ° C ) suggest that the post-depositional change

127

M i o c e n e - P l i o c e n e Hydrothermal Ore Deposits

-5

-10

Y~

t~

0 _..1

-15

i

v

I

200

I

t

i

I

!

250

i

i

i

Temp.(°C)

I

300

i

Figure 1.89. Activity of S2(as~)-temperature diagram showing possible as2 and temperature ranges for epithermal Au disseminated-type (hot spring type), epithermaI Au-Ag vein-type and epithermaI base metal vein-type deposits in Japan (Shikazono 1986; Shikazono and Shimizu, 1988b). TABLE 1.13 Filling temperature and NaCI eq. concentration of fluid inclusions from epithermaI gold-silver and base-metal vein-type deposits (Shikazono and Shimizu, 1992) Deposit

Filling temprerature (°C)

NaC1 eq. concentration (wt%)

Gold-silver type

Sado Seigoshi Yatani Ohguchi Todoroki Koryu Chitose Kushikino

305-190 243-178 273-209 265-184 24(~ 122 300-140 300-220 250-210

2.5-1.0 2.8~0.0 1.5~).5 1.64).0 1.7-0.4 1.4-0.0 2.0-1.0 1.1~0.6

Base-metal type

Toyoha Oppu Osarizawa Hosokura Nanetsu Taisyu Asahi Ani Oe Jokoku

300-150 330-170 268-156 231 - 130 306-250 376-150 275 250 266-207 310-145 250-125

4.24).2 i8.3-1.7 7.5~0.0 9.1 ~).0 7.9-4.4 33.5~.0 10.3-8.8 12.0 2.0-0.0 6.0-3.0

128

Chapter 1

in electrum composition may be negligible. One of the most likely explanations for the compositional zoning in a grain is that it reflects the change in chemical parameters of ore fluids during the precipitation of electrum. (3) The value of the activity coefficients of FeS in sphalerite determined for temperatures above 300°C can be extrapolated to lower temperatures. As stated by Barton and Toulmin (1966), ?/FeS does not depend on temperature above about 270°C. However, the activity coefficient below 270°C has not been studied. Scott and Kissin (1973) have stated that activity coefficients for FeS in sphalerite at low temperatures may be substantially different from those at higher temperatures. (4) The effects of impurities such as Mn and Cd in sphalerite on the equations are negligible. Generally, the concentrations of minor elements in sphalerite from the epithermal A u - A g vein-type deposits in Japan, except for iron, are very small (Cd less than 1 wt%, Mn less than n 1 x 10 -1 w t % , Cu less than n x 10 -1 wt%, the concentrations of other elements are also less than n x 10 .2 wt%; e.g., Shikazono, 1978b). These low concentrations affect the fs2-temperature relations in the F e - Z n - S system (e.g., Barton and Toulmin, 1966). Impurities in electrum such as copper and antimony are also very small, generally less than 1 wt%. Therefore, it is likely that these elements do not affect the thermochemical properties of electrum. Argentite from epithermal Au-Ag vein-type deposits in Japan contains sometimes up to 10 wt% selenium (Shikazono, 1978b). However, if the selenium contents are small, it is likely that the activity coefficient for Ag2Se in argentite does not deviate significantly from unity, because complete solid solution between Ag2Se and Ag2S exists above ca. 180°C (Sugaki et al., 1982); However, the absolute value of the activity coefficient for Ag2S in argentite cannot be determined. Impurities in pyrite such as nickel and cobalt are also very low (less than n x 10 -1 wt%), and thus, an activity coefficient of FeS2 in pyrite equal to one can be safely assumed. (5) The effect of pressure is negligible. These epithermal A u - A g vein-type deposits have formed in a shallow and low-pressure environment and pressure correction, such that any correction to the homogenization temperatures will be small (probably less than 20°C). The temperatures estimated from the electrum-sphalerite-pyrite-argentite assemblage are plotted versus homogenization temperatures of fluid inclusions as shown in Fig. 1.88. Although the electrum-sphalerite temperatures and homogenization temperatures from given veins for each deposit have some variation, the averages of the electrum-sphalerite temperatures show a good correlation with the average homogenization temperatures. Most electrum-sphalerite average temperatures correlate to within 30°C of the respective average homogenization temperatures. This good correlation can be seen in Fig. 1.88. The good correlation between homogenization temperatures and electrumsphalerite temperatures suggests several points: (1) the uncertainties of electrumsphalerite temperatures are less than 20 ° to 30°C, even at temperatures from ca. 180 ° to 300°C, (2) the electrum-sphalerite-pyrite-argentite assemblage was formed close to equilibrium in Japanese epithermal Au-Ag vein-type deposits, and (3) the pressure corrections to homogenization temperatures for Japanese epithermal Au-Ag vein-type deposits is small, less than 20°C to 30°C. t All n in this text book as natural number 1 through 9.

Miocene-Pliocene Hydrothermal Ore Deposits

129

1.4.3.3. Gas fugacities As will be mentioned in section 2.4.3, i s 2 c a n be estimated based on the Ag content of electrum coexisting with argentite (or acanthite), the FeS content of sphalerite coexisting with pyrite and temperature estimated from homogenization temperatures of fluid inclusions. Figures 1.68 and 1.69 show the FeS content of sphalerite and the Ag content of electrum from epithermal Au-Ag vein-type, epithermal base-metal vein-type and Kuroko deposits, indicating different fs2-temperature regions for these types of deposition. Figure 1.89 shows typical range of fs2 and temperature for epithermal basemetal vein-type and Au-Ag vein-type deposits. It is noteworthy that the ranges of fs2 for epithermal Au-Ag, epithermal Au-bearing base-metal, and epithermal Au-free base-metal vein-type deposits are different, while temperatures are not different. As mentioned already, small amounts of electrum occur in epithermal base-metal vein-type deposits. Electrum is not observed in the epithermal base-metal vein-type deposits in which pyrrhotite occurs (e.g., Toyoha-Soya, Oizumi, and Hosokukura Pb-Zn deposits). However, electrum is found in epithermal base-metal vein-type deposits in which hematite is commonly observed (e.g., Osarizawa and Ani C u - P b - Z n deposits). This indicates that electrum precipitates in relatively high fs2 and f02 condition.

Sulflirfugacity (fs2)

Oxygenfugacity (fo2). The fo2-pH diagrams (Figs. 1.90 and 1.91) were constructed at 200°C and 250°C based on the homogenization temperatures and electrum-sphalerite temperatures (Shikazono, 1985d). The ionic strength and activity coefficients of aqueous species are estimated from the freezing temperature of fluid inclusions. The ionic strength is assumed to be 1. Estimates of the total dissolved sulfur concentrations of ore fluids (ZS) responsible for several veintype deposits in Japan are of the order of 10-2-10 -3 mol/kg H20 (Shikazono, 1974b; Hattori, 1975). Chemical analyses of the hot springs accompanied by epithermal basemetal depositions (White, 1967; Browne and Ellis, 1970; Mayhon and Finlayson, 1972; Weissberg et al., 1979) give the values of 10 2-10-3 mol/kg H20. Therefore, the total dissolved sulfur concentration (ZS) was assumed to be 10-2-10 -3 mol/kg H20 for the construction of the diagrams. Gangue minerals and salinity give constraints on the pH range. The thermochemical stability field of adularia, sericite and kaolinite depends on temperature, ionic strength, pH and potassium ion concentration of the aqueous phase. The potassium ion concentration is estimated from the empirical relation of Na+/K + obtained from analyses of geothermal waters (White, 1965; Ellis, 1969; Fournier and Truesdell, 1973), experimental data on rock-water interactions (e.g., Mottl and Holland, 1978) and assuming that salinity of inclusion fluids is equal to mNa+ + mK+ in which m is molal concentration. From these data potassium ion concentration was assumed to be 0.1 and 0.2 mol/kg H20 for 200°C and 250°C. By giving the values of temperature, ZS, ionic strength, FeS content of sphalerite and Ag content of electrum, we can place a limit on fo2. However, we cannot know whether fo2 lies in the predominance field of reduced sulfur species or that of oxidized sulfur species from the constraints mentioned above.

130

Chapter 1 log fo 2 -35

i I

"".,,..

!

Ka ~, Se IAd i Hm -~,~,~_ i ,.~-r ---4 Mt I

B n,, Py I

~"~:L~."~-,~~ /

..A ~g - .

)

:%sit

|

3

.

PY . . . . .-.'.- - - F Po

I

4

I

5

,

l i

I

, ,

-o"~. ~"

o,

I

I

6

!

I

?

, l

/

8

I

9

I

10

pH

Figure 1.90. fo2-pH diagram constructed for temperature = 200°C, ionic strength = 1, and ]ES = 0.01 moI/kg H20. Solid lines (1), (2), (3), and (4) are, respectively, (1) 10 tool% FeS in sphalerite; (2) 70 tool% Ag in electrum, (3) 60 mo1% Ag in electrum, and (4) 1 tool% FeS of sphalerite; 0.1 tool% FeS sphalerite is nearly coincident with the Bn + Py/Cp boundary. The dotted and shaded areas A and B represent the possible foz-pH regions for Au-Ag vein-type (Yatani, Seigoshi, Omidani, Asahi) and Pb Zn vein-type (Toyoha, Yatani, Ikuno) deposition, respectively. Ka: kaolinite, Se: sericite, Ad: adularia, Hm: hematite, Mt: magnetite, Bn: bornite, Py: pyrite, Cp: chalcopyrite, Arg: argentite, Nsil: native silver, Po: pyrrhotite (Shikazono, 1978b). The fo2 o f ore fluids responsible for the epithermal base-metal veins might have been in the predominance field o f reduced sulfur species because ( l ) pyrrhotite is occasionally found in these deposits, (2) selenium content of argentite is very low and (3) H2S is dominant in the present-day epithermal base-metal fluids. Implication of selenium content of sulfides will be considered later. Barite is sometimes found in the late-stage of mineralization. Thus, it is likely that fo2 of barite stage lies in the predominance field of oxidized sulfur species. As already discussed, fo2 o f Kuroko ore fluids is considered to lie in the predominance field of reduced sulfur species from the following two reasons; (1) Selenium content of sulfides is very low (Yamamoto, 1974) and (2) H2S is dominant in hydrothermal solution venting from back-arc basins (section 2.3) from which hydrothermal ore deposits being similar to Kuroko deposits form. On the other hand, the ore fluids responsible for epithermal A u - A g vein-type deposits contain appreciable amounts of oxidized sulfur species, together with reduced sulfur species. Oxidized sulfur species/reduced sulfur species ratio is considered to be greater than 1, that is, f o 2 lies in the predominance field of oxidized sulfur species. The reasons for this estimation are: (1) hematite is c o m m o n in the deposits, (2) barite is found

131

Miocene-Pliocene Hydrothermal Ore Deposits

log Io 2 -30 ~

Ka] Se i Ad |

~"'-

'

Hm

~ + ~ - ? - 4 5 - , , , B n + P y "~,

.-'-(3)_ -(2)

Mt

CpN ~

--(I) - 4 0 ;__Arg____~' . Nsi

I

3

.

I PY

.

.

I

, I I

I I 11

4

Po

.

\

,

=

l

I

5

I

6

I

7

I

8

I

\

9

f

10

pH

Figure 1.91. fo2-pH diagram constructed for temperature = 250°C, ionic strength = 1, and lgS = 0.01 mol/kg H20. Is~FeS mole percent lines for sphalerite and the stability relations of some hydrothermaI minerals are also given. The solid lines (1), (2), and (3) are respectively 10, 1, and 0.! mol% FeS in sphalerite; 70 and 60 mol% lines of Ag in electrum are nearly identical to lines (i) and (2). The dotted and shaded areas of A and B represent the possible fo2-pH regions for Au-Ag vein4ype (Yatani, Seigoshi, Omidani, Asahi) and Pb-Zn vein-type (Toyoha, Yatani, Ikuno) deposition, respectively. Bn: bornite, Py: pyrite, Cp: chalcopyrite, Arg: argentite, Nsi: native silver, Hm: hematite, Mr: magnetite, Po: pyrrhotite, Ka: kaolinite, Se: sericite, Ad: adularia (Shikazono, 1978b). in Te-type deposits, although this mineral formed at late stage and not found in Se-type deposits, (3) selenium is usually contained in the deposits, (4) tellurium is concentrated in Te-type deposits, and (5) sulfate ion is abundant, together with H2S in geothermal waters, associated with epithermal A u - A g depositions (section 2.1). Selenium and tellurium contents of sulfides are useful indicators for estimating f Q (Shikazono, 1974a, 1978b). The selenium content o f sulfide is governed by the reaction, (MS) + Se 2 - = (MSe) + S 2 -

(1-37)

where (MS) and (MSe) are MS and MSe components in the sulfide-selenide solid solution M ( S I - x , Sex), respectively. Equilibrium constant o f this reaction (K1-37) is, K1-37 = ( a M S e a s 2 ) / ( a M s a s e 2

).

(1-38)

where a is activity. Therefore, if the activity coefficient ratio, gMSe/gMS, does not deviate from unity and the temperature is constant, the ratio, mMse/mMS, correlates to ase2-~as2- which is represented as functions of f02, pH, temperature, and l ~ S e / 2 S (Shikazono, 1978b).

132

Chapter 1

Log (asga~-)

-1 -2 -3

1

-4

!

-5 -6

A

I

,

I

-7

J

-8 -9

i

(~)

!(2)

(3)

I

I

(4)

I(5) I

I

! i

i

i

11

i

7 2345678 Figure 1.92. Dependence of the ratio aSe2 ~as2

i

|

i

,

,i

i

i

~

,

91011121314

,

pH

pH under the conditions: temperature = 150°C, ionic strength = 1, li~S = 0.01 mol/kg H20, ~2Se = 10-7 mol/kg H20, and ZS = ]~rnr (total reduced sulfur content). (1) H2S-H2Se region, (2) H2S-HSe region, (3) HS HSe region, (4) S2 -HSe- region and (5) S2--Se 2 region (Shikazono, 1978b). on

Figure 1.92 represents the pH dependence of ase2 /as2- in the predominance region of reduced sulfur species. Considering the equation mentioned above, it is thought that this activity ratio correlates well with the selenium content of sulfide, if the effects of the activity of the MSe and MS components in sulfide (M: metal element) are neglected. The activity ratio, ase2 ~as2 , is expected to be relatively high in the sericite region and low in the adularia region. However, the selenium content of acanthite from epithermal A u - A g vein-type deposits in which adularia is a common constituent is higher than that from the epithermal Pb-Zn vein-type deposits in which sericite is the common wall-rock alteration product. Therefore, the difference in the selenium content of acanthite from these deposits cannot be explained by the difference in pH. As the values of the activity ratio, ase2 ~as2 , on the f o 2 - p H diagrams at 150°C (Fig. 1.93) and 300°C (Fig. 1.94) are nearly identical, changes in temperature alone cannot be responsible for the variation in the selenium content, if the temperature dependence of the equilibrium constant, KI-37, is not large. In contrast to the effects of pH and temperature, fo2 has a great effect on the selenium content of sulfides, especially in the predominance region of oxidized sulfur species (Figs. 1.93 and 1.94), where the activity ratio is relatively high and increases rapidly with increasing fo2 at a fixed pH. In the predominance region of reduced sulfur species, the ratio is constant at a fixed pH and has a relatively low value. The wide variation in the selenium content of acanthite from epithermal A u Ag vein-type deposits can be explained by assuming that the acanthite formed within the predominance region of oxidized sulfur species at a constant I2Se/~2S ratio, ionic strength, temperature and pH (Figs. 1.93 and 1.94). The very low selenium content of acanthite from epithermal Pb-Zn vein-type deposits suggests the predominance region of reduced sulfur species (Figs. 1.93 and 1.94).

133

Miocene-Pliocene Hydrothermal Ore Deposits Io Ifo2 -30

-35 HSO~ (Na, K)SO4

-40 _ ~ -4s

__

~50

H2S

"-'

',HS~ S~-

H2SIHS

HS S; ;I

. . . . . . . . . .

i

i

2-

[

i

1 2 3 4 5 6 7 8 9 1011 121314

pH

Figure 1.93. foz-pH diagram with the stability fields of aqueous species in Na-K-H-S Se-O system for the conditions: ]ES = 10 -2 moI/kg H20, ~2Se = 10 7 mol/kg H20, ionic strength = 1, and temperature = 150°C. Dashed lines are the ratio iso-ase2 /as2 in logarithmic units. Stability fields for native sulfur and native selenium and the boundaries between predominance regions of oxidized and reduced selenium species are omitted for clarity (Shikazono, 1978b).

Iogfo2

-20 -25

HSO~ (Na, K)SO~

-30 -35 -40 -45

_0o

[

/-ii I--

. f

r

. =

i

HS:;

- -,

1 2 3 4 5 6 7 8 9 1011 121314

pH

Figure 1.94. fo2-pH diagram with the stability fields of aqueous species in the N a - K - H - S - S e - O system for the conditions: NS = 10 2 mol/kg H20, ]ESe = 10 7 mol/kg H20, ionic strength = 1, and temperature = 300°C. Dashed lines indicate iso-ase2 ~as2 contours in logarithmic units (Shikazono, 1978b).

134

Chapter 1

Log

a Carb

ZnCO3

1

O. -2' -4" -6' -8" -10,

i

U (2) -5o

-~

(3) t 6g fo 2

Figure 1.95. Activity of component ZnCO3 versus fo2- Carbonate containing ZnCO3 is in equilibrium with sphalerite. Thermochemical calculation was made under the following conditions: temperature = 200°C, ionic strength = 1, E;S = 10 2 m, and pH = 5. (1) CH4 and H2S region. (2) H2CO3 and H2S region. (3) H2CO3 and (Na, K) SO]- region (Shikazono, 1977b). In contrast to the estimation of total dissolved sulfur concentration (ES), the total dissolved selenium concentration (ESe) has not yet been estimated. The Beppu hot springs which are accompanied by A u - A g siliceous sinter, contain about 10 -7 tool/1 Se (Uzumasa, 1965). For the construction of Figs. 1.93 and 1.94, the Se/S ratio of the ore fluids is assumed to be 10 -5 . The difference in selenium content of acanthite from the different types of ore deposits can, of course, also be explained by the difference in Z S e / E S in the ore fluids; i.e., the ore fluids responsible for the formation o f epithermal A u - A g vein-type deposits may have had a higher E S e / E S ratio than that for epithermal P b - Z n vein-type deposits. It is likely that considerable amounts of sulfur were derived from marine rocks (Green tuff) and were incorporated into ore fluids for base-metal veins. The zinc content of siderite was applied to estimate fo2 of ore fluid by Shikazono (1977b). The analytical data on the late-stage siderites coexisting with barite, pyrite and hematite from the Ohmori epithermal A u - A g vein-type deposits that zinc content is high in the range of 0.8-5.8 wt% as ZnO, but the siderite of the early stage of mineralization from the Ohmori and from the Toyoha epithermal P b - Z n - A g vein-type deposits does not contain such a large amount of zinc. The relation between zinc content in carbonate and many physicochemical variables such as f % , temperature, pH, and so on was derived on the basis of the equilibrium between coexisting carbonate and sphalerite (Figs. 1.95 and 1.96). It is predicted theoretically that zinc content increases with increasing fo2 in the oxidized sulfur species and oxidized carbon species region (Fig. 1.95). This estimate is consistent with mineral assemblage containing the siderite. The iron content of sphalerites coexisting with siderite and pyrite from the Toyoha deposits is high (3-12 wt%). On the other hand, sphalerite of later stage Ohmori

135

Miocene-Pliocene Hydrothermal Ore Deposits Log fo 2 -35 HSO41(Na,K)SO 4

HCO~ CO~-

H2C03 HC03

I

+F;; • p ;..=,~~ . ,.,-:..:. . . . . . . . . . . . ... .t. . . . . . .

;

-40

. . . ~ . : ~ - - -- 2(2)

/

Bn+P, ~ "L ""~; ' ~ ' ' " (3) | ..................... ,,. . . . ~ '.':..( Na, K)S02 Cp H~CO~ ! i I I "~ ~ Z-45 D,,

"I

--

: :, I ' .....

"°

-5o

1

i

i

Ka Se Se:A i

4

~.-.aE~!Bn 2 ',~.-... CO.C ',,1.'>E-..J ~I ;! C ~" HL 4 TM Po',M t

H2s HS- Hg S2-

t 2

} ] CH4 : I -"~. :, j ._,. I i J-~ ~" '

5

6

7

\

\

! g

lb

pH

Figure 1.96. Log foz~H diagram constructed for temperature = 200°C, ionic strength = 1, }IS 10 -2 m, and P,C = 10 tm. Solid line represents aqueous sulfur and carbon species boundaries which are loci of equal molalities. Dashed lines represent the stability boundaries for some minerals. Ad: adularia, Bn: bornite, Cp: chalcopyrite, Ht: hematite, Ka: kaolinite, Mt: magnetite, Po: pyrrhotite, Py: pyrite, Se: sericite. Heavy dashed lines (1), (2), and (3) are iso-activity lines for ZnCO3 component in carbonate in equilibrium with sphalerite: (1 "~ a carb = 0 . 1 . (2) carb carb -(Shikazono,1977b). J ZnCO3 azncO3= 0 . 0 1 . (3) azncO 3-0.00I =

deposits contains small amounts of iron (0.1-0.4 wt%), and the content is narrow in range (although the iron content in early-stage is higher; 0.4-6.2 wt%). The decrease o f iron content from the early to the later stage indicates that oxidation occurred a n d / o r temperature decreased during the formation of the Ohmori deposits. The difference in iron content for these deposits means that the Toyoha deposits formed under relatively lower fo2 conditions than the early stage Ohmori deposits. These conclusions derived from the analytical data on coexisting siderite and sphalerite are consistent with the studies on the chemical environment of epithermal vein-type deposits in Japan (Shikazono, 1973, 1974a; Hattori, 1975).

Carbon dioxidefugacity (fc02), The fCO2 values can be estimated from (1) gangue mineral assemblages including carbonates and (2) fluid inclusion analyses. Shikazono (1985b) summarized the assemblage and mode of occurrence o f common gangue minerals from more than 70 Neogene epithermal vein-type deposits in Japan.

Chapter 1

136

3

J

2 1

~o 0

M , , ~ k , , ~_

/i

.,d

-2 -3 -4

150

200

250

300 Temperature (°C)

350

Figure 1.97. log fco2-temperature diagram showing the univariant equilibrium curves for some gangue minerals. A. 2Ca2AI3Si3012(OH) (clinozoisite) + 3 SiO2 (quartz) + 2CACO3 (calcite) + 2 H20 = 3 Ca2Al2 Si3OI0(OH)2 (prehnite) + 2CO2 (Xpis = 0.30, Xpis: mole fraction of pistacite component in epidote). B. Ca6Si60]7(OH)2 (xonotlite) + 6CO2 = 6CACO3 (calcite) + 6SIO2 (quartz) + H20. C. CaCO3 (calcite) q- TiO2 (rutile) + SiO2 (quartz) = CaTiSiO5 (sphene) + CO2. D. MnSiO3 (rhodonite) + CO2 = MnCO3 (rhodochrosite) + SiO2 (quartz). E. 3KA13Si3OIo(OH)2 (K-mica) + 4CACO3 (calcite) + SiO2 (qua~z) = 2Ca2A13Si3OI2(OH) (clinozoisite) + 3KAISi308 (K-feldspar) + 4CO2 4- 2H20 (Xpis = 0.3). P. 3KAI3Si3Om(OH)2 (K-mica) + 4CACO3 (calcite) + SiO2 (quartz) = 2Ca2AI3Si3OI2(OH) (clinozoisite) + 3KAISi308 (K-feldspar) + 4CO2 Jr- 2H20 (Xpis = 0.25). G. 3FeCO3 (siderite) + (1/2) 02 = Pe304 (magnetite) + 3CO2, C (graphite) + 02 = CO2. H. 3CaMg(CO3)2 (dolomite) + KAISi308 (K-feldspar) 4- H20 = 3CACO3 (calcite) + 3CO2 4- KMg3(AISi3Om)(OH)2 (phlogopite). I. C (graphite) + 02 = CO2, FeS (pyrrbotite) + 1/2S2 = FeS2 (pyrite), 2H2S(aq) + 02 = $2 4- 2H20(I). J. FeCO3 (siderite) + Fe203 (hematite) = Fe304 (magnetite) + CO2. K. CaA12Si4OI2-2H20 (wairakite) + KAISi308 (K-feldspar) + CO2 = CaCO3 (caIcite) + KAI3Si3Oto(OH)2 (K-mica) + 4SIO2 (quartz). L. CaAI2Si4OI22H20 (wairakite) + CO2 = CaCO3 (calcite) + AI2Si2Os(OH)4 (kaolinite) + 2SIO2 (quartz). M. CaAI2Si4OI2-4H20 (laumontite) + CO2 = CaCO3 (calcite) + AI2Si205(OH)4 (kaolinite) + 2 SiO2 (quartz) + 2 H20 (Shikazono, 1985b).

As already mentioned the major gangue minerals vary with different deposit types; quartz, chalcedonic quartz, adularia, calcite, smectite, interstratified mica/smectite, interstratified chlorite/smectite, sericite, zeolites, and kaolinite in Au-Ag deposits, chlorite, quartz, sericite, calcite, rhodochrosite, siderite and (magnetite) 2 in Pb-Zn-rich deposits, chlorite, sericite, siderite, hematite, magnetite and (epidote) in Cu-rich deposits. Based on the gangue mineral assemblage (Fig. 1.97), homogenization temperatures of fluid inclusions, thermochemical calculations (Fig. 1.98), and analytical data on fluid inclusions (Fig. 1.99), typical ranges of fco2 for the Au-Ag, Pb-Zn-Mn and Cu-Pb-Zn vein-type deposits were determined to be 10 .3 to 1 atm (190°C-250°C) l0 -I to 10 atm (200-250°C) and 10- I to 103 atm (200°C-350°C), respectively (Fig. 1.100). The fco2 of Kuroko ore fluids is close to that of Cu-Pb-Zn vein-type deposits. The fco2 values could be estimated based on FeCO3 content of carbonates coexisting with iron minerals (pyrite, hematite, magnetite, pyrrhotite) and minerals containing 2 Mineral in parentheses occurs in small amounts in each deposit type.

137

Miocene-Pliocene Hydrothermal Ore Deposits 3 2

# o

1 0

~-1

._1

-2 -3 -4

150

200

250

300 Temperature (°C)

350

Figure 1.98. Summary of fcQ-temperature ranges for Au-Ag-rich, Pb-Zn Mn-rich, and Cu-Pb-Zn-rich vein-type deposits based on gangue mineral assemblages and fluid inclusion data. Line a-b: Equilibrium among graphite, pyrite, and pyrrhorite. Line c~l: fc02 vs. temperature curve for fluid containing i tool% CO2 with temperature of first boiling of 300°C. Line e-f: "Plagioclase" + CO2 = calcite + "kaolinite". Line g ~ : fco2 temperature relation for geothermal reservoir waters obtained by Arn6rsson (1984). Line i-j: f c Q vs. temperature curve for fluid containing 1 tool% CO2 with a temperature of first boiling of 350°C. Line ~1: 2Ca2A13Si3012(OH) + 3KAISi308 + 4CO2 + 2H20 = 4CACO3 + 3KAI3Si3OI0(OH)2 + 6SIO2; mole fraction of KAI3Si3010(OH)2 in mica = 0.6, mole fraction of 2Ca2Fe3Si30[2(OH) in epidote = 0.2 (Shikazono, 1985b). 31

yA

2! 04

O o

1

S

0

C

P

o

_.1 -2 -3 -4

1;0

2;o

2;0

330

T e m p e r a t u r e (°C)

Figure 1.99. Estimated fco2-temperature ranges from anaytical data on fluid inclusions and homogenization temperatures (Shikazono, 1986). T: Taishu (Pb, Zn), O: Ohizumi (Cu, Pb, Zn), Y: Yatani (Pb, Zn), Os: Osarizawa (Cu, Pb, Zn), H: Hosokura (Pb, Zn), C: Chitose (Au, Ag), S: Seigoshi (Au, Ag). iron as solid s o l u t i o n ( i r o n c o n t e n t in s p h a l e r i t e ) . S h i k a z o n o ( 1 9 7 4 a ) t h e o r e t i c a l l y d e r i v e d the r e l a t i o n s h i p b e t w e e n iron c o n t e n t o f c a r b o n a t e s c o e x i s t i n g w i t h i r o n m i n e r a l s as f u n c t i o n s o f f c Q , f Q , a n d others. W e c o u l d e s t i m a t e f c Q u s i n g his m e t h o d . B u t n o q u a n t i t a t i v e a p p l i c a t i o n h a s b e e n c a r r i e d out.

138

Chapter i 3 2

1

020 o,-1 -2 -3 -4

150

200

250

300 350 Temperature (°C)

Figure 1.100. Typical fco2-temperature ranges for Au-Ag-rich, Pb-Zn-Mn-rich, and Cu-Pb-Zn-rich veintype deposits estimated from gangue mineral assemblages, homogenization temperatures of fluid inclusions, and thermochemical calculations (Shikazono, 1985b).

Seleniumfugacity (fse2).

If coexisting electrum, sphalerite, pyrite, argentite and galena are in equilibrium, the relationship between the Ag content of electrum, selenium contents of argentite and galena, iron content of sphalerite, temperature, fsz and fse2 can be derived from the equilibrium relations for the chemical reactions given in Table 1.14. Scott and Barnes (1971) have obtained an equation representing FeS content of sphalerite in equilibrium with pyrite as functions of fsz and temperature. By combining equations (1)-(5) in Table 1.14 and the relation between activity coefficient of Ag in electrum, Ag

TABLE 1.14 Chemical reactions and equations representing the equilibrium relations used for drawing Fig. 1.101 Chemical reactions 4 Ag + 82 (g) = 2 Ag2S (acanthite)

Equilibrium relations (I)

Temperature range (°C)

log f s 2 = ( - 9 7 9 0 . 2 I / T ) + 4.83

41ogaAg

(1)

25-176

(2)

log fs2 = (--9173.95/T) + 3.61 + 2 log aAg2s - 4 log aAg

(2)

I76-804

4 A g + Se2 (g) = 2Ag2Se (naumannite) (3)

Iog fSe2 = (-- 10644.67/T) + 3.12 + 21ogaAg2sc 4IogaAg

(3)

133-727

4PbS + Se2 (g) = 2Pb2Se

log iS2 = log fSe2 q- 755.68 -- 0.24 (4)

25-327

+ 21ogaAg2s 4 Ag + $2 (g) = 2 Ag2S (argentite)

(4)

- 2 Iog(apbse/apbs)

FeS + 1/2S2 (g) = FeS2 (pyrite)

(5)

log fs2 = -- 15460/T + i4.32 - 2 log XFcS (5)

ca. 400-700

T: absolute temperature; fs~: sulfur fugacity; fSe2: selenium fugacity; aAg2S: activity of Ag2S in argentite; aAg2Se: activity of Ag2Se in argentite; aAg: activity of Ag in electrum; apbse: activity of PbSe in galena; apbs: activity of PbS in galena; XFeS: FeS mole fraction of FeS in sphalerite (Shikazono and Takeuchi, 1984).

139

Miocene-Pliocene Hydrothermal Ore Deposits -10 o=

-10 8~

-15

-15

~ 2 1

~:~

o' o

~-"'~'-

~,--- 22

-20 H.T

-25 . . . . . . . 150

!. . . 200 . . . . . . . . . . 250 300 Temperature (°C)

-25

. . . . . . . 150 . . . . . . . . . . 200 ....

250

300

Temperature (°C)

Figure 1.101. Selenium fugacity-temperature diagram, i3: Argentite (or acanthite)-electmm-galena-Se2(g) equilibrium curve for XAg2S (X: mole fraction) = 0.8, Xpbse = 0.05 and XAg = 0.4. 14: Argentite (or acanthite)-electrum-S2(g) equilibrium curve for XAg2Se = 0.2 and XAg = 0.4. 15: Argentite (or acanthite)--electrum-galena-Se2(g) equilibrium curve for XAg2S= 0.8, XpbSe= 0.05, and XAg = 0.8. 16: Argentite (or acanthite)-electrum-Se2(g) equilibrium curve for XAg2Se = 0.2 and XAg = 0.6. 17: Sphalerite-pyrite-galena-Se2(g) equilibrium curve for Xr,bSe= 0.05 and X F e S = 0.0l. 18: Sphalerite~yrite-galena-Se2(g) equilibriumcurve for Xpbse = 0.05 and XFeS= 0.02. 19:Sphalerite-pyrite-galena-Se2(g) equiiibriumcurve for Xpbse= 0.02 and XFeS= 0.004. 20:Sphalerite-pyrite-galena-Se2(g) equilibriumcurve for Xpbse = 0.02 and XFeS= 0.007. 21: Argentite (or acanthite)-electrum-Se2(g) equilibrium curve for XAg = 0.5 and XAg2Se = 0.5. 22: Argentite (or acanthite)-galena-electrum-Se2(g) equiIibrium curve for Xpbse = 0.02, XAg2S~---0.5 and Xag = 0.5. 23: Argentite (or acanthite)-electrum-Se2(g) equilibrium curve for XAg2Se = 0.5 and XAg = 0.7. 24: Argentite (or acanthite)-galena--electrumSe2 (g) equilibrium curve for XAg2S= 0.5, XpbSe= 0.02 and XAg = 0.7. Estimated ranges of temperature and fSe2 for Kushikino and Takatama are shown in this figure as shaded areas. H.T.: homogenizationtemperature of fluid inclusions.

content of electrum and temperature obtained by White et al. (1957), the relationships between temperature, fs2, fse2, Ag content of electrum, FeS content of sphaterite, selenium contents of argentite (or acanthite), and of galena can be derived (Fig. 1.101). From these relations and analytical data on coexisting galena, sphalerite, argentite and electrum, the formation temperature, fs2, and fse2 can be estimated (Shikazono and Takeuchi, 1984) (Fig. 1.101). To derive the relations shown in Fig. 1.101, unity of activity coefficients of FeS2 in pyrite, Ag2S, and Ag2Se in argentite (or acanthite) was assumed. Activity coefficient of FeS2 in pyrite should be very close to unity, because the concentrations of minor elements (e.g., Ni, Co) in pyrite are very low, less than n x 10 . 2 wt%. Bethke and Barton (1971) have suggested in their experimental study on the distribution of selenium between coexisting galena and sphalerite that the P b S - P b S e system behaves as ideal solid solution at least above 600°C. Ag2S in argentite (or acanthite), FeS in sphalerite and Ag in electrum are 5 mol%, ca. 20 mol%, 1.2-2.4 mol%, and 4 6 - 6 0 atomic%, respectively, for the Kushikino deposits and 2 mol%, ca. 50 mol%, 0.4-0.7 mol%, and 5 0 - 7 0 atomic%, respectively for the Takadama deposits. Detailed descriptions on these deposits can be referred to in Kitami (1973), Sukeshita and Uemura (1976), Yamaoka and Nedachi (1978b), Izawa et al. (1981) and Takeuchi and Shikazono (1984). Based on these analytical data on the minerals mentioned above and thermochemical consideration formation temperature, fs2, and fse2 for these deposits are

140

Chapter 1

estimated. Temperature and fs2 estimated on the basis of fs2-temperature diagram is ca. 220-300°C and 10-9-10 -13 atm for the Kushikino and ca. 150-250°C and 10-11-10 -18 atm for the Takadama. Temperature and fse2 estimated on the basis of fSez-temperature diagram (Fig. 1.101) is ca. 200-300°C and 10-12-10 -18 atm for the Kushikino and ca. 150-250°C and 10-14-10 -23 atm for the Takadama. Homogenization temperatures of fluid inclusions in quartz for the electrumsphalerite-pyrite-argentite-galena stage of the Kushikino and Takadama deposits are in the range of ca. 180-250°C (Takeuchi, 1979; Izawa et al., 1981) and ca. 160-240°C (Yamaoka and Nedachi, 1978b; Watanabe, 1979), respectively. Wright et al. (1965) have found from hydrothermal experiments that continuous solid solution exists in the PbS-PbSe system at 300°C. Coleman (1959) has described specimens covering the entire range of solid solution between galena and clausthalite from vanadium-uranium deposits of the Colorado Plateau type. Thus, it seems likely that the departure from ideal solid solution in the PbS-PbSe system is not large in the temperature range considered here (ca. 180-300°C). However, unfortunately, absolute values of VPbS and YPbSe (Y: activity coefficient) in the temperature range considered here cannot be estimated. It is known that continuous Ag2S-Ag2Se solid solution exists in the temperature range concerned (ca. 180-300°C), although this solid solution is unquenchable. Deviation from ideality for AgzS-Ag2Se solid solution is also not studied. Analytical data on galena, argentite (or acanthite), electrum and sphalerite which all coexist with pyrite in the Se-rich Au-Ag vein-type deposits are available from Kushikino and Takadama Au-Ag vein-type deposits (Kawai, 1976; Takeuchi, 1979). PbSe in galena, Ag2Se in argentite (Takeuchi, 1979) and thermochemical calculations are used to estimate temperature and fs2 and Jse2- The temperature estimated from this mineral assemblage for the Kushikino seems to be slightly higher than the homogenization temperature of fluid inclusions. There are several possible reasons for this discrepancy. They are: (1) uncertainties of free energy changes for the reactions in Table 1.14, (2) changes of chemical compositions and phases during the post-depositions stage, (3) not simutataneous precipitations of electrum, sphalerite pyrite, galena, argentite (or acanthite) and quartz which are studied for the chemical composition and fluid inclusions, and (4) deviation from ideality of Ag2S-Ag2Se and PbS-PbSe solid solutions. It is difficult to determine which is the main cause for this discrepancy. In order to solve this, more detailed studies on this assemblage and thermochemical properties of PbS-PbSe and Ag2S-AgzSe solid solutions are required. Although such a discrepancy exists, it was for the Se-rich A u - A g vein-type deposits (Kushikino and Takadama) that (l) the mineral assemblage of sphaleriteelectrum-argentite-galena-pyrite assemblage is a useful indicator of environmental condition (fse~, fs2 and temperature), (2) precipitation temperature for this mineral assemblage is in the range of 150-300°C, and (3) fse2 is lower than fs2. Based on the .[Se2, fs2 and temperature estimated from this mineral assemblage, we can place a limit on the ranges of the other important chemical parameters such as total dissolved selenium and sulfur contents in ore forming solution responsible for the Se-rich A u - A g vein-type deposits.

141

Miocene-Pliocene Hydrothermal Ore Deposits

1.4.3.4. Chemical composition of ore fluids The solubilities of ore metals depend on several variables such as fs2, fo2, pH, salinity, NS and temperature (Barnes and Czamanske, 1967; Shikazono, 1972b; Barnes, 1979). In the earlier sections, we estimated the ranges of these variables. Therefore, it is possible to calculate chemical compositions of ore constituent elements in ore fluids. For instance, Au concentration of ore fluids responsible for epithermal A u - A g deposits could be estimated from the following reaction. Au + HS + H2S + l / 4 02 ----Au(HS) 2 + 1/2 H20

(1-39)

Seward (1973) experimentally determined the solubility of Au due to this complex and equilibrium constant for the above reaction. Figure 1.102 shows the solubility of Au on l o g f o z - p H diagram calculated based on the thermochemical data by Seward (1973). The solubility of gold is high in neutral and near oxidized sulfur species/reduced sulfur species boundary (Fig. 1.102). It is noteworthy that this region corresponds to the f o z - p H region of epithermal A u - A g vein-type depositions (Figs. 1.90 and 1.91) (Shikazono, 1974a; Hattori, 1975). The solubility of pure gold is shown in Fig. 1.102. However, gold does not occur as pure gold in A u - A g deposits, but as electrum. Therefore, the effect of Ag content of electrum on gold solubility must be considered. Shikazono and Shimizu (1987) estimated NAu/NAg ratio in ore fluids responsible for Au-Ag veins based on the following reaction. (Au)e] + AgC12 + 2H2S = (Ag)el + Au(HS) 2 + 2C1- + 2H +

(1-40)

where (AU)el and (Ag)el are the Au and Ag components of electrum, respectively. Au

I HS021 S0z~

30

HEMATITE

250 °C m:~S = 3 x 10-3

I

40 PYRRHOTITE MAGNETITE

45 2

I

4

I

6 pH

H2SI HS- \ i I

8

Figure 1.102. fo2-PH diagram at 250°C showing the stability fields of the principal sulfur species and solubiiity contours for gold in mg/kg as Au(HS)~-(HenIey,1984).

142

Chapter 1 From the equilibrium relation for equation (1-40), we obtain,

aAg/aAu =

2 2 2 (au2smagcl~K1-40)/ (mau(us)~aci_aH+)

(1-41)

where K1-40 is equilibrium constant for the reaction (1-40), and ?/AgCl~/)/Au(rtS)~ is assumed to be unity. The equation is controlled by temperature, acl , aH2S, pH and mAgc12/mAu(HS)~ratio. We can calculate the E A u / E A g ratio in ore fluids responsible for Au-Ag veins, by giving temperature, salinity, pH and H2S concentration. The temperature is estimated based on fluid inclusion studies and electrum-sphalerite-argentite-pyrite assemblage. The NaC1 equivalent concentration of ore fluids is approximated from freezing temperature data on inclusion fluids, though the final melting temperature of fluid inclusion ice is also affected by CO2 concentration in epithermal ore fluids (Hedenquist and Henley, 1985). The pH values are estimated assuming the equilibrium among K-feldspar, K-mica, and quartz; this in turn allows a calculation of the potassium ion activity. The activity of HzS is estimated based on the equation showing the relation between the partial pressure of H2S gas and temperature of active geothermal waters (Giggenbach, 1980; Arn6rsson, 1985). Using the typical value of these variables and XAg (mole fraction of Ag in electrum) = 0.5, which is a typical value for electrum from epithermal Au-Ag vein-type deposits, E A u / E A g is calculated to be about 0.1, which is similar to that of Broadlands geothermal water (New Zealand) which is associated with gold precipitation. The above calculation is based on the assumption that AgC12 is the predominant Ag species in ore fluids. However, it is possible that silver bisulfide complex could contribute significantly to the transportation of Ag (Henley, 1985; Brown, 1986). If gold and silver bisulfide complexes are the dominant gold and silver aqueous species in Broadlands geothermal water and ore fluids responsible for Japanese epithermal Au-Ag vein-type deposits, the Ag/Au ratio of ore fluids responsible for Japanese Au-Ag epithermal vein-type deposits may be similar to that of Broadlands geothermal water which is about 0.1. tf mci- and pH are assumed to be 1-5 molal and lower than that for K-feldspar-Kmica-quartz equilibrium, respectively, E A u / E A g is estimated to be considerably lower than 0.1. Therefore, E A u / E A g of ore fluids for epithermal base-metal vein-type deposits is thought to be considerably lower than 0.1. The concentrations of base-metals (Cu, Fe, Pb, and Zn) in hydrothermal solution in equilibrium with sulfides (chalcopyrite, pyrite, galena and sphalerite) depend on several variables such as pH, mcl- concentration, temperature, mN2S, and fo2. The relation between the concentrations and these variables can be derived based on the chemical equilibrium for the following reactions. CuFeS2 + 2 C1- ÷ 2 H + + 1/2 02

~---FeS2 +

CuC12 ÷ H20

(1-42)

FeS2 + 2 C1- + 2 H + + H20 = FeCI2 + 2 H2S + 1/2 02

(1-43)

ZnS + 2C1- + 2H + = ZnCI2 + H2S

(1-44)

The ranges of these variables of epithermal ore fluids were estimated in section 1.4.3. Although the ranges of these variables are wide, we could estimate the concentrations of base metal elements if we took typical ranges of these variables.

143

Miocene-Pliocene Hydrothermal Ore Deposits

1.4.3.5. Stable isotopes 6D and 3180. 3D and 3180 of the ore fluids responsible for epithermal A u - A g and basemetal vein-type deposits in Japan have been estimated from analyses of fluid inclusions (Hattori and Sakai, 1979) and minerals (Watanabe et al., 1976). These data are shown in Fig. 1.103. 3D values of ore fluids for epithermal A u - A g vein-type deposits are similar to those of present-day meteoric water values. 3D values of epithermal ore fluids for base-metal vein-type deposits are slightly higher than those of epithermal A u - A g vein-type deposits. This may be due to the boiling of epithermal base-metal ore fluids and involvement of seawater. 3180 values of ore fluids are higher than those of meteoric water values. This is considered to be due to oxygen shift, which was caused by meteoric water-rock interaction. 3180 of ore fluids responsible for epithermal base-metal ore fluids is higher than that for epithermal A u - A g ore fluids. This is due to high extent of water-rock interaction and probably involvement of seawater and igneous water in the ore fluids. In individual deposit, 3180 of minerals varies widely. For example, 3180 of quartz and adularia increases with the stage of mineralization, and correlates to adularia/quartz ratio (A/Q), and Au and Ag grades of ore (Fig. 1.104). Such increase in 8180 is found also in carbonates in the Seigoshi A u - A g deposits and quartz and adularia in the Hishikari A u - A g deposits (Shikazono, 1988a; Shikazono and Nagayama, 1993). This increase is interpreted in terms of boiling of fluids at late-stage (Shikazono, 1989), decrease in water/rock ratio (Shikazono and Nagayama, 1993) and/or an influence of sedimentary rocks (Shikazono, 1999b).

~

. -20-40t~

J

-60-

SG

S.W.

YUG

-80-100

I

-10

1

-8

I

-6

I

-4

I

-2

I

0

1

2

4

Figure 1.103.3D and 3180 of ore fluids responsibIe for epithermai Au-Ag vein-type deposits in Japan (Hattori and Sakai, I979: Imai et aI., 1998). S.W.: seawater, M.W. line: meteoric water line, KK: Kushikino, SG: Seigoshi, YUG: Yugashima, TK: Takatama, FUK: Fuke, YN: Yatani, KN: Kanisawa (Yatani), HK: Hishikari.

144

Chapter 1 ©

=

=

0

>

-1-

< 1" 0"

-2

-3. 0.5"

i "

B

-4

-5

L

O, -2.

-3.

• Quartz-Adularla ratio • 6180 Fluid

© ,.d

Pb

•

-4 3-1T I

1 3 6-IT

i 6 [

, 9 6-2T

i

/ 12 6-3T

~Au

I

i I5 64T

J I thickness 18 21 I 5-1T I 5-2T I (cm)

Figure 1.104. The relationship among 8180 (in permil) of fluids, minor element contents, and A/Q (adularia/quartz ratio) in the vein from wail rock side (3-IT) to central part (5-2T) (Ryosen No. 5 vein, Hishikari mine, 85 ml E50) (Shikazono and Nagayama, 1993).

8D and 8180 values for epithermal deposits from other countries are summarized in Fig. 1.105 (Field and Fifarek, 1985). The oxygen shift away from the meteoric water line is always observed, but 8D is similar to meteoric water value, suggesting meteoric water source of epithermal ore fluids. Magmatic contribution to ore fluids has not been found except in some ore fluids responsible for the deposits in the other countries; Tui

145

Miocene-Pliocene Hydrothermal Ore Deposits

8180 %0 -20 0

,

-10

0

i " /; Meteoric Water LinV

/

-20 -40 -60 -80 -100

Walrakel/Broadlandsl'.~ LarderelloQ~". . . . . Gov~,ar~/aAW C

-160

+20

Ocean Water iSMOW;

/

//

Magmatic water [ 1 4 &GS I•CP

I II

/ -

// / ,/.~

~ao/%~

/

_~/ oO~/e¢~

/,o ,G ,,~ . ~~/ : ,~q# =~_~--oSteamboat Springs /~-~ / / R&•A I " " CO& / / - " /

;Y~owstone i

CA~/'/ I

f

I

/

/_® t____J ~,~ # /..~-

.-e

Mt. Lassen(~-----.-~ . . . . . . . II/~T .Uu ATE =11

I

=

&TF

Salton SeaZ.~- . . . . . . . .

-120 -140

+10

6"g"

Geothermal H2' • Surface * Subsurface Hydrothermal H2' • Epithermal

// /

I

I

I

I

Figure 1. i05. Distributions of 3D and 3]80 in various waters, minerals, and hydrothermal fluids of epitherma] deposits (FieId and Fifarek, 1985).

(New Zealand) (Robinson, 1974); Finlandia (Kamilli and Ohmoto, 1977), and possibly Comstock Lode (Taylor, 1973). ~13C and 3180 of carbonates. 313C and 3]80 of carbonates have been obtained by Osaki (1973), Matsuhisa et al. (1985), Shikazono (1988a, 1989), and Morishita (1993) (Fig. 1.106). Roughly, the data lie between the igneous value (-7%0 to -8%o), and marine carbonate value (-1%o to +4%0). These are similar to the variations in carbonates from the Kuroko mine area. However, 313C and 3180 of carbonates from the epithermal vein-type deposits vary more widely than the carbonates from the Kuroko mine area. Some data plot in the region below the igneous carbonate-marine carbonate mixing line. This suggests that meteoric water was involved in the ore fluids. On the other hand, some data deviate from igneous carbonate-marine carbonate mixing line to higher 3180 region. This higher ~180 values are explained in terms of boiling of ore fluids. High 3180 values are obtained for carbonates of late-stage of mineralization (Shikazono, 1988a). Figure 1.107 shows the frequency of 313C of carbonates from epithermal A u Ag vein-type deposits and that from base-metal vein-type deposits. The carbonates are divided into two types: type A and type B. Type A is characterized by: (1) abundant occurrence in each deposit; (2) coexistence with sulfide minerals; and (3) large grain size. Main carbonate minerals are rhodochrosite and Mn calcite, whereas calcite is the main carbonate mineral for type B. Mn-carbonates of type A occur in P b - Z n - M n vein-type deposits. Type B is characterized by; (1) poor amounts in each deposit; (2) coexistence

146

Chapter 1

25 o

o

2O

oe oo

o

og

o -to

o

o

o ee

o

g15

o oo

•

o

.

o

•

o g

o

E O .IQ

e~

•

~dl0

oS t ~ . ~ , , o ~ . " , % , " Oo°o o

-,..;:

o•

o

~5

°o•

°o.**

Oo , *.~

t..~

o

o

o .

• o

ee

|

°o

o

o

oe

~o

-5 $13C of C a r b o n a t e s ( ' / . , ) Figure 1.]06. 5180-513C of carbonates from Neogene vein-type deposits in Japan (open circle = calcite; solid circle = rhodochrosite and Mn-ca]cite; solid triangle = dolomite; cross = siderite) (Shikazono, ]989).

with late-stage quartz occurring in vugs; and (3) small grain size. Type B carbonate occurs in Au-Ag vein-type deposits and in Cu-Pb-Zn vein-type deposits but not in Pb-Zn-Mn vein-type deposits. The range of 313C of Japanese epithermal veins is similar to that of the other countries which is in a range of -10%o to 0%0 (Fig. 1.108 (Field and Fifarek, 1985). 313C and 3180 of carbonates from southern Kyushu (Hokusatsu gold district) have been studied in detail (Matsuhisa et al., 1985; Morishita, 1993). Morishita (1993) found that the 313C values of hydrothermal solution in the district during the mineralization stages were low (-11%o), compared with that of average crustal carbon (-7%o), suggesting that 313C of hydrothermal solution is controlled by organic carbon in widely distributed sedimentay rocks of the Cretaceous Shimanto Supergroup basement. 313C and 3]80 of carbonates in the Seigoshi mine district, Izu Peninsula, middle part of Honshu were determined by Shikazono (1988a). He showed that early-stage carbonates have 313C and 3180 values of -2.9%0 to +0.6%0 and +1.7%o to +10.2%o, respectively, suggesting a contribution of marine carbonate in Miocene marine sediments (Green tuff), but late-stage carbonates have high 3180 and low 313C, suggesting the effect

147

Miocene-Pliocene Hydrothermal Ore Deposits 25 20

z

5

1 -10

25

-8

-6

-4

-2

a'~c (%o)

(b)

> , 20 E

<

0

15

-Q 10

E

z

5

8130

(%0)

Figure 1.107. Frequency histogram for ~ 13C of type-A(a) and type B (b) carbonates (Shikazono, 1989).

of boiling. Osaki (1973) suggested that the higher 3~3C (-2.5%0 to -7.5%0) and ~]80 values (+7.0%o to + 15.0%o) of rhodochrosites from the Jokoku Mn vein in Southern Hokkaido are due to contamination with limestone. In addition to boiling and origin of carbon, temperature and chemical states of dissolved carbon influence 313C of carbonates (Matsuhisa et al., 1985). These detailed studies on individual mine district suggest that carbon in carbonates was derived from the country rocks underlying the ore deposits and oxygen in ore fluids is controlled by origin of ore fluids (mostly meteoric water) and boiling of ore fluids.

~34S of sulfides. A large number of ~34Sdata on sulfides from epithermal base-metal and A u - A g vein-type deposits are available (e.g., Shikazono, 1987b). The 334S data are summarized in Fig. 1.109. The ~34S values of A u - A g deposits range from -7.5%0 to +5%o. However, majority of the ~34S values fall in a narrow range from -1%o to +3%e. ~34S

148

Chapter 1 Geothermal Systems Geysers:

002 HCO 3 whole rock calcite

a'~c (%0) -20

I

-10

0

I

10

I

Salton Sea/Cerro Prieto: Carbonate host altered carb. host carbonates

Broadlands/Wairakei: calcite

Sediment-Hosted Cortez: carbonate host altered carb. host calcite Carlin: carbonate host altered carb. host calcite

g !

Volcanic-Hosted

Pueblo Viejo: carbonaceous sed.

Zoned Polymetallic Veins

Creede: Sunnyside: Tui: Casapalca:

carbonates carbonates carbonates calcite

Figure 1.108. Distributions of ~13C in epithermaI deposits (Field and Fifarek, 1985).

values of sulfide sulfur from Green tuff-type epithermal Au-Ag deposits are higher than those from Non-Green tuff-type (Shikazono, 1999b) (Fig. 1.110). The other geochemical and mineralogical features (313C and 3180 of carbonates, salinity of inclusion fluids, Ag/Au total production ratio, association of metals, gangue minerals) are also different in Green tuff-type and Non-Green tuff-type of epithermal Au-Ag deposits (Table !.15). These differences can be explained by the influences of country rocks. The Green tuff-type deposits are affected by marine rocks containing marine sulfates and carbonates and interstitial water of seawater origin with high chloride concentration, resulting in higher ~34S, ~13C, salinity, and Ag/Au total production ratio and enrichment of base metals (Cu, Pb, Zn, Mn, Fe), whereas the Non-Green tuff-type deposits are by subaerial rocks. Sedimentary sulfur and organic carbon were involved in ore fluids responsible for the Non-Green tuff-type deposits, resulting in relatively low ~348 of sulfides and 3 I3C of carbonates. 334S values of epithermal base metal deposits are higher than those of the epithermal Au-Ag deposits and range mostly from +3%o to +7%o (Fig. 1.111 ). Although most of 334S values for base-metal deposits lie in this range, ~34S of composite sample of sulfides from the Motokura Cu-Pb-Zn deposits, Ohmori Cu-Ag deposits, Hosokura Pb-Zn deposits, Sasayama Cu-Pb-Zn deposits and Imai-Ishizaki Cu-Pb-Zn deposits are low, that is, +0.1, +1.8, +2.2, -0.9 and -2.1%o, respectively (Shikazono, 1987b; Shikazono and Shimizu, 1993).

149

Miocene-Pliocene Hydrothermal Ore Deposits [ ] Ginguro Type

if)

Base metal rich

[] Type Disseminated [] Type

Ct~

t~ e-

~[~

<

"6 E z 1

[~

-7

N]

-6

-5

IoIX~lxlX'gqN]X~ I I ~ I,[ 1,1 I,I 1,

.5~.~?~NNN

-4

-3

-2

-1

0

1

2

3

4

5

6

-l

8~s(%o)

Figure 1.109. Sulfur isotopic compositions of Neogene Au-Ag vein-type and disseminated-type deposits. Sulfur isotopic compositions on the samples from the Yatani deposits (Sample No. YT26 from Zn-Pb vein ~34S = +3.3%0), and HS72050305-YT1, YT24 and NS-3 from Au-Ag vein (average ~34S = -t-3.3%o)) by Shikazono and Shimazaki (1985) are also plotted. "Base-metal rich" implies the sample containing abundant sulfide minerals but no Ag-Au minerals from base-metal rich deposits and also from Ginguro-type deposits (Shikazono, 1987b). TABLE 1.15 Some characteristic features (~34S, ~I3c, Ag/Au total production ratio, metals produced, gangue mineraIs) of epithermal Au-Ag vein-type deposits of the Green tuff-type and the Non-Green tuff-type in Japan (after Shikazono, 1996)

334S 3~3C Au produced (metric tons) Ag produced (metric tons) Ag/Au (in wt. ratio) Amount of sulfides (Cu, Pb, Zn)

Green tuff-type

Non-Green tuff-type

high (-1%o to +6%o) high (-7%o to 0%o) 136 2586 high (average 19.1) large

low (-7%o to +2%0) low (-12%o to +2%0) 420 1506 low (average 10.7) small

common common common present

abundant rare rare absent

Gangue minerals

Calcite Rhodochrosite Manganoan calcite Barite

These data indicate that (1) ~348 values of sulfides are different in different mine districts; in the region where thick Green tuff occurs (e.g., Hosokura Pb-Zn deposit, ore deposits in Southwest Hokkaido) some values are relatively high, and ~34S values of the ore deposits in the Non-Green tuff region (e.g., Northeast Hokkaido) (Motokura, Khonomai, etc.) are low, (2) ~34S values of sulfides from small base-metal vein-type deposits (e.g., Imai-Ishizaki, Sasayama) are low, but ~34S values of relatively large base-metal deposits (Taishu, Toyoha, Ani, Osarizawa, Ohe, Jokoku) are high. These data suggest that ~34S values of large deposits are affected by reservoir sulfur in deep part, but sulfide sulfur of small ore deposits is influenced by the surrounding rocks having low ~34S values (basement sedimentary rocks).

150

Chapter 1 15

Non-Green tuff-type

>,1o.

0 c-

OLL

5"

0 q 15" Green tuff-type

o>,10c-

O"

E

LI.. 5-

-15

-5

-10

0

5

53'S (permil)

10

Figure 1.110. Frequency of ~348 values of sulfides from the Green tuff-type and the Non-Green tuff-type deposits (Shikazono, 1999b).

Cu.Pb.Zn Vein-Type Deposits

¢--

<

"6 $ E Z i

-8

,

-7

a

i

-6

-5

I

-4

t

i

-3 -2

-1

0

2

3

4

5

6

7 8 8~s (%o)

Figure 1.111. Sulfur isotopic values tbr sulfides from base-metal vein-type deposits (Shikazono, 1987b).

There is another explanation for the variations in 334S values of sulfide sulfur. It was cited that oxidation state (Jb2) and pH of ore fluids are important factor controlling ~34S values of ore fluids (e.g., Kajiwara, 1971). According to the sulfur isotopic equilibrium model (Kajiwara, 1971; Ohmoto, 1972), ~34S of sulfides in predominance

151

Miocene-Pliocene Hydrothermal Ore Deposits

]

Pyrrhotite-bearing

-2-1

0 1 2 3 4 5 6 7

8910

~34s (%°)

Hematite-bearing

•

-2-1

j

•

•

•

.

,

,

,

,

0 1 2 3 4 5 6 7

F]

89'16

8~s (%0)

Figure 1.112. Sulfur isotopic compositionof pyrrhotite-bearing(solid) and hematite-bearing (open) samples from base-metal-richdeposits in Greentuffregion (Shikazono, 1987b).

region of oxidized sulfur species is lower than that in predominance region of reduced sulfur species. Figure 1.112 shows that ~34S of sulfides containing pyrrhotite which formed in the predominance region of reduced sulfur species is lower than that containing hematite which formed in predominance region of oxidized sulfur species; ~348 values for pyrrhotite-bearing samples are less than +4.5%o, while those for hematite-bearing samples are more than +4.5 %o (Fig. 1.112). This is not consistent with equilibrium model (Kajiwara, 1971; Ohmoto, 1972). The ~34Svariation in individual deposits such as Yatani (Shikazono and Shimazaki, 1985), Hosokura, Toyoha (Kiyosu, 1977a; Hamada and Imai, 2000), Ohe (Kojima and Sugaki, 1989) and Taishu (Kiyosu, 1977b) have been studied. These data indicate that ~348 of sulfides do not vary widely in individual deposits. Figure 1.113 shows the distributions of ~34S in epithermal deposits in other countries summarized by Field and Fifarek (1985). ~34S of sulfides from volcanic-hosted epithermal deposits and zoned polymetallic veins are mostly in a range of -5%o to +5%o, typically 0%o (Creede, Sunnyside and Cotqui). This range is similar to that of Japanese epithermal Au-Ag deposits, but lower than that of base-metal deposits. Different interpretations of origin of sulfide sulfur with 0%o have been proposed. For example, Casadevall and Ohmoto (1977) suggest that the sulfur of sulfides in Sunnyside deposits (U.S.A.) originated from evaporite-bearing sedimentary rocks based on the assumption of equilibrium fractionation between H2S and SO ] . in the dominant region of oxidized sulfur species. Kamilli and Ohmoto (1977) also prefer the sedimentary sulfate source at Colqui (Peru), but suggest that an igneous origin is possible. For the Creede deposits, sulfate-sulfide sulfur isotopic equilibrium is not attained and igneous origin seems more likely. Distributions of major epithermal Au-Ag vein-type deposits are shown in Fig. 1.114. Green-tuff-type deposits are defined as the deposits occurring in the Green tuff region and Non-Green tuff type as those occurring in the regions other than the Green tuff region.

152

Chapter I (~34s

Geothermal Systems Iceland:

magmaticsulfur SO4 sulfates H2S pyrite

-10 I

0

(%0) 10

I

-20 I ! m

Yellowstone: SO4

sulfur H2S Salton Sea: sulfides BroadIands: sulfides Wairakel: SO4 sulfates H2S sulfides

B

B m

I m

Sediment-Hosted

Cortez: Carlin:

barite diagenetic py sulfides barite diagenetic py sulfides

32

Volcanic-Hosted Tolfa:

sulfates sulfides Pueblo Viejo: sulfates sulfur sulfides Goldfield: alunite pyrite Z o n e d Polymetallic Veins

San Juan Moutains: Creede: barite sulfides Sunnyside: sulfates sulfides Rico: sulfides Ouray: sulfides Tui: barite sulfides Guanajuato: sulfides in country rock sulfides in volc. sulfides in ore Casapalca: sulfides Finlandia: barite sulfides Western Cascades: sulfides Golden Sunlight: barite sulfides

m i

i

n B

45

_.B

m

. . . . . . . . . .

Figure 1.113. Distributions of ~34S in epithermal deposits (Field and Fifarek, 1985).

Figure 1.114 demonstrates that the deposits in northeastern Hokkaido, central Honshu, Sado Island and Kyushu deposits are Non-Green tuff type and those in southwestern Hokkaido, Northeast Japan (Tohoku), the Izu Peninsula, and San-in are Green tuff-type.

153

Miocene-Pliocene Hydrothermal Ore Deposits 133 °

o,

138 °

200 ~

,

Hokka~

~

40 ° ~

35°

%

SW Honshu

~

]

,¢ hu '

~

--

/ ~lzu peninsula

e tuff region t I (submarine volcanic region) ~ Non Green tuffregion I.. • 4 (subaerial volcanic region) Figure 1.114. Distributions of major Green tuff-type (solid circle) and the Non-Green tuff-type (open circle) Au-Ag vein-typedeposits in Japan (Shikazono, 1999b). ~ , , f " / ) MY Kyushu~ /

The Green tuff-type deposits are sometimes hosted not only by the Green tuff formation but also by the rocks overlying the Green tuff formation, implying they are younger than the Green tuff age (ca. 61-15 Ma). Examples of these deposits (Seigoshi, Yugashima, and Toi) occur in the Izu Peninsula, central Honshu. Base metal-rich deposits including epithermal vein-type and Kuroko deposits occur in the Green tuff, while base metal-rich deposits are few in the Non-Green tuff region. Sedimentary rocks often occur as host rocks, footwall rocks and basement rocks in the Non-Green tuff mine area. For example, in southern Kyushu, the Shimanto Supergroup shale is dominant as basement and a host rock for epithermal A u - A g vein-type deposits (e.g., Hishikari). ~34S values of sulfide sulfur from epithermal A u - A g vein type deposits obtained are summarized in Table 1.16 and Fig. 1.109. ~34S value of sedimentary rocks of the Shimanto Supergroup which hosts the Hishikari deposits, southern Kyushu is -12%o (Ishihara et al., 1986) which is considerably lower than those of the Hishikari deposits (+1%o to +2%0) (Shikazono, unpublished. This suggests that sedimentary sulfide sulfur is one of the sources of the sulfides and probably igneous sulfide sulfur is the dominant source. Morishita (1993) showed based on carbon isotopic composition of carbonates that carbon of carbonates in the gold-bearing quartzvein in southern Kyushu was derived from the Shimanto Supergroup shale. Imai et al. (1998) considered that hydrogen in the ore fluids was derived from the Shimanto Supergroup shale based on 3D (--60%0 to -100%o) of inclusion fluids in quartz and adularia of the Hishikari veins. These isotopic

Chapter 1

154 TABLE 1.16

~348 values of epithermal Au-Ag deposits in Japan (Shikazono, 1999b) Green-tuff-type (average +2.9%~, n = 40)

Non-Green tuff-type (average -1.4%~, n = 49) Mine

~348 (%0)

Metals produced

Kohnomai Ryuo Koudou SW H o k k a i d o Katsuyama

-3.5 -7.1 -0.7 -0.9 -7.0 -2.3 -7.2 -5.8

Au, Ag Au, Ag Au, Ag

- 1.3

Au, Ag Au, Ag

-2.8

Au

Sado

Kasuga Hishikari

Iwato

Sappro Eniwa Todoroki Chitose Teine Date Mutsu Innai

-0.2 -0.2 -14.6 -7.2 -4.2

Au, Au, Au, Au, Au,

Ag Ag Ag Ag, Cu Ag, As

+5.1 +4.5 +3.2 +3.2 +3.0 +2.5 +1.8 +4.3 +5.9

Au, Ag

+4.4 +5.2 +4.4 +2.1 +2.9 +3.5 +4.5 -0.4

Au, Ag

Handa Yatani

Takatama

-1.5 -0.9 +1.6 +0.7

Au, Ag

Central Honshu

Au, Ag, (Cu)

Izu Peninsula

-3.6 +1.0 -7.4 +0.2 +3.4 +5.6 +5.0 -0.5 +1.6 -2.0 --2.8, -1.5 -0.8, +5.5, +3.2, -0.2, -0. l, -2.3 +1.8, +0.4, +0.6, +1.5 +0.3, +0.2, 0.0, -1.1 -0.5, -0.8, -3.7, -5.4

Au Au Au Au Au Au Au Au Au Au

Ashiyasu Ohito

Kyushu

Bajo Asahi Hoshino Taio Ohkuchi Yamagano Yamada Arakawa Kushikino Akeshi

Metals produced

Au, Ag Au, Ag Au, Ag Au, Ag Au, Ag, Cu Au, Ag

NE Honshu

Sado Island

Takachi

Shizukari

Au, Ag

Central Honshu

Nebazawa Haguro Koei Masutomi Suzukura

~34S (%o)

NE Hokkaido

NE H o k k a i d o

Kami-oumu Hokuryu Sanru

Mine

Ag Ag Ag Ag Ag Ag Ag Ag Ag Ag

Au, Ag

Yugashima Nawaji Jyoren Omatsu Rendaiji Kawazu Okuyama Suzaki Mochikoshi Seigoshi Toi

Au, Ag Au, Ag

Au, Ag

+ 1.9

Au, Ag, Cu

+2.9 +2.0 +3.6 +4.9 +3.8 +2.4 +3.2 +5.8 +4.3 +3.2 +3.2 +3.1 +3.1 + 1.0 +0.4 +0.0

Au, Ag Au, Ag Au, Ag Au, Ag Au, Ag Au, Ag, Cu, Mn Au, Ag Cu Au, Ag, Cu Au, Ag Au, Ag Au, Ag Au, Ag

+0.1

-0. I Au, Ag

S W Honshu (San- in)

Takeno

+2.4

Au, Ag

Miocene-Pliocene Hydrothermal Ore Deposits

155

data support the view that a part of sulfide ore sulfur originated from sedimentary sulfur in the Shimanto Supergroup shale, although sedimentary sulfide is probably not a dominant source, and large amounts of igneous sulfide sulfur were involved. 3180 values of the Hishikari ore fluids estimated from 3180 values of quartz and adularia and homogenization temperature of fluid inclusions are -6%o to 0%0 (Shikazono and Nagayama, 1993) which are higher than those for the other Japanese epithermal A u Ag ore fluids (Fig. 1.103). These data are interpreted in terms of maturity of hydrothermal system (Shikazono and Nagayama, 1993), boiling of ore fluids (Nagayama, 1993b), a contribution of magmatic water (Matsuhisa and Aoki, 1994), or hydrothermal fluids whose oxygen isotopes exchanged for those of surrounding rocks at a low water/rock ratio (Shikazono and Nagayama, 1993), and a contribution of sedimentary heavy oxygen (Imai and Uto, 2001). Combined isotopic data (3180, 3D, 313C, 334S) indicate that the interaction of hydrothermal solution with the Shimanto Supergroup shale is likely cause for the higher 3180 of the Hishikari epithermal A u - A g ore fluids, compared with those of the ore fluids at main stage of A u - A g mineralization for the other epithermal ore deposits. Shikazono (1988a) has found that ~180 values of late-stage calcite in epithermal ore fluids are higher (0%0 to +10%o), compared with fluid inclusion data (Hattori and Sakai, 1979). Shikazono (1988a, 1989) considered that boiling of ore fluids is important for causing higher 3180 of late-stage calcite. However, it is also likely that this high value seems due to the interaction of hydrothermal solution with sedimentary rocks with high 3~80 values. 334S values of the Green tuff-type are higher than those of the Non-Green tuff-type and are close to ~34S of the Green tuff region (Kuroko and skarn-type deposits) fall in a narrow range of -t-2%o to +7%o and the sulfide sulfur in these ore deposits is called as the "Green tuff sulfur" by Shimazaki (1985). It is inferred by him that the "Green tuff sulfur" was derived from a deep and homogenous reservoir such as magma. However, two sources (igneous source and seawater sulfate) of the "Green tuff sulfur" are also likely because 334S values of base metal deposits (epithermal base metal vein-type and Kuroko deposits) lie between those of igneous sulfur (0%0) and seawater sulfate (+20%o). Kawahata and Shikazono (1988) calculated 334S of hydrogen sulfide (H2S) in ore fluids for midoceanic ridge deposits as a function of seawater/basalt ratio at constant temperature by giving 334S and sulfur content of original basalt. Following their calculation method, we could reasonably explain 3348 of base metal deposits in the Green tuff region (Shikazono, 1987b. However, 334S of the Green tuff-type A u - A g veins are slightly lower than those of the "Green tuff sulfur", indicating that sedimentary sulfur was involved, and degree of contribution of igneous sulfide sulfur was not large. Shikazono (1987b) pointed out a possibility of igneous origin for sulfide sulfur of epithermal A u - A g ore deposits based on average 3348 values of sulfides from these deposits ( - 1%o to +3%0). Smaller involvement of seawater sulfate into epithermal A u - A g ore fluids for the Green tuff-type than the ore fluids for base metal vein-type and Kuroko deposits in the Green tuff region is possible. The mixing of the "Green tuff sulfur" and sedimentary sulfur is likely cause for the difference in 334S of sulfide sulfur of two types of deposits, considering stable isotopic and geologic characteristics of the Green tuff-type and Non-Green tuff-type epithermal A u - A g veins.

156

Chapter 1

25 du N

20,

0

"~ 15. i •

~, 10-

0 a

D o

5-

-15

0 o

-10

Non-Grenn tuff type 0 Sado A Fuke Hishikari

~

•

.~

•i •

•

Im~ •

mOj mmn--~ m •

613C of Carbonates (%~)

b

Green tuff type • Nawagi • Kawaza-Kakehashi • Yugashima

• •

Yatani Seigoshi

Figure 1.115.3tSO-~t3Cof carbonates fromepithermal vein-typedeposits in Japan (Shikazono, 1999b).

313C values from both types of deposits are different (Fig. 1.115). 313C of the Non-Green tuff-type carbonates are relatively low, suggesting a contribution of organic carbon in sedimentary rocks in the mine areas. Although the ~13C values of the Green tuff-type vary widely, the data lie between igneous value (-7%0 to -8%0) and seawater value (-1%o to -4%o), suggesting a contribution of these two sources. Total metal (Au, Ag, Cu, Pb, Zn, Mn) production during the past, and Ag/Au total production ratio of major epithermal Au-Ag deposits (Green tuff-type and Non-Green tuff-type) are summarized in Table 1.17. Large amounts of Au and Ag have been produced from the Non-Green tufftype (e.g., Hishikari, Sado, Kohnomai, Kushikino). Ag/Au total production ratio of the Non-Green tuff-type (average 10.7) is lower than for the Green tuff-type (average 19.1). These differences can be explained by the HSAB principle by Pearson (1963, 1968). This principle indicates that HS- and H2S are likely to form complexes with the metals enriched in the Non-Green tuff-type (Au, Hg), whereas CI prefers to form complexes with the metals concentrated in the Green tuff-type (Ag, Pb, Mn, Fe, Cu). Occurrence of gangue minerals in both types of deposits is different. For example, Mn minerals (Mn carbonates, Mn silicates) occur abundantly in the Rendaiji, Yugashima, Yatani, and Todoroki epithermal Au-Ag vein-type deposits in the Green tuff region but not in the Non-Green tuff-type. 334S values of barite from these deposits are high (+18%~

Miocene-Pliocene Hydrothermal Ore Deposits

157

TABLE 1.17 Tonnages of gold, silver and the other associated metal, silver/gold ratio, K-Ar ages and host rocks for the Te-type and Se-typc epithermal gold deposits (Shikazono et ah, 1990) Deposit

Au (ton)

Ag (ton)

Ag/Au

Sanru

6.7

40

6.0

Koryu Chitose Yatani

0.76 22.8 1.7

22.2 105 64

29.4 4.7 39

Takadama Nebazawa Seigoshi Omidani Sakoshi-Odomari Kushikino Hishikari

28.8 1.0 13.5 0.3 1.1 55.2 21.7

279.9 65 455 79 9.7 456.2 14.3

Date Kobetsuzawa Teine Chitose Mutsu Osorezan Sado

<1 <1 10.4 22.8 <1

Other metals

Year of production (ton)

K-Ar ages

Host and country rocks

1925-1974

12.4 4- 0.6

1903-1957 I93(>-1974 1870-1974

1.0 4- 0.3 4.7 3.34.0.3

Rhyolitic tuff, tuff breccia, shale Shale Propylite Acidic tuff, shale

9.7 65 34 267 8.8 8.3 0.6

1429-1974 1942-1974 1935-1976 1914-1974 1977-1982 1965-1986 1983-1990

8.4 5.0-5.7 1-3.7 66-68

<1 <1 62.6 105 <1

0.7-2 48.1 6.0 4.7 0.1

1932-1975 1955 1932-1971 1936-1974 1940

5.2:t:0.4

82.9

2404.3

33

Cu: 5400

1601-1988

Kawazu

5.4

272.8

1-5

1915-1959

Okuyama Suzaki Chugu Takeno

<1 2.0 <1 4.6

<1

Cu: 1054 Mn: 15840 Cu: 1250

<1 90.0

Agawa Kato

<1 <1

<1 <3

lriki Yamada Fuke Okuchi Kushikino

<1 <1 2.4 21.2 55.2

<1 <7 1.5 I5.7 456.2

Se-type

Cu: 70 Cu: 1270 Pb: 29210 Zn: 58420

4.0 4- 0.3 0.97 4- 0.041.54.0.3

Tuff, shale Rhyolite Andesite, diorite porphyry Slate, sandstone Dacitic tuff Andes±re, propylite Shale, andesite

Te-type Cu: 7560.8 Cu: 70

0.1 20

10-I2 I054 0.7 0.7 8.3

Cu: 200 Pb: 500 Zn: 100 Cu: 1.3

1912-1963 1914-1941 1938 1868-1949

4.7

13.44-0.514.5±0.5 22.1 4-0.7 24.4 4- 0.8 1.44-0.31.5 4- 0.3

17.9 4- 4.518.24-4.6

1933-1945 1933-1954

1933-1937 1916-1955 1896-1947 1905-1974 1865-I 986

Tuff Propylite Propylite Propylite Liparitic tuff Dacite Dacite, rhyolite shale, dacitic tuff

Propylite, rhyolite Propylite, dacite, basalt Propylite, rhyolite Tuff Granite, tuff, andesite Rhyolite, dacitic tuff Porphyrite

0.453 -t-0.018 1.4 4- 0.2 1.i :t- 0.5 4.0 :k:0.3

Tuft, andesite, liparite Propylite Propylite Andesite, rhyolite Propylite, andesite

158

Chapter 1

to +29%o) (Watanabe and Sakai, 1983; Shikazono et al., t990), suggesting that sulfate sulfur of barite was derived from sulfate in the submarine volcanic and pyroclastic rocks (gypsum, anhydrite, sulfate ion in interstitial water) (Watanabe and Sakai, 1983). Carbonates are found both in the Green tuff-type and the Non-Green tuff-type, but large amounts of carbonates (dominantly calcite) occur in the Non-Green tuff-type such as the Kushikino and Kohnomai epithermal Au-Ag vein-type deposits. ~13C and 8180 of carbonates from epithermal vein-type carbonates are plotted in Fig. 1.106. 313C values from both types of deposits are different, gl3C values of the Non-Green tuff-type carbonates are relatively low, suggesting a contribution of organic carbon in sedimentary rocks in the mine areas. Although the ~13Cvalues of the Green tuff-type vary widely, the data lie between igneous value (-7%o to -8%0) and seawater value (-1%o to +4%o), suggesting a contribution of these two sources.

834S and 8180 of sulfates. Watanabe and Sakai (1983) have analyzed 3348and 3180 of barite, anhydrite and gypsum in the epithermal vein-type deposits. They found that (1) the hydrothermal vein sulfates are characterized by more scattered distributions of 334S values ranging mostly from -t-10%o to .1.20%o and 8180 mostly from 0%0 to +14%o than those of Kuroko sulfates (section 1.3.3) (Fig. 1.43). In Fig. 1.43, 334S and 8180 of sulfates from epithermal vein-type deposits (Watanabe and Sakai, 1983) are plotted. These data show that 334S (mostly from .1.24%o to -1-t-37.8%o)and 3180 of barite (0.1%o to ,1,1,18.7%o)from epithermal Au-Ag-Te vein-type deposits are higher than that of epithermal base-metal vein-type deposits (~34S; .1.16.0%o to +24.6%o, 3180; +2.1%o to +12.1%o). These data could be explained by: the sulfur of barite from epithermal Au-Ag-Te deposits came both from volcanic gas (SO2) and marine sulfate, but that of epithermal base-metal deposits came from marine sulfate and oxidation of H2S. 1.4.3.6. Lead isotopes Lead isotopic data on the epithermal deposits together with Kuroko deposits are plotted in Fig. 1.116 (Sato and Sasaki, 1973; Sato et al., 1973, 1981; Sato, 1975; Sasaki et al., 1982; Sasaki, 1987; Fehn et al., 1983). It is evident that lead isotopic compositions of epithermal vein ores are more scattered than Kuroko ores, although averaged values are similar to the Kuroko ores. This variation seems to be due to the difference in crustal materials underlying the ore deposits; Lead isotopic compositions of different ore deposits which formed at different ages in the same district show the same values (Sasaki, 1974). 1.4.3.7. Rare earth elements (REE) Kato et al. (1990) analyzed carbonates from epithermal Cu, Zn-Pb and Au-Ag vein-type deposits in Japan for rare earth elements (REE) and found that (1) calcite from Cu-type is characterized by high La/Yb, high REE concentration and negative Ce anomaly is small, and (2) calcite from Zn-Pb type is characterized by high La/Yb, low REE and negative Ce anomaly. These data suggest that ore fluids for Cu-type were generated under low water/rock ratio condition or were influenced by magmatic water, while meteoric water component was dominant for the Zn-Pb type.

159

Miocene-Pliocene Hydrothermal Ore Deposits 39.0

,

,

%

' !

i

: ............. l "CO

io

C, .04

uo

:d

t°

38.5

' o i: °"~

:

L .............

.t

SRM-981 ,. ,tie

[.r,

206/204

t*"

15.7

!

i

!

i

/k.

!

!

!

I

o TA

'O

0

15.6 • o e4 SRM-981

F~"7"';q :oL:z.?.°?...}

n W

H

.d"

15.5

206/204 15./,

I

16.9

,

I

17.0

~

t

18.3

I

18./,

,

I

l&5

I.

187

Figure 1.116. Lead isotopic variation in Japanese Neogene ores. The majority of data fall in a relatively narrow range which is no more than twice the experimental uncertainty indicated by the replicate analyses of NBS-SRM-981 standard (Sasaki et aI., 1982).

The REE characterics of calcite from the A u - A g type are variable. For example, calcites from Sado A u - A g vein, one of the largest A u - A g deposits in Japan have both signatures of meteric water and magmatic (or igneous) contributions. Positive Eu anomaly is only found in calcite containing low REE from A u - A g type (Seigoshi deposit) (Shikazono, unpublished). 1.4.4. Se- a n d T e - t y p e A u - A g d e p o s i t s As shown in Fig. 1.117, Se-type and Te-type epithermal A u - A g vein-type deposits are located in the Cretaceous-Quaternary volcanic terrane of Japan (e.g., northeast and southwest Hokkaido, middle Honshu, south Kyushu). Some Te-type deposits are located in regions similar to the Se-type deposits. Sometimes, Te mineralization is associated with the Se-type deposits, though Te minerals usually do not coexist with Se minerals. However, rarely, Te minerals coexist with Se minerals in the Te-type deposits (e.g., Teine, Suzaki, Kawazu, Iriki) on a polished section scale. For example, coexistence of native Te and Se-bearing tetrahedrite is found at Teine. Generally, Te mineralization occurs at

160

Chapter 1

Teine

I oh,os ( 200km

-Kobe~uzawa Date

Osorezan

Mutsu ."

4o~u

Sado

a :no

Ornidani

Chugu

0 0 Kawazu Iriki

okuchl MTL

I

~ . 133°

Solgoshi

I

138°

1 ~=1 2

~

4

s ['7"1

3 r!'N 615 1

Hlshikari Yamanda Figure 1.117. Map showing the distribution of Se-type and Te-type epithermal gold deposits in Japan. 1: Green-tuff and subaerial volcanic region of Tertiary/Quaternary ages, 2: Main Paleozoic/Mesozoic terrane, 3: Main metamorphic terrane, 4: Te-type deposits, 5: Se-type deposits, 6: Te- and Se-bearing deposits, ISTL: Itoigawa-Shizuoka tectonic line, TTL: Tanakura tectonic line, MTL: Median tectonic line (Sbikazono et al., 1990). Kushikino

~

a higher level than Se mineralization in the same mine district. Aoki (1988) reported that gold precipitation is currently taking place from the Osorezan hot springs, north Honshu, Japan. At Osorezan A u - T e minerals (e.g., krennerite, coloradoite) are found at very shallow levels from the surface but no Se minerals have yet been identified (M. Aoki, personal communication, ] 988). K - A r ages data on adularia and sericite in the veins and altered host rocks indicate that ages of mineralization vary widely, ranging from 1 Ma to 68 Ma and from 1 Ma to 24 Ma for the Se-type and Te-type, respectively (Tables 1.17 and 1.18). The total production of gold, silver and other associated base metals and silver/ gold production ratios from these deposits are summarized in Table 1.17. In addition to gold and silver, lead, zinc and manganese have been produced from some of the Se-type (e.g., Yatani) and copper has been produced from some of the Te-type (e.g., Teine, Kawazu). Total tonnage of production of Au and Ag from the Se-type is greater than

161

Miocene-Pliocene Hydrothermal Ore Deposits TABLE 1.i8 Summary of geologic, mineralogic and geochemical characteristics of the Te-type and Se-type deposits (Shikazono et al., 1990) Se-type Associated metals Ag/Au Host rocks

Au, Ag, Pb, Zn, Mn More than l0 Sedimentary rocks (dominantly shale), volcanic rocks (dacite and andesite) Form of deposits Vein Age of mineralization Late Cretaceous-Quaternary Opaque minerals Argentite,polybasite, naumannite, aguilarite, Ag-rich electrum, pyrargylite, sphalerite (FeS; 1-5 wt. %), galena, tetrahedrite, chalcopyrite, pyrite Gangue minerals Quartz (fine-large grained), adularia, illite/smectite, chlorite/smectite, calcite, rhodochrosite Homogenization 150-270°C temperatures ~34S of sulfides - 8 to +5% ~348 of sulfates Sulfur activity low Oxygen activity low pH high

Te-type Au, Ag, Cu, Bi, (Hg), (TI) Less than I0 Volcanic rocks (dacite and andesite) Vein, massive Miocene-Present Hessite, petzite, native Te, sylvanite, Au-rich electrum, chaIcopyrite, pyrite, marcasite, tetrahedrite, enargite, bismuthinite, sphalerite (FeS: 1 wt. %), galena Quartz (very fine-grained), barite, illite, kaolinite, adularia 200-300°C - 3 to +7%0 +18 to +29%o intermediate intermediate intermediate

those of the Te-type. A g / A u production ratio (in weight) from the Te-type and Se-type vary widely, generally less than 10 for the Te-type and more than 10 for the Se-type. Host rocks for the Se-type and Te-type epithermal A u - A g deposits are summarized in Table 1.17. In general, the dominant host rocks for both deposits types are intermediate and felsic volcanic rocks. Sedimentary rocks (usually shale) sometimes host the Se-type (e.g., Sanru, Kohryu, Takadama, Ohmidani, Hishikari), but never host the Te-type deposits. All o f the Se-type are vein in form, and most of the Te-type are also vein in form (Table 1.18). However, some Te-type (e.g., Kobetsuzawa, Date, Suzaki, Iriki) are stratabound or massive. Total production of Au from the Te deposits is small. Size (length and width of the vein) of the Se-type vein is generally greater than the Te-type vein. Temperatures of formation can be estimated from homogenization temperatures of fluid inclusion. The typical range o f temperature o f formation for the Se-type and Te-type is 150-270°C and 200-300°C, respectively (Tables 1.18 and 1.19). The pressure correction to obtain true formation temperatures based on homogenization temperatures of fluid inclusions has not been carried out because evidence for boiling is found in the fluid inclusions from some deposits such as Chitose (Yajima, 1979), Nebazawa (Enjoji and Nakayama, 1982); furthermore, homogenization temperatures of fluid inclusions are in good agreement with electrum-sphalerite temperatures (Shikazono, 1985d), indicating that the pressure o f formation is close to the vapor saturation curve. Electrum-sphalerite geothermometer can be used to estimate the temperature of formation for some deposits

162

Chapter 1

TABLE 1.19 Temperatures of formations of the Te-type and Se-type epithermal gold deposits estimated from fluid inclusion homogenization temperatures (Shikazono et al., 1990) Deposit

Homogenization (°C) Range

Average

Sanru Koryu Chitose Yatani Takadama Nebazawa Seigoshi Omidani Hishikari Kushikino

200-300 190-340 145-349 209-273 162-240 215-260 !95-243 148-178 90-260 210-250

250 255

419

241 200 250 225 160 195 235

30 100 3600 10 232 100

Te-type deposits Teine Chitose Sado Kawazu Okuyama Takeno Kato Fuke Okuchi Kushikino

180-240 220-260 247-305 206-281 193-284 210-270 190-330 220-280 164-240 260-310

265 236 261 250 250 245 230 280

34 31 53 110 16

Se-type deposits

No. of data

17

of the Se-type but not for the Te-type, because the assemblage of electrum-argentite (or acanthite)-sphalerite-pyrite is absent in the Te-type (Shikazono et al., 1990). The dominant opaque minerals from the Se-type deposits are Se-bearing A g minerals (argentite, acanthite, polybasite, naumannite, aguilalarite, pyrargyrite), electrum, tetrahedrite, sphalerite, Se-bearing galena, chalcopyrite and pyrite (Table 1.20). The dominant opaque minerals from the Te-type are hessite, native Te, petzite, pyrite, marcasite, enargite and bismuthinite (Table 1.21 ). These Te minerals are not present in the Se-type deposits. Very rarely Se-bearing minerals (tetrahedrite, bismuthinite) occur in the Te-type deposits. Generally, sulfide minerals except for pyrite and marcasite are very poor in amount in the Te-type. On the other hand, sulfide minerals such as argentite, sphalerite and galena are abundant in the Se-type. The chemical compositions of opaque minerals (sphalerite, electrum) are different in two types o f deposits. The FeS content of sphalerite from vein deposits of the Te-type is generally lower than that of the Se-type (Fig. 1.118). However, FeS content of sphalerite from massive deposits of the Te-type (Kobetsuzawa, Suzaki) is high, ranging from 1 to 7 mol%. The Ag content o f electrum from the Se-type is higher than that from the Te-type (Fig. 1.119).

TABLE t.20 Dominant opaque and gangue minerals from the Se-type epitb_ermal gold deposits (Shikazono et al., 1990)

¢5

Deposit

Se-bearing minerals

Opaque minerals

Gangue minerals

Sanru

aguilarite, naumannite, poiybasite, pyrargyrite, stephanite

electrum, miargyrite, chalcopyrite, lahore, arsenopyrite, marcasite, pyrite, sphalerite, stibnite cinnabar

quartz, adalaria, kaolinite, seficite, calcite

Koryu

aguilarite, pearceite, polygasite, proustite, pyrargyrite

electrum, miargyrite, native silver, chalcopyrite, lahore, hematite, magnetite, pyrite, galena, sphalerite

quartz, adularia, johannsenite, chlorite, kaolinite, vermiculite, Mn-calcite

Chitose

aguilarite, argentite, pearceite, polybasite, proustite

electrum, chalcnpyrite, lahore, pyrite, galena, sphalerite

quartz, adularite, chlorite, sericite, calcite

Yatani

argentite

electrum, chalcupyrite, marcasite, pyrite, pyrrhotite, galena, sphalerite

quartz, adularia, chlorite, sericite, rhodochrosite

Takadama

aguilarite, freibergitc, naumannite, pearceite, polybasite, pronstite, pyrargyrffc, stephanite

electrum, sternbergite, chalcopyrite, fahore, marcasite, pyrite, galena, sphalerite

quartz, adularita, kaolinite

Nebazawa

argentite, polybasite, proustite, pyrargyrite

electrum, mirargyrite, native silver, chalcopydte, covellite, fahore, arsenopyrite, marcasite, pyrite, galena, sphalerite

quartz, adularia, sericite, calcite

Seigoshi

argentite

electrum, pearceile, polybasite, pyrargyrite, stephanite, chalcopyrite, fahore, galena, sphalerite

quartz, adularia, inesite, xonotlite, chlorite, mixed layer clay mineral, sericite, calcite, rhodochrosite

Omidani

aguilarite, argentite, naamannite

electrum, native silver, pearceite, polybasite, chalcopyrite, lahore, pyrite, galena, sphalerite

quartz, adularia, chlorite, ankerite, calcite

SakoshiOdomari

aguilarite, argentite, jalpaite, naumarmite, polybasite, pyrargyrite

electrum, mckinstryite, native silver, chalcocite, chalct~pyrite, covellite, fahore, famatinite, native copper, arsenopyrite, hematite, pyrite, pyrrhotite, galena, sphalerite

quartz, kaolinite, sericite

Hishikari

naumannite

electrum, chalcopyrite, marcasite, pyrite, galena, sphalerite, stibnite

quartz, adularia, montmorillonite

Kushikino

aguilarite, argentite, naumannite, polybasite, pyrargyrite, stephanite

electrum, pyrostilpnite, chalcocite, ehalcopyrite, covellite, lahore, marcasite, pyrite, galena, sphalerite

quartz, adularia, mixed layer clay mineral, smectite, calcite, zeolite, truscottite

d

r~ ,~

~"

TABLE 1.21

g

Dominant opaque and gangue minerals from the Te-type epithermal gold deposits (Shikazono et aL, 1990) Deposit

Te-bearing minerals

Date

electrum, lahore, pyrite calaverite, coloradoite, tellurantimony altaite, frohbergite, hessite, native dyscrasite, chalcopyrite, hematite, tellurium, rickardite, sylvanite, marcasite, pyrite, pyrrhotite, galena, sphalerite, realgar, stibnite tellurantimony lhhore, gold-fieldite, hessite, native bismuthinite, lahore, hakite, native electrum, bornite, chalcocite, tellurium, rickardite, stuetzite, tellurium, stuetzite chalcopyrite, emplectite, enargite, lahore, luzonite, hematite, sylvanite marcasite, pyrite, galena, sphalerite, orpiment, realgar, stibnite hessite, petzite electrum, cha]copyrite, lahore, pyrite, galena, sphalerite krennerile, native tellurium chalcopyrite, fahore, arsenopyrite. marcasite, pyrite marcasite, pyrite, orpiment, realgar coloradoite, krennerite electrum, polybasite, pyrite hessite chalcopyrite, fahore, hemusite, empressite, fahore, gold-fieldite. kawazulite, native tellurium, stannoidite, hematite, pyrite, hessite, kawazulite, native poubaite sphalerite tellurium, poubaite, nckardite, stuetzite, sylvanke, tellurobithmuthite, tetradymite altaite, hessite, tetradymite electrumm chalcopyrite, native copper, hematite, pyrite, sphalerite calaverite, hessite, kawazulite, clausthalite, kawazulite, native chalcopyrite, marcasite, pyrite, kostovite, krennerite, native tellurium, poubaite, pyrite sphalerite tellurium, petzite, poubaite. sylvanite electrum, argentite, chalcopyfite, hessite electrum fahore, marcasite, pyrite, galena, sphalerite hessite, petzite

Kobetsuzawa Teine

Chitose Mutsu Osorezan Sado Kawasu

Okuyama Suzaki

Chugu Takeno

Se-bearing minerals

Opaque minerals

Gangue minerals quartz, anatase quartz, scricite, zeolite, anatase quartz, sericite, calcite. rhodochrosite, barite

quartz, adularia, chlorite, sericite, calcite quartz quartz, barite quartz, adularia quartz, sericite, barite, anatase

quartz, chalorite, sericite quartz, barite, anatase

quartz. quartz, adularia, chlorite, kaolinite, sericite, calcite, dolomite

e~

t"5.

TABLE 1.21 (continued) Dominant opaque and gangue minerals from the Te-type epithermal gold deposits (Shikazono e~ al., 1990) Deposit

Te-bearing minerals

Agawa

calaverite, telluro-bismuthite, tetradymite hessite, sylvanite

Kato lriki

fahore, gold-fieldite, native tellurium

Yamada

hessite, tetradymite

Fuke

hes~ite

Okuchi

hessite

Se-bearing minerals

lahore, famatinite

Opaque minerals

Gangue minerals

electrum, bornite, chalcopyrite, hematite, pyrite, molybdcnite electrum, chalcopyrite, pyrite, galena, sphalerite argentite, chalcopyrite, fabore, famatinim, arsenopyrite, marcasite, pyrite, pyrrhotite electrum, chalcopyrite, pyrite, sphalerite electrum, chalcopyrim, hematite, pyrite, galena, sphalerite electrum, argentite,

quartz, sericite, calcite quartz quartz, kaolinite, anatase quartz, calcite quartz, adularia, chlorite, kaolinite, smectite quartz, adularia,

166

Chapter 1

In the Se-type gangue minerals comprise quartz, adularia, illite/smectite interstratified mixed layer clay mineral, smectite, calcite, Mn carbonates (manganoan calcite, rhodochrosite), Mn silicates (inesite, johansenite) and Ca silicates (xonotlite, truscottite). In comparison, the Te-type contains fine-grained, chalcedonic quartz, sericite, barite, adularia and chlorite/smectite interstratified mixed layer clay mineral. Carbonates and Mn minerals are very poor in the Te-type and they do not coexist with Te minerals. Carbonates are abundant and barite is absent in the Se-type. Grain size of quartz in the Te-type is very fine, while large quartz crystals are common in the Se-type. Hydrothermal alteration patterns in the Se-type and Te-type deposits are not well defined due to a lack of detailed studies of hydrothermal alteration. However, it is evident that propylitic alteration is widely developed in Se-type gold mine districts (Shikazono, 1988a). For example, chlorite and chlorite/smectite interstratified mixed layer clay mineral are abundant in the country rocks which host the Seigoshi, Fuke, Kushikino and Hishikari deposits (Inome et al., 1981; Shikazono, 1985a; Shikazono et al., 2002). Alteration characterized by the presence of adularia is well developed in some Se-type mine districts (e.g., Seigoshi, Hishikari) (Shikazono, 1985b). Argillic alteration overlies the Kushikino, Seigoshi and Takadama vein deposits (Yagyu, 1954a,b; MITI, 1979). The country rocks in Te-type mine districts suffered argillic and sericite alterations and silicification. For example, such alterations are developed in the Kawazu Te-type gold 90 80

E] Te type deposits (vein) [] Te type deposits (disseminated) [] Se type deposits

70 60

~'50 ~ 4o Ii 30 20 ¸ 10 2

3

4

s

6

7

8

1'0 1112 13t;

is

1617

geS Mol%

Figure 1.118. Frequency (number of analyses) of FeS content (mole fraction of FeS) of sphalerite from the Te-type and Se-type deposits (Shikazono et al., i990).

167

Miocene-Pliocene Hydrothermal Ore Deposits Te Type deposits • n=5 Date ~l-~n=10 Okuyama • n=2 Kushikino Agawa Chugu ,~'H. n=15 Fuke = ~n=69 Okuchi n=10 en=17 Chitose : z = n=6 Takeno == n=l Sado Kato SeType deposits

• n=l ~ n=61

Sakoshi Hishikari Koryu Chitose Sanru Kushikino Takadama Yatani Nebazawa Omidani

n=20~

x =

=

~ n=22

4

=

n=41 =

I

0

I

I

I

I

I

50

~ n=8

: .~ n=44

I

I

I N Ag

I

I

100

Figure 1.119. Ag content (atomicfractionof Ag) of electrum from the Te-typeand Se-typedeposits (Shikazono et al., 1990). n: number of analyses. mine district in Izu Peninsula (Watanabe and Nagai, 1986). Advanced argillic alteration closely related to gold mineralization has not yet been recognized in the Se- and Te-type gold deposits. Sulfur fugacity can be estimated from the homogenization temperatures of fluid inclusions, FeS content of sphalerite coexisting with pyrite, Ag content of electrum coexisting with argentite (or acanthite) and the assemblages of sulfide minerals. Typical ranges of sulfur and oxygen fugacities for the Se-type and Te-type are shown in Fig. 1.120. It is likely that sulfur and oxygen fugacities for the Te-type are higher than those for the Se-type at the same temperature conditions. Probable ranges of oxygen fugacity and pH for the Se-type and Te-type are shown in Fig. 1.121. The pH and oxygen fugacity of ore fluids responsible for the formation of the Se-type appears to be lower and higher, respectively, than for those of ore fluids responsible for the formation of the Te-type (Fig. 1.121). Sulfur isotopic compositions (~34S) of sulfides and sulfate (barite) from the Setype and Te-type are summarized in Fig. 1.122. Almost all ~34S values from the Se-type and Te-type fall in a range from -3%o to +6%o (Fig. 1.122). In general, the ~34S values from the Se-type are similar to those of the Te-type. However, some ~348 values from the Se-type are lower than those from the Te-type. The ~34S values of sulfate (barite) sulfur from the Te-type range from +18%o to +29%0 (Fig. 1.122).

168

Chapter 1 -6

#

!

i

i

t~

COVELLITE DIGENITE

~_\'~

PYRITE+BORNITE

I/

ENARGITE

!

FAMACHtNITE F , \

,o.;"~

Te-type

......~

~ /

--

\/ \

1 mole%FeS in SPHALERITE ['-.t~~',"~'~-~-I" ," • ,it."

0

4%

. _//"

i"" s -tyr, e

7-'- % . . ; % . . d . .14

-38

-37

g:

-35

,

Log ao2

-33

i

-31

Figure 1.I20. Probable ranges of sulfur and oxygen activities for the Te-type and Se-type deposits. The diagram was constructed mainly based on Heald et al. (1987). Temperature = 250°C (Shikazono et aI., 1990).

The above-mentioned geologic, mineralogic and geochemical characteristics of the Se-type and Te-type are summarized in Table 1.18. These characteristics can be used to reconstruct the structure of the fossil geothermal system responsible for the formation of the Se-type and Te-type epithermal gold mineralizations. Figure 1.123 shows the relative sites of Se-type and Te-type epithermal gold deposition in a fossil geothermal system. The Te-type is considered to form a a site closer to the volcanic centre than the Se-type. SO2 gas derived from magma disproportionates in the presence of water to form sulfate ion and hydrogen ion according to the following reactions (Holland and Malinin, 1979). 4SO2 + 4 H 2 0 --+ 3 H2SO4 q--H2S

(1-45)

H2SO4 --+ 2H + + SO 2-

(1-46)

The sulfate ion generated by these reactions precipitated as barite in some Te-type deposits. This hydrolysis reaction can explain the 334S values of barite and sulfides in the Te-type deposits. Some of the lower 334S values of sulfide sulfur in the Se-type deposits than the Te-type deposits could be due to the incorporation of sedimentary sulfur in the host rocks of the Se-type deposits. However, sulfide sulfur of the igneous source was also involved in the ore fluids responsible for the Se-type. Relatively higher sulfur and oxygen fugacities and lower pH and higher temperature of ore fluids responsible for the Te-type than the Se-type may have been caused by a contribution of magmatic (or igneous) sulfur to the ore fluids responsible for the Te-type. However, it cannot be ruled out that seawater sulfate was involved in the ore fluids responsible for the Te-type, because ~34S of barite

169

Miocene-Pliocene Hydrothermal Ore Deposits -30

I

W!W

_

I

--I>zl~

#.,

~

~L 7 -~,,~

I

J

J ....

, "1

....._ ~ / ~

......

,.<&

Te-type

I

l

ANGLESITE GALENA

'

~O.,

"V',"<'.#_ TENNANTITE

"

I

]

-

/t

#'; {/

I

l

--

- , o

-

.

.

.

s

.¢"

~N'~'~ .

10m?Ie%FeSInSPHA2ERITE I ~/ '

I

,I

0,1%

~,~~t ' E

.

i 2

4

pH

6

8

1

10

Figure 1.121. Probable ranges of oxygen activity and pH for the Te-type and Se-type deposits. The diagram was constructed mainly based on Barton et al. (1977) and Heald et al. (1987). Temperature = 250°C, ICS = 0.02 molal, Salinity = 1 molaI with N a / K (atomc ratio) = 9. Dotted area: Te-type, Hatched area: Se-type (Shikazono et al., 1990).

are close to seawater sulfate value and Te-type deposits tend to occur in Green Tuff regions characterized by thick marine sedimentary and volcanic rocks. Advanced argillic alteration is not found in either the Te-type or Se-type deposits. Sericitization, silicification and argillic alteration are common in the Te-type mine districts. Adularia is a characteristic mineral of adularia-sericite-type (Heald et al., 1987). Thus, the Te-type appears to be intermediate type between the acid-sulfate-type and adularia-sericite-type. The Se-type belongs to low sulfidation-type, essentially adulariasericite-type. Almost all Te-type deposits belong to low sulfidation-type but some have characteristics of the high sulfidation-type. Acid-sulfate, high sulfidation deposits form just above the volcanic centre. However, the Te-type form further from the site of the volcanic centre, though close to the centre than the Se-type. The Se-type may occur in basement (mainly sedimentary rocks) as well as young volcanic rocks. The models for epithermal system by Buchanan (1981), Giles and Nelson (1983), and Berger (1985) emphasize vertical change from base metal-type deposits at the deeper part to precious

170

Chapter 1 (a)

[ ] Te type deposits [ ] Se type deposits

¢o

R

"6 E

Z

-8

8

.I.I "1"1 "1-1 [ • [/Ij1 "1 "1" • I . I , I . I A J~J I - V I'1/1"1"1"1" • I-I "I "Pq [2t , I Z I ~ I - I ,, V l t l - I - . I - I t l - J : l " - I . I . I . I - I , I , . I . I , 1 2 3 4 5 6 7 -6 -5" -4 -3 -2 -1 0 •

..Q

-7

(b)

BARITE

"6 E

[]

Teine

]

Kawazu

]

osorezan

27

F~. 28 29 ~'S(%o)

y.

, l'~, 18 19

20

. if'l, 21 22

23

. ITIQ~TI 24 25 26

Figure I. 122. Frequency (numberof analyses) of (a) ~348 of sulfide and (b) sulthte (barite) minerals(Shikazono et al., 1990). metal-type deposits at the shallower part. However, the above discussion emphasizes lateral changes of epithermal ore-types in the geothermal system, as shown in Fig. 1.123. It is worth comparing the fossil geothermal system associated with epithermal gold deposition with active geothermal systems. Henley and Ellis (1983) and Henley (1985) showed that sulfate-type hot springs occur at the volcanic centre and upper level of geothermal system, while bicarbonate hot springs occur on the margins of the geothermal system. Lateral flow of acidic hot waters from upper level of active geothermal system of high relief, with significant lateral flow, appears to be consistent with the distribution of the Te-type and Se-type. Carbonates are common in the Se-type, and barite is found in the Te-type. The occurrence of carbonates and barite suggests that ore fluids for the Se-type and Te-type are bicarbonate-rich and sulfate-rich, respectively. 1.4.5. Depositional mechanism and origin of ore fluids

1.4.5.1. Depositional mechanism It is widely accepted that boiling of ore fluids took place in the epithermal A u Ag mineralization system from the fluid inclusion data (Nakayama and Enjoji, 1985;

Miocene-Pliocene Hydrothermal Ore Deposits

Sotfatara "

Propylltlc / Z

171

""

~'

Figure 1.123. Schematic model for the formations of the Te-type and Se-type epithermal gold depositions in the fossil geothermal system. Reference: Henley and Ellis (1983) (Shikazono et al., 1990).

Shikazono, 1985a) and stable isotope data (Shikazono, 1988a, 1989). Generally, f s 2 and fo2 increase in response to boiling (Drummond, 1981). If boiling occurs, H2 gas escapes from ore fluids faster than other gaseous species, resulting to higher f02 and fs2- Thus, it seems likely that Au-rich electrum tends to precipitate under the high fs2 and f02 conditions from the ore fluids from which substantial amounts of vapor released. However, if boiling of ore fluids having high CO2 and CH4 concentrations occurs, fs2 and f02 do not considerably increase (Drummond, 1981). This type of boiling or gas loss could lead to the deposition of sulfide minerals (sphalerite, galena, argentite etc.) due to a pH increase (Drummond, 1981; Hedenquist and Henley, 1985). This pH increase is not favorable for the deposition of Au when thio-Au species are dominant among dissolved Au species because solubility of Au due to thio complexes increases with the pH increase (Fig. 1.102). If electrum coprecipitates with argentite under low fs2 and f02 conditions, although the amount of electrum precipitated might be small, considering the solubility of electrum under such conditions, electrum tends to contain high Ag content. Not only the deposition of Au, but also the deposition of Ag has to be taken into account in order to consider the depositional mechanism of electrum. Deposition of Ag from the ore fluids in which Ag chloro complexes (e.g., AgCI~-) are dominant Ag species may be controlled by the following reaction. AgCI~- + 1/2 H20 -+ Ag ÷ 2 C1- + H + ÷ 1/4 02

(1-47)

Thus, a pH increase could be favorable for the deposition of Ag. This may imply that the Ag content of electrum is high when substantial amounts of CO2 loss occur. The CO2 concentration of ore fluids responsible for epithermal base-metal vein-type deposits

172

Chapter 1

NAg 1.0

"3

o2

d

0.8

0.6

23 0

•

•

11 e7 • elO o6 120

*17

25 • 22 200 ~8 e24

26

0.4

16 21 15 • • j4 19 ,w

9

o8

5 e4

13

27° 0.2

30

r'l

-,

,28

29 Q

;

+,

i

~

log ( A g / A u )

i

Figure 1.124. Ag/Au total production ratio from each mine and Ag content of electrum. Solid circle: epithermal Au-Ag vein-type deposits. Open circle: epithermaI base metal vein-type deposits. Solid square: hypo/mesothermal polymetallic vein-type deposits. Open square: epithermal Au disseminated-type deposits. I: Tada, 2: Toyoha, 3: Omidani, 4: Innai, 5: Ikuno, ~" qe-Inakuraishi, 7: Nebazawa, 8: Kawazu, 9: Todoroki, 10: Yatani, 11: Seigoshi, 12: Sado, 13: Takeno, 14.145 ,,awaji, 15: Yugashima, 16: Takadama, 17: Handa, 18: Konomai, 19: Sakoshi-Odomari, 20: Toi, 21: Sanru, 22: Arakawa, 23: Taio, 24: Chitose, 25: Hokuryu, 26: Okuchi, 27: Fuke, 28: Yamagano, 29: Akeshi, 30: Kasuga (Sbikazono, 1986).

in Japan is generally higher than that for epithermal Au-Ag vein-type deposits in Japan (section 1.4.3). It is also inferred that the deposition of argentite occurs by a pH increase, considering the following reaction. 2AgC12 + H2S --+ Ag2S + 4C1- + 2 H +

(1-48)

Therefore, it is likely that Ag-rich electrum and large amounts of sulfide minerals including argentite could precipitate due to CO2 loss and pH increase under low f s 2 and fo2 conditions. Therefore, this mechanism (boiling and gas loss from the hydrothermal solution with different fs2, J%2, CO2 concentration and pH) could explain why the Ag content of electrum correlates with Ag/Au total production ratio (Fig. 1.124). The above mechanism (boiling, loss of CO2 and increase in pH) could also lead to the deposition of other sulfides. The reactions causing sulfide depositions by this mechanism are written as, ZnC12 + H2S -+ ZnS + 2 CI + 2 H +

(1-49)

PbC12 +H2S --+ PbS + 2C1- + 2 H +

(1-50)

CuC1 + 2H2S + FeC12 --+ CuFeS2 + 3 C1- 4- 1/2 H2 + 3 H +

(1-51)

Miocene-Pliocene Hydrothermal Ore Deposits

173

As noted already, Ag content of electrum from the epithermal Au-Ag vein-type deposits is higher than that of the electrum from Kuroko deposits, and Ag/Au total production ratio of epithermal Au-Ag vein-type deposits (average 18) is lower than that of Kuroko deposits (average 76). Therefore, this relation is different from that found in the epithermal vein-type deposits. Ag/Au ratio of electrum may be controlled by the following reaction (Shikazono, 1981): (AU)el + 2H2S + Age12 = (Ag)el + 2 e l - + Au(HS)2 + 2 H +

(1-52)

It is thought from this reaction, that Au-rich electrum precipitates from ore fluids with high C1- concentration and low pH. Therefore, it is considered that different C1concentration and pH are important factors causing different relationship between Ag content of electrum and Ag/Au total production ratio of Kuroko deposits and epithermal vein-type deposits. Temperature is also important factor controlling Ag content of electrum. Shikazono (1981) showed that Ag content of electrum increases with decreasing temperature, assuming that EAg/EAu in ore fluids is constant, and dominant Au and Ag species are Au (HS)2 and AgC12. Figure 1.102 shows the iso-Au concentration contours (solubility of Au) on log fo2-pH diagram (Henley, 1984). The f Q - p H region for maximum Au concentration is close to H2S/HS- boundary and SO2-/reduced sulfur species (HzS, HS-) boundary. The most probable /O2-pH region of epithermal Au-Ag vein formation is also close to SO2-/reduced sulfur species boundary and the pH in equilibrium with adularia/sericite boundary, but the pH lower than adularia/sericite boundary is estimated based on the minerals coexisting with electrum: (1) adularia is abundant in the epithermal Au-Ag vein-type deposits but not in direct contact with electrum; (2) the minerals in direct contact with electrum are quartz, sericite/smectite and chlorite/smectite interstratified clay minerals. The most likely explanation for the variation of f Q - p H region of different epithermal Au deposition is the mixing of ascending hydrothermal solution whose foz-pH region is close to the region showing maximum Au solubility and descending acid sulfate hot spring whose temperature is lower than that of ascending hydrothermal solution. Acid sulfate solution contains appreciable amounts of SO]- but no H2S. This mechanism can explain the formation of Te-bearing Au-Ag veins in which sulfides are poor in amounts. The deposition of sulfides is generally difficult by this mechanism because solubilities of sulfides generally increase with decreasing of pH. However, if temperature of mixed fluid decreases considerably by this mechanism, the deposition of sulfides may be possible, because solubilities of sulfides due to chloro complexes decrease with decreasing of temperature. The above interpretation of formation of epithermal Au-Ag vein-type deposits is supported by (1) thermochemical calculations on this type of mixing (Reed and Spycher, 1985), and (2) the geological occurrence of epithermal Au-Ag vein-type deposits and associated advanced argillic alteration. Reed and Spycher (1985) pointed out based on thermochemical calculations that gold does not precipitate due to boiling because pH increases by boiling, leading to an increase in gold solubility due to gold-thio complexes. They indicated that acidification

174

Chapter 1

,,"0 ~

Y

. ~,~ --: [ , 2

I

f~)

J

~

.,

," ',.,, Advanced Argillic ," /~ ,' "~" Alteration , ....,_Funabara area ,,s ~

"Seigoshi

".

-- ,~'~oi ,~l~ ~ ) m i n e , mine ~ ~. ,,1 '/

,"

,"

,"

', J"

"",

,',

F~ "

, ,,.

t

-

,'o

-

'~ 5 km

Proov -"

tic

: ~,~ Alteration

:. mine

\

Amagi &

J

Mochikoshi

,, Ugusu mine .

~_.J

"

mine '

,'

~

Diorite Porphyry Silicified Rock of Advanced Argillic Alteration r - ~ Gold-SilverQuartz Vein Fault

Figure 1.125. Distribution of epithermal Au-Ag vein-type deposits, propyIitic and advanced argillic alterations and intrusive rocks of diorite porphyry (Shikazono, 1985a).

of gold-bearing boiled waters by descending acid-sulfate waters is possibly of great significance to near-surface gold precipitation. The acidification destroys the dominant gold complex, Au(HS) 2, forcing gold precipitation by: H + + Au(HS)2 + 1/2 H2 ~ Au + 2 H2S

(1-53)

Shikazono (1985a) has studied hydrothermal alterations in the epithermal AuAg mine district in Izu Peninsula, middle part of Honshu, and indicated that (1) the propylitic alteration occurs widely in the district; (2) at the centre of the district and stratigraphicalty upper horizon, there exists advanced argillic alteration; (3) epithermal Au-Ag vein-type deposits are distributed at marginal zone in the district (Fig. 1.125); (4) the age of formations of advanced argillic alteration and epithermal Au-Ag veins are nearly the same (Au-Ag vein: 1.8-1.1 Ma; advanced argillic alteration: 2.2-1.2 Ma); and (5) epithermal Au-Ag deposition (Seigoshi) occurred in intermediale f Q - p H region of advanced argillic and propylitic alterations.

Miocene-Pliocene Hydrothermal Ore Deposits

175

By considering the features of hydrothermal system associated with epithermal Au-Ag mineralization, a model for the hydrothermal system is constructed in Fig. 1.123. Schematic diagrams showing the sites of Te-bearing and Se-bearing Au-Ag vein-type deposits in hydrothermal system are constructed based on the mineralogic, geologic and geochemical features of these deposits (Fig. 1.123). SO ] . in the acid sulfate hot springs responsible for advanced argillic alteration and/or epithermal Au-Ag ore fluids is considered to be of volcanic SO2 gas origin because ~348 of sulfates (alunite, barite) in advanced argillic alteration zone and in Te-bearing veins is high, more than +20%0. This is reasonably explained by the generation of SO ] - and H2S by the following hydrolysis reaction of volcanic SO2 gas. 4802 + 4 H20 "-'+ 3 H2SO4 + H2S

(1-54)

It seems likely that the mixing of acid sulfate solution with nearly neutral low salinity Au-bearing fluids seems the most likely mechanism for the formation of epithermal Au-Ag vein-type deposits. This mechanism as a main cause for epithermal-type Au deposition is supported by sulfur isotopic data on sulfides. Shikazono and Shimazaki (1985) determined sulfur isotopic compositions of sulfide minerals from the Zn-Pb and Au-Ag veins of the Yatani deposits which occur in the Green tuff region. The values for Zn-Pb veins and Au-Ag veins are ca. +0.5%0 to +4.5%o and ca. +3%o to +6%0, respectively (Fig. 1.126). This difference in 334S of Zn-Pb veins and Au-Ag veins is difficult to explain by the equilibrium isotopic fractionation between aqueous reduced sulfur species and oxidized sulfur species at the site of ore deposition. The non-equilibrium rapid mixing of H2S-rich fluid (deep fluid) with SOl--rich acid fluid (shallow fluid) is the most likely process for the cause of this difference (Fig. 1.127). This fluids mixing can also explain the higher oxidation state of Au-Ag ore fluid and lower oxidation state of Zn-Pb ore fluid. Deposition of gold occurs by this mechanism but not by oxidation of HzS-rich fluid. This type of mixing could reasonably explain the occurrence of acidic alteration minerals such as kaolinite and alunite in the low-sulfidation epithermal gold vein district (e.g., Seta in northeast Hokkaido, Hishikari in southern Kyushu) (Yajima et al., 1997) Mixing of high temperature hydrothermal solution with high salinity and low temperature solution with low salinity of meteoric water origin seems the most likely mechanism for the base-metal vein-type deposition. The salinity-temperature relationship obtained by fluid inclusion study on the Toyoha and Ohe P b - Z n - M n vein-type deposits and Fujigatani-Kiwada W skarn-type deposits indicates positive correlation (Fig. 1.128). This indicates that the mixing of high temperature fluids of deep-seated origin with low temperature fluid of meteoric water origin, together with the conductive cooling of mixed fluids caused the formation of these ore deposits (Fig. 1.129) (Shibue, 1991). Another possible mechanism for the ore deposition is boiling of fluids. For example, the solubility of galena is controlled by, PbS + 2H + + 2C1- = PbC12 + H2S

(1-55)

If boiling took place, H2S gas removes from the fluids. But CO2 loss causes an increase in pH, leading to the deposition of galena. However, no evidence of boiling of

176

Chapter 1

Galena

[]

Au-Ag

•

Zn-Pb vein

I

I

I

ll,ll

vein

i

,P1

,I"1 i

Chalcopyri le

Sphalerite

!

I

IL

~ 1

I

Pyri{e

i

~

Bulk

'

6 ].1

1

!

in

i

i

|

l

i

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,i

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I .,.I,,. ,

n

2

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,

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,nR

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, J

4

5

6 (~ 34 S ( ° / o o

)

Figure 1.126. Sulfur isotopic compositions of sulfide minerals from Au-Ag veins and Zn-Pb veins from the Yatani mine (Shikazono and Shimazaki, 1985).

fluids in inclusion fluids from the base-metal vein-type deposits is observed. Therefore, boiling seems unlikely as a depositional mechanism for galena as well as other common sulfides (chalcopyrite, sphalerite, pyrite).

1.4.5.2. Origin of ore fluids 3D and ~180 data on fluid inclusions and minerals at main stage of epithermal A u A g mineralization clearly indicate that the dominant source of ore fluids is meteoric water. Meteoric water penetrates downwards and is heated by the country rocks a n d / o r intrusive rocks. The heated water interacts with country rocks a n d / o r intrusive rocks and extracts sulfur, Au, A g and other soft cations (e.g., Hg, T1) from these rocks. If hydrothermal solution boils, it becomes neutral or slightly alkaline, leading to the selective leaching of soft cations such as Au, Ag, Hg and T1 from country rocks. However, a contribution of sulfur gas and other components from m a g m a cannot be ruled out.

177

Miocene-Pliocene Hydrothermal Ore Deposits

Au.Ag. ore fluid "~Zn.Pb

(a)

ore fluid

ore fluid

A

~-Zn.Pb

(b) 8 a4Sr.s.s.

ore fluid

D 8 34Sr.s.s.

Figure 1.127. (a) Schematic representation of change in ~SO42-/EH2S ratio (concentrationratio of oxidized sulfur species and reduced sulfur species) and 334S value of reduced sulfur species (334Sr.s.s.) in ore fluids during oxidation of reduced sulfur species at constant temperature and ES (total dissolved sulfur concentration). Dotted area represents the possible region for the fluids accompanied by oxidation of reduced sulfur species. A: initial values of 334Sr.s.s. and zsoZ-/EH2S in fluids prior to the oxidation. B: final values of 834Sr.s.s. and ~SO2 /EH2S in fluids when ~)34Sr.s.s. remains constant during oxidation. C: final values of 834Sr.s.s. and ESO2 /EHzS in fluids when isotopic equilibriumbetween reduced sulfur species and oxidized sulfur species is attained during the oxidation. Hatched areas show 334Sr.s.s. and ZSO]-/EH2S estimated for Zn-Pb ore fluid and Au-Ag ore fluid. (b) Schematic representation of change in ESO]-/EH2S ratio of ore fluids and 334Sr.s.s. value in ore fluids during mixing of t-I2S-richfluid (D) and SO42--rich fluid (E) at constant temperature. Dotted area represents the probable region resulted from the mixing of two fluids (D and E). F: final values of 834Sr~.s. and ESO]-/EH2S of the mixture of C and D fluids when isotopic equilibriumis attained betweenreduced sulfur species and oxidized sulfur species during the mixing. Hatched areas represent ~34Sr.s.s. and ESO]-/EH2S estimated for Zn-Pb ore fluid and Au-Ag ore fluid (Shikazono and Shimazaki, 1985).

8180 of late-stage hydrothermal solution is high (0%o to +3%o), as recognized in the Seigoshi and Hishikari A u - A g veins (Shikazono, 1988a; Shikazono and Nagayama, 1993). This increase in 8180 with the stage of hydrothermal system may be due to the change in water/rock ratio, boiling and kind of rocks interacting with fluids. ~D and 3180 data on fluid inclusions and minerals, 313C of carbonates, salinity of inclusion fluids together with the kind of host rocks indicate that the interaction of meteoric water and evolved seawater with volcanic and sedimentary rocks are important causes for the formation of ore fluids responsible for the base-metal vein-type deposits. High salinity-hydrothermal solution tends to leach hard cations (base metals, Fe, Mn) from the country rocks. Boiling may be also the cause of high salinity of base-metal ore fluids. However, this alone cannot cause very high salinity. Probably the other processes such as ion filtration by clay minerals and dissolution of halite have to be considered, but no detailed studies on these processes have been carried out. Origin of sulfide sulfur of epithermal base-metal veins is thought to be same as that of Kuroko deposits because average ~34S value of base-metal vein-type deposits is +4.7%o which is identical to that of Kuroko deposits (+4.6%o) (Shikazono, 1987b). Namely, sulfide sulfur of base-metal veins came from igneous rocks, sulfate of trapped seawater in marine sedimentary rocks, calcium sulfate (anhydrite, gypsum) and pyrite. ~34S of sulfide sulfur of epithermal base-metal vein-type deposits can be explained by the interaction of seawater (or evolved seawater) with volcanic rocks. There are the following three possibilities for the origin of sulfide sulfur of epithermal A u - A g vein-type deposits:

178

Chapter 1 t-" Q)

5-

a3

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Toyoha

Q)

o Z

32-

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•

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•

•

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tco

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I

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100

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200

3O0

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350

Homogenization (°C) t-8-

~r

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6-

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•

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lafe •

2o

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o~

o

sal-o e~.at. (198oI

8

0

6

./.

150

ii 4ko4 •

•

• I

I

[

100

10-

5

04

•

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J

300

250

Fujigaf'ani-Kiwada •

•

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z,~

:.

i ||~'!% '°z ;• " .:. r'.'i ..

Z

•

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.

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0 loo

150

200

Homogenization (°C)

250

300

Figure 1.128. Plots of homogenization temperature against salinity of fluid inclusions from the •he (Mn-PbZn), Toy•ha (Pb-Zn-Mn), and Fujigatani-Kiwada (W) deposits (Shibue, 1991).

(1) Leaching of sulfide sulfur from subaerial young (Miocene-Pliocene) volcanic rocks (Shikazono, 1987b). (2) Same origin of sulfur for base-metal vein-type deposits and Kuroko deposits. (3) Contributions both from sulfide sulfur leached from volcanic rocks and marine sulfate in Green tuff.

179

Miocene-t'liocene Hydrothermal Ore Deposits c-

"-~

o - - o C o r r e t a f i o n I.ine

6

OQ

•

Conductive tooting

0

~0

0 0 ~

z 0

•~ co

0

100

1

150

0

000

I

O

I

200 250 Homogenization temperature (°C)

I

300

I

350

Figure 1.129. Illustrationof the regressionprocedureof fluid inclusion data on the Ohe deposit(Shibue, 1991).

Possibility (1) was proposed by Shikazono (1987b) who considered that the lower 3348 values of sulfide sulfur than base-metal vein-type deposits and Kuroko deposits can be explained by the leaching of sulfide sulfur from volcanic rocks with lower 334S values (0%0 to +5%o) (Uyeda and Sakai, 1984). (2) is also possible. It has been shown that a large amount of SO ] - together with reduced sulfur species such as H2S and H S - were present in epithermal Au-Ag ore fluids. Therefore, it is likely that 334S of reduced sulfur species whose values are typical Green tuff value decreases due to SO]--H2S(HS - ) fractionation. As noted already, Shikazono (1999b) divided epithermal Au-Ag deposits into Green tuff-type and Non-Green tuff-type based on their distributions. 334S values of Green tuff-type are higher than those of Non-Green tuff-type (Fig. 1.110). Host rocks, basement rocks, distribution of ore deposits, total production of Au and Ag, Ag/Au total production ratio, metals produced, 313C and 3180 of carbonates from the two types of deposits are different (Table 1.18). For instance, 334S values of sulfides from Green tuff-type are higher than those from Non-Green tuff-type. These differences are clearly interpreted by that sulfide sulfur of Green tuff-type was contributed by marine sulfate sulfur but that of Non-Green tuff-type is by igneous and sedimentary sulfide sulfur. The above-mentioned consideration on the origin of Neogene epithermal vein-type and Kuroko deposits is roughly consistent with the view by Mosier et al. (1986). They examined grades, tonnages and basement rocks for 88 epithermal precious and base-metal quartz-adularia-type districts in North and Central America and Japan and revealed that the type of basement rock below the mineralized veins is useful for predicting grade and size of deposits. Epithermal districts overlying basement with salt and evaporites of rocks with trapped sea water have a median tonnage (production and reserves) of 1.4 million metric tons and median grades of 1.5 g/ton Au, 130 g/ton Ag, 2.5% Pb, 1.7% Zn and 0.16% Cu, and districts overlying sedimentay basements have a median tonnage of 0.77 million metric tons and median grades of 7.5 g/ton Au, 110 g/ton Ag, less than 0.025% Zn, less than 0.005% Cu and less than 0.001% Pb, and districts overlying igneous basements have a median tonnage of 0.3 million metric tons and median grades of 5.9 g/ton Au, 38 g/ton Ag, tess than 0.25% Zn, less than 0.002% Cu, and less than

180

Chapter 1

0.003% Pb. Their results clearly demonstrate that basement rocks affect vein components in epithermal precious and base-metal quartz-adularia-type deposits. As noted already, epithermal vein-type deposits are classified primarily on the basis of their major ore-metals (Cu, Pb, Zn, Mn, Au and Ag) into the gold-silver-type and the base-metal-type. Major and accessory ore-metals from major vein-type deposits in Japan were examined in order to assess the possible differences in the metal ratios in these two types of deposits (Shikazono and Shimizu, 1992). Characteristic major ore-metals are Au, Ag, Te, Se and Cu for the A u - A g deposits, and Pb, Zn, Mn, Cu and Ag for the base-metal deposits (Shikazono, 1986). Accessary metals are Cd, Hg, T1, Sb and As for the A u - A g deposits and In, Ga, Bi, As, Sb, W and Sn for the base-metal deposits (Table 1.22, Shikazono and Shimizu, 1992). Minerals containing Cu, Ag, Sb and As are common in both types of deposits. They are thus not included in Table 1.22. As already noted in section 1.4.3, geochemical features of ore fluids responsible for base-metal and gold-silver types of deposits are distinct. They are summarized in Table 1.22. The differences in metals concentrated to the deposits and geochemical fectures of ore fluids responsible for both types of deposits are interpreted in terms of HSAB (hard, soft, acids and bases) principle by Pearson (1963, 1968) below. Ahrland et al. (1958) classified a number of Lewis acids as of (a) or (b) type based on the relative affinities for various ions of the ligand atoms. The sequence of stability of complexes is different for classes (a) and (b). With acceptor metal ions of class (a), the affinities of the halide ions lie in the sequence F - > CI- > Br- > I - , whereas with class (b), the sequence is F - < C1- < B r - < I - . Pearson (1963, 1968) classified acids and bases as hard (class (a)), soft (class (b)) and borderline (Table 1.23). Class (a) acids prefer to link with hard bases, whereas class (b) acids prefer soft bases. Yamada and Tanaka (1975) proposed a softness parameter of metal ions, on the basis of the parameters En (electron donor constant) and H (basicity constant) given by Edwards (1954) (Table 1.24). The softness parameter ~ is given by c~/(oe +/~), where c~ and/~ are constants characteristic of metal ions. They indicated that the softness parameter may reasonably be considered as a quantitative measure of the softness of metal ions and is consistent with the HSAB principle by Pearson (1963, 1968). Wood et al. (1987) have shown experimentally that the relative solubilities of the metals in H20-NaC1-CO2 solutions from 200°C to 350°C are consistent with the HSAB principle; in chloride-poor solutions, the soft ions Au + and Ag + prefer to combine with the soft bisulfide ligand; the borderline ions Fe 2+, Zn 2+, Pb 2+, Sb 3+ and Bi 3+ prefer water, hydroxyl, carbonate or bicarbonate ligands, and the extremely hard Mo 6+ bonds only to the hard anions O H - and 0 2 - . Tables 1.23 and 1.24 show the classification of metals and ligands according to the HSAB principle of Ahrland et al. (1958), Pearson (1963, 1968) (Table 1.23) and softness parameter of Yamada and Tanaka (1975) (Table 1.24). Comparison of Table 1.22 with Tables 1.23 and 1.24 makes it evident that the metals associated with the gold-silver deposits have a relatively soft character, whereas those associated with the base-metal deposits have a relatively hard (or borderline) character. For example, metals that tend to form hard acids (Mn 2+, Ga 3+, In 3+, Fe 3+, Sn 4+, MoO 3+, WO 4+, CO2) and borderline acids (Fe 2+, Zn 2+, Pb 2+, Sb 3+) are enriched in the base-metal deposits, whereas metals that tend to form soft acids

181

Miocene-Pliocene Hydrothermal Ore Deposits TABLE 1.22 Accessory metals from vein-type deposits in Japan (Shikazono and Shimizu, 1992) Mine

Metal Hg

T1

Cd

Bi

Gold-silver type Konomai Kitanoou Showa Meiji Takadama Hirukodate Okuchi Hazami Osorezan Tsugu Chitose Todoroki Seigoshi Kushikino Teine Nishizawa Kawazu Ifiki Yatani Base-metal type Jizo Akenobe Toyoha Ashio Goka Miyatamata Agenosawa Akarimata Hosokura Nakanosawa Fukoku Ikuno Taishu Suttu Hayakawa Akagane Kishu Kutosan Kurokawa Goka Omidani Imaiishizaki Inakuraishi Tada Fukoku Omodani

Konjo Ryujima

Mo

Sn

W

X X

X

X X

X

X

X

X

X

X

X

X

X

X X X X X X X X X

X

Chapterl

182 TABLE 1.23 Classification of metals and Iigands according to the HSAB principle Mn2+, Ga3+, In3+, Co2+, Fe3+, As3+, Sn4+' MoO3+, WO4+, Co2+ Cu+ , Ag+ , Au+ , T1+, Hg+, Cd2+, Hg2+, Te4+ , TI3+ Fe2+, Co2+, Ni2+, Cu2+, Zn2+, pb2+, Sn4+, Sb3+, Bi3+, SO2 OH , CI-, CO2, SO2 H2S, HS-, S 2

Hard acids: Soft acids: Borderline acids: Hard bases: Soft bases:

TABLE 1.24 Softness parameter of the various metal ions Metal ions

Softness parameter

Metal ions

Softness parameter

Ag+ Hg+ TI+ Cu+ Cd2~ Ni2+ Zn3+ Co2+ Zn2+

1.03 1.01 0.98 0.96 0.96 0.94 0.93 0.92 0.91

CU 2+

0.89 0.87 0.85 0.84 0.82 0.78 0.73 0.58

Bi3+ Pb2+ Fez+ Mn2+ Fe3+ Sn2+ Ga3+

(Ag +, Au +, T1+, Te 4+, T13+) are enriched in the g o l d - s i l v e r deposits. Metals that have high values of the softness parameter (Ag +, Hg +, T1+, Cd 2+) are associated with the g o l d - s i l v e r deposits, whereas those that have low values o f the softness parameter (Zn 2+, In 3+, Bi 3+, Pb 2+, Te 4+, Mn 2+, Sn 4+, Ga 3+) are found with the base-metal deposits. These correlations mean that the HSAB principle could be a useful approach to evaluate the geochemical behavior o f metals and ligands in ore fluids responsible for the formation of the epithermal vein-type deposits. A m o n g the ligands in the ore fluids, H S - and H2S are the most likely to form complexes with the metals concentrated in the gold-silver deposits (e.g., Au, Ag, Cu, Hg, TI, Cd), whereas C1- prefers to form complexes with the metals concentrated in the base-metal deposits (e.g., Pb, Zn, Mn, Fe, Cu, and Sn) (Crerar et al., 1985). Generally, the complexes with intermediate or hard ligands (e.g., chloro complexes) should become more stable with increasing temperature than complexes with soft ligands (e.g., bisulfide complexes) (Seward, 1981 ; Crerar et al., 1985). The higher temperatures of formation of the base-metal deposits (Table 1.13) also are in accordance with the HSAB principle. It was shown in Table 1.13 that the base metal and g o l d - s i l v e r types of deposits formed at different temperatures and concentrations of C I - , H S - , and CO2. Thus, it could further be inferred that the H S A B principle can be successfully applied to the genesis of these vein-type ore deposits, formed mostly at less than ca. 250°C. Crerar et al. (1985) noted that Pearson's rule (the HSAB principle) successfully describes speciation to about 250°C, but may break down at higher temperatures, as all metals become harder.

Miocene-Pliocene Hydrothermal Ore Deposits

183

A few appilications of the HSAB principle to hydrothermal ore deposits have been carried out (Crerar et al., 1985; Wood et al., 1987; Shikazono and Shimizu, 1992). These studies demonstrate that the HSAB principle is useful in interpretations of the metal ratios in ore deposits. For example, Wood et al. (1987) has shown that gold is transported by lower salinity fluids than the base metals, and this difference in salinity is a significant factor in the separation of gold and base metals in Archean deposits in greenstone belts. Cathles (1986) also has indicated that the solubility of gold is much greater in lowsalinity solutions; this can explain the bimodal populations of base-metal-rich, gold-poor deposits (stratiform deposits similar to Kuroko deposits in Japan) and base-metal-poor, and gold-rich deposits in greenstone belts (lode gold deposits). Such differences in the salinity of ore fluids responsible for the epithermal gold deposits and Kuroko and base-metal vein-type deposits in Japan have been pointed out also by Shikazono and Shimizu (1993). Therefore, it could be inferred that the difference in salinity is a main cause for a separation of gold and base-metals in mineralized zones in young (Tertiary to Quaternary) Japanese epithermal vein-type and Kuroko deposits in volcanic terranes and in Archean deposits in greenstone terranes. However, the concentration of CO2 in ore fluids responsible for Japanese gold-silver deposits and Archean lode gold deposits in greenstone belt seems to be different; CO2 concentration is low (0.01-0.1 molal) for Japanese deposits (Shikazono, 1985a), whereas it is high (0.05-2 molal) for Archean deposits (Wood et al., 1987). The reason for this difference is not known.

1.4.6. Hishikari deposit: an example of Japanese epithermal Au-Ag vein-type deposits One of the most important steps made in the research of Japanese epithermal gold deposits during the last two decades is the comprehensive study of the Hishikari gold deposits which occur in southern Kyushu (Fig. 1.130). The deposits are characterized by high gold grade and large amounts of gold reserve. Average gold grade in the Honko (Main) deposits is ca. 70 g/metric ton which is enormously high, compared with the other Japanese epithermal gold deposits. Ore reserve is estimated to be about 250 metric tons Au, which is the largest among the Japanese epithermal gold deposits. Ag/Au production ratio is about 0.7, which is relatively low. The Hishikari deposit is only one which belongs to a giant (Laznicka, 1983) and bonanza (Sillitoe, 1993) epithermal-type deposit (Izawa, 2001). Much research on the Hishikari deposits has been carried out since the discovery of gold veins in 1981. Urashima and Izawa (1983) reported fluid inclusion studies using core samples. Abe et al. (1986) made a detailed description of the veins. The regional geology of this district is described in MMAJ and SMM (1987) and Urashima et al. (1987). Izawa et al. (1990) reviewed the studies on the geology, geophysics and geochemistry which had been done during 1980s. The Special Issue of Resource Geology on the Hishikari deposits (Shikazono et al., 1993) includes various aspects of the Hishikari deposits (oxygen isotopes of gangue minerals, hydrothermal alteration, precipitation sequence, fluid inclusions, vertical electric profiling and electric sounding surveys, structural geological analysis, opaque minerals,

184

Chapter 1

(/-f / -~ k~

~

k ) ,

[ | ]

.~v-/

• Hot springtype O EpithermalAu-Agvein-type A Activevolcano o

. , ~ /

r

,

20 I

,

40

I(km)

00kuehi

/

OYamagano

Miyazaki0 (

Kagoshimao) t. A /.j

e /

131E Figure ].130. A map showing the location of the Hishikati mine.

prospecting, etc.) which had been studied during late 1980s and early 1990s. No review on the studies during the last decade has been published in English. Thus, a review of the studies done during the last decade and recent studies by the author will be summarized below.

1.4.6.1. Geology and vein system A detailed general geology of the Hishikari district has been done by many investigators (Abe et al., 1986; MMAJ and SMM, 1987; NEDO, 1991; Izawa et al., 1993). The general geology of this district is briefly described below. The district is composed of sedimentary rocks of the pre-Paleogene Shimanto Supergroup (dominantly shale and sandstone) and Quaternary andesitic and dacitic volcanic rocks. The Shimanto Supergroup is comprised of shale, sandstone and their alternations. Although no fossil data are available, the age of sedimentation is thought to be middle to upper Cretaceous age from its lithology (Izawa et al., 1990). The Shimanto

185

Miocene-Pliocene Hydrothermal Ore Deposits

Supergroup rocks in the Hishikari district suffered hydrothermal alteration. Chlorite, quartz and sericite occur abundantly near the veins. The other constituents are pyrite, albite, calcite and organic matter. Quaternary volcanic rocks unconformably overlie the Shimanto Supergroup. Quaternary volcanic rocks consist of Hishikari Lower Andesites (0.98-1.62 Ma), Hishikari Middle Andesites (0.78~0.79 Ma), Kurozonsan Dacites (0.95-1.56 Ma), Shishimano Dacites (0.66-1.6 Ma), and Hannyaji Welded Tuff (0.56-0.731 Ma) (Izawa et al., 1990, 1993). The Hishikari Lower Andesites consist of hypersthene-augite andesite lava flows in the upper horizon and andesitic pyroclastic rocks in the lower horizon. The Kurozonsan Dacites overlying the Hishikari Lower Andesites dominantly consist of hyperstheneaugite-bearing biotite-hornblende dacide lava flows. The Hishikari Middle Andesites overlying the Hishikari Lower Andesites consist of hypersthene-augite andesite lava flows and pyroclastic rocks. The Shishimano Dacites overlying the Hishikari Middle Andesites consist mainly of biotite-hornblende dacite lavas. The Hishikari Upper Andesites overlying the Hishikari Lower Andesites and the Shishimano Dacites consist of hypersthene-augite andesite lava flows and their pyroclastic rocks. The deposit consists of three vein systems, the Honko, Sanjin and Yamada veins (Fig. 1.131). The veins strike N30°E to N50°E and dip 70-90°NW. The veins are hosted 0

0.5

I

1 km k,

,

I

\

. . . .

, ,. __,, " - ~ ' ~

25

J

" P" 7XX, X~,

25

I I

N

-'SF

°

•,y •

'-..

f -,~j

~

, , ~

~.

"

X

°" .Z.

S'~"

~ "

I:zonel

I1:zoneII

II1: zone III

Figure 1.131. Plan of alteration mineral zonation of the volcanic rocks in the Hishikari area (Izawa et al., 1990).

186

Chapter 1

in hydrothermally altered Shimanto Supergroup (Honko (Main) and Sanjin veins), and Quaternary andesite (Yamada veins).

1.4.6.2. Hydrothermal alteration Izawa et al. (1990) recognized the following alteration zones from the vein towards margin of the Hishikari A u - A g mine area, chlorite-sericite zone (zone IV), interstratified clay mineral zone (zone III), quartz-smectite zone (zone II) and cristobalite-smectite zone (zone I) and least altered zone (L.A. (least altered) zone) (Fig. 1.131).

1.4.6.3. Mineralogy Quartz and adularia are dominant gangue minerals in the veins. Small amounts of sericite/smectite, calcite, and truscottite occur as vein staff. Fine-grained electrum, naumannite, Ag-sulfosalts (pyrargyrite, polybasite), chalcopyrite, galena, pyrite, marcasite and stibnite are found. Detailed mineralogical descriptions are given in Urashima and Izawa (1983), Urashima et al. (1987), Izawa et al. (1990), and Nagayama (1993a). The sequence of mineralization has been studied by Nagayama (1993a) and Takahashi et al. (1998) (Fig. 1.132). Quartz is the most abundant mineral occurring throughout the vein. Adularia tends to occur at earlier stage. Smectite is the earliest mineral. Electrum tends to occur in early and middle stages. Naumannite occurs after the early-stage

"•,•quenceveil,of

t

miner~,.,

a

b

c

d

I

e

f

g

h

II

i

Quartz

111

m

m

Adularia -i

~0

Smectite Electrum &"

.k"

:6_'

e6_

Naumannite

"E ©

Ag-Sb-S minerals

Pyrite Chalcopyrite Sphalerite m

abundant ~

rich

~

common-

poor - - f e w

Figure 1.132. Paragenetic sequence of the Hosen No. 5 vein (Takahashiet ai., 1998).

Miocene-Pliocene Hydrothermal Ore Deposits

187

electrum. Ag-Sb-S minerals are found both in early- and late-stages. Abundant pyrite occurs in early-stage. Small amounts of chalcopyrite are widely distributed. Sphalerite is found in late-stage. Stibnite and realgar are found in the latest stage. The sequence of mineralization from early to late is consistent with that of the other deposits (Asahi, Ogane, Takadama) sulfide stage (pyrite, marcasite, chalcopyrite, sphalerite, galena, electrum, naumannite, polybasitc, quartz, adularia), sulfosalt stage (pyrite, chalcopyrite, electrum, tetrahedrite, pyrargyrite, miargyrite) adularia, quartz (calcite), and barren stage (adularia, quartz, pyrite, realgar, stibnite) (Fig. 1.133). Adularia/quartz ratio decreases with the stage of mineralization and correlates with Au content of ore (Shikazono and Nagayama, 1993).

1.4.6.4. Geochemicalfeatures Oxygen and hydrogen isotopes. Oxygen and hydrogen isotopic studies on the altered rocks and gangue minerals were reported by Shikazono and Nagayama (1993), Naito et al. (1993), Matsuhisa and Aoki (1994), and Imai et al. (1998). Shikazono and Nagayama (1993) found that 8180 of adularia and quartz and adularia/quartz ratio decrease from early- to late-stage of the Hishikari mineralization, gold contents of ore samples in a vein positively correlate with 3180 and adularia/quartz ratio, as well as, the Pb, Cu, Ag and Se contents of vein samples, and Hg and As contents of the same samples increase towards the late-stage of mineralization (Fig. 1.104). Using fluid inclusion data and 3180 of adularia and quartz, it is inferred that 8180 of ore fluids decreases with the stage of mineralization (Shikazono and Nagayama, 1993). The early-stage ore fluids were of highly exchanged meteoric waters or magmatic water origin, while late-stage ore fluids were dominated by meteoric water (Shikazono and Nagayama, 1993). 8180 and 3D of ore fluids were influenced by clay minerals in the Shimanto Supergroup sedimentary rocks (Imai et al., 1998). The 813C value of H2CO3 and the 8~80 value of the hydrothermal solution using the data on vein calcite coexisting with electrum and assuming isotopic equilibrium with calcite and the temperature obtained by fluid inclusion microthermometry range from -14.4%o to -9.1%o, and from -6.2%o to +5.4%o, respectively, suggesting that the hydrothermal solutions isotopically equilibrated with the sedimentary basement rocks were responsible for the gold mineralization associated with calcite (Imai and Uto, 2001). Naito et al. (1993) showed that oxygen isotopic composition (81SO) of altered volcanic rocks in the Hishikari mine area varies systematically; +5.9%o to +15.9%o (zone I), +7.1%o to +12.4 %0 (zone II), +2.8 to +11.7 %o (zone III), and +2.1%o to +8.2 %o (zone IV) (Fig. 1.134). They calculated the change in 8280 values of hydrothermally altered volcanic rocks as a function of water to rock ratio by weight and temperature, assuming that oxygen isotopic equilibrium is attained in a closed system, and demonstrated that the increase in 8180 values of altered andesitic rocks from the veins towards peripheral zones can be interpreted as a decrease in temperature from the vein system (Fig. 1.135). In their calculations, the effect of mixing of hydrothermal solution with groundwater was not considered.

! 88

Chapter 1

Later

Earlier

Sulphide stage Sulphosalt stage Adularta Quartz Calcite Pyrite Marcasite Sphalerlte Chalcopyrlte Galena Gold Argentita Polybasite Tetrahedrlte Pyrargyrlte Mlargyrlte Realgar

mare

Barren stage

i

I

iwua=

I lib i

milm i

-? n

.,,?

Paragenetic sequence of the Asahi deposit Olta pref. Matsukuma (1951)

II Quartz Adularla Pyrite Chalcopyrlte Sphalerlte Galena Ag-mlnerele Gold Rhodochroalte Barite

~

III

-

~

--

-.-,

Psregenetic sequence of the Nishltanl vein, Nlahltanl vein group, the Ogane mine, Hokkaldo. Aklba (1957)

A

B

C

D

Quartz Adularla A:Big adularla crystals with fine quartz B:Aggregatee of adularla and quartz C:Masslve quartz with adulsrla l D:Drusy quartz General feature of paragenetlc sequence of gangue minerals, Tekatama mine, Fukushlma pref, modified from Yagyu (1954) Figure 1.133. "The normal order" in other epithermal gold deposits in Japan (Nagayama, 1993a).

189

Miocene-Pliocene Hydrothermal Ore Deposits

I 8.9+-1.4%-,n=50)

10

L.A. (least altered)

III IIII IIIIII

iI

Zone 110.9:1:2.7%.,n:~,l Cr-Sm 5 Cr-Ka(Ko) 6

5

6

g

I

/ CM-Ser l |/

. o

I I III

5

I II

II

lo

I"l R - F ~

Zone ( 6.9-+ 1,8% , n=32} 5 t Int.ClaYmirerolc I-13

i 1 i i I-I-I

I11

IK~KaI

t I[ Zone (9.0±1.8%o, n=ll) Oz-Srn 5 Qz-Ko(Ko)

6

llll

I I I

Vl~ lo

15

VI

1~

I , r-I I-I

lo

1is

[7 Illllll [-'1 i i Ii Ii Ii Ii Ii 'i 'i I I I rl s

[-1

lb

1~

a~8o (%0) Figure I.i34. Histograms of 3[80 values of the volcanic rocks for five distinct alteration zones (Naito et aI., 1993). Shikazono et al. (2002) interpreted the ~180 zonation based on hydrothermal solution-groundwater mixing model. For the calculations the following assumptions were given: (1) Hydrothermal system at discharge zone is composed of five reservoirs such as ore deposit/zone IV boundary, zone III/II boundary, zone I I / I boundary, zone I/fresh country rocks boundary and temperature of each reservoir is 250°C, 220°C, 150°C, 100°C, and 25 °C, respectively.

190

Chapter 1

250C-

+20.C

+15. O

~

-

65°C 100°(:;_ 113°C_

O 0 +10.0 ,.,f 0

t43

1500C +5.0

200°C 250°(3 300°(3 I

1

i

2

i

3

I

4

i

5

1

6

Water/Rockratioin Weight

I

7

9

Figure 1.I35. Change in 5180 values of hydrothermally altered rocks as a function of water to rock ratio with several different temperatures. The initial 3180 values of rock and water were taken as +8.5%o and -5.0%o, respectively (Naito et al., 1993).

(2) Initial 3180 of hydrothermal solution and groundwater are 0%0 and -7%0, respectively. (3) Oxygen isotopic equilibrium between mixed fluid and alteration minerals is attained. (4) Minerals in oxygen isotopic equilibrium with mixed fluid are feldspar for ore deposit/zone IV, and zone III/II, montmorillonite, and kaolinite for zone II/I, and montmorillonite for zone I/fresh rocks. Oxygen isotopic fractionation factors used for the calculation were taken from Taylor (1997). Initial 8180 value of hydrothermal solution (0%0) was estimated from 8180 values of K-feldspar and quartz in the veins and homogenization temperatures (Shikazono and Nagayama, 1993), and that of groundwater (-7%0) was estimated from meteoric water value of the south Kyushu district (-7%0) (Matsubaya et al., 1975). The mixing ratio of hydrothermal solution and groundwater was calculated based on the temperature of each reservoir. Using the estimated 81SO of fluid and oxygen isotopic fractionation between water and mineral, 3180 of altered rocks were estimated. The calculated result, together with the average 8180 of hydrothermal alteration zone by Naito et al. (1993), is shown in Fig. 1.136. This shows a fairly good agreement

191

Miocene-Pliocene HydrothermaI Ore Deposits 16

X 14

- • - Analytical value

12

~

--X-- Calculated value

.~ ~o

.....

o

£

• x. , -~~ . ~

,;o

,;o

=;o

=;o

aoo

Temp.(*(;)

Figure 1.136. 3180 change due to the mixing of hydrothermal solution and groundwater (Shikazono et al., 2002). between the model calculation and analytical data on 8180 except for the data on the most peripheral zone, suggesting the mixing of hydrothermal solution and groundwater is important for causing of 8180 variation in hydrothermally altered andesite.

~34S of sulfides. 834S values of sulfides in the veins are -1%o to +2%0 (Shikazono, unpublished). This is very close to average 834S value of Japanese epithermal A u - A g vein-type deposits which is ca. +2%o. In contrast to the 834S of hydrothermat solution for the vein, that of pyrite in hydrothermally altered rocks (Shimanto Shale) varies very widely, ranging from -5%o to +15%o. Based on the microscopic observation, pyrite with low 834S values less than 0%o is usually framboidal in form, suggesting that low 834S was caused by bacterial reduction of seawater sulfate. There are two possible interpretations of high 834S values (+10%o to 4-15%o). One is the reduction of seawater sulfate in a relatively closed system. The other one is a contribution of volcanic SO2 gas. As noted already, volcanic SO2 gas interacts with H 2 0 to form H2SO4 and H2S. 834S value of SO 2- formed by this reaction is generally high although the 834S value depends on the sulfur isotopic fractionation between H2SO4 and H2S and temperature. It is worth noting that in both cases a considerable degree of sulfate reduction took place. This may be due to the small water/rock ratio in the altered rock system compared to open crack systems through which large amounts of hydrothermal solution migrate. Bulk compositions of altered andesite. Bulk compositions of altered andesite were obtained by XRF analysis (X-ray fluorescence analysis) (Shikazono et al., 2002). Figures 1.137 and 1.138 show the change in elemental contents (SiO2, K20, CaO, MgO) of altered andesitic rocks away from the quartz vein system. Figure 1.139 shows the relationships between (CaO + Na20) content and K 2 0 content. These data indicate the following features of compositional variation of altered rocks.

I'o

90

B

A

8o

°l l • 91i !

8t-

70

~5o

A

•

• •

.,, "

~8 •

•

•

••

~_ _

.

7t-

0 • 60 i

C

5F

•

i

.0_40

,, • •

.

4!

•

1

30

4

3 i-

•

20

2["

•

2

•

•

•

•

i

A

•

• ti

10 0

240 4()0 660 D i s t a n c e l r o m the vein (m)

800

00

200 400 6()0 800 D i s t a n c e from the vein (m)

0 oi

260

460

600

D i s t a n c e f r o m the vein (m)

~00

Figure 1.137, The variations of major element contents in andesite (drilling core, 8-MAHAK-4 and underground samples) away from the vein system (Shikazaoo et al., 2002), Diamond: Hishikari Lower Andesitic tuff (underground samples): square: Hishikari Lower Andesitic tuff (drilling core samples); triangle: Hishikari Lower Andesitic lava (drilling core samples); ×: relatively fresh Hishikari Lower Andesitic lava (drilling core samples). (A) SiO2 content variation. (B) K20 content variation. (C) CaO content variation.

193

Miocene-Pliocene Hydrothermal Ore Deposits 4

3.5 3 o~ 2.5

*

|

2*

4'

•

0

if2 1.5 1 0.5 0

0

Distance from the vein (m)

Figure 1.138. MgO content variation for underground samples. Diamond: Yusen No. 7 vein; square: Seisen No. 8 vein.

14

12

10

1(

"x ×

•

~8 o Abbreviations: A zone I, × zone II, , zone III, • zone tV, + z o n e IV ( n e a r the quartz vein), • fresh andesite (Hishikari Lower Andesite)

ll•

Z +

O6 o

AA

X &

2

• • • ÷

o

0

÷

•

+%

÷ #

÷

04.+~

++'~÷

5 K20 (wt%)lO

÷

15

Figure 1.139. The relationship between K20 content and (Na20 + C a • ) content in andesite (Shikazono et al., 2002).

194

Chapter I

(1) SiO2 content of andesite tends to be high near the veins. High SiO2 content may be due to the presence of quartz veinlets (Fig. 1.137A). (2) K20 and MgO contents of altered andesite decreases away from the vein (Fig. 1.137B and Fig. 1.138). (3) Analytical data on andesite are plotted on (CaO 4- N a 2 0 ) - K 2 0 diagram (Fig. 1.139). This shows that (CaO 4- Na20) content inversely correlates with K20 content. These results are consistent with XRD (X-ray diffraction) results. The amounts of K-feldspar, K-mica and chlorite are higher in the altered rocks closer to the veins and Ca-zeolites and smectite decrease in amounts towards periphery of the alteration zones. Although data are scattered and mostly no systematic correlation of Na20, CaO and MgO contents with 8t80 exist, K 2 0 content and 8~80 seem negatively correlated; K20 content is low and high for zone I and IV and 8180 is high and low for zone ! and IV, respectively. This correlation suggests that mineralogical alteration zoning correlates to 8180 variation that has been cited by Naito et al. (1993). The variations in K20, Na20 and CaO contents could be interpreted by thermodynamic consideration. The dependence of concentration of K +, Na +, Ca 2+ and H4SiO4 in equilibrium with common alteration minerals (K-feldspar, Na-feldspar, quartz) on temperature is shown in Fig. 1.140 (Shikazono, 1988b). This figure demonstrates that (1) chemical compositions of hydrothermal solution depend on alteration minerals, temperature and chloride concentration, and K + and H 4 S i O 4 concentrations increase and Ca 2+ concentration decrease with increasing of temperature. In this case, it is considered that potassic alteration adjacent to the gold-quartz veins occurs when hydrothermal solution initially in

B

Na+

H

E ..go

Ca2+

"

.A N

-1

-2

-3

1;0

200

2;0

Temp. (°C)

300

Figure 1.140. The dependence of concentration of K +, Na +, Ca 2+ and H4SiO4 in equilibrium with common alteration minerals (K-feldspar, Na-feldspar, quartz) with temperature (Shikazono, 1988b). Thermochemical data used for the calculations are from Helgeson (1969). Calculation method is given in Shikazono (1978a). Chloride concentration in hydrothermal solution is assumed to be 1 mol/kg H20. A-B: Na + concentration in solution in equilibrium with low albite and adularia. C-D: K + concentration in solution in equilibrium with low albite and adularia. E-F: H4SiO4 concentration in solution in equilibrium with quartz. G-H: Ca 2+ concentration in solution in equilibrium with low albite and anorthite.

195

Miocene-Pliocene Hydrothermal Ore Deposits

Total mKCJmHCI

600 0 500 o

400

Q. 300 E 200

Andalusite~ ~. ..... ~

Pyrophylli~__ ~ ~

H~A

Kaolinite

100

K-feldspar

Acidic GW 0

2 4 Total mKcJmHcI

GW I 6

Figure 1.141. Temperalure-aK+/aH+ (a: activity) trend due to the mixing of hydrothermal solution and groundwater accompanied by hydrothermal alteration (Hemley and Jones, 1964). HS: hydrothermal solution, GW: groundwater, A: hydrothermaIsolution-groundwater mixing line, B: hydrothermal solution acidic groundwater mixing line (Shikazonoet al., 2002). equilibrium with propylitic alteration minerals (albite, K-feldspar, quartz) ascends rapidly and interacts with country rocks at lower temperature ( I - J - K in Fig. 1.140). It is also considered that K + is added from hydrothermal solution to the rocks as K-mineral such as K-feldspar and K-mica at the site of ore deposition, accompanied by the destruction of plagioclase in the country rocks and liberation of Ca 2+ and Na + to the fluid. It is expected that SiO2 content of the country rocks increases with proceeding of the alteration because solubility of SiO2 decreases with a decrease in temperature ( O - P - Q in Fig. 1.140). Figure 1.141 shows that the hydrothermal alteration zoning from K-feldspar through K-mica to kaolinite from the vein towards marginal part of alteration zones can be explained by a combined change of temperature and m~+/mH+ (m: molal concentration) ratio. Decreases in temperature and mK+/mH+ ratio of hydrothermal solution cause the above alteration zoning. This trend suggests that ascending hydrothermal solution mixed with acidic (low pH) and low-temperature descending groundwater (B in Fig. 1.141). Acidic groundwater may be formed by the interaction of groundwater with acidic alteration zone at shallow part, input of volcanic SO2 gas into groundwater, condensation of steam, or oxidation of H2S. If K + is added to the rock accompanied by the destruction of feldspar, the following reactions occur. CaO (Ca-feldspar) + 2 K + ~ K 2 0 (K-feldspar) + Ca 2+

(1-56)

Na20 (Na-feldspar) ÷ 2 K + --+ K 2 0 (K-feldspar) + 2 Na +

(1-57)

It is expected from (1-56) and (1-57) that CaO and Na20 contents of altered rocks inversely correlate with K20 content of altered rocks with a negative slope of - 1 . Analytical results show the negative correlation on (CaO + N a 2 0 ) - K 2 0 diagram

196

Chapter 1

(Fig. 1.139). However, the low CaO 4- Na20 content and K20 content data on Fig. 1.139 cannot be explained only by these reactions. The relatively low contents data could be attributed to the dissolution of silicates by acidic solution which was generated by the disproportionation reaction, 4SO2 (volcanic gas) + 4 H 2 0 ~ 6H + + 3 SO 2- + HzS

(1-58)

The following dissolution of silicates such as feldspar occur by the interaction of H + with feldspar, CaO (Ca-feldspar) + 2H + ~ Ca 2+ 4- H20

(1-59)

Na20 (Na-feldspar) 4- 2H + ~ 2Na + 4- H20

(1-60)

K20 (K-feldspar) 4- 2H +

(1-61)

--+ 2K + 4- H20

Therefore, it is thought that the mixing of acidic solution with hydrothermal solution occurred and andesite near the gold-quartz veins suffered superimposed potassic and advanced argillic alterations.

1.4.6.5. Interpretation of Si02 mineral zoning in terms of kinetics-fluidflow mixing model

Shikazono et al. (2002) considered the depositional mechanism of quartz and cristobalite and the change in silica concentration of fluid migrating through the altered rocks in the Hishikari mine district based on kinetics-fluid flow mixing model. Their discussion is summarized below. Hydrothermal solution containing appreciable amounts of dissolved silica migrates through andesitic volcanic rocks, accompanying SiO2 precipitation. Figure 1.142 shows the dependence of solubility of SiO2 minerals (quartz, cristobalite) on temperature. As described already, cristobalite occurs in peripheral and shallower part of hydrothermal alteration zone. Quartz is present in zones occurring in deeper and closer to the gold-quartz veins. Such zoning from quartz to cristobalite is also common in main active geothermal systems (Hayashi, 1973; Takeno et al., 2000). Precipitations of quartz and cristobalite occur due to a decrease in temperature that is caused by heat conduction, and fluid mixing. The changes in the concentration of dissolved silica during these processes are shown in Fig. 1.142. Decrease in temperature due to heat conduction alone cannot explain the distributions of quartz and cristobalite. The temperature at which the heat conduction trend crosses the cristobalite saturation curve is ca. 200°C, which is higher than the 100°C corresponding to the cristobalite/quartz boundary in active geothermal system (Hayashi, 1973; Takeno et al., 2000). The curve for the mixing of hydrothermal solution and groundwater always lies below the cristobalite saturation curve if the hydrothermal solution is in equilibrium with quartz. Therefore, the heat conduction and mixing of fluids in equilibrium with quartz are considered to be not main causes for the precipitation of quartz and cristobalite. Therefore, in order to know the change in dissolved silica concentration and temperature during the precipitation of quartz and cristobalite and mixing of fluids, the following equation could be used:

197

Miocene-Pliocene Hydrothermal Ore Deposits 800

800i

i (B)

(A)

700

o

600

[3

~

0 H.S.--

500

400

g o5

8 O300

E~ 0

6OO

[]

500

[] H.ST"-

4O0

300' 200 i

200 100

700!

[]

~.~.

/

._~ ~ 2

~

I

~

100 150 Temp (°C)

~/~

[3

200

250

0.1

A / M ~ 1.00

l°° i ~ D D - ^ A

50

o

%ee~%5

AA-

~

550

Temp (°C)

200

250

Figure 1.142. The computed result of the relationship between dissolved silica (H4SiO4) concentration of mixed fluid and temperature based on four reservoirs model (Shikazono et aI., 2002). Open triangle: solubility curve for quartz, Open square: solubility curve for c~-cristabalite,Solid triangle: Hishikari Lower Andesite lava (drilling core), Cross: Relatively fresh Hishikari Lower Andesite lava (drilling core). H.S.: hydrothermal solution; G.W.:ground water.

dC/dt = k(A/M)(Co - C) ÷ ( q l C 1 / V ) + (q2C2/V) - (ql + q2) C~ V

(1-62)

where C = concentration o f dissolved silica of output fluid ( m o l / k g H 2 0 ) , t = time (s), k = rate constant (mol m - 2 s 1), Co = concentration of dissolved silica saturated with respect to quartz ( m o l / k g H 2 0 ) , C1 = concentration of H4SiO4 o f input hydrothermal solution ( m o l / k g H 2 0 ) , C2 = concentration of H4SiO4 of input groundwater ( m o l / k g H 2 0 ) , ql = volume flow rate of hydrothermal solution (m 3/s), q2 = volume flow rate of groundwater (m3/s), M = mass of aqueous solution in a system (kg), V = volume of aqueous solution in a system (m3), and A = surface area of rocks which contacts with aqueous solution (m2). If steady state is attained, we obtain,

C = {k(a/M)Co + (ql C1 + q2C2/V } / { k ( a / M ) + (ql + q2)/V }

(1-63)

As shown in Fig. 1.143, the system is divided into four reservoirs. Each reservoir corresponding to alteration zone IV, III, II and I is assumed to be homogeneous with respect to temperature and concentrations of dissolved silica in aqueous solution. The temperature of the initial hydrothermal solution is assumed to be 250°C from homogenization temperature of fluid inclusions in vein quartz (Shikazono and Nagayama, 1993). Temperature o f each reservoir was estimated from the assemblage of hydrothermal alteration minerals and temperature of alteration zone in active geothermal system (e.g., Hayashi, 1973; Takeno et al., 2000).

198

Chapter 1 G°W.

H.S. ~ T ~250~

T ----220

T ~150

[ T ~-100

T =25

Figure 1.143. Four reservoirs model. H.S.: hydrothermal solution, G.W.: groundwater (Shikazono et at., 2002).

Volume flow rates of incoming hydrothermal solution and groundwater to each reservoir were estimated from the temperature of reservoirs and initial hydrothermal solution (250°C) and groundwater (25°C) Volume of solution (V) in a reservoir is expressed as, V = Vbox~b/100

(1-64)

where Vbox = volume of reservoir as a rock, and q~ = porosity. Volume flow rate (q) is expressed as, q = A'v(a/lO0

(1-65)

where A I = cross-section area of box (m 2) and v = velocity of fluids. It is assumed that the following simple relation is approximately established in a case of two fluids (fluid 1 and fluid 2) mixing. ql 7"i ÷qzT2 = (qI + q z ) T

(1-66)

where T = absolute temperature. The mixing ratio for fluid (R) is approximated as, R = q2/(ql + q2)

(1-67)

The calculations based on four reservoir models were made using equations (162)-(1-67) and precipitation rate constant (k) for SiO2 minerals by Rimstidt and Barnes (1980). Volume of each reservoir was calculated from the volume of each alteration zone and assuming porosity of alteration zone to be in a range of 3-1%. Dissolved silica concentration of groundwater was assumed to be 60 mg/kgH20 that is average dissolved silica concentration of groundwater in andesitic volcanic region in Japan (Shikazono, unpublished). Assuming the ranges of A / M , flow rate of mixed fluid, porosity and giving the precipitation rate of SiO2 minerals (Rimstidt and Barnes, 1980), the relationship between dissolved silica concentration of mixed fluid and temperature was obtained (Fig. 1.142). It was found that the porosity does not change the results of calculations. Figure 1.144 shows the results of calculation based on multireservoirs (40 reservoirs) model in which each reservoir corresponding to each alteration zone is divided into

199

Miocene-Pliocene Hydrothermal Ore Deposits 800

800

700

[2

600

C] H.S.-- ,

~ 400

[]

500

[] H.S.~

400

[3 °

8

E3 E3

600

[]

5O0

G

700 :

[]

oo

G

b5 3OO

300 A/M~0.0

V = l O -3 ~

200

100200 G

-4¢dM~1.00

1O0 50

100

150

Temp (°C)

200

250

~

50

I

100 150 Temp (°C)

__10~.2

°~'a

200

250

Figure t. 144. Computedresults for multireservoirs (40) model. Open triangle: solubility curve for quartz, Open square: solubility curve for c~-cristabalite, Cross: no precipitation, Open circle: computedresult (Shikazono et al., 2002). ten reservoirs and temperature of each reservoir was given. The results of calculation indicate that the flow rate, 10.4.2 m/s, is the best estimate for A/M = 0.1 that is estimated from the width of quartz veinlets in altered andesite, and this rate is in agreement with the flow rate of geothermal water in active geothermal system ( 1 0 - 6 - t 0 .4 m/s) (Fujimoto, 1987). In this model, the area considered is (2 km x 0.5 km) and flow rate is 10 -4.2 m/s. When porosity is 2%, total mass flow rate is estimated as 10 42 m x 106 m 2 × 2/100 = 1.4 × 106 g/s. This is very similar to that in the Wairakei geothermal area in New Zealand which is 1.3 x 106 g/s (Elder, 1966).

1.4.6.6. Goldprecipitation due to mixing of fluids in epithermal system The mineralogy of hydrothermal alteration zoning, bulk compositional variation of altered rocks, thermodynamic consideration, mass transfer, and oxygen isotope computations, and the sulfur isotope study mentioned above all suggest that the mixing of two fluids (hydrothermal solution and acid groundwater) is the main cause for the geochemical and mineralogical variations in the Hishikari gold mine district. A large number of studies on the depositional mechanism of gold in epithermal system have been carried out (e.g., Shikazono, 1974a, 1986; Shikazono et al., 1990, 1993; Drummond, 1981; Reed and Spycher, 1985; Spycher and Reed, 1989). Depositional mechanism of gold in the Hishikari deposit has been discussed by several investigators (Izawa, 1988; Izawa et al., 1990; Shikazono and Nagayama, 1993; Nagayama, 1993b; Hayashi et al., 2000a,b). For example, Izawa (1988) and Izawa et al. (1990) thought that

200

Chapter 1

the mixing of ore fluids with groundwater, boiling and oxidation of ore fluids due to the interaction of ore fluids with oxidized hematite-rich paleosol are the main causes for the deposition and enrichment of gold in the veins. Hayashi et al. (2000a,b) suggested from a drastic change in 3180 of quartz at the precipitation sequence that electrum precipitated due to the mixing of fluids. The previous studies clearly demonstrated that Au thio complex is dominant among dissolved Au species in ore fluids responsible for Japanese epithermal A u - A g vein-type deposits (e.g., Shikazono, 1974a), considering the estimated f Q - p H range of Japanese epithermal gold deposits and the gold solubility due to thio complex (Seward, 1973, 1981). Thus, accepting the assumption that the Au thio complex is dominant among dissolved Au species in the ore fluids responsible for the Hishikari deposit, the precipitation reaction of gold in electrum is expressed as, Au(HS) 2 + H + + 1/2H2 ~ Au + 2H2S

(1-68)

This reaction suggests that a decrease in H2S concentration, and increases in H + concentration and fH2 (H2 fugacity) and temperature variations are important causes for the deposition of gold in electrum. It is also likely that the deposition of gold in electrum occurs by the following oxidation reaction. Au(HS) 2 + 1 5 / 2 0 2 + 1 / 2 H 2 0 --+ Au+2SO42 + 3 H +

(1-69)

This reaction suggests that oxidation of fluids is important as a depositional mechanism. However, this oxidation reaction seems difficult to explain the gold deposition from the following reasons. (1) The rate of oxidation of H2S to S O ] - is slow at a site of gold deposition. (2) O2 concentration of groundwater is very low. (3) HzS concentration in epithermal ore fluids seems higher than SO 2-. It is also possible that the following reaction is important for the deposition of gold. Au(HS)2 + 8 H20 ~ Au + 2 SO,]- + 3 H + + 15/2 H2

(1-70)

This reaction proceeds due to the degassing of H2, decreasing of H + and probably decreasing of temperature. We cannot evaluate this reaction as an important cause for gold deposition because no study on this reaction has been done. The importance of reaction (1-68) has been already cited by Reed and Spycher (1985) and Shikazono and Nagayama (1993). Shikazono and Nagayama (1993) favored the two fluids mixing as a cause for the gold deposition by the reaction (1-68) from the presence of acidic alteration zone overlying the Hishikari gold-quartz veins. For example, alunite and silicified rocks occur in higher elevations in the north-eastern part of the Hishikari district (Izawa et al., 1990). Shikazono (1985a) suggested based on the studies of wall rock alterations associated with Japanese epithermal-type Au-Ag deposits that acid sulfate fluids coexist with near-neutral chloride-rich fluids at relatively shallow zone from the surface (less than 1 kin) at the time of epithermal-type A u - A g ore formation. Reed and Spycher (1985) made a computation on the gold deposition from the mixed fluids. Their results support that gold deposition occurred in the Hishikari veins

Miocene-Pliocene Hydrothermal Ore Deposits

201

due to the mixing of two fluids (hydrothermal solution and acidic sulfate groundwater). However, it is uncertain that the alteration minerals (K-feldspar, K-mica and kaolinite) formed in the same stage of hydrothermal alteration. It may be possible that descending acid low-temperature solution (groundwater) formed kaolinite and cristobalite at different stage of gold-quartz mineralization associated with K-feldspar precipitation (Hedenquist et al., 1996). If such kind of mixing occurs, pH decreases, H2S concentration decreases, fH2 decreases and temperature decreases. These changes except the decrease of H2 proceed reaction (1-68). Another important depositional mechanism for gold in electrum is boiling of ore fluids, as inferred by many investigators (e.g., Drummond, 1981). Boiling of ore fluid causes an increase in pH. If dominant Ag species is AgCIf, this pH change as well as decreasing temperature, H2 degassing and dilution (decreasing of C1 concentration) causes the deposition of Ag in electrum according to the following reaction. AgC12 + H20 --+ Ag + 2 C1- + H + + 1/2 H2

( 1-71)

Galena, sphalerite and chalcopyrite precipitate also due to increase of pH due to the breakdown of base metal chloro complexes accompanied by the boiling by the reactions such as, PbC12 + H2S --+ PbS + 2C1- + 2H +

(1-72)

ZnC12 + H2S --+ ZnS + 2C1- ÷ 2H +

(1-73)

CuC12 + 2H2S + FeC12 --+ CuFeS2 + 4C1 + 4H +

(1-74)

However, in order to clarify the depositional mechanism of electrum and sulfides, more detailed description of alteration minerals, 3180, 3D data and the salinity (C1concentration)-enthalpy relationship are clearly required, and the two fluids mixing model has to be evaluated based on these data. It is difficult to evaluate that the boiling is an important mechanism for depositions of sulfides and electrum because of scarce data on fluid inclusions (salinity, temperature, enthalpy, 3180 and ~D). However, the above consideration and previous studies on epithermal Au-Ag deposits stressed an importance of boiling for gold depositions (Seward, 1991; Hayashi et al., 2000b). Therefore, it cannot be ruled out that the boiling and H2S loss are main causes for the depositions of electrum and sulfides (spbalerite, galena, and chalcopyrite) in the Hishikari deposit. However, preliminary study on the fluid inclusions indicates that boiling zone and ore zone containing high gold content are different (Etoh et al., 2001). Probably, the mixing of boiled fluid with acid groundwater caused efficient deposition of electrum in the Hishikari hydrothermal system.

1.5. Evolution of tectonics and hydrothermal system associated with epithermal and Kuroko mineralizations Numerous studies on the geologic and tectonic evolution in and around the Japanese Islands from Miocene to present have been carried out (e.g., Kitamura, 1959;

202

Chapter 1

Ozawa, t963; Sugimura and Uyeda, 1973). For example, Sugimura and Uyeda (1973) summarized volcanic rocks, degree of deformation of sediments, structural trends of sediments deposited, and amount of uplift and subsidence since Miocene. However, in contrast to these geologic and tectonic studies, very few studies on the relationship between tectonics and hydrothermal system in Neogene age have been carried out. Therefore, these studies are briefly summarized and then the relationship between geologic and tectonic evolution and evolution of hydrothermal system associated with the mineralizations (Kuroko deposits, epithermal veins) are considered below. 1.5.1. P a l e o g e o g r a p h y and stress field

Geologic environments and tectonic evolution of Green tuff region in Northeast Honshu have been studied by many workers (e.g., Kitamura, 1959). Paleogeography of the Japanese Islands and surrounding areas during a period of middle Miocene age (Kuroko-stage) and late Miocene-Pliocene age (vein-type stage) have been reconstructed in Fig. 1.145, mainly based on the paleontological studies (e.g., Ikebe, 1973, 1978; Chinzei, 1991). According to these studies, Green tuff region of middle Miocene age has been submarine environment in most areas (Fig. 1.145). Rapid subsidence occurred at middle Miocene age, which was contemporaneous with the age of Kuroko mineralization (Yamaji, 1990). From about 5 4-2 Ma uplift of Green tuff region in Inner Honshu province, Northeast Japan took place and this region became subaerial (Sugi et al., 1983; Otsuki, 1989, 1990).

(a)

(b)

o

K T 1 o*

,,¢(~

; @ushO

,.

L.V

0

200km

Land area -, -,~KKfi;hu Y £;k Coastal line *,v ~ ' ~,' (present-day) "~""

~3 Land area I

0

I

O

200km

Coastal line (present-day)

Figure 1.145. (a) Middle Miocene (Kuroko-stage)(solid circle: Kurokodeposits, open circle: Au-Ag vein-type deposits). (b) Late Miocene-Pliocene (vein-stage)(open circle: Au Ag vein-typedeposits) (Shikazono, 1987b).

Miocene-Pliocene Hydrothermal Ore Deposits

203

It is worth noting that the age of uplift is nearly coincident with the age of epithermal vein-type deposits in Northeast Japan including the Chitose, Yatani, Takadama, and Nebazawa Au-Ag vein-type deposits and Hosokura Pb-Zn vein-type deposits. However, the ages of some large base-metal vein-type deposits in Northeast Honshu (Ani, Osarizawa and Nikko Cu deposits) are older than this uplift, suggesting that unconsolidated marine sediments containing interstitial seawater were distributed in the mine area at that time. 334S data on sulfides and sulfates and high salinity of fluid inclusions from these large base-metal vein-type deposits support the above argument. Unfortunately tectonic situations of the regions other than Northeast Honshu of Neogene age are not well understood. However, it seems evident that even in the regions other than Northeast Honshu epithermal Au-Ag vein-type deposits formed when the uplift started and the area of land expanded. In addition to the paleontologic data, the country rocks of epithermal Au-Ag mine districts also suggest that epithermal Au-Ag vein-type deposits have formed under the subaerial condition: welded tuff occasionally occurs in the mine area (e.g., Sado, Nebazawa, Northeast Hokkaido) and in general submarine sedimentary rocks and volcanic rocks are poor or absent in the Au-Ag mine districts (e.g., epithermal Au-Ag vein-type deposits in Kyushu). As already noted, most epithermal Au-Ag vein-type deposits are hosted by young (late Miocene-Pliocene) volcanic rocks and by sedimentary rocks, but dominant host and country rocks for base-metal vein-type deposits are submarine sedimentary and volcanic rocks. Submarine felsic tuft', tuff breccia, dacite lava, intrusive rocks and mudstone are dominant host and country rocks of Kuroko deposits. Detailed studies on the change of stress field during Miocene to present in Northeast Japan have been carried out based on (1) the measurements of direction of dike (Nakamura, 1977; Kobayashi, 1979; Takeuchi, 1980, 1981, 1987), and (2) measurements of direction and age of epithermal vein-type deposits (Horikoshi, 1975b; Otsuki, 1989). Otsuki (1989) recognized the Neogene tectonic stress fields of the Northeast Honshu Arc by analyzing the epithermal veins and showed that in middle to late Miocene, ENE-trending al and a2 coexisted and a3 had a NNW trend, and it was replaced by E - W compression at 7 Ma in Southwest Hokkaido and at 5 Ma in Northeast Honshu. This view is supported by the analysis of trend of dikes (Tsunakawa and Takeuchi, 1986; Takeuchi, 1987) which indicates that the stress field changed from extensional to compressional during 8-6 Ma in southern Northeast Honshu (Figs. 1.146 and 1.147). These analyses of stress field suggest that the dip of subduction of Pacific plate during Miocene might have been steeper than during Plio-Pleistocene considering the mode of subduction of Pacific plate in Northeast Honshu (Niitsuma, 1979; Niitsuma and Akiba, 1984). Usually active plate subduction with gentle dip causes uplift of land and expansion of land area (Uyeda and Kanamori, 1979). Such change from steep subduction to a gentle one caused the changes in site of hydrothermal activity from submarine area during Miocene (Kuroko mineralization) to subaerial area by Plio-Pleistocene to present (epithermal vein-type mineralization). It is also noteworthy that major igneous and hydrothermal activities in the Japanese areas seems likely to have taken place almost at same times as the stress changes. The stress changes occurred at about 22 Ma, 15 Ma, 12 Ma and 8 Ma in the southern part of

204

Chapter I

3)

2)

...........

;~...... ..................... L .....

~

3

Ma

16-12

22~18 N

-- ~fl

i' ' .;! ~

12-8

p , 6~(.~i.;....,,!S T11

6-(0)

0I

(I)

200km 1

TRENDOF DIKES NORMALFAULT •'~ REVERSEFAULT STRIKESLIPFLT. ~.','~,~STRESSTRAJECT.

Figure 1.146. Stress trajectory maps of southern Northeast Honshu in the late Cenozoic period, after Tsunakawa and Takeuchi (1986) with a slight addition. C~H...... trajectory is drawn by smoothing the inferred stress orientations from the selected dike-swarms with K Ar dates. Selected major faults with age estimation are also shown for indicating types of stress fields. T: Extensional stress field, where O'v > O-Hm,x >> OHma., and normal or gravity faulting is preferable. P: CompressionaI. O-H,,,~,~>2> CrHm,,~> OV, reverse or thrust faulting (Takeuchi, 1987).

Northeast Japan and at about 15 Ma in the eastern part of Southwest Japan (Tsunakawa, 1986; Takeuchi, 1987) (Fig. I. 146). The changes in stress fields, and intensities of igneous and hydrothermal activities seem to correlate to oscillatory motion of the Pacific plate (Jackson's episodes) (Jackson et al., 1975; Jackson and Shaw, 1975) (Masuda, 1984). Masuda (1984) and Takeuchi (1987) pointed out that the oscillatory motion of Pacific plate during the least 42 Ma correlates with magmatism, the intensity of tectonism, the change of stress field and the history of sedimentary basin in arc-trench system (Fig. 1.147). The above arguments also suggest that the mineralizations in arc and back-arc systems relate to the oscillatory motion of the Pacific plate.

1.5.2. Volcanic activity The type of volcanic activity in and around the Japanese Islands changed throughout Tertiary. In early Tertiary subaerial andesitic activity was intense. For example, in

205

Miocene-Pliocene Hydrothermal Ore Deposits AGE

EO.

Ma BP

40

OLIGOCENE 30

Hltl

'I'"

I

15

I

10

tl /1"'

"

]

ci

cSua

I

5

1.0 . 5

I

II/ll II

c2

d

~ AVE,AoErH

0

20

I"

~

.~ ,-,

EARLYMIOCENE MIDDLE MIOCENE LATEMIO. PLIOCENEiQUAT.

90t

NO

OF HAWAIIANCHAIN

Episodes

o [ I

Inner Stress Field

r Hmax ORIENTATION

cou,to,~lO~,,i,~

n

T .

"

(~n(P)

r .

"

P transverse

~

i

]

Figure 1.147. Jackson's curve and arc stress reorientations. Apparent swing motion of Pacific Plate (Jackson et al., 1975) and regional stress orientation at the Northeast Honshu convergent margin are illustrated in order to show their synchronous relationship. Dashed line represents the average trend of the Hawaiian volcanic chain. Pacific plate moves along the direction with fluctuation in reference to Hawaii Hot Spot. Vertically shaded parts of the graph indicate the climax phases of "clockwise episodes". Lower part of the figure shows the phases and reversals in orientation of tectonic stress fields on the inner zone of Northeast Honshu Arc (Takeuchi, 1987). Northeast Honshu, subaerial andesitic volcanic activity was dominant at Daijima stage (mostly early Miocene). F r o m middle Miocene b i m o d a l basaltic-dacitic activity started with rapid subsidence in Northeast Japan. The production o f volcanic activity was probably greater than today. Basalt at middle Miocene age erupted at Northeast Japan was studied by Shuto (1989) and Tsuchiya (1988, 1989) who showed that basaltic m a g m a generated in deep mantle. Dudfis et al. (1983) showed that p r e - K u r o k o ore basalt in the Hokuroku district has Mg number ( M g O / ( M g O + FeO*)), where FeO* is total iron content expressed as FeO, in the range 0.85-0.67 4-0.01, suggesting that the basalt is a relatively primitive, unfractionated, and mantle-derived melt. From late Miocene to present, subaerial a r c - v o l c a n i c activity ( c a l c - a l k a l i rocks, andesite, tholeiitic and high alumina basalt) started associated with uplift of the Japanese Islands. This volcanic activity is different from that at middle Miocene age. The above-mentioned changes in paleogeography, volcanism, crustal movement (subsidence and uplift), and stress field clearly demonstrate that these features of back-arc volcanism in e a r l y - m i d d l e Miocene are quite different from those o f Island arc volcanism in late Miocene to present. According to Yoshida and Yamada (2001), the age of change in volcanism from back-arc type to Island arc type in Northeast Honshu was 12.7 M a and this age corresponds to the age of Kuroko formation.

206

Chapter I

1.5.3. Tectonic influence on temporal and spatial relationships in Kuroko and vein-type deposits in southern Hokkaido, Japan It is worth studying Kuroko and vein-type deposits occurring in one metallogenic province. Shikazono and Shimizu (1993) carried out integrated geological, geochemical and mineralogical studies on Kuroko and vein-type deposits in a southwest metallogenic province which is described below. Geology of the province is composed of Paleozoic basements, Tertiary altered submarine volcanic and sedimentary rocks (Green tuff) and Quaternary volcanic rocks. The basements are shale, tuff, limestone and chert of unknown ages. A simplified geologic map is shown in Fig. 1.148. Tertiary rocks are distributed widely. They are composed of alternations of sandstone, mudstone, andesitic and dacitic tuff, tuff breccia and lava. These rocks are intensively and extensively altered and are called as Green tuff. Tertiary volcanic rocks are variable in composition. Andesite, dacite and basalt are found. Quaternary volcanic rocks are dominantly andesite lava and are abundantly distributed in the northern part of the province (Fig. 1.148). The vein-type deposits can be divided into two based on the metals produced; precious (Au, Ag) and base metal (Pb, Zn, Ag, Mn, Cu, Fe) vein-types. There are two sub-types of the base metal vein-type deposits, the C u - P b - Z n sub-type and the P b - Z n - M n - A g sub-type. C u - P b - Z n veins occur in southern part of the province. Large P b - Z n - M n - A g veins and A u - A g veins are distributed in northeastern part. In the northeastern part, A u - A g vein-type deposits occur in marginal zones of the province, while the base metal-rich deposits ( P b - Z n - M n veins and Kuroko deposits) in central zone (Fig. 1.149). The marginal zone is characterized by exposure of Quaternary volcanic rocks and Plio-Pleistocene volcanic rocks in which Au-Ag veins occur, whereas the central zone is by thick submarine volcanic rocks (Fig. 1.150), in which base metal-rich deposits (base metal veins and Kuroko deposits) occur (Fig. 1.150). Tertiary volcanic rocks, Quaternary volcanic rocks and faults are distributed, trending generally from NW to SE. Some C u - P b - Z n veins in southern part are hosted by basement rocks. On the other hand, P b - Z n - M n - A g and Au-Ag veins occur in Tertiary and Quaternary volcanic rocks. Recently K - A t age dating of ore deposits, associated volcanic rocks and plutonic rocks, have been carried out. These data are summarized in Table 1.25. It is obvious that the Kuroko deposits have been formed at middle Miocene, being very similar to the ages of the Kuroko mineralization at Hokuroku district, northeast Honshu where large Kuroko deposits occur. On the other hand, the vein-type deposits were formed during Plio-Pleistocene ages. It is also worthwhile to note that the ages and distribution pattern of the vein-type mineralization in this province are very similar to those of andesite which overlies the vein-type deposits (Watanabe, 1990b; Sawai et al., 1992b). Although ages of volcanic rocks which host the Kuroko deposits have not been determined, it is obvious that the Kuroko mineralization took place at the time of formation of volcanic rocks (mainly dacite). Therefore, it is clear that felsic volcanic activities were related to the mineralizations of the Kuroko deposits, while andesitic volcanism to the vein-type deposits.

Miocene-Pliocene Hydrothermal Ore Deposits 140°E

I

N

207

141,E

Inakuraishi

l

Akaiwa vein

ocene ks

Imai-lshizoki ~km

Basemen[ Rocks

Figure 1.148. Simplified geologic map and distribution of the Kuroko-typeand vein-type deposits in south Hokkaido (Shikazono and Shimizu, 1993). Open circle: precious vein-type deposits. Solid circle: base metal vein-type deposits. Solid square: Kurokodeposits. Minor elements associated with the vein-type and Kuroko deposits are different. Characteristic minor elements concentrated to the ore deposits are Se, Te, Hg, As, Sb and Bi in the A u - A g vein-type deposits, Ag, Bi, As, Sb, Sn, W and Mo in the base metal vein-type deposits, and Au, Ag, Sb, As, Mo and Bi in the Kuroko deposits. This difference in minor elements is consistent with that found in the other epithermal vein-type deposits in Japan (Shikazono and Shimizu, 1992). Analytical results of sulfur isotope previously obtained are summarized in Fig. 1.151.

208

Chapter 1

°; \ o ,"o o Q "

\

\.

/

10oB,o"~,~ s A P P o R o

6 Hg

~

ell O"a~\-

Sn

AgSn"13

~%'Mo \

/~~/

TeSe~~

°

\.

o°

o.o

• CuPbZnMnVeins~ O AuAgVeins • Kurokoor massivebaritedep.

k

',Bi \.

,4"

/:%

•

•

0

J

1119

18

'\ \

\/

2"0 , . - " 1

\

Figure 1.149. Distribution of ore deposits in northeastern part of the province (modified after Yajima, 1979). 1: Suttsu, 2: Kutosan, 3: Toyoha, 4: Akaiwa, 5: Matsukura, 6: Meiji, 7: Todoroki, 8: Teine, 9: Kobetsuzawa, 10: Otoyo, 11: Inatoyo, 12: Toyohiro, 13: Jozankei, 14: Toyotomi, 15:Koryu, 16: Eniwa; 17: Chitose, 18: Shiraoi, 19: Morono, 20: Minami-Shiraoi (Shikazonoand Shimizu, 1993).

It is found that ~348 values for different types of ore deposits are different; 4-1.8%o to 4-5.1%o for the A u - A g vein-type deposits, -2.8%0 to +7.5%o for the base metal veintype deposits and +4%o to 4-5%o for Kuroko deposits. 334S of the base metal vein-type deposits is widely variable, but generally higher than that of the A u - A g vein-type and nearly similar to that of Kuroko deposits. It is worthwhile to note that average ~34S values for the base metal vein-type deposits (except for the deposits in basements), A u - A g metal vein-type and Kuroko deposits in the province are identical to those previously obtained for the other metallogenic provinces in Japan (Shikazono, 1987b). It is also found that 334S values of the ore deposits are related to the host rock types; 334S of ore deposits hosted by basement rocks are relatively low. For example, 334S of the Imaiishizaki and Sasayama is -2.1%o and -0.9%0, respectively. 334S values of relatively large ore deposits such as the Toyoha, Jokoku, and Ohe are high in a range of 4-4%~ to 4-8%0, compared with those of small ore deposits. It is likely that sulfide sulfur in ore fluids responsible for small ore deposits is influenced by the surrounding rocks having low 334S values.

209

Miocene-Pliocene Hydrothermal Ore Deposits 141 °

140 ° 43 °

i~.! ¥ Vv

Quaternary

Pliocene

Miocene

Pre-Tertiary

subaerial volcaniclastic deposits subaerial lava & volcaniclastic rocks coarse laminated sandstone, subaerial to submarine lava & volcaniclastic rocks (Setana E) massive mudstone, submarine lava & volcaniclastic rocks (Kuromatsunai F.) mudstone & shale, submarine lava & volcaniclastic rocks (Yakumo E) granitic rocks sandstone, submarine lavas & volcaniclastic rocks (Kunnui E) mudstone & non marine sediments (Yoshioka F.) subaerial lava & volcaniclastic rocks (Fukuyama F.) ~ basement . . . . . . . rocks

Figure 1.150. Simplified geologic map of northeastern part of the Hokkaido metatlogenic province (modified after Yamagishi, 1989) (Shikazono and Shimizu, 1993).

Available homogenization temperatures of fluid inclusion from the base metal vein-type, A u - A g vein-type, and Kuroko deposits are summarized in Fig. 1.152. Salinity (NaC1 equivalent concentration) of inclusion fluids is 1-6 wt%, 1-14.5 wt% and 0 - 3 wt% for Kuroko deposits, base metal vein-type deposits, and A u - A g vein-type deposits, respectively. These data clearly demonstrate that the salinity of inclusion fluids for the base metal-rich deposits (base metal vein-type deposits, Kuroko deposits) is higher than that of the A u - A g vein-type deposits, while homogenization temperatures of fluid inclusion for these three types of ore deposits do not show a wide

210

Chapter I Precious veins (/3

_>,

c-< o .13

E

Z

-5 -4 -3 -2 -1 0

1 2 3 4

5 6

7 8 9 10

(~ 34 8 ( % 0 )

B a s e metal veins 03 0'1 t~ c.< 0 ..Q

E

Z

, l;>4Xb
t

-5-4-3-2-10

1 234

5 678910

1~ 34 8 ( % 0 )

Kuroko >09, c--

< "5

..Q

E

Z

,

1

L

i

-5-4-3-2-10

~

i

1 2 3 4

56

7 8

910

1~ 34 8 ( % o )

Figure 1.151. Sulfur isotopic compositions of sulfides in the vein-type and Kuroko deposits. Solid box represents sulfur isotopic data from the ore deposits occurring in basement rocks (Shikazono and Shimizu, 1993).

211

Miocene-Pliocene Hydrothermal Ore Deposits

TABLE 1.25 K-Ar age data on the vein-type and Kuroko-type deposits in southern Hokkaido (Shikazono and Shimizu, 1993) Name of mine Vein-type deposits Chitose Koryu Toyoha Ohe lnakuraishi Todoroki Teine Ofukeshi Yakumo Jokoku Shizukari

K-Ar age (Ma) 4.7 3.6 3.5, 3.4, 3.3 1.1 2.2 3.3 3.3, 3.4 4.8, 4.9, 2.8, 2.7 2.1, 3.1 2.9 4.0 2.3 2.3 1.3 5.2 2.4

Name of mine Kuroko-typedeposits Kunitomi Horobetsu Kagenosawa Toya-Takarada Minamishiraoi Minamishiraoi

K-Ar age (Ma) 12.6 12.3 14.2 14.0 13.6 13.6

Disseminated-type deposits Hakuryu 6.5 Date 5.2

variation. Homogenization temperatures for the precious vein-type, base metal vein-type and Kuroko deposits are 180-280°C, 200-250°C, and 180-250°C, respectively. Ishiyama et al. (1987) estimated hydrogen and oxygen isotopic compositions of ore fluid responsible for the base metal vein-type deposits in Jokoku-Katsuraoka area, southwestern part of the province. Estimated 3180 and 3D values range from -49%o to -88%o and from -11.6%o to +5.5%o, respectively. Hattori and Sakai (1979) analyzed inclusion fluids from the Chitose A u - A g vein-type deposits and indicated that 3D and 31SO for the ore fluids are -65%o to -75%o and -4.8%o to -6.2%o, respectively. Interpretation of the geological and geochemical characteristics of the three types of ore deposits are given below. It is inferred that in the northern part of the province submarine volcanic rocks are thick in the central zone, while at marginal zone it is thin and the Plio-Pteistocene subaerial volcanic rocks are exposed. The vein-type deposits occur widely in the province. The precious vein-type deposits occur in relatively young (Plio-Pleistocene) volcanic rocks, while large base metal vein-type deposits (e.g., Toyoha, Inakuraishi, Ohe) and Kuroko deposits (e.g., Kunitomi) occur in central zone where thick Miocene submarine volcanic rocks are distributed (Figs. 1.149 and 1.150). Small base metal vein-type deposits occur in Paleozoic rocks in the southern part. 3348 of the vein-type deposits hosted by sedimentary rocks of the basement are low (less than 0%o), reflecting low 3348 of country rocks. However, they are scarce in number. 334S of the base metal vein-type deposits and Kuroko deposits are relatively high (average value; +4%0 to +5%o) and most probably influenced by sulfate sulfur of seawater

212

Chapter 1

15

d- 10 0 Z .=_

5

O9

0

2OO

30O

Filling Temp. (~C) Figure 1.152. Salinity (NaCl eq.wt.%) and filling temperatures for the base metal vein-type (solid circle), Kuroko-type (solid square) and precious vein-type deposits (open circle) in southwest Hokkaido (Shikazono and Shimizu, 1993).

trapped in the Green tuff rocks. ~34S of the precious metal vein-type deposits closely associated with Plio-Pleistocene subaerial volcanic rocks have igneous value (ca. +2%o), suggesting that sulfide sulfur was extracted from subaeriat country rocks. The causes for the different site of hydrothermal activity (submarine and subaerial environment) could be considered in terms of tectonic and geologic evolution of this metallogenic province from middle Miocene to Pleistocene. During Miocene age most of this province was in a submarine environment. Violent submarine volcanism (bimodal and basic type) took place at Miocene age in this province. This geologic environment may be related to an extensional stress regime (Uyeda and Kanamori, 1979). The Kuroko deposits have been formed related to this tectonic situation. From late Miocene, uplift took place due to the collision of Pacific plate to North American plate under the Kuril arc and Japanese island arc. Watanabe (1986, 1989, 1990a,b, 1991) studied the vein pattern, the age of veintype deposits and the volcanic rocks in southwest Hokkaido and showed that the major veins such as those at the Toyoha and Chitose have been formed at dextral strike-slip movement of an E - W trend, and those veins are situated at the west-southwest extension of the maximum displaced zone within the dextral shear belt along the Kuril arc. Watanabe (1990b) also showed that the veins in the Sapporo-Iwanai district strike E - W and are oblique to the N W - S E volcanic chains which are sub-parallel to the maximum principal stress estimated in southwest Hokkaido during Late Miocene to Holocene and oblique subduction of Pacific Plate was active during the Plio-Pleistocene age. Yamagishi and Watanabe (1986) studied the geologic faults, the dykes, the veintype deposits and active faults in the rocks of the middle Miocene to Quaternary in

Miocene-Pliocene Hydrothermal Ore Deposits

213

this province and clarified that in the middle to late Miocene, southwest Hokkaido was characterized by a tensional stress field of an E - W to N W - S E direction, while in the late Miocene to Quaternary this area was in a compressional stress field. They showed that the main ore veins are recognized as sheared fractures arranged mostly in an E - W direction indicating the signal trajectory lines of an E - W to NW direction. It is noteworthy that the stress field of N W - S E direction estimated from a province of dikes during Miocene is parallel to the distribution trend of the Kuroko deposits in northeastern part of the district. Yamagishi and Watanabe (1986) suggest that the dip of subduction of Pacific plate during Miocene might have been steeper than during Plio-Pleistocene, considering the mode of subduction of Pacific plate in Northwest Honshu (Niitsuma, 1979). Usually active plate subduction with gentle dip causes uplift of land and expansion of land area (Uyeda and Kanamori, 1979). Such change from steep subduction to gentle one caused the changes in site of hydrothermal activity from submarine area during Miocene to subaerial area by Plio-Pleistocene to present. This change from extensional stress regime to compressional regime may correspond to the change of mode of subduction from Mariana-type to Chilean-type as defined by Uyeda and Kanamori (1979). It has been pointed out that this change has occurred at about 5 Ma in Northeast Honshu (Sugi et al. 1983; Ohmoto et al., 1983), and about 7 Ma in southwest Hokkaido (Otsuki, 1989). 1.5.4. Geochemical features of sedimentary rocks formed in the Japan Sea as a proxy for hydrothermal activity As noted already, intense submarine hydrothermal activity took place in the Japan Sea in 15-12 Ma, associated with Kuroko mineralization. However, it is uncertain that submarine hydrothermal activities associated with the Kuroko mineralization took place in the other periods from middle Miocene to present in the Japan Sea. Therefore, the geochemical features of sedimentary rocks which formed from the Japan Sea at these ages have been studied by the author because they are better indicator of age of hydrothermal activities than those of hydrothermally altered igneous rocks because the samples of continuous age of sedimentation are able to be collected and the ages are precisely determined based on microfossil data (foraminiferal, radioralian and diatom assemblages). Thick sedimentary pile from middle Miocene to late Pliocene is exposed in the Oga Peninsula, northern Honshu, Japan (Fig. 1.153). Age of the sedimentary rocks has been determined by microfossil data. Thus, the sedimentary rocks in the Oga Peninsula where type localities of Miocene sedimentary rocks in northern Japan are well exposed have been studied to elucidate the paleoenvironmental change of the Japan Sea (Watanabe et al., 1994a,b). Kimura (1998) obtained geochemical features of these rocks (isotopic and chemical compositions) and found that regional tectonics (uplift of Himalayan and Tibetan region) affect paleo-oceanic environment (oxidation-reduction condition, biogenic productivity). However, in their studies, no detailed discussions on the causes for the intensity and periodicity of hydrothermal activity, and temporal relationship between hydrothermal activity, volcanism and tectonics in the Japan Sea area were discussed. They considered only the time range from ca. 14 Ma to ca. 5 Ma.

214

Chapter 1

NYUDOZAKI

NISHIKUROSAWA ~

NOMI/RA

i

K1TAURA AIKAWA

SARUgAWA

ANDEN

L"

' RECLAIMED

LAND

TOGABAY~_I~j ~NINIOME.G ..A. .TAICHLNOMEG . ATA S A N ~ O M E -GATA

KAMO Mr, K E N A S H I Y A M A JAPAN SEA

MONZEN

5kin

['~

~

SUGOROKU

MINAMIH IRt..qAWA 0NNAGAWA

Kampu-zanVOlCaNOs

Megata volcanic ejecta

~ KatanishiFormation ~ ~'/tJ Togapumicebed i

FunakawaFormation OnnagawaFormation NishikurosawaFormation

~rl~ ShibikawaFormation ~

DaijlmaFormation

~

WakimotoFormation ~

MonzenFormation

Kitaura Formation

AkashimaFomation

~

Figure 1.153. Geologic sketch map of the Oga peninsula.

Therefore, the wider time range from middle Miocene to present is considered below based on available age data on hydrothermal ore deposits (Kuroko deposits, epithermal vein deposits) and hydrothermal alteration in the mine areas in Northeast Japan.

Miocene-Pliocene Hydrothermal Ore Deposits

215

The geochemical features of the sedimentary rocks in the Oga Peninsula and the hydrothermal activity in Japan Sea deduced from these features are described below. The area is located at the west of Oga Peninsula (Fig. 1.153) and is composed of Miocene sedimentary rocks (Nishikurosawa, Onnagawa and Funakawa Formations). The Nishikurosawa Formation is composed of siltstone, mudstone, conglomerate and sandstone. Siltstone and mudstone contain foraminiferal fossil such as GIoborotalia birnageae, and G. denseconnexa, indicating Zone N. 9 by Blow (1969). The upper part is characterized by glauconite-bearing sedimentary rock. The total thickness is about 150 m. The Onnagawa Formation conformably overlies the Nishikurosawa Formation and is composed of siliceous shale and shale. The rocks are characterized by organicrich laminated diatomaceous deposits, siliceous microfossils, and fish bones, while foraminiferal fossil is poor in amounts. The total thickness is about 300 m. The age of base of the Onnagawa Formation is estimated to be 12.9 Ma based on diatoms (Koizumi and Matoba, 1989). The age of the top of the Formation is 5.8 Ma. The Funakawa Formation conformably overlies the Onnagawa Formation. The thickness is about 1000 m. The Formation is composed of siltstone, intercalated by tuff and tuffaceous siltstone. Foraminiferat fossil and siliceous microfossil exist in the formation. Average analytical data on the sedimentary rocks are compared with NASC (North American Shale Composite; Gromet et al., 1984) in Fig. 1.154. Average chemical compositions of the Nishikurosawa are very similar to NASC. However, Fe203, MgO, TiO2, MnO and P205 contents are higher and K20 content is lower than those of NASC, respectively. SiO2 and A1203 contents are similar with each other. The Onnagawa Formation is characterized by very high SiO2 content (average 89.12 wt%). Most of the other major element contents are lower than those of NASC, although MnO and P205 contents of the Onnagawa are higher. Average major element contents except SiO2 content of the Funakawa are very similar to NASC, but SiO2, Na20 and MnO contents are higher and the other element contents are lower than the NASC. Analytical data on minor element contents are compared with Average Shale by Turekian (1972) (Fig. 1.155). The average contents of most of the elements except Zn and Ba of the Nishikurosawa are same in orders to NASC. Average contents of minor elements of the Onnagawa are mostly lower than Average Shale. However, Cu, Zn, Mo, Ba and U contents are anomalously high compared with Average Shale. Average contents of minor elements of the Funakawa are lower than NASC which is due to slightly higher SiO2 content of the Funakawa. REE pattern normalized by NASC (REE contents; Goldstein and Jacobsen, 1988) are shown in Fig. 1.156. Nishikurosawa Formation is characterized by Eu positive anomaly which is calculated by the following equation: Eu/Eu* = (2Eu/Eu~ASC)/(Sm/SmNASC -k-Gd/GdNASC)

(1-75)

Chapter 1

216

I00

(wI%}

~A)

8

N/shik~t,

shale

rasa~;,~

~g 40 NK¢-~eries s v e r a z e

2{) l{

-

@rome~. e~

~L 1984]

t

:
.---

10E

(B) O n n a g a w a harod s h a l e

......~

OO-ser es @r~met

\

N

80

4O

~ver~e

et d

.

i9~

\

\

-

\

\

I

2

Miocene-Pliocene Hydrothermal Ore Deposits

217

4~ 'i } t{

k

Figure 1.154. Major element composition of studied rock samples. (A) Nishikurosawa shale, (B) Onnagawa hard shale, (C) Funakawa shale. Shaded area represents ranges.

The REE pattern for the Funakawa plot within 0.1-1 of NASC normalized value (Fig. 1.156). Ce and Eu anomaly for the Funakawa are weak (Ce/Ce* = 1.20 4-0.01; Eu/Eu* = 1.09 4- 0.05) (Fig. 1.157). The REE pattern for the Nishikurosawa exhibits slightly light REE enriched one (La/Yb = 0.55 4- 0.09) and plots close to the line of NASC normalized value -----1. Positive Ce anomaly is also found in the Nishikurosawa. This anomaly is defined by the following equation. Ce/Ce* = (3 Ce/CeNAS¢)/(2La/LaNASC + Nd/NdNAsC)

(1-76)

Hydrothermal solution venting from midocean ridges and back-arc basins has positive Eu anomaly (Klinkhammer et al., 1983; Michard et al., 1983; Mitra, 1994; Shikazono, 1999a) (Fig. 1.158). Therefore, the positive Eu anomaly of the sedimentary rocks is thought to be due to a contribution of hydrothermal solution. In order to know the contribution of hydrothermal solution the positive Eu anomaly of seawater (Eu/Eus*awater) is useful.

Chapter 1

218

' ppm~ f1

C

i

]ktrcB~ 1~;2]

i

I

[

~y °

~0~

m a h a ~ d sAale

fB~ 0 7 ~

i0/

° b" ......

%#

{ J

V

~n

I

,¢

{ h

~r ~e~

Zr y

?~ Nb

A ..........

Hf

R

Ba

Ph %~

U

219

Miocene Pliocene Hydrothermal Ore Deposits (ppm) }£

[ ........ r - - - 7 - - 7 - - - ]

.............................................................

~ z t ;k~ a~

"

~

...........

192°9.

i

[ I} ~--

~

ot

Figure 1.155. Trace element composition of studied rock samples. (A) Nishikurosawa shale, (B) Onnagawa hard shale, (C) Funakawa shale. Shaded area represents ranges. The anomaly at that time is estimated based on the following equation. gu/gUr*oc k = n (gu/gut%rrigeneous) -Jr (1 - n)(gu/gus%awater )

(1-77)

where n is proportion of terrigeneous component in sediments, Eu/Eur*ck is Eu positive anomaly of sedimentary rocks, and Eu/Eut*rrigeneous is positive Eu anomaly of terrigeneous component. Eu/Eus*awater estimated based on the above equation is shown in Fig. 1.159. It is evident in this figure that intense hydrothermal activity occurred in the Nishikurosawa stage (14-13 Ma), it was very high " (Eu/Euseawater * -- 2.4) at the Nishikurosawa/Onnagawa boundary (12.6 Ma) which is similar to that of the sediments below the hydrothermal pool of Red Sea (Fish debris, Eu/Eu* = 2.4-4.9; Oudin and Cocherie, 1988), the intensity of hydrothermal activity decreased in the Onnagawa stage towards the Funakawa stage, and intense hydrothermal activity occurred intermittently at ca. 12.6, ca. 10.5 and ca. 8.2 Ma. High Zn, Cu and Ba contents of the Nishikurosawa Formation also indicate that the intense hydrothermal activity occurred at these ages. Magnetic susceptibility data are inferred to have been reflected by hydrothermal activity. The magnetic susceptibility data on the sedimentary rocks are shown in

220

Chapter 1 10

Funakawa shale

0.1

0,01

l

i

l

i

l

l

l

l

l

l

l

l

l

l

10 Onnagawa hardshale N

o~

@

=, 0.1

r/3

.< Z

10

l

l

l

f

i

l

i

l

l

l

l

I

0,01

Nishikurosawa shale

O.I

(1,(11

I

La Ce

I

t

I

I

v

NdSmEuGd

,

~

Dy

I

i

Er

I

E

YbLu

Figure 1.156. NASC-normalizedREE patterns. REE data of NASC obtained by Goldsteinand Jacobsen (1988) are used as normalizationvalue.

Fig. 1.160. The magnetic susceptibility values of the Nishikurosawa sedimentary rocks are high but those of the Onnagawa are relatively low. The values of the Funakawa are lower than those of the Nishikurosawa but higher than those of the Onnagawa. The values around the Nishikurosawa/Onnagawa boundary are the highest and they are variable in the Nishikurosawa and in the Onnagawa. The values of the Onnagawa are relatively high at 14-13 Ma, 12.6 Ma, 10.5 Ma, and 8.2 Ma. This variation correlates to Eu/Eus*awater

221

Miocene-Pliocene Hydrothermal Ore Deposits 1.4 1.3

m

Eu/Eu, 1 J:o

o

ml

|E

o

E

1.2

o

0

o

m

® B

o

0

o

1.1 o no a n o m a l y ~

1.0

e! a

o o

o

O e

Pl )sitive negative

. . . . . . . . . . . . . . . . . . . . .

0.9 NK-shale

0t-t3

OG-hard

! .... l iFK-shale 01-04

shale 01-21

Figure 1.157. Eu/Eu* values of studied rock samples. I

w -E

0.01

I

I

I

I

Z

I

~ 13*N

-~

o

J

~.

O

1

:

_

#.°

-

oc-

-

E

-

(xlO-3t

\

_

1984

13°N F982

m 2 1 * N 1981

~ \

o,oo

J

/\

J

\

/

Cl~

0 0.0001

-

-

-

i

Ce

t

Nd

I

I

L

Sm Eu Gd

L

Oy

I

Er

I

Yb

Figure 1.158. Rare earth elements in vent fluids from the East Pacific Rise (redrawn from Michard and Albarede, 1986). The Sm values for the 1984 I3°N set are thought to be somewhat too high because of contaminationfrom the samplingsyringes.Data for basalt from 13°N East Pacific Rise are from Michard et at. (1983) (Scott, 1997). variation, suggesting that high magnetic susceptibility may be due to high content of magnetite in the sedimentary rocks. Magnetite is considered to have formed biogenetically because magnetite is very fine-grained. If the grain size of magnetite is less than 0.03 Ixm, magnetic susceptibility of sediments containing magnetite is very high (Ioka and Yamazaki, 1994) and magnetite is thought to be of biogenic origin (Yamazaki et al.,

222

Chapter 1

(A)

(B)

- 2.0 " ~ ~ [ ' ~ t ~

......., ~

H ydrothermaI sediment

Seawater

d~tal

* - l.OI--n..... 0.8 I

Modern*

]

[

'v~"~O.~Q~Q " "~O'

aly

negative >12.9

., ,

12

,

11

,

10

,

9

"

% ,

8

7

5.8>

Middle Miocene to early Pliocene

J~l~r'. (Ma)

]

Figure 1.159. Eu/Eu* values of (A) modern sediment, hydrothermal solution and seawater and (B) midMiocene to early Pliocene Japan Sea (gu/Eus*awater, see in text). Modern data are from the Pacific ocean, except fish debris in the hydrothermal sediment (Red Sea). Modern* is defined here as including Quaternary time. Data sources: Hydrothermal solution (Michard and Albarede, 1986), seawater, fish debris in the Atlantic II deep (Oudin and Cocherie, 1988), hydrothermal sediment (Ruhlin and Owen, 1986).

1991). Probably, iron of biogenic magnetite was originated from hydrothermal solution. It is considered that ferric iron of hydrothermal solution was oxidized by iron oxidizing bacteria to form magnetite. The ages of Neogene mineralization and hydrothermal alteration in and around the Northeast Honshu and Hokkaido have been determined by K - A r data on K-minerals (K-feldspar, sericite). These data are summarized in Fig. 1.147 and Table 1.26. It seems clear by comparing Fig. 1.159 with Table 1.26 that the ages of hydrothermal mineralization and alterations determined by K - A r age dating are consistent with those of sedimentary rocks affected by hydrothermal activity in the Oga. Hydrothermal activities were intense at ca. 14-13 Ma, 12.6 Ma, 10.5 Ma, and 8.2 Ma. As already noted, major igneous and hydrothermal activities in the Japanese Islands seems likely to have taken place nearly at same times as the stress changes (Fig. 1.147). The stress field changes occurred at about 22 Ma, 15 Ma, 12 Ma and 8 Ma in the southern part of Northwest Japan and at about 15 Ma in the eastern part of Southwest Japan (Takeuchi, 1987) (Fig. 1.147). Kuroko and vein-type mineralization occurred at 23 Ma (Sado), 15-14 Ma (Kuroko in Hokuroku, Sado), 12.8 Ma (Karuizawa), 8 Ma (Takadama), 5 Ma (Hosokura), 3 Ma (Yatani) (Shikazono, 1985e, 1987b; Shikazono and Tsunakawa, 1982; Sugaki et al., 1986; Sawai and Itaya, 1996; Otsuki, 1989).

223

Miocene-Pliocene Hydrothermal Ore Deposits (10 -3 SI) m!

1.0 t

Magnetic

2

susceptibility

0.8 0.3 Q

m

0.2 []

WW

0.I

0 ® 0

0 0 i

J

t

t

I

i

I

I

I

t

NK-shale 01-13

i

;

I+

I ~79

®®

0

@ ®

I ®I [ I I I I I I l0I T I I I I

OG-hard shale 01-21

0 fin Oi @ II FK-shale 01-04

Figure 1.160. Magnetic susceptibility value of studied rock samples. Most younger two samples (NK-12 and -13) of Nishikurosawa shale have anomalously high magnetic susceptibility value.

The changes in stress fields, and intensities of igneous and hydrothermal activities except 12.6 Ma seem correlate to oscillatory motion of the Pacific plate (Jackson's episodes; Jackson and Shaw, 1975; Jackson et al., 1975) (Masuda, 1984). Masuda (1984) and Takeuchi (1987) pointed out that the oscillatory motion of Pacific plate during the last 42 Ma correlates with magmatism, the intensity of tectonism, the change in stress field and the history of sedimentary basin in arc-trench system (Fig. 1.147). The above argument also suggests that the hydrothermal mineralizations in arc and back arc systems relate to the oscillatory motion of the Pacific plate. It is generally accepted that Kuroko deposits formed under the submarine environment, while polymetallic vein-type deposits in central and Northwest Japan (Ashio, Tsugu, Kishu, Obira, etc.) under the subaeriat environment. This spatial difference in the distribution of back-arc and vein-type deposits is found also for present-day mineralizations. For example, back-arc deposits are forming at Okinawa Trough, while base-metal and precious-metal precipitations are occurring on land such as at Ibusuki and Beppu geothermal area, Kyushu. Recent gold mineralizations (ca. 1 Ma) occurred at Noya and Hishikari areas in southern Kyushu. These mineralizations can be regarded as almost contemporaneous mineralization with the Okinawa mineralization. This kind of temporal and spatial relationship between epithermal Au vein-type mineralization and back-arc mineralization are found also in the Izu-Bonin area. Seafloor

224

Chapter 1

TABLE 1.26 Isotopic ages of metalliferous veins, altered country rocks and the igneous rocks which have genetical relations to the mineralization (Otsuki, 1989) Mine

Age (Ma) Mineral

Sanru Ohe Toyoha Inatoyo-Toyohiro Chitose Horobetsu Kagenosawa Koryu Hakuryu Date Ani Koaizawa Kameyamamori Arakawaohsawa Rata Sugisawa Kakkonda Hosokura Yatani

Takatama Karuizawa Ohizumi Sado

Ashio Nehazawa Chichibu

12.4 3.3 2.2 4.7 3.6 12.3 14.2 1.1 6.5 5.2 11 12.6 8.1 13.4 9.4

5.8 3.3 3.4 3.6 8.4 12.8

Metals

Vein strike

Au.Ag Mn>>Zn-Pb Zn > Pb>>Mn

N70°E, N85W; N15W N55°-85°W N40°-50°W, EW NS N65°E-N80W

Rock 6.6 8.4 8.5 3.4 10.3

12-15 23.6 15.2 8.7 13.6 12.0 4.9 9.7 9 5.2

5.3 13.4 14.5 22.1 24.4 14.8 5.0-5.7

Au-Ag Au-Ag-Cu.Pb.Zn Au.Ag.Cu-Pb.Zn Au-Ag Au.Ag.Cu.Pb.Zn Au-Ag-Cu-Pb.Zn Cu Ag.Cu.Pb.Zn Au.Ag.Cu.Pb-Zn Zn.Pb Au.Ag.Cu.Pb-Zn Au-Ag-Cu-Pb.Zn

N70°E, N55°-80°W N70° 90°E N80°-90°E ENE ENE N23°-70°E, N80°W N50°-60°W N25°W WNW-EW-NS N45°E, EW, NS NW N30°-80°E NNW-NW

Zn.Pb Zn.Pb

ENE, EW, NW EW, NE, NW

Pb.Zn Au-Ag Au.Ag Au.Ag Ag.(Pb-Zn) Zn.Pb.Cu Au.Ag

EW, 70W N70.W-EW-N45E

Cu Au-Ag 6.6

WNW, ENE, NNW N60°W, N60°E EW-ENE + NS + NW N65°-90°E, NW

N50°-70°E, EW N70 °, 85°E NS

b a s e - m e t a l m i n e r a l i z a t i o n s i m i l a r to the K u r o k o m i n e r a l i z a t i o n is f o u n d in I z u - B o n i n s e a f l o o r ( l i z a s a et al., 1999), w h e r e a s e p i t h e r m a l g o l d v e i n - t y p e m i n e r a l i z a t i o n o c c u r r e d at v e r y r e c e n t age (ca. 1 M a ) at Izu P e n i n s u l a , c e n t r a l Japan. T h i s spatial d i f f e r e n c e is c o n s i s t e n t w i t h the d i s t r i b u t i o n w i t h that o f h y d r o t h e r m a l d e p o s i t s o f m i d d l e M i o c e n e ( K u r o k o a n d p o l y m e t a l l i c v e i n - t y p e d e p o s i t s in Japan). I f the s u b m a r i n e vs. s u b a e r i a l h y p o t h e s i s m e n t i o n e d a b o v e is correct, it is e x p e c t e d t h a t K u r o k o d e p o s i t s are f o u n d in the m a r i n e s e d i m e n t a r y a n d v o l c a n i c h o r i z o n s w h o s e

Miocene-Pliocene Hydrothermal Ore Deposits

225

ages are 12.6 Ma, 10.5 Ma, 8.5 Ma and 5 Ma as well as 15-14 Ma (Kuroko deposits in Hokuroku area) and present-day (Okinawa and Izu-Ogasawara deposits). Large Kuroko deposits formed at 15-14 Ma (ore deposits in Hokuroku and Hokkaido) and at present (Okinawa Trough, Izu-Ogasawara). These ages correspond to rapid change of oscillatory motion of Pacific plate (Jackson's episodes). The change in the motion of the Pacific plate occurred at 15-14 Ma, 10.5 Ma, 8 Ma, 6 Ma, 3-2 Ma and probably present (Fig. 1.161). Thus, small Kuroko deposits probably formed at submarine environment at 10.5 Ma, 8 Ma, 6 Ma and 3-2 Ma, although these ore deposits have not been discovered. However, it is noteworthy that the 12.6 Ma corresponding to the Nishikurosawa/Onnagawa boundary when intense igneous and hydrothermal activities occurred does not correlate to Jackson's episode. The geochemical characteristics of the Nishikurosawa/Onnagawa boundary (12.6 Ma) are distinct from those of the other ages affected by hydrothermal activity (15-14 Ma, 10.5 Ma, 8.5 Ma); The positive Eu anomaly and magnetic susceptility were the highest compared with the other ages, glauconite is abundant, paleoenvironment was extremely reducing, and mass extinction occurred. This difference suggests that the cause for the volcanism and hydrothermal activity at that time were different from the other ages. The cause for hydrothermal and igneous activities at 12.6 Ma is uncertain because of no relationship with Jackson's episode. However, it is speculated that bolide impact may affect igneous and hydrothermal activities at this age. In addition to no relationship with Jackson's episode, this speculation comes from that the Earth's surface environmental change at the other geologic boundaries such as Eocene/Oligocene which is considered to have been influenced by bolide impact (Keller, 1986; Montanari et al., 1993; Vohof et al., 2000). Bolide impact may be important. But, further work should be clearly necessary to evaluate this hypothesis which is highly speculative.

1.5.5. Mode of subduetion and formation of back-arc basin

Uyeda and Kanamori (1979) divided mode of subduction into two types: Marianatype characterized steep subduction and Chilean-type characterized by gentle subduction estimated from the dip of Benioff-Wadati zones (Fig. 1.161). The geological phenomena associated with these subductions are shown in Fig. 1.162. It is inferred that the change in mode of subduction from Mariana-type to Chilean-type occurred at ca. 5 Ma in Northeast Honshu. Kuroko deposits are associated with Mariana-type, whereas epithermal vein-type deposits (particularly Au-Ag deposits) with Chilean-type. The above-mentioned view on the evolutionary history in metallogenic provinces from the formation of massive sulfide deposits at early-stage to that of the vein-type deposits in late-stage is consistent with the view by Cathles (1986) who has shown that massive sulfide deposits have been formed at an earlier stage than gold-quartz veins in Archean greenstone belt. Therefore, it is likely that the change in style of mineralization from massive sulfide deposits to gold vein-type deposits caused by the change in mode of subduction is common geologic phenomenon from Archean to Quaternary age in an Island arc-back-arc geologic setting. However, the relationship between the mode of

226

Chapter i . Andaman M. America (12 N), C. Aleution, L. Antitles 100 km -

"~--'~

~

S.Chile(40%) Sumatra(5~S)

/S.Chile (30"S)/

I~~__~.~ru

~,laska,~.

'

New

~

~ ~

~.

j

N.Chile(20°S)

Hebridesd'~ ~k'~~.~ "% ~

.yo.yu.p'%\'<,

'New Zealand

Ik ~

• Solomons

~ ~It~ , ~

Philippine-- - - ~ N e w Britain

~ ~

~

Depth, km 100 200

300

'~ '~

NEJalpan 400 ~-,,~Kurl~,

Java"

Mariana

'

(10°S)

0

'

700

Figure1.161.DipofBenioff-Wadatizones(UyedaandKanamori,1979). subduction and style of mineralization during the geologic history is not clear and should be studied in future. Horikoshi (1975a) proposed the hypothesis that Kuroko deposits formed at the time changing from extensional environment to compressional environment. He thought that metal-bearing fluids in the crust were pushed out by the compression. In contrast, Otsuki (1990) and Yamaji (1990) thought that back-arc rift activity and rapid subsidence occurred at ca. 15 Ma under the extensional environment. Bimodal volcanism was dominant at that time. Therefore, it is inferred that Kuroko formation took place at extensional environment, rather than compressional environment favored by Horikoshi (1975a); formation of Kuroko deposits may be related to the opening of the Japan Sea. However, the relationship between the opening of the Japan Sea and Kuroko mineralization is unclear, and there are many opinions on the genesis of the Japan Sea and timing of opening of the Japan Sea. For instance, Otofuji et al. (1987) and Otofuji (1996) showed that Southwest Japan Sea opened at middle Miocene very rapidly and widely based on paleomagnetic data. However, the rate of opening of the Japan Sea estimated by Otofuji et al. (1987) is too high (60 cm/year) and seems to be unreasonable. Therefore, many workers think that opening of the Japan Sea did not occur very rapidly and widely at middle Miocene age, but it occurred from earlier age (25 Ma) and opening occurred several times by the slower rate (Nakajima and Nakagawa, 1994). Twice openings are favored by several workers (Nakajima and Nakagawa, 1994). It is considered that there were several axes of the opening based on K - A r ages of volcanic rocks (Kobayashi, 1983).

227

Miocene-Pliocene Hydrothermal Ore Deposits

Ce_,

¢

-,,,

d

.,o*o

Mar~aria

"~~~ "

Figure 1.i62. Two kinds of subduction boundaries (Uyedaand Kanamori, 1979).

In contrast to Southwest Japan, opening of northeastern part of the Japan Sea is unclear, compared with the southwestern part of Japan. Tosha and Hamano (1988) made a paleomagnetic study of Tertiary rocks of Oga Peninsula (northern Honshu) and considered that a counterclockwise rotation of Northeast Japan with respect to eastern Asia took place between about 22 Ma and 15 Ma and the before the rotation, Northeast Japan was situated along the east coast of the Asian continent. Lallemand and Jolivet (1986) have interpreted the opening of the Japan Sea as a pull-apart basin between two right-lateral strike-slip fault zones; the Yagsan-Tsushima fault to the west and the Tartary-Hidaka shear zone to the east. Southward migration of Southwest Japan as a drawer between a right-lateral system to the west and a left-lateral one to the east along the Tanakura Tectonic Line (Otsuki and Ehiro, 1978) or a clockwise rotation of Southwest Japan about a pole located in the Tsushima Strait (Otofuji et al., 1987), or a combination of both (Kobayashi, 1983) have been proposed.

228

Chapter 1

Although many discussions on the origin of the Japan Sea have been carried out since the pioneer works by Tokuda (1927), and Terada (1934) who considered that the Japan Sea formed by southward migration from the Asian continent, the problem of the mechanism of the opening of the Japan Sea remains unsolved as mentioned above. Kuroko formation occurred at middle Miocene (15-16 Ma). But the opening of Southwest and Northeast Japan Sea occurred probably from the age earlier than this age. Therefore, it is thought that the beginning of the opening of the Japan Sea did not directly relate to the Kuroko formation. Cathles (1983a) proposed failed rift hypothesis to explain unusual features of the Green tuff region and Kuroko deposits. He mentioned that this hypothesis can account for the distribution of mine districts in the Green tuff region, the observed extensive and substantial premineralization subsidence and postmineralization uplift in the region, and its volcanic evolution. He further pointed out that the Green tuff belt of Japan is strikingly similar in geology, tectonics, and mineral deposits to Archean greenstone belts, suggesting that Archean belts may be failed rifts. As mentioned above, formation of back-arc basins and marginal seas may be important for the formation of Kuroko and vein-type deposits, although genetic relationship between Kuroko formation and opening of the Japan Sea is not clear. For example, Horikoshi (1977) insists that vein-type deposits in Northeast Hokkaido did not form without the opening of Ohotsuku back-arc basin. In section 2.3 and in Chapter 3, it is shown that the formation of back-arc basins take important role for the mineralization (back-arc deposits (Kuroko deposits), epithermal Au veins) and global geochemical cycle. Thus, it must be worth considering the formation mechanism of back-arc basins. The origin of the back-arc basins has been investigated considerably (Karig, 1971; Sleep and Toks6z, 1971; Uyeda and Kanamori, 1979; Tamaki and Honza, 1991; Uyeda, 1991; Tamaki, 1995) and various explanations for the origin have been proposed. Tamaki and Honza (1991) summarized the previously proposed models and the currently plausible models of back-arc spreading (Fig. 1.163). Model 1 shown in Fig. 1.163 is a slab-induced upwelling model. Upwelling generated along the down going slab-mantle boundary (Karig, 1971) or by secondary convection that is introduced by the slab (Sleep and Toks6z, 1971) causes back-arc formation (Karig, 1971 ). Model 2 is a plume injection model. Miyashiro (1986) showed that a northward migration of "hot region" (a large high-temperature region at a deeper level) governed the formation of back-arc basins in the Western Pacific. Tatsumi et al. (1990) examined the distributions of Cenozoic basalts and active rift systems in the Northeast China region and Quaternary subduction unrelated volcanism in the northwestern part of New Zealand (Fig. 1.164) and concluded the following: (1) Injection of the asthenosphere into the mantle wedge and the dam-up effect of the subducted slab explains the rifting process in the Japan Sea (Fig. 1.165). (2) Injection of the asthenosphere was associated with the eastward horizontal convective flow caused by the upwelling of asthenosphere beneath the northeastern China region (Fig. 1.165).

Miocene-PIiocene Hydrothermal Ore Deposits

229 , J / ~ - ""--I

I=IS!~!I, ! i I ~ k ~ I m tb mOdi~ £ 8 (le

Mode 1i Ssb4ii :d c@d pw@t gmsd~i!

XXXX

X

XX

M~de 4. ILIt,q!! £ iO@l£

m~d~l ? I~+I

M s I e 2 P!ume i i ~ t i ~ n msd@

!\

j y' /' / i ( i s<S

Figure 1.163. ModeIs of backarc spreading (Tamaki and Honza, 1991).

(3) The active back-arc extension in the Okinawa Trough and the Taupo Depression in New Zealand can be explained by the injection model. Otofuji (1996) proposed a "double door" opening mode with a fast spreading rate of 21 cm/year for the evolution of the Japan Sea, caused by the injection of asthenosphere with a low viscosity beneath the Japan Sea area. Model 3 is a plate kinematic model. The retreat of a back-arc plate forms a back-arc basin (Dewey, 1980). Model 4 is also a plate kinematic model. The retreat of a fore arc plate forms a back-arc basin. This model seems attractive. Jackson et al. (1975) found the periodicities of rotational motions of the Pacific plate. When the direction of the Pacific plate changed and obliquely subducted, the compressional force of oceanic plate to continental plate decreases. That means that the retreat of fore arc plate occurs.

Chapter 1

230

L NE CHINA UPWELLING 40N

L/ j~jl "~ • ".i I i ,

,

, °

°.

t~,,, e

:.,

•

B I":

140E

20N 120E Figure 1.164. Distribution of Cenozoic basalts and active rift systems in the northeast China region. Arrows indicate the horizontal convective current in the upper mantle associated with the upwelling of the asthenosphere beneath the region. A: Baikal Rift; B: Shanxi Graben; C: Tancheng-Lujiang Fault; D: Okinawa Trough (Tatsumi et aI., 1990).

INTRA-PLATE VOLCANISM CONTINENTAL RIFT SYSTEM P

+

+

÷

•

° E4

+'m,(÷

+'lrff+

l lll {11{

BACKARC BASIN

o,

;

÷

+

VOLCANIC BASIN

. .

°

5rrrvrrrrrrff+

÷

,

.4, y ~ .

~ /~/;~'0"0 Asthen°sPhi re.~q~

Asthenosphere

Upwellingof Asthenosphere ~

f

e,~,~

Figure 1.165. Schematic diagram showing the effect of upwelling of asthenosphere beneath a continent which is adjacent to an arc-trench system (Tatsumi et al., i990).

Miocene-Pliocene Hydrothermal Ore Deposits

231

In any model, back-arc basins form under the extensional stress regime and are associated by Mariana-type subduction by Uyeda and Kanamori (1979) rather than Chilean-type subduction. In this tectonic situation, intense bimodal volcanism and associated seawater circulation occur, resulting to the formation of Kuroko deposits on the seafloor and formation of vein-type mineralization under subaerial condition and intense hydrothermal and volcanic CO2 fluxes to ocean and atmosphere. Such fluxes affect the long-term environmental changes (see Chapter 4).

1.6. Other hydrothermal ore deposits 1.6.1. Polymetallic vein-type deposits The Honshu arc is divided into Northeast and Southwest Honshu by the Tanakura Tectonic Line or the Itoigawa-Shizuoka Line (Fig. 1.166). There are many geological and geochemical differences in these two geologic provinces. Neogene altered volcanic rocks (Green tuff) are widely distributed in Northeast Japan, whereas Cretaceous-Neogene granitic rocks and metamorphic rocks are in Southwest Japan. 87Sr/S6Sr ratio of volcanic rocks and granitic rocks (Shibata and Ishihara, 1979) (Fig. 1.166) and ~348 and sulfur content of Quaternary volcanic rocks (Ueda and Sakai, 1984) are different in two provinces (Fig. 1.167). As for the mineralization at the middle Miocene age in the Northeast Japan, Kuroko and epithermal base-metal veins have been formed. No enrichment of Sn and W is found in these deposits. In contrast, in Southwest Japan, polymetallic veins (so-called xenothermal-type deposits in the sense of Buddington (1935) or subvolcanic hydrothermal type in the sense of Cissartz (1928, 1965) and Schneiderhrhn (1941, 1955) occur. Examples of these deposits are Ashio, Tsugu, Kishu and Obira. All these vein-type deposits have formed at middle Miocene age in western part of Tanakura Tectonic Line under subaerial environment. In these deposits, many base-metal elements (Sn, W, Cu, Pb, Zn) and small amounts of Au and Ag are concentrated. These deposits are associated with felsic volcanic and plutonic rocks along the Median Tectonic Line (MTL) or south of MTL. According to Ishihara (1977), these granitic rocks are ilmenite-series while granitic rocks in Green tuff region are magnetite-series. 3 1.6.1.1. Ashio deposit The Ashio mine which is located in middle Honshu is one of the largest Cu ore producer in Japan until it was shut down in 1973 (Kanehira, 1991). The ore deposits

3Ilmenite-series and magnetite-series granitic rocks are defined as follows (Ishihara, 1977): the magnetiteseries and ilmenite-series granitic rocks are distinguished by the presence or absence, respectively, of magnetite in polished sections.

232

Chapter 1

HOKKAIDO

TANAKURA TECTONIC LINE ,"

0

0

o

o.,d KYUSHU

MEDIAN TECTONIC LINE

d

O U

I

300 KM ,

i

I

Figure 1.166. Distribution or initial 878r/86Sr (r i) for the Cretaceous-Paleogene plutonic rocks. Open squares: gabbros; open circles: granites. Numbers indicate the last two or three digits of ri values. Solid squares and solid circles imply the magnetite-series; open squares and open circles the ilmenite-series. (Shibata and Ishihara, 1979).

formed at middle Miocene age, associated with rhyolitic intrusive activity (Nakamura, 1970). The deposits occur in rhyolitic body as a funnel-shaped mass (Figs. 1.168 and 1.169). The deposits are characterized by conspicuous metal zoning and polymetallic mineralization. From the centre to margin of the mine district, the following zonings are recognized; S n - W - B i - C u zone, Cu-As-Zn zone, and Z n - P b - C u - A s zone (Nakamura, 1970). 3180 values of quartz are +12.8%o to 12.2%o (SMOW) (early-stage ( S n - W - B i Cu) quartz with cassiterite), +11.5%o to +11.4%o (quartz with stannite), +10.3%o to +8.8%0 (middle Cu-As-Zn stage), +9.4%e to +6.4%0 (late Z n - P b - C u - A s stage) (Shimazaki et al, 1993). If temperature is 300°C estimated from homogenization temperatures

233

Miocene-Pliocene Hydrothermal Ore Deposits 240

70

WEST JAPAN

420

6

60

Southwest Japan (SW) • RyukyuArc (RY)

A[

50

v High 834S rock (HI)

RY

* Alkaline (AL)

~" 4O

@

0..

~" 30 20

®

HI

10

+5

501

+10

EAST JAPAN

~ I M

30

20

0

-2

;H

HK

~'

_.~_~

0

+20

o Akita-Komagatake(AK) Hakone (HK) * Nasu zone (NA) , Chokai zone (CH) = Izu-Mariana(IM)

4O

(3.

+1'5

NA

+5

,

+10 ~34s (%,,)

+15

Figure 1.167. Sulfur content v s . 3 3 4 8 value for Quaternary volcanic rocks of Japan. Field bounded by solid lines show two volcanoes (AK and HK), two volcanic zones (NA and CH) in Northeast Japan, three volcanic belts (IM, SW and RY), alkaline rocks (AL) and volcanic rocks of unusually high 334S values (HI) in Ryukyu belt. Symbols surrounded by small circles show 334S values of Satsuma-Iwojima volcanic rocks in West Japan (Ueda and Sakai, 1984).

of fluid inclusion throughout all stages of mineralization, ~180 values of hydrothermal solution are early-stage: +4%° to +6%o; middle stage: +2%o to +3%0; and late-stage: - 1 % o to +2%0 (Shimazaki et al., 1995). This variation indicates that hydrothermal solution at early-stage was dominantly o f magmatic origin (or igneous origin), and a contribution of meteoric water component increased with the stage of mineralization. Main opaque minerals are chalcopyrite, cassiterite, stannite, arsenopyrite, bismuthinite, pyrrhotite and sphalerite. The FeS content of sphalerite is high (about 18 tool% FeS). Dominant gangue minerals are quartz, sericite, kaolinite and siderite. The Fe content of chlorite is very high (Fig. 1.170). All these data suggest low fs2 and f o 2 conditions.

234

Chapter 1 A V Jv

V

V

V

V

V,"

v

V

V

,

I1[

"","7"

/

\

II1

~°~°~

~_-'f.~" B

I: (Central) Sn-W-Bi-Cu zone I1: (Intermediate) Cu-As-Zn zone Ill: (Marginal) Zn-Pb-Cu-As zone

Figure 1.168. Geologic map of the Ashio mining area, showing the mineral zoning in the Ashio rhyolitic body and the location of the "Kajika" deposits (alter Ashio mine, partly revised) (Nakamura, 1970).

~34S values o f sulfides are close to 0%~ (Shimazaki, 1985), suggesting magmatic or igneous origin and no contribution of seawater sulfur. These values are quite different from those o f Green tuff sulfur (Kuroko and epithermal base-metal vein-type deposits; +2%~ to +7%o).

1.6.1.2. Tsugu deposit The Tsugu g o l d - a n t i m o n y deposit is located in the central part o f Honshu, Japan (Fig. 1.171). Geologically, the Tsugu deposit occurs in the Ryoke metamorphic terrane,

235

Miocene-Pliocene Hydrothermal Ore Deposits m

1200 A 1000 8O0 600 400

I [ ~~TM,. ~

•

\I

~

/

2OO

0t Rhyoliteweldedtuff ['L~

~

Rhyolite

S,a,e

~

OreVei. ' ai,k; depos,,s

lternationof sandstone

and slate

Basalbreccia

H : Hotel vein,

lO00m

~Chert

[ q ~ " - I Rhyolite(lithoklite) dike. ~

500

J : Jimbovein, S : Shinseivein, Y : Yokomabuvein,

Figure 1.169. Geologic section along the line A B in Fig. 1.168 showing the mineral zoning in the Ashio rhyolitic body (Nakamura, 1970).

Mg4AI/ /

/"\ \

/ • ', /

/

L/'_ _~1.42 / . . . .

/',,,,,' Mg6

/"\ \

z13\

', ../.

/"k 06 l U ~ Fe(Mn)]4AI2

\,it/

/

"x,~.T12 - IlL/ -

\ 10_4/03\

\~.~/ i \

--$~--2_~

',, .,, ,,, ,,, ,,, ,,, ,,, ,,, \ [Fe(Mn)]6

Figure 1.170. Diagram showing the octahedraI composition of chlorites from the subvolcanic hydrothermal deposits, propylite, and Kuroko deposits in Japan (Nakamura, 1970). Chlorite occurring as a gangue mineral in the subvolcanic hydrothermal deposits: Nos. 1, 2, 3 and 4: Chlorite from the Ashio copper mine. Nos. 5, 6, and 7: Chlorite from the Kishu mine. No. 8: Chlorite from the Arakawa mine. Nos. 9 and 10: Chlorite from the Ani mine. No. 11: Chlorite from the Osarizawa mine. Chlorite from the so-called propylite: No. 12: Chlorite from the Yugashima mine. No. 13: Chlorite from the Budo mine. Chlorite from the Kuroko deposits: No. 14: Chlorite from the Wanibuchi mine.

adjacent to the outer z o n e o f southwestern Japan. In this z o n e and in the R y o k e m e t a m o r phic terrane, A u deposits are rare, w h e r e a s there are m a n y H g - S b deposits. O n the other hand, as already noted, in the Tertiary s u b m a r i n e v o l c a n i c r e g i o n ( G r e e n tuff region) and Quaternary v o l c a n i c region, m a n y e p i t h e r m a l v e i n - t y p e deposits occur. M i n e s w h i c h h a v e p r o d u c e d b o t h Sb and A u are very rare. T h e r e f o r e , the T s u g u deposit is an unusual type o f A u deposit in Japan. I g n e o u s activity related to the m i n e r a l i z a t i o n indicates that the age o f m i n e r a l i z a t i o n was m i d d l e M i o c e n e and this age is different f r o m that o f e p i t h e r m a l A u - A g mineralization, m o s t l y late M i o c e n e to Pleistocene, as already noted.

236

Chapter 1

17= z,0o - - -

35 °

~.

Tsugu D e p o s i t

i 133°

0

200kin

138°

Figure 1.171. Map showing location of the Tsugu gold-antimony deposit. I: Green-tuff region, 2: outer zone of southwestern Japan, TTL: Tanakura tectonic line, 1STL: Itoigawa-Shizuoka tectonic line, MTL: Median tectonic line (Shikazono and Shimizu, 1988a).

The previous studies on the Tsugu deposit (Tsuboya, 1936; Tatsumi, 1948; Shikazono and Shimizu, 1988b) demonstrated that: (1) Opaque minerals include stibnite, jamesonite, cinnabar, gold, pyrite, pyrrhotite, arsenopyrite, marcasite, sphalerite, galena and chalcopyrite. (2) The ore minerals display a zonal distribution: gold and cinnabar are enriched in the upper part of the veins, and sphalerite, galena and chalcopyrite are more abundant in the deeper parts. Pyrrhotite and arsenopyrite are distributed throughout the veins. (3) From the mode of occurrence of opaque minerals it is considered that pyrrhotite and sphalerite were precipitated at an early-stage, gold, pyrite, marcasite, stibnite and cinnabar were precipitated at a late-stage, and arsenopyrite was precipitated throughout the mineralization period. (4) The accessory minerals are tourmaline and apatite. Hg content of electrum ranges from nil to 8.5 wt%, and Ag content ranges from

237

Miocene-Pliocene Hydrothermal Ore Deposits Au

(z-solid /

so, t,o

Ag

10

20

j

•

30

•

40

50

•

60

70

80

90

Hg

Hg atomic % Figure 1.172. Chemical composition of mercurian gold plotted on a Au-Ag-Hg triangular diagram. Data sources: Tsugu (open square; Shikazono and Shimizu, 1988b); L~ngsele (solid circle; Nysten, 1986); Aitik, northern Sweden (solid triangle; Nysten, 1986); Kazakhstan (cross; Naz'mova and Spiridonov, 1979); Bajupura-Dariba (solid square; Basu et al., 1981)• Phase boundaries (solid lines) at 450°C were drawn from Basu et al. (1981) (Shikazono and Shimizu, 1988b).

6.0 to t 1.5 wt%. The chemical compositions of gold from the Tsugu deposit are plotted in terms of Au, Ag and Hg, together with data from the literature (Fig. 1.172). The Au/Ag ratios of electrum from the Tsugu deposit are high compared with those from the other mercurian gold occurrences. Note that Hg content varies inversely with Au content (Fig. 1.173), implying that Hg substitutes for Au. A frequency histogram of the Ag content of electrum from epithermal gold-silver vein-type deposits and the Tsugu deposit (Fig. 1.174) clearly indicates that the Au/Ag of electrum from the Tsugu deposit is higher than that from epithermal vein-type deposits. Electrum and cinnabar occur in the same part of the vein, suggesting that these minerals were in equilibrium. The following reaction can be used to determine which factors are important for controlling mercury content of electrum in equilibrium with cinnabar: (Hg) + 1/2S2 = HgS

(1-78)

where (Hg) denotes the Hg component in electrum. From the equilibrium relation for reaction (1-78), we obtain, NHg

1/(gl-78fs2 YHg) 1/2

(1-79)

where NHg is the atomic fraction of the Hg component in electrum, K1-78 is the equilibrium constant of reaction (1-78), i s 2 is the sulfur fugacity, and gHg is the activity coefficient of the Hg component in electrum.

238

Chapter 1 8

6

E

4

-r

2

0

70

75

80

85

90

Au atomic % Figure 1.173. Relationship between mercury and gold contents of mercurian gold from the Tsugu goldantimony deposit (Shikazono and Shimizu, 1988b).

From (1-79) it is clear that the Hg content of electrum is related to f s 2 and K1-78. Using the thermochemical data for reaction (1-78) by Barton and Skinner (1979), isoactivity lines for Hg in electrum may be drawn on a log fsz-temperature diagram (Fig. 1.175). At a given temperature, the activity of Hg in electrum increases with a decrease in fs2. Therefore, the f s 2 for the mercurian gold of the Tsugu deposit is inferred to be relatively low. Microprobe analyses of sphalerite indicate that sphalerite from the Tsugu deposit is markedly higher in iron than that from the epithermal Au-Ag vein-type deposits in Japan. Measurements of homogenization temperatures of fluid inclusions that were obtained from quartz coexisting with sphalerite, pyrrhotite and chalcopyrite of the earlystage of sulfide mineralization, and from quartz coexisting with electrum, arsenopyrite, pyrite and sericite of the late-stage of mineralization. Homogenization temperatures of fluid inclusions are 281-345°C for the early sulfide stage, and 275-295°C for the latestage of gold mineralization. These fluid inclusion data, stability fields of arsenopyrite, pyrite and pyrrhotite, and the iron content of sphalerite can define the possible ranges of sulfur fugacity and temperature for the Tsugu mineralization (Fig. I. 176). The ranges of is2 and temperature for epithermal Au-Ag vein-type deposits in Japan have been clearly defined based on the chemical composition of sphalerite and electrum, and homogenization temperatures of fluid inclusions (Shikazono, 1985d). Values of fs2 for the Tsugu deposit are lower than the typical ranges of values for the epithermal Au-Ag vein-type deposits in Japan (Fig. 1.176). Such a low i s 2 is in accord with the high Hg content of electrum in the Tsugu deposit.

239

Miocene-Pliocene Hydrothermal Ore Deposits 70 ¸

60

50 O ©

40

(1) ~4 30

20

10

20

40

60

Ag a t o m i c

80

%

I00

Figure 1.174. Frequency (number of analyses) histogram for Ag (atomic %) of gold from the Tsugu deposit (solid) and epithermal gold-silver vein-type deposits in Japan (open). I: sampte I; II: sample II. Data sources: Tsugu deposit (Shikazono and Shimizu, 1988b); epithermal gold-silver vein-type deposits in Japan (Shikazono 198l, 1986; Shikazono and Shimizu, 1988b). ~xx%

o .a -15

150

200

250

300 Temp.(°C)

Figure 1.175. Activity of S2-temperature diagram showing iso-Hg contents contours for gold. The calculations were carried out using thermochemical data of Craig and Barton (1973).

240

Chapter 1

/-

~

-10

o

-

-1 ~

'

J

'

,

,

r

150

,

,

,

,

I

200

c

i

~

I

250

I

i

t

i

i

O k/~

J

I

300

.

e

%t-.

,. ~,.o-

//...~,%.-'//

//////" ,

/

-

i

i

i

i

1

350 Temp,(°C)

Figure 1.176. Possible ranges of activity of $2 (as2) and temperature for the Tsugu deposit (solid and dotted areas) and epithermal gold-silver vein-type deposits in Japan (shaded area) (Shikazono and Shimizu, 1988a).

Mercurian gold has been reported from several Au deposits: Longsele, Aitik, Kazakhstan, and Bajpura-Dariba (Naz'mova and Spiridonov, 19797 Basu et al., 1981, Nysten, 1986). The nature of the opaque minerals coexisting with mercurial gold in these minerals are pyrrhotite, galena, sphalerite, stibnite, and jamesonite, and less common are cubanite, arsenopyrite, pyrite, and owyheeite. Pyrrhotite, cubanite, and arsenopyrite suggest relatively low fs2, but the temperatures of formation of these deposits have not been estimated. 1.6.1.3. Kishu deposit The Kishu deposit, located at middle Honshu, is associated with Kumano acidic rocks intruded into the Shimanto and Kumano formations. K-Ar age determination indicates that the activity of Kumano acidic rocks occurred at middle Miocene (144-1 m.y.) (Ikebe, 1973). The stage of ore mineralization from early to late is (1) Cu-Fe (chalcopyrite, pyrite, chlorite, quartz), (2) Cu-Zn-Pb (chalocopyrite, sphalerite, galena, sericite, quartz, siderite, calcite, fluorite, chlorite), (3) Au-Ag (native gold, argentite, pyrite, sphalerite, chalcopyrite, chlorite, quartz, adularia, sericite), (4) calcite (Nakamura and Miyahisa, 1976). ~348 of sulfides from Kishu is very low (-14.2%o) (Shimazaki, 1985), indicating sedimentary sulfide sulfur source. 1.6.1.4. Obira deposit Obira deposit is located at south Kyushu. It occurs in Shimanto Group rocks composed of shale, and limestone, associated with middle Miocene granitic intrusive rocks. Polyascendant zoning and mineralization are observed in the Obira deposit. From early- to late-stage mineralizations are:

241

Miocene-Pliocene Hydrothermal Ore Deposits cassiterite-apatite-tourmaline --+ cassiterite-wolframite-molybdenite-pyrite --+ pyrite-arsen•pyrite-pyrrh•tite-magnetite-cha•c•pyrite-stannite-bismuthinite-ga•enatetrahedrite --+ marcasite-pyrite --+ marcasite-Sb minerals (stibnite, jamesonite, berthierite, boulangerite).

Occurrence of pyrrhotite, arsenopyrite and high iron content of sphalerite indicate low i s 2 and / O 2 conditions. Decrepitation temperatures of fluid inclusions in quartz from S n - C u - A s stage are 403-322°C (Enjoji and Takenouchi, 1976). ~34S values of sulfides are low (-6.7%o) (Shimazaki, 1985), suggesting a contribution of sedimentary sulfide sulfur. Distinct features of the above-mentioned polymetallic deposits are summarized as follows: (1) The FeS content of sphalerite and Fe content of chlorite from these deposits are high. (2) Temperature of formations is high and variable. (3) fs2 and f02 are low. (4) Many metals (Cu, Pb, Zn, Au, Ag, Sn, W, Bi) have been produced. (5) Orebody and regional metal zonings are conspicuous. (6) Salinity of inclusion fluids varies widely from high to low. (7) Common gangue minerals are kaolinite and sericite, but K-feldspar is not found, suggesting low pH of ore fluids. (8) The deposits have been formed under the subaerial environment. No contribution of seawater and low fs2 and f02 conditions are consistent with the geologic environment.

1.6.1.5. Temperature and sulfur fugacity estimated from iron and zinc partitioning between coexisting stannite and sphalerite and coexisting stannoidite and sphalerite Stannite is the most common tin sulfide mineral in the ore deposits associated with tin mineralization. This mineral sometimes contains appreciable amounts of zinc, together with iron. Several workers have suggested that the zinc and iron contents of stannite are related to temperature. With respect to the study of the phase relationships in the pseudobinary stannite-kesterite system, Springer (1972) proposed zincic stannite as a possible geothermometer mainly based on the chemical compositions of the two exsolved phases (stannite and kesterite). Nekrasov et al. (1979) and Nakamura and Shima (1982) experimentally determined the temperature dependency of iron and zinc partitioning between stannite and sphalerite. Because natural stannite contains a considerable amount of zinc, sphalerite contains a considerable amount of iron, and these contents can be easily analyzed using an electron microprobe, a stannite-sphalerite pair is expected to be a useful indicator of formation temperature and sulfur fugacity. Iron and zinc partitioning between stannite and sphalerite is represented by the exchange reaction. Cu2FeSnS4 (in stannite)+ ZnS (in sphalerite) = Cu2ZnSnS4 (in stannite)+ FeS (in sphalerite)

(1-80)

242

Chapter 1

"" 1.0 cn o

I .

~.

1

Nekrasov el 0t.0979),

looKd =l.27&.lOd'T-1-l.174

7 Nokomuro & Shimo ( 1982 ), log Kcl = ZS.103.T- 1_3fi

2.0

250

l~a

300

350

1.6

400 450

~4

590~C

f.2

103"T -1 Figure 1.177. Comparison between the stannite-sphalerite geothermometer after Nekrasov et al. (1979) and one after Nakamuraand Shima (1982). Crossbars indicate experimental uncertainties (Shimizu and Shikazono, 1985). 153

The partition coefficient, Kd, for the above reaction is expressed by,

Kd = (Xcu2FeSns4/Xcu2znsnS4)stannite/(XFeS/XZnS)sphalerite

(1-81)

where X denotes mole fraction of a given component in stannite and sphalerite. Nekrasov et al. (1979) and Nakamura and Shima (1982) reported a temperature dependency of iron and zinc partitioning between stannite and sphalerite (Fig. 1.177) logKd = 1.274 x 103T -I - 1.174 (Nekrasov et al., 1979)

(1-82)

logKd = 2.8103T -1 - 3.5 (Nakamura and Shima, 1982)

(1-83)

Both geothermometers are in agreement being close 380°C (Fig. 1.177). However, at the lower and higher temperatures the difference between the temperatures estimated from equations (1-82) and (1-83) becomes larger. Figure 1.178 represents a comparison between the stannite-sphalerite temperatures and homogenization temperatures of fluid inclusions or sulfur isotope temperatures. It can be seen in Fig. 1.178 that Nakamura and Shima's geothermometer would be rather consistent with the temperature estimated based on the fluid inclusions or sulfur isotope studies. It is notable that almost all stannite-sphalerite temperatures are within 30°C of average homogenization temperatures and sulfur isotope temperatures. Temperatures estimated based on equations (1-82 and 1-83) range from ca. 250°C to 350°C for both skarn-type and vein-type deposits except the Yatani epithermal Au-Ag vein-type deposits.

243

Miocene-Pliocene Hydrothermal Ore Deposits °C

°C

400.

400 /

E O

E 300,

IoT ~K

or' (N

tO

200

200

100

100

@

C O D

"6 2

0

Nekrasov et al (1979) 0

100

200

300 °C

0

Nakamura and Shima (1982) t

100

200

i

300 °C

stannite-sphalerite geothermomentry Figure 1.178. Comparison between the stannite-sphalerite temperatures and fiIIing temperatures of fluid inclusions or sulfur isotope temperatures.NT: Nakatatsu, OB: Obira, KN: Kant, KG: Kuga, TM: Tsumo, KM: Kamioka, OT: Ohtani, KU: Kaneuchi, Ak: Akenobe, TT: Takatori,YT: Yatani (Shimizu and Shikazono, 1985). The chemical compositions of the coexisting stannite and sphalerite are plotted on log(XFeS/XZnS)sphaterite--log(Xcu2Fe2SnSg/XCu2ZnSnS4)stannite diagrams (Figs. 1.179 and 1.t80). Iso-logfs2 lines on these diagrams are based on equation (1-83), and thermochemical data on the F e - Z n - S system by Scott and Barnes (1971). The line 20.8 mol% FeS in sphalerite, which corresponds to the composition of sphalerite in equilibrium with pyrite and pyrrhotite at 1 bar as determined by Boorman (1967), is also given on the diagrams. It is evident in Figs. 1.179 and 1.180 that the FeS content of sphalerite coexisting with pyrite is generally lower than that of sphalerite coexisting with pyrrhotite or both pyrite and pyrrhotite. Iso-FeS content lines for sphalerite in equilibrium with pyrite or pyrrhotite were drawn on the log fs2-temperature diagram (Figs. 1.179 and 1.180) using thermochemical data by Scott and Barnes (1971) and Barton and Skinner (1979). The relationship between the iron content of stannite in equilibrium with sphalerite and pyrite or with sphalerite and pyrrhotite was derived based on thermochemical data by Scott and Barnes (1971), Barton and Skinner (1979) and Nakamura and Shima (1982). These types of deposits are skarn-type polymetallic (Sn, W, Cu, Zn, Pb, Au, Ag) vein-type and S n - W vein-type deposits. As shown in Fig. 1.181, the fs2-temperature range for each type of deposits is different; at a given temperature, fs2 increases from S n - W vein-type through skarn-type to polymetallic vein-type deposits. It is interesting to note

244

Chapter 1

250°C

¥r

].

Ti.0

am

/ ~ 1 2 C U')

u7 rt/~

U

(D

450°C

0

0

-t

-t

0

log XFeS in spholerite XZnS Figure 1.179. log(XFeS/XZnS)sphalerite--Iog(XCuzFeSnS4/Xcu2ZnSnN4)stannite diagram showing that sphalerite and stannite are associated with pyrrhotite (Po) and/or pyrite (Py). Temperature lines are based on data by Nakamura and Shima (I982). Solid curves show logfsz based on data by Scott and Barnes (1971) in the pyrrhotite field. Abbreviations are the same as in Fig. 1.178 (Shimizu and Shikazono, 1985). that i s 2 associated with gold-silver mineralization in polymetallic deposits is relatively high compared with those of the other types of deposits. This difference is similar to that found in epithermal vein-type deposits; fs2 of epithermal A u - A g veins is higher than that of base metal veins. Shimizu and Shikazono (1987) studied the compositional relations of coexisting stannoidite, sphalerite and tennantite-tetrahedrite (Fig. 1.182). Based on these data they estimated the sulfur fugacity of stannoidite-bearing tin ore. Considering the complementary work on stannite-bearing tin ores from Japanese ore deposits (Shimizu and Shikazono, 1985), a comparison between environmental conditions of these two types of tin sulfides was made. Their study is described below. The chemical compositions of coexisting sphalerite and tennantite-tetrahedrite from the mines were determined. Except the Ashio polymetallic deposits, the other deposits have been formed at late Cretaceous related to felsic magmatism. Scanning patterns and many analytical data obtained by electron-microprobe analyzer reveal that most stannoidite grains are compositionally homogeneous. There is

245

Miocene-Pliocene Hydrothermal Ore Deposits

YT

250oc 9'/', ,

I

/iK@iL,

C 0

/

C

3sooc ,..5ooc

o

-1 -2

,

/

.

Py

1

tog XFeS XZnS

in sphalerite

Figure 1.180. log(XFeS/XZnS)sphalerite Iog(Xcu2FeSnS4/XCu2ZnSnS4)stannitediagram showing each deposit is skarn-type or vein-type. Abbreviations are the same as in Fig. 1.178 (Shimizu and Shikazono, 1985).

a wide range in extent of Fe and Zn substitution in stannoidite (NFe/l~Zn between 2.03 and 14.4). The stannoidite from the Tada deposit is the richest in Zn (5.24 wt%, Zn, 9.09 wt% Fe) and has the approximate formula Cu8Fe2ZnSn2S12; that from the Konjo deposit is the richest in Fe (11.90 wt% Fe, 1.69 wt% Zn) and has the formula Cu8Fe2(Fe0.77Zn0.19)Sn2S12. From the Fe2+/Zn 2+ ratio of stannoidite, a continuous solid solution is inferred to exist between Cu8Fe2ZnSn2S12 and CusFe2FeSn2SI2. Iron and manganese contents of sphalerite coexisting with stannoidite are in the range from 0.17 to 1.93 wt%, and from 0.02 to 0.16 wt%, respectively. Lee et al. (1974) conducted an experimental study on the equilibrium for the assemblage of stannoidite-chalcopyrite-bornite-mawsonite-S2 (gas) in a temperature range from 430 to 300°C. Curves A and B in Fig. 1.183 correspond to fsz-temperature relationships for this equilibrium assemblage for aFe = 1 and aFe = 0.1, where aFe is the activity of the Cu8Fe2FeSn2S12 component in stannoidite solid solution. As mentioned already, Shimizu and Shikazono (1985) have estimated the is 2temperature range for stannite-bearing assemblages from Japanese vein-type and skarntype tin deposits. This estimated fs2-temperature region is also shown in Fig. 1.183. The fsz-temperature range for the formation of these two types of tin sulfides is different.

246

Chapter 1

Cn 0 I

10

11

Temperoture, *C

250 30O 35O *C Figure 1.181. Temperature-log fs: diagram. Skarn-type deposits (solid squares) are considered to be formed under lower fs2 condition than vein-type deposits (open squares). Abbreviations are the same as in Fig. 1.178 (Shimizu and Shikazono, 1985).

Microscopic observation suggests that stannite formed earlier than stannoidite on the scale of one polished section, although in general coexistence of stannoidite and stannite cannot be observed in the same polished section. If stannite formed at an earlier stage than stannoidite, it could be inferred that fs2 increased or that temperature decreased (or both) with evolution of tin mineralization. The Fe2+/Zn 2+ ratio of coexisting stannoidite, sphalerite and tennantite-tetrahedrite from the Tada, Omodani and Ohmidani deposits is low, compared with that from the other deposits such as the Ashio, Akenobe and Ikuno deposits (Fig. 1.182). The Tada, Omodani and Ohmidani deposits are characterized by polymetallic (Zn-Cu-Ag-Au; Zndominated) mineralization, and tungsten is not recovered from these deposits. On the other hand, the Akenobe, Ikuno and Ashio deposits are characterized by polymetallic (Cu-AnPb-Sn-W-Ag-Au-Bi) mineralization, and tungsten is recovered from these deposits. The Fe/Zn ratio of coexisting stannoidite, sphalerite and tennantite from the tungsten-bearing polymetallic deposits is higher than in those from the tungsten-free polymetallic deposits. High content of indium in chalcopyrite from tungsten-beating polymetallic deposits probably indicates a high temperature of formation. Sphalerite rarely appears where the Fe2+/Zn 2+ ratio of the coexisting stannoidite exceeds approximately 1. These differences imply that the tungsten-bearing deposits formed under lower fs2 or higher temperature conditions (or both) than tungsten-free deposits.

247

Miocene-Pliocene Hydrothermal Ore Deposits

KONJO

2-

2

I~KUNO A ¢N

ASHIO

1

~OMO_~I

N'ENOE

AKENOBEI

~,

/ FUKOKU

0

I 2 3 4 ( Fe 2" I Zn 2" )sp x 10 2

ASHIO

=~"KOKU

tOMO' ~DANI OHMIDANI 0

1 2 3 ( F e z. / Zn 2" )tenn-tetr

Fig 1.182. Fe2+/Zn2+ of sphalerite (left) and of tennantite-tetrahedrite-series mineral (right) as a function of FeZ+/Zn2+ in stannoidite (atomic proportions) (Shimizu and Shikazono, 1987).

Figure 1.183 shows i s 2 and temperature region for stannoidite-bearing tin deposits is different from that for stannite-bearing tin deposits.

1.6.2. Hg and Sb deposits Hg deposits are distributed along the MTL (Median-Tectonic Line) in Southwest Japan and in Northeast Hokkaido (Fig. 1.184). The deposits are vein or disseminated in form. The deposits are hosted by sedimentary and igneous rocks. No K - A r age data on the deposits are available. However, from the age of host rocks, the age of igneous activities along MTL and the studies on the movement of MTL and K - A r ages of Au-veins associated with Hg mineralization in Northeast Hokkaido (e.g., Khonomai), it is likely that these deposits formed at middle Miocene age. However, mercury mineralization in Kitami Province (north Hokkaido) occurred at approximately the same age as the epithermal gold-silver mineralization in the same district (4.5-5.3 Ma) (Maeda, 1997). Main ore minerals are few. They are cinnabar, metacinnabar, stibnite and pyrite. Gangue minerals are quartz and calcite. Fluid inclusion decrepitation temperatures

248

Chapter 1

0001

Stcmnoidite r e g i o n

/ 0005

0

~'-. . . jf

10 ," ///

15

B

A

20

Stotnnite r e g i o n (Shimizu & 5hikazono 1985)

./

"

200

Temperature 250

3 0

350"C

Figure 1.i83. Temperature-log fs2 diagram for stannoidite-bearing ores from Japanese veto-typedeposits. MW = Mawsonite, Sd = Stannoidite, Bn = Bornite, Cp = Chalcopyrite(Shimizu and Shikazono, 1987).

for quartz and calcite from the Itomuka Hg deposit are 328-337°C and 201-242°C, respectively (Enjoji and Takenouchi, 1976). ~34S of cinnabar and stibnite are wide in a range from -16%e to +15%o, and from -15%o to +4%0 for cinnabar and metacinnabar, and stibnite, respectively (Shikazono and Shimizu, unpublished). Usually, 334S values of cinnabar and stibnite are low less than 0%o except for these minerals from Green tuff region. These low 334S values imply that sulfide sulfur of Hg and Sb deposits came dominantly from sedimentary sulfur, and no contribution of seawater sulfur. The sulfur isotopic data are consistent with geologic environments of Hg and Sb deposits; Sedimentary rocks are dominant and marine rocks are not present in Sb-Hg mineralization districts. However, a few samples of stibnite and cinnabar from the deposits in Green tuff region display high ~34S values. In contrast of this interpretation on the origin of sulfur, Ishihara and Sasaki (1994) thought that sulfur came from ilmenite-series granitic rocks. However, these rocks are not found in the north Hokkaido. Horikoshi (1995) showed areal extension of volcanic belts since middle Miocene in the Japanese Islands. He suggested that Hg and Sb mineralization in outer zone of Southwest Japan and north Hokkaido (Kitami district) related to forearc igneous activities in trench side. ~34S values become heavier across the Northeast Japan arc from the trench to the back-arc side. He thought that the changes in tectonic character from forearc to arc environment controlled the ~34S values.

249

Miocene-Pliocene Hydrothermal Ore Deposits

p

R

i Sado

O0

,

g

Sanin

Kyoto s

%

o okue~_ ~

~o~s p~o.4\~ce

e~ 're'=

\\ra

zone,

o~te~

qxl'l ,~*~"

,0

0 I • e • e o

I

I

3 0 0 km I

Hg Ore Deposits Sb Ore Deposits Au-Ag Ore Deposits Pb-Zn Ore Deposits Sn-W Ore Deposits

o Vein o

m Skarn

D

Figure 1.184. Distribution of major sulfide deposits in Miocene volcanic and plutonic provinces. Class i and 2 are ore deposits (that is, the largest and second largest are shown together with the larger symbols). Class 3 is shown as a smaller circle and box. Tin ores occur associated with base-metal deposits at Mt. Okue, and tungsten deposits are concentrated in Yakushima, Kyushu (Ishihara and Sasaki, 1991).

1.6.3. Gold-quartz vein-type deposits (mesothermal-hypothermal vein-type deposits) 1.6.3.1. Geology, mineralogy and geochemistry G o l d - q u a r t z v e i n - t y p e ( m e s o t h e r m a l - t y p e a n d h y p o t h e r m a l - t y p e d e p o s i t s in t h e s e n s e o f L i n d g r e n ( 1 9 2 8 ) ) o c c u r in s e d i m e n t a r y t e r r a n e a s s o c i a t e d w i t h C r e t a c e o u s f e l s i c

250

Chapter 1

0

100

200km i

40° Nishizawa.~°h °

k

/ Hayochine ~Z~;i!~v//Owashi t - - - % ~ ~//Nojiri u ~ Shishiori

K i n k e i \ F::~::/\ ~ ~ i ' ) o Amo---------~E~r,', ~ ~ _ 35°

~3 0

_/ .....

133°

/__..~

~~b'~--.

x AikGwa

hio owo

:~; \ ~

~

Suwa

AShiyasu

138°

Figure 1.185. Locations of hypo/mesothermal vein-type deposits in Japan. Abbreviations are the same as those in Fig. 1.62. Open circle: hypo/mesothermal Au vein-type deposits. Solid circle: hypo/mesothermal po]ymetallic vein-type deposits (Shikazono and Shimizu, 1988a).

plutonic rocks or in regionally metamorphosed rocks (Fig. 1.185). Deposits occur along the Tectonic lines such as Hidaka (Hokkaido), Kitakami (northeast Honshu), Tanakura (middle Honshu) and Itoigawa-Shizuoka (middle Honshu) Lines (Fig. 1.185). The deposits occurring in the sedimentary rocks (mainly black shale) are distributed mainly in three districts: Kitakami, Yamizo and Koma (Fig. 1.185). The few deposits in the metamorphic region are the Suwa, Kinkei, Amo and Hashidate (Fig. 1.185). Ages of mineralization in the Hidaka and Kitakami regions may be Cretaceous, considering the ages of associated granitic rocks. Isozaki and Maruyama (1990) thought that the Tanakura Tectonic Line moved at middle Miocene (ca. 15 Ma) related to the opening of Japan Sea. The time of movement of Itoigawa-Shizuoka Line is thought to be also middle Miocene. Ages of granitic rocks near the itoigawa-Shizuoka Line are middle Miocene (Shimizu, personal communication, 1998). These lines of evidence suggest that ages of mineralization along the Tanakura and Itoigawa Tectonic Lines are probably middle Miocene.

251

Miocene-Pliocene Hydrothermal Ore Deposits AURIFEROUS VEIN Type 1

60

Type 2

50

F7~J

f° ~

after Nedachi (1974) after Yarnaoka (1981)

GOLD-SILVER VEIN E~ after Shikazono (1981) >, 40

O c-

$ u_ 30

20

10

0

0

20

40

60

80

100

Ag atomic % Figure 1.186. Frequency (number of analyses) histogram of Ag content (atomic %) of auriferous vein and gold-silver vein deposits in Japan. Frequency means numbers of analyses (Shikazono and Shimizu, 1987).

Main opaque minerals include native gold, electrum, pyrite, pyrrhotite, chalcopyrite, cubanite, sphalerite, arsenopyrite and tellurobismutite. The amounts of these sulfide minerals are poor, compared with those in epithermal A u - A g vein-type deposits. It is noteworthy that silver minerals are abundant in epithermal A u - A g vein-type deposits, whereas they are poor in gold-quartz veins. The Ag content of electrum is very low (Fig. 1.186) and FeS content of sphalerite is high (6-17 FeS mol%) (Fig. 1.187) (Shikazono and Shimizu, 1987). Combining these compositional data with homogenization temperatures of fluid inclusions, fs2 of ore fluids was estimated (Fig. 1.188). Estimated fs2 range is lower than that of epithermal A u - A g vein ore fluids. Shikazono and Shimizu (1987) calculated Z A u / I ; A g of ore fluids using fluid inclusion data (salinity, temperature), pH, mH2s and mK+. The procedure for estimating Z A u / Z A g by them is given below.

252

Chapter 1

100

•

AURIFEROUSVEIN

[--] GOLD-SILVERVEIN 80 O ¢-

60

t:r

1.1_ 401

20

--

0

2

4

6

8

10 12 14 16 18

FeS mole % of sphalerite Figure 1.187. Frequency (number of analyses) histogram of FeS content (mol %) of sphalerite from auriferous vein deposits in Japan (Shikazono and Shimizu, 1987).

^~,'-~*~//7..: . ~

~,

-10 d

-15 ,

~O~:5,,,,,.,~ ~

150

~' S,4~.,~

~'3P''f:££'O'a

200

250

300

350

Temp (°C)

Figure 1.I88. Typical sulfur activity and temperature ranges for Japanese auriferous vein (dotted) and goldsilver vein (hatched) deposits. Iso-FeS content curves for sphalerite were drawn based on the equation of Barton and Skinner (1979). py: pyrite, po: pyrrhotite (Shikazono and Shimizu, 1987).

It is essential to know the mode of transport o f Au and Ag in ore fluids to consider the factors which control the A g / A u ratio o f native gold and electrum. Many studies on Au and Ag complexes in ore fluids have been conducted and reviewed by several workers (Barnes and Czamanske, 1967; Barnes, 1979; Seward, 1981; Shenberger, 1986).

253

Miocene-Pliocene Hydrothermal Ore Deposits

According to these previous studies, the most dominant dissolved states of Au and Ag in ore fluids are considered to be bisulfide and chloride complexes, depending on the chemistry of ore fluid (salinity, pH, redox state, etc.). However, very few experimental studies of Au solubility due to chloride complex and Ag solubility due to bisulfide complexes under hydrothermal conditions of interest here have been conducted. Thus, it is difficult to evaluate the effects of these important species on the Ag/Au of native gold and electrum. Other Au and Ag complexes with tellurium, selenium, bismuth, antimony, and arsenic may be stable in ore fluids but are not taken into account here due to the lack of thermochemical data. Assuming that Au is transported dominantly as bisulfide or chloride complexes, the following reaction can be used to determine which species are dominant under near neutral conditions• AuC12 + 2H2S -----Au(HS) 2 + 2C1- + 2H +

(1-84)

The equilibrium relation for reaction (1-84) is expressed as 2 2 2 aauci 2/aau(HS) ~ = mauci 2/mAu(HS) ~ = (aci- all+)/(K1-84aH2S)

(1-85)

where a is activity, m is molality, and K1-84 is the equilibrium constant for equation (1-84)• It is assumed that the activity coefficient ratio, YauC12/yAu(US)2, is one. Several investigations to elucidate the geochemical environment of Japanese epithermal A u - A g vein-type deposits have bee,, ,'onducted (Shikazono, 1974a, 1977a, 1978b, 1985a,b). Based on these studies, the valu306of mauci-/mAu(HS) as a function of tem• perature may be calculated (Fig. 1.189), by taking the average values for a c v , all+, au2s, and temperature. The ratio of mAuci- /mAu(HS) increases with increasing temperature and acl (Fig • 1" 189), while an increase m pH causes a decrease in m I-A u t~. / m ~. U. [.r l.~.) ~. ratio " ~l 2 It is apparent in Fig. 1.190 that Au bisulfide species are more abundant than Au chloride species under the conditions common for ore fluids responsible for Japanese A u - A g veins. However, Au chloride species may dominate Au bisulfide species in ore fluids responsible for the gold-quartz (auriferous) vein deposits, as shown in Fig. 1.189. Assuming that AgC12 and Au(HS)~- are the predominant Ag and Au species, the following reaction may be used to derive the relationship between the Ag/Au ratio of native gold or electrum and temperature or other variables: •

•

.

.

2

.

2

2

(Au) + AgC12 + 2H2S = (Ag) + Au(HS)~ + 2C1- + 2H +

2

.

•

(1-86)

where (Au) and (Ag) are the Au and Ag components of native gold and electrum, respectively. From the equilibrium relation for equation (1-86) we obtain,

aAg/aAu ----(a22smagc12 KI-86)/mau(Hs)2a21 a2+

(1-87)

where K1-86 is the equilibrium constant for equation (1-86)• Equation (1-87) implies that the aAg/aAu of native gold and electrum is controlled by temperature, an.-, aN-S, pH, and m -A g t A~.2 / m A. .u t.t a.~.) 2. " ~1 2 The curves representing the relationship between temperature and aAg/aAu (aag, activity of Ag component in native gold or electrum; aAu, activity of Au component in native gold or electrum) are shown in Fig. 1.190. It is assumed that aK+/aN+ is controlled

254

Chapter 1

6 5 t~

3

z

I::::: i : ' ' ~ : : : :"::::::::' :": : :

£ fi 131 O --

T-,

0

......

-2

-3 -4 -5 ,

200

,

,

.

,

250

.

.

.

.

i

300

T(oC )

Figure 1.I89. The relationship between m .... /m . . . . . . . and temperature• Hatched and dotted areas represent • • , AUk.12 . AU[IJ~)2 the probable geochemical envlronment for typical Japanese gold-silver vein and auriferous vein deposits, respectively• A, mcl = 10, mK+ = 2, aH2S = l0 3, K-feldspar/K-mica/quartz equilibrium; B, reel = 1, mK+ = 0.2, aH2S = 10 -35, K-feldspar/K-mica/quartz equilibrium; C, mci = I, inK+ = 0•2, aH2S = 10 -2, K-feldspar/K-mica/quartz equilibrium; D, mcl- = 0.2, inK+ = 0•04, aH2S = 10 -2, K-feldspar/K-mica/quartz equilibrium; E, mcl = 0.2, inK+ = 0.04, aH2S = 10 -3, K-feldspar/K-mica/quartz equilibrium; F, mcl = 0.2, mK+ = 0.04, aH2S = 10 -2, K-feldspar/K-mica/quartz equilibrium• Thermochemical data for the calculations were taken from Helgeson (1969), Seward (1973), Drummond (1981), and Henley et al. (I984)• (Shikazono and Shimizu, 1987).

by the K-feldspar-K-mica-quartz assemblage which commonly occurs in Au-Ag vein deposits in Japan (Shikazono, 1974a). The values of HzS activity in ore fluids responsible for A u - A g veins are assumed to be 10 - 2 - 1 0 - 3 , which are estimated from sulfide mineral assemblage and chemical compositions of minerals (e.g., Fe content of sphalerite; e.g., Shikazono, 1974a). The transport of Au in the ore fluids responsible for the gold-quartz veins probably takes place as Au chloride complexes rather than as Au bisulfide complexes (Fig. 1.190). This suggestion is based on the fact that the temperature of formation of gold-quartz vein deposits is high (probably 250-350°C) and the chloride concentration of the ore fluids is high (probably more than 1 molal and less than 10 molal) based on the fluid inclusion studies (e.g., Nedachi, 1974; Shikazono, unpublished). Based on the experimental data on the solubility of Au due to chloride complexes (Henley, 1984) Au is highly soluble at high temperatures and/or in fluids of high chloride concentration. In fluids where precious metal transport is dominated by AgCI 2 and AuC12, the following reaction may be written to consider the compositional variation of native gold or electrum: (Au) + AgCI~- = (Ag) -t- AuCI~-

(1-88)

255

Miocene-Pliocene Hydrothermal Ore Deposits 5 4 3

A

C

~2

%1 t~

~o,o

-1 -2 -3 200

250

300

Temp. (°C)

Figure 1.190. The relationship between aAg/aAu of native gold and electrum and temperatures. A, aH2s = 10 2, mcl_ = 0.2, inK+ = 0.04, aAuC]£/aAu(HS)£= 102"°, K-feIdspar/K-mica/quartz equlibrium; B ' aH~S = 10 - 2 , m c i - = 0.2 , mK+ = 0.04 ' aAucI /aAu~HSa = 10 T M , C , aH-S = 10 -3 ~ mcl = 0.2 , z ~ /2 z mK+ = 0 . 0 4 , aauCl~-/ a Au(HS)~ = 10 2.0., D, a H 2 S = 10 -~ , me1 = 0 . 2 , mK+ = 0 . 0 4 , aAuCl~/aAu(HS)2 = 10 1.5., E, aAuCI~/aAu(HS)~= 1020; F, aAuci2/aAu(HS)~= 101"5. Calculations were m a d e a s s u m i n g K - f e l d s p a r / K m i c a / q u a r t z equilibrium for A, B, C, and D. T h e r m o c h e m i c a l data for the calculations were taken f r o m Helgeson (1969), Seward (1973, 1976), D r u m m o n d (1981), a n d Henley et al. (1984). (Shikazono a n d Shimizu, 1987).

From the equilibrium relation for equation (1-88) we obtain

aAg/aAu ~ ( a A g c 1 2 K1-88) /aAucI 2 (]-89) Curves E and F in Fig. 1.190 were drawn assuming that the value of N A g / E A u (NAg, total dissolved Ag concentration, NAu, total dissolved Au concentration) is approximately 102-1015 , which corresponds to the ratio for crustal rocks. These curves indicate that the aAg/aAu ratio of native gold or electrum does not change with temperature at a constant E A u / N A g ratio. It is noteworthy that the temperature dependency of the aAg/aAu ratio of native gold and electrum in AuC12 dominant ore fluids (Fig. l. 190). This difference may account for the range of variation and the Ag/Au ratios of native gold and electrum from gold-quartz vein and epithermal A u - A g vein deposits. However, it has to be noted that uncertainties of thermochemical data on gold chloride species might be large. As suggested by Henley et al. (1984) and Brown (1986), it is possible that Ag(HS)2 is dominant in low-salinity and near neutral ore fluids. Under this condition the Ag/Au ratios of electrum and native gold are controlled by the following reaction: (Au) + Ag(HS)~ = (Ag) + Au(HS)2

(1-90)

It is expected from the equilibrium ratios for equation (1-90) that the Ag/Au ratios of native gold and electrum are controlled by temperature and E A g / Z A u in ore fluids, although the equilibrium constant for equation (t-90) has not experimentally been determined. The E A u / E A g ratio in ore fluids responsible for epithermal A u - A g veins could be calculated from equation (1-87). The temperature of formation is taken from fluid

256

Chapter 1 N I

Fukushirr~ K.oganesawa

-~.~Oshima Tochigi Pref.~.'~'--. . ~ ~

Pref.

rune om, o

Nokoz~mo...{/2/~."

Nosu~ ~ ' . " .... ~_-~z~"'o

• / ' . - Y.,

[: "("Ibaragi Pref. . .) H .) i--Tarozawa Tokakum~... - . . ~ t~etsukuji Otori_.===:/V~N..t---'/:~-'.-. . ." ~--'~Hmaga BATO,~J - .'X.) • xx[ . ,'..qh-.~Shiozawa ~.."-'~.;y /..~'_-_-_-_-_-_-_-.~Kuryu

a UTSUNOMIYA Figure 1.I91. Locationof gold vein-typedeposits in the YamizoMountains. Dotted area represents the Yamizo Mountains except the Toriashi and Tsukuba blocks. Open circle: Au vein-typedeposits, solid circle: studied Au vein-typedeposits, square: city or town (Shimizu and Shikazono, 1987). inclusion studies (Enjoji and Takenouchi, 1976; Shikazono, 1985b) and the electrumsphalerite-argentite-pyrite assemblage (Shikazono, 1985d). The NaCI equivalent concentration of ore fluids is approximated from freezing data on inclusion fluids (Enjoji and Takenouchi, 1976), though the final melting temperature of fluid inclusion ice is also affected by CO2 concentration in epithermal ore fluids (Hedenquist and Henley, 1985). The pH values are estimated assuming the equilibrium among K-feldspar, K-mica, and quartz, which occur commonly in these deposits; this in turn allows a calculation of the activity of K +. The activity of HzS is estimated based on the equation showing the relation between the partial pressure of H2S gas and temperature for active geothermal waters (Giggenbach, 1980; Arn6rsson, 1985). Using the typical values of these variables and XAg = 0.5, which is a typical value for electrum from Au-Ag veins, E A u / E A g is calculated to be about 10 - I , which is similar to that of Broadlands geothermal water (Table 1.27). The above calculation is based on the assumption that AgCI~- is the predominant Ag species in ore fluids. However, Brown (1986) and Henley (1985) suggested that silver bisulfide complex (Ag(HS)2) could contribute significantly to the transportation of Ag in low-salinity geothermal waters (e.g., Broadlands, New Zealand). Thus, Ag bisulfide complex is also probable as dominant Ag species in ore fluids responsible for Au-Ag veins in Japan, though the salinity of ore fluids (0.1-0.3 molal) is generally higher than that of Broadlands geothermal water. If Au and Ag bisulfide complexes are the dominant Au and Ag species in Broadlands geothermal water and ore fluids responsible for Japanese gold silver veins, the Ag/Au ratio of ore fluids responsible for Japanese

257

Miocene-Pliocene Hydrothermal Ore Deposits

TABLE 1.27 EAu/NAg in ore fluidsestimated from chemical compositionof nativegold and electrum in active geothermal waters and in crustal rocks. XAg (X: atomic fraction) = 0.2 for auriferous veins and XAg = 0.5 for gold-silver 0.1, pH = 5.6; (2) 250°C, veins are assumed. Assumed values include: (1) 200°C, aH2S 10-35, mcl aHzS = 10 2, mcI =0.1, pH = 5.4; (3) 300°C, aH2s = 10-1, mcl- =0.i, pH = 5.4; (4) 300°C; (5) 300°C =

=

Au/Ag log (atomic ratio) (Au/Ag)

References

0.03~0.01 0.013~0.004

-1.6-2.0 -1.1-2.2

Wehdepohl(1978) Holland(1978)

0.i0 0.11 0.03

-0.99 - 1.00 - 1.6

Brown (1986) Henley et al. (1984) Koga (1957, 1961)

0.09

-1.0

(2)

0.09

- 1.0

(3)

0.08

- 1.1

Thermochemicaldata from Helgeson (1969), Seward (1973, 1976), Henley et aI. (1984) Thermochemical data from Helgeson (1969), Seward (1973, 1976), Henley et aI. (1984) Thermochemical data from Helgeson (1969), Seward (1973, 1976), Henley et al. (I984)

Auriferous vein (4)

0.007

--2.2

(5)

1.48 × 10 7

-6.8

(6)

0.007

-2.2

Crustal rocks (average) Seawater Active geothermal water

Broadlands (BR 22) Imperial Valley Beppu Ore fluids

Gold-silver vein (1)

Thermochemical data from Drummond(1981), Seward (I976) Thermochemical data from Helgeson (1969), Seward (1976) Thermochemical data from Drummond(198I), Seward (1976)

gold-silver veins may be similar to that of Broadlands geothermal water which is about 0.1 (Table 1.27). If AgCI~ and AuC12 are the predominant Ag and Au species, we can calculate E A u / E A g in ore fluids by using thermochemical data on gold chloride complexes by D r u m m o n d (1981), Helgeson (1969), and thermochemical mixing properties of A u - A g alloy by White et al. (1957)• The ratio of XAg/XAu (X: atomic fraction) for native gold or electrum from gold-quartz vein deposits is taken to be 0.25. The calculated value of m AuCl£/mAgCl2 based on thermochemical data on AuC12 by D r u m m o n d (1981) is about 10 -2. Using thermochemical data for Au chloride complexes compiled by Helgeson (1969), we obtain mAuci /mA_C1 which is very different from the A u / A g ratios of • 2 g 2 crustal rock and active geothermal waters (Broadlands, New Zealand; Imperial Valley Magmamax 1, California; Beppu, Japan), though the A u / A g values for the Imperial Valley Magmamax 1 and Beppu geothermal water do not reflect those in the deep fluid. ~348 of sulfide (pyrite) from this type of deposit varies widely, ranging from - 5 to +7%0 (Shikazono, unpublished), but average ~34S value is 0%0, suggesting igneous origin from deep part, but some from reduction of seawater sulfate in sediments. 313C values of

258

Chapter i

carbonates from these deposits are mostly from -5%0 to -7%0, suggesting igneous origin (Shikazono, unpublished).

1.6.3.2. Gold-quartz vein-type deposits in Yamizo Mountains, central Japan Shimizu and Shikazono (1987) studied gold-quartz vein-type deposits in Yamizo mountain, central Japan. Their study is summarized below. The Yamizo Mountains extend over the eastern part of Tochigi Prefecture, and the western part of Ibaragi Prefecture, central Japan, occupying an area ca. 20 km by 50 km. A lot of hypo/mesothermal (in the sense of Lindgren, 1933) gold vein-type deposits exist there (Fig. 1.191). The Yamizo Mountains are mostly occupied by Paleozoic-Mesozoic sedimentary rocks, mainly of Jurassic age (e.g., Sashida et al., 1982) with a small amount of intrusive granitoids of unknown ages. The Paleozoic-Mesozoic sedimentary rocks have been called the Yamizo Formation (Kanomata, 1961). It is composed chiefly of shale, sandstone, alternating beds of shale and sandstone, and a small amount of limestone and chert. The succession of the geologic units and geologic structure of the Yamizo Formation have been left pending due to complex structure such as upturned beds (Kasai, 1978) and submarine land sliding (Aono et al., 1985). Many small gold-quartz veins occur in the Yamizo mountains (Fig. I. 191). Some of them were already undertaken under the Satake clan in the Tokugawa era (more than 150 years ago). Most of them were productive during the 1930s to 1950s in several ten tons per month with a grade of ten to 50 grams per metric ton of gold and silver, respectively. Gold-quartz veins ranging in width from several centimeters to several ten centimeters are situated in shale and/or sandstone of the Yamizo Formation and rarely in granitoids. The strike and dip of the veins are various. For example, the veins of the Kuryu, Shiozawa, and Daigo deposits strike N40°W, N20-30°W, and N30°E, respectively, and dip 60°E, 30-50°W, and 20°W, respectively. The veins are composed mostly of quartz and a small amount of sulfide minerals (pyrite, pyrrhotite, arsenopyrite, chalcopyrite, sphalerite, and galena), carbonate minerals (calcite, dolomite) and gold, and include breccias of the host rocks with carbonaceous matters. Layering by carbonaceous matters has been occasionally observed in the veins. Banding structure, wall rock alteration and an evidence of boiling of fluids that are commonly observed in epithermal veins have not been usually found. Gold is present either as isolated grains in quartz or in association with sulfide minerals. In the case of the latter, gold is generally in direct contact with pyrite and/or galena. Gold is irregular in shape with 0.01 to l mm in size. Some of gold grains can often be observed by naked eyes. Analytical results on gold are shown in Table 1.28. Chemical compositions of gold are relatively homogeneous and low Ag contents. Frequency histograms for Ag atomic% of gold (Nag) from the Yamizo Mountains and other regions are summarized in Fig. 1.192. Values of NAg of gold from the Yamizo Mountains range from 5.9 to 18.4, and this is almost the same as those from the other regions (Fig. 1.192). However, NAg of hypo/mesothermal gold vein-type deposits in metamorphic rocks tends to be slightly lower than that in sedimentary rocks.

Miocene-Pliocene Hydrothermal Ore Deposits

259

TABLE 1.28 Representative chemical compositions of gold from the Yamizo Mountains, Japan (Shimizu and Shikazono, 1987) Ore deposit

Weight % Au

Atomic % Ag

Cu

Total

Au

Ag

Cu

Ag/Au

Kuryu

96.70 96.31 96.15 96.13 96.20 95.66 95.89 96.05 95.36 95.24

3.31 3.41 3.50 3.6I 3.77 3.79 3.90 3.97 4.12 4.26

0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.00

100.01 99.72 99.65 99.74 99.97 99.46 99.79 100.02 99.48 99.50

94.11 93.93 93.78 93.58 93.33 93.22 93. l0 92.98 92.69 92.45

5.89 6.07 6.22 6.42 6.67 6.74 6.90 7.02 7.31 7.55

0.00 0.00 0.00 0.00 0.00 0.04 0.00 0.00 0.00 0.00

0.063 0.065 0.067 0.069 0.071 0.072 0.074 0.076 0.079 0.082

Shiozawa

92.83 92.10 92.24 92.15 92.22 92.62 92.20 91.96 91.55 91.64

7.00 7.06 7.11 7.16 7.24 7.37 7.37 7.45 7.48 7.55

0.05 0.05 0.02 0.04 0.06 0.03. 0.06 0.06 0.04 0.00

99.88 99.21 99.37 99.35 99.52 100.02 99.63 99.47 99.07 99.19

87.76 87.60 87.61 87.47 87.32 87.24 87.12 85.98 86.92 86.92

12.09 12.25 12.33 12.42 12.52 12.67 12.71 12.86 12.96 13.08

0.15 0.15 0.06 0.11 0.17 0.09 0.17 0.17 0.11 0.00

0.138 0.140 0.141 0.142 0.143 0.145 0.146 0.150 0.149 0.150

Daigo

92.32 91.68 91.65 91.77 91.54 91.64 91.27 91.41 91.33 91.16

7.54 7.5I 7.69 7.78 7.83 7.93 8.04 8.11 8.18 8.37

0.02 0.01 0.02 0.02 0.02 0.03 0.02 0.04 0.00 0.04

99.88 99.20 99.36 99.57 99.39 99.60 99.33 99.56 99.51 99.57

86.97 86.96 86.66 86.55 86.44 86.28 86.10 85.96 85.95 85.54

12.97 13.00 13.28 13.40 13.50 13.63 13.85 13.93 14.05 14.35

0.06 0.04 0.06 0.06 0.06 0.09 0.06 0.11 0.00 0.11

0.149 0.149 0.153 0.155 0.156 0.158 0.161 0.162 0.163 0.168

Saigane

89.73 89.40 89.56 88.69 89.19 89.12 88.42 88.99 88.30 88.34

9.68 9.71 10.19 10.28 10.46 10.54 10.66 10.78 10.87 10.91

0.07 0.06 0.06 0.06 0.08 0.08 0.07 0.08 0.08 0.09

99.48 99.17 99.81 99.03 99.73 99.74 99.15 99.25 99.25 99.40

83.38 83.31 82.67 82.39 83.17 82.05 81.79 81.46 81.46 81.31

16.42 16.52 17.17 17.44 17.59 17.72 18.01 I8.30 18.30 18.44

0.20 0.17 0.16 0.16 0.24 0.24 0.20 0.24 0.24 0.25

0.197 0.198 0.208 0.212 0.214 0.216 0.220 0.225 0.225 0.225

260

Chapter 1 N=144

In Metamorphic R. Suwa, Kinkef, Amo, Hashidate

50 4,0 30 20 10

El N: 1[

70

Sh

6O 5O D

o 40

o-

~-

In Sedimentary R. Yamizo Mtns, Kuryu, Daigo Shiozawa, Saigane

30

u_ 2O

10 ~

~Sa

S. Koma Region Ho Gohaku, Koei

[

N=1'57j~ 30. A~erlgel)

' Kitakami Mtns. Oya, Shishiori,

20

Hayachine, Kohoku Nedachi(1974) Yamaoka(1981)

10

0

10 20

50

NAg

100

Figure 1.192. Frequency (no. of analyses) histograms for Ag content (in atomic %) of gold from "hypo/ mesothermal" gold vein-typedeposits in Japan. K: Kuryu deposit, Sh: Shiozawa deposit, D: Daigo deposit, Sa: Saigane deposit. Data are from: the Yamizo Mountains:Shikazonoand Shimizu (1988a); the South Korea Region: Shikazono and Shimizu (1987); the Kitakami Mountains: Nedachi (1974); Yamaoka (1981), Abe (1981) and Shikazono and Shimizu (1987); metamorphicregions: Shikazono and Shimizu (t987).

In the Yamizo Mountains, the correlation between NAg and a kind of the host rocks is found: NAg of gold from gold-quartz veins in shale is lower than that in sandstone. Gold negatively correlates with 834S of the associated sulfide minerals from the veins, and that 334S of sulfide minerals in sandstone is lower than that in alternating beds of shale and sandstone, probably than that in shale.

Miocene-Pliocene Hydrothermal Ore Deposits

261

Homogenization temperatures of fluid inclusions mainly in quartz closely associated with gold and sulfide minerals in hypo/mesothermal gold vein-type deposits in Japan are in a range from 250 ° to 350°C (Shikazono and Shimizu, 1987), higher than those in epithermal gold-silver vein-type deposits in Japan. Iron contents of sphalerite and thermochemical data by Barton and Toulmin (1966) and Barton and Skinner (1979), possible sulfur fugacity (fs2) of the ore fulids can be estimated: log fs2 = - 1 5 at 250°C to - 3 at 350°C. This means that fs2 for hypo/mesothermal gold vein-type deposits is lower than that for epithermal gold-silver vein-type deposits in Japan. Factors in controlling chemical compositions of gold in equilibrium with the ore fluids are temperature, pH, concentration of aqueous H2S and C1- in the ore fluids, concentration ratio of Au and Ag species in the ore fluids, activity coefficient of Au and Ag components in gold, and so on (Shikazono, 1981). In the Yamizo Mountains, as a result, Ag/Au ratios of gold are correlated with a kind of the host rocks and sulfur isotopic compositions of the deposits. This correlation could be used to interpret Ag/Au ratios of gold. There are two possibilities here to explain this correlation. One is that isotopically heavy sulfide sulfur derived from seawater sulfate was fixed in shale because reducing agency of shale with carbonaceous matters is thought to be stronger than that of sandstone. The ore fluids extracted this sulfur. Gold of low NAB precipitated in shale like the Kuryu deposit under more reducing environment than in sandstone like the Saigane deposit. Shikazono and Shimizu (1987) concluded that Ag contents of gold precipitated from low-salinity fluids is higher than that prediction and the relationship between Nag of gold and salinity of fluid inclusions estimated from freezing temperature data. Therefore, another interpretation is that Nag of gold from shale-hosted deposits is lower than that from sandstone-hosted deposits, because shale is expected to be richer in C1- mainly due to adsorption by clay minerals included in shale than sandstone. It is concluded that, in either case, in situ interaction of the ore fluids with the host rocks controls the chemical compositions of gold in the Yamizo Mountains.

1.6.4. Hot spring-type gold deposits Hot spring-type gold deposits (Nansatsu-type by Urashima et al. (1981, 1987), high sulfidation-type by Hedenquist (1987), epithermal Au disseminated-type) are distributed in the Nansatsu district of southern Kyushu (Fig. 1.193). The deposits (Kasuga, Akeshi, Iwato) were formed at Pliocene age (5.5-3.7 m.y.) in the calc-alkaline volcanic rocks of nearly same age (Togashi and Shibata, 1984). The deposits, which are similar to Nansatsu-type deposits, occur in Southwest Hokkaido (Date, Hakurhu). The deposits are characterized by conspicuous alteration zoning from the centre (orebody) to margin (Tokunaga, 1955; Doi, 1972; Urashima et al., 1981, 1987). They are siliceous zone, alunite zone, kaolinite zone, sericite zone and montmorillonite zone. In the siliceous body, electrum and Cu minerals (enargite, luzonite, covelline), and native sulfur occur. The Ag content of electrum is lower (0.0-5.3 wt%) than that from epithermal A u - A g vein-type deposits (Fig. 1.194) (Shikazono and Shimizu, 1987). Low Ag content of electrum and sulfide mineral assemblage (enargite, native sulfur, covellite, pyrite) indicate high i s 2 condition (Fig. 1.194).

262

Chapter I

J~ f J

~ ~

\

o

~._~v-

~

[ [

• Hot springtype ~ EpithermalAu-Agvein-type ~ Active volcano

o

- ~

I,

~

20

~,

4o

I
/

o o~uo~i

)

/

o,.o..=o

(

j

/

Kagoshima

Mt.Kaimondake

31N

~

I

l 131E

Figure 1.193. Locality map of the hot spring-type deposits in Kyushu. 4O

=~ 3o

g

2O

10

0

0

20

40

60

80

1oo NAg

Figure 1.194. Frequency histogram for the Ag content of eIectrum from epithermal Au disseminated-type deposits in Japan (Iwato and Kasuga). (Shikazono and Shimizu, 1987).

Miocene-Pliocene Hydrothermal Ore Deposits

263

/

Steam-heated alunite fluid / / / I .i~<-~/

•

SMOW

K ~

I 0 Alunite

•

0 Clays

.~'V Alunitefluid '@ ~ O

-50

{ 1._.: ~s~ ~

volcanic fluids

O-shift

a

Clayfluid 0 ~ 0 0

-100

.2

I

I

t

I

Residual silica • I

(

I

H

•

silica

Kasuga

and

0

O Iwato

Deepveins at Kasuga

(~

Peripheralveins Kago veins

~

vein

0

quartz

Other localities Calculated fluids

I

-10

•

0 Residual

Quartz veins in residualsilica

I

8180

+10 %o

I

+20

Figure 1.I95. Diagram showing the results of isotopic analysis of residual silica (now recrystaltized to quartz) and quartz crystals or veins (3180 only, plotted on the bottom), and clays and alunite (~D vs. 3180, plotted on the top). The calculated composition of the fluids in equilibrium with these minerals, plus the compositions of local meteoric water (small open square) and high-temperature volcanic fluids (large open square) are also shown (Matsuo et al., 1974; Matsubaya et al., 1975). A least-squares fit line (dashed) through the hypogene alunite fluids and the high-temperature volcanic fluid end member extends to an isotopic composition similar to present local meteoric water and identical in gD values to the clay fluid. Thus the isotopic composition of alunites record mixing of magmatic and meteoric waters, whereas the marginal clays (and possibly the vein quartz and residual silica quartz) indicate that meteoric water dominated on the margins of the quartz bodies, variably shifted in oxygen isotope composition due to exchange with wall rocks. Some of the deep fluids responsible for vein formation are identical in isotopic composition to the 180-enriched fluid that equilibrated with the residual silica quartz, whereas other samples indicate a greater proportion of unshifted meteoric water. One vein alunite from Iwato has an isotopic composition which indicates that it formed in a steam-heated environment after evaporation of local ground water (fluid composition calculated for a steam-heated temperature range of 100-180°C). K and I denote Kasuga and Iwato, respectively. (Hedenquist et al., 1994).

264

Chapter 1

Hedenquist et al. (1994) indicated that homogenization temperature of postmineralization of the Akeshi deposit and salinity of ca. 1 wt% NaCl eq. and the presence at times of a two-phase fluid in the centre of ore zones indicating boiling of ore fluids. 834S of sulfides (pyrite, enargite, luzonite) and sulfate(alunite) are -5%0 to 0%0 and +24%0 to +35%0, respectively (Shikazono, 1987b; Hedenquist et al., 1994), suggesting that sulfide sulfur came both from igneous and sedimentary sources and sulfate sulfur is generated by the hydrolysis reaction of 4 S O 2 -k- 4H20 ~ H2S + 3 H 2 8 0 4 . SO2 gas may come from volcanic gas. 3D and 8180 of ore fluids were estimated from these data on quartz, clay minerals and alunite (Fig. 1.195) (Hedenquist et al., 1994), indicating higher 8D and 8180 values compared with epithermal Au-Ag vein ore fluids. These isotopic data (834S, 8D, 8180), fluid inclusion data and alteration studies suggest that (1) the gas-rich (HCl, SO2, H2S) vapor phase separated from magma ascent to the surface and condense into meteoric water, (2) this acid chloride-sulfate fluid leached elements from the host rocks, forming a permeable zone of mixing, (3) then, the dense metal-rich liquids ascent into the leached zone and precipitate Cu-minerals and Au (Hedenquist et al., 1994). The fo2-PH conditions for this type of deposits and other types of epithermal deposits and hydrothermal alteration are shown in Fig. 1.196 (Shikazono and Aoki, 1981). It is worth noting that the f o 2 - P H ranges for the epithermal Au deposits (hot spring-type, Te-type, and Se-type) lie between the points A (high fo2, and low pH) and B (low f o e 0

Log fo 2 'o~ vm, e 'r

High sulfidation type ,

-30 • s,

-35

I

I

/ll

~AIIKal , / 1 , Y"

."--. ,,

l./I

/'

/ s -

Bn+Py Cp

I . ~ t - " ",. I "3-.

D,,

--I---I Kal

1

I

I

2

3

4

I I

~.

I

~

Po

Hm Mt

' ""!"';"c'--

I 1

~40

Low sulfidation type

/I

Se I

5

4

i

1

1

e4~CO

~

-r ', -r Kf ' 1

I

I I

2~

'

1

I

I~

|

I

6

7

8

9

pH

Figure 1.196. fo2 pH ranges for hot-spring-type deposits and low sulfidation-type deposits. Temperature = 250°C, £S = 0.01 mol/kg H20, ionic strength = I. Ka: kaolinite, AI: alunite, SI: liquid sulfur, Kf: K-feldspar, Hm: hematite, Mt: magnetite, Py: pyrite, Po: pyrrhotite, Bn: bornite, Cp: chalcopyrite.

265

Miocene-Pliocene Hydrothermal Ore Deposits Assumed direction of the Pacific Plate Deviation form the north 90 ° 70 ° 50o 30 °

Ma 0

O

2 -

f

N

3-

5-

Q.R intrusion

~

1-

4 -

EpithermaI system

.

- Kobui H.S.S. .

O

- Tokiwamatsu H.S.S.

--

Q.R

Hokko-Minami LS.S.

- - Yamato L.S.S.

6--Mitsumoriyama

7 -

Q.R

H.S.S.

89-

N

10-

H.S.S.: High sutfidation s y s t e m L.S.S.: L o w sulfidation s y s t e m Q.R: Q u a r t z p o r p h y r y N: n o r m a l s u b d u c t i o n , O: o b l i q u e s u b d u c t i o n Figure 1.197. Subduction mode of the Pacific plate beneath the Northeast Japan arc, and style of the epithermal systems in the Kameda peninsula, Hokkaido. Normal, oblique and highly oblique subduction are tentatively defined by the angle (c0 between the assumed subduction direction of the Pacific plate and the trend of the Northeast Japan arc. Normal subduction (60 ° < c~ = 90°), oblique subduction (30 ° < c~ = 60 °) and highly oblique subduction (0° < c~ = 30°). Highly oblique subduction has not occurred along the Northeast Japan arc since 11 Ma. (Watanabe et al., I996).

and high pH) in Fig. 1.196. These difference in physicochemical nature of fluids can be explained by space-dependent evolution of a hydrothermal system relative to the centre of volcanic activity (Shikazono and Aoki, 1981; Shikazono, 1985a; Shikazono et al., 1990; Sillitoe, 1991) or a time-dependent evolution of a hydrothermal system (early-stage; high sulfidation-type, late-stage; low sulfidation-type) (Aoki et al., 1993). For example, Shikazono (1985a) has shown that advanced argillic alteration (Ugusu silica deposit in west Izu Penninsula, central Honshu) occurs at the centre and

266

Chapter 1

stratigraphically higher position in the mine district, while epithermal Au-Ag veintype deposits at the margin and lower position (Fig. 1.76). Shikazono et al. (1990) compared the characteristic features of Te-type with Se-type and suggested that the Te-type formed at the position closer to volcanic centre in a hydrothermal system than Se-type (Fig. 1.123). It is probable that the difference in fo2 and pH in several types of epithermal Au deposits is due to the different ratio of meteoric water component (low fo2 and high pH) to magmatic water (or igneous water) component (high fo2, low pH), related to the position from volcanic centre. Watanabe et al. (1996) and Watanabe and Ohta (1999) pointed out that the high sulfidation (advanced argillic alteration) and low sulfidation (sericitic and propylitic alteration) system in Southwest Hokkaido (Minamikayabe area, Jozankei-Zenibako district) occurred during normal and oblique subduction of the Pacific plate beneath the Northeast Japan arc, respectively (Fig. 1.197). These studies suggest that the difference in two styles of epithermal gold mineralization (epithermal vein (low sulfidation-type) and hot spring-type (high sulfidation-type)) is caused by the subduction mode of the plate and not by the temporal and spatial differences in a given hydrothermal system. Figure 1.198 shows grade-tonnage relationship for low sulfidation-type (Se-type) and high sulfidation-type (Te-type and Akeshi-type) in Japan and deposits in Circum Pacific region summarized by Hedenquist et al. (1996). This clearly indicates that grade and tonnages for Akeshi-type and Te-type in Japan are low and small, compared with low sulfidation-type in Japan. It is obvious that low sulfidation-type is dominant Au-Ag deposits in Japan. The plot of K20 versus SiO2 for volcanic rocks thought to be genetically related to Au mineralization indicates that high sulfidation deposits (hot spring-type deposits appear

ll-t

100

I

Illllll

I Xl~

"'.z°o, -. 7 "za

"

-.e %

""-

]0

o", ",

eO OO

lllllll

I

I

~

I

[[IFF[

I

I

F lIIII

• Low sulfidation • High sulfidation

%0

,.L.,

. ' ~ "" ~

•

•

""

" "•-o RO • • "~'oEmMc oIB~'" "PV La "e. • • p p

] 0.l

[

"" f.ooa "- z Hi "'.

1

10

,e. A 100

RM ". 1000

Tonnage, Mt Figure 1.198. Grade-tonnage plot for epithermal Au deposits. LS: Ba, Baguio (Philippines); CC, Cripple Creek; Mc, McLaughlin and RM, Round Mountain (USA); Era, Emperor (Fiji); Hi, Hishikari (Japan); Ke, Kelian (Indonesia); La, Ladolam and Po, Porgera Zone VII (Papua New Guinea). HS: Ch, Chinkuashih (Taiwan); EI, EI Indio (Chile); Go, Gotdfieid and PP, Paradise Peak (USA); Le, Lepanto (Philippines); PV, Pueblo Viejo (DominicanRepublic); Ro, Rodalquilar (Spain) (Hedenquist et al., 1996).

267

Miocene-Pliocene Hydrothermal Ore Deposits 8

i

J

i

i

• LS(calc-alkalic) • LS(shoshonitic-alkalic) • HS(caIc-alkalic)

6

jO 4~

4

4 ~

•~

• ~0

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Figure 1.199. Plot of K20 versus SiO2 for volcanic rocks thought to be genetically related to Au mineralization. High sulfidation (HS) deposits appear to be associated with a narrow range of igneous rock composition, dominated by dacite and andesite. Low sulfidation (LS) deposits are largely associated with andesite and rhyolite, but major deposits such as Ladolam, Porgera and Cripple Creek are associated with shoshonitic and alkalic rocks (Hedenquist et al., 1996).

to b e a s s o c i a t e d w i t h a n a r r o w r a n g e o f i g n e o u s r o c k c o m p o s i t i o n , d o m i n a t e d b y d a c i t e a n d a n d e s i t e , w h i l e l o w s u l f i d a t i o n d e p o s i t s are w i t h a n d e s i t e a n d r h y o l i t e , b u t m a j o r d e p o s i t s (e.g., L a d o l d m , P o r g e r a a n d C r i p p l e C r e c k ) are a s s o c i a t e d w i t h s h o s h o n i t i c a n d alkalic r o c k s ( H e d e n q u i s t et al., 1996) (Fig. 1.199).

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295

Chapter 2

Present-day Mineralization and Geothermal Systems in and around the Japanese Islands

2.1. Subaerial geothermal system and mineralization Since the start of operation of the first geothermal power station at Matsukawa, Northeast Japan in 1966, many geothermal areas were explored for developing geothermal energy in the Japanese Islands. The geothermal power stations operating and in development include Matsukawa, Otake, Ohnuma, Onikobe, Hachobaru, Kakkonda, Mori at Nigorikawa, Okuaizu, Uenotai, Fushime, Takigami, Ogiri at Kirishima and Oguchi (Fig. 2.1 ). Various aspects (geology, geophysics, mineralogy, hydrothermal alteration, stable and radiogenic isotopes, chemistry of geothermal waters) of these active geothermal areas have been investigated during the past three decades. These studies are briefly summarized below, focusing on the characteristic features of present-day base metal and precious metal mineralizations associated with the geothermal activities.

2.1.1. Chemical compositions of geothermal waters controlled by hydrothermal alteration mineral assemblage Previous studies clearly indicated that the chemical compositions of geothermal waters are intimately related both to the hydrothermal alteration mineral assemblages of country rocks and to temperature. Shikazono (1976, 1978a) used a logarithmic cation-C1concentration diagram to interpret the concentrations of alkali and alkaline earth elements and pH of geothermal waters based on thermochemical equilibrium between hydrothermal solution and alteration minerals. Several studies of geothermal areas have shown that alteration assemblages regularly change with increasing depth and temperature. These changes have been observed in Otake (Hayashi, 1973), Nigorikawa (Yoshida, 1991), Sumikawa (Sakai et al., 1986), Okuaizu (Nitta et al., 1991) and Onikobe (Seki, 1991). Alteration minerals at the greatest depth in areas of relatively high temperatures (200-300~ are quartz, chlorite, mica, anhydrite, K-feldspar, calcite, pyrite, albite, wairakite, laumontite, and minor amounts of epidote and prehnite. Hence, it is assumed that the minerals in equilibrium with geothermal waters are albite, K-feldspar, muscovite, quartz, calcite, anhydrite, chlorite and wairakite.

296

Chapter 2

NI 3

e

9 1~

8 "22 24'" r ~

"

"~20~~ ~ / 21 '~23 ~2s

170

Pacific Ocean o 170

30okra

Fig. 2.1. Map of Japan showing the location of geothermal drilling for the year ending 31 March 1990 (from the Japan Geothermal Energy Association). Open circles denote geothermal survey drilling sponsored by NEDO, while closed circles are for systems with drilling sponsored by private enterprise. Open squares designate areas where drilling was conducted by private companies under partial sponsorship of NEDO. 1: Akan, 2: Noboribetsu, 3: Mori, 4: west of Kakkoda, 5: Matsukawa, 6: Ohnuma and Sumikawa, 7: Kakkonda, 8: east of Lake Tazawa, 9: west of Mt. lwate, 10: Minase, I1: Uenotai, 12: Onikobe, 13: Mogami Akakura, 14: east of Obanazawa, 15: Inawashiro, 16: Okuaizu, 17: Hachijojima Island, 18: upper Oita River, 19: Takigami, 20: Otake and Hatchobaru, 21: Oguni, 22: Hishikari, 23: Kirishima (Ogiri), 24: Kirishima International Hotel, 25: Fushime (Hedenquist, 1991 ).

In geothermal waters, the condition of electroneutrality must be fulfilled: ~Zcimci = ~Zajmaj

(2-1)

where Zci and mci refer, respectively, to the charge and molality of cation i, and Zaj and maj to the charge and molality of anion j, respectively. It is found that Na + and C1- are the predominant cation and anion in geothermal waters of high temperatures (200-300~ If the concentrations of the other species are negligible, compared with those of Na + and CI-, equation (2-1) is approximated by, mNa+ = m C l -

(2-2)

where m is concentration. This relationship is shown in Fig. 2.2 in which data on Na + and C1- concentrations of geothermal waters in Japan and in other countries are plotted. Albite and K-feldspar are commonly observed to coexist. If the following reaction is in equilibrium, K-feldspar + Na + = albite + K +

(2-3)

297

Present-day Mineralization and Geothermal Systems Log tuna+ 1C

0

.c

-1

-2

W

2'

9

cS e

9

"D

"

Log mc~i

-2

!

-1

I

!

0

1

Fig. 2.2. Relation between the Na + and C1- concentrations of geothermal waters and of inclusion fluids. The solid line indicates the condition of electroneutrality approximated by the equation mNa+ - - m c 1 - . Solid and open circles mean the chemical analytical data on inclusion fluids and geothermal waters, respectively. S = Salton Sea; R -- Reykjanes; W = Wairakei; B = Broadlands; O = Otake; H -- Hveragerdi; C -- Climax; D = Darwin; P -- Providencia (Shikazono, 1978a).

Then the equilibrium constant, K2-3, is expressed as, K2-3 = (aAb a K + ) / ( a K - f a N a + )

(2-4)

where a A b , aK_f, aK+ and aNa+ are activity of albite component in feldspar, activity of K-feldspar component in feldspar and activities of K + and Na+, respectively. Combining (2-4) with (2-2), we obtain, logmK+ = logmcl- + log(C1)

(2-5)

where C1 = (YNa+aK-fK2-3)/(YK+aAb)and gi = the activity coefficient of the ith species. In a plot of log(mK+) vs. log(mcl-) (Fig. 2.3), a slope of about +1 is obtained. Fig. 2.3 was constructed using a K2-3 value at 250~ extrapolated from hightemperature data by Orville (1963), Iiyama (1965) and Hemley (1967). Ion activity coefficients were computed using the extended Debye-Htickel equation of Helgeson (1969). The values of effective ionic radius were taken from Garrels and Christ (1965). In the calculation of ion activity coefficients, ionic strength is regarded as 0.5 (mNa+ -Jr-mc1-) ( - mCl-). The activity ratio, aK-f/aAb, is assumed to be unity. As in the case of K +, if the concentration of alkali elements (Cs +, Rb +, Li +) are controlled by feldspars, it is convenient to take into account of the following exchange reaction, albite + X + = X-feldspar + Na +

(2-6)

where X denotes an alkali element. The equilibrium constant of reaction (2-6) is K2-6 ~- ( a x A 1 S i 3 0 8 a N a + ) / ( a A b a X + )

(2-7)

298

Chapter 2 / C/S

Log mK+

S.P"

.

c pcc

2

-3 -

!

-2

i

-1

!

0

Log mcl1

Fig. 2.3. Relation between the K+ and C1- concentration of geothermal waters and inclusion fluids. The solid line defines the equilibrium condition between the solution and the assemblage albite-K-feldspar at 250~ For symbols used, see caption to Fig. 2.2. (Shikazono, 1978a). From (2-2) and (2-7) we obtain, log(mx+) = logmcl- + logC2

(2-8)

where C2 -- (aXAISi308 YNa+ K2-6)/(aAbYX+). The concentrations of Rb and Cs in feldspars in geothermal areas have not been studied. However, if the concentration ranges are similar to those observed for feldspars in common igneous rocks, we can construct curves to show the relationship of the concentration of an alkali element X + and the C1- concentration of geothermal waters (Fig. 2.4 and Fig. 2.5). It is assumed that the concentrations of Cs and Rb in feldspars range from 10 -3 to 10 -4 wt% and 10-1 to 10 -2 wt%, respectively. The exchange reactions between aqueous solution and feldspars for Cs and Rb were studied by Lagache and Sabatier (1973) and Volfinger (1975). The partition coefficients below 400~ are not available. Therefore, using the value for 400~ by Volfinger (1975) and the relationship between the K + and C1- concentrations shown in Fig. 2.3, the relationship between the alkali element (Rb +, Cs +, Li +) concentrations and C1- concentration was derived. It is apparent from Figs. 2.4 and 2.5 that a linear relationship with a slope approximately +1 exists between the alkali elements concentration and the CI- concentration. The partition coefficient for Li between feldspar and aqueous solution has not been determined, but it is predicted that in a l o g ( m L i + ) log mcl- diagram a slope of approximately +1 would also be obtained, if the exchange reaction of Na + and Li + between aqueous solution and feldspar is in equilibrium (Fig. 2.6). If muscovite, K-feldspar and quartz are saturated with geothermal waters, the following reaction can be written: 3 K-feldspar + 2 H + = muscovite + 6 quartz + 2 K +

(2-9)

299

Present-day Mineralization and Geothermal Systems Log mcs+

(1)

-4

(2)

i

-6

H

Log mcjFig. 2.4. Relation between the Cs + and C1- concentrations of geothermal waters. Solid lines 1 and 2 indicate the equilibrium condition between feldspars and solution, assuming that the Cs content of the feldspar is between 10 -3 wt% and 10 -4 wt%. For symbols used, see caption to Fig. 2.2 (Shikazono, 1978a).

Log mRb+ -3

-

.(1) -4 (2)

,W

-5

-6

|

-2

I

-1

I

Log mcr

-0

Fig. 2.5. Relation between the Rb + and C1- concentrations of geothermal waters. Solid lines 1 and 2 indicate the equilibrium condition between feldspars and solution, assuming that the Rb content of the feldspar is between 10 -1 wt% and 10 -2 wt%o. For symbols used see caption to Fig. 2.2. (Shikazono, 1978a).

Using the equilibrium constant for (2-5) and (2-9), the relationship between pH and C1- concentration can be derived as shown in Fig. 2.7. The line shown in Fig. 2.7 has a slope of approximately - 1 , indicating that the pH of the geothermal waters decreases with increasing C1- concentration.

Chapter 2

300

Log m Li+

-2 W o

-3

0 o

-4

I

,Log mc~-

!

-2

-1

0

Fig. 2.6. Relation between the Li + and C1- concentrations of geothermal waters (Shikazono, 1978a).

pH oO oH

?

6

s 5

~ -

'

,

-1

0

Log mc~-

Fig. 2.7. Relation between the pH and CI- concentration of geothermal waters. The solid line indicates the albite-K-feldspar-muscovite-quartz-solution equilibrium at 250~ For symbols used see caption to Fig. 2.2. (Shikazono, 1978a).

If calcite is saturated with geothermal waters, the following reaction is written" CaCO3 + 2 H + -- Ca 2+ 4- H2CO3

(2-10)

Combining the equilibrium relation for reaction (2-10) and the relationship between pH and the C1- concentration derived already, we can derive the dependence of

301

Present-day Mineralization and Geothermal Systems

Log mca2+ 1

(1) //(2)

0

-5 -6

-;

o

Log mc,-

Fig. 2.8. Relation between the Ca 2+ and C1- concentrations of geothermal waters and inclusion fluids. Solid lines indicate: (1) albite-K-feldspar-muscovite-quartz-calcite-solution equilibrium at aHzCO3 = 10-2"5; (2) albite-K-feldspar-muscovite-quartz-calcite-solution equilibriumn at an2co3 = 10-2; (3)anhydrite-solution at ES0 (total dissolved sulfate concentration) = 10-3; and (4) anhydrite-solution equilibrium at ES0 -- 10 -2. For symbols used see caption to Fig. 2.2 (Shikazono, 1978a).

Ca 2+ concentration on C1- concentration (Fig. 2.8). At constant temperature and activity of H2CO3, log(mca2+) plots linearly as a function of the log mc1- with a slope of about +2. The saturation curve for anhydrite can be constructed from the following reaction: CaSO4 + Na + -- Ca 2+ + NaSO 4

(2-11)

From this equilibrium relation and (2-2), we obtain, logmca2+ = logmc1- + logC3

(2-12)

where C3 = (YNa+K2-11)/(mNaSO4 FNaSO4 YEa2+). The anhydrite saturation curve is drawn at constant total sulfate concentration and temperature. Chlorite commonly coexists with albite, K-feldspar, muscovite and quartz. The equilibrium relation among these minerals and aqueous solution is given in Fig. 2.9. The logmMg2+ plots linearly as a function of logmc1- with a slope of about +2. Generally, considerable amounts of Fe are contained in chlorite. The Fe content of chlorite varies with temperature, and oxygen and sulfur fugacities, if it is in equilibrium with Fe-oxides and Fe-sulfides. Therefore, the Mg 2+ concentration will vary not only as a function of the C1- concentration and temperature but also with fo2 and fs2 (Shikazono and Kawahata, 1987). If these variables change considerably in the geothermal system, the Mg 2+ concentration deviates from the equilibrium curve in Fig. 2.9. For the construction of Fig. 2.9, activity of Mg-chlorite is taken to be unity. The chemical compositions of geothermal water plot in Figs. 2.3-2.9. All data plot near the equilibrium curves. This indicates that the geothermal waters considered

302

Chapter 2 Log

mMg2+

..9/

oCC~ D

-2 -3

79

-4

-5

B .W ~

-6 I

-2

I

-1

!

Log mc,-

0

Fig. 2.9. Relation between the Mg 2+ and Cl- concentrations of geothermal waters. The solid line indicates albite-K-feldspar-muscovite-quartz-Mg-chlorite-solution equilibrium at 250~ For symbols used see caption to Fig. 2.2 (Shikazono, 1978a).

are nearly in equilibrium with the alteration assemblages commonly found in geothermal reservoirs. The above discussions are based on the assumption of constant temperature. However, temperature varies widely. The chemical compositions of geothermal waters intimately relate to temperature. For example, the correlation between Na/K ratio in geothermal waters and temperature has been interpreted as indicating that this ratio is controlled by albite and K-feldspar (White, 1965; Ellis, 1969, 1970). The Na/Li ratio of geothermal waters decreases with increasing temperature and has been used as a geothermometer (Fig. 2.10) (Fouillac and Michard, 1981), suggesting that this ratio is controlled by feldspar-solution equilibrium (Shikazono, 1978a). Therefore, it is worth considering the effect of temperature on the chemical composition of geothermal waters in Japan. Chiba (1991) summarized the chemical compositions of Japanese geothermal waters (Tables 2.1 and 2.2), derived the relationships between cation/proton ratios, pH and temperature based on thermodynamics and compared his calculated results with the chemical compositions of geothermal waters (Figs. 2.11-2.13; Tables 2.1 and 2.2). His results are consistent with the view that geothermal waters are saturated with common alteration minerals (K-feldspar, K-mica, quartz, albite, anorthite (or Ca-zeolites), Mg-chlorite). The partial pressure of CO2 is controlled by the reaction involving calcite, K-bearing silicates (K-mica, K-feldspar) and albite or Ca-zeolites (wairakite, laumontite). 2.1.2. N a - K - C a geothermometer It is well known that the atomic ratio of Na/K in natural waters depends on temperature (White, 1965, 1968; Ellis, 1969, 1970). This relationship has been also

303

Present-day Mineralization and Geothermal Systems

y

oETB

a

~q 3 o _J

G#

~

!

I

2

I

3 I/Tx

4 iO

-3

Fig. 2.10. Relationship between Na/Li and temperature (Fouillac and Michard, 1981). KCS: Kettelman interstitial solutions; concentrated samples. KDS: Kettelman interstitial solutions, diluted samples. CB: Cesano Brine. ETB: E1 Tatio Brine. WBIS: Basalt in interaction resulting solution. H: Hveragerdi and G: Geysir.

TABLE 2.1 Summary of fluid compositions of Japanese geothermal systems (Chiba, 1991) Geothermal system

Well name

Nigorikawa Kakkonda Sumikawa Okuaizu Takigami

ND-7, NF-1 KA-la, KB-3, KC-3, KE-1, K7-2 S-4, KY-2 1t NE-2, NE-3, TT-1, TT-2, TT-7, TT-8, TT-14 H-15, H-20 KEI-3, KEI-5, KEI-6, KEI-7, KEI-17, KEI-19S, KEI-22 TJC32, SYC18, SYC37, SYC50, SYC112, SYC235, SKC13, KSC36, KSC69, KSC180, KSC204, SNC53, SNC85, SNC100, INC6, INCI5, INC65, INC86, INC99, INC102, ISC16, ISC72, ISC73, STC71

Hatchobaru Kirishima Hot springs

Reservoir condition T (~

pH

C1

Na + K

Total C02

247-252 215-253 240-259 240 174-245

8.09-8.11 8.50-8.96 5.59-5.96 5.59 8.7-9.4

200-203 14.4-26.4 5.7-7.6 164 11.7-17.1

212-213 16.2-27.7 8.7-10.2 193 16.9-18.9

131-139

276-279 233-242

7.70-7.79 5.46-7.57

47.7-50.2 12.7-22.4

50.4-52.2 15.4-22.5

12.4-13.5 0.7-11.1

58-211

6.56-8.53

1.8-299

8.5-202

0.7-10.7

1.7-5.6

2.8-6.7 153,500 4.7-11.4

TABLE 2.2 Chemical features of Japanese geothermal systems (Chiba, 1991)

W

0 P

Locality:

Hatchobaru

Well name:

H-15

H-20

KEI-3

Reservoir temp., "C Enthalpy, kJ/kg pH at reservoir

276 1191 6.38

279 1154 6.59

242

24 1

233

-

-

-

7.12

6.50

6.31

Liquid phase Pressure of liquid sep. pH of liquid phase C1- mg/l SO:-, mg/l HCO;, mg/l S;. mg/l S i O z W , mg/l Na+, mg/l K+. mg/l Ca'+, mg/l Mg2+, mg/l Fe2-, mg/l ~ 1 ~mg/l + ,

0.0 7.70 2710 129 32.3 n.d. 977 I640 321 17.3 0.04 0.02 0.31

0.0 7.90 2550 154 39.7 n.d. 997 1570 298 10.7 0.55 0.02 0.69

0.0 8.45 660 145 63.7 3.8 567 420 55.6 8.3 n.d.

0.0 8.70 625 142 44.2 4.2 652 455 59.5 9.7

0.59

3.7 0.096 92.5 5.8 1.2 n.d. n.d. ad. n.d.

4.4 0.093 92.6 6.1 0.8 n.d. n.d. n.d. n.d.

0.0 0.02 75.5 3.9 13.4 6.8 0.4 n.d. n.d.

.

Steam phase Pressure of steam sep. Noncondensable gas. vol.% c02,o/o

H2S, % Nz. % CH4.% H2, % Ar, % He, % Geothermometer Quartz, "C Na/K. "C Na-K-Ca, "C

276 282 270

Kirishima

279 279 213

<0.01

236 242 224

KEI-5

KEI-6

KEI-7

KEI-19s

KEI-22

237

234

238

-

-

-

7.12

7.33

6.93

7.26

<0.01 1.4

0.0 8.30 1075 79.9 5.1 18.7 508 653 78.4 19.5 n.d. 0.14 0.69

0.0 8.60 640 I94 21.1 3.0 618 489 61.6 14.1 n.d. <0.01 1.1

0.0 8.70 640 194 21.1 3.0 616 492 61.6 14.1 n.d. 10.01 1.1

0.0 8.70 662 203 21.1 3.0 574 506 61.0 16.9 n.d. 10.01 1.1

0.0 8.70 645 192 24.5 3.0 606 488 61.6 14.4 n.d. tO.O1 1.1

0.0 0.087 78.74 14.96 3.72 2.43 0.14 n.d. n.d.

0.0 0.043 70.8 27.0 2.0 0.1 0.1 n.d. n.d.

0.0 0.009 49.45 47.42 2.94 0.1 0.11 n.d. n.d.

0.0 0.006 51.52 46.06 2.00 0.14 0.16 n.d. n.d.

0.0 0.02 56.62 29.94 12.63 0.52 0.3 n.d. n.d.

0.0 0.008 48.72 47.21 3.60 0.22 0.26 n.d. n.d.

n.d.

246 24 1 223

228 233 218

238

KEI- 17

~

242 238 218

242 237 218

236 234 214

240 238 218

9

4

B h,

TABLE 2.1 (continued) Locality:

Sumikawa

Okuaizu

Takigami

Well name:

KY-2

It

NE-2

NE-3

TT- 1

TT-2

TT-7

TT-8

TT-14

Reservoir temp.. "C Enthalpy, kJ/kg pH at reservoir

240 1147 6.39

219 996.4 4.91

I74 636.3 6.5 I

184 724.3 6.55

225 870.8 7.36

213 820.5 6.90

220 946.1 6.68

225 925.2 6.72

245 1013 6.90

Liquid phase Pressure of liquid sep pH of liquid phase C1-. mg/l SO:-, mg/l HCO;. nig/l s'-, mg/l SiO? (aq). mg/l ma+. mg/I K+, mg/l Ca'f. m g / ~ Mg'+. mg/l Fe'-, mg/l Al'+. mg/l

0.0 8.00 798 160 104 7.3 606 268 45. I 9.7 0.04 10.1 0.1

0.05 7.98 7780 77 1 956 n.d. 450 5560 644 76.0 3.22 0.69 0.87

0.0 8.70 458 267 125 n.d. 784 466 27.0 34.1 0.05 0.1 0.7

0.0 9.20 489 198 59.9 n.d. 322 461 30.4 33.5 0.37 0.1 0.4

1.0 9.40 525 26 1 50.8 1 0 .I 437 443 4 1 .o 20.0 0.10 0. I 0.1

1.3 9.00 547 23 8 53.9 <0. 1 365 466 47.1 18.8 0.08 0.1 0.22

0.5 9.10 640 154 63.9
0.0 9.70 585 25 1 60.8
1.3 9.10 785 95 56.9 <0. 1 593 502 89.0 8.4 0.08 0 .I 0.1

Steam phase Pressure of steam sep. Noncondensable gas. vol.%j CO?. % H:S. % N?, c/r CHd, 5% Hl. %, Ar, C/r He. %

0.0 0.089 34.6 40.6 21.9 0.269 1.90 0.299 5.19 x lop4

0.0 9.88 99.0 0.1 0.887 0.0108 0.0036 n.d.

0.0 0.108 88.5 1.4 6.3 2.7 0.3

0.0 0.171 69.0 1.3 18.7 1.7 0.03 n.d.

I1.d.

n.d.

n.d.

11.d.

1.3 0.068 75.7 7.0 15.9 0.9 0.2 n.d. n.d.

0.5 0.084 84.7 9.0 5.3 0.77 0.06 n.d. n.d.

0.0 0.120 71.7 5.9 21.0 0.5 1 0.10

11.d.

1 .o 0.053 85.1 5.5 7.9 0.8 0.04 n.d.

n.d.

n.d. n.d.

Geothermometer Quartz, "C Na/K. "C Na-K-Ca. "C

240 266 224

219 230 261

I69 174 165

178 184 171

225 210 192

213 218 199

220 243 215

225 23 1 213

245 272 245

I1.d.

1.3 0.10'

83.1 11.0 4.5 1.2 0.09

TABLE 2.2 (continued) w

0

Locality:

Nigorikawa

Kakkonda

Well name:

ND-7

NF- 1

KA- 1a

Reservoir temp.,"C Enthalpy, kJ/kg pH at reservoir

252 5.88

247 5.81

-

215

222

253

-

-

-

6.39

6.41

6.47

6.86

6.34

Liquid phase Pressure of liquid sep. pH of liquid phase CI-, mg/l SO:-, mg/l HCO;, mg/l s2-,mg/I SiO?(aq), mg/l Na+, mg/l K', mg/l Ca", mg/l MgZf, mg/l Fe2-, mg/l A I ~ +mg/l ,

0.0 8.09 10100 255 n.d. t o .1 704 6410 923 25.5 0.2

0.5 0.10

0.0 8.11 10100 264 n.d. tO.l 664 6360 809 30.8 0.1 0.4 0.19

0.01 8.59 1120 106 n.d. n.d. 462 727 84.6 24.7 tO.O1 tO.O1 0.42

0.05 8.57 1120 106 n.d. n.d. 426 728 83.0 25.4 0.02 0.14 0.34

0.01 8.50 1I40 109 n.d. n.d. 475 740 86.6 22.9 tO.O1 0.02 0.44

0.05 8.96 729 90.6 n.d. n.d. 712 492 69.5 7.7 tO.O1 0.03 1.14

Steam phase Pressure of steam sep. Noncondensable gas, vol.% c02, % H2S,% Nz, % CH4, % H2,% Ar, % He, %

7.2 1.36 97.5 1.1 0.839 0.541 0.0150 5.37 1 0 - ~ 2.14 x

6.8 1.34 97.5 1.o 0.801 0.517 0.0180 4.03 x 2.35 1 0 - ~

0.05 0.03 80.4 13.6 1.872 0.282 3.846 n.d. n.d.

0.05 0.02 77.9 17.1 3.19 0.625 0.585 n.d. n.d.

0.05 0.02 72.3 20.0 3.49 0.685 3.526 n.d. n.d.

Geothermometer Quartz, "C Na/K, "C Na-K-Ca, "C

252 250 282

241 249 278

220 230 216

215 228 214

223 23 1 217

220

Sumikawa

KB-3

KC-3

KD-6

KE- 1

K7-2

s-4

220

219

-

-

6.34

259 1225.4 6.94

0.01 8.50 1220 124 n.d. n.d. 454 780 89.0 28.3 tO.O1 tO.O1 0.37

0.05 8.58 1080 125 n.d. n.d. 449 705 81.7 23.5 to.01 10.01 0.42

0.0 7.70 42 1 132 74.5 t0.5 767 331 59.5 1.1 0.03 t o .1 1.8

0.05

0.05

0.05

0.05 67.3 26.1 4.17 1.102 1.333 n.d. n.d.

0.02 65.5 29.4 3.01 0.372 1.719 ad. n.d.

0.03 66.2 23.7 2.59 0.434 7.080 n.d. n.d.

0.0 0.019 39.7 39.2 18.53 0.190 1.462 0.197 2.32 1 0 - ~

253 249 232

220 229 214

219 230 215

255

o\

tQ

307

Present-day Mineralization and Geothermal Systems A -

4-r"

6

oo

Z ...... o

~

o I

o%_.

D /

4,

I

;"" "1-

o

_ K-fe/dspar

I

'

o

~_

B []

4

+ v

_o

3

O

I

O

9

I

I

I

J'

C

~

10

~

o ~

oo

5:

.

6

- 1 4 150

200

300

Temperature, ~

Fig. 2.11. The temperature dependence of cation/proton activity ratios of geothermal well discharges in Japan. The lines in the figure are recalculated temperature dependences of cation/proton ratios in Icelandic geothermal waters. The dashed curve in B represents the reaction: 1.5 K-feldspar + H+ -- 0.5 K-mica § 3 quartz (or chalcedony) + K+ (Chiba, 1991). Open circle: Takigami, open triangle: Kakkonda, open square: Okuaizu, solid circle: Kirishima, solid triangle: Sumikawa, solid square: Nigorikawa.

experimentally obtained (Orville, 1963; Ellis and Mahon, 1967; Hemley, 1967). This is interpreted by albite-K-feldspar-aque-ous solution equilibrium (White, 1965; Ellis, 1970; Fournier and Truesdell, 1973). Fournier and Truesdell (1973) showed that the ratio N a / K of solution generally yields the temperature much higher than actual temperature, and in addition to the concentrations of Na and K, the concentration Ca has to be considered in order to estimate the temperature of geothermal water. They called the relationship between the concentrations of Na, K and Ca and temperature " N a - K - C a geothermometer".

308

Chapter 2 i

i

i

"

+

"r~

o ~

7

o

6'

a

o A

5

least squaras

fit 4

9

I

;

"

I

'

B

=

I

+

[

+

o c~ 0 ...I

0

-2

A o

~

150

o

'

200

250

300

Temperature, ~

Fig. 2.12. The temperature dependence of cation/proton activity ratios of geothermal well discharges in Japan (Chiba, 1991). The solid curves are recalculated temperature dependences of cation/proton ratios of Icelandic geothermal waters. Symbols are as in Fig. 2.11. Shikazono (1976) attempted to interpret this N a - K - C a geothermometer based on thermodynamic equilibrium calculation. As noted already, the first approximation of electroneutrality relation is mNa+ -- m c l -

(2-13)

where m is molality. Helgeson (1967) constructed an activity diagram depicting chemical equilibrium points (albite-sericite-K-feldspar and albite-sericite-Na-montmorillonite) of N a 2 0 K 2 0 - S i O z - A 1 2 0 3 - H 2 0 system at elevated temperatures. At these points, aNa+/aH+ -- FNa+mNa+/FH+mH+ = constant

(2-14)

where V is activity coefficient. For these points, the relationship between pH and the concentration of C1- can be derived from equations (2-13) and (2-14). Figure 2.7 shows that silicate mineral assemblages comprise a pH buffer system and pH decreases with increasing concentration of C1-. In a similar manner to that described above, the relationship between the concentrations of K + and C1- is derived (Fig. 2.3). In this case, logmK+ increases with increasing log mc1-.

309

Present-day Mineralization and Geothermal Systems 9

,

I

i

I

8

~-to/~ ~-~~

'

A [

7

~7

9 ,...-",:~o.ol --.211,

L

[

~176

6

Japan

chalcedony . ~

,

t7

quartz

[

8

7 9

Iceland

51 0

.

'

"

o.,

chalcedony-~ quartz 1O0

200

300

Temperature, ~ Fig. 2.13. (A) Temperature dependence of pH in Japanese thermal waters. Lines indicate the temperature dependence of pH when pH is buffered by the K-feldspar-K-mica-quartz (or chalcedony at less than 200~ assemblage at a Na + K concentration of 0.1 and 0.01 mol/kg H20. Symbols are as in Fig. 2.11. (B) Temperature dependence of pH of Icelandic thermal waters. Large circles indicate well discharges. Small dots represent hot spring waters (Chiba, 1991).

If calcite is in equilibrium with the aqueous solution, the relationship between pH and the concentration of Ca 2+ can be obtained from reaction (2-10). This equation shows that activity of Ca 2+ is related to pH, concentration of H2CO3 and temperature. Because pH is related to the concentration of C1- for the equilibrium curves 1 and 2 in Fig. 2.14, the relationship between the concentrations of Ca 2+ and C1can be derived for calcite-albite-sericite-K-feldspar-quartz equilibrium (curves 4 and 7 in Fig. 2.14) and calcite-albite-sericite-Na-montmorillonite-quartz equilibrium (curves 5 and 8 in Fig. 2.14) with constant mHzCO3- The range of mHzCO3 in the solution in equilibrium with calcite is assumed to be 10 -2 to 10 -1. The other equilibrium curves for the assemblage including Ca minerals are also drawn (Fig. 2.14). These assemblages are wairakite-albite-sericite-K-feldspar-quartz (curve 3), Ca-montmotillonitealbite-sericite-Na-montmorillonite-quartz (curve 6), Ca-montmorillonite-albite-sericiteK-feldspar-quartz (curve 9) and anhydrite (curve 10). The effect of solid solution on the equilibrium curves is not considered because of the lack of thermochemical data of solid solution. Fournier and Truesdell (1973) showed that logmNa+/mK+ +4/310gmcaz+/mNa+ is constant at constant temperature over 100~ This value is about 0.8 at 250~

3 lO

Chapter 2

Log mca2+ 1

1

2

3 4

8

Log mc~-1

0

1

Fig. 2.14. The variation of concentration of Ca2+ with concentration of Cl- in aqueous solution in equilibrium with a given mineral assemblage at 250~ l" Equilibrium curve based on albite-sericite-Na-montmorillonitequartz-aqueous solution equilibrium and Na-K-Ca relationship obtained by Fournier and Truesdell (1973). 2: Equilibrium curve based on albite-K-feldspar-aqueous solution equilibrium and Na-K-Ca relationship obtained by Fournier and Truesdell (1973). 3: Wairakite-albite-sericite-K-feldspar-quartz. 4: Calcite-albitesericite-K-feldspar-quartz (mH2r = 10-2) 95: Calcite-albite-sericite-Na-montmoriilonite-quartz (mr~2co3 -10-2). 6: Ca-montmorillonite-albite-sericite-Na-montmorillonite--quartz. 7: Calcite-albite-sericite-K-feldspar-quartz (mHzCO3 = 10-1). 8: Calcite-albite-sericite-Na-montmorillonite-quartz (mH2CO3= 10-1). 9: Ca-montmorillonite-albite-sericite-K-feldspar-quartz. 10: Anhydrite (mso]- = 10-3). (Shikazono, 1976)

Therefore,

logmNa+/mK+ + 4 / 3 logmca2+/mNa+ = 0.8

(2-15)

Substitution of tuNa+ for mCl- can be approximately made. The concentration of K + is expressed as a function of that of C1- (Fig. 2.14). Therefore, the empirical relationship obtained by Fournier and Truesdell (1973) is changed into, logmca2+ = 2 1 o g m c l - + 6(logvK+) + K

(2-16)

where K is 10 -0.9 and 10 l ~ for albite-K-feldspar and a l b i t e - s e r i c i t e - N a - m o n t m o r i l l o n i t e quartz equilibrium, respectively and Yi denotes the activity coefficient of i species. These curves are drawn in Fig. 2.14. The slopes of the curves of 1 and 2 are nearly the same as those of the other curves calculated. It is very notable that the curves for the c o m m o n mineral assemblages at elevated temperatures (calcite-albite-sericite-K feldspar-quartz (mHzCO3 --- 10-2) and wairakitealbite-sericite-K-feldspar-quartz assemblage) are close to curve 2 in position. To the contrary, the curve for the mineral assemblage, albite-sericite-Na-montmorillonite-quartz, derived from the N a - K - C a relationship obtained by Fournier and Truesdell (1973) is quite different in position from the other curves calculated and the Ca 2+ concentration under high salinity conditions. These results indicate that the chemical composition of geothermal water at 250~ is largely controlled by such minerals c o m m o n l y occurring in geothermal area as albite, K-feldspar, sericite, calcite, wairakite and quartz.

311

Present-day Mineralization and Geothermal Systems 2.1.3. Present-day mineralization in subaerial geothermal areas in Japan

Base-metal (Pb, Zn, Cu) mineralizations were reported from Okuaizu, Nigorikawa, Arima and Ibusuki geothermal areas as scale products and precipitates from hot springs. Gold precipitations from hot springs are known from the Osorezan volcano and Beppu. Chemical compositions of hydrothermal solutions are summarized in Table 2.2 and chemical compositions of precipitates are presented in Table 2.3. C1- concentration of hydrothermal solutions precipitating base metals is high compared with hydrothermal solution not associated with base metal precipitations.

2.1.3.1. Nigorikawa The Nigorikawa (Mori) geothermal system is located in southwest Hokkaido (Fig. 2.15). The area is composed of basement rocks (sedimentary rocks such as limestone), and Tertiary andesitic rocks (Yoshida, 1991). The area is characterized by a Krakatoan-type caldera which was formed by volcanic activity about 12,000 to 20,000 years ago (Sato, 1988). The caldera is filled with sediments consisting of clays, conglomerate, and fall back materials (tuff and tuff breccia). The rock formations surrounding the caldera consist of pre-Tertiary rock which are unconformably overlain by the Neogene Tertiary formation in thicknesses of 300-700 m. The Nigorikawa geothermal water is characterized by the high CO2 content

N

140OE

o

43~

Sapporo

~

-~f

HOKKAIDO ISLAND

Nlgo Kawa

,20.

I~

omagatai. ?'Fi'";"CANT

~

'~

'

50km I Fig. 2.15. Locality map of the Mori geothermal power plant (Sato, 1988).

312

Chapter 2

compared with other geothermal systems and the equilibrium with silicate minerals except for (mca2+)/(mH+) (Chiba, 1991).

2.1.3.2. Osorezan Gold precipitations occur from hot springs at the Osorezan volcano which is located at the northern end of the volcanic front of the northeast Honshu (Fig. 2.16). The gold content of the precipitates is anomalously high, compared with those from the other geothermal areas (Table 2.3). The basement rock of the area is composed of the Mesozoic (limestone, chert, and late Tertiary volcanic formation. Osorezan is a composite volcano with caldera structure and post caldera domes. The K-Ar age of strata volcano in the earliest stage and the latest domes are 1 Ma and 0.2 Ma, respectively (Aoki, 1991). Dominant sulfides and sulfate minerals in the Osorezan area are pyrite, marcasite, orpiment, realgar, stibnite, krennerite, coloradoite, jordanite, wurtzite, sphalerite, cinnabar, and barite (Aoki, 1992b). HzS concentration of the Osorezan hot spring is very high, compared with the other Japanese geothermal waters (Table 2.4).

41o21 ' N i|i. Ill

~1

il

I

"~"-"~ "'i--~ I1 ~,r ,,,&,$

'.400,,, A ~

624m ,k"

~176 9

.,,'~ r, "

321 r n ~ . , , ~

/

~'t:

Q~/~'~'

/

1

.! 807m

'~

(#t

= I

:

%, "~176176 99 ,,

2krn

;'

'

,,:

."

LLI

..

"-

~176 ..~

~

I i

9 '.......'

"r-

I .o.. /

/

/

~o

'-,

~

", i ,0 ,.

I

/

-'

9LAKE USORI

"I--"

/

."

_

ILl 0

-%"

" q"

'

j /

826m J,

"..:

J

* 513m

"..,

"r'-

580rn

'-, A '

;."~ ,.. :646m

...-- - , , . . , . . . , , . . . . ,,. , . , . . , . - , , , , . . . , " '%. ,o

A

-

41~

N

Fig. 2.16. Location of Osorezan at the extreme north end of Honshu (inset), and the position of thermal area on the north shore of Lake Usorizan (box B) (Aoki, 1992b).

Present-day Mineralization and Geothermal Systems

313

TABLE 2.3

Chemical composition of hydrothermal gold-bearing precipitates from some geothermal waters (Izawa and Aoki, 1991) Onuma (pipeline)

Okuaizu

(control valve)

Fushime (pipeline)

Osorezan (hot spring)

Au (ppm) Ag (ppm) Hg (ppm) As (%) Sb (%) Pb (%) Zn (%) Cu (%) Fe (%)

0.86 23 0.005 0.001 0.004 0.35

116 34,900 2.75 11.1 13.1 13.4 16.0 1.56

1.4 2225 0.03 1.23 40.16 17.15 7.79 2.92

6510 0.4 5520 0.37 0.10 0.14 0.26 0.007 3.34

Te ( % )

-

-

-

1.05

The Okuaizu fluid may account for the more efficient transport and deposition of gold. - = not determined.

Chemical analytical data on the hot spring waters are given in Table 2.5 (Aoki and Thompson, 1990). Gold is precipitating from HzS-rich, diluted chloride and neutral hot springs (Aoki, 1992b). 3D and 3180 values of hot spring waters clearly indicate a mixing of groundwater and geothermal water with high 3180 and 3D (Fig. 2.17). Extrapolation of this mixing trend projects to the range of isotopic compositions of high temperature volcanic gases separated from andesitic to dacitic magma, suggesting a magmatic contribution to the hot spring waters (Aoki, 1992a,b). Another interesting characteristic of the Osorezan hydrothermal system is that it is located at a volcanic front. This is different from low sulfidation epithermal Au-Ag veins

----

I.TVG I

-20 o

v

s -40

-60 i

-8

i

!

-4

I

6~0

i

0

i

4

"

(%0)

Fig. 2.17. 3D and 3180 of Osorezan hot springs (Aoki, 1992b). P: deep hotwater, HTVG: high temperature

volcanic gas, solid circle: groundwater, the other symbols: hot springs in this area.

w

c

P

TABLE 2.4 Range of chemical compositions of geothermal waters discharged from some production wells and Osorezan hot spring (Izawa and Aoki, 1991)

Mori PH 8.0-9.3 SiOz (mg/l) 407-747 Ca 3.4-25.5 M& 0.14.7 Na 30884660 K 373-881 so4 90-593 c1 5160-7490 HC03 H2S 0.0-9.8

Onuma Matsukawa Kakkonda Onikobe Sumikawa

7.1-7.6 3.5-8.6 410419 163-1987 0-11 6.4-301 2.0-79.7 368-520 80-800 5342 25-250 102-221 338-1855 499493 5.4-14.9 40-65

tr.. trace; -, not determined.

Okuaizu

6.3-8.0 8.5-9.1 3.3-8.1 421-789 360-690 439-990 5.5-10.9 120-520 42-1750 0.1-20.8 0.01-0.34 0.2-39 342-500 960-1400 2700-10600 180-2550 34-63 88-200 51-107 20-92 4.0-771 428492 2290-5 100 6900-21800 4.3-956 tr.-6.1 tr-1.5

Takigami

Otake

Hatchobaru Kirishima

8.1-9.2 6.7-8.4 4.3-8.0 365-661 142468 509-994 8.4-20.6 8.4-30.4 9.8-80 0.10-0.38 2.3-14.4 0.1-1.9 466-509 670- 1060 1820-2590 47.1-89.0 70-140 289-383 95-252 96460 54-266 547-785 460- 1760 26504475 11-51 5-57 -

2.4-9.4 238-1036 5.6-30 0.04.8 85-8 18 13.2-1 90 79.9497 38-1180 0-42.9 tr.-18.7

Fushime

Osorezan (hot spring)

3.9-8 .O 540-1250 903-2480 0.83-23.2 7570-17400 1380-4970 <10-61 15000-32700 12-36

6.8 422 483 0.52 3830 489 48 6560 69 7.11

7.0 180 167 41.1 835 133 327 1380 635 75.5

Present-day Mineralization and Geothermal Systems

315

TABLE 2.5

Chemical composition of some typical hot spring waters at Osorezan (concentration in ppm; Aoki, 1992b) Location:

Stop 2-10

Stop 2-13

Stop 2-14

Stop 2-11

Date Temperature Lab. pH

3 Oct. 89 94.0 6.8

2 Oct. 89 96.8 5.8

4 Oct. 89 64.2 7.0

4 Oct. 89 65.6 2.0

Na K Li Ca Mg Sr Ba Fe Mn As SiO2 C1 F SO4 HCO3 B H2S

3830 489 6.04 483 0.52 4.58 0.94 0.009 0.069 29.0 422 6560 1.61 48 69 348 7.11

1440 244 2.96 212 15.6 2.22 0.41 0.004 3.40 47.8 142 2580 2.21 201 148 147 40.1

835 133 1.46 167 41.1 1.68 0.045 0.003 2.06 13.0 180 1380 1.22 327 635 73.5 75.5

400 30 0.22 63.5 5.19 0.47 0.059 2.04 0.64 0.35 171 458 0.36 1070 24.5 25.6

C1/B

18.9

17.6

18.8

18.7

(Se-type) that have been formed at extensional environment which is similar to back-arc environment. Hydrothermal alteration in the Osorezan area is extensive. At the foot of the lava dome, highly silicified alteration occurs. From this zone towards marginal parts, kaolinite zone and montmorillonite zone exist. This type of alteration was caused by the acid hydrothermal solution. But at present such acid hot solutions are not present in the Osorezan area. The acid solution is considered to be of volcanic origin. It is therefore thought that the water chemistry evolved from extremely acid at the early stage to neutral pH at present (Aoki, 1992a). Such evolution of a hydrothermal system from acidic sulfate hydrothermal solution to neutral is common in the epithermal system associated with precious metal mineralization. For example, advanced argillic alteration and intense silicification occurred at earlier stage of hydrothermal system in the Seigoshi A u - A g mine area. The mineralogical, isotopic and geologic characteristics of Osorezan are similar to Te-type epithermal A u - A g vein-type deposits. These characteristic features of Osorezan and Teine (Te-type) are similar to each other.

2.1.3.3. Okuaizu The Okuaizu geothermal area is located in Northeast Japan (Fig. 2.18). Pyroclastic rocks are the oldest ones of early Miocene (18-16 Ma). These Miocene and Pliocene

Chapter 2

316

T A B L E 2.6 Analytical results and discharge data of geothermal wells (SEN, 1991) Well:

OA-4

OA-6

84N- 1t

84N-2t

84N-3t

84N-5t

85N-6T

Sampling date:

84/2/3

83/12/26

85/6/14

85/11/8

86/1/16

84/12/11

86/12/20

W.H.P. (b.g.)

0.2

0.2

1.1

6.2

0.6

0.0

21.0

E sep (b.g.)

0.04

0.0

0.04

0.14

0.20

0.01

4.53

E stm (t/h) E wat (t/h)

5.8 4.7

0.6

5.9

24.1

2.8

1.9

59.4

1.1

20.9

33.4

3.5

3.4

135.2

H. td (kJ/kg)

1660

1210

908

1365

1415

1220

1295

Steam

sample

(vol. %)

H20 Gas

96.60 3.40

Gas CO2

96.5

H2S R. gas

2.7 0.8

R. gas H2 N2 CH4 He Ar Water

sample

90.12

96.37

95.85

96.79

90.67

9.88

3.63

4.15

3.2

9.33

99.0 0.1

96.3 3.0

99.1 0.3

98.1 1.3

96.6 2.6

0.9 0.4

0.7 7.10

0.6 2.0

0.6 6.0

0.8 3.10

98.4

92.3

97.0

92.0

96.1

1.2

0.3

0.6

2.0

0.50

-

0.05 0.27

0.03 0.41

-

0.05 0.31

7.98 -

6.44 9.8

6.30 16.9

6.19 -

6.37 5.3

(mg/kg)

pH (at 25~ Li +

7.05 -

5.50 -

Na +

5300

2700

5560

7120

10600

6990

3600

K+ Mg 2+

1400 1.10

i 80 0.10

644 3.22

2050 8.05

2550 20.8

1350 15.4

1030 8.9

Ca 2+

230

42

26.0

1160

1750

788

609

AI 3+ Mn 2+

0.43 -

0.10 -

0.87 -

0.19 419

0.40 185

0.09 -

0.12 129

0.2

0.93

Fe 2+

0.92

Cu Pb

. .

. .

70.0 . .

0.69 . .

0.31 . .

. .

1.47 . .

Zn

.

.

.

.

.

.

.

Sb

.

.

.

.

.

.

NH4

-

-

-

2.90

4.40

-

1.93

C1-

8800

6900

7780

15000

21800

13400

7930

.

SO 2-

36.0

191

771

4.0

HCO 3

83.0

4.3

956

41.1

6.2 5.2

50.2 12

61 .7 27.0

FBr-

-

-

-

2.70 29.1

1.94 43.0

-

1.45 11.2

I-

-

-

-

4.1

5.9

-

2.0

SiO2

680

439

450

990

866

662

605

HBO2

568

-

304

628

846

-

337

As

8.00

2.20

8.2

12.4

14.6

6.4

6.52

Hg

<0.001

<0.001

.

.

.

.

.

H2S

1.5

<0.1

.

.

.

.

.

Present-day Mineralization and Geothermal Systems

317

85N- 10T 87/9/26

86N- 11T 89/1/8

87N- 14T 87/10/30

87N- 15T 87/12/7

87N- 16T 89/1/8

87N- 17T 89/1/8

89N-21T 89/12/26

89N-22T 89/12/23

10.0 6.50 49.9 34.5 1920

8.3 6.42 42.6 55.7 1596

9.6 6.57 48.9 67.2 1565

62.7 6.50 87.7 134.0 1519

8.6 6.46 35.8 63.2 1447

8.0 6.45 41.7 38.6 1773

8.5 6.42 51.8 25.2 2092

15.3 6.42 63.3 28.9 2121

97.16 2.84 96.5 2.8 0.7 3.80 95.2 0.66 0.05 0.32

96.62 3.38 96.6 2.7 0.7 10.30 89.0 0.42 0.05 0.20

97.37 2.63 97.8 1.6 0.6 3.40 95.7 0.58 0.05 0.24

89.94 10.06 97.4 2.0 0.6 18.5 77.9 3.28 0.04 0.26

97.82 2.18 96.3 3.1 0.6 11.7 87.4 0.515 0.047 0.254

95.21 4.79 95.6 3.8 0.6 21.4 77.9 0.448 0.047 0.196

6.70 6210 1640 3.63 1030 0.02 125 0.02

6.56 7170 2000 9.81 1060 0.64 290 1.69

6.67 12.0 8020 1940 20.4 1330 0.14 311 0.61

7.68 3.9 3010 547 3.84 81.2 0.21 0.9 0.11

6.38 10.5 7770 1950 9.63 1220 0.55 266

5.64 10.7 5310 1780 6.55 647

Steam

sample

96.68 3.32 97.0 2.3 0.7 9.50 89.5 0.67 0.05 0.21 Water

7.03 3770 940 5.16 385 0.28 17.0 0.42

(vol.

%)

92.88 7.12 97.6 1.9 0.5 4.70 94.6 0.54 0.05 0.11 sample

(mg/kg)

6.90 7.33 5170 1060 8.11 586 0.29 46.9 0.49

1.07

0.25 1.8

7360 102 53.2

2.69 10000 48.1 23.5

13300 62.7 44.3

15300 7.1 35.9

-

1.78

-

-

-

17.3

-

-

-

2.8

-

-

6.70 -

771 462 7.73 <0.0005

787 11.7 -

1010 17.2 -

-

6.7

-

-

642

2.55 17000 65.4 <4 1.90 34.0 4.3 821 624 12.6 <0.0005 2.5

1.53 5300 34.5 145 1.81 10.8 1.5 632 370 2.20 0.0006 13.3

1.47

356 30.1 0.18 1.5

1.01

1.28

2.9 4.84 16500 22.4 20 2.64 31.4 2.8 887 693 17.4 <0.0005 1.9

2.7 5.86 11700 34.3 11 3.25 22.6 2.9 952 693 0.87 <0.0005 <0.5

318

Chapter 2 30~

'

./

='~G

YON EZAWA ATA

9Mt IIDE

,zT: 9

El

FUKUSHIMA

I,f

Mt. BANDA!

KORIYAWA

t. N A S U A

37 ~

I~ "~ SHIRAKAWA

IWAKI ta

A Mt. YAMIZO 9 Mt. NANTAi

139~

140 ~

141 ~

Fig. 2.18. Location of the Okuaizu geothermal system (Seki, 1991).

rocks are extensively altered. Rhyolite (0.21-0.59 Ma) intruded the early Miocene volcanic rocks. The Okuaizu geothermal system is characterized by high temperatures (maximum 340~ high salinity (about 2 wt% total dissolved solids (TDS)) and large amounts of non-condensable gases (1 wt% CO2 and 200 ppm H2S). The pH of the hydrothermal solution measured at 25~ is 6.44 (Table 2.6). However, the pH of the original fluid in the reservoir is computed to be 4.05. This pH as well as alkali and alkali earth element concentrations are plotted near the equilibrium curve of albite, K-mica, anhydrite and calcite (Fig. 2.19) (Seki, 1991 ). The dominant alteration minerals at the deeper part of the well include anhydrite, epidote, sericite, chlorite, calcite, dolomite, rhodochrosite, kutnahorite, zeolites (mordenite, clinoptilorite), chlorite and sericite/smectite interstratified clay mineral with subordinate amounts of kaolinite in the shallower part (Imai et al., 1996). Sulfide scales deposited on the casing wall were found. The minerals of the scales are listed in Table 2.7 (Imai et al., 1988, 1996; Nitta et al., 1991). The chemical compositions of sphalerite, chalcopyrite, and tetrahedrite are given in Table 2.7. 3D, 3180 and C1- concentration data suggest the mixing of meteoric water, connate seawater and magmatic gas (Seki, 1991) (Fig. 2.20). Br/C1 and B/C1 ratios are different from those of seawater (Fig. 2.21). This difference and Nz-Hz-Ar gas composition indicate a contribution of magmatic gas (Seki, 199 l, ! 996). Okuaizu geothermal area is located in Green tuff region. C1- might have been derived from evolved seawater. The high concentration of C1- in Okuaizu may be due to a contribution of evolved seawater. High C1- concentration, kinds of minerals, temperature and geologic environments are similar to those of Miocene-Pliocene epithermal basemetal vein-type deposits. Therefore, it is likely that the origin of geothermal water is the same as that of fossil epithermal base-metal vein-type deposits. Mixture of evolved seawater and meteoric water is thought to be the most likely origin.

319

Present-day Mineralization and Geothermal Systems '

I

'

"

'

0

'

I

[]

+5

0

_J I

)II

-5

i

!

9

250 300 Temperature (~ Fig. 2.19. Reservoir temperature versus saturation indices (log Q/K) for calcite, anhydrite, K-feldspar and K-mica based on the estimated composition of reservoir fluid (SEN, 1991). Estimation based on gas results of Seki (1990), with saturation calculations carried out by PECS (Takeno, 1988). Gas concentrations were assumed to be 1 wt% of CO2 and 250 mg/kg for HzS for all wells (SEN, 1991).

Cl- (xl000 mg/kg)

..JL

" SMOW/~"

(a)

/

///

/

*" SMOW '

."....,,.

i ell "I ' /~1~ ~s~harges 1 7 j.HTVG 6

~I~Hot spring waters 10

_,l_

(b)'

_,,/

~II/We'i discharges

Surficial waters

5180 (%0)

/,~

Hot spring waters

, if7 Surficial waters

20

-10

0

.40 -60

. +10

Fig. 2.20. (a) 3180-chloride diagram for geothermal well discharges, hot spring waters and local meteoric waters (modified from Nitta et al., 1991) (SEN, 1991). (b) 3180-3D diagram for geothermal well discharges, hot spring waters and local meteoric waters. The compositional range of Japanese high temperature volcanic gas (J-HTVG) is also shown (modified from Nitta et al., 1991).

320

Chapter 2

TABLE 2.7 EPMA analyses of sulfide minerals of precipitates from Okuaizu geothermal water (in wt%; Nitta et al., 1991) Sphalerite

1 Cu Ag Fe Zn As Sb Pb Cd Mn S Total

2

3

0.88 0.00 3.75 58.52 . 0.09 0.00 0.10 3.84 35.20

2.54 0.28 3.65 53.63 . . 1.53 0.34 0.15 3.52 35.16

3.57 0.53 3.40 53.55 . 2.37 0.57 0.18 3.29 33.88

102.33

100.80

101.34

Chalcopyrite

Tetrahedrite

1

1

2 31.42 0.10 32.28 0.48

26.73 0.08 36.06 0.42

0.14 0.15 0.00 0.03 36.25 100.85

Tetrahedrite

2

3

1

2

3

0.09 0.00 0.00 0.24 37.18

34.74 4.19 2.41 5.82 5.17 20.39 0.29 0.07 0.29 27.11

33.92 3.77 2.44 5.91 4.89 20.61 0.44 0.07 0.27 26.88

34.64 4.08 2.37 5.79 4.93 20.05 0.38 0.00 0.27 27.46

38.4 4.6 0.0 4.4 6.4 16.6 24.8

35.1 2.1 5.8 5.9 3.7 15.7 27.7

35.1 1.7 5.4 7.0 5.0 16.3 28.4

100.80

100.48

99.20

99.97

95.2

96.0

98.9

.

0.0100 0

, m

m L_

r~ m 0.0010

seawater i

i

i

,I

i

i

iJ

0.001

i

i

i

i

i

1

1 i

i

0.010

Br/CI ratio Fig. 2.21. B/CI and Br/CI weight ratio of Okuaizu geothermal water prior to boiling, showing that both ratios in geothermal water are different from those of seawater (Seki, 1996).

2.1.3.4. Sumikawa The Sumikawa geothermal system is located on the northern flank of Mt. Yakeyama, northeast Honshu. The district is composed of andesitic to acidic volcanic and sedimentary rocks of Miocene to Quaternary ages. The chemical compositions of geothermal waters are characterized by low total dissolved solid (<300 mg/1), near neutral pH, and Na-C1 domination (Sakai et al., 1986; Ueda et al., 1991). Inoue et al. (1999) carried out a detailed investigation of hydrothermal alteration

Present-day Mineralization and Geothermal Systems

321

minerals. They found a simple zonal distribution from the bottom to the surface, biotite (+ chlorite) --+ illite + chlorite --+ illite/smectite + corrensite --+ kaolinite --+ smectite --+ halloysite. They considered the formation mechanism of these minerals. Their conclusions are: (1) The compositions of the Sumikawa discharge fluids are controlled by Kfeldspar, illite, calcite and Ca-aluminosilicates (epidote, prehnite, wairakite). (2) The redox condition of the deep fluids may be buffered by the chlorite-pyrite equilibrium. (3) Extensive boiling and gas loss at deeper levels shift the fluid composition locally from the illite stability field to the K-feldspar, epidote, prehnite, and/or wairakite stability fields. At shallower levels, the vapors separated during boiling mix with cold groundwaters to form a CO2-rich steam-heated water and it strongly hydrolyzed plagioclase and volcanic glass to form kaolinite, dolomite, and siderite. (4) Cold groundwaters descend to the permeable zones. Mixing of the deep fluids with the cold waters brings about a rapid decrease in temperature and there by results in the precipitation of Fe-rich chlorite of about 0.5 of Fe/Fe + Mg ratio.

2.1.3.5. Arima hot springs Arima hot springs are located at middle Honshu. The geology of this area is composed of upper Cretaceous, and Paleogene granitic rocks. The Arima hot spring waters are classified into three types" (1) Na-Ca-C1 type brine which is high in salinity and CO2 and medium to low in temperature; (2) highly saline Na-Ca-Cl-type water of high temperature and low CO2 concentration and (3) dilute and CO2-rich water of low temperature (Table 2.8). The hot spring waters are notable for their high content of Fe (up to 200 ppm), Mn (up to 67 ppm) and base-metals (Cu, Pb, Zn) (0.1-0.4 ppm) (Table 2.8) (Ikeda, 1955). The Arima hot spring waters are significantly enriched in the heavy isotopes of oxygen, indicating mixtures of up to 80% of a saline and CO2-rich brine of 8180 = +8%0 and 8D -- -25%0 to -30%0 and local meteoric water (Fig. 2.22) (Sakai and Matsubaya, 1974). The chemical and isotopic compositions and 3He/4He observed in the gases of Arima brines suggest that the deep brines originate from the deep-lying magmas underneath the Arima Spa or from sedimentary rocks during metamorphism induced by the magma. Sulfide minerals (galena, sphalerite, pyrite) were reported by Nakamura and Maeda (1961). Masuda et al. (1986) recognized two stages of alteration of host rocks. The earlier regional alteration stage is characterized by the presence of 2M- and 1M-type muscovite, albite, chlorite, calcite and epidote. Muscovite and Fe-chlorite formed in the late hydrothermal stage. Oxygen and hydrogen isotopic compositions of secondary muscovite show that the early-stage alteration proceeded under the meteoric water circulation, and the second-stage hydrothermal activity had chemical and stable isotopic characteristics of non-meteoric water origin similar to the present-day Arima brine. The present-day Arima hydrothermal system is thought to be a remnant of the second-stage hydrothermal activity.

TABLE 2.8 Chemical composition of geothermal water associated with base-metal deposits (White. 1967)

Tamagawa Springs, Japan Arima Springs, Hyogo Pref., Japan Salton Sea geothermal area, Imperial Co. Calif. East Tintic district, Utah Co., Utah San Juan Mountains. Colo.. Springs, Ouray Co.. Colo. Centennial mine, Houghton Co.. Mich.

Tamagawa Springs, Japan Arima Springs, Hyogo Pref., Japan Salton Sea geothermal area. Imperial Co. Calif. East Tintic district, Utah Co., Utah San Juan Mountains, Colo., Springs. Ouray Co., Colo. Centennial mine, Houghton Co., Mich.

Temp. 3"(

pH

Na

K

NH4

Ca

Mg

98 90.5 -220 60 62 Cold

1.2 6.4 -5.5 6.1 6.8 6.5

1I4 16,200 51,000 2.270 111 11.900

65 3,160 25,000 228 8.0 38

44 482 2.4 110

210 3.530 40,000 375 376 62,900

83 24 35 53 6.1 179

HCO3

CI

SO4

H2S

B

SiOl

Others

-

146 -

749 128 24

3,240 33,260 185,000 3,670 45 128,000

1.330 1.0 56 391 1,030 88

-

0.3 0.0 -

-

23 106 300 5.6 0.2 2.5

324 69 z 110 48 49 <10

Fe

105 187 3,200 1.o 0.4 1.O

Mn

4.2 61 2,000 0.6 0.9 2.5

Cu 0.01; Pb 1.0: Zn 2.8; Hg 0.01 Cu 0.1: Pb 0.4; Zn 0.2 Zn 970; Pb 104; Cu 10; Ag 1 As 0.6 As 0.00; Cu 0.02; Ag 0.002 Cu 0.7: Sr 320

323

Present-day Mineralization and Geothermal Systems

/,,~Oj

|

+5 (~18 0

%

J t"l

-5

0o

/

-10

~

9

0

.

.

f

o.Oj

cm"

O " Arima, high temperature brines 9 Arima, 9 low to medium temperature brines and meteoric waters (at CI'=0) : Arima, the brines of the maximum salinity @ : Takarazuka brine I"1 : Ishibotoke brines and meteroric waters (CI-=0)

. 500

.

1000

CI" m mol/i

Fig. 2.22. Relationship between ~180 and CI- concentration of Arima, Ishibotoke, and Takarazuka brines (Sakai and Matsubaya, 1974).

2.1.3.6. Beppu hot springs The Beppu hot springs are located at the eastern end of the Beppu-Shimabara Graben, Kyushu. The basement rocks are composed of Paleozoic crystalline schists and Cretaceous granitic rocks. Miocene to early Pleistocene andesitic rocks occur mainly in the southern part of the area, while lava domes of hornblende andesite which are younger than 100,000 years occur in the western part. The total amount of discharged hot spring water is 50,000 ton/day, indicating a huge geothermal system (Taguchi et al., 1988). The hot spring waters are divided into sulfate-rich steam heated water, chloriderich deep-water, bicarbonate-dominated water and their intermediate types. The red precipitates in the Chino-ike Jigoku contain high concentration of base and precious metals (Au, 23; Ag, 383; As, 4440; Sb, 180; Pb, 14442; Zn, 104; Cu, 57 ppm, respectively) and consist of low cristobalite, tridymite, kaolinite, hematite and montmorillonite (Koga, 1961, 1972; Yamashita, 1977; Yoshida et al., 1978). The hot water of Jigoku is a mixture of sulfate-rich water and deep chloride-rich water.

324

Chapter 2

2.1.3. 7. Fushime The Fushime geothermal area is located at the southeastern part of the Satsuma Peninsula, Kyushu, Japan. The area consists of Quaternary late Pliocene pyroclastics and sedimentary rocks. Marine mudstone and sandstone of Mesozoic-lower Tertiary Shimanto Supergroup are overlain by these rocks. Thick (more than 1,000 m) dacitic tufts interbedded with marine sedimentary rocks of late Pliocene-early Pleistocene age occur. These rocks overlie altered andesite lava and dacitic pyroclastics of Miocene-late Pliocene (Yoshimura et al., 1988). 3D and 8180 of geothermal waters demonstrate a mixture of modified seawater and local meteoric water (Sakai and Matsubaya, 1974). The reservoir waters are saline and the maximum C1- concentration is similar to that of seawater (Akaku, 1988). Thus, a source of C1- in the Ibusuki hot spring which is located at southern Kyushu is thought to be seawater (Chiba, 1991; Akaku et al., 1991). The depletion of Mg and SO 2- and enrichments of K, Ca, SiO2 and base metals (Fe, Mn, Zn, Pb) in geothermal waters indicate that they originated from seawater-rock interaction (Akaku, 1988). Thermochemical calculation indicates that the geothermal waters are nearly in equilibrium with anhydrite and Na- and K-feldspars (Chiba, 1991). The alteration is divided into four zones with depth: montmorillonite zone, transition zone, chlorite zone and epidote zone (Yagi, 1990; Yagi and Kai, 1990).

2.2. Comparison of active geothermal systems with epithermal vein-type deposits Close similarities between epithermal vein-type deposits and active geothermal systems have been cited by various authors (e.g., White, 1955, 1981; Henley and Ellis, 1983; Shikazono, 1985a,b; Izawa and Aoki, 1991). In this section (2.2), geochemical, mineralogical and geological characteristics of epithermal vein-type deposits summarized in section 1.4 will be compared with subaerial active geothermal systems associated with base metal and Au-Ag mineralizations mentioned in sections 2.1.1 and 2.1.2.

2.2.1. Distribution Fig. 2.23 shows the distributions of major geothermal systems and epithermal gold deposits of Japanese Islands. It is interesting to note that their distributions are similar and they are distributed close to the volcanic front.

2.2.2. Metals Based on the study of precious-metal vein-type deposits in western U.S.A., White (1981) classified the epithermal precious-metal vein-type deposits into an Au-dominated and an Ag-dominated groups. White (1981) considered that the Ag-dominated group

325

Present-day Mineralization and Geothermal Systems

Hokkaido

,L epithermal gold deposits 0

Chitose

active gold depositing hot springs

n geothermal w p~176 ,/" volcanic front Bajo .

.

Taio ~

Teine Mori -----~f'~'l~ Sumikawa

~~-Osorezan

On um a ' ~ l ~ l ~ , Matsu kawa ' Okua,zu /O~/Kakk~ S a do.--=/.1~ "JIk\ [ ll r -~~ "Onikobe ~ ~ , v j : f N ~ ' T a katama

~I'st"-~-z'" "~D.

KyUShU d&'~ - - ~ _

.~.~ J "

(, /"- ~=~,,,~,, "Seigoshi _~~176 Mochiko'shi

Okuchi'~-//fH~tc~;;~ru Shikoku Hishikari----~~'X Otake, Takigami i Yamagan~~~na .... =,..... ~ Kushikino AP~Fushime ! Kasuga, Iwato, Akeshi

. . . . . u

.yuKm

Fig. 2.23. The major geothermal system and epithermal gold deposits of Japan. Epithermal gold deposits are represented by gold mines or mining areas that have produced more than 10 metric tons of gold (Izawa and Aoki, 1991). has formed at a lower depth than the Au-dominated group from the comparison of the characteristics of epithermal Au-Ag vein-type deposits with those of active geothermal systems accompanying the precious and base metal precipitations. He mentioned that the Ag-dominated group is rich in base metals, Se, Te and Bi and the Au-dominated group is characterized by T1 and B. This correlation between the Ag/Au total production ratio and metals concentrated in the epithermal vein-type deposits of the western U.S.A. seems to exist also in the Japanese epithermal vein-type and hot spring-type deposits. Fig. 2.24 indicates the relationship between the Ag/Au total production ratios for Japanese epithermal deposits, and the metals concentrated in the deposits (Shikazono, 1986). In Fig. 2.24, Cu, Pb, Zn and Mn are also plotted if the metal has been produced from that particular mine. Sn, W, Cu, Bi, Pb, Zn, Mn, Se, Ge, Te and Hg is plotted if the presence of the following minerals has been established: cassiterite and stannite for Sn, naumannite (Ag2Se), aguilarite and Se-bearing argentite for Se, hessite (Ag2Te) for Te, and cinnabar (HgS) for Hg, argyrodite (AgsGeS6) for Ge, wolframite (FeWO4) for W, and bismuthinite (Bi2S3) for Bi. This figure clearly demonstrates that base metals (Sn, W, Cu, Pb, Zn, Mn, and Bi) tend to be concentrated in the deposits with high Ag/Au total production ratios, whereas Te- and Hg-bearing deposits tend to have low Ag/Au total production ratios. Se and Ge, although there is not much data, tend to be concentrated in the deposits with intermediate Ag/Au total production ratios. As considered by White (1981), the depth of ore formation relates to the type of ore deposits. Although the estimate of depth of formation for Japanese epithermal

326

Chapter 2

Sn

9 m . .

.m 9

ml

W

., m m i m l _ . . t 9 ml

Cu --

Bi

Pb

. _m___m.

Or) (1) r

.__

9

O

Zn . . . .

(D ..Q

m ~

E

JI

m m m

z m

i'l

m

m

0

9

Mn

Se

m

Ge

m

Te

m

_Z

+1

+2

+3

+4

Hg

Log(Ag/Au) Fig. 2.24. Relationship between the A g / A u total production ratios and metal elements from epitherma] deposits (Shikazono, ] 986).

deposits is difficult, fluid inclusion studies give some constraints on the estimate of depth. The fluid inclusion studies of epithermal Au-Ag deposits, although data are scarce because of the very fine-grained nature of the minerals, indicate that boiling took place during the mineralizations in some deposits (e.g., Enjoji and Nakayama, 1982; Nakayama and Enjoji, 1985). This type of deposit is similar to the epithermal Au-Ag bonanza-type deposits such as Finlandia (Peru) and Tayoltita (Mexico) studied by Kamilli and Ohmoto (1977), and Smith et al. (1982). These authors indicated that boiling took place during the precipitation of Au and Ag minerals. The fluid inclusion studies of hot spring-type deposits (Takenouchi, 1981) show a wide range of homogenization temperatures in a given quartz crystal which suggests the boiling of ore fluids. These fluid inclusion studies demonstrate that the hot spring-type formed under shallow depth from the surface. In contrast to epithermal precious-metal vein-type and hot spring-type deposits, no report on the evidence of boiling in inclusion fluids from epithermal base-metal vein-type deposits in Japan has been published. Therefore, from the fluid inclusion studies, it is suggested that the Au-rich and base-metal-poor deposits formed under a shallower environment than Au-poor base-metal-rich deposits. This tendency is in agreement with that found in some active geothermal systems. As summarized by White (1981), Weissberg et al. (1979), Hedenquist (1983) and others, Au-rich siliceous sinters are forming at shallow depths from the surface as observed in Steamboat Springs (U.S.A.), Ohaki Pool and Waiotapu (New Zealand), Beppu and Osorezan (Japan), but base metal

Present-day Mineralization and Geothermal Systems

327

(Pb, Zn) sulfides occur in deeper part as in Broadlands (New Zealand) and in Okuaizu (Japan). Pb and Zn sulfides occur dominantly at 400 m from the surface in Broadlands. In the boreholes of Broadlands, concentrations of precious and base metal elements in the boreholes change with depth as studied by Ewers and Keayse (1977), Au, As, Sb and T1 decrease with depth, but Ag increases with depth. In active geothermal systems in Japan, Au, Te, Se, T1 and Hg are enriched on the surface in the Osorezan hydrothermal system.

2.2.3. Mineralogy 2.2.3.1. Opaque minerals Opaque minerals identified from active geothermal areas are pyrite, sphalerite, galena, chalcopyrite, and tetrahedrite from Okuaizu, Fushime, and Nigorikawa (Japan), Salton Sea (U.S.A.) and Broadlands (New Zealand). These opaque minerals are common in Pb-Zn vein-type deposits in Japan Opaque minerals identified from silica sinter containing gold and mercury are krennerite, coloradoite and metacinnabar (Osorezan, Japan; Waiotapu, New Zealand). These minerals are not found in low sulfidation-type Au-Ag deposits in Japan but are reported from Kobechizawa, and Date in south Hokkaido which are massive and disseminated types and similar to hot spring type deposits. 2.2.3.2. Gangue and alteration minerals The generalized sequence of alteration minerals from shallower to deeper portions and/or from lower to higher temperatures in active geothermal systems, which is constructed mainly based on the work by Henley and Ellis (1983), is given in Fig. 2.25. It is shown in Fig. 2.25 that the change in alteration and gangue minerals largely depends on temperature as well as on the other physicochemical parameters such as fs2, fo2, fr and pH. It has been pointed out by Giggenbach (1981) on the basis of thermochemical calculations that epidote occurs at higher temperatures of at least more than 240~ and K-feldspar occurs at restricted temperatures, i.e. below ca. 250~ in active geothermal systems. These theoretical results seem to be consistent with those observed in epithermal vein-type deposits in Japan. Epidote tends to occur in high-temperature Cu deposits and adularia occurs abundantly in precious deposits whose average homogenization temperature is ca. 240~ or slightly less. Although broad similarities are found between epithermal vein-type deposits in Japan and active geothermal systems, there is a significant difference between their alteration and gangue minerals, and those from active geothermal systems; Mn- and Fe-minerals (rhodochrosite, rhodonite, pyroxmangite, inesite, johannsenite, bustamite, siderite, Fe-chlorite) are common in epithermal vein-type deposits in Japan, but they are rare in active geothermal system except for siderite which occurs in the lower temperature zone of some active geothermal systems (Browne, 1978). The iron content of chlorite from the epithermal vein-type deposits is highly variable (Shirozu, 1969; Shikazono and

328

Chapter 2 A m o r p h o u s Silica Quartz K-feldspar Albite Calcite Dolomite/Ankerite Kaolinite Montmorillonite Mont.-Illite Illite Chlorite Chl.-Mont. Vermiculite Biotite Actinolite Tremolite Diopside/Hedenbergite Garnet Epidote Prehnite Heulandite Stilbite Ptilolite Laumontite Wairakite Sphene

9

ml

,,,- ,m I

,B o ,,,,....,

,,

~. __

I

100

I

150

I

200

!

I

250

300

Temperature (~

Fig. 2.25. Generalized zonal sequence of alteration minerals in active geothermal areas (Henley and Ellis, 1983" Shikazono, 1985b).

Kawahata, 1987), but is generally higher than that of chlorite from active geothermal systems such as the Larderello (Cavaretta et al., 1982) and the Salton Sea (McDowell and Elders, 1980). This difference may be explained by the fact that gangue minerals in epithermal vein-type deposits were precipitated from ascending ore fluids which contained high amounts of Fe and Mn together with other base metal elements under an environment of a high water/rock ratio, whereas alteration minerals (excluding vein minerals) in active geothermal systems were formed mainly in geothermal reservoirs at a low water/rock ratio.

2.2.4. Geochemical features of hydrothermal fluids

2.2.4.1. Gas fugacities C02 fugacity (fco2). In Fig. 2.26 /CO2 calculated by several workers for geothermal reservoir fluids in active geothermal systems are summarized. Giggenbach (1981) summarized the values of partial pressure of CO2 gas (Pco2) of geothermal waters in New Zealand (Broadlands, Wairakei, Kawerau) and suggested that the dominant chemical reaction determining Pco2 of the geothermal waters is, plagioclase +

C02

--

clay + calcite

(2-17)

where "plagioclase" and "clay" are anorthite and kaolinite components in unspecified mineral phases. He derived the semi-empirical relationship between temperature and

329

Present-day Mineralization and Geothermal Systems T.Alfina B

Broadlands ..L' Niland

agnore

d' ,

ii

Kizildere d

,

Rotorua

@J o 0 c.)

o

[|

/ ~

.

..

i~rafla

I l~,..,"~'Jt"~1 1 phase Walrakel ~__~'~_~=,t~l;-~,_~_.,~---------------~Krafla ( . . ~ n j ~ ~ ' ~ ~ 2 phase ...........

= = o n = .

-I

-2

/

/

,

f

/

. .......

"

=.i0.==

-3 -4 i

i

i

|

150

200

250

300

T(~

Fig. 2.26. Range of carbon dioxide fugacity (fco2) and temperature for the propylitic alteration (epidote zone) in the Seigoshi area and same active geothermal systems. Seigoshi = propylitic alteration of the Seigoshi district. The curves A-B and A'-B t are equilibria for epidote ( X p i s - " 0 . 3 0 ) - K-mica (aK-mica = 0 . 9 ) -K-feldspar (aK-feldspar = 0 . 9 5 ) -- calcite assemblages at saturated water vapor pressure condition (Shikazono, 1985a). PCO2 for the above chemical reaction,

log P c Q = 15.26 - 7850/(T + 273.15)

(2-18)

where T is the temperature in ~ Bird and Helgeson (1981) derived a relationship between /CO2 and temperature for the equilibrium assemblage of K-feldspar-K-mica-epidote-calcite as a function of the mole fraction of pistacite component in epidote assuming that the activity of KAlz(A1Si3010)(OH)2 in K-mica is equal to 0.6. They have shown that the fco2 in the Salton Sea geothermal system is controlled by this equilibrium assemblage. Arn6rsson and Gunnlaughsson (1983) have suggested that the f c Q of the Icelandic geothermal system is controlled by secondary mineral assemblages containing epidote. Fig. 2.26 shows that the fco2 of geothermal reservoir waters and vein-ore fluids increase with increasing temperature, although at lower temperatures fco2 of some geothermal systems (Alfina, Bagnore, Kizildere, Ngawha) gives higher values than the fco2 values controlled by silicate and carbonate minerals. This relationship may suggest that some buffering assemblages (i.e., plagioclase-clay-calcite and epidote-K-feldspar-K-mica) in country rocks play an important role in controlling the fco2 of the epithermal vein-ore fluids as well as in the geothermal waters in active geothermal systems. There are several processes which may cause deviations from the equilibrium between ore fluids and alteration minerals in country rocks. They are boiling of the fluids, conductive cooling of the fluids, mixing of the ore fluids with cold waters such as extensive groundwater, and adiabatic expansion of the fluids. As suggested by Grant (1982) and Giggenbach (1980), if the fluids which are initially in equilibrium with the plagioclase-clay-calcite assemblage at high temperatures

Chapter 2

330

(more than ca. 300~ rise along the adiabatic expansion curve, then the Pco2-temperature path has a position similar to the plagioclase-controlling path at temperatures higher than ca. 250~ (Fig. 2.26). Below this temperature, in general, the adiabatic expansion path deviates significantly from the plagioclase-controlling path. Therefore, adiabatic expansion may also explain the fcoz-temperature relationship for the epithermal vein ore fluids, although it is necessary that the fr is controlled initially by the alteration mineral assemblages at high temperatures in the reservoir rocks. Fluid inclusion data from some epithermal Au-Ag vein-type deposits in Japan show that the average homogenization temperature of the main-stage of Au-Ag mineralization is slightly less than 250~ Therefore, it would be expected that the fco2 deviates from the silicate-carbonatecontrolling curve due to the effects of adiabatic expansion as well as other processes such as mixing, boiling, and conduction.

H2S fugacity (fH2s). It is thought that /H2S is controlled by pyrite, Fe-oxides such as magnetite, hematite and Fe-silicates such as chlorite. For instance, the following reaction is important for controlling fHzs (Giggenbach, 1997): 2FeS2 + 2(FeO) + 4H20 -- 4(FeOi.5) + 4H2S

(2-19)

where (FeO) and (FeO1.5) mean FeO component in Fe-silicates such as fayalite and (FeOl.5) in hematite. Temperature dependence of the reaction (2-19) corresponds to (Giggenbach, 1997), log fHzS = 6.05 -- 3990/(T + 273)

(2-20)

where T is temperature in ~ The relation reproduces very satisfactory /H2S observed for Icelandic, New Zealand, Philippine and other geothermal well discharges (Arn6rsson, 1985). The equilibrium relations of epidote-K-mica-K-feldspar-pyrite-chlorite, hematite -+- Sliq = pyrite + H2S, anhydrite-magnetite-pyrite-clinozoisite and pyrite-hematitemagnetite assemblage are shown in Fig. 2.27. Based on the equilibrium curves and analytical data on epidote and chlorite, fHzs of the epithermal Au-Ag vein ore fluids for some propylitic alterations is also estimated (Shikazono, 1985a). The fI-IzS values of epithermal vein ore fluids were estimated also based on the equilibrium for the reaction, H2Sgas -k-(FeO)chl -k- 1/2S2 = H2Oaq -k-FeS2

(2-21)

where (FeO)chl is FeO component in chlorite. The equilibrium constant for (2-21), K2-21, is expressed as, K2-21 --

1/(fx-IzsaFeof~f 2)

(2-22)

Since temperature, fs2, and iron content of chlorite were determined (section 2.4.2), we can estimate fH2s range from the following reaction: (FeO)ch! + H2S + 1/2S2 ----FeS2 + H20

(2-23)

From the H2S concentration of hydrothermal solution, we can estimate the temperature of hydrothermal solution buffered by these assemblages to be ca. 350~ for

331

Present-day Mineralization and Geothermal Systems

+2

+1 if)

A

| Broadlands .,/~ The Geysers ~, ( w e l l II) . o., " " ~, 9Mexicali (well 5! L %, 9 .."e o " " 9Nasu-ChausudaKe Larderello \ \. .. ~ ~' t ~ ; " - " ~ Hv-r-- -rdi k \ , ~"e"~ W a i r a k e e a g e o~ ~ "~.,,,,,~ ~ ' . (average,) , ~; ;~,~f,,':" " - - C ' . Showashinzan

"14.-

0 _J

-1 -2

Showashinzan'~..-". %

-3

~..~a

s a 9 lie "

t"

&

SatsumaIwojima

. ;.~:4.~" .~

," -1" A h u a c h a p a n ,, - " El Tatio,, , , " =,o j

~':'f*.:

.....

""

g

J

-4 C,"

150

150

150

150

150

150

150

T(Oc)

Fig. 2.27. Range of hydrogen sulfide fugacity ( f H 2 S ) and temperature for the propylitic alteration (epidote zone) and the advanced argillic alteration (silica-rich zone) at Seigoshi and some active geothermal systems. A -- propylitic alteration; B = advanced argillic alteration. Data on active geothermal sytems are taken from Ellis and Mahon (1977). The calculation of fIazS of Showashinzan for saturated water vapor pressure condition was based on the analytical data on volcanic gas by Matsuo (1961). The calculation of fHzs of Satsumaiwo-Jima was based on the analytical data on volcanic gas by Matsuo et al. (1974) assuming PI-]2o -- 0.5 kbar. Thermochemical data necessary for estimating fHzs for the propylitic and advanced argillic alterations are taken from Bird and Helgeson (1981) and Giggenbach (1981). The curves A-A f, B-W and C-C ~represent the equilibrium o f e p i d o t e (Xpis = 0 . 3 0 ) - K - m i c a (aK-mica = 0 . 9 ) - K - f e l d s p a r (aK-feldspar = 0 . 9 5 ) - p y r i t e - c h l o r i t e

(aFeo = 0.5)

where Xpis, aK-mica, aK-feldspar and aFeo are mole fraction of pistacite component in epidote, activity of K-mica component in mica, activity of K-feldspar component in K-feldspar and activity of FeO component in chlorite, and hematite + liquid sulfur ~ pyrite + HzS respectively (open circle = vapor-dominated system; solid circle --hot water-dominated system; solid triangle --volcanic gas)(Shikazono, 1985a).

the O k u a i z u g e o t h e r m a l water in g e o t h e r m a l reservoir. This t e m p e r a t u r e is in a c c o r d a n c e with the m a x i m u m t e m p e r a t u r e m e a s u r e d for the O k u a i z u g e o t h e r m a l water. T h e r e f o r e , it is h i g h l y likely that the H z S c o n c e n t r a t i o n is buffered by this a s s e m b l a g e . O n the other hand, Seki (1991) t h o u g h t that the h i g h H2S c o n c e n t r a t i o n of g e o t h e r m a l reservoir water of the O k u a i z u is c a u s e d by the injection of volcanic gas f r o m a d e e p e r part to the reservoir. H o w e v e r , it s e e m s likely f r o m the above r e a s o n that this injection is not n e c e s s a r y for a h i g h c o n c e n t r a t i o n of H2S. C h i b a (1991) has s h o w n that m o s t h i g h - t e m p e r a t u r e g e o t h e r m a l waters in Japan are saturated with a n h y d r i t e (Fig. 2.28). Pyrite, m a g n e t i t e and e p i d o t e are v e r y c o m m o n in the g e o t h e r m a l areas. T h e r e f o r e , it is h i g h l y likely that h i g h - t e m p e r a t u r e g e o t h e r m a l waters are c o n t r o l l e d by a n h y d r i t e - p y r i t e - m a g n e t i t e - e p i d o t e a s s e m b l a g e .

2.2.5. Geological and tectonic environment and volcanism T h e r e are two opinions c o n c e r n i n g the type of v o l c a n i s m g e n e t i c a l l y related to p r e c i o u s - m e t a l and b a s e - m e t a l v e i n - t y p e m i n e r a l i z a t i o n s . O n e is that the m i n e r a l i z a t i o n is related to andesitic calc-alkaline v o l c a n i c activity and the f o r m a t i o n of c a u l d r o n s

332

Chapter 2

0 ~ 0 --I

-6

9 .;

-8

9

.;.',,

supersaturated

-

t ~

"10

~.o=o 6

-12 undersaturated

-14

Japan

-16 supersaturated

*

"8

0 ~

0 .J

-10 -12

Iceland

-8

9

o ~...~

-I0

o

"12

9

9 o

~

undersaturated

-14 -16

supersaturated

C

-6

~

9

."

undersaturated

-16

0 ..i

OooQ, 9

~

-14

0

9

oe 9

Broad~ands, N . Z o

I

'~~ o Q

360 Temperature, ~

Fig. 2.28. Plot of logarithm of activity product (log Q) versus reservoir temperature. Solid circles stand for activity products of Ca 2+ and SO 2- in geothermal fluids, and open circles for Ca 2+ and CO~-. Small and large circles indicate hot spring and geothermal well samples, respectively. The upper line is the solubility

product of anhydrite and the lower is calcite. From top to bottom, figures are for Japanese, Icelandic and Broadlands, N.Z., thermal waters, respectively (Chiba, 1991).

(Kubota, 1991). For example, in the south Kyushu (Nansatsu district), the southwest Hokkaido and the north Hokkaido epithermal Au deposits are hosted by andesitic rocks (calc-alkaline rocks) and ages of mineralization are close to those of andesitic volcanic activity (Watanabe, 1990). The other opinion is that the mineralization is related to bimodal volcanism. For example, in west Izu Peninsula, epithermal A u - A g vein type deposits (e.g., Seigoshi) are associated with basic intrusive rocks and acidic volcanic rocks.

Present-day Mineralization and Geothermal Systems

333

Miyashita (1995) proposed that the epithermal mineralizations in Hokusatsu district (Kyushu) are related to strata volcanoes in the volcanic depression and not to caldera formation, by compiling the data on the ages of mineralization and volcanic activities, gravity and geology. He considered that the Hokusatsu volcanic depression zone is an extension of the Okinawa Trough. This volcanic depression started from 5-6 Ma (Kamata and Watanabe, 1985), where bimodal volcanism is found. Considering the above argument it is likely that the major epithermal mineralizations (Se-type and low sulfidation-type) are generally related to bimodal volcanism and not to calc-alkaline volcanism even in north and south Hokkaido and south Kyushu where andesite is dominant in the mine area, although it is thought some deposits (e.g., Takadama in northeast Honshu) formed closely related to the caldera volcanism (Seki, 1993).

2.3. Submarine geothermal systems and associated mineralization 2.3.1. Submarine metal precipitation at back-arc basins around the Japanese islands Recently, several submarine hydrothermal sites have been discovered from the seafloor of back-arc depression zones and volcanic fronts near the Japanese Islands (Okinawa Trough and Izu-Bonin) (Fig. 2.29). The studies on these areas are described below.

2.3.1.1. Okinawa Trough The Okinawa Trough (Fig. 2.30) is thought to be a back-arc basin. Sibuet et al. (1995) considered from reflection and magnetic data that the Okinawa Trough is still in a rifting stage. Normal faultings are widespread, supporting that bulk compositions of Okinawa Trough basalt, especially their large-ion lithophile element (LILE) abundances are transitional between the Ryukyu 1A basalts (LILE-enriched) and normal-type midocean ridge basalts (N-Morb, LILE-depleted) (Ishizuka et al., 1990). Honma et al. (1991) have shown that the Okinawa Trough basalts have significantly high K, Rb and Sr contents and D/H, 180/160 and 87Sr/86Sr ratios than N-Morb have and these are due to generation of magma from normal-type mantle peridotite modified by component from the subducted slab and crustal contamination. Halbach et al. (1997) reported lead isotope data on volcanic rocks, sediments and ores from the hydrothermal JADE field in the Okinawa Trough and pointed out that lead isotopic compositions of Okinawa JADE ores are very similar to Kuroko ores (Fig. 2.31) and both sediments and volcanic rocks contributed comparable amounts of lead to the deposit. Several types of mineralizations, massive-type stockwork mineralization and a sulfide-bearing sediment layer were described (Halbach et al., 1989). The outer portion of the massive sulfide sample (late-stage) is composed of barite, realgar, orpiment, amorphous silica and hydrous Fe-Mn oxides, small amounts of sphalerite, galena and pyrite. The central portion consists mainly of sphalerite, pyrite and galena with small

334

Chapter 2 80~

100 ~

120 ~

140 ~

160 ~

180 ~

160~ 60 ~

60 ~

I l e 26

40 o

4o ~

19.20 21

2o o

20 o

13"14q

24

4 ~5"6

2o o

2o~

4o o

40~

80~

100 ~

120 ~

140 ~

160 ~

180 ~

160~

Fig. 2.29. Hydrothermal mineral occurrences related to a r c - b a c k - a r c systems of the western Pacific. Numbers are listed in Table 2.12 (Ishibashi and Urabe, 1995).

amounts of chalcopyrite. The other massive sulfide samples contain sphalerite, galena, chalcopyrite and pyrite accompanied by anglesite. Sr isotopic compositions of halloysite, barite and anhydrite are similar to seawater value. These Sr isotopic data and 3D and ~180 data on kaolinite (8180 - +7.4%o, ~D = -23%o), chlorite, montmorillonite (3180 -- +7.0%o, 3D -- -32%o), and mica (~180 = +5.4%o to +9.9%o, ~D = -30%o to -26%o) suggest fluids of heated seawater origin (Marumo and Hattori, 1997).

2.3.1.2. Izu-Bonin Arc Several hydrothermal sites have been discovered on the seafloor of Izu-Bonin arc that is located at the eastern margin of the Philippine Sea plate (Fig. 2.32). This arc has been formed, related to the westward subduction of the Pacific plate (Fig. 2.32). Hydrothermal mineralization occurs both in back-arc depression and volcanic chain (Shichito-Iwojima Ridge). Hydrothermal venting and mineralizations are found

335

Present-day Mineralization and Geothermal Systems 126 ~

130~

I

I

~A

Tunghaishelf r - - "

(

-

"

.,

r

17

~/~V ~

...2,~ e , ' ,

".-

ZX

0

I , , ,,~.~';18 ~ ~ o , : : . ~ . .,, ,,. ~_, ~

,:

g., #~

,"

30 ~

,/

,=

;

/"

~

./

,7 /

r/

.~-

-

- / S o" " 7._o,

j_.~-"

26~

k. ~ i I I ! Fig. 2.30. Location and tectonic setting of Okinawa Trough. The solid circles mark the hydrothermal fields listed in Table 2.12. The broken lines show contours of 1000 m water depth. The triangles indicate Quaternary volcanoes (Ishibashi and Urabe, 1995).

in submarine caldera of calc alkaline rocks (e.g., Kita-Bayonnaise Caldera Myojinsho Caldera). Iizasa et al. (1999) reported a sulfide body which has features similar to Kuroko deposits from Myojin Knoll Caldera. The ore body is located on the caldera floor at a depth of 1210-1360 m and is notably rich in gold and silver (Table 2.9 and Table 2.10). Iizawa et al. (1999) reported native arsenic, native sulfur, orpiment, realgar, and wurtzite from active chimneys but not from the inactive Sunrise deposit chimneys (Table 2.11). This mineralogical difference indicates that these minerals dissolved during the postdepositional stage. In the Kuroko deposits, these minerals are very rare, suggesting such mineralogical modification after the primary deposition.

2.3.2. Characteristics of back-arc deposits in the Western Pacific Recently, 27 back-arc deposits have been discovered in the Western Pacific (Ishibashi and Urabe, 1995) (Fig. 2.29). The summary of geologic structure, type of deposits and mineralogy of the back-arc deposits in Western Pacific are given in Table 2.12 (Ishibashi and Urabe, 1995).

336

Chapter 2

~,

I

'

l

'

A

'

I

I

!

I

|

"39.0

,,~

I

r

38.0

38.0

I

I

208 Pb/204 Pb

208 Pb/204 Pb 39.0

i

,.

I

I

I

I

I

[

t

I

I

I

I

I

I

I

I

I

B 207 Pb/204 Pb

B 207 Pb/204 Pb

- 15.65

Kuroko JADE field

15.50

206pb / 204 Pb 18.4

I

18.6 I

I

18.8 I

I

I

I

18.8 I

206 Pb/204 Pb I

I

t

Fig. 2.31. Comparison of JADE field data with Kuroko (Halbach et al., 1997).

2.3.2.1. Tectonic settings, geologic structure and volcanic rocks Tectonic settings, geologic structure and volcanic rocks associated with the backarc deposits are summarized in Table 2.12 (Ishibashi and Urabe, 1995). Most of the deposits are located at the active back-arc spreading centres (Mariana Trough, Andaman Sea Marius Basin, Woodlark Basin, North Fiji Basin, Lan-Havre Basin). A small number of hydrothermal deposits were discovered from back-arc riffs (e.g., Sumisu rift, Northern Okinawa Trough, Middle Okinawa Trough) and volcanic fronts (e.g., Middle Okinawa Trough, Izu-Ogasawara arc).

2.3.2.2. Metal contents Chemical analytical data are summarized in Table 2.13 and Table 2.14. The back-arc deposits are characterized by higher Pb, Ba Ag, Au, As and Sb contents than midoceanic ridge deposits. This difference is due to different mineralogy which is described below.

337

Present-day Mineralization and Geothermal Systems 140 ~ '

'

'

145~

Ip?'

-

'

...,,.,

:..:;;,,-

. j #..."

'

;

I

~

-

35 ~ N

~, tt~._ \ ~ ,'

"

'.'.,

,1 er

30 ~

i

-

i

"le. t , " l

(a

"';J

#i.~, i l ,

~'.'~s

i

{;

'<'=#

-

"'~.~

.

zx

""'

i.."

t 4, ~

,~

i~'

i

i

i

I

"N.J

-

/" ,,,."

.

zx Quaternary volcano . = Backarc depression Trench

-

25~

. \

"',

.

,~f,~. \.,~,, i

, ,,, ,

,

,

I

Fig. 2.32. Location and tectonic setting of I z u - B o n i n arc. The solid circles mark the hydrothermal fields listed in Table 2.12. The broken lines delinate contours of 3000 m water depth (Ishibashi and Urabe, 1995).

2.3.2.3. Mineralogy Main opaque and gangue minerals are barite, amorphous silica, and opal-CT (Sumisu rift, Izu-Ogasawara (Bonin) arc), anhydrite, sphalerite, wurtzite, tetrahedritetennantite, chalcopyrite, galena, barite, bornite, and covellite (Northern Okinawa Trough, "Minami-Ensei Knoll"), calcite, rhodochrosite, anhydrite, wurtzite, pyrrhotite, galena, isocubanite, chalcopyrite, and argentite (Middle Okinawa Trough, "Iheya Ridge (CLAM site)"), chalcopyrite, galena, jordanite, marcasite, native arsenic, native sulfur, pyrite, realgar, sphalerite, tetrahedrite-tennantite, and wurtzite (Myojin Knoll Caldera), anhydrite, amorphous silica, pyrite, marcasite, chalcopyrite, sphalerite and wurtzite (North Fiji). Mica, kaolins (kaolinite and halloysite), Mg-rich chlorite, talc, montmorillonite and chlorite/montmorillonite are abundant in the host epiclastic sediments and pumiceous tufts at Jade in the Okinawa Trough (Marumo and Hattori, 1997).

2.3.2.4. Chemical and isotopic compositions of hydrothermal solution Chemical compositions of end member hydrothermal solution venting from backarc basins and midoceanic ridges are summarized in Table 2.14 and Table 2.15 (Ishibashi and Urabe, 1995; Gamo, 1995).

W W 00

TABLE 2.9 Average content (in parts per million or percent) of selected metals and SiO? in 37 hydrothermal sulfide samples from the Sunrise deposit (Iizasa et al., 1999)

Sulfide chimney Chalcopyrite-rich conduit wall ( N = 2) - Sphalerite-rich interior and conduit wall ( N = 22) - Anhydrite-rich rind ( N = 1) - Silica-rich rind ( N = 1) Layered sulfide ( N = 1 ) Brecciated massive ( N = 4) Disseminated sulfide ( N = 6) Maximum Average ( N = 37) -

a

Au“ Ag” As” ( P P ~ ) ( P P ~ ) (ppm)

Sb” (ppm)

Cdb (ppm)

Cub

Pbb

Znb

Fe”

Baa

SiOzb

(%)

(%I

(%)

(%I

(%)

(%)

2.5 31.0 3.5 2.9 0.8 4.0 5.1 49.0 20.0

33 5,759 380 210 65 130 658 28,000 3.564

6 2.991 102 28 145 62 401 8,930 1,858

26.70 2.94 0.29 0.12 24.60 11.92 2.40 30.70 5.54

0.08 3.64 0.87 1.42 0.03 0.04 0.22 11.00 2.27

0.34 33.90 3.79 1.36 1.77 0.82 8.85 52.10 21.89

25.30 6.45 3.42 0.99 21.20 30.50 7.92 32.30 10.48

0.10 4.03 8.50 9.90 0.70 0.07 23.26 45.50 6.70

0.11 4.31 0.33 49.40 16.00 8.55 21.06 49.40 8.68

Analyses (dry basis) by neutron activation. Analyses (dry basis) by inductively coupled plasma emission.

83 1,609 180 36 100 75 1,447 4,530 1,213

745 7,366 1,300 1.800 330 1.575 2.228 19.000 5,044

339

Present-day Mineralization and Geothermal Systems TABLE 2.10

Average content (in parts per million or wt%) of selected metals in samples from the Sunrise deposit compared with other modem seafloor deposits and an average Kuroko massive sulfide deposit (Iizasa et al., 1999) Au (ppm)

Ag (ppm)

Cu (wt%)

Fe (wt%)

Zn (wt%)

Pb (wt%)

20.0 27.7

1213 190

5.5 11.9

10.5 14.6

21.9 20.6

2.27 0.77

4.7 3.8

2806 178

2.0 5.2

10.3 10.6

20.9 21.6

11.49 0.38

2.9 2.0

80 87

7.8 9.2

24.0 34.4

11.1 6.2

0.04 0.05

0.8 4.7 1.2 0.1 0.5

124 175 80 17 50

3.4 0.3 6.0 4.7 1.5

26.8 4.9 23.0 28.7 15

5.0 18.8 14.8 0.8 3

0.11 0.36 0.08 0.01 1

Arc -fron t Sunrise deposit (N -- 37) Suiyo seamount (N -- 11)

Back-arc basin Izena cauldron (N = 14) Lau (N ---46)

Mid-ocean ridge, slow spreading TAG (N -- 31) Snake Pit (N = 36)

Mid-ocean ridge, intermediate to fast spreading Explorer (N -- 48) Axial seamount (N = 16) East Pacific Rise, 13~ (N = 6) Galapagos (N -- 17)

Average Kuroko deposit

TABLE 2.11 Selected minerals identified in the Sunrise deposit a (Iizasa et al., 1999) Active chimneys Anhydrite Barite Cerussite Chalcopyrite Covellite Galena Jordanite Marcasite Native arsenic Native sulfur Orpiment Pearceite Pyrargyrite Pyrite Realgar Sphalerite Tennantite Tetrahedrite Wurtzite

E -A E-A E-A E-A, Collo. E-A, Collo. E-A E-A E-A E-A E-A

E-A, Pell. E-A E-A, Col., Den. E-A E-A E-A

Inactive chimneys

Mounds, massive and layered

Disseminated in tuff breccia

E-A

E-A

E-A

E-A, Collo. E-A E-A, Collo. E-A E-A

E-A E-A E-A

E-A E-A E-A

E-A

E-A

E-A, Pell.

E-A, Pell.

E-A E-A E-A, Collo., Pell.

E-A, Col., Den. E-A E-A

E-A E-A E-A

E-A E-A E-A

a Present in euhedral to anhedral forms (E-A), colloform (Collo.), pelletal (Pell.), columnar (Col.), dendritic (Den.).

w

TABLE 2.12 Hydrothermal mineral occurrences in arc-back-arc systems in the Western Pacific (Ishibashi and Urabe, 19951 ____

~~

~

~ _ _ _ _ _ _ _ _ ~ _ _ _ _ ~

_ _ _ _ _ ~

Water depth (m)

Geologic structure

Type of deposit

Mineralogy

1980

Axial graben at topographic high of north-central segment near triple junction. Sheet lava floor.

Active (T = 290°C) anhydrite chimneys standing on dead sulfide mound. Forest of dead sulfide chimneys.

Anhydrite, amorphous silica in dead chimneys; pyrite, marcasite, chalcopyrite, sphalerite, wurtzite, goethite.

2 North Fiji Basin. 'Station 14' (18"50'S, L73"30'E)

2720

Collapsed lava lake on flat rise crest of fast spreading south-central segment. No sediment cover.

Warm ( T = 5.2"C) fluid discharge through mussel bed. No hydrothermal minerals. Site of megaplume (Noriji et al., 1989).

None.

3 Fiji Transform Fault 'Extensional Relay Zone A' (16"lO'S. 177"25'E)

1869-2335

Short spreading ridge axis which displaces Fiji uansform fault as interpreted by Jarvis et al. (1994).

Hydrothermal sulfide impregnation in MORB-like basalt dredged from axial valley.

Magnetite, pyrrhotite, chalcopyrite and opal on fracture surface.

4 Central Manus Basin 'Vienna Woods' (3"10'S, 150"17'E)

2500

2-km-wide axial rift graben of the northeast spreading center. Mostly massive pillow lava floor.

Sulfide chimneys up to 20 m high are venting clear. milky and black fluids. Sulfate smokers are also present.

Sphalerite. wurtzite, pyrite. marcasite, chalcopyrite, galena, amorphous silica, barite. Sulfate chimney; anhydrite, silica, barite.

5 Eastern Manus Basin 'Desmos cauldron' (3"42'S, 151"52'E)

2000

Ferruginous oxide deposits. Caldera of basalt/basaltic andesite a1 Sulfide ores were not recovered. an intersection of a spreading center Megaplume-like methane anomalies in Pyrite and native sulfur water column over the caldera. disseminated in basaltic andesite. and a transfonn fault?

6 Eastern Manus Basin 'Pacmanus field' (3"42'S, 15Io42.6'E)

1650

Crest of a prominent ridge of dacite flows and domes called Pual Ridge which exists in pull-apart basin.

4-m-high sulfide chimneys venting 'smoke' (no temperature information). Only one ore chip (1 cmj was recovered.

Cohesive anhydrite, chalcopyrite, bornite, tennantite and sphaletite.

7 Valu Fa Ridge, Lau Basin 'Vai Lili site' (22"18'S, 176"35'W)

1700

300-m-long zone along a normal fault on a ridge crest. Basaltic andesite to rhyodacitic lava.

Black and white smokers venting up to 400°C fluid. Height up to 17 m. Massive Zn- and Cu-sulfide ores are uresent.

Pyrite, chalcopyrite, marcasite, sphalerite, barite, tennantite, galena. aragonite, native gold.

Location and number

Back-arc spreading center 1 North Fiji Basin, 'Station 4' (16"59'S, 173"55'E)

9

8'3 -?

TABLE 2.12 (continued)

m4

~~

Location and number

Bark-arc spreading center 8 Valu Fa Ridge, Lau Basin 'Hine Hula site' (72"33'S, 176V3'W) 9 Valu Fa Ridge, Lau Basin 'White Church site' (21"55'S, 176"32'W)

WLtter depth (nil

Geologic structure

1900

On south Valu Fa Ridge crest dominated by andesite and dacite. Strongly bleached volcanic rocks.

1946- 1966

Type of deposit

Mineralogy

E-

'1

At the top and along the flanks of northern Valu Fa Ridge. Brecciated and pillow lava, and pumice field.

Low-temperature Mil-oxide crust is covering high-temperatuue fossil sulfide deposits. Diffusive fluid iT = 40T).

10-1 5-m-high inactive barite-sulfide chimney field extends over 300 ni along nornial faults.

d-MnO:. birnessite. todorokite. amorphous Fe-oxide. goethite, sphalerite. barite. pyrite. chalcopyrite. Barite and sphalerite are dominant with lesser galena. tennantite, chalcopyrite, gold and pyi-rhotite.

2

2

t l

8. 9 ... n %

4

6 9 s

Dredged hyaloclastite basalt fragincnts which are cemented by opal and euhedral barite crystals.

Opal. barite, montmorillonite. philipsite (oxygen-isotope study suggests temp. of 117-131"C).

17 Western Woodlark Basin 2113-2366 'Franklin Seamount' (9"55'S. 15I"50'W)

Wcsternmost propagating tip of spreading center. Basaltic andesite and inferred sodic rhyolite.

Spires and mounds of Fe-Mn-Si oxide up to several meters thick and 200 in in extent. Venting 20-30°C clear solution.

Inactive barite silica chimneys contain up to 21 ppm Au. Si-braring Fe oxyliydroxide.

13 Central Mariana Trough 'Alice Spring field' and other (18"12'N, 143"10'E)

3600-3700

Flank of an axial volcauo of basaltic andesite on back-arc spreading center. Volcanics are fractionated.

Active barite-rich sulhte chimneys. I m high. venting clear fluid ( T = 287°C).

Barite chimneys. Sulfides; abundant sphalerite with lesser amounts of galena, chalcopyrite, pyrite.

11 Central Mariana Trough ( I S'02'N. 144"15'E)

3675

Crest of an axial ridge where the relief is greatest (ca. 800 ni). A low

Several active chimneys surrounded by a low inound of hydrothermal

Sphalerite, barite. amorphous silica.

mound. 20-30 m in diameter.

precipitates.

1664-1 900

m

barite. and amorphous silica. Thin film of Mil-oxyhydroxide.

At a fracture zone ofrsetting northwestcrly trending active spreading center of the northern Lau Basin.

I 1 Peggy Ridge, Lau Basin (66'35's. 176'49.5'W)

3

Wurtzite, pyrite, chalcopyrite.

Dredged black smohrr chimney samples.

2100

%

3

Axial region of northeasterly trending active spreading ridge of the northern Lau Basin.

10 Northeastern Lau Basin 'Papatua expedition site' i IS"17'S. 171"35'W)

w

7

2 i

5

h '+

2

2

'A

5

342

Chapter 2

TABLE 2.13 Average bulk compositions of samples from seafloor sulfide deposits at seamounts and back-arcs (Scott, 1997) Axial Seamount CASM

North Fiji Basin

Mariana Trough at 18~

ValuFa, Lou Basin

Eastern Manus Basin

Jade Okinawa Trough

14 B

24 B

11 A

47 B,A,D

26 D

17 R

Zn Cu Pb Fe SiO2 Ba Ca

22.2 0.4 0.35 5.6 28.1 9.6 0.21

6.6 7.5 0.06 30.1 16.2 0.8 0.2

10.0 1.2 7.4 2.4 1.2 33.3 3.7

16.1 4.6 0.3 17.4 12.5 11.6 0.6

26.9 10.9 1.7 14.9 0.8 7.3 0.3

24.5 3.1 12.1 4.8 10.2 3.4 -

(ppm) Ag Au Hg Cd Sn As Sb

189 4.9 20.2 522 7 569 349

151 1.0 260 < 10 -

184 0.8 22 465 126 190

256 1.4 >1 482 4 2,213 51

230 15 17 1,155 11,000 1,130

1,160 3.3 620 31,000 -

No. a:

Host b:

(wt%)

a Number of samples analyzed. b A = andesite; B = basalt; D = dacite; R -- rhyolite.

The chemistry of h y d r o t h e r m a l solutions from midoceanic ridges has been reasonably explained by the effect of buffering by alteration minerals (Seyfried, 1987; Berndt et al., 1989). Therefore, it might be worth explaining the chemical composition of h y d r o t h e r m a l solutions from back-arc basins in terms of chemical equilibrium b e t w e e n h y d r o t h e r m a l solutions and alteration minerals. The H2S concentration of h y d r o t h e r m a l solution is plotted in Fig. 2.33. Based on these data, we can estimate the temperature of h y d r o t h e r m a l solution buffered by alteration mineral a s s e m b l a g e s such as a n h y d r i t e - p y r i t e - c a l c i t e - m a g n e t i t e and p y r i t e p y r r h o t i t e - m a g n e t i t e for O k i n a w a fluids. For example, a s s u m i n g anhydrite-magnetite-calcite-pyrite-pyrrhotite buffers redox in sub-seafloor reaction zones and a pressure of 500 bars, dissolved H2Saq concentrations of 21 ~ E P R fluid indicate a temperature of 3 7 0 - 3 8 5 ~ However, the estimated temperatures are higher than those of the m e a s u r e m e n t . This difference could be explained by adiabatic ascension and probably conductive heat loss during ascension of hydrothermal solution from d e e p e r parts where chemical c o m p o s i t i o n s of h y d r o t h e r m a l solutions are buffered by these assemblages. It is well k n o w n that anhydrite, pyrite, magnetite, and epidote are widespread in the country rocks in the K u r o k o mine area. Therefore, it is likely that this assemblage controls the c h e m i s t r y of the h y d r o t h e r m a l solution.

343

Present-day Mineralization and Geothermal Systems TABLE 2.14 Chemical compositions of hydrothermal fluids in the Western Pacific (Ishibashi and Urabe, 1995) Location No."

Temp. (~ pH Li (~tM) Na (mM) K (mM) Rb (~M) NH4 (mM) Mg (mM) Ca (mM) Sr (IxM) Ba (txM) Mn (~M) Fe (BM) Cu (~tM) Zn (~M) Pb (nM) S04 (mM) C1 (mM) Br (~M) Alkalinity (meq) B (raM) A1 (IxM) Si (mM) H2S (mM) CO2 (mM) 87Sr/86Sr 834S (H2S) (%0)

1 7 13 16 N. Fiji White Lau Vai Mariana Alice O k i n a w a Lady Lili Spring Minami-Ensei

18 21 O k i n a w a Izu-Bonin Izena Suiyo

285 4.7 200 210 10.5 8.8 0 0 6.5 30 5.9 12 13

320 4.7 600 446 73.7 28 5.32 0 23.2 110 60 370 21 0.003 7.6 36 0 550 1045 0.88 3.41 4.9 12.5 13.1 200 0.7089 7.4-7.7

0 255 306 0.12 0.47 6.0 14.0 2.0 14.4 0.7046

334 2.0 623 590 79.0 68

285 4.4 5800 431 31.0 30

0 41.3 20 >39 7100 2500 34 3000 3900 0 790 1140

0 23.6 73

0.83 14.5 <5 15 0.7044

0 544 866 0.43 0.81 14.0 2.6 0.7033 3.6-4.8

278 5.0 1860 430 50.9 360 4.70 0 22.1 227 55 94

0 527 3.51 4.0 10.8 2.44 96 0.7100 3.6

311 3.7 891 432 29.7 <0.1 0 89.0 300 100 587 435

0 658 -0.20 1.43 17.0 13.2 1.6 40

OBS EPR 350 3.4

23.2

0 15.6 81 8 960 1660 35 106 308 0 489 802 -0.40 0.51 5.2 17.6 7.3 8 0.7031 1.3-5.5

A few R E E data on h y d r o t h e r m a l solutions are available (Fig. 2.34). Chondrite n o r m a l i z e d R E E patterns of h y d r o t h e r m a l solutions from Vienna Wood, P a c m a n u s and D e s m o s , M a n u s Basin exhibit positive Eu a n o m a l y and L R E E e n r i c h m e n t are similar to midoceanic ridge solution and Kuroko ore fluids. This positive Eu a n o m a l y (Fig. 2.35) m a y have b e e n caused by the selective leaking of Eu due to the interaction of an ascending h y d r o t h e r m a l solution and footwall volcanic rocks (Gena et al., 2001). It is interesting to note that altered basaltic andesite has a negative Eu a n o m a l y and this feature is the same as that found in the Kuroko m i n e area (Shikazono, 1999). Isotopic compositions of h y d r o t h e r m a l solutions are s u m m a r i z e d in Table 2.15. It is notable that 834S of HzS of the h y d r o t h e r m a l solution at O k i n a w a Trough (Jade site) is high (+7.4%0 to +7.7%0) which is similar to that from the K u r o k o deposits. 313C of H2CO3 in h y d r o t h e r m a l solution is - 6 . 2 % 0 to -3.6%0. This range is higher than igneous value (-7%0) and lower than marine value (+1%~), suggesting both (igneous source, marine source) contribution.

TABLE 2.15 Submarine hydrothermal fluid end-member data observed at various plate boundary regions (end-member values are those obtained by extrapolation to an assumed value of zero magnesium, except for the values at site #2). BAB: back-arc basin (Gamo, 1995) Site No.: Name of site:

Type: Sediment: Sampling year (submersible): Depth (m): Location (Lat.): Location (Long.): Smokers: Temp. ( " 0 : Spreading half rate (cmlyear): pH (25°C): Li (mM) Na (mM) K (mM) Rb (wM) Cs (nM) Be (nM) Mg (mM) Ca (mM) Sr (wM) Ba (IN Mn (IW Fe (LM ) Co (nM) Cu ( L W Ag (nM) Zn (I*.M) Cd (nM) B (mM) A1 (LM) SiO? (mM) Ge (nM) Pb (nM)

-7

1 Okinawa Trough JADE BABorArc -hosted 1989 (Sh2K)

Okinawa Trough CLAM BAB -hosted 1989 (Sh2K)

3 Okinawa Trough South Ensei BAB -hosted 1992 (Sh2K)

1340 27"16'N 127"04'E Black 320 2

I390 27"33'N 126"58'E Clear 220 2

710 28"23'N 127"38'E Clear 267- 178 2

1380 28"34'N 140"39'E Clear 296-3 I I

4.7 2.5 425 72 360

5.3

4.9-5.1 5.4-5.8 410431 49-5 1

3.7 0.6 446 30

0 22 94 59 110 2.8 2.7

(-20) (-20)

0 21-22 2 15-227 53-56 88-94

0 89 303

(34) (5040) (180-220)

(400-500)

4 Izu-Bonin Suiyo SM. Arc - starved 1992 (Sh2K)

100 587 435

5 Mid-Mariana Trough Alice Springs BAB -starved 1987 (Alvin) 1992 (Sh6.5K) 3600 18"13'N 144"42'E Clear 238-287 3 3.94.4 0.594.83 438 3148 30 800 27 0 22 72-90 24 295 6.4

6 South Mariana Trough Forecast vent Arc - starved 1993 (Sh6.5K)

Vienna Woods BAB - starved 1990 (Mir)

1490 13"24'N 143"SS'E Clear 220 3

2500 3"IOS 150" 17'E Grey 275 6

0.3 438 26

0.7

0 61 165 < 10 300

0 73

11

0.003 >7.6 1.3 3.4

3.74.0

12.9

10.4-10.8

36

1.4 17 13.2

0.71-0.83 7.0 12.3-14

7.7

7 Manus Basin

297 77 6 0.04 6 24 9

15

30

W

%

8 9 North Fiji Basin Lau Basin

BAB -starved 1989 (Nautile) 1991 (Sh6.5K) 2000 16"59'S 173"55'E Clear 285-29 I 3.5

Vai Lili BAB - starved 1989 (Nautile) 1720 22"15'S 176"35'W White 280-334 3

4.7 0.20-0.28 2 10-239 10.5-14.5 8.8-17

2 0.58-0.75 520-615 55-80 59-75 1060-1560

0 6.5-9.0 30-43 5.3-5.9 12-26 9-13

0 28-41 20-135 >20 to >60 5800-7 100 1 160-2900 15-35

0.46-0.47 6 13.3-14.0

1200-3100 700-1500 0.77-0.87

n

0.1s

0.4 42.;

7. _ i

-I

1.1-1.5

0c u

I1

0.20

n.?- I

~

Kanc c?f ,?e: T.

F.

Jlinn ce F ~ s Ride< a Sliddlc \2l.. Dezd Dog. Best Hiii XIOR

de Fuca RiCee Endesrour J 3 x

110K

Juan d e Fuc3 Ridge Juan de Fuca Ridge luzn de Fuza Ridgs Junn dc F u G ASHES .ASHES .ASHES Rides :Clsnriihrd! :CI-norma: : :c1-dspklsdI Y. CItfI SSE.

510R

hfOR

%TOR

154'

1.v2 45'56.5 I 3o"o 1' w Clear ?k-299 3

15'56 K I?o-vl w

Clear-BkcL 293-323 7

--

1 :->.I

4.4

I2 :9-0.5s 391-M 2 I -:5

0 1 8 4 29

145-109 -.0-11.i

0 1C-3 4&s I

6 14€!-?9cl

12-3 I

0.147

99 1001i-I

2.2-2 6 0.454.33

13.5 I

3 10

510R

Jcnn de FUC3 Ridgs s.CIS? SSS. \iOR

Ex-nnabz Trollgh \10R

I.-I-'.@

4-12

7.1 50

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87Sr/86Sr of hydrothermal solution is in a range of 0.7044-0.7093, suggesting igneous (0.706-0.707) and marine (0.7091)origin.

2.3.3. Comparison of present-day back-arc deposits with Kuroko deposits Among the back-arc deposits, the features of the Okinawa Trough deposits are similar to those of Kuroko deposits, but North Fiji and Mariana deposits are different. For instance, gold, silver, zinc, lead and arsenic are rich in the Okinawa Myojin-Knoll Caldera and some Kuroko deposits but poor in the North Fiji and Mariana deposits. The difference in mineralogy of the Kuroko and present-day back-arc deposits are: (1) metastable phases such as native sulfur, wurtzite, and amorphous silica are poor in the Kuroko deposits; (2) arsenic minerals such as jordanite, tetrahedrite-tennantite, native arsenic, and realgar are common in the present day back-arc deposits (Okinawa Myojinsho Knoll Caldera), but rare in Kuroko deposits except tetrahedrite-tennantite; (3) secondary minerals such as cerussite and covellite are common in present day back-arc deposits (e.g., Okinawa, Myojinsho Knoll Caldera); (4) Dendritic texture is common in the present day back-arc deposits. This difference indicates that primary texture and mineral assemblages in the Kuroko deposits were modified after the formation of ore deposits.

2.3.4. Spatial relationship between back-arc deposits and epithermal gold deposits As noted already, the formation of polymetallic vein-type deposits and Kuroko deposits occurred under the subaerial and submarine environments, respectively, at nearly the same time (middle Miocene).

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This spatial difference in the distribution of back-arc and vein-type deposits is also found for present day mineralizations. For example, back-arc deposits are forming at Okinawa Trough, while base-metal and precious-metal precipitations are occurring on land such as at Ibusuki and Beppu geothermal area, Kyushu. Recent Au mineralizations (ca. 1 Ma) occurred at Noya and Hishikari areas in south Kyushu (Fig. 2.36). These mineralizations can be regarded as almost contemporaneous mineralization with the Okinawa mineralization. This kind of temporal and spatial relationship between epithermal Au vein-type

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mineralization and back-arc mineralizations are also found in the I z u - B o n i n area. Seafloor base metal mineralization similar to Kuroko mineralization is found in the I z u - B o n i n seafloor (Iizasa et al., 1999), whereas epithermal gold vein-type mineralization occurred at a very recent age (ca. 1 Ma) at the Izu Peninsula. For example, the mineralization ages are 1.1-1.8 Ma (Seigoshi), 1.0 Ma (Toi), 2.5 Ma (Yugashima), 2.0 Ma (Daimatsu), and 1.4-1.5 Ma (Rendaiji), respectively. It seems likely that epithermal Au mineralization (low sulfidation-type) forms in volcanic depression zone (back-arc on land) under the extensional stress regime rather than compressional stress regime. This spatial difference is consistent with the distribution of hydrothermal deposits of middle M i o c e n e (Kuroko and epithermal vein-type deposits in Japan). As mentioned in section 1.3, Kuroko deposits formed under the submarine environment, while polymetallic

Present-day Mineralization and Geothermal Systems

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vein-type deposits formed in central and Northwest Japan (Ashio, Tsugu, Kishu, Obira, etc.) under the subaerial environment. This submarine vs. subaerial hypothesis for the origin of the two types of deposits (Kuroko deposits, epithermal vein-type deposits) can reasonably explain the difference in metals enriched into the deposits by HSAB (hard-soft acids and bases) principle proposed by Pearson (1963) (Shikazono and Shimizu, 1992). Relatively hard elements (base metal elements such as Cu, Pb, Zn, Mn, Fe) are extracted by chloride-rich fluids of seawater origin, while soft elements (Au, Ag, Hg, T1, etc.) are not. Hard elements tend to form chloro complexes in the chloride-rich fluid, while soft elements form the complexes in HzS-rich and chloride-poor fluids. C1 in ore fluids is thought to have been derived from seawater trapped in the submarine volcanic and sedimentary rocks.

Chapter 2

354

2.4. Comparison of back-arc deposits with midoceanic ridge deposits 2.4.1. Hydrothermal solution 2.4.1.1. Chemical compositions Chemical compositions of major elements (alkali, alkali earth elements, Si) in back-arc and midoceanic ridge hydrothermal solutions are not so different (Table 2.15). This is thought to be due to the effect of water-rock interaction. For example, Berndt et al. (1989) have shown that mca2+ / m2+ of midoceanic ridge hydrothermal fluids is controlled by anorthite-epidote equilibrium (Fig. 2.37). Figure 2.37 shows that mca2+/m2+ of back-arc hydrothermal fluids is also controlled by this equilibrium. However, there are slight differences in inK+/tuNa+, tuBa2+/msr2+, and msr2+ ~inCa2+. These differences are considered to be attributed to the differences in compositions of rocks and alteration minerals interacted with circulating seawater or modified seawater at elevated temperatures. For example, high K and Li concentrations in the hydrothermal solution in the M i d - O k i n a w a Trough back-arc basin (Jade site) are due to the interaction of hydrothermal solution with acidic volcanic rocks (Sakai et al., 1990). It is evident that the chemical compositions of hydrothermal solution are largely affected by water-rock interaction at elevated temperatures. The concentrations of elements in hydrothermal solution depend not only on the compositions of rocks, but also on temperature. It is well known that SiO2 concentration in hydrothermal solution increases with increasing of temperature and it can be used as a geothermometer.

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355

Present-day Mineralization and Geothermal Systems

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Fig. 2.38 shows the temperature estimated by silica geothermometer for backarc basin and midoceanic ridge hydrothermal solutions. This shows that the SiO2 concentration of back-arc basin fluids is lower than that of midoceanic fluids (e.g., East Pacific Rise). This relation indicates that the SiO2 concentration is controlled by quartz in the rocks. Ascending hydrothermal solution interacts with wall rocks and the SiO2 concentration in fluids change, depending on temperature. However, SiO2 concentration deviates from the equilibrium concentration. This deviation relates to adiabatic ascending, flow rate of fluids and kinetics of SiO2 precipitation and dissolution of silicate minerals in the wall rocks (Wells and Ghiorso, 1991). Na-K and N a - K - C a geothermometers are useful for the estimate of temperature of geothermal waters in terrestrial geothermal systems (Fournier and Truesdell, 1973). According to Na-K geothermometer, mNa+/mK+ in hydrothermal solution decreases with increasing temperature. Fig. 2.39 summarizes the relationship between mNa+/mK+ and temperature in basalt-hosted hydrothermal sites in midoceanic ridges and back-arc basins. This indicates that mNa+/mK+ and temperature of midoceanic ridge fluids are higher than those of back-arc fluids. If mNa+/mK+ is controlled by the equilibrium between Na-feldspar and K-feldspar, NaA1Si308 + K + -- KA1Si308 + Na +

(2-24)

Equilibrium constant for reaction (2-24) is expressed as, K2-24 -- (aNa+aK-f)/(aK+ / mNa-f)

(2-25)

where aNa+, aK+, aK-f, and aNa-f are activities of Na +, K +, K-feldspar component and Na-feldspar.

356

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Fig. 2.39. Na+/K + atomic ratios of well discharges plotted at measured downhole temperatures. Curve A is the least squares fit of the data points above 80~ Curve B is another emperical curve (from Truesdell, 1976). Curves C and D show the approximate locations of the low albite-microcline and high albite-sanidine lines derived from thermodynamic data (from Fournier, 1981). Small solid: subaerial geothermal water; Solid square: Okinawa Jade; Open square: South Mariana Through; Solid circle: East Pacific Rise ll~ Open circle: Mid Atlantic Ridge, TAG.

It is reasonably approximated that ~,,la+/YK+ (where y - activity coefficient) is one. Thus, mNa+/mK+ = K2-24(aNa-f/aK-f)

(2-26)

The aK-f of plagioclase from MORB is generally lower than that of back-arc basin igneous rocks. Therefore, the higher mNa+/mK+ of midoceanic ridge fluids compared with back-arc basin fluids could be explained in terms of Na-feldspar/K-feldspar/hydrothermal fluids equilibrium. The differences in base metal concentrations in the two types of hydrothermal solution are unclear, probably because of scarcity of data. However, it seems obvious that mFe/mMn ratio of midoceanic ridge hydrothermal solution is higher than back-arc hydrothermal solution (Table 2.15). This may be due to the differences in Fe and Mn contents of volcanic rocks at back-arc basin and midoceanic ridges and temperature of fluids. Gamo (1995) showed that mFe/mMn of hydrothermal solution from sedimenthosted hydrothermal site is high (Fig. 2.40). This suggests that Mn concentration of hydrothermal solution increased by the interaction with sediments (probably dissolution of Mn-oxides and hydroxides). Gamo (1995) showed that Rb, Li and Cs concentrations of sediment-hosted fluids are apparently higher than those of sediment-starved fluids (Fig. 2.41). This suggests that Br, Rb, and Cs are supplied from sediments. The difference in 87Sr/86Sr ratios between the sediment-hosted systems (0.70420.7100) and the sediment-starved systems (0.7028-0.7063) (Table 2.15) is also interpreted in terms of the difference in sediment contribution. For example, the 87Sr/86Sr ratios at

357

Present-day Mineralization and Geothermal Systems I

I

I

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7

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Okinawa Trough and at Escanaba Trough are even higher than that of seawater (0.7092), strongly suggesting the contribution of sedimentary material with high 87Sr/S6Sr ratios (Gamo, 1995). The pH of hydrothermal solution in sediment-hosted sites is higher than that of the other volcanic rock-hosted sites. This might be due to the interaction of hydrothermal solution with sediments (probably carbonates). If pH increases by the interaction with sediments, sulfides tend to precipitate at the subsurface environment due to an increase in pH. The low concentration of base metals in hydrotherrnal solution in sediment-hosted deposits could be explained by such subsurface depositions. Although the data are not plentiful, it is clear that the hydrothermal solutions of sediment-hosted ridges and back-arc basin covered by sediment (Okinawa Trough) contain high amounts of ammonium (2.8-13.6 Ixmolal) (Scott, 1997). This means that ammonium was derived by thermal maturation reaction of organic matter in sediments by the following reactions (Gamo et al., 1991), (CH20)106(NH3)16(H3PO4) -- 53CO2 + 53CH4 -4- 16NH3 + H3PO4

(2-27)

NH3 + CO2 Jr- H20 -- NH + + HCO 3

(2-28)

358

Chapter 2 400

(a)

300

./

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::I.

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100

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0

~,

1

l

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~"

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Phase separation and segregation are occurring in some hydrothermal systems (Gamo, 1995) which modify the chemistry of initial hydrothermal solution. Von Datum and Bischoff (1987) and Butterfield et al. (1994) obtained high chloride concentration of more than twice of the seawater value for the hydrothermal solution at the North Cleft Segment and the South Cleft Segment of the Juan de Fuca Ridge. Low chloride concentrations of hydrothermal solutions were obtained from North Fiji Basin (Grimaud et al., 1991) and the Endeavour Segment of the Juan de Fuca Ridge (Butterfield et al., 1994). This wide variation in chloride concentration could be explained in terms of gas-liquid separation at the shallow part from the seafloor (Von Damm and Bischoff, 1987; Von Damm, 1988; Cowan and Cann, 1988). Ishibashi et al. (1994a) analyzed the chemical composition of white smoker from which anhydrite is precipitating at the North Fiji Ridge and showed that the chloride

Present-day Mineralization and Geothermal Systems

359

concentration is about a half of seawater and iron and manganese concentrations are low. They considered from these chemical features that the gas phase separation and segregation controlled the compositions of the fluids. The pH of hydrothermal solution of white smoker from which anhydrite is precipitating shows very low: 2 for Lau Basin Vail Lili fluid. This low pH cannot be explained only by water-rock interaction process. One likely explanation is decreasing of pH due to precipitation of sulfides. The pH decreases by the following reaction, M 2+ nt- H2S ~ MS + 2 H +

(2-29)

where M is metal element such as Zn and Pb. However, this mechanism seems to be unlikely, because hydrothermal solution originated from vapor phase does not contain appreciable amounts of base metals. The more likely process causing low pH is injection of volcanic SO2 gas. SO2 gas reacts with H20 to generate H + by the reaction,

SO2 nt- H20 -k- 1/2 02 --+ 2 H + + SO 2-

(2-30)

4 802 -t- 4 H20 --+ H2S nt- 3 SO 2- + 6 H +

(2-31)

Gena et al. (2001) reported advanced argillic alteration of basaltic andesite from the Desmos caldera, Manus back-arc basin which was caused by interaction of hot acid hydrothermal fluid originated from a mixing of magmatic gas and seawater. It is noteworthy that the acid alteration is found in back-arc basins (Manus, Kuroko area) but not in midoceanic ridges.

2.4.1.2. Isotope data (3180, tSD, t$34S~t~13C~87Sr/86S~ 3He/4He, t$11B, ~$4Li) Isotope data on hydrothermal solution from back-arc basins and midoceanic ridges are summarized in Table 2.15 (Gamo, 1995; Scott, 1997). 3180 and ~D values are close to seawater values. However, some data deviate from seawater values. This is caused by the separation of vapor and liquid and interaction of fluids with rocks. ~34S data on H2S and sulfides from Okinawa Trough (Okinawa Izena) show a high ~34S value (+8%0 (Sakai, H. pers. comm cited in Ishibashi and Urabe, 1995). ~348 values from other districts are similar to those of midoceanic ridges: +2.1%~ to +3.1%o, Mariana Trough (Kusakabe et al., 1990); +0.3%o to +2.2%0, Minami Ensei Knoll (Nedachi et al., 1992; +2.1%o to +2.8%~, Kita-Bayonnaise caldera (Iizasa et al., 1992); +0.9%~ to +1.2%o (Kaikita caldera) (vein part) (Ishibashi and Urabe, 1995). ~348 values from the Okinawa (Izena) sulfides are higher than any of ~34S data from midoceanic ridge sulfides and H2S of hydrothermal solutions. Kawahata and Shikazono (1988) summarized ~348 of sulfides from midoceanic ridge deposits and hydrothermally altered rocks (Fig. 2.42). They calculated the variations in ~348 of H2S and sulfur content of hydrothermally altered basalt as a function of water/rock ratio (in wt. ratio) due to seawater-basalt interaction at hydrothermal condition (Fig. 2.43) and showed that these variations can be explained by water/rock ratio. The geologic environments such as country and host rocks may affect ~348 variation of sulfides. For example, it is cited that a significant component of the sulfide sulfur could

360

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Fig. 2.42. Sulfur isotope values of seawater, fresh basement rocks, and sulfides from various submarine hydrothermal areas (Kawahata and Shikazono, 1988).

5

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Fig. 2.43. Graphical illustration of sulfur isotope values of H2S (left axis and solid line) produced during basalt-seawater interaction at various water/rock ratios. Calculations assume that seawater sulfate is mostly removed as anhydrite, that any residual sulfate is reduced by iron oxidation in reacting basalt, and that there is quantitative leaching of basaltic sulfide (+0.3%(0 and homogeneous mixing of both sulfides. Dashed line represents the relation between sulfur content of greenstones (right axis) and water/rock ratio (Kawahata and Shikazono, 1988). be derived from seawater sulfate which was reduced under the h e m i p e l a g i c sedimentary pile and ~34S values of sulfides under these e n v i r o n m e n t s are higher than those of igneous rock piles ( G o o d f e l l o w and Blaise, 1988). ~13C data on CO2 in h y d r o t h e r m a l solution from back-arc and midoceanic ridge are not different except G u a y m a s Basins data. ~13C of h y d r o t h e r m a l solution from G u a y m a s Basins are in a range o f - 6 . 0 % , ~ to +2.7%,~. L o w ~ 13C value is probably due to the dissolution of carbonates in sediments.

Present-day Mineralization and Geothermal Systems

361

87Sr/86Sr of hydrothermal solution from Okinawa Trough is high (0.7089, JADE; 0.7100, South Ensei) and quite different from that of midoceanic ridge hydrothermal solution (0.7028-0.7041 for sediment-starved ridge and 0.7052-0.7070 for sediment ridge) (Scott, 1997). However, 87Sr/86Sr ratios of other back-arc hydrothermal solutions are not different from midoceanic ridge data: Valu Fa (0.7044), Mariana (0.7053-0.7093, av. 0.7076), N. Fiji Basin (0.7046) (Scott, 1997). Higher 3He/4He ratio of hydrothermal solution of hydrothermal solution in the Western Pacific than midoceanic ridges is probably due to the influence of the ocean island basalt (OIB) component into the magma source (Ishibashi et al., 1994b). 3He/4He ratios of hydrothermal solutions from the Izu-Bonin arc Suiyo seamount and Mariana Trough (Central Peak) are similar to those of midoceanic ridge hydrothermal solution (East Pacific Rise, Mid-Atlantic Ridge). Tsunogai et al. (1994) thought that this similarity is due to minor contamination by crustal radiogenic helium from the thin continental crust of the Izu-Bonin intraoceanic island arc. ~llB values of hydrothermal solution from back-arc basin are lower than those from midoceanic ridge hydrothermal solution (Gamo, 1995). ~11B values of Okinawa Trough hydrothermal solution are particularly low (-5%0 to - 1 0 % 0 (Ishikawa and Nakamura, 1993), suggesting a contribution of sedimentary boron to hydrothermal solution. 34Li values of Mariana hydrothermal solution (-8.5%0 are similar to average value ( - 9 % 0 of midoceanic ridge hydrothermal solution (Elderfield and Schultz, 1996). This value can be explained by the constant mixing ratio of basaltic Li (36Li = -4%0) and seawater Li (~4Li = -32.3%0 (Elderfield and Schultz, 1996). 2.4.2. Metal ratios and mineralogy Bulk compositions of midoceanic ridge deposits and back-arc deposits are summarized in Tables 2.16 and 2.17. It is clear that midoceanic ridge ores contain higher amounts of Fe, Mn, Zn, Co, Ni, Se and Pt but lower amounts of Au, Ag, Cu, Pb, Ba, As and Sb compared with back-arc deposits (Tables 2.18 and 2.19). These differences are caused by the influences of fo2, fs2, pH, temperature of ore fluids and chemical compositions of rocks interacted with seawater in hydrothermal system. Although mineralogy is different in different site, the more abundant minerals in back-arc deposits than in midoceanic ridge deposits are barite, anhydrite, electrum, As-minerals (realgar, orpiment), tetrahedrite-tennantite and galena. The iron content of sphalerite from back-arc deposits is lower than midoceanic ridge deposits. Pyrrhotite and wurtzite are not common in back-arc deposits, although they were identified from Iheya Ridge, Middle Okinawa Trough, and Mariana. Based on iron content of sphalerite, assemblage of iron minerals (magnetite, chlorite, hematite), assuming SS and pH values, we can place a limit on typical ranges of /O2 and /82 of back-arc and midoceanic ridge fluids which are shown in Fig. 2.44. Generally, the /O 2 and fs2 of back-arc fluids are clearly higher than those of midoceanic ridge fluids (E.ER. N21 ~ etc.).

TABLE 2.16 Chemical composition of selected mineral assemblages from midocean ridge sulfide deposits (Hannington et al.. 1995). N : number of analyses

Galapagos Rift Pyrite-marcasite Pyrite-sphalerite

(22) (6)

5.1 2.2

36.1 19.7

1.4 13.8

0.04 0.06

35.4 24.5

16.9 33.8

13"N, East Pacific Rise 284-335°C chimneys ( 5 ) 245-284°C chimney (1) Pyritexhalcopyrite (2) Sphalerite-pyrite (5)

9.1 0.8 6.0 0.9

13.7 0.2 16.7 24.6 0.4 35.7 20.1 31.4

0.01 0.08 0.02 0.18

27.9 33.5 39.9 35.5

-

21"N,East Pacifc Rise Black smoker (2) Sphalerite-pyrite (2)

0.2 1.1

2.3 21.5

0.9 31.1

0.05 0.18

24.1 38.8

-

0.02 0.05

-

4.7

0.03 0.06

0.1 0.8

0.21 25 0.85 122

152 94

16.5 6.7 <0.l <0.01 0.6

0.22 3 1.20 155 0.05 19 0.45 163

10 119 86 412

26.1<0.10 6 1.9 <0.20 118

7 493

-

8 24

360 24

78 18

17 13

50 323

141 87

413 1015

26 33

- 174 - I90 - 3500

-

-

-

-

-

9 817 19 835

-

38

96 3 186 8

72

310

1678 832
2 33

29 4

5 91

2 4

40 840

1

45

36 296

3418 14

16

27

-

886

30

775

5

1I

13

8

519

13

970

53

~

11"N, East Pacific Rise Sphalerite-pyrite ( 1 1) Endeavour Ridge Black smoker (4)

1.9

22.4

28.0

0.07

35.7

1.2

0.06 ~ 0 . 1 0.15

38

399

30

8.8

-

0.3

0.03

-

3.8

-

5.0

22

-

-

Southern Juan de Fuca (3) Sphalerite-pyrite

0.2

19.7

36.7

0.26

39.3

5.1

0.06 <0.1

0.11 178

359

18

0.12

-

-

-

TABLE 2.16 (continued) IN)

Cu

Fs

i%) iai

Axial Seamount Sphalerite-marc3site Barite-silica

(9)

0.5

6.2

7

0.1

3.3

P!~ritc-chalco~!a-ii~ 115) (7) Pyrite-mllrcazitc is) Marcasitr-sphalerite Sphalerltr-iiiarc~ihite (121 Barite-silica (6)

8.0

37.7

Zn (5%)

Ph

S

iaj)

(5?)

S102 (%I

Ba

Ca

(%)

('%)

28.5 6.4

0.33 0.39

20.7 13.3

27.9 28.4

3.3 32.5

0.1 0.9

1.7 27.1 1.5 16.0 0.3 6.7

14.3 1.7

0.02 0.03 0.18 0.19 0.15

30.4 35.5 28.5 25.7 17.0

0.5 0.9 15.7 16.9 6.0

<0.01

1.5 39.1

( 7 ) 13.4 21.2 ( 2 ) 8.6 1 I . S 1 2 ) l2.S 37.3 (1) 3.6 32.4 (10) I 5 21.6 (2) 22.5

0.6 39.0 1.4 3.6 21.3 5.5

0.01 0.02

32.6

0.5

4.01
34.4 2.0 41.5 0.S 44.0

1.1 6.3 5.3

4.01 0.05 0.02

2.0
36.2

11.1

0.11

Au (ppm)

Ap

As Sb Co Se Ni Cd Mu Mn Sr (ppm) ippm) (ppm) (ppm) ippm) (ppm) (ppm) (ppnij (ppmi (ppm)

4 3

3 2

22

13

957 7S8

610 350

64 80 76

102 84 2.5

190 1 13 75 6 3

7 5 8 9 15

32 30

2 59

531 s3

201

48

4

-

61 122

6 10

75 13

98

33

19 4

S <2 c2 c2

42 76

0.2 0.3

5.17 4.12

199 145

561 605

232

c0.I <0.l 4 . 1 0.7 3.3

0.1s 0.52

18

24

432 7S4

1.40 1.33 0.83

393 212 92

10.4

0.48

13

4.18 1.13 2.29 3.2s 0.41

158

375

-

693 193

37 31

567 650

521 4300

Explorer Ridge

TAG Mound Black .;nioher White rmoher Pyrilc-chalcopyrite Marcasite-rphalcrite Sph3lt.ri(c.-niarcasite Silica-sphaleritr Snakepit Field Blsch smohcr 335 C chimney 227 C mound Pyl-ire-pyrrhorite Pyrite-mal-casite Sphalerite-tiiai-cabi~c

(4) (1)

(3) (1) (2) (2)

3.7 1.7 0.6

2.1

16.2 34.7

5.S

-

0.02 38.9 0.05 0.06 -

3Y.6 33.1 -

-

2.7 3.6 7.4 -

30.3 27.4 40.5

0.9 0.2 4.8

0.02

31.1

0.02

39.6 42.2

<().I 0.4 0.7

0.07

3.S 9.1 35.1

<0.01 ~0.01

d.01 4.01

c0.01

-

4 . 1 7.7 5.3 -

5.0

~0.1 4 . 1

~ 0 . 0 1
0.34 2.35 1.45

3s 35 151 25

19 69 52

0.39

10

0.92 4.62

<5 90

594

82

48 307

2S8 122 324 918

4

10

5 29

24 10 3 46

-

6

290

s5

15

21

I90

202

25

168 506

79 59 60

600 405 1847

463 1011 1417

17 1-148

IIS

52

38

I44 S2

128

-

73 650 151

17

-

260 18

77

-

220

35 35

I255

2277 1451

-

-

I08 8 51

224

34 44 45

34 216 173

11 23 33

77 232 542

482

I15 29

100

136 62 <2

31 48 64

74

<5

14 42 28

77 155

7 7 10

20 6

21Y

542

2 9

W W

a.

364

Chapter 2

TABLE 2.17 Base and precious metals in modern seafloor hydrothermal systems (Hannington et al., 1995). N" number of analyses Depth (m)

Dimension of largest deposit (approx.)

Cu (wt%)

Zn

Pb

Ag (ppm)

Mid-ocean ridges Southern Explorer Ridge Endeavour Ridge Axial Seamount Southern Juan de Fuca North Gorda Ridge 21~ East Pacific Rise 14~ Seamount 13~ Seamount 13~ East Pacific Rise 1 I~ East Pacific Rise 18~ to 26~ Galapagos Rift TAG Mound, MAR Snakepit, MAR

1800 2100 1500 2200 2700 2600 2500 2500 2600 2600 2600-2900 2700 3600 3400

250 m x 200 m up to 200 m small vent fields small vent fields small vent fields small vent fields small fields 800 m x 200 m small vent fields small vent fields small vent fields 100m x 100m 250 m x 200 m up to 100 m

3.6 3.0 0.8 1.4

Sedimented rifts Atlantis II Deep, Red Sea Middle Valley Guaymas Basin Escanaba Trough

2000 2400 2000 3200

90 million tonnes large fields small vent fields up to 250 m

0.5 0.4 0.2 1.0

2.0 3.4 0.9 11.9

<0.1 <0.1 0.4 2.0

Back-arcs Mariana Trough Valu Fa Ridge, Lau Basin Okinawa Trough Manus Basin North Fiji Basin Woodlark Basin

3600 1700 1400 2500 2600 2500

small vent fields 200 m x 100 m 300 m x 100 m up to 150 m small fields up to 200 m

1.2 4.2 3.7 1.0 8.6 <0.1

10.0 11.8 20.1 16.6 4.9 0.4

7.4 184 0.3 155 9.3 1900 0.6 51 <0.1 0.2 295

Au

(N)

1.0 6.1 0.1 132 <0.1 4.3 <0.1 188 2.6 23.2 <0.1 203 34.3 0.2 169 0.1 relict chimneys, 247~ sulfide vent 1.3 19.5 0.1 157 0.1 2.8 4.7 <0.1 48 0.5 massive pyritic sulfides and gossan 7.8 8.2 <0.1 49 0.4 1.9 28.0 <0.1 38 0.2 6.8 11.4 <0.1 121 0.5 4.1 2.1 <0.1 35 0.2 6.2 11.9 <0.1 78 2.2 2.0 6.3 <0.1 119 2.2 39 l0 78 187

0.5 <0.2 <0.2 < 10

(66) (31) (64) (11) (14) (5) (33) (11) (61) (73) (40) (16)

(39) (14) (7)

0.8 (11) 2.9 (44) 4.8 (9) (126) (24) 13 (34)

Solubility of Au is higher in higher fo2 and fs2 condition as shown in Fig. 2.44. The reasons for higher fo2 of back-arc fluids are that Fe3+/Fe 2+ of the country rocks in the back-arc basin area are higher than that of midoceanic ridge rocks and probably an injection of SO2 volcanic gas into back-arc fluids. This higher fo2 condition causes enrichment of Au. Higher concentrations of Au are interpreted by the following reaction: Au + 2H2S + 1/2 02 -- Au(HS) 2 + H + + H20

(2-32)

This reaction means that higher /H2S causes higher concentration of Au(HS) 2 as well as higher .[b2- Higher fH2S, /82 and fo2 of back-arc hydrothermal fluids can explain higher Au, Hg, As and Sb whose dominant dissolved species are probably thio complexes. Midoceanic ridge deposits are divided into volcanic-type and sedimentary-type (Gamo, 1995) or sediment-starved type or sediment-covered type (Scott, 1997). Metals concentrated to two types are distinct. In the sulfide deposits at Escanaba Trough,

Present-day Mineralization and Geothermal Systems

365

TABLE 2.18 Chemical composition of Kuroko ore and MORB (midoceanic ridge basalt) (logarithmic unit in wt%) (Shikazono, 1988)

Fe Zn Cu Pb SiO2 CaO Ag As Au Ba Bi Cd Co Ga Ge Hg Mn Mo Ni Pd Pt Rh Sb Se Sr T1 W

Kuroko

MORB

1.5 0.5 0.5 0 1.5 -0.5 to 0.5 -2.5 to - 2 - 1.5 to -0.5 -4 -0.5 to 0.5 -2.5 - 2 to - 1.5 -2.5 -2.5 -3.5 to -2.5 -4.5 to - 4 - 3 to - 2 -3.5 to -2.5 -2.5

0.5 to 1.5 0.5 to 1.5 - 0 . 5 to 0.5 - 2 to 0.5 - 1 to - 0 . 5 - 1.5 - 2 to - 3 - 3 to - 2 - 6 to - 4 . 5 - 3 to - 1 -4 - 2 . 5 to - 1 - 3 . 5 to - 1 - 4 to - 2 . 5 - 4 to - 2 - 5 to - 4 - 2 to - 1 - 3 . 5 to - 2 -3.5 < - 7 to - 6 . 5 < - 7 to - 6 . 5 < - 7 to - 6 . 5 - 3 . 5 to - 2 . 5 - 3.5 to - 2.5 - 4 to 0.5 - 4 to - 3 <3

- 2 to - 2 - 3.5 to - 3 - 0 . 5 to 0.5 - 3 . 5 to - 3 - 3 . 5 to -2.5

TABLE 2.19 Elements enriched into Kuroko and midoceanic ridge ores (Shikazono, 1988) Cu, Pb, Si, Ag, As, Au, Ba, Bi, Ga, Ge, Hg, Sb, Sr, T1 Fe, Cd, Mo, Ni, W, Zn, Mn, Co, Se, Pd, Pt, Rh

Elements enriched into Kuroko Elements enriched into MORB

s i g n i f i c a n t e n r i c h m e n t s in t r a c e m e t a l s s u c h as P b , Sn, A s , Sb, Bi, a n d S e a p p e a r to b e a c o n s e q u e n c e o f the interaction o f the h i g h e r t e m p e r a t u r e solutions with the s e d i m e n t s ( K o s k i et al., 1 9 8 8 ; Z i e r e n b e r g et al., 1 9 9 3 ) . T h e h i g h e r P b c o n t e n t s r e f l e c t t h e d e s t r u c t i o n o f f e l d s p a r s f r o m c o n t i n e n t a l l y - d e r i v e d t u r b i d i t e s . T h i s p r o c e s s is s u p p o r t e d b y P b i s o t o p e s t u d i e s ( L e H u r a y et al., 1 9 8 8 ; H a n n i n g t o n et al., 1 9 9 5 ) . Sediment-covered

deposits

(Guaymas)

contain

higher

Ba,

Pb

and

carbonates

( H a n n i n g t o n et al., 1 9 9 5 ) . B a c k - a r c d e p o s i t s a r e g e n e r a l l y s i m i l a r to s e d i m e n t - c o v e r e d d e p o s i t s at m i d o c e a n i c r i d g e . F o r e x a m p l e , P b , B a a n d A s a r e e n r i c h e d in t h e s e d e p o s i t s .

366

Chapter 2

sol

1

HM

I'~I

e4

oo

32

34

Ckl

o

(.9 O ._1 i

MT

\

-loQ

36

-11 38

12

t

,0.01

, 0.1 '~ 1 I

I

\i

II

x

PY

PO

\ I

,, ',

;

;

42

,

',I;

9

44

4

5

6

1

I

'l

3

i

i -[

LOG aS 2 \ I

40

i

~ H2S ,~HS\ , / ! ~1

iw

7

8

9

pH Fig. 2.44. logao2-pH ranges for Kuroko ore fluids and midoceanic ridge hydrothermai solution, l" Kuroko; 2: Axial Explorer; 3" 21~ Southern Juan de Fuca; 4: 21~ Endeavour; 5" Guaymas. Temperature = 250~ ~Sr (total reduced sulfur concentration) - 6.6 x 10 -3 m. HM: hematite, MT: magnetite, PY: pyrite, PO: pyrrhotite. Dotted line: Au solubility (ppm)(Shikazono, 1988).

2.4.3. M e c h a n i s m of f o r m a t i o n of c h i m n e y and ore deposits

The above argument clearly indicates that the back-arc deposits have been formed by a mechanism very similar to the Kuroko deposits. However, the form of the ore body for two deposit types (back-arc deposits and Kuroko deposits) is distinct: back-arc deposits occur as chimney and mound, while Kuroko deposits are mostly strata-bound and massive and chimney-like structure is uncommon. Farrell and Holland (1983) cited based on Sr isotope study on anhydrite and barite in Kuroko deposits that the most appealing model for the formation of Kuroko strata-bound ores would seem to entail precipitation of the minerals from a hydrothermal solution within the discharge vent or in the interior of a hydrothermal plume formed immediately below above the vent exit in the overlying seawater (Eldridge et al., 1983). The study on the chimney ores from Kuroko deposits support this model which is discussed below.

367

Present-day Mineralization and Geothermal Systems ANHYDRITE BARITE AMORPHOUSS l L I C A ~ DENDRITICSPHALERITECOLLOFORMSPHALERITE

]

I

MARCASITE PYRITE WURTZITE ~

]

CHALCOPYRITE PYRRHOTITE ISOCUBANITE < 100

150

250

350

350

< 100

Temperature~

Fig. 2.45. Hypothetical mineral paragenesis for a sulfide-sulfate-silica chimney. Approximate temperature ranges for different minerals are based on direct measurements of vent fluid temperatures and analyses of fluid inclusions in real chimneys (Hannington et al., 1995). Recently, chimney-like ores have been described from Kuroko deposits (Matsukuma, 1989; Shimazaki and Horikoshi, 1990; Shikazono and Kusakabe, 1999). The formation mechanism of chimney from the hydrothermal ore deposits at midoceanic ridges was clarified. Thus, these studies constrain the formation mechanism of Kuroko ore deposits. The characteristic features of chimney-like ores and chimney from Kuroko deposits and Mariana Trough are summarized and mineralogical and geochemical characteristics of Kuroko and Mariana chimneys and those of the midoceanic ridge chimneys are compared below.

2.4.3.1. Zonation and sequence of mineral precipitation Mineralogical zonations of the midoceanic ridge chimneys have been described by many investigators (e.g., Haymon and Kastner, 1981; Goldfarb et al., 1983; Haymon, 1983; Tivey and Delaney, 1986; Graham et al., 1988; Peter and Scott, 1988; Hannington and Scott, 1988; Paradis et al., 1988) (Fig. 2.45). For example, the studies on the chimneys from the Juan de Fuca Ridge showed that anhydrite, barite and amorphous silica precipitated at an early stage of hydrothermal activity and formed the outer zone of the chimney. These minerals were replaced successively by zinc-iron and copper-iron sulfides (Paradis et al., 1988). The core is composed of chalcopyrite or isocubanite which formed from hot hydrothermal solution at a late-stage of hydrothermal activity. This Cu-rich zone is common also in East Pacific Rise chimneys (e.g., Goldfarb et al., 1983). However, no chalcopyrite zone occurring in inner parts of Kuroko and Mariana chimneys has been found (Shikazono and Kusakabe, 1999). Graham et al. (1988) established 8 facies from the exterior to interior for the midoceanic ridge chimneys: 1, anhydrite-rich; 2, marcasite + wurtzite; 3, pyrite; 4, bornite-rich; 5, chalcopyrite; 6, pyrite; 7, marcasite; 8, wurtzite. This study also indicated

368

Chapter 2

that in the midoceanic ridge chimneys sulfates (anhydrite and barite) precipitated earlier than sulfides. Amorphous silica precipitated after anhydrite but before sulfides. This sequence of mineral precipitation is roughly consistent with that of Kuroko and Mariana chimneys (Shikazono and Kusakabe, 1999). 2.4.3.2. Mineral composition Galena is abundant in the inner parts of Kuroko chimneys. Galena occurs in association with sphalerite in Kuroko and Mariana chimneys. On the other hand, galena is rarer in the midoceanic ridge chimneys, though in the caldera of Axial Seamount, Juan de Fuca Ridge, Pb-minerals such as galena and jordanite precipitated at a latestage (Hannington and Scott, 1988). Barite is abundant or rare in Kuroko and Mariana chimneys, whereas it is common but not abundant in midoceanic ridge chimneys. Cuminerals (e.g., isocubanite), pyrrhotite, wurtzite and marcasite are absent in Kuroko and Mariana chimneys, whereas they are found in some midoceanic ridge chimneys. A little tetrahedrite-tennantite is found in Kuroko and Mariana chimneys but is not common in the midoceanic ridge chimneys. Secondary sulfide minerals (e.g., bornite, chalcocite and covellite) are not found in Kuroko and Mariana chimneys but they are present in some midoceanic ridge chimneys. 2.4.3.3. Ore texture Replacement texture is common in midoceanic ridge chimneys. For example, sulfides (marcasite, wurtzite) replace anhydrite and barite (EPR 21~ l l~ 13~ (Haymon, 1983; Graham et al., 1988). This texture is also common in Kuroko and Mariana chimneys. Dendritic texture is common in the outer parts of midoceanic ridge chimneys. However, this is not observed in Kuroko and Mariana chimneys, though dendritic-like texture of sphalerite, chalcopyrite and galena is found in Mariana chimney and Kuroko ores. Granular and framboidal textures of sulfides are common in Kuroko and Mariana chimneys as well as in midoceanic ridge chimneys. Chalcopyrite disease in sphalerite which was first reported for Kuroko ores from the Furutobe mine by Barton (1978) is relatively common in midoceanic ridge chimneys, as reported from EPR 21~ (Goldfarb et al., 1983) and from Escanaba Trough, Gorda Ridge (Koski et al., 1988), but not observed in Mariana chimneys, but observed in Kuroko chimneys. Watermelon texture (Barton and Bethke, 1987) was reported from Escanaba Trough, Gorda Ridge (Koski et al., 1988), but not from Kuroko and Mariana chimneys. Overgrowth of sulfides on barite is common in Mariana chimney, and sometimes observed in Kuroko chimneys, but not commonly reported in many midoceanic ridge chimneys. 2.4.3.4. Grain size A few studies on grain size of minerals in midoceanic ridge chimneys have been published. Feely et al. (1987) described the grain size of mineral particles in the smoke and sediment samples from southern Juan de Fuca Ridge. They report the following grain sizes: sphalerite: 0.3-100 gm in diameter (usually less than 20 txm); pyrite: 0.1-10 gm; Fe-Si, Ca-Si phases: 5-150 gm. Converse et al. (1984) reported the grain size

Present-day Mineralization and Geothermal Systems

369

of pyrrhotite in the hot springs at EPR 21~ to be less than 5-10 Ixm. According to microscopic studies on sulfides at EPR 11~ and 13~ by Graham et al. (1988), grain size of sulfides is as follows: euhedral cubic and octahedral pyrite crystals: 5-50 txm across (usually less than l0 txm); subhedral to euhedral pyrite crystal occurring in the inner parts of the chimney: 100-500 txm; marcasite: 5-500 Ixm. Previously reported grain sizes of sulfates are very few: grain size of anhydrite from Guaymas basin is 0.2-0.5 mm (Peter and Scott, 1988). The previous studies suggest that the grain size of sulfides is smaller than anhydrite. This difference in grain size is consistent with that found in Kuroko ores and chimneys. Very fine-grained sulfides (n x 10 -1 Ixm), which are common in Kuroko ores (Shikazono, unpublished), have not been reported from midoceanic ridge chimneys. However, SEM (scanning electron microscope) observations of Kuroko and Mariana chimneys indicate that the minerals are aggregates of very fine-grained crystals. Therefore, SEM observation is necessary to measure grain size of individual mineral crystals. However, data from SEM observations of midoceanic ridge chimneys are scarce. 2.4.3.5. Sulfur isotope data Sulfur isotopic studies on midoceanic ridge chimneys show a wide variation in 834S of sulfides from the exterior to the interior of the chimney. The 834S values of sulfides in outer parts are generally higher than that of sulfides from inner parts (Shanks and Seyfried, 1987; Woodruff and Shanks, 1988). This variation trend is also observed in Kuroko chimneys and could be explained by the different degree of reduction of seawater sulfate by Fe-minerals (pyrite, pyrrhotite), or by H2S of hydrothermal solution (Shanks and Seyfried, 1987). There are other explanations for 3345 variation in the chimneys. Graham et al. (1988) thought that ~34S of vent fluid of early stage of hydrothermal activity was low, while that of late stage increased. This was caused by the change in contribution of H2S derived from reduction of seawater sulfate compared to H2S from basalt by leaching of basalt at deeper parts of the plumbing system or by rapid chemical reactions of dissolution, reprecipitation, and replacement between hydrothermal fluids and earlier sulfide minerals in the chimney. Local seawater sulfate reduction in chimney conduit is also possible. However, this mechanism seems unlikely because a rate of sulfate reduction is too slow at the site of chimney formation (Graham et al., 1988).

2.4.3.6. Mineral particle behavior in hydrothermal plumes It is worth elucidating mineral particle behavior in hydrothermal plumes in order to consider the formation mechanism of chimney and massive ores on the seafloor. Using the grain size data on sulfides and sulfates, the density of the fluids and of the minerals, the relationship between vertical settling rate and grain size of sulfides and sulfates can be derived based on the following Stokes equation: Vs -- 2rZ g(ps - p) o~/9#

(2-33)

where Vs = settling velocity of fine particles, r = grain size of particle, g = acceleration due to the force of gravity, ot = shape factor, Ps = density of fine particles, p = density of fluid, and/z = viscosity of fluid.

370

Chapter 2

We can estimate minimum grain size of minerals settling onto the seafloor, if flow rate of the ascending solution is known. Although flow rate of hydrothermal solutions responsible for the formation of Kuroko deposits cannot be determined directly, we can get information on the flow rate of hydrothermal solutions issuing on midoceanic ridges. Estimated or measured flow rates are 1-5 m / s at National Geographic vent area (Macdonald et al., 1980), 2-10 liters/s at Galapagos geothermal area (Corliss et al., 1979; Turekian et al., 1981), and 0.7-2.4 m / s at EPR 21~ (Converse et al., 1984). Minimum grain size of the fine particles estimated using equation (2-33) which can settle from the hydrothermal plume is estimated about 100 Ixm. However, this estimate is not consistent with the observed grain size (n x 10 -1 btm) of sulfides overgrown on barite from Kuroko deposits (Shikazono, unpublished). Sulfides with very fine grain size observed in Kuroko and Mariana chimneys are expected to not settle onto the seafloor based on the calculations of settling velocity of sulfides and sulfates by Converse et al. (1984), Feely et al. (1987), and Shikazono and Kusakabe (1999). Converse et al. (1984) calculated that less than 3% of pyrrhotite particles with 30 t,m in diameter that are entrained in the hydrothermal plumes at EPR 21 ~ settle from the plume before the dispersal by a lateral submarine current at a height of ca. 250 m above the seafloor. According to the calculation by Feely et al. (1987) the majority of large-grained black smoker particles should be deposited within a few hundred meters of the vent and the very fine sulfide particles (30 Ixm) can be expected to be dispersed farther than a few kilometers from the vent. These studies are consistent with the calculations but not with the observation of grain size of sulfides by Shikazono and Kusakabe (1999). The disagreement between the calculation on the minimum grain size of mineral particles settling onto the ocean floor near the hydrothermal vent and observation of grain size of sulfides found in Kuroko and Mariana chimneys suggest that fine grained sulfides did not precipitate homogeneously and directly from aqueous solutions, but heterogeneously nucleated and grew on the large grained barite. If this process takes place efficiently, lots of sulfides precipitate from hydrothermal solution, leading to a formation of large hydrothermal ore deposits on the seafloor.

2.4.3.7. Model for the formation of sulfate-sulfide chimneys and massive deposits on the seafloor The above summarized mineralogical and geochemical studies on Kuroko and Mariana chimneys (Shikazono and Kusakabe, 1999), and previous studies on midoceanic ridge chimneys, combined with the studies of mineral particle behaviors in the plumbing system, are used to develop the following plausible model for the growth history of sulfate-sulfide chimneys on the seafloor (Shikazono and Kusakabe, 1999). Sulfates (barite and anhydrite) precipitate due to the mixing of discharging hydrothermal solution with cold seawater above the seafloor at an early stage of hydrothermal activity. Ca and Ba in hydrothermal solution react with SO42- in cold seawater, leading to the precipitations of anhydrite and barite. It is observed that anhydrite precipitated earlier than barite. This may depend on the initial Ca and Ba concentrations of end member hydrothermal solutions, temperature and degree of mixing of hydrothermal solutions and

Present-day Mineralization and Geothermal Systems

371

ambient seawater. This precipitation mechanism for sulfates can explain 87Sr/86Sr ratios of anhydrite from the midoceanic ridges, Mariana chimney and Kuroko ores (Albarbde et al., 1981; Farrell and Holland, 1983; Kusakabe et al., 1990) and 834S values of sulfates and sulfides from these areas (e.g., Kawahata and Shikazono, 1988), but a simple heating of seawater cannot explain these isotopic data. Precipitation mechanisms of anhydrite and barite for midoceanic ridge chimneys have been discussed by Janecky and Shanks (1988). They indicated that anhydrite precipitates at an earlier stage than sulfides and barite. Calculation of S.I. (saturation index of minerals) as a function of temperature during the mixing of hydrothermal solution with seawater in Kuroko-forming system indicates that S.I. for anhydrite decreases with decreasing temperature, whereas it increases with decreasing temperature for barite, suggesting that anhydrite precipitates at higher temperatures than barite (Ohmoto et al., 1983). This calculation is consistent with the results for the midoceanic ridge system computed by and Janecky and Shanks (1988). As noted already, the grain size of anhydrite and barite is relatively larger than sulfides. Large-grained sulfate particles can settle onto the seafloor, forming sulfate mounds and/or chimneys on the seafloor. Fine grained sulfide particles precipitate due to the rapid mixing of hydrothermal solution with cold seawater at a low degree of mixing of hydrothermal solution with seawater (Janecky and Shanks, 1988). At early stage the buoyancy of very fine grained sulfide particles takes place and they do not settle near the hydrothermal vent.

2.4.4. Hydrothermal alteration As already noted, intense bimodal volcanic activity occurred in the Kuroko mine area at middle Miocene age and dacitic and basaltic rocks suffered hydrothermal alteration. The midoceanic ridges basalt (MORB) is widespread and sometimes hydrothermally altered. Shikazono et al. (1995) compared hydrothermally altered basalt from the Kuroko mine area and MORB and clarified the differences in the characteristics of these basaltic rocks. The differences between the Kuroko basalt alteration and MORB alteration are: (1) late-stage Fe-rich epidote and hematite filling vesicles are common in the basalt studied but not in MORB, (2) the CaO concentration of bulk rock is relatively lower and the Fe203 and Na20 concentrations are relatively higher in the Kuroko basalt, (3) the Na20 concentration of the basalt studied correlates inversely with their MgO concentrations, and (4) the MgO/FeO ratios of chlorite and actinolite and the Fe203 concentrations of epidote in the Kuroko basalt are higher. The original rock composition seems to be an important cause for these differences. Higher FezO3/FeO and MgO/FeO ratios of the original Kuroko basalt led to higher MgO/FeO ratios in chlorite and actinolite than in MORB minerals (Fig. 2.46). High Fe203 concentration in epidote and the presence of hematite, high Na20 and low CaO concentrations in the Kuroko basalt can all be explained as being inherited from the fresh Kuroko basalt (Table 2.20). The vesicle volume of the Kuroko basalt is large (average 20%) and vesicles are filled with hydrothermal minerals (epidote, calcite, chlorite, pyrite, quartz) which formed

372

Chapter 2 15

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Fig. 2.46. Frequency histogram for MgO/FeO ratios (in wt%) of chlorite from the Kuroko basalt (A) and MORB (B). Data sources are: Shikazono et al. (1995), Humphris and Thompson (1978) (M: Mid-Atlantic Ridge) and Kawahata (1984) (C: Costa Rica Rift, Galapagos Spreading Centre). The data on chlorite from MORB are taken from typical metabasait and not from quartz-chlorite breccia and veins which formed in a hydrothermal upflow zone(Shikazono et al., 1995).

during the late stage. The abundant vesicles provided high permeability and thus also strongly controlled the bulk rock compositions, the stable isotope composition, and the alteration mineralogy. Superimposed alteration of chlorite- and epidote-types may be caused by high permeability of basalt owing to abundant vesicles. Such vesicular basalt (vesicle volume 2 0 - 4 0 % has been reported in several back-arc basins (e.g., Lau Basin, Sumisu Rift, Fiji Basin). In contrast, the vesicle volume in MORB is less than 10% (Humphris and Thompson, 1978), suggesting that vesicles are not an important factor in MORB alteration. Although several differences are observed in both basalts from Kuroko mine area and midoceanic ridges, the hydrothermal alteration mineral assemblages as a function of

373

Present-day Mineral&ation and Geothermal @stems TABLE 2.20

Characteristics of the hydrothermally altered basalt from Kuroko mine area and MORB (Shikazono et al., 1995) Kuroko basalt

Midocean ridge basalt

Absent Abundant (20 vol.%)

Present Few (less than 10 vol.%)

MgO/FeO: 1.1-1.6 Fe203 content: 11-16 wt% MgO/FeO: 1.3-2.0 Common

MgO/FeO: 0.5-1.7 Fe203 content: 5-15 wt% MgO/FeO: 0.7-1.2 Rare

Positive correlation with MgO and H20 Negative correlation with MgO Original composition (wt%) of basalt

FeO, ~Fe CaO, SiO2, Na20 CaO: 4.6; Na20: 3.6; MgO: 6; FeO: 5

FeO CaO, SiO2 CaO: 9-13; Na20: 2-3; MgO: 7-8; FeO: 7-9

Stable isotopes:

313C of calcite: - 8 to -3%0 3180 of calcite: +5 to +12%o ~D of epidote: about 40%0

313C of calcite: - 7 to -3%o 3180 of calcite: +7 to +12%o

Temperature:

Chlorite alteration event: 230-250~ Epidote alteration event: 250-280~

Epidote-actinolite, albite-quartz, chlorite assemblage: 250-350~

Seawater/basalt ratio:

Chlorite: high (up to 40)

Chlorite-quartz-rich rock: high (more than 50) Epidote-actinolite, albite- quartz, chlorite assemblage: low (1.6)

Petrography Chlorite-quartz-rich rock Vesicles

Mineralogy Chlorite Epidote Actinolite Hematite

Bulk composition

Calcite: low (0.1-1.8)

s e a w a t e r / r o c k ratios are generally consistent with each other. For instance, epidote appears in low w a t e r / r o c k ratio while chlorite in high w a t e r / r o c k ratio (Mottl, 1983) (Fig. 2.47).

2.5. Besshi-type deposits in comparison with Kuroko deposits and midoceanic ridge deposits The characteristic features of Besshi-type deposits are similar to those of m i d o c e a n i c ridge deposits and back-arc deposits. The c o m p a r i s o n a m o n g these features is given below.

2.5.1. General features and classification P a l e o z o i c - M e s o z o i c v o l c a n o g e n i c stratiform Cu deposits in Japan which are generally m e t a m o r p h o s e d have been called Besshi-type deposits (Kato, 1937) or b e d d e d cupriferous iron sulfide deposits ( K a n e h i r a and Tatsumi, 1970).

374

Chapter 2 100

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Fig. 2.47. Model predicting mineral assemblages and proportions produced when basalt reacts with seawater in different water/rock mass ratios. The model is based on experimental data but is close to actual observed assemblages in recovered greenschist facies metabasalts (Mottl, 1983). General descriptions of Besshi-type deposits have been given by Kanehira and Tatsumi (1970), Franklin et al. ( 1981 ) and Fox (1984). Typical Besshi-type deposits occur in Sanbagawa metamorphic Belt in Japan. But the deposits considered to belong to the Besshi-type occur widely in the world. Fox (1984) thought that the Besshi-type deposits include, in addition to the deposits of the Sanbagawa Belt, the deposits of the Trondheim region in Norway (Killingdal, Tverfjell, etc.), the Matchless belt in Namibia (Otjihase, Matchless, etc.), Blue Ridge of eastern USA (Ore Knob, Cherokee, etc.), the Outokumpu region of Finland and the Goldstream area of Canada. The other deposits such as the Kieslager deposits of the eastern Alps (Tarkian and Garbe, 1988), the Shimokawa deposits in Hokkaido (Mariko, 1984; Mariko and Kato, 1994), cupriferous pyritic lenticular deposits in Abukuma, Shimanto and Maizuru in Japan (Kanehira and Tatsumi, 1970) and the Kiliembe deposit of Uganda (Warden, 1985; Maiden, 1993). The Besshi-type deposits in Japan are characterized by the following features (Kanehira and Tatsumi, 1970): (1) They are composed largely of massive ores consisting of pyrite with some amount of chalcopyrite. In some deposits a considerable amount of pyrrhotite occurs. (2) They are generally bed-like or lenticular in form, and lie conformably in crystalline schists. (3) Most are in association with the products of basic submarine volcanic activities or their metamorphosed equivalents. (4) The occurrence of the deposits in a limited area is usually confined to a definite stratigraphic horizon. Besshi-type deposits in Japan are divided into Besshi-subtype and Hitachi-subtype based on geological, mineralogical and geochemical characteristics (Sato and Kase,

375

Present-day Mineralization and Geothermal Systems

1996). Sato and Kase (1996) summarized these characteristic features and divided Besshisubtype into Group A (sediment-barren type) and Group B (sediment-covered type) (Table 2.21). These characteristic features of these Besshi-type deposits in Japan, mainly focusing on the Besshi-subtype and comparing of these features with those of Kuroko and midoceanic ridge deposits are described below.

2.5.2. Geological characteristics Sato and Kase (1996) summarized geological characteristics and inferred tectonic settings of the Besshi-type deposits as shown in Table 2.21. These geological characteristics are summarized below.

2.5.2.1. Distribution More than 100 Besshi-subtype deposits occur in the Sanbagawa Belt in the southwest Japan (Fig. 2.48). Representative large deposits are the Besshi, Sazare, Shirataki, Okuki in Shikoku, Iimori in Kii and Kune and Minenosawa deposits in central Honsyu. Most are Group A deposits (e.g., Besshi, Sazare, Shirataki) but some are Group B deposits (e.g., Kune, Minenosawa). A few deposits are distributed in the Inner Zone of southwest Japan (e.g., Tsuchikura), in the Shimanto terrain (e.g., Makimine) and in the Hidaka Belt in Hokkaido (e.g., Shimokawa) (Fig. 2.49). Hitachi-subtype deposits occur in the Abukuma (Hitachi) and in the Maizuru zone (e.g., Yanahara) and in Kitakami (e.g., Taro) (Fig. 2.49).

2.5.2.2. Age of formation of ore deposits Kojima et al. (1956) found that most of the Besshi-subtype deposits in the Sanbagawa metamorphic terrain of Shikoku occur in the Minawa Formation characterized by the basic schists, metamorphic equivalents of basaltic lavas and hyaloclastics, based on the summary of the stratigraphic horizon of about one hundred Besshi-subtype deposits. TABLE 2.21 Geological characteristics and inferred tectonic settings of Paleozoic to Mesozoic volcanogenic Cu sulfide Besshi-type deposits of Japan (Sato and Kase, 1996) Subtype

Hitachi

Volcanism Age

Bi-modal(rhyolite, basalt) Carboniferous to Permian

Host rocks

Group A Rhyolite,basalt, shale, sandstone Basalt, chert, shale

Tectonic setting Back-arcrift or continental rift Example

Hitachi, Yanahara

Besshi Uni-modal (basalt) Carboniferous to Cretaceous Group B Basalt, shale (chert absent or minor) Sediment-barrenmidoceanic Sediment-covered ridge or hot spot midoceanic ridge Iimori, Besshi, Sazare, Shimokawa, Makimine, Shirataki, Okuki, Tsuchikura Minenosawa,Kune (Nago)

376

Chapter 2 N

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f

i Fig. 2.48. Location of the representative bedded cupriferous iron sulfide deposits (Besshi-type) and some other deposits. 1: Shimokawa, 2: Taro, 3: Hitachi (Fujimi and Fudotaki deposits), 4: Kune (including Nago and Ohniwa deposits), 5: Minenosawa (including Oi deposit), 6: Tsuchikura, 7: Iimori, 8: Yanahara, 9: Kotsu, 10: Shingu, 11: Hino-oku, 12: Sazare, 13: Shirataki, 14: Ikadazu (including Yokei and Sekizen deposits), 15: Besshi, 16: Chihara, 17: Okuki, 18: Makimine, 19: Iwami (Miocene Kuroko deposit) (Kase and Yamamoto, 1988).

Isozaki and Itaya (1990) suggested from biostratigraphic data that the age of Minawa Formation is Triassic to Jurassic, indicating Besshi-subtype (mostly Group A) formed at these ages. Age of formation of Group B of Besshi-subtype are variable, early Cretaceous (Shimokawa), late Early Cretaceous to early Late Cretaceous (Makimine) and Triassic to Jurassic (Minenosawa, Kune). Age of formation of Hitachi ~.ubtype is Middle Permian (Yanahara), and carboniferous-Permian (Hitachi).

2.5.2.3. Host rocks and tectonics Besshi-type deposits in Sanbagawa metamorphic terrain occur in the Minawa Formation which is composed of basic schist. Sometimes, they are associated with quartz schists. Probably, quartz has been originally formed from hydrothermal solution like siliceous ore in Kuroko deposits. Original rocks of basic schists are basaltic lava and hyaloclastics. Detailed geochemical investigation on the basic schists in the Sanbagawa

377

Present-day Mineralization and Geothermal Systems !

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Fig. 2.49. Map of pre-Neogene tectonic division of northern Japan and locations of the Shimokawa and the Hitachi mines. 1: Abukuma Belt, 2: South Kitakami Belt, 2h: Hayachine Structural Belt, 3: North Kitakami Belt, 4: Iwaizumi Belt, 5: Taro Belt, 6: Ishikari Belt, 7: Kamuikotan Belt, 8: Hidaka Belt (a: Hidaka Western Greenstone Belt, b: Greenstone-Chert-Turbidite Facies Belt, c: Hidaka Metamorphic Belt, d: Turbidite Facies Belt), 9: Tokoro Belt, 10: Nemuro Belt (Mariko and Kato, 1994).

has not been carried out. However, recent analyses of the basic schists near the ore deposits in Shikoku indicate MORB-like chemistry (Isozaki, 1995, cited in Sato and Kase, 1996). Detailed investigation of the Shimokawa mine area in the Hidaka, Hokkaido (Mariko, 1984; Mariko and Kato, 1994) has been carried out during the last two decades. The Shimokawa deposits are associated with pillow basalt, diabase sheets and black slate. The basaltic rocks are of MORB-like compositions, suggesting the deposits formed at midoceanic environment. Mariko (1984) suggested that the tectonic setting of the Shimokawa deposits is an ancient analogue of the Guaymas Basin. However, Miyake (1988) thought that the Shimokawa ore-forming environment was different from the Guaymas Basin, considering an absence of high temperature hydrothermal alteration in the sedimentary rocks in Hidaka Belt which is typical accretion complexes on the Cretaceous convergent margin and is thought to represent the oceanic trench. Mariko and Kato (1994) analyzed 171 samples and 43 samples from the Shimokawa area for major elements and minor elements, respectively, and indicated that the

378

Chapter 2

tectonic setting of the Shimokawa basic rocks and massive sulfide deposits belongs to pull-apart basin as the Gulf of California and the Andaman Sea among the marginal basin. Hitachi deposits, representative deposits of the Hitachi subtype, are distributed in southern Abukuma region, northern Honshu, consisting of metamorphosed volcanicsedimentary sequence of Late Paleozoic age. The volcanic rocks are bimodal type (felsic and basaltic rocks) and tholeiitic and alkaline (Tagiri and Okura, 1979). The basic rocks in these regions as well as in the Kuroko mine area (e.g., Takara in central Honshu, Furutobe, and Ainai in Hokuroku) (Shimazu, 1971) suffered metamorphism of prehnite-pumpellyite facies, but this facies is very rare in the present-day midoceanic ridge. Mariko and Kato (1994) thought mainly from major and minor elements geochemistry of the host rocks that submarine volcanism took place in island arc and back-arc environments. Shinozuka et al. (1999) analyzed the host volcanic and intrusive rocks in the Minamidani mine district in the Maizuru tectonic Belt and found that these rocks formed in an island arc back-arc system near Laurasia during late Paleozoic. Probably the Yanahara deposits, one of the representative Hitachi subtype deposits, were formed in an island arc back-arc system as same as the Minamidani. Sato and Kase (1996) thought that the Hitachi-subtype deposits formed in back-arc rift or continental rift (Table 2.21).

2.5.3. Metamorphism and hydrothermal alteration The host rocks for the Besshi-type deposits are generally suffered regional, hydrothermal and contact metamorphims. The Sanbagawa basic and pellitic rocks suffered high P - T metamorphism, Sanbagawa metamorphism, that is green schist facies metamorphism (Miyashiro, 1967). It is thought that submarine hydrothermal alteration of basalt took place in the Besshi-subtype mineralization area in Sanbagawa. However, no study on the hydrothermal alteration of basalt (basic schist) has been carried out and it is clearly necessary. The contact metamorphosed rocks and sulfide ores caused by Tertiary granitic intrusion are found in deeper part of the Besshi mine and Makimine mine (Tatsumi, 1952; Kase, 1972). The basic rocks in the Shimokawa area are metamorphosed up to green schist to amphibolite facies which are considered to be due to submarine hydrothermal activity (Miyake, 1988; Mariko, 1984; Mariko and Kato, 1994). Host rocks in the Hitachi area suffered regional metamorphism, contact metamorphism by Cretaceous granitic rocks and hydrothermal alteration associated with sulfide mineralization. Host rocks in deeper parts of the Besshi mine suffered contact metamorphism probably due to Cretaceous granitic intrusion. Kase (1972, 1977, 1988) suggested the contact metamorphism in the Besshi mine from the occurrence of large amount of pyrrhotite and As, Zn, Ag and Sn vein mineralization in deeper levels. In the other areas, generally, the host rocks weakly suffered regional and contact metamorphisms but suffered ocean-floor hydrothermal alteration. For example, hydrothermal alteration mineral assemblages in the Minamidani mine district in the Maizuru range from prehnite-pumpellyite facies to a transition state from green schist to amphibolite

Present-day Mineralization and Geothermal Systems

379

facies (Shinozuka et al., 1999). Shinozuka et al. (1999) suggested that the ocean-floor alteration in island arc back system such as the Minamidani district is characterized by the lack of zeolite facies.

2.5.4. Mineralogical characteristics 2.5.4.1. Opaque and gangue minerals Mineralogical characteristics of Besshi-subtype deposits (e.g., Besshi deposit) were described by Kase (1972, 1977, 1986) and Kase and Yamamoto (1988). Their studies are described below. Main opaque minerals are chalcopyrite, pyrite, pyrrhotite, sphalerite and bornite (Table 2.22). These minerals commonly occur in massive, banded and disseminated ores and are usually metamorphosed. Hematite occurs in "red chert" which is composed of fine grained hematite and aluminosilicates (chlorite, stilpnomelane, amphibole, quartz) and carbonates. The massive sulfide ore bodies are overlain by a thin layer of red ferruginous rock in the Okuki (Watanabe et al., 1970). Minor opaque minerals are cobalt minerals (cobaltite, cobalt pentlandite, cobalt mackinawite, carrollite), tetrahedrite-tennantite, native gold, native silver, chalcocite, acanthite, hessite, silver-rich electrum, cubanite, "valleriite", and mawsonite or stannoidite (Table 2.22).

TABLE 2.22 List of metallic minerals from the Besshi-type deposits (Kase and Yamamoto, 1988) Elements and alloy Native gold Native silver Electrum Native bismuth Sulfates Barite Anhydrite Gypsum Oxides Hematite Magnetite Ilmenite Rutile

Sulfides and sulfosalts Pyrite Marcasite Pyrrhotite Troilite Mackinawite Cobaltite Carrollite Cobalt pentlandite Chalcopyrite Bornite Cubanite Idaite Chalcocite Covellite Sphalerite Galena Tetrahedrite Acanthite Stromeyerite Molybdenite Stannoidite Hessite Colusite

380

Chapter 2

Dominant gangue minerals are quartz, muscovite, chlorite, actinolite, hornblende, epidote, and biotite (Table 2.22). Minor minerals are rutile, illite, sphene, and glaucophane. It is interesting to note that silicate minerals such as chlorite, epidote, pumpellyite, and albite are common and actinolite has been reported from the basalt near the Ainai Kuroko deposits (Shikazono et al., 1995) and they are also common in the basic schist which host the Motoyama Kuno deposits (Yui, 1983). A pyrite-chalcopyrite-sphalerite(-bornite) assemblage is ubiquitous in most of the Besshi-subtype deposits. Pyrite of the Besshi-type deposits is characterized by high Co content ranging from n x 10 - 2 to n x 10 -1 wt% (Itoh and Kanehira, 1967), which is quite a bit higher than that of Kuroko deposits. Chemical compositions of sphalerite from the Besshi-Honko, Besshi-Ikadatsu, Sazare, Hitachi, and Shimokawa have been obtained by EPMA analyses (Mizuta, 1988; Mariko, 1988). FeS mol% of sphalerite from these deposits is 4.0-10.3, 4.7, 5.1, 2.23.9, 8.7-18.1 and 14.3-19.8 for Besshi-Honko, Besshi-Ikadatsu, Sazare, Hitachi, and Shimokawa, respectively. This variation range is simi|ar to that of midoceanic ridge deposits, although usually the FeS contents of sphalerite from the midoceanic ridge deposits are higher (Shikazono, 1988). For example, sphalerite and wurtzite coexisting with hexagonal pyrrhotite from Guaymas Basin has a widely variable FeS content of 25.9-54.8 mol%, implying either fluctuation of i s 2 (10-12-10-23), at the vent temperature, or remnant disequilibrium (Peter and Scott, 1988). Kase and Horiuchi (1996) obtained a large number of analytical data on sphalerites from sixteen Besshi-type deposits, mainly at Besshi and its vicinity, Hitachi, and Shimokawa. They revealed that: (1) the Mn/Zn and Co/Zn ratios of sphalerite may have markedly increased during contact metamorphism, while the Cd/Zn ratios remained unchanged; (2) the Emco/~mzn (Em: total dissolved concentration in ore fluids) and Zmco/Zmzn ratios in the initial ore solutions responsible for the mineralizations at Besshi which was calculated based on the equilibrium fractionation model between hydrothermal solution and sphalerite and analytical data on sphalerites are quite similar to the ratios of hydrothermal solutions at EPR 21 ~ (3) however, these ratios for the Hitachi solutions are very low and different from those of the Besshi-subtype solution. Some sphalerites from the Shimokawa and Sazare deposits contain high amounts of Co up to 0.9% (Mariko, 1988 Kase, 1988), suggesting high 52mco/~mzn in the Shimokawa and Sazare solutions. No compositional zoning in sphalerite grain is observed, suggesting original zoning has been homogenized by thermal metamorphism (Mizuta, 1988). In general electrum is rare, but it is found in some deposits such as the Shimokawa, Besshi, and Yanahara (Urashima, 1974). A detailed study on the mode of occurrence of electrum has been carried out only on the Shimokawa deposit (Maeda et al., 1981). Electrum occurs in compact massive or banded ores in the upper horizon. This mineral is intimately associated with sphalerite and pyrrhotite. Analytical data on electrum are available only from the Shimokawa (Maeda et al., 1981). The Ag content ranges from 22.2 to 71.6 atomic %, but average content is ca. 30 atomic %. This variation range is

Present-day Mineralization and Geothermal Systems

381

similar to that of the Kuroko deposits but the Ag content is lower than that of epithermal veins (Chapter 1). Most of other opaque minerals are of secondary origin. For example, carrollite occurs in interstices of pyrite grains in ores from the Sazare and Shirataki mines which are situated in the area of higher metamorphic grade, suggesting its secondary origin (Itoh et al., 1973; Tatsumi et al., 1975). It is thought that the Co minerals may have been formed by reaction between chalcopyrite and Co-beating pyrite during the metamorphic processes (Kase and Yamamoto, 1988). Tetrahedrite-tennantite has been reported from several deposits (e.g., Kune, Okuki, Besshi, Ikadatsu, Sazare) (e.g., Kase, 1986). This mineral occurs in chalcopyrite-rich chalcopyrite-bornite-pyrite ores. Tellurium-bearing tetrahedrite-tennantite (Te: 11.2 wt%) was found from the Sazare and Ikadatsu deposits (Kase, 1986). Galena, tetrahedrite-tennantite, mawsonite and native silver occur in the copper rich ores but not in ordinary pyritic ores and copper rich ores most commonly occur as offshoots, tongues and veins in the deformed deposits. This suggests that these minor minerals formed during the metamorphic deformation stage accompanied by recrystallization. The assemblages of gangue minerals (muscovite, chlorite, actinolite, hornblende, epidote, and plagioclase) are compatible with the regional grade of metamorphism (Takeda and Sekine, 1960), suggesting these silicate minerals formed by regional metamorphism. On the contrary, small amounts of sulfate minerals (anhydrite, gypsum, barite) occur in several deposits. They are thought to be primary minerals precipitated by the mixing of hydrothermal solution and seawater, although these minerals recrystallized. Barite and anhydrite occur in the massive pyritic ores. Barite occurs frequently in the sphalerite-rich layers of the massive ores in the Chihara deposit (Kanehira, 1959). Detailed investigation of mineralogy of the Shimokawa deposit has been carried out by Mariko (1984). The ore is divided into massive, banded and disseminated types. Dominant sulfides are chalcopyrite, pyrite, pyrrhotite and sphalerite. Dominant gangue minerals are calcite, quartz, siderite, chlorite and muscovite. Siderite, epidote, apatite, actinolite and sphene are minor. Cobalt contents of pyrrhotite and pyrite are high, 0.1-0.6 wt% and 0.1-5.7 wt%, respectively. Sphalerite contains high amounts of Fe (14.3-19.8 FeS mol%). The ores of the Hitachi deposits are generally medium to coarse-grained, massive, and composed mainly of pyrite, pyrrhotite, chalcopyrite, sphalerite, and gangue minerals and some ores contain magnetite or galena (Kanehira and Tatsumi, 1970). Quartz, biotite, muscovite, and hornblende are common and barite and cordielite are also found in some ores (Kanehira and Tatsumi, 1970). The data on the chemical compositions of gangue and alteration minerals are scarce. Shinozuka et al. (1999) analyzed amphibole, chlorite and pumpellyite in the basic rocks from the Minamidani area in Maizuru Belt (Fig. 2.50 and Fig. 2.51). In general, chemical compositions of these minerals are more widely variable than those of midoceanic ridges (Humphris and Thompson, 1978). For example, Fe/(Fe + Mg) of chlorite from the Minamidani is generally higher than that of the midoceanic ridge, but some plot in lower Fe/(Fe + Mg) region (Fig. 2.51). This chlorite composition is quite different from that for the Kuroko chlorite, which is characterized by high Mg content (Fig. 1.83).

382

Chapter 2 AI

9 404

/

~

"3008

Fe

Mg

Fig. 2.50. Plot of amphibole compositions on the A1203-FeO*-MgO (AFM) diagram, together with the compositional range of amphibole in the hydrothermally metasomatized basalt from midocean ridges (Shinozuka et al., 1999). 1.0

9Type1

0.9 0.8

~

0.7

"~ -I-.

0.8

o Type2

m n brunsveigite O= r'l

=

LL 0.5

oceanic basalt

LL 0,4

A Type3

B

diabantite

0.3 .,,.,,

0.2 0.1

10 4.0

.

.

.

5.0

.

.

Si

6.0

7.0

Fig. 2.51. Plot of chlorite compositions on the diagram proposed by Hey (1954). The compositional range of chlorite in the metasomatized basalt from midocean ridges is taken from Humphris and Thompson (1978) (Shinozuka et al., 1999).

2.5.4.2. Ore texture C o l l o f o r m textures o f sulfide m i n e r a l s w e r e f o u n d in ores o f the m e t a m o r p h i c g r a d e s l o w e r than the g r e e n schist facies (Kase, 1988; W a t a n a b e et al., 1993). T h e s e

Present-day Mineralization and Geothermal Systems

383

textures are considered to be relict of primary ore texture. It is generally thought that the colloform textures caused by rapid precipitation from the supersaturated solution are probably due to the mixing of hydrothermal solution and ambient cold seawater. Framboidal textures which are common in Kuroko deposits are found in the deposits in less metamorphosed rocks (Watanabe et al., 1993). Yui (1983) reported graded bedding of pyrite observed in the Motoyasu mine (located about 15 km southwest of the Besshi deposit) and thought that the graded bedding of pyrites, neither size grading nor density grading but is reflected by rhythmic changes in the degree of supersaturation due to intermittent discharge of hydrothermal solution on the seafloor. Grain size of pyrite changes with the metamorphic grade. For example, Doi (1961, 1962) found that grain size of pyrite from many deposits in Shikoku in lower grade, intermediate and higher grade zones is less than 0.15 mm, 0.1-0.2 mm and 0.2-0.4 mm, respectively. This grain size change is caused by the recrystallization of fine-grained pyrite by regional metamorphism. Recrystallization texture is also common for chalcopyrite and sphalerite. The occurrence of chalcopyrite and sphalerite filling the interstices between the pyrite cubes and chalcopyrite inclusion within them is considered to be due to the recrystallization of pyritic ores containing chalcopyrite and sphalerite (Yui, 1983). Yui (1983) suggested that these textures found in the Besshi-type deposits are useful in interpreting ore textures of the Kuroko ores, particularly their diagenetic recrystallization features because such textures are commonly observed in the Kuroko ores (Yui, 1983; Eldridge et al., 1983). Chalcopyrite disease in sphalerite which is common in Kuroko deposits has not been reported from the Besshi-type deposits. 2.5.5. Geochemical features

2.5.5.1. Sulfur isotopes A large number of sulfur isotope data on the Besshi-type deposits are available, although the variation in individual deposit has not been studied well (Yamamoto et al., 1968, 1984a,b; Kajiwara and Date, 1971; Miyake and Sasaki, 1980) (Fig. 2.52). The sulfur isotopic compositions of sulfides are different in different regions. 334S for the Besshi-subtype deposits in the Sanbagawa in Shikoku (e.g., Besshi, Shirataki) and in Kii (Iimori) range from 0%o to +4%0 (av. +2%0 to 3%0) which is very close to those of sediment-barren midoceanic ridge deposits (e.g., East Pacific Rise 21 ~ etc.). 334S for the Besshi-subtype deposits in eastern part of Sambagawa (Kune, Minenosawa), and in the Shimanto belt (Makimine) are high, in a range of +8%0 to +11%o. Sato and Kase (1996) pointed out that these values are similar to those of sedimentcovered oceanic ridge deposits (Middle Valley at the northern end of Juan de Fuca Ridge, Escanaba Trough in southern Gorda Ridge). They thought that 334S values for the Besshisubtype in the Hidaka (Shimokawa) and Shimanto (Makimine) are also high, ranging from +5%o to +10%o. Sato and Kase (1996) thought that a significant component of

Chapter 2

384 Massieve sulfides in active spreading ridges

Sediment-covered ridges - 17.8

Sediment-barren ridges ,,

MORB -

-

17.6

0

+5

+10

Hidaka & Shimanto belts

Shimokawa Gojo 9

Makimine

_.

i

.......

..

,

i

,

I-1

I

i " r ' l r

i

.

,

9

Sanbagawa belt

..., n J7

Kune, Honzan w

!

Kune, Nako

r-1

i

Minenosawa

--

!

,

!

. . . . . . . . .

r-1

~

!

[-h

,

Limori Shingu

Sazare

Shirataki '" l l

.

.

.

.

.

.

l

"

"

Besshi "'i

Okuki

._

1

Nanogawa

.

~

~

'

t"l

I

1

t-""l |

9

9

Mino-Tanba belt

Tsuchikura . . . . .

"~

"'1

'

i

'r

w

Hitachi & Maizuru belts

Hitachi

Yanahara "

i

0

i

,

,

.

+5

,5 34S (%o)

,

'

,

!

!

'i

+10

|

Present-day Mineralization and Geothermal Systems

385

the sulfur derived from seawater sulfate deduced in the hemipelagic sedimentary pile in basalt-shale sequence in the formation of the Group B deposits, as well as the sedimentcovered midoceanic ridge deposits (e.g., Escanabat, +7.8%0). However, 334S values of sedimented ridge deposits according to Scott (1997) such as Guaymas Basin are low (-2.3%0 to + 1.1%0, av. +0.4%o), suggesting incorporation of sedimentary sulfur with low ~34S into ore fluids. 834S values for the Hitachi-subtype (e.g., Hitachi) are wide in range from 0 to +6%0. It is interesting to note that this variation range and average value (+4%0 to +5%0) are very similar to those of Kuroko deposits (av. +4.5%0) and active back-arc deposits (Mariana: av. +3.5%0). This 334S characteristic and other geochemical and geological data (metal ratio, igneous activity, mineralogy) clearly indicate that the Hitachi deposits formed under the island arc back-arc tectonic setting similar to Kuroko deposits (Kase and Yamamoto, 1985). 334S values of sulfides from the other Hitachi-subtype deposits (e.g., Yanahara, Tsuboi in the Maizuru) show a narrow range from +0.6%0 to +2.9 %0 (Yamamoto et al., 1968, 1984a). These values are similar to those of midoceanic ridge deposits and of some back-arc deposits (e.g., Okinawa Trough, Clam Site). The 334S data on barites from the Yanahara and Hitachi (Yamamoto et al., 1984b; Kase and Yamamoto, 1985) are +12%o to +15%o which is similar to those of Late Paleozoic seawater sulfate, indicating that barite formed by the mixing of seawater and hydrothermal solution as same as Kuroko barite (Kusakabe and Chiba, 1983).

2.5.5.2. Metal ratios Bulk chemical composition data of the Besshi-type deposits are summarized in Tables 2.23-2.25. Besshi-type deposits are characterized by high Cu/Zn (2-3 in weight ratio) compared with most of midoceanic ridge deposits (EPR21~ 13~ ll~ Explorer Ridge, S. Juan de Fuca, Galapagos Rift, TAG MAR (Mid-Atlantic ridge), Snake Pit MAR, Guaymas Basin, Middle Valley, etc.) (Tables 2.26-2.28) (Scott, 1997). The elements more concentrated to the Besshi-type ores compared with most of midoceanic ridge deposits are Cu, Fe, Co, Ni and probably Te. The contents of Zn, Ba, Ag, Cd, Sn, As, Se and probably Hg of the Besshi-type ores are lower than those of midoceanic ridge ores. Kase and Yamamoto (1988) pointed out that the cobalt-rich and lead- and nickel-poor ore natures of the Besshi-type deposits are recognized also in the copper-rich MAR and GSC (Galapagos Spreading Centre) deposits. Particularly the MAR sulfide deposits (the average of five friable unconsolidated sulfide samples taken from drilling from sulfide mound) are quite similar in composition to the Besshi-type deposits, but zinc and barium-rich deposits such as EPR deposit and the Juan de Fuca deposit are significantly different in composition from the Besshi-type deposits (Table 2.24).

Fig. 2.52. Histograms showing sulfur isotope data for sulfide minerals from the major volcanogenic Cu sulfide deposits within the Jurassic to Cretaceous accretionary terrains in Japan (Sato and Kase, 1996).

Chapter 2

386 TABLE 2.23 Chemical composition of massive ores (Kase and Yamamoto, 1988) wt% Cu Zn Fe S SiO2 A1203 CaO MgO Na20 K20 C

4.77 1.80 39.61 42.91 5.67 1.68 0.78 0.42 0.03 0.11 0.20

Ratios 870 260 7.1 1700 0.09 0.07 0.14 0.11 0.01 0.04 10

Mn Co Ni Pb Se Te As Sb Sn Au Ag Ge Ba Cd V Cr Ti

ppm

Ratios

271 558 34 100 52 13 51 5 25 0.44 23.4 3.0 35 3.0 110 3.0 396

0.29 22 0.45 8.0 1000 260 28 25 13 440 330 2.0 0.08 15 0.81 0.03 0.07

TABLE 2.24 Bulk chemical composition data of the Besshi-type deposits (Besshi), the seafloor sulfide deposits from the Mid-Atlantic Ridge at 23~ (MAR), the Galapagos Spreading Center at 86~ (GSC) and the East Pacific Rise at 21~ (EPR) (Kase and Yamamoto, 1988) Besshi

MAR

GSC

EPR

wt% Cu Zn Pb Fe S

4.77 1.80 0.01 39.61 42.91

5.96 2.28 0.035 37.40 33.47

4.98 0.14 <0.07 44. I 52.2

0.81 32.3 0.32 19.2 35.3

ppm Mn Co Ni Se Te As Sb Au Ag Ge Ba Cd

271 558 34 52 13 51 5 0.44 23.4 3 35 3

132 618 10 280

140 482 3.1 100

18

125 1.8 0.05 < 10 <1 16 < 32

390 <3.5 3 63 <2 489 37 <0.2 1.59 65 2315 600

255

Present-day Mineralization and Geothermal Systems

387

TABLE 2.25 Average chemical compositions of Besshi-type, Besshi-subtype, and midoceanic ridge (MAR, GSC, EPR) ores Besshi-type (av.) a

Besshi-subtype b

MAR

GSC c

EPR

wt%

Zn Cu Pb Fe SiO2 Ba Ca

1.5 5.9 0.05 36.4

1.8 4.8 0.01 39.6 5.7 0.0035 e 0.78 (CaO)

0.03 0.86

ppm Ag Au Hg Cd Sn As Se Sb Mn Co Ni Te Ge V Cr Ti Mo Bi

32.7 814.6 (ppb) 3394.4 (ppb) 47.3 7.1 113.2 52 f 338.3 354.6 1221.4 28.5 4.6 62.1 8.0 78.5

23.4 0.44 3.0 25 51 5 271 558 34 13 3.0 110 3.0 396 -

2.28 5.96 0.035 37.40 18 255 280 132 618 10 -

0.14 4.98 <0.07 44.1 16 (ppm) -

32.3 0.81 0.32 19.2 19.0 d 0.23 11.2 d

< 10 0.05 (ppm) <32 125 1.8 100 140 482 3.1 <1 -

159 0.15 d 2d 600 489 37 63 390 <3.5 3 <2 65

-

a MITI, 1993. b Doi, 1962; Kase and Yamamoto, 1988. c Bischoff et al., 1983. d Hannigton et al., 1991. e Mean of 12 analyses from the Sasase and Besshi deposits. f Kase and Yamamoto, 1988.

T w o p o s s i b l e e x p l a n a t i o n s f o r t h e d i f f e r e n c e in t h e c h e m i c a l c o m p o s i t i o n s o f t h e B e s s h i - t y p e o r e s a n d z i n c - r i c h a n d b a r i u m - r i c h E P R - t y p e r i d g e o r e s a r e (1) t h e d i f f e r e n t site o f o r e d e p o s i t i o n , a n d (2) t h e d i f f e r e n t c h e m i c a l n a t u r e o f o r e fluids. Barite drothermal

and

sphalerite

solution

mixed

(seawater/hydrothermal

tend with

to p r e c i p i t a t e a large amount

at l o w e r of cold

temperature seawater

from

the

(but mixing

hyratio

solution) m a y be less than 0.2). T h e s e m i n e r a l s precipitate on

t h e s e a f l o o r a n d / o r at v e r y s h a l l o w s u b s u r f a c e e n v i r o n m e n t . H o w e v e r , c h a l c o p y r i t e t e n d s to p r e c i p i t a t e f r o m h i g h t e m p e r a t u r e s o l u t i o n s in o r e b o d i e s a n d / o r s e d i m e n t s . U s u a l l y s h a l e w h i c h is r e l a t i v e l y i m p e r m e a b l e bodies. This suggests that hydrothermal

at t h e s u b - s e a f l o o r

overlies the Besshi-type ore

solution could not issue from the seafloor and

388

Chapter 2

TABLE 2.26 Average bulk compositions of samples from seafloor sulfide deposits at sediment-starved midocean ridges in host basalts 21 ~ EPR

13~ EPR

11~ EPR 11

Explorer Ridge 48

S. Juan de Fuca 3

Galapagos TAG MAR Rift 18 59

SnakePit MAR 31

Na:

5

33

wt% Zn Cu Pb Fe SiO2 Ba Ca

19.8 0.6 0.21 12.4 19.0 0.15 11.2

8.2 7.8 0.05 26.0 9.2 0.4 4.5

28.0 1.9 0.07 22.4 1.2 0.06 0.02

5.3 3.2 0.11 25.9 9.1 7.4 1.2

36.7 0.2 0.26 19.7 5.1 0.06 0.05

4.0 4.5 0.04 32.6 21.5 0.04 0.3

25.2 6.3 0.04 15.4 6.0 <0.01 5.0

7.0 12.4 0.07 35.5 4.2 0.0 0.6

98 0.15 2 376 296 23

49 0.42 233 <10 154 -

38 0.15 3 886 5 399 30

97 0.63 11 199 10 544 27

178 0.11 2 519 359 18

46 0.35 5 109 18 139 11

378 5.7 25 701 2 66 95

111 2.2 270 33 364 32

ppm

Ag Au Hg Cd Sn As Sb

a Number of samples analyzed.

probably sulfides precipitate at the sub-seafloor e n v i r o n m e n t at high temperatures in the Besshi-type ore forming environment. Generally, the sediments at midoceanic ridges (EPR, etc.) are not thick and sulfide-sulfate chimneys and m o u n d s form on the seafloor by the mixing of hydrothermal solution and seawater above the seafloor. In such open space environments, the temperature of hydrothermal solution rapidly decreases by the rapid mixing of hydrothermal solution with cold ambient seawater and low temperature minerals such as sphalerite and barite precipitate from the mixed fluid. However, at the sub-seafloor e n v i r o n m e n t high temperature minerals such as chalcopyrite and pyrrhotite containing Co tend to precipitate instead of low temperature minerals (sphalerite and barite). The possibility (2) is difficult to evaluate. In order to evaluate it, the chemical nature of ore fluids have to be estimated, as has been done by Kase and Horiuchi (1996) who estimated C d / Z n , M n / Z n and C o / Z n based on the chemical composition of sphalerite. A c c o r d i n g to their calculations, C o / Z n of ore fluid responsible for the formation of Besshi-type deposits is higher than that of the EPR hydrothermal solution. The C o / Z n of ore fluids may be related to the chemical composition of host rocks, degree of water-rock interaction, temperature, etc. Thus, we need the study on the water-rock interaction experiments using the host rocks in Besshi mine area and n u m e r o u s analytical data on the host rocks in the Besshi-type mine districts.

Present-day Mineralization and Geothermal Systems

389

TABLE 2.27 Average bulk compositions of samples from sediment-hosted seafloor sulfide deposits

N a:

Guaymas Basin 14

Middle Valley 43

Escanaba Trough 32

Atlantic II Red Sea 53

1.0 0.2 0.4 5.9 28.4 0.15 11.2

2.5 0.4 0.04 30.3 17.4 14.9 6.7

5.1 2.3 1.1 31.2 2.2 2.0 0.17

5.8 1.0 0.1 21.1 3.2 0.07

69 0.2 5 29 6 113 158

9 0.14 40 227 33

117 1.7 18.5 308 187 4600 567

85 2.0 6 194 345 32

wt%

Zn Cu Pb Fe SiO2 Ba Ca ppln

Ag Au Hg Cd Sn As Sb

a Number of samples analyzed.

TABLE 2.28 Average bulk compositions of samples from seafloor sulfide deposits at seamounts and back-arcs

Na:

Axial Seamount CASM 14

N. Fiji Basin

Mariana Trough at 18~ 11

Valu Fa, Lau Basin 47

Eastern Manus Basin 26

Jade, Okinawa Trough 17

22.2 0.4 0.35 5.6 28.1 0.8 9.6 0.21

6.6 7.5 0.06 30.1 16.2 10.2 0.8 0.2

10.0 1.2 7.4 2.4 1.2

16.1 4.6 0.3 7.4

26.9 10.9 1.7 14.9

24.5 3.1 12.1 4.8

33.3 3.7

11.6 0.6

7.3 0.3

3.4 -

189 4.9 20.2 522 7 569 349

151 1.0 260 <10 -

184 0.8 22 465 126 190

256 1.4 >1 482 4 2213 51

230 15 17 1155 11,000 1130

1160 3.3 620 31,000

24

wt%

Zn Cu Pb Fe SiO2 12.5 Ba Ca ppm

Ag Au Hg Cd Sn As Sb

a Number of samples analyzed.

-

390

Chapter 2 Iwami E.

. .

Fujimi L.

77

! i i >, 3. o . Fudotaki ! I I I c IIII1 1l-'l I I I I I cr a~

I-I r 7 . 1

I-

Qkuki

I-I I-I

u_ 2~, Oi, Nago. 4

~

17

F~

-

2-Kotsu 2~ Sazare 2-

..

~ 1!

,aOaz. n R ....

~,~7[ -3

~ .... -5 -4 Log Se/S

Fig. 2.53. Se/S atomic ratios of pyrite from the Besshi-type deposits, the deposits of the Fujimi and Fudotaki groups of the Hitachi mine and the Kuroko-type lwami deposit (Kase and Yamamoto, 1988).

2.5.5.3. Se/S of sulfide ore Kase and Yamamoto (1985, 1988) analyzed pyrite, compiled Se/S of pyrite from the Hitachi deposits, and found that the Se/S ratio of pyrite is significantly higher in the deposits of basic volcanic affinity (the Fudotaki group deposits of the Hitachi mine and the Besshi deposits) than in the deposits in felsic volcanic sequence (the Fujimi group deposits of the Hitachi mine and the Iwami Kuroko deposit) (Fig. 2.53). These data may indicate that Se/S of pyrite was reflected by that of ore fluids and not significantly influenced by other physicochemical parameters (fo2, pH, temperature). Particularly, in high fo2 region (oxidized sulfur species and reduced sulfur species dominant region), Se/S of ore fluids varies significantly and probably affect that of pyrite (Shikazono, 1978b). This may suggest that fo2 of ore fluids responsible for the Besshi-type deposits associated with basaltic volcanic activity was low (reduced sulfur species region). The common occurrence of pyrrhotite in the deposits in low-grade metamorphic terrane (Doi, 1962; Takeda and Sekine, 1960) support this view. 2.5.5.4. Co and Ni of sulfide ore According to Kase and Yamamoto (1985), Co and Ni contents of pyrite are higher in the Fudotaki deposits than in the Fujimi deposits in the Hitachi mine district (Fig. 2.54

391

Present-day Mineralization and Geothermal Systems 6 5

Fudotaki

3 2

>,1 ~

[-1

i1111111

o- 61 I - - I

Fujimi

100

300

pprn

M

500

7o0

"

900

Fig. 2.54. Cobalt contents of pyrite from the deposits of the Fujimi and Fudotaki groups of the Hitachi mine. Crossed, solid and hatched squares correspond to pyrite of ores involving large amounts of pyrrhotite, pyrite from sphalerite predominant ores, and pyrite from chalcopyrite predominant ores, respectively (Kase and Yamamoto, 1985).

Fudotaki

>, (D c(D

1

I i!1

[1

z 7 6 5 4 3

n

Fujimi

13" LL

F/'A

2 1

P. 80

100

120

140 ppm

. 160

180 260

"

Fig. 2.55. Nickel contents of pyrite from the deposits of the Fujimi and Fudotaki groups of the Hitachi mine. Symbols as defined in Fig. 2.54 (Kase and Yamamoto, 1985).

and Fig. 2.55). These features together with S e / S ratio of ore, ~34S and occurrence of ore constituent minerals (galena, barite in Fujimi deposit) indicate that the Fudotaki and Fujimi deposits are similar to the Besshi-type and Kuroko deposits, respectively (Kase and Yamamoto, 1985) (Fig. 2.56).

2.5.5.5. Gold in ore As discussed in previous chapters, gold deposition in epithermal systems and backarc basins (Kuroko deposits) occurs in relatively higher fo2 conditions than base metal

392

Chapter 2 5

Fudotaki

3

'2

_[-1

~-~ o 5

N

I I i I I

I !

Fujimi ,,

2

fIR 8a4Spy

Fig. 2.56. Sulfur isotope ratios of pyrite from the deposits of the Fujimi and Fudotaki groups of the Hitachi mine (Kase and Yamamoto, 1985).

(Cu, Pb, Zn, Fe) deposition in epithermal systems. Thus, it is interesting to summarize the gold distribution in Besshi-type deposits, considering the geochemical environment of gold deposition. Native gold is found as minute grains in the Okuki, Sekizen and Yokei deposits (Horikoshi, 1959; Takeda and Sekine, 1960; Tsunori, 1962). Among them the Okuki deposit is the largest and well studied. The average assay of massive ores is reported as 7% Cu, 34% S, 4 g Au/ton, and 60 g Ag/ton. Au content of the ore is very high compared with other Besshi-type deposits. Characteristic ore and gangue minerals in the Okuki deposit are tetrahedrite-tennantite, cobaltite, garnet, anhydrite, and barite. The red-chert containing fine grained hematite, and silicates (garnet, epidote, amphibole, chlorite and stilpnomelane) overlies the massive sulfide ore. These mineralogical data suggest that sulfides and gold precipitated in relatively high fo2 conditions, compared with pyrrhotite-rich ore deposits such as in low-grade metamorphic terrain. It is noteworthy that bornite, chalcocite and tetrahedrite-tennantite which are common minerals in Kuroko deposits occur in gold bearing Besshi-type deposits. Although these minerals are considered to be secondary minerals, depositional environments of these minerals are characterized by higher fs2 and fo2 conditions. It is also noteworthy that these deposits are rich in pyrite rather than pyrrhotite. Probably, Besshi-subtype deposits in Shikoku formed under the higher fQ and /82 conditions than the deposits characterized by pyrrhotite (Maizuru, Hidaka, Kii, east Sanbagawa). Such typical Besshitype deposits (Besshi-subtype deposits in Shikoku) are characterized by simple sulfide mineral assemblage (chalcopyrite, pyrite, small amounts of sphalerite). Inclusion of bornite in pyrite is also common in these deposits.

2.5.5.6. Lead isotopes Sato and Kase (1996) summarized lead isotope data of ores from 12 major Besshi-type deposits (11 Besshi subtypes and 1 Hitachi subtype) (Fig. 2.57).

Present-day Mineralization and Geothermal Systems i

393

|

o,..o~

39

~.o s ~

..Q

o,,O*

K9,_

t9

,,oo..io

.o

.

.-

{~ KI,,!v

...:;';-" -*s

~', ~/ <> ', .,,J ....ii "5' ,,,...,.. r~ 3 _..,...,.......:,,..:........... .,,,. S h i m o k a w a ..,,.. o 9 ,, ,,. ~ ,..,/'

o ..Q c0 oo4

o

I

'Ca ._nahara,%

38

~

"

II 9~

O,e '~

~'

I Ill

910

"'" ii

~,-,*

J ~,, 1 1 s

.,.s

~,, ,~

s"

/"

"

o J

.--"" ""

j~'

MORB

-"""

37 16.0

,

,

.....

Cretaceous-Tertiary vein, s k a m & kuroko deposits in Japan

0_

....

o ~x:~t5.6 n r

15.2

.,.._. 9

.

o o4

t7.5

.......

~e"u 9

.......................... 18.0

9

o

Oceanic sediments /

.................../.. ., ,~.,_...~....='.'i'.2:'2""'-'" . . . . .

:=_= : . . . . =

_

...........

.

9

dLO_. . . . . . . .

t8.5

'

t9.0

206Pb/204Pb

Fig. 2.57. Pb isotopic compositions in volcanogenic Cu sulfide deposits in Japan. Diamond shape: K], K2 and K3 indicate average values of Kuroko-type deposits: K1 Taro (Cretaceous); K2 Northeast Japan (Miocene); K3 southwest Japan (Miocene). Note that the Shimokawa ores (solid square, Besshi subtype) and Yanahara ores (open square, Hitachi subtype), closely associated with sediments, tend to have slightly more radiogenic values than the others (solid circles) (Sato and Kase, 1996).

It is shown in Fig. 2.57 that the lead isotopic variation of the Besshi-subtype is similar to that of midoceanic ridge basalt, suggesting the lead in the Besshi-subtype was derived from mantle. The data from the Shimokawa, and Yanahara deposits (Group B) are slightly more radiogenic than Group A, suggesting that crustal lead was involved in the formation of the S h i m o k a w a deposit, and lead isotopic values for the S h i m o k a w a and Yanahara plot between M O R B and Cretaceous-Tertiary deposits in Japan (Kuroko, skarn, vein-type deposits).

2.5.5. 7. Rb/Sr and Nd/Sm isotopic compositions Watanabe et al. (1993) have determined a whole rock isochron age of 107 4- 15 M a (Vttb -- 1.42 x 1 0 - l l / y ) for well-preserved pillowed basalts in Western Shikoku with their basalt initial ratio of 0.70401.

394

Chapter 2 MORB

//

0.5132~,~ : ,

~

0

30 '

,I

II ' II I

ESr

6O I

Jlo -8

-6 - 4

ENd

-12

"~" Z

Z

t'O Z

OIB

.

BULK EARTH

0.5126

0 Present

--2 ~'---107Ma

. . . .

-

0.5124

L,--...130Ma -4

--'-6

o.5~22 -

! 0.703

--8

I ~-~

z

- -I0

,,

- -12

!: 0.705

- -14

, 0.707

0.709

0.711

87Sr/SSSr

Fig. 2.58. Relationships between eSr and eNd of the pillowed basalts. MORB and OIB indicate midoceanic ridge basalt and ocean island basalt, respectively (Watanabe et al., 1993).

Watanabe et al. (1993) indicated end and eSr for the pillowed basalt plot within the OIB region, but very close to the MORB region (Fig. 2.58), assuming 130 Ma is the age of the original basalt lavas (Isozaki and Itaya, 1990).

2.5.5.8. Geochemical environment of ore deposition The geochemical environment of ore deposition for Besshi-type deposits is generally difficult to estimate because of the effect of metamorphism. As described above, abundant pyrrhotite is associated with pyrite, chalcopyrite and sphalerite in the lower grade deposits, suggesting that this assemblage could be of primary as same as in the modern geochemical submarine hydrothermal deposits associated with midoceanic ridges (Watanabe et al., 1993). Watanabe et al. (1993) described the relic bornite included in pyrite in the lower-grade deposits. They estimated fsz-temperature conditions for pyrrhotite-pyrite-bearing deposits and bornite-pyritechalcopyrite-bearing deposits (Fig. 2.59). If iron content of sphalerite in association with pyrite and chalcopyrite is primary one (0.04-2.25 mol% FeS) (Watanabe et al., 1993), we could define fsz-temperature (Fig. 2.59). However, it is probable that bornite

395

Present-day Mineralization and Geothermal Systems

.

~o

s 9

LL

Highly metamorphosed B: Weakly

'

(~

8

/t'

12

/ 200

/"

B

. . . . Cooling and uplift (Retrograde) < Heating(Prograde)

/

/ /

A:

,,. ,~/" ~ O +

~ PRIMARY ORE-FORMING ENVIRONMENT / (~]]]]]])METAMORPHIC ENVIRONMENT i

I

250

300

I

400

~

,I

500

-

I

600

: 700

Temperature (~ Fig. 2.59. Metallogenic evolution of the Besshi-type deposits expressed on a log fs2-T diagram (Watanabe et al., 1993).

formed later stage than the other opaque minerals (chalcopyrite, sphalerite, pyrite) and the bornite-chalcopyrite-pyrite assemblage and iron content of sphalerite does not reflect initial geochemical environments.

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Chapter 3

Hydrothermal Flux from Back-Arc Basin and Island Arc and Global Geochemical Cycle

3.1. Major element (alkali, alkali earth, silica) flux It was shown in previous chapters that intense hydrothermal activities occurred in the Neogene age in and around the Japanese Islands under the submarine and subaerial environments. In this chapter the influence of these hydrothermal activities on the seawater chemistry, and the global geochemical cycle are considered. The studies on the hydrothermal systems at midoceanic ridges during the last three decades clearly revealed that the seawater-basalt interaction at elevated temperatures (ca. 100-400~ affects the present-day seawater chemistry (Wolery and Sleep, 1976; Edmond et al., 1979; Humphris and Thompson, 1978). For example, a large quantity of Mg in seawater is taken from seawater interacting with midoceanic ridge basalt, whereas Ca, K, Rb, Li, Ba and Si are leached from basalt and are removed to seawater (Edmond et al., 1979; Von Damm et al., 1985a,b). As mentioned already in Chapter 2, submarine volcanism occurs not only at midoceanic ridges but also at subduction-related tectonic settings such as the Shikoku and Daito Basins, Parce Vela Basins, and Mariana Trough, Okinawa Trough and Izu Bonin Arc (e.g., Wood et al., 1980; Dick, 1982; Delaney and Boyle, 1986). We saw in section 2.3.2 that present-day hot spring venting and sulfide-sulfate depositions have been discovered in back-arc basins in the Western Pacific. These intense hydrothermal activities indicate that seawater-volcanic rock interactions are taking place at these environments. Bulk rock chemistry of hydrothermally altered midoceanic ridge basalt has been well studied and used to estimate the geochemical mass balances of oceans today (Wolery and Sleep, 1976; Humphris and Thompson, 1978; Mottl, 1983). In contrast, very few analytical data on hydrothermally altered volcanic rocks that recently erupted at back-arc basins are available. However, a large number of analytical data have been accumulated on the hydrothermally altered Miocene volcanic rocks from the Green tuff region in the Japanese Islands which are inferred to have erupted in a back-arc tectonic setting (section 1.5.3). The age of Green tuff volcanic activity ranges widely from ca. 25 Ma to 2 Ma. Volcanic activity during the early to middle Miocene (25-15 Ma) was intensive, whereas it was weak during the late Miocene to early Quaternary (Sugimura et al., 1963) (Fig. 3.1). The production of lavas and other effusives per unit time reached five or six time more

408

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(3-1)

A large number of analytical data on the altered volcanic rocks from Green tuff regions in Japan are summarized in Fig. 3.2. Although data are scattered, the data

409

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demonstrate that the Mg in seawater exchanged for Ca of basalt on a molar basis. This relationship between the MgO and CaO contents of the basalt from Green tuff region is similar to that for the hydrothermally altered midoceanic ridge basalt (Mottl, 1983). The relationships between MgO content and the contents of the other constituents (K20, SiO2) of the hydrothermally altered volcanic rocks from Green tuff region are also similar to those for the hydrothermally altered midoceanic ridge basalt. It is indicated that SiO2 and K20 contents of basalt and dacite are taken up by the cycled seawater. From the difference between the MgO content of fresh and altered volcanic rocks, Mg uptake by volcanic rocks is estimated to be ca. 1-10 g per 100 g basalt and 1-5 g per 100 g dacitic rocks. It is notable that the MgO uptake by basalt from Green tuff region is similar to that by midoceanic ridge basalt at ridge axis, but that by dacitic rocks is smaller. Gain and loss of other elements (Ca, K, Si) can be also estimated based on the value of Mg uptake and the relationship between MgO content and the contents of these elements. The total volume of volcanic rocks that erupted during the Green tuff volcanic activity (25-2 Ma) is estimated to be 617,000 km 3 (Sugimura et al., 1963). During the early to middle Miocene age (25-15 Ma) the volume was large (15,000 km3), whereas it was small (2,000 km 3) during the late Miocene to Pliocene (Fig. 3.1) (Sugimura et al., 1963). The total volume ratio of acidic to basic volcanic rocks that erupted in the early to middle Miocene is 3 : 2 (Sugimura et al., 1963). Therefore, the most reasonable average Mg uptake is 2-3 g/100 g volcanic rocks. Using this value and duration of volcanic

410

Chapter 3

activity in the early to middle Miocene (about 10 million years), the average rate of annual Mg removal from seawater to volcanic rocks during 25-15 Ma is estimated to be 1 • 0.2 x 1012 g/year. The total mass of volcanic rocks that erupted per year during 25-15 Ma is estimated to be 4 • 1013 g/year. The rate of seawater cycling is 7.7 4- 1.5 • 1014 g/year, if all of the Mg in cycled seawater is removed to volcanic rocks by the reaction of cycled seawater to rocks at elevated temperatures. Thus the average seawater/volcanic rock ratio (by weight) is calculated to be 14 4- 7 which is roughly similar to that estimated for the recharge zone of the present-day midoceanic ridge hydrothermal system (e.g., 10 + 8 by weight; Humphris and Thompson, 1978). Intense submarine and subaerial volcanic activities during the Tertiary at Green tuff regions took place not only at the Japan Sea but also at marginal basins in the circum-Pacific Region. According to the summary of the development of back-arc basins in the Cenozoic age by Tamaki and Honza (1991) (Figs. 3.3 and Fig. 3.4) and Kaiho and Saito (1994) (Fig. 3.5), many back-arc basins (Japan Sea, Kuri, Shikoku, Parece Vela, South China, Sulu, Makassar, Central Scotia, Cayman) widely and rapidly developed during 30-15 Ma. The total volume of volcanic rocks that erupted at back-arc basins in the circumPacific region during the early to middle Miocene is difficult to estimate. However, it is likely that the total volume of submarine volcanic rocks that erupted during early to 65 ~

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411

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middle Miocene age in the circum-Pacific region is at least 20-30 times that of Green tuff activity near the Japanese Islands at that time, if one considers the total area of Green tuff regions in the circum-Pacific region (Yano, 1985), and total volume of oceanic production rate at back-arc basins during that time (Kaiho and Saito, 1994). Therefore, the total eruption rate at back-arc basins is (8-12) x 1014 g/year which is consistent with that by Kaiho and Saito (1994) (7.4 x 1014 g/year). Therefore, it is estimated that the quantity of Mg removal by the reaction of seawater with volcanic rocks at back-arc basins in the circum-Pacific region during 25-15 Ma is 2.6 i 1 x 1013 g/year (Table 3.1). Annual gain for Ca, K, and Si in seawater due to seawater cycling through back-are basins in Green tuff region in the circum-Pacific region during the early to middle Miocene and those through present-day midoceanic ridges are shown in Table 3.1.

Chapter 3

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Hydrothermal Flux from Back-Arc Basin and Island Arc

413

TABLE 3.1 Annual gain and loss for some constituent in seawater due to seawater cycling through back-arc basins in the Green tuff region in the Circum-Pacific region during the early to middle Miocene, and those through present-day midoceanic ridges (in g/year) (Shikazono, 1994) Green tuff region

Midoceanic ridges

Mg

- 2 . 6 4-1 x 1013

Ca

+4.3 4-1.7 • 1013

K

+4.2 • 1.6 x 1013

Si

+ 1.0 4- 0.4 x 1014

- 2 . 4 • 1013 (Mottl, 1983) - 6 . 5 • 1013 (Wolery and Sleep, 1976) +2.3 x 1013 (Edmond et al., 1979) +94-3 x 1013 (Wolery, 1979) +8.1 • 1012 a +3.9 • 1013 b +0.4 • 1013 (Hart and Staudigel, 1982) +2.3 4-1.2 • 1013 (Holland, 1984) + 1.5 • 1013 (Palmer and Edmond, 1985) +6 4- 8 • 1013 (Wolery and Sleep, 1976)

a Estimation assuming molar exchange of Mg and Ca using 2.4 x 1013 g/year for Mg flux. Estimation assuming molar exchange of 1 mol K and using 2.4 x 1013 g/year for Mg flux.

b

Mg removal by seawater cycling through midoceanic ridges has been estimated by several workers. They vary widely and are: - 2 . 4 x 1013 g/year (Mottl, 1983), - 6 . 5 x 1013 g/year (Wolery and Sleep, 1976), and - 3 . 9 x 103 mol/year (Elderfield and Schultz, 1996). Kaiho and Saito (1994)estimated 2 0 x 106 km3/m.y, and 2 x 106 km3/m.y. for present-day midoceanic ridge crustal production rate and back-arc basin crustal production rate, respectively. If their estimates are correct, Mg removal to midoceanic ridge basalt during early-middle Miocene age is estimated to be 2.6 • 1 x 1013 g/year. Although estimates of annual Mg removal by interaction of circulating seawater with midoceanic ridge basalt are uncertain, it seems likely that Mg removal by seawatervolcanic rock interaction at back-arc basins corresponds to that of Mg removal at midoceanic ridge axis.

3.2. Volatile element (CO2, S, As) flux 3.2.1. CO2 flux

Previous studies demonstrated that the C O 2 fluxes by hydrothermal solution and volcanic gas from midoceanic ridges play an important role in the global CO2 cycle and affect the CO2 concentration in the atmosphere (e.g., Javoy, 1988). However, submarine volcanism and hydrothermal activity occur not only at midoceanic ridges but also at island arc and back-arc basins as already noted. The CO2 concentrations of present-day hydrothermal solutions venting from backarc basins and midoceanic ridges are summarized in Table 3.2. The data show that the CO2 concentrations of hydrothermal solution from back-arc basins and midoceanic

414

Chapter 3

ridges range mostly from 30 to 200 mmol/kg. H20 and from 5 to 30 mmol/kg. H20, respectively. The CO2 concentrations of hydrothermal solutions at Guaymas Basin vary widely and some data show high CO2 concentrations. These high CO2 concentrations and low 313C of fluids (-10.5%o) are considered to be caused by the effect of decomposition and dissolution of organic matters and carbonates in the sediments overlying basalt (Simoneit et al., 1984). The CO2 concentrations of hydrothermal solution from back-arc basins can be also estimated from the fluid inclusion data on Kuroko deposits (Table 3.2). In order to estimate hydrothermal CO2 flux, fluxes of hydrothermal solution into the ocean have to be estimated. As discussed in sections 1.5.3, 2.3 and 2.4.1, the hydrothermal solutions both from back-arc basins and midoceanic ridges are dominantly of seawater origin. Therefore, the fluxes of hydrothermal solution are estimated from seawater cycling rate. This rate is considered to be equal to oceanic production rate times seawater/rock ratio. Kaiho and Saito (1994) estimated the crustal production rate at back-arc basins (Okinawa, Mariana, Andaman, Manus, Woodlark, North Fiji, Lau-Havre, East Scotia and Cayman) based on the spreading rate, thickness of crust and length of ridge axis. Their estimated oceanic crustal production rate is 8.5 x 106 km3/m.y, which is roughly equal to 2.5 x 1015 g/m.y. The seawater/rock ratio can be estimated from the chemical and isotopic compositions of rocks altered by seawater-rock interaction and fresh rocks (Table 3.3). Particularly Mg concentration of altered rocks is useful for the estimation because Mg in seawater removes almost completely into the rocks through the reaction of seawater with rocks at elevated temperatures. Shikazono (1994) summarized the chemical compositions of altered igneous rocks in the Green tuff region of Miocene age in Japan and estimated seawater/rock ratios (by weight) to be 12-18. Oxygen isotopic compositions of altered rocks in the Kuroko mine area can be used to estimate seawater/rock ratio (by weight) to be more than 1 to less than 40 (Green et al., 1983; Shikazono et al., 1995). Therefore, seawater/rock ratio for the discharge zone is assumed to be 2. The average seawater/rock ratio for the submarine hydrothermal system is higher, probably 5-20. Using 2.5 x 1015 g/m.y, as oceanic production rate and 5-20 as seawater/rock ratio and assuming that 30% of oceanic crust interacts with circulating seawater, and the crustal production rate is (0.8-1.1) • 1019 kg/m.y., then the rate of seawater cycling through back-arc basin is estimated to be ( 4 - 2 2 ) x 1019 kg/m.y. Using this value and the CO2 concentration of hydrothermal solution ((0.05-0.3) mol/kg. H20) (Table 3.2), hydrothermal CO2 flux into the ocean is estimated to be (0.2-6) x 1019 kg/m.y. The CO2 flux by hydrothermal solution from midoceanic ridges can be estimated based on the similar procedure mentioned above. The crustal production rate at midoceanic ridges (discharge zone) is 5 x 1022 g/m.y, by Kaiho and Saito (1994) and 4.5 x 1022 g/m.y, by Holland et al. (1996). Seawater/rock ratios for midoceanic ridge hydrothermal systems previously estimated vary widely: Humphris and Thompson (1978), 2-17 (by weight); Wolery and Sleep (1976), 3.5 (by weight); Holland (1978), 10 (by weight). If we accept 5-20 for the seawater/rock ratio, we can calculate the seawater cycling rate at midoceanic ridges as (0.8-4.6)x 1017 g/year. This value is consistent

415

Hydrothermal Flux from Back-Arc Basin and Island Arc TABLE 3.2

CO2 concentration of hydrothermal solutions venting from midoceanic ridges (MOR) and back-arc basins (BAB) Locality

CO2 (mmolal)

Temperature (~

209 160-200 64-96 34-42 43.4 42.1 11.1-14.4 196 134

320 220 267-278 296-311 238-287 220 285-291 250-280 250-280

5.7-8.0 11-18 2.6-6.5 16-24 more than 9 8-22 3.7-4.5 179-285 4-12 (90-115) 50

273-355 354-381 less than 403 270-315 220 345--400 140-332 28-299 265-276 283-321 328

10 140 350 240 210 230 150 90 160 60 100 130 310 222 835 1001 134 302 223 190 257

300-400

BAB (back-arc basins)

Okinawa Trough JADE CLA South Ens lzu-Bonin Suiyo SM Mid-Mariana Trough Alic Springs South Mariana Trough Forest vent North Fiji Basin Myojinsyo

MOR (midoeeanic ridges) East Pacific Rise 21~ 13~ 9-10~ Guaymas Basin Escanaba Trough Juan de Fuca Endeavour Juan de Fuca Rise S. Cleft. Seq. Juan de Fuca Rise ASHES Juan de Fuca Rise Middle Val Dead Dog Bent Bent Ril Juan de Fucca ASHES (Cl-normal) Juan de Fuca Rise ASHES (Cl-enriched)

Kuroko deposits Uwamuki No. 2 Uchinotai W Fukazawa

Matsumine

Shakanai Kosaka Matsumine Nurukawa

300-400 200-300 200-300 300-400 200-500 200-500 200-300 300-500 200-300 230-270 230-270 230-270 230-270 230-280 230-280 230-280 230-280 230-280

416

Chapter 3

TABLE 3.3 Water/rock ratio in midoceanic ridge and Kuroko hydrothermal system (Shikazono, 1988) 1. Geochemical estimate

Midoceanic ridge 2

3-7 5 7-16 1

1-5 1-5 0.5 Cyprus

5.6 0.75 0.43

87Sr/86Sr 87Sr/86Sr Chemical composition of hydrothermal solution Chemical composition of hydrothermal solution 8180 87Sr/86Sr 818O 818O 87Sr/86Sr 8180 S content

Ku roko 1

>2 >1.4 >2.3

Chemical composition of ore fluids and rocks Mg content Mg content Ca content

2. Geophysical estimate

Midoceanic ridge

3.5 0.5 <6.6 2 2-10 2-2.5

with 3 4-1 x 1016 g/year (axial flux) by Elderfield and Schultz (1996). The seawater cycling rate at the hydrothermal system (axis, flank and off axis) is higher than this value: Wolery and Sleep (1976), 2.9 x 1017 g/year; Holland (1978), 1.0 x 1017 g/year; Holland (1984), 0.5 x 1017 g/year; Lister (1973), (0.6-3.6) x 1017 g/year; Palmer and Edmond (1989), 1.2 x 1017 g/year; Kadko et al. (1995), 2.9 x 1017 g/year; Holland et al. (1996), 4.1 x 1017 g/year. The CO2 concentration of the hydrothermal solution of midoceanic ridges is ca. (0.005-0.01) mol/kg. H20 (Table 3.2). Thus, hydrothermal CO2 flux from midoceanic ridges is estimated to be (2.2-18.0)x 102~215 0 . 0 0 5 - - ( 0 . 4 - 9 ) x 1018 mol/m.y. This is consistent with the estimate of Elderfield and Schultz (1996). The value of (0.4-9.0) x 10 Is mol/m.y, can be compared with the CO2 flux by volcanic gas from midoceanic ridges: Marty and Jambon (1987), 2.2 x 1018 mol/m.y.; Des Marais (1985), 1-8 x 1018 mol/m.y.; Tajika and Matsui (1990), 4.0 x 1018 mol/m.y.; Tajika (1998), 2.0 x 1018 mol/m.y. Javoy et al.'s (1982) estimate (20.3 x 1018 mol/m.y.) is quite different and higher than those by other researchers.

Hydrothermal Flux from Back-Arc Basin and Island Arc

417

TABLE 3.4 Estimates of present-day global fluxes (mol/year) of CO2 to and from the atmosphere (Seward and Kerrich, 1996; Shikazono and Kashiwagi, 1999) Geologic process

Global flux

Chemical weathering Midoceanic ridge system Subaerial volcanism Subaerial + submarine volcanism Anthropogenic Back-arc Subaerial hydrothermal

-6.7 x 1012 +0.7 to 1.5 x 1012 +1.2 to 1.8 x 1012 -+-2to 4 x 1012 +5 x 1014 +5 to 16 x 1012 > 1012

The above argument indicates that hydrothermal C02 flux from back-arc basins is similar to or greater than that from midoceanic ridges, and thus the hydrothermal flux from back-arc basins as well as hydrothermal flux from midoceanic ridges have to be taken into account when we calculate global geochemical CO2 flux. At the convergent plate boundaries, CO2 degasses not only from back-arc basins by hydrothermal solutions but also from terrestrial subduction zones by volcanic gases and hydrothermal solutions. However, the studies on CO2 degassing from terrestrial subduction zones are not many. Seward and Kerrich (1996) have shown that hydrothermal CO2 flux from terrestrial geothermal system (such as Taupo volcanic zone in New Zealand) exceeds 1012 mol/year which is comparable to that of midoceanic ridges (Table 3.4). Sano and Williams (1996) calculated present-day volcanic carbon flux from subduction zones to be 3.1 x 1012 tool/year based on He and C isotopes and CO2/3He ratios of volcanic gases and fumaroles in circum-Pacific volcanic regions. Williams et al. (1992) and Brantley and Koepenich (1995) reported that the global CO2 flux by subaerial volcanoes is (0.5-2.0) x 10 is mol/m.y, and (2-3) x 10 is mol/m.y. (maximum value), respectively. Le Guern (1982) has compiled several measurements from terrestrial individual volcanoes to derive a CO2 flux of ca. 2 x 10 is mol/m.y. Le Cloarec and Marty (1991) and Marty and Jambon (1987) estimated a volcanic gas carbon flux of 3.3 x 1017 mol/m.y, based on C/S ratio of volcanic gas and sulfur flux. Gerlach (1991) estimated about 1.8 x 10 is mol/m.y, based on an extrapolation of measured flux. Thus, from previous estimates it is considered that the volcanic gas carbon flux from subduction zones is similar to or lower than that of hydrothermal solution from back-arc basins.

3.2.2. Causes for high C02 concentration and origin of C02 of hydrothermal solution from back-arc basins The main alteration minerals surrounding Kuroko ore body are K-mica, K-feldspar, kaolinite, albite, chlorite, quartz, gypsum, anhydrite, and carbonates (dolomite, calcite, magnesite-siderite solid solution), hematite, pyrite and magnetite. Epidote is rarely found in the altered basalt (Shikazono et al., 1995). It contains higher amounts of ferrous iron (Fe203 content) than that from midoceanic ridges (Shikazono, 1984).

418

Chapter 3

1 o o

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150

I

200

I

250

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300 350 Temperature('C)

Fig. 3.6. log fco2-temperature diagram showing the univariant equilibrium curves for some gangue minerals. A: 2Ca2A13Si3OIz(OH) (clinozoisite) + 3SIO2 (quartz)+ 2CACO3 (calcite) + 2 H 2 0 - - 3CazA12Si30]0(OH)2 (prehnite) + 2CO2 ( X p i s - - 0 . 3 ) . B: Ca6Si6OI7(OH)2 (xonotlite) + 6CO2 -- 6CaCO3(calcite) + 6SiO2(quartz) + H20. C: CaCO3 (calcite) + TiO2 (rutile) + SiO2 (quartz) --CaTiSiO5 (sphene) + CO2. D: MnSiO3 (rhodonite) + CO2 -- MnCO3 (rhodochrosite) + SiO2 (quartz). E: 3 KAI3Si3010(OH)2 (Kmica) + 4CACO3 (calcite) + SiO2 (quartz) = 2CazA13S3OIz(OH)(clinozoisite) + 3KAISi308 (K-feldspar) + 4CO2 + 2 H 2 0 ( X p i s = 0 . 3 ) . F: 3KAI3Si3OI0(OH)2 ( K - m i c a ) + 4CACO3 (calcite)+ SiO2 (quartz)= 2CazA13Si3OI2(OH) (clinozoisite) + 3 KA1Si308 (K-feldspar) + 4CO2 + 2 H 2 0 ( X p i s = 0 . 2 5 ) . G" 3 FeCO3 (siderite) + (1/2)O2 = Fe304 (magnetite) + 3CO2, C (graphite) + O2 = CO2. H: 3 CaMg(CO3)2 (dolomite) + KAISi308 (K-feldspar) + H 2 0 - - 3CACO3 (calcite) + 3CO2 + KMg3(AISi3OI0)(OH)2 (phlogopite). I: C (graphite) + 02 = CO2, FeS (pyrrhotite) + 1/2S2 = FeS2 (pyrite), 2H2S(aq) -Jr- O2 -- $2 + 2H20(1). J: FeCO3 (siderite) + Fe203 (hematite) = Fe304 (magnetite) + CO2. K: CaA12Si4OIz2H20 (wairakite) + KAISi308 (K-feldspar) + CO2 = CaCO3 (calcite) + KAI3Si3Oj0(OH)2 (K-mica) + 4SIO2 (quartz). L: CaAlzSi4OIz2H20 (wairakite) + CO2 = CaCO3 (calcite) + AIzSi2Os(OH)4 (kaolinite) + 2SIO2 (quartz). M: Ca(A12Si4Ol2.4 H20 (laumontite) + CO2 = CaCO3 (calcite) + AlzSi2Os(OH)4 (kaolinite) + 2 SiO2 (quartz) + 2 H20. Solid star: midoceanic ridge, solid circle: back-arc basin (modified after Shikazono, 1985).

Hydrothermal alteration minerals from midoceanic basalt are analcite, stilbite, heulandite, natrolite-mesolite-scolecite series, chlorite and smectite for zeolite facies, prehnite, chlorite, calcite and epidote for prehnite-pumpellyite facies, albite, actinolite, chlorite, epidote, quartz, sphene, hornblende, tremolite, talc, magnetite, and nontronite for green schist facies, hornblende, plagioclase, actinolite, leucoxene, quartz, chlorite, apatite, biotite, epidote, magnetite and sphene for amphibolite facies (Humphris and Thompson, 1978). The fco2-temperature relationships for the above-mentioned mineral-fluid equilibria are shown in Fig. 3.6. Based on the thermochemical calculations, minerals summarized above and temperatures estimated, we could estimate typical fco2-temperature ranges for hydrothermal solutions from midoceanic ridges and back-arc basins. The analytical data on CO2 in hydrothermal solutions and fluid inclusions and measured temperatures (Table 3.2) are consistent with the thermochemical calculations mentioned above.

Hydrothermal Flux from Back-Arc Basin and Island Arc

419

Ishibashi and Urabe (1995) considered that the high volatile concentrations (C02, etc.) of hydrothermal solutions venting from back-arc basins are due to higher contribution of magmatic fluids containing high concentrations of volatiles than hydrothermal solutions at midoceanic ridges. However, the high concentration of CO2 of hydrothermal solution from back-arc basins does not imply larger contribution of magmatic fluids to hydrothermal solutions at back-arc basins, but is seems more likely from the abovementioned reasons that the alteration minerals buffering fluid chemistry are different for midoceanic ridges and back-arc basins and CO2 in hydrothermal solution was derived from carbonates in altered rocks and not directly from magma. 813C data on the carbonates in altered volcanic rocks in back-arc basin (Kuroko mine area) are -5%~, indicating that most carbon in carbonates were of igneous origin (Shikazono et al., 1995). Stable isotopic studies of 8180 and 8D of hydrothermal solutions venting from back-arc basins show no evidence of contribution of magmatic fluids to the hydrothermal solutions at back-arc basins and midoceanic ridges. As noted already, the stable isotopic data (834S, 813C, 8180, 8D) all indicate that hydrothermal solutions in submarine hydrothermal system in back-arc basins and midoceanic ridges were generated by seawater-rock interaction at hydrothermal conditions. The CO2 concentrations of present-day geothermal waters in terrestrial environment have been also interpreted in terms of the interaction of hydrothermal solutions with country rocks (Giggenbach, 1981; Shikazono, 1978,1985). For example, as noted in section 2.4.3, Shikazono (1985) estimated fco2 for epithermal Au-Ag and base-metal veintype deposits in Japan which formed in terrestrial environments at Miocene-Pliocene age and showed that fco2 is controlled by the alteration minerals (Fig. 3.6). Estimated fcoztemperature range for epithermal Cu-Pb-Zn vein-type deposits are clearly similar to those for the Kuroko and back-arc deposits in which base metals (Cu, Pb, Zn) are concentrated. It is likely that the minerals controlling fco2 of hydrothermal solution at back-arc basins are dolomite, siderite, calcite, hematite, magnetite, graphite, K-mica and kaolinite. Most of these minerals are not found in altered ridge basalt. Among the minerals mentioned above calcite and kaolinite may be important for controlling fco2 of terrestrial geothermal waters. It was cited by Giggenbach (1981) that fco2 (or Xco2, mole fraction of CO2) of terrestrial geothermal waters is controlled by "plagioclase" + CO2 -- calcite + "kaolinite". Berndt et al. (1989) have indicated that acaz+/a2+ and aNa+/aH+ of midoceanic ridge hydrothermal fluids is controlled by clinozoisite, Ca-feldspar, and Na-feldspar. In addition to these assemblages, calcite is in equilibrium with fluids. Therefore, we can derive the fcoz-temperature relationship from the following equilibrium relations. Electroneutrality relation is approximated by, mNa+ -- me1-

(3-2)

The mc1- of midoceanic ridge hydrothermal solution is generally close to that of seawater. If the above argument is correct and the fco2 of hydrothermal solutions from backarc basins is in equilibrium with alteration mineral assemblage including dolomite and calcite, mgg2+/inCa2+ of fluids can be estimated to be 0.03-0.055 from dolomite-calcite-

420

Chapter 3

fluid equilibrium (250-300°C) (mMg2+/mca2+-temperature diagram). Using estimated of hydrothermal solution from back-arc basins (10-30 mmol/kg. H20) (Gamo, 1995), mMg2+is estimated to be 0.1-2 mmol/kg. H20. Usually concentration of endmember hydrothermal solution is estimated by the extrapolation of mMg2+ to be zero. However, the above argument demonstrates that Mg2+ concentration of end member hydrothermal solution prior to the mixing with ambient cold seawater is not zero. This is supported by the thermochemical consideration on the equilibrium between chlorite and hydrothermal solution (Tamura, 1982). Figure 1.84 shows the variation of aFe2+/aMg2+ of hydrothermal solution in equilibrium with chlorite having constant Fe2+/Mg2+ (= 0.6) as a function of temperature (Shikazono and Kawahata, 1987). The Fe concentration of hydrothermal solution from back-arc basins is usually very low. For example, the Fe concentration of hydrothermal solution from Okinawa Trough (JADE), Manus, Izu-Bonin (Suiyo, SM), Mid-Mariana Trough (Alice Springs), South Mariana Trough (Forecast vent), Lau Basin (Vaiu Fa) is 2.8, 435, 6.4, 1 1, 77, 9-1 3 (kmol/kg. H20), respectively. Although Fe concentration varies widely, however, if we take Fe concentration as (10-100) Fmol/kg. H2O (= 0.010.1 mmol/kg.H20), and aFe2+/aMg2+ = 10-3-10-4 from Fig. 1.84, aMg2+is estimated to be (0.1-10) mmol/kg.H20. A 10 mmol/kg.H20 Mg concentration seems too high. However, one to several mmol/kg. H20 seems likely and is consistent with the estimate of a@+ from dolomite-calcite equilibrium. Shikazono ( 1978) theoretically derived that the concentrations of alkali and alkali earth elements in chloride-rich hydrothermal solution are nearly in equilibrium with hydrothermal alteration minerals such as albite, K-feldspar, K-mica, quartz, calcite, wairakite, and Mg-chlorite. If we use 500 mmol/kg. H 2 0 as the average C1- concentration of hydrothermal solution from the back-arc basin, mMg2+, which is in equilibrium with albite, K-feldspar, K-mica, quartz and Mg-chlorite at 250"C, is estimated to be 10-4-10-3 mol/kg. H20. Therefore, it is likely that end-member hydrothermal solution contains 0.1-10 mmol/kg. H20 Mg. Therefore, strictly speaking, the method previously used to estimate the chemistry of end-member hydrothermal solution is not correct, if equilibrium between alteration minerals and hydrothermal solution is attained at deep and high temperature conditions.

mCa2+

3.2.3. S flux Although the global S cycle has been studied by several investigators (e.g., Holser and Kaplan, 1966; Holland, 1978), hydrothermal S flux has not been considered. Thus, hydrothermal S flux is estimated below. H2S concentration in midoceanic ridge hydrothermal solution is mostly in a range of (1-6) mmol/kg. H 2 0 (Gamo, 1995). Using this concentration and rate of seawater cycling (= (1 -4) x lo2" kg/m.y.), we obtain hydrothermal S flux from midoceanic ridge as (2-24) x I O l 7 mol/m.y. The most likely flux is 6 f2 x lo'' mol/m.y. This estimated value is consistent with the value by Elderfield and Schultz (1996) which is (0.85-9.6) x 10'' mol/m.y. kg/m.y. (mostly ( 6 f Shikazono and Kashiwagi (1999) estimated (4-22) x 2) x 1019kg/m.y.) as a rate of seawater cycling at present-day back-arc basin. Using these

Hydrothermal Flux from Back-Arc Basin and Island Arc

421

values, we obtain hydrothermal S flux as (0.8-3.2) x 1017 mol/m.y. This estimated value is lower than that from midoceanic ridges and seems to be comparable to volcanic gas flux from terrestrial island arc ((1.5-5) x 1017 mol/m.y.) (Wolery and Sleep, 1976). Sulfur in the sediments and oceanic crust which is derived from seawater subducts to deeper parts. This subduction flux is estimated to be ca. 4 x 1017 mol/m.y. (Shikazono, 1997). Therefore, degassing S flux from back-arc and island arc ((2.3-8.2)x 1017 mol/m.y.) seems to be not different from the subduction flux, although uncertainty of estimated degassing and subduction flux is large. We speculate from the above argument that primordial sulfur degasses from midoceanic ridges even at present time as well as He, because subduction flux to mantle seems to be small. However, we need more detailed study on long-term S cycle including hydrothermal S flux to evaluate this speculation. 3.2.4. As flux

The As (arsenic) concentration of seawater is controlled by input of rivers, sedimentation on the seafloor, weathering of the seafloor, exchange between atmosphere and seawater, volcanic gas input, and hydrothermal input. Previous studies on the geochemical cycle of As have not taken into account the hydrothermal flux of As. Therefore, hydrothermal flux of As from back-arc, island arc and midoceanic ridges to ocean is considered below. As concentrations of submarine and subaerial hydrothermal solutions are summarized in Table 3.5, which clearly shows that the As concentration of hydrothermal solutions from back-arc basins and from subaerial island arc are higher than those from midoceanic ridges. Many analytical data on As concentration of hot springs in subaerial island arc are available (Ellis and Mahon, 1977; Weissberg et al., 1981). These data clearly indicate high concentration of As in these hot springs. An experimental study at 350~ on the interaction between NaC1 solution and graywacke which occurs widely in island arc geologic setting indicates that the final solution contains (0.6-0.7) ppm As (Bischoff et al., 1981). Analytical data on As concentration of hydrothermal solution at back-arc basins are few. Arsenic concentration of hydrothermal solution at Lau Basin is 6.0-8.2 ppm (Foquet et al., 1991). We can also estimate As concentration of hydrothermal solution based on the solubility data on orpiment and realgar because these As-bearing minerals are common in back-arc basin deposits (e.g., Okinawa Trough, Kuroko deposits). Based on the above data and argument, it is inferred that hydrothermal solution venting from back-arc basins contains appreciable amounts of As (1-5 ppm). Using 2.5 x 1019 kg/m.y, as the ocean crustal production rate at back-arc basins (Kaiho and Saito, 1994), 1-10 as seawater/rock ratio, and 1-5 ppm As concentration, we can estimate hydrothermal As flux from back-arc basins to be (1.3-0.13)x 1018 g As/m.y. It seems unlikely that all of the oceanic crust produced interacts with seawater (Holland, 1978). Accepting that 30% of oceanic crust interacts with circulating seawater, hydrothermal As flux is estimated to be (3.8-0.1) x 1011 g As/year. This flux, although

422

Chapter 3

TABLE 3.5 Arsenic concentrations of hydrothermal solutions issuing at midoceanic ridges, back-arc basins and island arcs (Shikazono, 1993) As concentration (ppm) Midocean ridges NGS 21 ~ OBS North SW HG Guaymas

0.002 0.02 0.02 0.03 0.02-0.08

Back-arc basins Lau Basin

6.0-8.2

Island ares Apapel Springs, Kamchatka Broadlands, New Zealand drill 2 Surface Deep aquifer Cheleken, USSR: i Cheleken, USSR: ii Cheleken, USSR: iii Dvukhyurtochnye Springs, Kamchatka Mendeleyev Volcano, Kurile Islands Caldera Springs, Kamchatka Waiotapu, New Zealand, Champagne Pool Wairakei, New Zealand, Hole 44 Pauzhetsk, Kamchatka Steamboat Springs, USA Ngawha, New Zealand

2.5-3.0 8.1 5.5 0.1 0.03 0.5 2.8 2.2 25 4.9 4.8 1.0 2.7 0.2

uncertainty is large, is comparable to riverine flux (7.8 x 10 I~ g/year). Hydrothermal As flux from midoceanic ridges can be also estimated using oceanic crustal production rate (= 5 x 10 22 g/m.y.) (Kaiho and Saito, 1994) and As concentration of hydrothermal solution. This estimated flux is (0.5-0.06) x 10 l~ g/year. This estimated flux is considerably small compared with hydrothermal flux from back-arc basin and riverine flux. The flux of volcanic gas to ocean has not been estimated. Walsh et al. (1979) estimated As flux of volcanic gas to atmosphere to be 2.8 x 109 g/year. Therefore, this flux to ocean is also small. Arsenic removes from basalt to seawater by the weathering of ocean floor basalt. Kawahata and Shikazono (1988) found that the sulfur content of the midoceanic ridge basalt at Galapagos rift decreases from ca. 1,000 ppm to ca. 400 ppm by the seafloor weathering. Average As/S ratio of pyrite is (8.7 4- 3) x 10 -4 (Fleisher, 1955; Utter, 1978; Huerta-Diaz and Morse, 1992). Using this ratio and assuming that all of As in pyrite remove to seawater by the seafloor, weathering and volume of basalt suffered by the seafloor weathering is (4.5-15) x 1014 g/year [this is estimated by assuming that 60-200 m thick basalt is weathered and ocean production rate is 3 x 10 l~ cmZ/year (Deffeyes,

Hydrothermal Flux from Back-Arc Basin and Island Arc

423

TABLE 3.6 Geochemical balance of arsenic in ocean and subduction flux (g/year) (Shikazono, 1993) Input flux to ocean (1) River flux (2) Hydrothermal flux: island arc-back-arc basin axis of midocean ridge (3) Volcanic gas (4) Atmosphere (5) Weathering of ocean floor of basalt

7.8 x 10 l~ (0.2-5.2) x 10 l~ (0.8-1.6) x 1011 2.8 x 109 (max.) 2 . 6 x 109 (2.7-9) x 10 8 Total: (1.0-6.1)x 1011

Output flux from ocean (6) Sedimentation (formation of pyrite) (7) Atmosphere

(1.3-2.9) x 10 ll 1.4 x 10 8 Total: (1.3-2.9) x 1011

Subduction flux

(4.0-8.2) x 10 l~

1970)]. Arsenic flux by the weathering is estimated to be (2.7-9) x 108 g/year. Total As input is a sum of hydrothermal, volcanic gas, riverine, and weathering fluxes which is equal to (1.1-6.1) x 1011 g As/year. Arsenic removal from seawater to sediments is mainly governed by pyrite formation in the seafloor sediments. Production rate of sedimentary pyrite is 2.5 x 1014 g S/year (Holland, 1978). Therefore, As removal by pyrite from seawater is (1.3-2.9) x 1011 g/year. This is the same order of magnitude as As input to ocean by river which is equal to 0.7 • 1011 g/year. Therefore, it is likely that the steady state is maintained with regard to As concentration in seawater. The geochemical balance of As in ocean and subduction flux of As are summarized in Table 3.6.

3.3. Other elemental flux 3.3.1. Hg flux Hg concentration in hydrothermal solution from back-arc basins and midoceanic ridges has not been determined. Experimental study on graywacke-water interaction suggests that the hydrothermal solution interacted with graywacke contains n x 10 -2 ppm Hg (Bischoff et al., 1981). Cinnabar and metacinnabar are not common but were reported from several Kuroko deposits (Urabe, 1974). From the solubility data on cinnabar and metacinnabar (Barnes and Czamanske, 1967), we can place a limit on the Hg concentration of ore fluids to be n • 10 -2 ppm. Using n x 10 -2 ppm concentration and seawater cycling rate at back-arc basins, hydrothermal Hg flux from back-arc

424

Chapter 3

basins is estimated to be n x (108-109 g)/year. Riverine Hg flux is estimated from average Hg concentration of river ((4.0-7.4)x 10 -8 mol/kg. H20) and riverine flux (-- 4.6 x 1022 g/year). This riverine Hg flux is (1.8-3.4) x 109 g/year. Therefore, it is likely that hydrothermal Hg flux from back-arc basins may be important for controlling Hg concentration of seawater, although we need more detailed investigation on the hydrothermal Hg flux.

3.3.2. Mn flux The Mn concentration of hydrothermal solution from back-arc basins varies widely from 12 I~mol/kg.H20 to 7100 I~mol/kg.H20 (Lau Basin, North Fiji Basin) (Gamo, 1995). But, it ranges mostly from 10 txmol/kg. H20 to 300 ~tmol/kg. H20. Using this range and seawater cycling rate (= (0.08-0.8)x 1017 g/year), we obtain Mn flux as (0.08-2.4) x 10 l~ mol/year. Present-day ocean production (back-arc)/ocean production (midoceanic ridge) ratio is about 0.1 (Kaiho and Saito, 1994). According to Elderfield and Schultz (1996), the best estimate of axial flux at midoceanic ridge is (3 4- 1.5) x 1013 kg. HzO/year. If we use this value and the production rate, H20 flux from back-arc basins is estimated at (3 4- 1.5) x 1015 kg-HzO/year. Thus, we estimate hydrothermal Mn flux from back-arc basins to be (1.35-0.15) x 109 mol/year. This hydrothermal Mn flux from back-arc basins is less than the value of riverine Mn flux (0.49 x 1010 mol/year). Hydrothermal Mn flux from midoceanic ridge is estimated as (1.1-3.4) x 10 I~ mol/year (Elderfield and Schultz, 1996) which is one to two order of magnitudes greater than the hydrothermal Mn flux from back-arc basins.

3.3.3. Ba flux Ba concentration of hydrothermal solution from back-arc basins ranges from 5.3 gmol/kg. H20 (North Fiji Basin) to 100 gmol/kg. H20 (Izu-Bonin Suiyo SM) (Gamo, 1995). Assuming that Ba concentration is (20-60) Ixmol/kg. H20 and seawater cycling rate is 1.8 x 1016 g/year, we obtain Ba flux as ( 3 - 6 6 ) x 10 l~ mol/year. This is greater than or comparable to that of midoceanic ridge flux (2.4-13 x 10 l~ mol/year) (Elderfield and Schultz, 1996) and is comparable to or greater than that of riverine Ba flux (1 x 101~ mol/year) (Elderfield and Schultz, 1996).

3.4. Comparison of back-arc hydrothermal flux with midoceanic ridge hydrothermal flux Elderfield and Schultz (1996) estimated midoceanic ridge hydrothermal fluxes using heat and water fluxes estimated by various data (3He/heat, Mg concentration, Sr isotopes, Li isotopes, Ge/Si ratio). Their estimated fluxes are presented in Fig. 3.7 and Table 3.7. Hydrothermal flux from back-arc basins estimated based on the H20 flux which was estimated from oceanic crust production rate and seawater/rock ratio at back-arc

425

Hydrothermal Flux from Back-Arc Basin and Island Arc

I

" 102:1

S0 4 M(

1012 9 Fe Mn LI

l

1"1 ~ ' " g l

Si I 9 " "" K +9 I

, j J

cD

9"

1010 9 jj

=

[ Alk

):102

C~a

I

I

9

10 8 Co ,'" Ag~ ." "

P

* 106 104 L" " 10 6

l

Se"l

I

10 8

! 10 TM

I

1012

1014

river flux (mol/year) Fig. 3.7. Comparison of hydrothermal fluxes and river fluxes (data from Table 3.7) (Elderfield and Schultz, 1996).

basins and the concentrations of elements in hydrothermal solution from back-arc basins is given above. The following points are inferred from the comparison of back-arc hydrothermal flux mentioned above with midoceanic ridge hydrothermal flux: (1) Back-arc basin hydrothermal flux of most elements is small, compared with midoceanic ridge hydrothermal flux. (2) Back-arc basin hydrothermal fluxes of CO2, As, and Ba are probably higher than the midoceanic ridge hydrothermal fluxes, although we need more detailed investigation. (3) The average chemical compositions of Kuroko ores and those of back-arc deposits suggest that Hg, As, Sb, T1 and Ba are concentrated to the ore fluids responsible for the Kuroko and back-arc deposits, suggesting that these fluxes from back-arc basins are high compared with midoceanic ridge fluxes. (4) Au, Sb and Hg are more enriched into the ores of back-arc basins compared with midoceanic ridge and thus it is likely that back-arc basin hydrothermal flux for these elements is higher than midoceanic ridge hydrothermal flux. However, the concentrations of these elements in back-arc basin hydrothermal solution have not been analyzed. Thus, we need to accumulate analytical data on the concentration of these elements in back-arc basin hydrothermal solution. Further, H20 flux from back-arc basin has to be estimated based on various methods (3He/He, Mg concentration, Li isotope, Sr isotope, Ge/Si ratio) which was argued for midoceanic ridge hydrothermal system by Elderfield and Schultz (1996), but not for back-arc basin by the previous workers.

426

Chapter 3

TABLE 3.7 Comparison of primary axial high-temperature hydrothermal chemical fluxes and fiver chemical fluxes (Elderfield and Schultz, 1996) Element

Li K Rb Cs Be Mg Ca Sr Ba SO4 Alk Si P

B A1 Mn Fe Co Cu Zn Ag Pb As Se CO2 CH4 H2 H2S

Chydrothermafluid (mol/kg) a

Cseawater (mol/kg) a

411-1322 Ix 17-32.9 m 10-33 Ix 100-202 n 10-38.5 n 0 10.5-55 m

26 Ix 9.8 m 1.3 Ix 2.0 n 0 53 m 10.2 m

87 Ix > 8 to >42.6 Ix 0-0.6 m - 0 . 1 to - 1 . 0 m 14.3-22.0 m 0.5 Ix 451-565 Ix 4 - 2 0 Ix 360-1140 IX 750-6470 Ix 22-227 n 9.7-44 Ix 4 0 - 1 0 6 Ix 26-38 n 9-359 n 30-452 n 1-72 n 5.7-16.7 m 25-100 Ix 0.05-1 m 2.9-12.2 m

87 Ix 0.14 Ix 28 m 2.3 m 0.05 m 2 Ix 416 Ix 0.02 tx 0 0 0.03 n 0.007 Ix 0.01 Ix 0.02 n 0.01 n 27 n 2.5 n 2.3 m 0 Ix 0 m 0 m

Fhydrothermal (mol/year) 1.2 2.3 2.6 2.9 3.0

to 3.9 • 10 l~ to 6.9 • 10 l~ to 9.5 x 104 to 6.0 x 106 to 12 • 105 - 1 . 6 • 1012 9.0 to 1300 • 109

0 >2.4 to 13 x 108 - 8 . 4 x 1011 - 7 . 2 to 9.9 x 101~ 4.3 to 6.6 • 10 II - 4 . 5 x l07 i. 1 to 4.5 • 109 1.2 to 6.0 • 108 1.1 to 3.4 x 10 l~ 2.3 to 19 • 10 l~ 6.6 to 68 x 105 3.0 to 13 • 108 1.2 to 3.2 x 109 7.8 to 11 x 105 2.7 to 110 x 105 0.9 to 140 • l05 3.0 to 220 x 104 1.0 to 12 • l0 II 0.67 to 2.4 • 1() l~ 0.3 to !.5 • 1() j~ 0.85 to 9.6 • 1() II

Frivers (101~ mol/year) 1.4 190 0.037 0.00048 0.0037 530 1200 2.2 1.0 370 3000 640 3.3 5.4 6.0 0.49 2.3 0.011 0.50 1.4 0.0088 0.015 0.072 0.0079

a m = 10-3; tx = 10-6; n = 10 -9.

References Barnes, H.L. and Czamanske, G.K. (1967) Solubilities and transport of ore minerals. In: Barnes, H.L. (ed.), Geochemistry of Hvdrothermal Ore Deposits. New York: Holt, Rinehart and Winston, pp. 334-381. Berndt, M.E., Seyfried, W.E. Jr. and Janeckey, D.R. (1989) Plagioclase and epidote buffering of cation ratios in midocean ridge hydrothermal fluids: Experimental results in and near the supercritical region. Geochim. Cosmochim. Acta, 53, 2283-2300. Bischoff, J.L. and Dickson, F.W. (1975) Seawater-basalt interaction at 200~ and 500 bars: Implications for origin of seafloor heavy-metal deposits and regulation of seawater chemistry. Earth Planet. Sci. Lett., 25, 385-397. Bischoff, J.L., Radke, A.S. and Rosenbauer, R.J. (1981) Hydrothermal alteration of graywacke by brine and seawater; role of alteration and chloride complexing on metal solubilization at 200~ and 350~ Econ. Geol., 76, 659-676.

Hydrothermal Flux from Back-Arc Basin and Island Arc

427

Brantley, S.L. and Koepenich, K.W. (1995) Measured carbon dioxide emissions from Oldoinyo Lengai and the skewed distribution of passive volcanic fluxes. Geology, 23, 933-936. Deffeyes, K.S. (1970) The axial valley; a steady-state feature of the terrain. Chapter 9 in: Johnson, H. and Smith, B.L. (eds.), The Megatectonics of continents and oceans. New Brunswick, New Jersey: Rutgers U. Press. Delaney, M. and Boyle, E.A. (1986) Lithium in foraminiferal shells: Implications for high-temperature hydrothermal circulation fluxes and oceanic crustal generation rates. Earth Planet. Sci. Lett., 80, 91-105. Des Marais, D.J. (1985) Carbon exchange between the mantle and crust, and its effect upon the atmosphere: Today compared to Archean time. In: Sundquist, E.T. and Broecker, W.S. (eds.), The Carbon Cycle and Atmosphere C02; Natural Variations Archean to Present. Washington, D.C.: Am. Geophysical Union, pp. 602-611. Dick, H.J.B. (1982) The petrology of two back-arc basins of the northern Philippine sea. Am. J. Sci., 282, 644-700. Edmond, J., Measures, C., McDuff, E. et al. (1979) Ridge crest hydrothermal activity and the balances of major and minor elements in the ocean: The Galapagos data. Earth Planet. Sci. Lett., 46, 1-18. Elderfield, H. and Schultz, A. (1996) Midocean ridge hydrothermal fluxes and the chemical composition of the ocean. Annu. Rev. Earth Planet. Sci. Lett., 24, 191-224. Ellis, A.J. and Mahon, W.A.J. (1977) Chemistry and Geothermal Systems. New York: Academic Press. Fleisher, M. (1955) Minor elements in some sulfde minerals. Econ. Geol., g0, 970-1024. Fouquet, Y., von Stackelberg, V., Charlou, J.L., Donval, J.E, Foucher, J.E, Erzinger, J., Herzig, P., Mfihe, R., Wiedicke, M., Soakai, S. and Whitechurch, H. (1991) Hydrothermal back-arc basin; sulfides and water chemistry. Geology, 19, 303-306. Gamo, T. (1995) Wide variation of chemical characteristic of submarine hydrothermal fluids due to secondary modification process after high temperature water-rich interaction: a review. In: Sakai, H. and Nozaki, Y. (eds.), Biogeochemical Processes and Ocean Flux in the Western Pacific. Tokyo: Terra Sci. Publ., pp. 425-451. Gerlach, T.M. (1991) Present-day CO2 emission from volcanoes. EOS (Trans Am. Geophys. Union), 72, 249-251. Giggenbach, W.F. (1981) Geothermal mineral equilibria. Geochim. Cosmochim. Acta, 45, 393-410. Green, G.R., Ohmoto, H., Date, J. and Takahashi, T. (1983) Whole-rock oxygen isotope distribution in the Fukazawa-Kosaka area, Hokuroku district, Japan, and its potential application to mineral exploration. Econ. Geol. Mon., g, 395-411. Hajash, A. and Chandler, G.W. (1981) An experimental investigation of high-temperature interactions between seawater and rhyolite, andesite, basalt and peridotite. Contrib. Mineral. Petrol., 78, 240-254. Hart, S.R. and Staudigel, H. (1982) The controls of alkalines and uranium in seawater by ocean crust alteration. Earth Planet. Sci. Lett., 58, 202-213. Holland, H.D. (1978) Chemistry of Atmosphere and Ocean. New York: John Wiley and Sons. Holland, H.D. (1984) The Chemical Evolution of the Atmosphere and Oceans. Princeton, N.J.: Princeton U. Press. Holland, H.D., Horita, J. and Seyfried, W.E. Jr. (1996) On the seawater variations in the composition of Phanerozoic marine potash evaporites. Geology, 24, 993-996. Holser, W.T. and Kaplan, I.R. (1966) Isotope geochemistry of sedimentary sulfates. Chem. Geol., 1, 93-135. Huerta-Diaz and Morse, J.W. (1992) Pyritization of trace metals in anoxic marine sediments. Geochim. Cosmochim. Acta, 56, 2681-2702. Humphris, S.E. and Thompson, G. (1978) Hydrothermal alteration of oceanic basalts by seawater. Geochim. Cosmochim. Acta, 42, 107-125. Javoy, M. (1988) Carbon geodynamic cycle revised. Chem. Geol., 70, 39. Javoy, M., Pineau, F. and Allegre, C.J. (1982) Carbon geodynamic cycle. Nature, 300, 171-173. Ishibashi, J. and Urabe, T. (1995) Hydrothermal activity related to arc-back arc magmatism in the Western Pacific. In: Taylor, B. (ed.), Back-arc Basins Tectonics and Magmatism. New York: Plenum Publ., pp. 451-496.

428

Chapter 3

Kadko, D., Baross, J. and Alt, J. (1995) The magnitude and global implications of hydrothermal flux. In: Humphris, S.E., Zierenberg, R.A., Mullineauz, L.S. and Thompson, E. (eds.), Seafloor Hydrothermal Systems. Washington, D.C.: Am. Geophysical Union, pp. 446-466. Kaiho, K. and Saito, S. (1994) Oceanic crust production and climate during the last 100 Myr. Terra Nova, 6, 376-384. Kawahata, H. and Shikazono, N. (1988) Sulfur isotope and total sulfur studies of basalts and greenstones from Hole 504B, Costa Rica Rift: Implications for hydrothermal alteration. Can. Mineral., 26, 555-565. Konda, T. (1974) Bimodal volcanism in the northeast Japan arc. Geol. Soc. Japan J., 80, 81-89. Le Cloarec, M.-F. and Marty, B. (1991) Volatile fluxes from volcanoes. Terra Nova, 3, 17-27. Le Guern, E (1982) Les d6bits de C02 et de 802 volcaniques dans l'atmosphbre. Bull. Volcanol., 45(3), 197-202. Lister, C.R.B. (1973) Hydrothermal convection at seafloor spreading centers; source of power or geophysical nightmare? Geol. Soc. Am. Abst. Programs, 5, 74. Marty, B. and Jambon, A. (1987) C/3He in volatile fluxes from the solid Earth: Implications for carbon geodynamics. Earth Planet. Sci. Lett., 83, 16-26. Mottl, M.J. (1983) Metabasalts, axial hot springs and the structure of hydrothermal systems at midoceanic ridges. Geological Soc. Am. Bull., 94, 161-180. Mottl, M.J. and Holland, H.D. (1978) Chemical exchange during hydrothermal alteration of basalt by seawater. I. Experimental results for major and minor components of seawater. Geochim. Cosmochim. Acta, 42, 1103-1115. Ozawa, A. (1963) Neogene orogenesis, igneous activity and mineralization in the central part of northeast Japan. I. On the Neogene igneous activity. J. Japan. Assoc. Mineral. Petrol. Econ. Geol., 50, 167-184 (in Japanese with English abst.). Palmer, M.R. and Edmond, J.M. (1989) The strontium isotope budget of the modern ocean. Earth Planet. Sci. Lett., 92, I 1-21. Reed, M.H. (1983) Seawater-basalt reaction and the origin of greenstones and related ore deposits. Econ. Geol., 78, 466-485. Sano, Y. and Williams, S.N. (1996) Fluxes of mantle and subducted carbon along convergent plate boundaries. Geophys. Res. Lett., 23, 2749-2752. Seyfried, W.E. Jr. (1987) Experimental and theoretical constraints on hydrothermal alteration processes at midocean ridges. Ann. Re~: Earth Planet. Sci., 15, 317-335. Seward, T.M. and Kerrich, D.M. (1996) Hydrothermal emission from the Taupo volcanic zone, New Zealand. Earth Planet. Sci. Lett., 139, 105-113. Shikazono, N. (1978) Possible cation buffering in chloride rich geothermal waters. Chem. Geol., 23, 234-259. Shikazono, N. (1984) Compositional variations in epidote from geothermal areas. Geochem. J., 18, 181-187. Shikazono, N. (1985) Gangue minerals from Neogene vein-type deposits in Japan and an estimate of their CO2 fugacity. Econ. Geol., 80, 754-768. Shikazono, N. (1988) Difference in elements enriched into Kuroko and ridge hydrothermal ore deposits. Ocean Science, 20, 217-222 (in Japanese). Shikazono, N. (1993) Influence of hydrothermal flux on arsenic geochemical balance of seawater. Chikyukagaku (Geochemistry), 27, 135-139 (in Japanese). Shikazono, N. (1994) Hydrothermal alteration of green tuff belt, Japan: Implications tor the influence of seawater/w)lcanic rock interaction on the seawater chemistry at a back arc basin. The Island Arc, 3, 59-65. Shikazono, N. (1997) Geochemistry ~?["Earth System. Tokyo: U. Tokyo Press, 319 pp. (in Japanese). Shikazono, N. and Kashiwagi, H. (1999) Carbon dioxide flux due to hydrothermal venting from back-arc basin and island arc and its influence on global carbon dioxide cycle. 9th Annual V.M. Goldschmidt Conference, August 22-27, Harvard, Abstr., p. 272. Shikazono, N. and Kawahata, H. (1987) Compositional differences in chlorite from hydrothermally altered rocks and hydrothermal ore deposits. Can. Mineral., 25, 465-474. Shikazono, N., Utada, M. and Shimizu, M. (1995) Mineralogical and geochemical characteristics of hydrothermally altered basalt in Kuroko mine area, Japan: Implications for the evolution of back arc basin hydrothermal system. Applied Geochemistry, 10, 621-642.

Hydrothermal Flux from Back-Arc Basin and Island Arc

429

Simoneit, B.R.T., Philip, R.E, Jeden, ED. and Galimov, E.M. (1984) Organic geochemistry of Deep Sea Drilling Project sediments from the Gulf of California - Hydrothermal effects on unconsolidated diatomic ooze. Org. Geochem., 17, 173-205. Sugimura, A. and Uyeda, S. (1973) Island Arc. Iwanamishoten. Sugimura, A., Matsuda, T., Chinzei, K. and Nakamura, K. (1963) Quantitative distribution of late Cenozoic volcanic materials in Japan. Bull. Volcanol., 26, 125-140. Tajika, E. (1998) Climate change during the last 150 million years: reconstruction from a carbon cycle model. Earth Planet. Sci. Lett., 160, 695-707. Tajika, E. and Matsui, T. (1990) Evolution of terrestrial proto-CO2 atmosphere with thermal coupled history of the Earth. Earth Planet. Sci. Lett., 113, 251-266. Tamaki, K. and Honza, E. (1991) Global tectonics and formation of marginal basins: Role of the western Pacific. Episodes, 14, 224-230. Tamura, M. (1982) Alteration minerals and mineralization in the Shakanai Kuroko deposit, Akita Prefecture. Mining Geology, 32, 379-390. Tsunakawa, H. and Takeuchi, A. (1986) Paleo-stress field and igneous activity of Japanese Island. In: Taira, A. and Nakamura, K. (eds.), Formation of the Japanese Island. Iwanamishoten, pp. 201-208 (in Japanese). Urabe, T. (1974) Mineralogical aspects of the Kuroko deposits in Japan and their implications. Mineralium Deposita, 9, 309-324. Utter, T. (1978) Morphology and geochemistry of different pyrite types from the upper Witwatersrand system of the Klerksdorp Goldfield South Africa. Geol. Rundschau, 67(2), 774-804. Von Damm, K.L., Edmond, J.M. Measure, C.I., Walden, B. and Weiss, R.E (1985a) Chemistry of submarine hydrothermal solutions at 21~ East Pacific Rise. Geochim. Cosmochim. Acta, 49, 2197-2220. Von Damm, K.L., Edmond, J.M., Measure, C.I. and Grant, B. (1985b) Chemistry of submarine hydrothermal solutions at Guaymas Basin, Gulf of California. Geochim. Cosmochim. Acta, 49, 2221-2237. Walsh, ER., Duce, R.A. and Fashing, J.C. (1979) Considerations of the enrichment sources, and flux of arsenic in the troposphere. J. Geophys. Res., 84, 1719-1726. Weissberg, B.G., Browne, P.R.L. and Seward, T.M. (1981) Ore metals in active geothermal systems. In: Barnes, H.L. (ed.), Geochemistry of Hydrothermal Ore Deposits. New York: Wiley, pp. 738-780. White, D.E. (1967) Mercury and base-metal deposits with associated thermal and mineral waters. In: Barnes, H.L. (ed.), Geochemistry of Hydrothermal Ore Deposits. New York: Holt, Rinehart and Winston, pp. 575-631. Williams, T.J., Schaefer, S.J., Calvache, U.M.L. and Lope, Z.D. (1992) Global carbon dioxide emission to the atmosphere by volcanoes. Geochim. Cosmochim. Acta, 56, 1765-1770. Wolery, T.J. (1979) Some chemical aspects of hydrothermal processes at midoceanic ridges: A theoretical study. I. Basalt-seawater reaction and chemical cycling between the oceanic crust and oceans. II. Calculation of chemical equilibrium between aqueous solutions and minerals. Ph.D. Thesis, Northwestern U., Evanston, Illinois (unpublished). Wolery, T.J. and Sleep, N.H. (1976) Hydrothermal circulation and geochemical flux at midocean ridges. J. Geol., 84, 249-275. Wood, D.A., Jordon, J.L., Marsh, N.G., Tarney, J. and Greuil, M. (1980) Major and trace element variations drilled in basalts from the North Philippine Sea during Deep Sea Drilling Project Leg 58: A comparative study of back-arc basin basalts with lava series from Japan and midocean ridges. Deep Sea Drilling Project, Initial Reports, 58, 827-894. Yano, T. (1985) Proposal of "Pacific Problems" and its development. Chikyukagaku (Earth Science), 39, 313-318.

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431

Chapter 4

Influence of Hydrothermal CO2 Flux on Tertiary Climate Change

4.1. Tertiary climate change in relation to COz flux by volcanic, hydrothermal and metamorphic activities In Chapter 3, it was cited that hydrothermal C02 flux from back arc basins is larger than that from midoceanic ridges in the Neogene age. The hydrothermal CO2 flux is probably similar in order of magnitude to that of riverine CO2 flux (weathering flux). Therefore, it is likely that hydrothermal CO2 flux from back arc basins as well as from midoceanic ridges takes an important role for global CO2 cycle and long-term climate change. Before considering the role of CO2 by hydrothermal, volcanic and metamorphic activities in Cenozoic climate changes, paleontological, geological and geochemical studies on the Cenozoic climate are summarized below. Cenozoic climate changes have been extensively studied from a large number of fossil and geologic data (e.g., Frakes, 1979; Crowley and North, 1991; Frakes et al., 1992). The climate changes can be estimated based on 3180 of foraminiferal shell, pollen assemblage, marine molluscs, paleobotanical analysis and computer simulation of the global carbon cycle. Oxygen isotopic analysis of foraminiferal shell is perhaps the most useful technique for estimating paleotemperatures (Douglas and Savin, 1973; Douglas and Woodruff, 1981; Miller et al., 1987, 1991; Savin et al., 1975; Savin, 1977; Savin and Yeh, 1981; Shackleton and Kennett, 1975a,b). Fig. 4.1 shows oxygen isotopic paleotemperature analysis of planktonic and benthic foraminifera from the sub-Antarctic Pacific (Shackleton and Kennett, 1975a). 3180 studies on foraminifera indicate (Savin et al., 1975) that (1) following an apparently small and short-lived drop in temperature near the Tertiary-Cretaceous boundary, temperature remained warm and relatively constant through Paleocene and early and middle Eocene time; (2) a sharp temperature drop in the late Eocene time was followed by a more gradual lowering of temperature; (3) temperatures rose through early Miocene time; (4) a significant temperature drop occurred at middle Miocene age and then temperature gradually decreased during the middle-late Miocene age. This tendency is generally consistent with paleontological studies on pollen (Yamanoi, 1993) (Fig. 4.2), and marine molluscs (Ogasawara, 1994) analyses, paleotemperatures derived from the 3180 values of coexisting carbonate and phosphate in marine sediments (Zheng, 1996), and Mg/Ca in benthic foraminifera calcite (Lear et al., 2000) (Fig. 4.3). In the previous studies it was considered that the climate change has been reflected by the changes in ocean circulation pattern, ice volume, albedo, and weathering. Many

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review papers on the relationship between these changes and Neogene climate change have been published (Tada, 1991; Kawahata, 1991 ). It has been pointed out that the decline in global temperatures from middle Miocene to Quaternary may have been related to the increased rates of silica weathering in the Himalayas, development of ice sheets, and corresponding increase in albedo,

Influence of Hydrothermal C02 Flux on Tertiary Climate Change

433

or change in the ocean circulation pattern. However, a few systematic studies on the change in CO2 flux from the earth's interior by volcanism, hydrothermal activity and metamorphism have been done. Budyko et al. (1987) pointed out that Eocene and Miocene volcanisms and related CO2 fluxes have been large. Kennett et al. (1977) showed the histograms of middle to late Cenozoic K-Ar dates with relative volumetric estimates of igneous rocks erupted in the circum-Pacific region (Fig. 4.4). The rates of basaltic, andesitic and rhyolitic eruptions at middle Miocene (16-15 Ma) were very high, but declined between 15 and 10 Ma, and since 10 Ma it increased. For example, at the middle Miocene age, huge amounts of Columbia basalt (Yellowstone mantle plume) erupted (Camp, 1995). It was estimated that there was an increase of 3~ by CO2 from this eruption (Coffin and Eldholm, 1993). In the Japanese Islands, Ti-rich Icelandite-like and Mg-rich basaltic activities occurred at middle Miocene (Dud~s et al., 1983; Shuto, 1989). This type of basaltic magma is thought to have been generated at the deep mantle, probably related to mantle plume activity (Tatsumi et al., 1989). Kaiho and Saito (1994) calculated oceanic crustal production rate (midoceanic ridges, oceanic plateaus, and back arc basins) during the last 100 m.y. and indicated that the oceanic crustal production rate correlates well to climate change (see Fig. 3.5). The oceanic production rate at back arc basins estimated by Kaiho and Saito (1994) seems to correlate to Neogene climate change. For example, rapid cooling events at the Eocene/Oligocene boundary and middle Miocene are due to decreasing of oceanic production rate of back arc basins (Fig. 3.5). Total crustal production rate in back arc basins during 15-8 Ma was very low, but in Oligocene and early to middle Miocene was high. However, the rate in Pliocene and Quaternary was high, but temperature during that time was probably low compared with Oligocene to middle Miocene (Fig. 4.1). Kaiho and Saito (1994) thought that the lower temperature in the Pliocene-Quaternary was caused by the combined effect of weathering at the Himalaya, increase in albedo by ice sheet development, and ocean circulation pattern (development of antarctic circumpolar current). During the Eocene the production rate of back arc basins was low (see Fig. 3.5), but the global average atmospheric temperature was high. However, the ocean crustal production rate in midoceanic ridges was high. Therefore, it is reasonable to infer that during the Eocene, hydrothermal activity at midoceanic ridges affected the global climate (Owen and Rea, 1985), and that the CO2 flux from both midoceanic ridges and back arc basins has played an important role in climate change in the Neogene age. In addition to affecting spreading rate and intensity of volcanism, the changes in the configuration of ridge-transform plate boundaries are important for controlling seafloor hydrothermal activity (Owen and Rea, 1985). A significant reorganization in plate motion occurred at this time and may have increased CO2 through enhanced volcanism. For example, the Norwegian-Greenland sea opened, and volcanoes were extruded throughout much of the northern North Atlantic (the Thulen basalts; Roberts et al., 1984). Roberts et al. (1984) estimated that the total volume of the Thulean basalts is comparable to late Cretaceous Deccan basalts in India. Climate change in the vicinity of the Japanese Islands during the Neogene has been

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studied considerably via 3180 of foraminifera, biobotanical analyses, marine molluscs, and pollen assemblage. Pollen assemblage analysis indicated that temperature at middle Miocene age was high and after that it decreased rapidly (Fig. 4.2) (Yamanoi, 1993). Temperature from the middle Miocene to present time decreased gradually, however; it increased from 4 to 3 Ma and then decreased (Fig. 4.2) (Yamanoi, 1993).

Fig. 4.3. (A) Composite multispecies benthic foraminiferal Mg/Ca records from three deep-sea sites: DSDP Site 573, ODP Site 926, and ODP Site 689. (B) Species-adjusted Mg/Ca data. Error bars represent standard deviations of the means where more than one species was present in a sample. The smoothed curve through the data represents a 15% weighted average. (C) Mg temperature record obtained by applying a Mg calibration to the record in (B). Broken line indicates temperatures calculated from the 3180 record assuming an icefree world. Blue areas indicate periods of substantial ice-sheet growth determined from the 3180 record in conjunction with the Mg temperature. (D) Cenozoic composite benthic foraminiferal 3180 record based on Atlantic cores and normalized to Cibicidoides spp. Vertical dashed line indicates probable existence of ice sheets as estimated by (2). ~w, seawater 3180. (E) Estimated variation in 3180 composition of seawater, a measure of global ice volume, calculated by substituting Mg temperatures and benthic 3180 data into the 3180 paleotemperature equation (Lear et al., 2000).

436

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The middle Miocene climatic optimum at about 16 Ma is well established by the studies on shallow marine molluscs (Ogasawara, 1994). Based on the data on Neogene shallow marine molluscs of the Japanese Islands, the paleoclimate was reconstructed by Ogasawara (1994). After the middle Miocene climatic optimum at 16 Ma, a gradual cooling commenced at about 14-13 Ma. The zoogeographic history of the Japan Sea proposed by Chinzei (1991) (Fig. 4.5 and Fig. 4.6) indicated that: (1) from 17 to 15 Ma, before the spreading of the Japan Sea, tropical and subtropical molluscs invaded the area; (2) the benthic and planktonic faunas changed sharply into low temperature faunas at about 15 Ma; (3) from about 10 to 5 Ma, sedimentary deposits formed that were barren of benthic molluscs as a result of the stagnation of the Japan Sea basin; (4) cold-water molluscan fauna reappeared in the coastal areas at about 5 Ma; and (5) this has been followed by a period of marked cyclicity of cold and warm water faunas since about 1.3 Ma. The change in temperature near the Japanese Islands mentioned above seems to be consistent with the global temperature change (Ogasawara, 1994), as well as with changes

Influence of Hydrothermal C02 Flux on Tertiary Climate Change

437

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in hydrothermal activity mostly from 15 Ma to the present inferred from the formation ages of Neogene hydrothermal ore deposits in Japan (Fig. 4.7). During the middle Miocene, Kuroko deposits, polymetallic vein-type deposits, gold-quartz vein-type deposits and Sb and Hg vein-type deposits were formed (see sections 1.3 and 1.6). Many vein-type deposits were formed not only in and nearby the Japanese Islands, but also at middle Miocene in northwest USA (Basin and Range; Lipman, 1982), and elsewhere in the circum-Pacific regions (e.g., Peru). It is probable that large amounts of CO2 effused into the atmosphere from hydrothermal solution associated with this widespread mineralization and volcanic gas from subduction zones, causing an increase in temperature. During middle to late Miocene (15-6 Ma), very little vein-type (Fig. 4.7) mineralization occurred in the Japanese Islands. During 5 Ma to present many vein-type deposits formed.

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This metallogenic epoch in and near the Japanese Islands correlates well with Jackson's episode (Masuda, 1984). Jackson's episode means the change in the direction of Pacific plate, which was estimated from the array and age of Hawaiian islands by Jackson et al. (1975) (Fig. 4.8).

Influence of Hydrothermal C02 Flux on Tertiary Climate Change

439

In addition to hydrothermal and volcanic activity, metamorphism may have influenced the CO2 levels of the atmosphere and caused climate changes. Based on a model of the Cenozoic extension in the North American Cordillera, Nesbitt et al. (1995) demonstrated that CO2 generation associated with crustal extension may have been a major contributor to the elevated CO2 levels of the Cenozoic atmosphere and the resulting global warming due to the CO2 greenhouse effect. Kerrick and Caldeira (1993) estimated the amount of metamorphic CO2 produced at depth in the Himalayan orogeny with data on the timing of the India/Asia collision, duration of regional metamorphism, and proportions and bulk compositions of metamorphic CO2 source rocks. They have shown that the extensive metamorphism associated with the India/Asia collision and the closing of the Tethys have contributed to CO2 greenhouse warming in early to middle Cenozoic. Their estimated CO2 metamorphic flux is 1018-1019 mol/m.y, which is high enough to warm the climate.

4.2. Computation on global long-term carbon cycle and climate change In the last two decades extensive studies on the computation of the long-term global carbon cycle have been carried out (Garrels and Lerman, 1983; Berner et al., 1983; Lasaga et al., 1985; Berner, 1987, 1990, 1991, 1994; Volk, 1987; Caldeira and Rampino, 1991; Berner and Rye, 1992; Tajika and Matsui, 1992, 1993; Godderis and Francois, 1995; Ishikawa, 1996; Tajika, 1992, 1998; Kashiwagi et al., 2000; Shikazono and Kashiwagi, 1999). These studies demonstrated that the numerical modeling provides a useful tool to estimate long-term CO2 change in atmosphere and climate change. Most of these works studied long-term (ca. 100 m.y.) changes. However, short-term carbon cycle and climate change (i.e., during Neogene) have not been quantitatively well studied. Recently, Ishikawa (1996), and Kashiwagi et al. (2000) calculated CO2 change in atmosphere during the last 30 Ma and the last 60 Ma respectively, mainly based on the method developed by Berner et al. (BLAG model) (1983), Berner (GEOCARB model) (1994), and Tajika (1998) taking into account hydrothermal CO2 flux from back arc basins and island arcs which were not considered in previous studies. In this section, Ishikawa (1996), and Kashiwagi et al. (2000) studies are described. Then, their calculated results are given, emphasizing the influence of hydrothermal and volcanic gas CO2 flux from back arc basins and island arc on CO2 concentration of atmosphere and climate change and are compared with the changes in CO2 and temperature obtained by analytical and paleontological data (3180 of foraminiferal shell, Ce anomaly, 313C, etc.). Ishikawa (1996) performed the calculations using a modified BLAG model (Berner et al., 1983) based on a simple box model which has five reservoirs (atmosphere, ocean, continental crust, oceanic crust and mantle) (Fig. 4.9) and estimated carbon flux during the last 30 Ma between one reservoir and the other, giving the values of external parameters such as continental surface area, seafloor spreading rate, back arc basin spreading rate, surface temperature, degassing rate from island arcs, back arc basins, and midoceanic ridges.

Chapter 4

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0

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,

,

5

10

|

!

|

15

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Age (Ma) Fig. 4.10. Atmospheric CO2 variation estimated by modified BLAG model including CO2 fhlx related to mantle plume activity (Ishikawa, 1996). Rco2 = Pr (P~o2" present-day Pr The calculated result assuming that degassing of C02 from the back arc basin is from the subducting plate and degassing from the mantle is given in Fig. 4.10. The calculation clearly indicates that the CO2 concentration in the atmosphere at the middle Miocene age was high and temperature decreased from the middle Miocene age. Therefore, these calculations indicate that CO2 degassing from back arc basins is important for controlling Neogene climate change. The temperature decreased from middle Miocene (15-17 Ma) to 10 Ma is 3-4~ Kashiwagi et al. (2000) calculated the change in atmospheric CO2 level during the last 70 Ma mainly based on the GEOCARB II model by Berner (1994) and the modified GEOCARB model by Tajika (1998) which consist of weathering factors, degassing factors, biogenic factors and solar factors. Weathering factors relate silicate weathering, carbonate weathering, weathering due to Himalayan uplift, carbonate land area and global runoff. Degassing factors are spreading and subduction rate. Kashiwagi et al. (2000)

441

Influence of Hydrothermal C02 Flux on Tertiary Climate Change

I

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.

.

.

=0.2

.

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i

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Age (my) Fig. 4.11. Atmospheric C02 variation estimated by modified GEOCARB II model including volcanic eruption rate of circum-Pacific region by Kennett et al. (1977) (Kashiwagi et al., 2000). y represents the contribution of the flux from back arc basin to that from subduction zones at present. Rc02 = Pr (P~o2: present-day PC02).

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Age (my) Fig. 4.12. Atmospheric C02 variation estimated by modified BLAG model including C02 flUX related to mantle plume activity (Kashiwagi et al., 2000). u represents the contribution of the flux from back arc basin to that from subduction zones at present. Rco2 = PCO2/P~O2 (P~O2: present-day Pco2).

modified this G E O C A R B II by adding the degassing from back arc basins and island arcs. The correction factors for back arc basins and island arcs considered by them were taken from Kaiho and Saito (1994) and Kennett et al. (1977), respectively. The computation results by Kashiwagi et al. (2000) are shown in Fig. 4.11 and Fig. 4.12. This model calculation cannot explain the decrease of temperature since the middle Miocene (ca. 14 Ma) to present. Kashiwagi et al. (2000) calculated the effect of albedo by ice sheets to explain this inconsistency. However, the albedo effect is less than 1 ~ suggesting that this effect is not important. The possible causes for the cooling from the middle Miocene to the present are: (1) the change in carbonate deposition with time; (2) the change in albedo due to topographic change (building of mountains such as the Himalaya, etc.); (3) the change

442

Chapter 4

in ocean circulation patterns such as destruction of the stratified ocean; (4) the change in CO2 exchange between the atmosphere and surface of the ocean; (5) environmental changes in the ocean such as ocean current and marine bioproductivity; (6) the change in terrestrial vegetation; and (7) the cooling effect by volcanic gas (SO2) and volcanic ash. It is uncertain which factor is important for causing cooling from the middle Miocene to present. Further work on the causes for cooling are necessary. Liu and Schmitt (1984, 1993a,b, 1996) and Liu et al. (1988) derived the relationship between the negative Ce anomaly of seawater and Pco2 from thermochemical calculations. Liu and his coworkers estimated Pco2 variation during the last 120 Ma based on this method. Pco2 at middle Miocene age estimated by Liu and his coworkers is high and seems to be consistent with that by Ishikawa (1996). However, Liu's method was criticized by Elderfield and Schultz (1987). Ce anomaly depends not only on Pco2, but on many other factors such as Po2, pH and ECe. Pco2 values were estimated by 313C method by Cerling (1984, 1991, 1992a,b) who used 813C of carbonates in terrestrial soil as an indicator of Pco2. However, these data on Cenozoic age are scarce and scattered. Pagani et al. (1999)estimated Pco2 during the Cenozoic based on 313C of organic matter in marine sediments. They used alkanone-based Pco2 estimates and indicated that Pco2 increased from 14 Ma to 9 Ma and stabilized at pre-industrial values by 9 Ma (Fig. 4.13 and Fig. 4.14). Their result suggests that the C4 plant expansion was likely driven by a tectonically-related episode of enhanced low-latitude aridity or changes in seasonal precipitation patterns on a global scale (or both). Pearson and Palmer (2000) used the boron isotope ratios of ancient planktonic foraminiferal shells to estimate the pH of surface layer seawater throughout the past 60 m.y. and to reconstruct atmospheric CO2

.2

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Time (my) Fig. 4.13. PCO 2 estimates from Oligocene to late Miocene (modified after Pagani et al., 1999).

443

Influence of Hydrothermal C02 Flux on Tertiary Climate Change

813C (%)

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I I Fig. 4.14. Tracing climate change in the Miocene. Shown here are records of ice volume and temperature (based on foraminiferal 3180) and relative organic carbon burial (based on foraminiferal 313C), compared with the CO2 estimates of Pagani et al. (1999), and tectonic events that may have affected ocean heat transport. Trends in CO2 are consistent with organic carbon burial and CO2 drawdown during the Monterey Excursion, but cannot explain the Miocene Climatic Optimum (MCO) or expansion of the East Antarctic Ice Sheet (EAIS).

concentrations. They found that since the earliest Miocene the system has been constant and closely comparable to the present (Fig. 4.15). The studies by Pagani et al. (1999), Pearson and Palmer (2000) and Kashiwagi et al. (2000) suggest that Pco2 did not decrease from the Miocene to the present time. This view is not in agreement with that was favored by previous works such as Raymo's hypothesis (Raymo et al., 1988) and Monterey's hypothesis (Vincent and Berger, 1985). If their estimates are correct, it is inferred that Pco2 is not an important factor affecting cooling from the middle Miocene to present. At present, the main cause for cooling since the middle Miocene is not elucidated. To solve this problem, we need to make computer calculations on the long-term global carbon cycle including the effect of terrestrial biota, ocean circulation pattern, and metamorphic activities which are not included in Kashiwagi et al. (2000)'s computation.

444

Chapter 4

(a) elPlil E

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Fig. 4.15. Atmospheric CO2 data for the past 60 m.y. (a) The entire record, (b) an enlargement of the past 25 m.y. (Pearson and Palmer, 2000).

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Berner, R.A. (1990) Atmosphere carbon dioxide levels over Phanerozoic time. Science, 249, 1382-1386. Berner, R.A. (1991) A model for atmospheric CO2 over Phanerozoic time. Am. J. Sci., 291, 339-376. Berner, R.A. (1994) GEOCARB II: a revised model of atmospheric CO2 over Phanerozoic time. Am. J. Sci., 294, 56-91. Berner, R.A. and Rye, D.M. (1992) Calculation of the Phanerozoic strontium isotope record of the oceans from a carbon cycle model. Am. J. Sci., 292, 136-148. Berner, R.A., Lasaga, A.C. and Garrels, R.M. (1983) The carbonate-silicate geochemical cycle and its effects on atmospheric carbon dioxide over the past 100 million years. Am. J. Sci., 283, 641-683. Budyko, M.I., Ronov, A.B. and Yanshin, A.L. (1987) History of the Earth's Atmosphere. Berlin: SpringerVerlag, 139 pp. Caldeira, K. and Rampino, M.R. (1991) The mid-Cretaceous super-plume, carbon dioxide and global warming. Geophys. Res. Lett., 18, 987-990. Camp, V.E. (1995) Mid-Miocene propagation of the Yellowstone mantle plume head beneath the Columbia River basalt source region. Geology, 23, 435-438. Cerling, T.E. (1984) The stable isotopic composition of modern soil carbonate and its relationship to climate. Earth Planet. Sci. Lett., 71, 229-240. Cerling, T.E. (1991) Carbon dioxide in the atmosphere; evidence from Cenozoic and Mesozoic paleosols. Am. J. Sci., 291, 377-400. Cerling, T.E. (1992a) Further comments on using carbon isotopes in paleosols to estimate the CO2 content of the paleo-atmosphere. J. Geol. Soc. London, 149, 673-676. Cerling, T.E. (1992b) Use of carbon isotopes in paleosols an indicator of the P(CO2) of the paleo-atmosphere. Global Biogeochemical Cycles, 6, 307-314. Chameides, W.L. and Perdue, E.M. (1997) Biogeochemical Cycles. Oxford University Press, 224 pp. Chinzei, K. (1986) Faunal succession and geographic distribution of Neogene molluscan faunas in Japan. In: Kotaka, T. (ed.), Japanese Cenozoic Molluscs Their Origin and Migration. Palaeont. Soc. Japan, Special Paper, 29, 17-32. Chinzei, K. (1991) Late Cenozoic zoogeography of the sea of Japan area. Episodes, 14, 231-235. Coffin, M.E and Eldholm, O. (1993) Large igneous provinces. Scientific American, October 1993, 42-49. Crowley, T.J. and North, G.R. (1991) Paleoclimatology. New York: Oxford University Press, 339 pp. Douglas, G. and Savin, S.M. (1973) Oxygen and carbon isotope analyses of Cretaceous and Tertiary foraminifera from the central North Pacific. Washington, D.C.: U.S. Govt. Printing Office, Initial Reports of the Deep Sea Drilling Project, 17, 591-605. Douglas, R.G. and Woodruff, F. (1981) Deep sea benthic foraminifera. In: Emiliani, C. (ed.), The Sea, 7. New York: Wiley, pp. 1233-1327. Dudfis, F.0., Campbell, I.H. and Gordon, M.P. (1983) Geochemistry of igneous rocks in the Hokuroku district, northern Japan. Econ. Geol. Mon., 5, 115-133. Elderfield, H. and Schultz, A. (1996) Midocean ridge hydrothermal fluxes and the chemical composition of the ocean. Annu. Rev. Earth Planet. Sci. Lett., 24, 191-224. Frakes, L.A. (1979) Climates Throughout Geologic Time. Amsterdam: Elsevier, 310 pp. Frakes, L.A., Francis, J.E. and Syktus, J.I. (1992) Climate Modes of the Phanerozoic. Cambridge University Press, 274 pp. Garrels, R.M. and Lerman, A. (1983) Coupling of sedimentary sulfur and carbon cycles an improved model. Am. J. Sci., 284, 989-1007. Godderis, Y. and Francois, L.M. (1995) The Cenozoic evolution of the strontium and carbon cycles: relative importance of continental erosion and mantle exchange. Chem. Geol., 126, 169-190. Ishikawa, Y. (1996) Global C02 cycle and climate change during 30 Ma-present. Masters thesis, Keio University (in Japanese). Jackson, E.D., Show, H.R. and Bargar, K.E. (1975) Calculated geochronology and stress field orientations along the Hawaiian chain. Earth Planet. Sci. Lett., 26, 145-155. Kaiho, K. and Saito, S. (1994) Oceanic crust production and climate during the last 100 Mys. Terra Nova, 6, 376-384. Karig, D.E. (1971) Origin and development of marginal basins in the western Pacific. J. Geophys. Res., 84, 6796-6802.

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Kashiwagi, H., Shikazono, N. and Tajika, E. (2000) Global carbon cycle model in the Cenozoic. 10th Annual V.M. Goldschmidt Conference, September 3-8, Oxford University; Abst. Kawahata, H. (1991) Oceanic and mass circulation. J. Geography, 100, 1007-1018 (in Japanese). Kennett, J.P., McBirney, A.R. and Thunell, R.C. (1977) Episodes of Cenozoic volcanism in the circum-Pacific region. J. Volc. Geotherm. Res., 2, 145-163. Kerrick, D.M. and Caldeira, K. (1993) Paleoatmospheric consequences of CO2 released during early Cenozoic regional metamorphism in the Tethyan orogen. Chem. Geol., 108, 205-230. Lasaga, A.C., Berner, R.A. and Garrels, R.M. (1985) An improved geochemical model of atmospheric CO2 fluctuations over the past 100 million years. In: Sundquist, E.T. and Broecker, W.S. (eds.), The Carbon Cycle and Atmospheric C02: Natural Variations Archean to Present. Washington, D.C.: Am. Geophysical Union, pp. 397-411. Lear, C.H., Elderfield, H. and Wilson, P.A. (2000) Cenozoic deep-sea temperatures and global ice volumes from Mg/Ca in benthic foraminiferal calcite. Science, 287, 269-272. Lipman, P.W. (1982) Volcano-tectonic setting of Tertiary ore deposits, southern Rocky mountains. Mining Geology, 32, 1-24. Liu, Y.-G. and Schmitt, R.A. (1984) Chemical profiles in sediment and basalt samples from Deep Sea Drilling Project Leg 74, Hole 525A (Moore, T.G. Jr., et al., eds.), 74, 713-730. Liu Y.-G. and Schmitt, R.A. (1993a) Earth's partial pressure of CO2 over the past 120 Ma: evidence from Ce anomalies in the deep (>60 m) Pacific ocean. Lunar Planet. Sci., XX IV, 883-884. Liu Y.-G. and Schmitt, R.A. (1993b) Earth's partial pressure of CO2 over the past 100-500 Ma: evidence from Ce anomalies to mostly shallow seas (<200 m) as recorded in carbonate sediments. Lunar Planet. Sci., XXIV, 885-886. Liu, Y.-G. and Schmitt, R.A. (1996) Cretaceous Tertiary phenomena in the context of seafloor rearrangements and P(CO2) fluctuations over the past 100 m.y. Geochim. Cosmochim. Acta, 6, 973-999. Liu, Y.-G., Miah, M.R.V. and Schmitt, R.A. (1988) Cerium: a chemical tracer for paleooceanic redox conditions. Geochim. Cosmochim. Acta, 52, 1361-1371. Masuda, F. (1984) Sedimentary basins in arc-trench system as a high sensitive recorder of oceanic plate motion. Mining Geology, 34, 1-20 (in Japanese). Miller, K.G., Fairbanks, R.G. and Mountain, G.S. (1987) Tertiary oxygen isotope synthesis, sea level history, and continental margin erosion. Paleoceanography, 2, 1-19. Miller, K.G., Feigenson, M.D. and Wright, J.D. (1991) Miocene isotope standard section, DSDP Site 608: An evaluation of isotope and biostratigraphic resolution. Paleoceanography, 6, 33-52. Nesbitt, B.E., Mendoza, C.A. and Kerrick, D.M. (1995) Surface fluid convection during Cordillera extension and the generation of metamorphic CO2 contributions to Cenozoic atmospheres. Geology, 23, 99-101. Ogasawara, K. (1994) Neogene paleogeography and marine climate of the Japanese Islands based on shallowmarine molluscs. Paleogeogr. Paleoclimatol. Paleoecol., 108, 335-351. Owen, R.N. and Rea, D.K. (1985) Seafloor hydrothermal links climate to tectonics: Eocene carbon dioxide greenhouse. Science, 227, ! 66-169. Ozawa, K., Kawahata, H. and Nakanishi, M. (1990) Formation of ocean floor and spreading rate. Kagaku (Science), 60, 661-669 (in Japanese). Pagani, M., Freeman, K.H. and Arthur, M.K. (1999) Late Miocene atmospheric CO2 concentrations and the expansion of C4 grasses. Science, 285, 876-879. Pearson, P.N. and Palmer, M.R. (2000) Atmospheric carbon dioxide concentrations over the past 60 million years. Nature, 406, 695-699. Pitman, W.C. (1978) Relationship between eustacy and stratigraphic sequences of passive margins. Geol. Soc. Am. Bull., 89, 1289-1403. Raymo, M.E., Ruddiman, W.F. and Froelich, P.N. (1988) Influence of late Cenozoic mountain building on ocean geochemical cycles. Geology, 16, 649-653. Roberts, D.G., Mortin, A.C. and Backmann, J. (1984) Late Paleocene-Eocene volcanic events in the northern North Atlantic ocean. Washington, D.C.: U.S. Gov. Printing Office. Initial Report Deep Sea Drilling Project, 81, 913-923. Savin, S.M. (1977) The history of the Earth's surface temperature during the past 100 million years. Annu. Rev. Earth Planet. Sci., 5, 319-355.

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Savin, S.M. and Yeh, H.-W. (1981) Stable isotopes in ocean sediment. In: Emiliani, C. (ed.), The Sea, 7. New York: Wiley, pp. 1521-1554. Savin, S.M., Douglas, R.G. and Stehli, EG. (1975) Tertiary marine paleotemperatures. Geol. Soc. Am. Bull., 86, 1499-1510. Shackleton, N.J. (1986)Paleogene stable isotope events. Paleogeogr. Paleoclimatol. Paleoecol., 57, 91-102. Shackleton, N.J. and Kennett, J.E (1975a) Paleotemperature history of the Cenozoic and the initiation of Antarctic glaciation: oxygen and carbon isotope analyses in DSDP Sites 277, 279, and 281. Initial Reports of the DSDP, 29, Washington, D.C.: U.S. Gov. Printing Office, pp. 743-755. Shackleton, N.J. and Kennett, J.E (1975b) Late Cenozoic oxygen and carbon isotopic ranges at DSDP Site 284: Implications for glacial history of the northern hemisphere and Antarctica. Initial Report DSDP, 29, Washington, D.C.: U.S. Govt. Printing Office, pp. 801-807. Shikazono, N. (1988) Isotopic composition of sulfide sulfur and origin of Neogene Au-Ag and base metal vein-type deposits in Japan. J. Fac. Sci. Univ. Tokyo, Sect. II, 239-255. Shikazono, N. and Kashiwagi, H. (1999) Carbon dioxide flux due to hydrothermal venting from back-arc basin and island arc and its influence on the global carbon dioxide cycle. 9th Annual V.M. Goldschmidt Conference, August 22-27, Harvard University, Abst., p. 272. Shikazono, N., Utada, M. and Shimizu, M. (1995) Mineralogical geochemical characteristics of hydrothermally altered basalt in Kuroko mine area. Applied Geochemistry, 10, 621-642. Shuto, K. (1989) Tertiary volcanism of the Northeast Japan arc in view of hypothesis of the spreading of the Japan sea. Earth Science (Japan), 43, 28-42 (in Japanese). Tada, R. (1991) Earth's surface environmental change in Cenozoic. J. Geography, 100, 937-950 (in Japanese). Tai, Y. (1963) Foraminifera in Tertiary age from Setouchi-Sannin and foraminiferal sharpline. Fossil, 5, 1-7 (in Japanese). Tajika, E. (1992) Evolution of the atmosphere and ocean of the Earth; global geochemical cycles of C,H,O,N, and S, and degassing history coupled with thermal history. Doctoral Thesis, University Tokyo, 416 PP. Tajika, E. (1998) Climate change during the last 150 million years: reconstruction from a carbon cycle model. Earth Planet. Sci. Lett., 160, 695-707. Tajika, E. and Matsui, T. (1992) Evolution of terrestrial proto-CO2 atmosphere with thermal coupled history of the Earth. Earth Planet. Sci. Lett., 113, 251-266. Tajika, E. and Matsui, T. (1993) Degassing history and carbon cycle: From an impact-induced steam atmosphere to the present atmosphere. Lithos, 30, 267-280. Takahashi, M. (1989) Volcanic activity in Fossa Magna region. Monthly Earth, 11, 544-551 (in Japanese). Takeuchi, A. (1987) On the episodic vicissitude of tectonic stress field of the Cenozoic Northeast Honshu arc, Japan. In: Nasu, N. et al. (eds.), Formation of Active Ocean Margins. Dordrecht, the Netherlands: Reidel, pp. 443-468. Tatsumi, Y., Otofuji, Y., Matsuda, T. and Nohda, S. (1989) Opening of the Japan sea by asthenospheric injection. Tectonophysics, 166, 317-329. Uyeda, S. and Kanamori, H. (1979) Back-arc opening and the mode of subduction. Geophys. Res., 84, 1049-1061. Vincent, E. and Berger, W.H. (1985) Carbon dioxide and polar cooling in the Miocene: The Monterey Hypothesis. In: Sundquist, E.T. and Broecker, W.S. (eds.), The Carbon Cycle and Atmospheric C02: Natural Variations Archean to Present. Washington D.C.: Am. Geophysical Union, pp. 455-468. Volk, T. (1987) Feedbacks between weathering and atmospheric CO2 over the last 100 million years. Am. J. Sci., 287, 763-779. Volk, T. (1989) Rise of angiosperms as a factor in long-term climatic cooling. Geology, 17, 107-110. Walker, J.C.G. and Hays, EB. (1981) A negative feedback mechanism for the long-term stabilization of earth's surface temperature. J. Geophys. Res., 86(C19), 9776-9782. Yamanoi, T. (1993) Vegetation in Neogene age in Japanese Islands. Monthly Earth, 16, 181-190 (in Japanese). Zheng, Y.E (1996) Oxygen isotope fractionations involving apatites: Application to paleotemperature determination. Chem. Geol., 127, 177-187.

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449

Chapter 5

Summary

In and nearby the Japanese Islands, which are situated close to plate boundaries (Pacific, Philippine, North American and Asian plates), many hydrothermal ore deposits were formed during the Miocene-present eras. Major deposit types are Kuroko and epithermal vein-types. Epithermal vein-type deposits are classified into precious (Au, Ag)-types and base-metal (Cu, Pb, Zn, Fe, Mn, Ag)-types. Geological, mineralogical and geochemical features of these deposit types (distribution, age, associated volcanism, host and country rocks, fluid inclusions, opaque, gangue and hydrothermal alteration minerals, chemical features of ore fluids (temperature, salinity, pH, chemical composition, gaseous fugacity, isotopic compositions (O, D, S, 87Sr/S6Sr' Pb), rare earth elements)) were summarized. Based on these summaries, the formation mechanism of ore deposits and origin of ore fluids were considered. Mixing of ascending hydrothermal solution and ambient cold water (seawater, groundwater) is considered to be an important depositional mechanism. Chemical and isotopic compositions of Kuroko ore fluids are reasonably explained by the seawater-rock interaction at elevated temperatures (200-350~ in submarine hydrothermal system at back-arc basins. The ore fluids responsible for epithermal base-metal vein-type deposits were generated predominantly by meteoric water-rock interaction at elevated temperatures (200-350~ Fossil seawater in marine sediments was also involved in the ore fluids responsible for this type of deposits. Epithermal precious metal ore fluids were generated by meteoric water-rock interaction at 150-250~ Small amounts of seawater sulfate were involved in the ore fluids responsible for epithermal precious metal vein-type deposits occurring in Green tuff region (submarine volcanic and sedimentary rocks). The difference in the kinds of metals enriched in Kuroko, base metal vein-type and precious metal vein-type deposits could be explained in terms of the HSAB (hard, soft, acids and bases) principle (Pearson, 1963). According to this principle, relatively hard cations (base metal (Cu, Pb, Zn, Fe, Mn, Ag) ions) tend to combine preferentially with chloride ion in hydrothermal solution, while soft cations (Au, Ag, T1, Hg ions etc.) combine with HzS and HS-. The differences in salinity of ore fluids in base-metal-rich deposits (base metal vein-type deposits and Kuroko deposits) and base-metal-poor deposits (precious metal vein-type deposits) is also in accordance with the HSAB principle. The K-Ar age and foraminiferal age of epithermal vein-type and Kuroko deposits clearly demonstrate that epithermal precious and base metal vein-type deposits formed

450

Chapter 5

at 15 Ma-present and Kuroko deposits formed at 15-12 Ma. Epithermal base metal veins formed generally before epithermal precious metal vein-type deposits and are younger than Kuroko deposits in the same metallogenic district. Paleogeography at the time of formation of Kuroko and epithermal vein-type deposits and country rocks of the metallogenic province indicate that Kuroko deposits occur at the centre of the Green tuff region and were formed in the submarine environment, while epithermal veins were formed in the subaerial environment. Kuroko and epithermal base metal vein-type deposits occur in the Green tuff region, which is thick submarine volcanic pile, and epithermal precious-metal vein-type deposits formed associated with subaerial igneous activities at the margin of the Green tuff region. Spatial and temporal distributions of these deposits were reflected by the change in tectonics near the Japanese Islands, which was intimately related to the mode of subduction of the Pacific Plate and Jackson's episode since the middle Miocene to present. Mariana-type and Chilean-type subductions according to Uyeda and Kanamori (1979) relate to the formation of Kuroko and epithermal vein-type deposits, respectively. Mariana-type subduction characterized by the subduction with steep angle causes extensional stress field, subsidence of land, formation of back-arc basins, enlargement of submarine areas, bimodal (basaltic and acidic) volcanism, and submarine hydrothermal systems. This geologic situation is favorable for the formation of hydrothermal ore deposits (Kuroko deposits) on the seafloor and/or sub-seafloor environment. Chilean-type subduction characterized by the subduction with gentle angle causes compressional stress field, uplift, calc-alkaline volcanism, enlargement of land, and subaerial hydrothermal system. Epithermal vein-type deposits were formed in relation to the geologic activities caused by the subduction similar to Chilean-type and oblique subduction. Epithermal precious-metal deposits are divided into Se-type (low sulfidation type), Te-type and hot spring type (high sulfidation type). Te-type and hot spring type deposits formed under compressional environments related to Chilean-type subduction, while Se-types formed under extensional environment. It is suggested that the mode of subduction of the Pacific Plate since the middle Miocene age related to Jackson's episode, hence oscillation of direction of lateral movement of Pacific plate. Synchronized igneous and hydrothermal activities and Jackson's episode indicate that the formations and characteristics of hydrothermal ore deposits (Kuroko and epithermal vein-type deposits) are largely influenced by plate tectonics (mode of subduction, direction of plate movement, etc.). For example, sulfur isotopic composition of sulfides is not controlled by /O2 and pH, but by E~34S of ore fluids. This means that the sulfur isotopic compositions of sulfides are reflected by origin of sulfide sulfur and global tectonic environment rather than the physicochemical environment at the site of ore deposition. The other important deposit types of the Neogene age are polymetallic (Cu, Pb, Zn, Sn, W, Au, Ag) (xenothermal or subvolcanic type deposits) vein-type deposits, hypo/mesothermal Au vein-type deposits, disseminated-type (hot spring type) Au deposits, and Hg-Sb disseminated and vein-type deposits. Characteristic features of these deposit types also indicate that geologic and tectonic environments affect the origin of ore fluids and formations of these deposits.

Summary

451

For instance, polymetallic vein-type deposits formed under the subaerial environment influenced by igneous and sedimentary components. The ore fluids responsible for Hg-Sb deposits are largely controlled by the interaction of hydrothermal solution with sedimentary rocks. However, further works on these deposits are clearly necessary. During the last three decades, many hydrothermal deposits have been discovered at midoceanic ridges, back-arc basins and subaerial active geothermal systems. Characteristic features of back-arc deposits at the western Pacific region (e.g., Okinawa Trough, Izu Ogasawara, North Fiji and Mariana deposits) are very similar to those of Kuroko deposits. The characteristic features of most Besshi-type deposits in Japan are similar to those of midoceanic ridge deposits but some (Hitachi, Yanahara) are similar to Kuroko deposits. These deposits are characterized by polymetallic (Cu, Pb, Zn, Au, Ag, etc.) mineralization and formation in extensional stress fields. Ore fluids responsible for these ore deposits are dominated by seawater origin, considering isotopic and chemical composition of ore fluids. Several active geothermal systems in Japanese Islands are associated with precious- and base-metal mineralization. Base metal mineralization takes place from hot springs containing high chloride concentration probably due to the contribution of seawater. Precious-metal mineralization occurs in the Osorezan hot springs which are characterized by neutral pH, high H2S concentration, and low salinity. These chemical features are similar to those of epithermal precious metal vein-type deposits in Japan. Intense hydrothermal activities of Neogene age from middle Miocene to present under the submarine and subaerial environments affect seawater chemistry and the global geochemical cycle. Summaries of chemical composition of hydrothermally altered volcanic rocks in the Green tuff region indicate that Mg in seawater exchanged for Ca of basalt and dacite. Annual Mg removal from seawater and Si, Ca and K inputs to seawater by the interaction of circulating seawater with igneous rocks in the Green tuff region were estimated. The estimates of Mg, Ca, Si, and K fluxes by this interaction suggest that the hydrothermal flux from middle Miocene back-arc basin at the circum-Pacific region is important as a factor controlling seawater chemistry. Fluxes of volatile elements (CO2, S, As) and other elements (Hg, Mn, Ba) due to hydrothermal activities at back-arc basins were calculated. Probably the hydrothermal flux of minor elements concentrated in Kuroko deposits (Sb, T1, etc.) is large compared with those from midoceanic ridges. CO2 flux from back-arc basins is estimated to be large compared with that from midoceanic ridges. Taking into account hydrothermal and volcanic gas, CO2 flux from back-arc basins and island arcs in the circum-Pacific region, the global carbon cycle from 60 Ma to present was computed based on modified BLAG and GEOCARB II models by Berner (1994) (Kashiwagi et al., 2000). The factors considered by these models are weathering, degassing from ridges and solar effects, but not degassing from back-arc basins and island arcs by hydrothermal solution and volcanic gas. The computations explain warming through Paleocene and early-middle Miocene and a temperature rise through early Miocene time; however, a significant temperature drop at middle Miocene age and

452

Chapter 5

gradual temperature decrease from middle Miocene to present cannot be explained by the change in CO2 concentration in the atmosphere. The CO2 concentration in the atmosphere is controlled by hydrothermal and volcanic gas fluxes from back-arc basins, island arcs and midoceanic ridges, chemical weathering of silicates, and carbonates related to the change in continental area, uplift of the Himalaya and runoff, depositions of marine carbonates and organic carbon. Probably climate change from the middle Miocene to the present were related to oceanic circulation, terrestrial biota, and albedo related to the change in topography and the area covered by ice sheets, and SO2 and volcanic ash originating from subduction volcanism. However, it is uncertain which are important factors causing the climate change. Thus, further works on climate change from the middle Miocene to present are clearly necessary.

References Berner, R.A. (1994) GEOCARB II: a revised model of atmospheric CO2 over Phanerozoic time. Am. J. Sci., 294, 56-91. Kashiwagi, H., Shikazono, N. and Tajika, E. (2000) Global carbon cycle model in the Cenozoic. 10th Annual V.M. Goldschmidt Conference, September 3-8, Oxford University, Abstract. Pearson, R.G. (1963) Hard and soft acids and bases. J. Am. Chem. Soc., 85, 3533-3539. Uyeda, S. and Kanamori, H. (1979) Back-arc opening and the mode of subduction. Geophys. Res., 84, 1049-1061.

Subject Index

alteration minerals, 23, 31-34, 59, 60, 95, 98, 99, 103, 104, 107, 110, 113, 119, 121123, 175, 185, 190, 194, 195, 201,295, 302, 318, 327-330, 342, 354, 381,417, 419, 420 alteration zone, 30, 34, 37, 38, 40, 105, 107, 109, 112, 186, 189, 194, 195, 197, 198, 200, 352 hydrothermal alteration, 23, 30, 36, 37, 51, 59, 60, 84, 98-100, 110, 113, 119, 122, 124, 166, 174, 183, 185, 190, 195-197, 201,214, 222, 264, 295, 315, 320, 371, 372, 377, 378, 408, 417, 420, 449 potassic alteration, 98-100, 123, 194 superimposed alteration, 36, 372 alunite, 36, 100, 108-110, 112, 113, 175, 200, 261, 263, 264, 352 andesite, 16, 18, 76, 102, 103, 108, 109, 124, 157, 161, 185, 186, 191-194, 196, 197, 199, 205, 206, 267, 323, 324, 333, 340-342, 351,359 anhydrite, 20, 23, 28, 29, 33, 35, 54, 55, 57-59, 61-65, 67, 80, 158, 177, 295, 301, 309, 310, 318, 319, 324, 331, 332, 334, 337, 339, 340, 342, 358-361, 366-371, 379, 381,392, 417 CaSO4, 20, 301 Sr content of anhydrite, 62 arc-trench system, 204, 223, 230 Archean lode gold deposits, 183 Arima hot springs, 321 As, 4, 14, 15, 18, 21-23, 25, 27, 30, 33, 34, 36, 3841, 44, 45, 47, 49, 50, 53, 57, 60-62, 6468, 71-77, 82, 84, 85, 88-100, 103-105, 107-112, 114, 116-121, 123, 124, 128132, 134, 136-142, 151, 153-155, 158, 159, 161, 162, 168-170, 172, 173, 175177, 179, 180, 182, 186, 187, 189-191, 195-201, 203, 205-208, 212, 213, 215, 220, 222, 223, 225, 227, 228, 231, 232, 235, 241, 243-247, 249, 250, 253-256, 258, 261, 267, 297, 299, 301, 302, 308311, 313, 315, 316, 318, 320, 323, 325330, 336, 340, 342, 350-355, 359-366, 368, 369, 371, 372, 374, 375, 378-381,

actinolite, 38, 99, 371,373, 380, 381, 418 active rift systems, 228, 230 activity coefficients, 47, 49, 50, 116, 128, 129, 131, 138-140, 237, 253, 261, 297, 308, 310, 356 activity of $2, 240 adularia, 84, 95, 98, 99, 103, 122, 123, 129-132, 135, 136, 143, 153, 155, 160, 161, 163166, 169, 173, 186, 187, 194, 240, 327 adularia/quartz ratio, 143, 187 Ag, 2, 4, 6-9, 12, 13, 20, 25-27, 41, 43, 44, 83, 84, 88, 89-94, 126, 130, 131, 137-139, 141143, 149, 154, 156, 157, 160, 161, 163, 167, 171, 173, 176, 179, 180, 182, 187, 201, 206, 207, 231, 236, 237, 239, 241, 243, 251-262, 313, 320, 322, 323, 327, 335, 336, 339, 342, 350, 353, 361-365, 378-381,385-389, 426, 449-451 acanthite, 94, 103, 129, 132, 134, 138-140, 162, 167, 379 Ag content of electrum, 26, 27, 43, 91-93, 124, 126, 129, 138, 139, 141, 162, 167, 171-173, 237, 251,261,262 Ag minerals, 89, 90, 162, 326 AgzS, 43, 128, 138, 139, 172 AgzSe, 128, 138-140, 325 Ag/Au ratio, 88, 142, 173, 252, 253, 256 Ag/Au total production ratio, 12, 148, 149, 156, 172, 173, 179, 325 AgC12 , 141, 171,253, 254 argentite, 41, 43-45, 47, 89, 94, 103, 124, 126, 128-131, 138-140, 161-165, 167, 171, 172, 240, 325, 337 age of mineralization, 161,235 Akenobe, 181,243, 246 alteration, 16, 19, 30, 33, 34, 36, 59, 64, 77, 95, 98-100, 103, 104, 107-113, 122-124, 132, 166, 169, 174, 185, 194, 195, 197, 201, 258, 261, 264, 265, 295, 302, 315, 320, 321, 324, 327, 328, 330, 331, 342, 352, 359, 371-373, 377-379, 381,419 advanced argillic alteration, 98, 100, 103, 107-113, 123, 167, 169, 173-175, 265, 266, 315, 331,359

453

454 385-389, 391, 392, 394, 407, 408, 413, 414, 417, 419-426, 431, 435, 436, 439, 441-443, 451 As flux, 421,422 As concentration of hydrothermal solution, 421 Au, 2, 4-10, 12, 13, 20, 45, 50, 76, 83, 84, 88, 96, 97, 113, 127, 137, 141-143, 146, 154, 156, 157, 160, 161, 164-168, 170, 171, 173-176, 179, 180, 182, 183, 187, 199201, 206, 207, 223-225, 228, 231, 236241, 243, 250, 252-261, 264, 266, 267, 311-313, 323, 325-327, 335, 336, 339, 341, 342, 351-353, 361-366, 386-389, 391,392, 425, 449-451 Au deposits, 10, 167, 183, 261 Au-Ag mineralization, 244, 247 Au-Ag vein, 251,252, 254 Au-Ag vein-type deposits, 26, 84, 85, 88-92, 94, 95, 98, 100, 103, 124, 126, 128130, 132, 134, 140, 142, 143, 145-147, 149, 151, 153, 156, 158, 159, 172-175, 177, 191,200, 202, 203, 206-209, 211, 238, 242, 251,253, 261,266, 315, 325, 330 Au-rich siliceous sinters, 326 Au precipitation, 142, 160, 174 AuCI~-, 254, 255, 257 electrum, 23, 25-27, 41, 43-45, 47, 8994, 103, 126, 128-131, 138-142, 161165, 167, 171-173, 186, 187, 200, 201, 236-238, 251-257, 261,361,379, 380 epithermal gold deposits, 10, 157, 160, 162, 163, 183, 188, 200, 324, 325 gold-quartz vein-type deposits, 258, 437 native gold, 240, 251-255, 257, 340, 379, 392 Ba, 4, 6, 30, 50, 70, 72, 76, 77, 79, 80, 215, 219, 266, 315, 336, 342, 361-363, 365, 370, 385-389, 407, 424-426, 451 Ba flux, 424 BaSO4, 20, 72, 74 barite, 15, 20, 23, 25, 28, 29, 40, 54, 55, 57-59, 61, 65-75, 77, 80, 98, 130, 134, 149, 156, 158, 161, 164, 166-168, 170, 175, 312, 333, 334, 337, 339-341,361, 366-368, 370, 371,379, 381,385, 387, 388, 391,392 barite ore, 23, 28, 67, 74, 75 precipitation mechanism of barite, 72 precipitation of barite, 68, 72, 74

Subject Index back arc, 223, 431,433, 439-441 back-arc depression, 15, 19, 333, 334 back-arc spreading, 6, 228, 336, 340, 341 back-arc spreading center, 341 back-arc basin, 4, 19, 66, 67, 71, 72, 75, 228, 229, 333, 354-357, 359, 361, 364, 407, 409, 411-415, 417-423, 425, 427, 451 basalt, 4, 17, 19, 35-38, 43, 44, 58-60, 76, 83, 118, 119, 121, 155, 205, 206, 221, 303, 333, 340-342, 360, 361, 365, 369, 371-375, 377, 378, 380, 382, 393, 394, 407-409, 414, 417, 419, 422, 423, 433, 451 benthic molluscs, 436, 437 Beppu hot springs, 134, 323 Besshi mine, 378, 388 Besshi-type deposits, 1, 2, 119, 373-376, 378-380, 383, 385, 386, 388, 390, 392, 394, 395, 451 Bi, 4, 88-90, 161, 180, 181, 207, 241, 325, 365, 387 black smoker, 67, 71,341,362, 363, 370 BLAG, 439-441, 451 boiling, 41, 61, 62, 81, 105, 106, 113, 123, 137, 143, 145, 147, 155, 161, 170-173, 175177, 200, 201, 258, 264, 320, 321, 326, 329, 330 bornite, 23, 25-27, 47, 130, 131, 135, 164, 165, 264, 337, 340, 368, 379, 392, 394 Broadlands, 119, 142, 256, 257, 297, 327, 328, 332, 422 C, 134, 135, 146-148, 153, 156, 158, 417, 419, 431,439, 440, 443, 451,452 global carbon cycle, 431,439, 440, 443, 451, 452 C isotopic composition, 41, 153 ~3C, 35, 145-149, 155, 156, 158, 179, 187, 257, 359, 360, 373, 414, 419, 439, 442, 443 CaO content, 38, 192, 408 carbonate, 34, 35, 41, 48, 81, 94, 95, 98, 99, 104, 105, 111, 134-137, 143, 145-148, 153, 156, 158, 166, 170, 179, 180, 258, 329, 357, 360, 365, 379, 414, 417, 419, 431, 440442, 452 ~13C of carbonate, 145, 147, 148, 177, 442 3J80 of carbonate, 35 calcite, 30, 35, 41, 43, 48, 64, 94, 95, 98, 103, 104, 107, 108, 111, 119, 121, 123, 136, 137, 145, 146, 149, 155, 158, 159, 161, 163-166, 185-187, 240, 247, 248, 258, 295,300, 302, 309, 310, 318, 319,

Subject Index 321,328, 329, 332, 337, 371,373, 381, 417-420, 431 dolomite, 35, 41, 48, 95, 136, 146, 164, 258, 318, 321,417-419 magnesite, 35, 41, 48 siderite, 94, 95, 98, 134-136, 146, 233, 240, 321,327, 381,418, 419 carbonate compensation depth (CCD), 81, 82 Ce anomaly, 58, 60, 158, 217, 439, 442 chalcedony, 307, 309 chemical composition of geothermal waters, 301, 302 chimneys, 6, 66, 67, 71, 75, 335, 338-341, 362364, 366-371,388 chloride complex, 253, 254, 257 chlorite, 23, 28, 30, 34, 36, 38, 43, 44, 60, 77, 94, 95, 98, 99, 104, 105, 107, 113-118, 136, 163-166, 185, 194, 233, 235, 240, 241, 295, 301, 318, 321, 324, 328, 330, 331, 334, 337, 361, 371-373, 379-382, 392, 408, 417, 418, 420 composition of chlorite, 116, 118, 381 Fe content of chlorite, 117, 327, 330 MgO/FeO of chlorite, 114 CI-, 30, 49, 77, 123, 156, 173, 180, 182, 261,296302, 304-306, 308-310, 316, 318, 324 C1- concentration, 49, 50, 77, 78, 122, 123, 173, 201,298-301,311,318, 323, 324 climate, 431,433, 435, 437, 439, 441,443 climate change, 431-433, 439, 440, 443, 452 CO2, 40, 47, 48, 76, 100, 111, 119, 123, 136, 137, 142, 171, 172, 175, 182, 183, 231, 256, 302, 304-306, 311, 316, 318, 319, 321, 328, 360, 413-419, 425, 426, 431, 433, 437, 439-444, 451,452 CO2 flux, 231,413, 414, 416, 417, 431,433, 439-441,451 CO2 fugacity (fco2), 47, 48, 107, 108, 111, 121, 135-137, 327-330, 419 hydrothermal CO2 flux, 414, 416, 417, 431, 433, 435, 437, 439, 441,443 Pr 329, 440-443 colloform textures, 382, 383 complexes, 1, 49, 60, 72, 119, 141, 142, 156, 171, 173, 174, 180, 182, 200, 201, 252, 253, 256-258, 353, 364, 377 Au thio complex, 200 bisulfide complex, 142, 182, 253, 254, 256 chloro complex, 49, 171, 173, 182, 201,353 thio complex, 171,200, 364 compositional zoning, 25, 27, 121, 126, 128, 380

455 country rock, 15, 22, 61, 64, 80, 101, 123, 147, 148, 166, 176, 177, 189, 195, 203, 211, 212, 224, 295, 329, 342, 364, 419, 449, 450 coupled fluid flow-precipitation model, 69 coupled fluid flow-reaction model, 67 cristobalite, 28, 30, 123, 124, 196, 201,323, 352 Cu, 2, 4, 6-9, 12-14, 20, 25, 27, 40, 50, 76, 77, 80, 88, 91, 92, 94-98, 128, 137, 142, 148, 149, 154, 156-158, 161, 179, 180, 182, 187, 203, 206, 215, 219, 224, 231, 241, 243, 259, 261, 311, 313, 316, 320-323, 325, 327, 339, 342, 353, 361-365, 373, 375, 385-389, 392, 393, 419, 426, 449451 chalcopyrite, 20, 21, 23, 26, 40, 47, 52, 53, 67, 74, 75, 88, 95, 97, 103, 130, 131, 135, 142, 161-165, 176, 186, 187, 201, 233, 236, 238, 240, 246, 251,258, 264, 318, 327, 334, 337, 339-341,367, 368, 374, 379, 381,383, 387, 388, 391,392, 394, 395 CuFeS2, 142, 172, 201 Cu-Pb-Zn vein-type deposit, 84, 114, 116, 118, 136, 146, 419 dacite, 16, 20, 21, 30, 31, 34, 40, 57-60, 71, 84, 157, 161, 185, 203, 206, 267, 340-342, 409, 451 degassing, 123, 200, 201,417, 421,439-441,451 dendritic crystal, 72-74 deposition, 10, 23, 30, 47, 61, 63, 64, 94, 99, 100, 116, 118, 124, 129-131, 168, 170-175, 195, 200, 201, 313, 335, 387, 391, 392, 394, 441,450 anhydrite deposition, 61, 62 depositional mechanism, 61, 171, 176, 196, 200, 201,449 dip of Benioff-Wadati zone, 225, 226 discharge zone, 31, 35, 38, 61, 117, 118, 123, 189, 414 disseminated-type deposits, 12, 149 East Pacific Rise, 38, 60, 121, 221, 339, 355, 356, 361,364, 367, 383, 386, 415 electrum, see Ag, Au, Hg epidote, 38, 43, 60, 77, 95, 98, 99, 104-108, 110, 111, 119-121, 136, 137, 295, 318, 321, 324, 327, 329-331, 342, 354, 371, 373, 380, 381,392, 417, 418 Fe content of epidote, 38, 119-121 epithermal base-metal deposits, 158

456 epithermal deposits, 99, 124, 144, 145, 148, 151, 152, 158, 264, 325, 326 epithermal gold deposits, 10, 157, 160, 162, 163, 183, 188, 200, 324, 325 epithermal gold vein-type deposits, 5 epithermal gold disseminated-type deposits, 85, 172, 262 epithermal precious-metal deposits, 10, 450 epithermal system, 113, 169, 199, 265, 315, 391, 392 epithermal vein deposits, 214 epithermal vein-type deposits, 1, 6, 7, 14, 27, 8385, 98, 100, 113, 124, 135, 142, 145, 156, 158, 173, 180, 182, 203, 207, 225, 235, 237, 244, 324, 325, 327, 328, 352, 353, 449, 450 estimate of temperatures, 355 Eu anomaly, 57-61, 159, 217, 219, 225 Fe, 6, 8, 9, 20, 25, 35, 40, 50, 76, 77, 79, 91, 94-97, l l4, 118, 142, 148, 156, 177, 182, 206, 233, 241, 245, 254, 301, 313, 315, 320-322, 324, 328, 339, 341, 342, 353, 356, 357, 361-363, 365, 381, 385-389, 392, 420, 426, 449 Fe content of sphalerite, 91, 126, 138, 238, 241,361,394, 395 FeS content of sphalerite, 45, 90, 124, 126, 129, 138, 139, 162, 167, 233, 241,243, 251 FeS2, 115, 128, 136, 138, 139, 418 feldspar, 61, 121, 190, 195, 196, 297-299, 365 ferruginous chert, 23, 57-61, 67, 71 fluid inclusion, 25, 38-41, 47-51, 61, 62, 65, 68, 71, 77, 80, 82, 105, 106, 109, 110, 112, 113, 116, 124-129, 135-140, 142, 143, 155, 161, 162, 167, 170, 175-179, 183, 187, 197, 201, 203, 209, 233, 238, 241243, 247, 251, 254, 256, 261, 264, 326, 330, 367, 414, 418, 449 filling temperature, 125, 212, 243 homogenization temperature, 25, 40, 41, 4548, 71, 105, 107, 109, l l0, 112, !13, 124, 126, 128, 129, 136-140, 155, 161, 162, 167, 178, 190, 197, 209, 211,232, 238, 242, 251,261,264, 326, 327, 330 inclusion fluids, 40, 46, 62, 81, 105, 124, 129, 142, 148, 153, 176, 177, 209, 211, 241,256, 297, 298, 301,326 NaC1 equivalent concentration, 124, !42, 209, 256

Subject Index salinity, 40, 41, 46, 49, 61, 67, 68, 81, 129, 141, 142, 148, 169, 175, 177, 178, 183, 201,203, 209, 212, 241,251,253, 256, 261,264, 310, 318, 321,449, 451 flux, 407, 409, 411, 413-415, 417, 419-425, 427, 431,439, 441 As flux, 421,422 Ba flux, 424 CO2 flux, 231,413, 414, 416, 417, 431,433, 439-441, 451 Hg flux, 423, 424 hydrothermal flux, 417, 421-425, 451 hydrothermal CO2 flux, 414, 416, 417, 431, 433, 435, 437, 439, 441,443 Mn flux, 424 S flux, 420, 421 foraminifera, 61, 62, 81,431,432, 435 foraminiferal shell, 431,439 framboidal textures, 368, 383 freezing temperature, 41, 129, 142, 261 fugacity, 25, 41, 47, 59, 107, 110, 112, 119, 121, 129, 135, 138, 167, 168, 200, 301, 328, 449 fco2, 47, 48, 107, 108, 111, 135-137, 327330, 419 fH2S, 107, 111, 112, 330, 331, 364 f02, 23, 47, 50, 59, 84, 94, 110, 111, 114, 116, 118-121, 129-132, 134, 135, 137, 141, 142, 171, 172, 241,264, 266, 301, 327, 361, 364, 390-392, 450 foz-pH region, 141, 173, 174 fs2, 25, 41, 43, 45, 47, 84, 94, 110, 116, 124, 129, 138-141, 167, 171, 172, 237, 238, 240, 241,243,244, 246-248, 251,261, 301,327, 330, 361,364, 380, 392 Fukazawa, 9, 20, 22, 23, 25, 27, 29-34, 37, 40, 55, 56, 59, 62, 63, 65,415 Furutobe, 9, 28, 37, 40, 115, 119, 120, 368, 378 Fushime, 295, 296, 313, 314, 324, 327 gangue minerals, 10, 28, 66, 94, 95, 98, 103, 129, 135-138, 148, 149, 156, 161, 163-166, 183, 186, 187, 233, 235, 241, 247, 327, 328, 337, 380, 381,392, 418 gas fugacities, 38, 110, 111, 124 GEOCARB, 439-441, 451 geothermal, 110, 113, 114, 117, 119-121, 137, 142, 168, 257, 295-303, 305, 307-313, 315317, 319, 321, 323-325, 327-333, 335, 337, 339, 341, 351, 353, 355, 357, 359, 361, 363, 365, 367, 369, 371, 373, 375,

Subject Index 377, 379, 381, 383, 385, 387, 389, 391, 393, 395 active geothermal systems, 23, 107, 108, 112, 170, 196, 197, 199, 324-329, 331,451 geothermal system, 117, 119-121, 168, 170, 171,295, 301,303, 304, 311,312, 318, 320, 323-325, 329, 333, 355, 417 submarine geothermal system, 60, 117 geothermal water, 49, 77-79, 119, 120, 129, 131, 142, 199, 256, 257, 295-302, 307, 308, 310-314, 318, 320, 322, 324, 328, 329, 331,355, 356, 419 geothermal area, 111, 119, 121, 199, 223, 310, 315, 318, 322, 324, 351,370 geothermal power stations, 295 geothermometer, 126, 161,241,242, 302, 354 Na-K geothermometer, 355 Na-K-Ca geothermometer, 308 silica geothermometer, 355 stannite-sphalerite geothermometer, 242 global long-term carbon cycle, 439 grain size, 75, 98, 126, 145, 146, 166, 221, 368371,383 granitic rocks, 2, 59, 81, 88, 231, 250, 321, 323, 378 granitoids, 2, 258 ilmenite-series granitic rocks, 4, 231 magnetite-series granitic rocks, 4, 231 Green tuff, 4, 16, 85, 134, 146, 149, 153, 155, 169, 178, 179, 206, 211, 228, 231, 234, 407-411 Green tuff activity, 411 Green tuff region, 4, 5, 15, 81, 84, 151, 155, 156, 175, 202, 228, 231,235, 248, 318, 407-409, 411, 413, 414, 449-451 Green tuff type, 148-150, 153, 155, 156, 158, 179 groundwater, 100, 110, 112, 113, 123, 187, 189191, 195-201,313, 329, 449 gypsum, 1, 15, 20-23, 28, 29, 33, 35, 54, 55, 57, 61, 98, 158, 177, 379, 381,417 H2S, 10, 30, 47, 49-51, 65, 66, 76, 80, 107, 111, 112, 116, 123, 130, 131, 134, 142, 151, 155, 156, 158, 173, 175, 179, 182, 191, 195, 200, 201, 254, 256, 261, 264, 304306, 312, 314-316, 318, 319, 330, 331, 342, 359, 360, 369, 420, 426, 449, 451 halloysite, 5, 321,334, 337 Hanaoka, 8, 16, 17, 20-22, 28, 75, 115 Hanaoka district, 22, 53

457 hematite, 23, 25, 43, 67, 89, 90, 94, 95, 98, 107, 108, 110, l l l , 120, 129-131, 134-136, 151, 163-165, 264, 323, 330, 331, 361, 366, 371,373, 379, 392, 417-419 heterogeneous nucleation, 75 Hg, 4, 14, 25, 27, 76, 156, 161, 176, 180-182, 187, 207, 237, 238, 247, 248, 303, 313, 316, 322, 325, 327, 342, 353, 364, 365, 385, 387-389, 422-425, 437, 449, 451 cinnabar, 163, 236, 237, 247, 248, 312, 325, 423 Hg-Sb deposits, 5, 235, 451 Hg content of electrum, 236, 238 Hg flux, 423, 424 Hg-type deposits, 84 mercurian gold, 237, 238, 240 high sulfidation type, 14, 113, 169, 261,265, 266 high sulfidation-type Au-Ag deposits, 113 Himalayas, 432 Hishikari deposit, 183, 199-201 Hishikari mine, 11, 144, 187, 196 Hitachi, 375-378, 380, 381,385, 390-393, 451 Hokuroku, 8, 15, 17-20, 22, 23, 30, 31, 34, 36-38, 55, 57, 61, 64, 65, 81,205, 206, 222, 225, 378 Hosokura, 12, 84, 88, 90, 99, 127, 137, 148, 149, 151, 181,203, 222, 224 host rocks, 84, 95, 103, 114, 117, 121, 153, 157, 160, 161, 168, 177, 179, 208, 247, 258, 260, 261,264, 321,359, 375, 378, 388 hot spring waters, 309, 313, 315, 319, 321,323 hot spring-type deposits, 14, 266, 325, 326 hot spring-type gold deposits, 261 hot-water-dominated system, 107 HSAB principle, 156, 180, 182, 183, 449 hydrothermal activity, 1, 4, 21, 60, 61, 203, 212, 213, 215, 219, 222, 225, 321, 323, 367, 369, 370, 378, 413, 433, 437 hydrothermal alteration, 23, 30, 36, 37, 51, 59, 60, 84, 98-100, 110, 113, 119, 122, 124, 166, 174, 183, 185, 190, 195-197, 201, 214, 222, 264, 295, 315, 320, 371, 372, 377, 378, 408, 417, 420, 449 hydrothermal alteration zoning, 104, 195, 199 hydrothermally altered basalt, 37, 38, 58, 60, 359, 371,373 hydrothermal deposits, 224, 336, 352, 394, 451 hypothermal-type deposits, 249 subvolcanic hydrothermal deposits, 235 hydrothermal flux, 417, 421-425, 451

458 hydrothermal ore deposits, 1, 6, 28, 66, 68, 69, 71, 87, 115, 118, 130, 183, 214, 367, 370, 437, 449, 450 hydrothermal plumes, 75, 366, 369, 370 hydrothermal solution, 19, 23, 30, 31, 34, 35, 47, 48, 50, 59-69, 72, 75-81, 83, 100, 114, 117, 118, 122, 123, 130, 142, 146, 155, 172, 173, 175-177, 187, 189-191, 194199, 201, 217, 222, 233, 295, 311, 315, 318, 330, 337, 342, 350, 354-361, 366, 367, 369-371, 376, 380, 381, 383, 385, 387, 388, 413-425, 437, 449, 451 hydrothermal systems, 38, 67, 71, 80, 114, 117, 118, 123, 155, 175, 177, 189, 201, 202, 265, 266, 313, 315, 321, 327, 358, 361, 364, 407, 410, 414, 416, 419, 425, 449, 450 hypo/mesothermal polymetallic vein-type deposits, 250 Ikuno, 13, 88, 90, 94, 130, 13 l, 172, 18 l, 246 island arcs, 6, 7, 10, 19, 119-121, 205, 212, 361, 378, 379, 385, 407, 413, 421, 422, 439, 441,451,452 isotope, 34, 38, 242, 263, 359, 365, 366, 369, 425, 442 ~13C, 35, 145-149, 155, 156, 158, 179, 187, 257, 359, 360, 373, 414, 419, 439, 442, 443 ~180, 33-35, 37-40, 51, 52, 54, 65, 80-82, 143-148, 155, 158, 176, 177, 179, 187, 189-191, 194, 200, 201, 21 I, 232, 233, 263, 264, 313, 318, 321,323, 324, 334, 359, 373, 419, 431,435, 439, 443 ~348, 1, 52-54, 65, 80, 81, 147-156, 158, 161, 167, 168, 170, 175, 177, 179, 191, 203, 208, 211,231,233, 234, 241,248, 257, 260, 264, 359, 360, 369, 371,383, 385, 391,419 hydrogen isotopes, 34, 187 lead isotopes, 333, 392 radiogenic isotopes, 295 Itoigawa-Shizuoka Line, 231,250 Izu peninsula, 6, 84, 87, 94, 95, 100, 101, I 13, ! 2 I, 146, 152, 153, 167, 174, 224, 332, 352 Izu-Bonin arc, 6, 334, 337, 361 Izu-Ogasawara, 225, 336, 337 Jackson's episode, 225, 438, 450 Japanese islands, l, 6, 7, 10, 14, 84, 201,202, 204, 205, 222, 248, 295, 324, 333, 407, 411, 433,436-438, 449-451

Subject Index K-Ar ages, 4, 10, 19, 84-87, 157, 160, 206, 21l, 222, 226, 240, 247, 312, 408, 438, 449 K-feldspar, 36, 48, 59, 98, 99, 104, 105, 107, 111, 119, 121, 123, 136, 142, 190, 194-196, 201, 222, 241, 256, 295-298, 301, 302, 307, 310, 319, 321, 327, 329, 331, 355, 417, 418, 420 K-mica, 48, 107, 112, 119, 121, 136, 142, 194, 195, 201, 256, 302, 307, 318, 319, 329, 331, 417-420 kaolinite, 4, 5, 30, 36, 48, 95, 98, 100, 112, 113, 119, 123, 129-131, 135, 136, 161, 163165, 175, 190, 195, 201, 233, 241, 261, 264, 315, 318, 321, 323, 328, 334, 337, 417-419 kinetics, 67-69, 71, 74, 123, 355 kinetics-fluid flow mixing model, 196 precipitation rate constant, 69, 71, 198 uniformly mixed kinetics model, 69 Kishu deposit, 240 Kosaka, 9, 20, 22, 33, 40, 55, 56, 415 Kosaka district, 21, 22, 53 Kuroko, 1, 4-7, 11, 14-16, 18-30, 33, 35-41, 43, 45, 46, 48, 50-68, 71, 72, 74, 75, 77, 80-82, 84-86, 91, 92, 114-116, 118-120, 124, 129, 130, 145, 153, 155, 158, 173, 177-179, 183, 201-203, 205-214, 222226, 228, 231, 234, 235, 333, 335, 336, 339, 342, 350, 352, 353, 359, 365-373, 375, 376, 378, 380, 381, 383, 385, 390393, 408, 414-417, 419, 421, 423, 425, 437, 449-45 I black ore, 7, 15, 20, 21, 23, 25, 27, 40, 47, 55, 56, 67, 71, 74, 75 Kuroko deposits, 4-7, 14-16, 19-28, 30, 33, 35, 36, 38-40, 45, 52-57, 59, 61, 62, 64, 66-68, 71, 72, 74, 75, 81, 82, 8486, 91, 92, i14-116, 118, 124, 129, 130, 153, 155, 158, 173, 177-179, 183, 202, 203, 206-214, 223-226, 228, 231, 235, 335, 350, 352, 353, 366, 367, 370, 373, 376, 380, 381,383, 385, 390-392, 408, 414, 421,423, 437, 449-451 Kuroko ore fluids, 41, 48-52, 77, 80-83, 116, 130, 136, 366, 449 Kuroko ore-forming fluids, 51 Kuroko ore solution, 50, 80 Kuroko-type deposits, 211,393 Kushikino, 13, 89, 94, 96, 99, 100, 113, 119, 127, 139, 140, 143, 154, 156-158, 162, 163, 166, 181

459

Subject Index laumontite, 95, 99, 104, 105, 110, 111, 123, 295, 302, 418 ligands, 180, 182 low sulfidation-type deposits, 14, 113, 169, 264266, 333, 352 low sulfidation-type Au-Ag deposits, 113, 327 magmatic fluids, 80, 81, 113, 419 magnetic susceptibility, 219-221,223 magnetite, 89, 94, 95, 98, 105, 110, 111, 130, 131, 135, 136, 163, 221, 222, 231, 264, 330, 331, 340, 342, 361, 366, 379, 381, 417419 Mariana, 66, 67, 75, 77, 82, 350, 356, 361, 367371,385, 414, 451 Mariana Trough, 336, 341,342, 359, 361,364, 367, 389, 407, 415, 420 Matsumine, 8, 21, 29, 40, 415 Median Tectonic Line, 3-5, 85, 160, 231,236, 247, 353 metal ratios, 83, 180, 183, 385 metal zoning, 88, 232 metallogenic province, 14, 206, 212, 450 metallogeny, 1 metamorphism, 321, 378, 380, 381,383, 394, 433, 439 meteoric water, 4, 81-83, 113, 114, 143-145, 147, 158, 175-177, 187, 190, 233, 263, 264, 266, 318, 321,324 MgO content, 37, 193, 408, 409 midocean ridge, 333, 339, 362, 373, 423 midoceanic ridge basalt, 38, 121, 365, 393, 394, 407-409, 413, 422 midoceanic ridge deposits, 1, 66, 155, 336, 359, 361,364, 373, 375, 380, 383, 385, 451 Miocene, 4-6, 10, 14, 15, 17, 19, 20, 38, 54, 61, 62, 64, 82, 84, 85, 101,146, 201-203,205, 206, 211-215, 224, 226, 228, 231, 232, 235, 240, 247-250, 315, 318, 320, 323, 350, 352, 371, 376, 393, 407-411, 413, 414, 431-433, 435--437, 440--443, 450452 mixing, 23, 34, 53, 59, 61, 63-65, 68, 71, 72, 74, 75, 77, 80, 81, 83, 100, 110, 113, 118, 145, 155, 173, 175, 177, 187, 189-191, 195, 196, 198-201, 257, 263, 264, 313, 318, 321, 329, 330, 359-361, 370, 371, 381,383, 385, 387, 388, 420, 449 mixing of fluids, 118, 123, 196, 199, 200 Mn, 3, 4, 12, 13, 60, 76, 77, 88, 90, 95-97, 128, 145, 147, 148, 156, 157, 161, 166, 177, 180, 182, 206, 315, 320-322, 324, 325,

328, 353, 356, 357, 361-363, 365, 386, 387, 424, 426, 449, 451 montmorillonite, 5, 30, 33, 34, 38, 100, 104, 108, 110, 112, 113, 190, 261, 315, 323, 324, 334, 337, 341 MORB, 19, 44, 356, 365, 371-373, 393, 394 morphology, 28, 53, 72 muscovite, 30, 295, 298, 301,321,380, 381 Na20 content, 31, 38, 195 Nigorikawa, 295, 303, 307, 311,327 O isotopes, 39, 155, 183, 199, 263 O isotopic composition, 187, 211, 414 ~180, 33-35, 37-40, 51, 52, 54, 65, 80-82, 143-148, 155, 158, 176, 177, 179, 187, 189-191, 194, 200, 201,211,232, 233, 263,264, 313, 318, 321,323, 324, 334, 359, 373, 419, 431,435, 439, 443 O isotopic equilibrium, 187, 190 Obira deposit, 240 Oga peninsula, 213-215, 227 Okinawa Trough, 6, 41, 81, 82, 223, 225, 229, 230, 333, 335-337, 350, 351, 357, 359, 361, 364, 385, 407, 420, 421,451 Oko, 15 Okuaizu, 295, 296, 303, 305, 307, 311, 313-315, 318, 320, 327, 331 opaque minerals, 105, 120, 161-165, 183,233, 236, 240, 251,327, 379, 381,395 opening of the Japan Sea, 19, 226-228 ore fluids, 10, 38, 40, 41, 45, 47, 49, 68, 77, 80, 81, 84, 114, 118, 124, 128-130, 134, 141-145, 147, 148, 150, 153, 155, 158, 167, 168, 170, 171, 173, 175-177, 179, 180, 182, 183, 187, 200, 201, 208, 211, 241, 251257, 261, 264, 326, 328-330, 353, 361, 380, 385, 387, 388, 390, 416, 423, 425, 449-451 chemical composition of ore fluids, 77, 80, 124, 451 Kuroko ore fluids, 41, 48-52, 77, 80-83, 116, 130, 136, 366, 449 origin of ore fluids, 77, 147, 449, 450 ore texture, 383 Osorezan, 6, 113, 157, 160, 164, 181, 311-315, 326, 327, 451 paleobathymetry, 82, 83 paleogeography, 202, 205, 437, 450 paragenetic sequence, 186 partition coefficient, 62, 242, 298

460 Pb, 2, 4, 6, 7, 12-14, 20, 30, 40, 50, 72, 76, 77, 88-90, 137, 142, 148, 149, 156, 157, 161, 179, 180, 182, 187, 206, 231, 241, 243, 311, 313, 316, 320-325, 327, 336, 339, 342, 353, 359, 361-365, 386-389, 392, 393, 419, 426, 449-451 Cu-Pb-Zn vein-type deposit, 84, 114, 116, 118, 136, 146, 419 galena, 15, 20, 23, 25, 40, 52, 53, 56, 67, 74, 75, 88, 94, 95, 103, 138-140, 142, 161-165, 171, 175, 176, 186, 187, 201, 236, 240, 258, 321,327, 333, 334, 337, 339-341,361,368, 381,391 PbC12, 172, 175, 201 PbS, 138, 172, 175, 201 Pb isotopic composition, 54, 55, 158, 333 Pearson's rule, 182 pH, 8-10, 30, 49-51, 59, 76, 81, 114, 116-118, 123, 129, 131, 132, 134, 141, 142, 150, 161, 167-169, 171-173, 175, 195, 201, 241, 251, 253, 256, 257, 261, 264-266, 295, 299, 300, 302-306, 308, 309, 314316, 318, 320, 322, 327, 357, 359, 361, 390, 442, 449-451 plagioclase, 30, 59, 104, 123, 195, 321, 328, 354, 356, 381,418 plate, 1, 2, 4, 203, 213, 223, 229, 266, 417, 433, 440, 449, 450 North American plate, 4, 212 Pacific plate, 1, 4, 203-205, 212, 213, 223, 225, 229, 265, 266, 334, 438, 450 Philippine Sea plate, 334 plate kinematic model, 229 plate tectonics, 450 Pliocene, 17, 84, 101-103, 213, 222, 261,315, 324, 409, 433 plume injection, 228 pollen assemblage, 431,435 polymetallic vein, 15 I, 231 precipitation, 61, 62, 64-69, 72-75, 111, 124, 126, 128, 140, 142, 160, 183, 196, 198-201, 321, 326, 333, 355, 359, 366, 368, 371, 383, 442 anhydrite precipitation, 62 Au precipitation, 142, 160, 174 precipitation mechanism of barite, 72 precipitation of barite, 68, 72, 74 precipitation of quartz, 196 precipitation rate constant, 69, 71, 198 sequence of mineral precipitation, 368 prehnite, 94, 95, 99, 104, 105, 111, 120, 121, 136, 295, 321,418

Subject Index propylite, 95, 99, 121, 157, 235 propylitic alteration, 98-100, 103, 104, 107, 108, 110, 111, 113, 123, 166, 174, 194, 266, 329, 331 pyrite, 21, 23, 25, 27, 43, 47, 52, 53, 67, 74, 80, 88, 90, 94, 95, 103-105, 107, 108, 110, 111, 114, 116, 120, 124, 126, 128-131, 134-140, 142, 161-165, 167, 176, 177, 185-187, 191, 236, 238, 240, 243, 244, 247, 251, 252, 257, 258, 261, 264, 295, 312, 321, 327, 330, 331, 333, 334, 337, 339-342, 352, 366-369, 371, 374, 379381, 383, 390-392, 394, 395, 417, 418, 422, 423 pyrrhotite, 89, 90, 94, 95, 105, 110, 111, 116, 120, 129-131, 135, 136, 151, 163-165, 233, 236, 238, 240, 241, 243, 244, 251, 252, 258, 264, 337, 340, 341, 361, 366, 368370, 374, 378-381, 388, 390-392, 394, 418 quartz, 10, 23, 28, 30, 40, 45, 48, 65-71, 75, 94, 95, 98, 103-105, 107-109, 111-113, 122, 123, 136, 140, 142, 143, 146, 153, 155, 161, 163-166, 173, 185-187, 190, 191, 194, 196, 197, 199, 200, 232, 233, 238, 240, 241, 247, 248, 256, 258, 261, 263, 264, 295, 298, 301, 302, 304-307, 310, 326, 352, 354, 355, 371, 373, 376, 379-381, 417, 418, 420 amorphous silica, 28, 66, 67, 71, 123, 333, 337, 340, 341,350, 367, 368 precipitation of quartz, 196 SiO2, 136, 215, 322 SiO2 mineral, 196 solubility of quartz, 66 solubility of quartz minerals, 196 rare earth elements (REE), 19, 57-61, 158, 159, 215, 217, 220, 221,351,352, 449 recharge zone, 117, 123, 410 reservoir temperature, 319, 332 rhyolite, 16, 76, 84, 157, 267, 318, 341,342, 375 S, 1, 2, 4-6, 8, 9, 14, 20, 25, 49, 51-54, 56, 61, 63, 65, 66, 73-75, 77, 80, 81, 96, 97, 100, 107, 108, 110-112, 116, 120, 129, 130, 132-135, 137, 140, 141, 143, 148151, 153, 155, 158, 161, 167, 168, 173, 175-179, 183, 184, 191, 197, 198, 208, 210-212, 234, 240-242, 248, 252, 258, 261, 264, 297, 301, 303, 304, 320, 324327, 331, 335, 337, 339, 340, 350, 359,

Subject Index 362, 363, 369, 371, 383, 385, 386, 388, 390, 392, 410, 413, 415-417, 420, 421, 449-451 S content, 132, 155, 231,233, 359, 360, 422 S cycle, 420, 421 S deposits, 6 S flux, 420, 421 S fugacity (fs2), 25, 41, 43, 45, 47, 84, 94, 110, 116, 124, 129, 138-141, 167, 171, 172, 237, 238, 240, 241,243, 244, 246-248, 251,261,301,327, 330, 361, 364, 380, 392 S02, 100, 123, 168, 175, 191, 195, 264, 359, 364, 442, 452 SO]-, 62-65, 72, 116, 134, 151, 173, 175, 179, 191,200, 304, 316, 324, 332, 370 total dissolved sulfur (ES), 49, 81, 114, 116118, 129-135, 169, 177, 264 S isotopes, 4, 53, 199, 207, 242, 243, 360, 383, 385, 392 S isotopic composition, 1, 52, 80, 149, 151, 167, 175, 176, 210, 261,383, 450 334S, 1, 52-54, 65, 80, 81, 147-156, 158, 161, 167, 168, 170, 175, 177, 179, 191, 203, 208, 211,231,233, 234, 241,248, 257, 260, 264, 359, 360, 369, 371,383, 385, 391,419 334S and 3180 of sulfate, 54, 158 334S of sulfate, 161, 175 334S of sulfide, 53, 148, 150, 151, 155, 161, 170, 177, 240, 257, 260, 264, 359, 369 S isotopic values, 4, 150, 360 Sado, 13, 54, 84, 126, 127, 152, 154, 156, 157, 159, 162, 164, 172, 203, 222, 224 saturation index, 66, 68, 73, 74, 371 Sb, 4, 13, 14, 25, 76, 88, 91, 92, 96, 97, 180, 207, 235, 247, 248, 313, 316, 320, 323, 327, 336, 342, 361-365, 386-389, 425, 437, 451 Se, 4, 89, 90, 96, 97, 130, 131, 134, 135, 159, 160, 180, 187, 206, 207, 325, 327, 361-363, 365, 385-387, 426 Se-bearing deposits, 14, 160 Se fugacity (fse2), 138-140 Se minerals, 89, 159, 160 Se-type deposits, 90, 91, 93, 98, 131, 157, 159-163, 166-171,264, 266, 315, 333, 450 seafloor spreading, 411,439 seawater, 2, 23, 34, 41, 50, 51, 53-55, 57, 5968, 72, 74-77, 79-83, 114, 118, 143, 148, 155, 156, 158, 168, 169, 177, 191, 203,

461 211, 217, 222, 231, 234, 241, 248, 257, 261, 318, 324, 334, 353, 354, 357-361, 366, 369-371, 374, 381, 383, 385, 387, 388, 407-411, 413, 414, 419-424, 435, 442, 449, 451 depth of seawater, 46 Japan Sea, 19, 213, 215, 222, 226-229, 250, 410, 436, 437 marginal seas, 228 seawater cycling, 41 O, 411, 413, 414, 416, 420, 423, 424 seawater/rock ratio, 60, 81,414, 421,424 seawater-rock interaction, 50, 63, 77, 80, 83, 324, 414, 419, 449 Se, 94, 128, 130-134, 138-140, 253 sedimentary rocks, 33, 38, 84, 143, 149, 151, 153, 155, 156, 158, 161, 169, 177, 184, 187, 203, 206, 211, 213, 215, 217, 219-222, 248, 250, 258, 311, 320, 321, 324, 353, 377, 449, 451 Seigoshi, 13, 84, 94, 95, 99-101, 103, 104, 106108, 110-112, 119-121, 126, 127, 130, 131, 137, 143, 146, 153, 154, 157, 159, 162, 163, 166, 172, 174, 177, 181, 315, 329, 331,332, 352 sericite, 4, 5, 16, 28, 30, 33, 35, 36, 38, 59, 84, 95, 98, 103, 108, 123, 129-132, 135, 136, 160, 163-166, 185, 222, 233, 238, 240, 241,261,310, 318 Shimokawa, 1, 119-121, 374-378, 380, 381, 383, 393 siliceous ore, 15, 21, 23, 26-28, 40, 53, 67, 71, 74, 376 SiO2 content, 100, 123, 124, 191, 192, 195, 215 skarn deposits, 26 skarn-type deposits, 4, 155, 175, 246 smectite, 94, 95, 98, 99, 105, 136, 163, 165, 166, 186, 194, 321,408, 418 Sn, 2, 4, 13, 88, 180-182, 207, 231,241,243, 325, 342, 365, 378, 385-389, 450 softness, 180, 182 solid solution, 94, 95, 117, 118, 128, 131, 137, 139, 140, 245, 309, 354, 417 solubility, 45, 49, 50, 62, 66, 67, 72-74, 80, 123, 141, 171, 173, 175, 183, 195, 197, 199, 200, 253, 254, 332, 364, 366, 421,423 solubility of electrum, 171 solubility of quartz, 66 solubility of SiOs minerals, 196 Sr, 29, 55, 57, 63, 75 87sr/g6sr, 19, 55, 57, 62-65, 80, 231, 232, 333, 350, 356, 357, 359, 361,371,449

462 87Sr/86Sr of barite, 57 Sr content of anhydrite, 63, 77 Sr isotopic composition, 55 stable isotopes, 124, 171,372 stannite, 232, 233, 241-244, 246, 325 stannoidite, 164, 241,244-247, 379 stress field, 203-205, 213, 222, 223, 450 stress trajectory maps, 204 subduction, 4, 203, 213, 225, 227, 228, 231, 265, 266, 334, 417, 421,423, 440, 450, 452 Mariana-type subduction, 4, 231,450 mode of subduction, 4, 203, 213, 225, 226, 450 normal subduction, 265 oblique subduction, 4, 212, 265, 266, 450 subduction zones, 2, 4, 417, 437, 441 subsidence, 15, 18, 19, 82, 202, 205, 226, 228, 450 sulfate-sulfide chimneys, 75, 370 Sumikawa, 295, 296, 303, 305, 307, 314, 320, 321 supersaturation, 64, 66, 74, 123, 383 Tanakura Tectonic Line, 85, 160, 227, 231, 236, 250 Te, 4, 91, 159, 161, 162, 180, 207, 313, 325, 327, 381,385-387 Te minerals, 90, 159, 162, 166 Te-bearing deposits, 14 Te-type deposits, 90, 93, 98, 131, 157, 159162, 164-171,264, 266, 315,450 tectonics, 1,201,202, 213, 228, 450 plate tectonics, 450 tectonic evolution, 201,202 tectonic setting, 1, 15, 19, 335-337, 353, 375, 377, 378, 385,407 tetrahedrite, 159, 161, 162, 187, 318, 327, 339 tetrahedrite-tennantite, 23-26, 91, 92, 94, 337, 350, 361,368, 379, 381,392 tetsusekiei, 28, 74 TI, 176, 180-182, 325, 327, 353, 365, 425, 449, 451 total dissolved sulfur concentration, 49, 114, 129, 134, 177 total production, 88, 160, 161, 179, 325, 326 Toyoha, 12, 28, 88-91, 94, 95, 99, 100, 114-117, 127, 130, 131, 134, 135, 149, 151, 172, 175, 178, 181,208, 211,212, 224 Tsugu, 181,223, 231,234-240, 353 Ugusu, 100-102, 107-109, 112, 113, 265 uplift, 4, 21, 202, 203, 205, 212, 213, 228, 440, 450, 452 upwelling of asthenosphere, 228, 230

Subject Index Uwamuki, 9, 16, 17, 20, 23, 34, 45, 53, 56, 415 vapor-dominated system, 107, 111, 112, 331 vein types, Au-Ag vein, 10, 84, 86, 88, 89, 94, 100, 105, 108, 121, 141, 142, 149, 153, 155, 159, 173-177, 206, 244, 251-256, 264, 313, 330, 332 Au-Ag vein-type deposits, 26, 84, 85, 88-92, 94, 95, 98, 100, 103, 124, 126, 128130, 132, 134, 140, 142, 143, 145-147, 149, 151, 153, 156, 158, 159, 172-175, 177, 191,200, 202, 203, 206-209, 211, 237-239, 240, 242, 251,253, 261,266, 315, 325, 330 auriferous vein, 251,252, 254 base-metal vein-type deposits, 7, 84, 88, 90, 91, 94, 95, 98, 124, 127, 129, 142, 143, 145, 149, 150, 158, 171, 176-179, 183, 203, 234, 318, 326, 419, 449 Cu-Pb-Zn vein, 84, 114, 117, 206 Cu-Pb-Zn vein-type deposit, 84, 114, 116, 118, 136, 146, 419 epithermal vein-type deposits, 1, 6, 7, 14, 27, 83-85, 98, 100, 113, 124, 135, 142, 145, 156, 158, 173, 180, 182, 203, 207, 225, 235,237, 244, 324, 325, 327, 328, 352, 353, 449, 450 Neogene vein-type deposits, 116, 117, 124, 125, 146 polymetallic vein-type deposits, 15, 172, 223, 224, 243, 250, 350, 353, 437, 451 vein pattern, 212 vein system, 187, 19 I, 192 vein-type deposits, 2, 4-6, 14, 25, 26, 54, 55, 84-86, 88-95, 98-100, 103, 114, 116, 118, 124, 126-130, 132, 134, 136-138, 140, 142, 143, 145-147, 149, 151, 153, 156, 158, 159, 172-175, 177, 180, 181, 191,200, 202, 203, 206-209, 211,212, 223, 225, 228, 231,237-240, 242, 243, 246, 248-251,253, 256, 258, 260, 261, 266, 315, 324, 325, 327, 330, 351,393, 419, 437, 438, 449-451 velocity of fluid, 70 volcanic activity, 4, 204, 205, 265, 311, 331, 332, 371,390, 407-410, 439 volcanic front, 6, 19, 312, 313, 324 volcanic gas, 81, 100, 107, 110-113, 123, 158, 264, 313, 319, 331, 364, 413, 416, 417, 421-423, 437, 439, 442, 451,452

463

Subject Index volcanism, 5, 15, 18, 19, 87, 205, 206, 213, 228, 331-333, 375, 417, 433, 449, 452 bimodal volcanism, 19, 87, 226, 231, 333 submarine volcanism, 212, 378, 407, 417

225, 450,

Yellow ore, 7, 15, 20, 21, 23, 40, 53, 55, 56, 67, 74, 75

332,

zeolites, 1, 5, 30, 33, 36, 39, 95, 98, 99, 101, 105, 106, 110, 111, 123, 136, 163, 164, 318, 379, 418 Zn, 2, 4, 6-9, 12-14, 20, 25, 30, 40, 49, 50, 76, 77, 80, 88-91, 94, 96, 97, 137, 142, 148, 149, 156, 157, 161, 179, 180, 182, 206, 215, 219, 231, 241, 243, 245, 311, 313, 316, 320-325, 327, 339, 342, 353, 359, 361-365, 378, 385-389, 392, 419, 426, 449-451 Cu-Pb-Zn vein-type deposit, 84, 114, 116, 118, 136, 146, 419 sphalerite, 15, 20, 23, 25, 26, 40, 45, 47, 49, 52, 53, 67, 74, 75, 88, 90-92, 94, 95, 103, 116, 124, 126, 128-131, 134, 135, 137-140, 142, 161-167, 171, 176, 187, 201,233, 236, 238, 240-247, 251,252, 254, 258, 261,312, 318, 321,327, 333, 334, 337, 339-341,361,368, 379-381, 383, 387, 388, 391,392, 394, 395 ZnCI2, 49, 142, 172, 201 ZnS, 25, 49, 92, 94, 142, 172, 201 zonal sequence of alteration minerals, 328 zoogeographic history, 436

413,

W, 2, 4, 56, 63, 73, 76, 80, 88, 89, 143, 175, 178, 180, 181, 198, 207, 224, 231, 241, 243, 297, 316, 325, 365, 410, 415, 436, 450 Wairakei, 110, 111, 119, 199, 297, 328, 422 wairakite, 95, 99, 104, 105, 111, 123, 136, 295, 302, 310, 321,418, 420 water/rock ratio, 33, 34, 59, 60, 82, 114, 119, 120, 122, 143, 155, 158, 177, 191, 328, 359, 360, 373 weathering, 417, 421-423, 431-433, 440, 451,452 wurtzite, 66, 312, 335, 337, 339-341, 350, 361, 367, 368, 380 xenothermal-type deposits, 231 Yatani, 12, 28, 88, 89, 94, 100, 113, 115, 127, 130, 131, 137, 143, 149, 151, 154, 156, 157, 160, 162, 163, 172, 175, 181, 203, 222, 224, 242, 243

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