ZEOLITES IN FISSURES OF

CRYSTALLINE BASEMENT ROCKS

INAUGURALDISSERTATION

zur Erlangung des Doktorgrades der Fakultät für Chemie, Pharmazie und Geowissenschaften der Albert-Ludwigs-Universität Freiburg im Breisgau

vorgelegt von

TOBIAS WEISENBERGER aus Emmendingen

2009

Vorsitzender des Promotionsausschusses: Prof. Dr. Rolf Schubert Referent: Prof. Dr. Kurt Bucher Korreferent: Prof. Dr. Reto Gieré Tag des Promotionsbeschlusses: 9.. Juli 2009

Saint Barbara statue in the Arvigo quarry– Patron saint of geologists and firemen Saints day: 4th December

Der Tag der heiligen Barbara! - Feierlich stehen sie alle da, die Männer, die aus des Berges Nacht - das schwarze Gestein zu Tage gebracht, das dort gelegen seit Urweltzeit; - bald wird es vom roten Feuer gefreit. Feierlich stehen sie alle da.- Es ist 4. Dezember: St. Barbara!

Du Schutzpatronin, St. Barbara, - Im Schmucke treten sie alle dir nah'; An dem Tschako wiegt sich die schwarze Feder, - Schwarz ist ja alles, Anzug und Leder. Dort sind die weißen, Musik trägt rot. - In Ordnung und Würde, wie nach Gebot beginnt der Zug, und wer ihn sah',- weiß, es ist heute St. Barbara!

Zurück von der Kirche St. Barbara. - Und es geschieht, was immer geschah, Musik spielt lustig, die Federn winken, - in Oberschlesien will man auch trinken, sorglos sich freuen, den Tag genießen, - wen sollte das heitere Volk verdrießen? Und es geschieht, was immer geschah! - Nur einmal im Jahr ist St. Barbara!

St. Barbara; Poem after Käthe Gutwein

ABSTRACT I

ABSTRACT

The goal of the thesis is to study the occurrences and formation of zeolites hosted in crystalline basement rocks. The low-grade fissure mineral assemblages including zeolites are the key to the appreciation of water-rock interaction in hydrothermal and geothermal systems at relatively low temperatures (< 250 °C) located in granites and gneisses of the crystalline basement. Extensive work is done on zeolite occurrences in sedimentary rocks often pyroclastic origin and volcanic rocks, whereas elements necessary for zeolite formation derive from primary glass of from feldspar. In contrast the formation of zeolites in granites and gneisses is poorly studied and no systematic of evaluation and spatial distribution are carried out either chemical studies on zeolites or formation consideration are done.

Therefore a systematic evaluation of zeolites in the Central Swiss Alps is presented. Ca-zeolites occur in various assemblages in late fissures and fractures in granites and gneisses. The systematic study of zeolite samples showed that the majority of finds originate from three regions particularity rich in zeolite-bearing fissures: (1) in the central and eastern part of the Aar- and Gotthard Massif, including the Gotthard road tunnel and the Gotthard-NEAT tunnel, (2) Gibelsbach/Fiesch, in a fissure breccia between Aar Massif and Permian sediments, and (3) in Penninic gneisses of the Simano nappe at Arvigo (Val Calanca). The excavation of tunnels in the Aar- and Gotthard massif give an excellent overview of zeolite frequency in Alpine fissures, whereas 32 % (Gotthard NEAT) and 18 % (Gotthard road tunnel) of all fissures are filled with zeolites. The number of different zeolites is limited to 6 species: laumontite, stilbite and scolecite are abundant and common, whereas heulandite, chabazite and epistilbite occur occasionally. Ca is the dominant extra- framework cations, with minor K and Na. Heulandite and chabazite additionally contain Sr up to 29 and 10 mole%, respectively. Na and K content of zeolites tends to increase during growth as a result of systematic changes in fluid composition and/or temperature. The K enrichment of stilbite found in surface outcrops compare to stilbite in the subsurface may indicate late cation exchange during interaction with surface water. Texture data, relative age sequences derived from fissure assemblages and equilibrium calculations shows that the Ca-dominated zeolites precipitated from fluid with decreasing temperature in the order (old to young = hot to cold): scolecite, ABSTRACT II laumontite, heulandite, chabazite and stilbite. The components necessary for zeolite formation are derived from dissolving primary granite and gneiss minerals. The nature of these minerals depends on the metamorphic history of the host rock. Zeolites in the Aar Massif derived from the dissolution of epidote or calcite and albite that were originally formed during Alpine greenschist metamorphism. Whereas albitization of plagioclase in higher grade rocks releases the necessary components for zeolite formation, a process that is accompanied by a distinct porosity increase. Zeolite fissures occur in the zone where fluid inclusions in earlier formed quartz contain H2O dominated fluids. This is consistent with equilibrium calculations that predict a low CO2 tolerance of zeolite assemblages particularly at low temperature. Pressure decrease along the uplift and exhumation can increase zeolite stability. The major zeolite forming reaction consumes calcite and albite; it increases pH and the total of dissolved solids. The produced Na2CO3 waters are in accord with reported deep groundwater (thermal water) in the continental crust, which are typically oversaturated with respect to Ca-zeolites.

A detailed local study of the mineralogical, chemical and petrological evolution of crystalline basement rocks in Arvigo was performed to assess information about the evolution of fluid-rock interaction during uplift of the Alpine orogen. The Arvigo fissures contain the assemblage epidote, prehnite, chlorite and various species of zeolites. Fluid rock interaction takes place along a retrograde exhumation path which is characterized with decreasing temperature by: (1) coexisting prehnite/epidote, that reveals temperature conditions of 330 – 380 °C, (2) chlorite formation at temperature of 333 ± 32 °C and (3) formation of zeolites <250 °C. The formation of secondary minerals is related to the hydrothermal replacement reaction during albitization and chloritization that releases components for the formation of Ca-Al silicates and form a distinct reaction front. The fluid-rock interaction is associated with a depletion of

Al2O3, SiO2, CaO, Fe2O3 and K2O in the altered wall rock. The reaction is associated with an increase in porosity up to 14.2 ± 2.2 %, caused by the volume decrease during albitization and the removal of chlorite. The propagation of the sharp reaction front through the gneiss matrix occurred via a dissolution-reprecipitation mechanism. Zeolite formation is tied to the plagioclase alteration reaction in the rock matrix, which releases components for zeolite formation to a CO2-poor, alkaline aqueous fluid. ABSTRACT III

A combined study of 40Ar/39Ar age dating, apatite fission track (FT) and chemical characterization of tunnel and surface samples are present to carry out the position of low-temperature water-rock interaction in respect to the Alpine history. Apatite FT analysis yields an exhumation rate of 0.45 mm a-1, a cooling rate of 13 °C Ma-1 and a geothermal gradient of 28 °C km-1. Combining these with the 40Ar/39Ar plateau age for apophyllite of ∼2 Ma, a minimum formation temperature and depth of 70 °C and 2800 m, respectively can be assumed. Temperature-time evolution of fissures in the Aar Massif and thermodynamic mineral evolution indicate that laumontite were formed between 7 and 2 Ma before present at temperatures between 150 and 70 °C.

ACKNOWLEDGMENTS IV

ACKNOWLEDGMENTS

It is a great pleasure for me to thank the many people who made this thesis possible.

I am most thankful to my advisor Prof. Dr. Kurt Bucher. I thank him for awarding the topic of this thesis and an outstanding supervision. I am glad that he gave me the opportunity to continue my research interests on zeolites that I experience during my diploma thesis. I always appreciated the discussion and constructive criticism with him. A special appreciation has to be mentioned that he gave me the freedom to develop and follow my own ideas. He enhanced me to educate myself during numerous DMG workshops and teached me to deal with thermodynamic calculations, which was not even easy with me. During theses years Kurt must have used up a lifetime supply of red ink pens to teach me geological common sense. This thesis would not have been possible without the great advise and trust in me. Thank you!

Also, I want to thank Prof. Dr. Reto Gieré for the additional supervision, numerous discussions and the takeover of the co-referee.

I want to thank my parents, who always gave me the liberty to follow my interest and supported me during my education in a loving environment.

A special thank to all the people who supported me during my search for Alpine zeolites and supplied samples: Peter Amacher, mineral representative of the NEAT Amsteg-Sedrun section who provided high-quality minerals specimen and who was always easy to contact for discussion. Beda Hoffmann and Peter Vollenweider from the Swiss Natural History Museum in Bern, giving me the possibility to study their mineral collection and their encourage during my work in the “dungeon”. Giovanni Polti and Alfredo Polti SA for permission to do field work in the active quarry in Arvigo, especially Luigi who took care about me during blasting, even we conduct our conversion by signs, due to my lack in Italian language.

I appreciate all the help and support that I get by using technical equipment in external research institutes: Prof. Dr. Stefan Graeser form the Mineralogical Institute Basel, who provided me the possibility to use the FTIR instrument; Dr. Egbert Keller from the Crystallographic Institute Freiburg, who guided me through DSC-TGA measurements and Andreas Leemann from the Swiss Federal Laboratories for ACKNOWLEDGMENTS V

Materials Testing and Research for impregnation of rock samples. Roelant van der Lelij from the Department of Mineralogy in Geneva for the apophyllite dating and helpful discussion and PD Dr. Meinert Rahn for helping with the apatite fission track analysis and the always profitable conversations.

I wish to thank Zeng Lu, Fleurice, Siggi, Zhou Wei, Hiltrud and Duy Anh Dao for their always open doors, where I find a sympathetic ear for discussion.

The Friedrich Rinne Stiftung of the Albert-Ludwigs-University, Freiburg for the financial support.

Last but not least I want to thank Simon and Rune. I am very happy to call them my friends. We had a wonderful time during our diploma thesis and they encouraged me during my PhD study whenever I needed them. Thanks Rune for the long insightful phone calls, the high email exchange rate and help when I need a cheer-up. Thanks Simon for the friendship during the last years and all the help, which contribute to the succeed of this thesis – feel free to ask me if you need a red ink pen again! TABLE OF CONTENTS VI

TABLE OF CONTENTS

ABSTRACT I ACKNOWLEDGMENTS IV TABLE OF CONTENTS VI

1. INTRODUCTION 1 1.1. LAYOUT OF THE THESIS 1 1.2. MOTIVATION OF THIS STUDY 3 1.3. ZEOLITES 4 1.4. ZEOLITE STRUCTURE AND CHEMISTRY 6 1.5. ZEOLITE OCCURRENCES, ZEOLITE FACIES AND ZEOLITE ZONES 7 1.6. ZEOLITE STABILITY 11 1.7. GEOLOGICAL SETTING 12 1.8. REFERENCES 15

2. ZEOLITES IN FISSURES OF GRANITES AND GNEISSES OF THE CENTRAL ALPS 20 2.1. ABSTRACT 21 2.2. INTRODUCTION 22 2.3. GEOLOGICAL SETTING 25 2.3.1. Metamorphic conditions during Alpine orogenesis 26 2.4. SAMPLING AND ANALYTIC METHODS 29 2.5. ZEOLITES IN THE CENTRAL ALPS 30 2.5.1. Spatial distribution 31 2.5.2. Field occurrences 35 2.5.2.1. General features 35 2.5.2.2. Aar Massif/Gotthard NEAT tunnel 43 2.5.2.3. Gotthard massif/ Gotthard road tunnel 45 2.5.2.4. Gibelsbach/Fiesch 46 2.5.2.5. Arvigo/Val Calanca 47 2.5.3. Mineralogy and crystal chemistry of zeolites and associated minerals 48 2.5.3.1. Chabazite-Ca 48 2.5.3.2. Heulandite-Ca 51 2.5.3.3. Laumontite 52 2.5.3.4. Scolecite 53 TABLE OF CONTENTS VII

2.5.3.5. Stilbite/Stellerite 54 2.6. DISCUSSION 55 2.6.1. Reactions and processes of zeolite formation 58 2.6.2. Assemblage stability and phase relationships involving zeolites 61 2.6.3. Fluid composition 64 2.7. CONCLUSIONS 67 2.8. ACKNOWLEDGMENTS 68 2.9. REFERENCES 68

3. POROSITY EVOLUTION, MASS TRANSFER AND PETROLOGICAL EVOLUTION DURING LOW TEMPERATURE WATER-ROCK INTERACTION IN GNEISSES OF THE SIMANO NAPPE - ARVIGO, VAL CALANCA, GRISONS, SWITZERLAND 76 3.1. ABSTRACT 77 3.2. INTRODUCTION 77 3.3. GEOLOGICAL SETTING 79 3.4. PREVIOUS WORK 82 3.5. SAMPLING AND ANALYTIC METHODS 83 3.6. RESULTS 84 3.6.1. Petrography 84 3.6.1.1. Unaltered rock 86 3.6.1.2. Altered rock 86 3.6.1.3. Fissure minerals 87 3.6.1.4. Changes in modal mineralogy 87 3.6.2. Mineralogy and mineral chemistry 87 3.6.2.1. Plagioclase and its alteration products 92 3.6.2.2. Biotite-chlorite 95 3.6.2.3. Muscovite 95 3.6.2.4. K-feldspar 95 3.6.2.5. Quartz 96 3.6.2.6. Epdiote 96 3.6.2.7. Prehnite 97 3.6.2.8. Zeolites 99 3.6.3. Porosity 102 3.6.4. Whole rock geochemistry and mass changes 103 3.7. DISCUSSION 105 3.7.1. Mineral reactions 105 TABLE OF CONTENTS VIII

3.7.2. Mass changes and element mobility 107 3.7.3. Mineral stability and mineral equilibria 110 3.7.3.1. Prehnite and epidote 110 3.7.3.2. Chlorite 112 3.7.3.3. Zeolites 113 3.7.4. Mineral evolution 117 3.7.5. Fluid accessibility and composition 120 3.8. CONCLUSION 121 3.9. ACKNOWLEDGMENTS 123 3.10. REFERENCES 124

4. TIMING AND MINERAL EVOLUTION DURING LOW- TEMPERATURE FLUID-ROCK INTERACTION ON UPPER CRUSTAL LEVEL: 40Ar/39Ar APOPHYLLITE-(KF) DATING AND APATITE FISSION TRACK ANALYSIS ON ALPINE FISSURES (CENTRAL ALPS/SWITZERLAND) 132 4.1. ABSTRACT 133 4.2. INTRODUCTION 133 4.3. GEOLOGICAL SETTING 135 4.4. MATERIAL AND METHODS 137 4.4.1. Analytic 137 4.4.2. Samples 140 4.5. RESULT 140 4.5.1. Petrography and geochemistry 140 4.5.2. Mineralogy and geochemistry 143 4.5.2.1. Laumontite 143 4.5.2.2. Apophyllite-(KF) 144 4.5.3. Ar/Ar age 146 4.5.4. Apatite FT analysis 147 4.6. DISCUSSION 149 4.6.1. Mineral reaction 149 4.6.2. Depth and temperature estimation 151 4.6.3. Thermodynamic approach 152 4.6.4. Alpine history 154 4.7. CONCLUSION 155 4.8. ACKNOWLEDGMENTS 156 4.9. REFERENCES 157

TABLE OF CONTENTS IX

APPENDIX I OWN CONTRIBUTION i II PUBLICATIONS ii III CURRICULUM VITAE v

INTRODUCTION 1

1. INTRODUCTION

1.1. LAYOUT OF THE THESIS

The thesis is divided into 4 chapters. The 1st Chapter gives an overview and structure of the thesis and the research interest that is associated with this thesis. Additionally general information about zeolites, zeolite occurrences and the geological setting of the working area are given to introduce into the topic of natural zeolites.

Chapter 2, 3 and 4 are separated chapters covering each a manuscript and therefore analytic methods are described in each chapter, including measuring methods, measuring conditions and standards that were used. The chapters are arranged to start with a general overview and evaluation of zeolite occurrence and formation in crystalline rocks and continue with two detailed local studies to understand the process of zeolite formation in crystalline basement rocks and the relation of zeolite formation in respect to the Alpine history (Table 1.1). At the beginning of each chapter an overview of my contribution to finalize each manuscript is given.

Chapter 2 is concerned with the documentation and compilation of all known zeolite occurrences in the Central Swiss Alps to get information about the spatial distribution. Zeolites hosted in granites and gneisses are chemically characterized to get information about chemical changes during growth as well as information about chemical variations in different settings within the Central Alps. Petrographic observation yields information about zeolite-forming reactions that are formulated in this chapter. Thermodynamic phase-diagram-modeling to approach the conditions at which zeolites are formed is used to discuss the zeolite appearance in respect to fluid composition variation in the Swiss Alps.

This work was presented at the “Zeolite06 Conference” in Socorro, USA in July 2006 and at the annual meeting of the “Deutsche Mineralogische Gesellschaft” in Hannover, Germany in September 2006. The manuscript was submitted to Journal of Metamorphic Geology (March 2009) as:

Weisenberger T. and Bucher K.: ZEOLITES IN FISSURES OF GRANITES AND GNEISSES

OF THE CENTRAL ALPS. Journal of Metamorphic Geology INTRODUCTION 2

Chapter 3 presents a detailed petrographic, mineralogical and geochemical study of the zeolite locality Arvigo. The active quarry gives good access to many factures filled with zeolites, which are characterized by a leaching zone from which elements for secondary mineral formation are derived. Mass balance calculation, element mobility and porosity measurements were done in combination with thermodynamic phase diagram modeling to show the approach and applicability of thermodynamic modeling at such problems, including PT-path modeling, mineral and porosity evolution and relation of fluid composition with respect to the zeolite formation.

This work was presented at the “17th Goldschmidt Conference” in Cologne, Germany in August 2007 and at the “Swiss Geoscience Meeting” in Geneva, Switzerland in November 2007. The manuscript will be shortly submitted to Contributions to Mineralogy and Petrology as:

Weisenberger T. and Bucher K.: POROSITY EVOLUTION, MASS TRANSFER AND

PETROLOGICAL EVOLUTION DURING LOW TEMPERATURE WATER-ROCK INTERACTION IN

GNEISSES OF THE SIMANO NAPPE - ARVIGO, VAL CALANCA, GRISONS, SWITZERLAND. Contributions to Mineralogy and Petrology

Chapter 4 is a combined study of apophyllite Ar/Ar age dating, apatite fission track analysis, chemical phase analysis and thermodynamic modeling in order to constrain the evolution of zeolites with respect to the Alpine history. This yields information about minimum temperature conditions of laumontite formation. Apophyllite age dating was performed by R. van der Lelij at the Mineralogical department of the University Geneva, Switzerland. Apatite fission track analyses were carried out by PD Dr. M. Rahn at the Mineralogical department of the University Basel, Switzerland.

Parts of this work were presented at the “33rd International Geological Congress” in Oslo, Norway in August 2008 and at the annual meeting of the “Deutsche Mineralogische Gesellschaft” in Berlin, Germany in September 2008. The manuscript will be shortly submitted to Mineralogical Magazine as: INTRODUCTION 3

Weisenberger T., Rahn M., van der Lelij R., R. Spikings and Bucher K.: TIMING

AND MINERAL EVOLUTION DURING LOW-TEMPERATURE FLUID-ROCK INTERACTION ON 40 39 UPPER CRUSTAL LEVEL: AR/ AR APOPHYLLITE-(KF) DATING AND APATITE FISSION TRACK

ANALYSIS ON ALPINE FISSURES (CENTRAL ALPS/SWITZERLAND). Mineralogical Magazine

Table 1.1: Overview of thesis and major topics and major questions addressed to each chapter. Zeolites in fissures of crystalline basement rocks

Chapter 2: Zeolites in fissures of Chapter 3: Porosity evolution, mass Chapter 4: Timing and mineral granites and gneisses of the Central transfer and petrological evolution evolution during low-temperature Alps during low temperature water-rock fluid-rock interaction on upper interaction in gneisses of the Simano 40 39 crustal level: Ar/ Ar apophyllite- Nappe - Arvigo, Val Calanca, (KF) dating and apatite fission track Grisons, Switzerland analysis on Alpine fissures (Central Alps/Switzerland)

• general overview and compilation • detailed study of a zeolite locality • detailed study of a zeolite locality of all zeolite localities in the to assess the formation of Ca-Al to assess the formation of Ca-Al Central Swiss Alps silicates during uplift silicates during exhumation • what zeolite species occur in • temporal evolution of secondary • timing of zeolite formation crystalline basement rocks? minerals • depth of zeolite formation • spatial distribution of zeolites • mineral reactions during water- • chemical change during formation • chemical characterization of rock interaction of secondary minerals zeolites • mass transfer during water-rock • relation of zeolite formation with • zeolite forming reactions interaction respect to the Alpine history • what factors control the zeolite • porosity evolution during water- • estimation of laumontite forming formation? rock interaction temperature • thermodynamic phase-diagram • fluid-composition in relation to calculation approaching physical the formation of zeolites and chemical conditions • thermodynamic phase-diagram calculation to approach physical and chemical conditions

1.2. MOTIVATION OF THIS STUDY

Already during my diploma thesis in 2004 and 2005 on zeolites in Iceland, one of the worldwide famous zeolite localities (e.g. Walker 1959, 1960) I got attracted to the zeolite mineral group. Simultaneous ongoing excavation of the New Gotthard Railway Base tunnel (NEAT) supplied a large amount of zeolites specimen during the drilling period. This finds, the possibility to get access to these specimen and the current research interest on deep continental fluids triggered Kurt Bucher, my adviser and me to found this PhD project about the zeolite formation in crystalline basement rocks that is driven by fluid-rock interaction. INTRODUCTION 4

But what makes zeolites in crystalline basement rocks so interesting?

Although zeolites are known in fissures and gashes of the crystalline basement from mineral collectors since about 150 years (e.g. Kenngott 1866; Parker 1922; Niggli et al. 1940; Huber 1943; Sigrist 1947; Stalder et al. 1998), they did not affect the interest of the scientific research community. So far, previous publications about zeolite occurrences and their genesis in the central Swiss Alps and other areas of crystalline basement rocks are limited, which easily can be count on one hand (Armbruster et al. 1996; Freiberger et al. 2001; Fujimoto et al. 2001; Ciesielczuk and Janeczek 2004). Considering the latest special edition on natural zeolites (Bish and Ming 2001), only two short notes were made on zeolites hosted in granites and gneisses and so far no research was done on the chemical characterization and spatial distribution of those zeolites, like in other environments. This lack of research interest may be related to the economically non-profitable occurrences of zeolites hosted in basement rocks compared to the well-known and widely used deposits of natural zeolites from zeolitized volcanic tuffs and sediments.

Nevertheless, geochemical studies of deep continental fluids suggest that many crystalline basement aquifers are oversaturated with respect to zeolites (e.g. Stober and Bucher 1999; Bucher and Stober 2000). Therefore zeolites play a role when considering mass transfer, porosity and permeability of these aquifers, which are important research areas in relation to geothermal energy production as well as the problematic storage of nuclear waste in crystalline basement rocks.

1.3. ZEOLITES

Zeolites are among the most common products of chemical interaction between groundwaters and the Earth’s crust during diagenesis and low-grade metamorphism (e.g. Bish and Ming 2001). Zeolite minerals occur in low temperature (<250 °C), low pressure (<200 MPa), water saturated environments. The required amount of silica, alumina, and alkali and alkali-earth cations necessary for the formation of zeolites is commonly derived from dissolution of volcanic glass and primary phases. INTRODUCTION 5

Zeolites are tectosilicates characterized by an open framework structure of Si and Al surrounding channels of ~2-10 Å in size which contain molecular water and charge-balancing cations of alkali and alkali-earth metals (e.g. Neuhoff et al. 2000; Armbruster and Gunter 2001). Their unique and distinct crystal structures result in a large molar volume, high cation-exchange capacities, and molecular sieve capabilities (e.g. Gottardi and Galli 1985; Bish and Ming 2001). These properties lead to widespread industrial application in water softening, catalysis, water and wastewater treatment, agriculture, nuclear waste storage, heating and refrigeration and construction industry (e.g. Murphy et al. 1978; Kalló 2001; Ming and Allen 2001; Tchernev 2001; Hauri 2006).

During the past decades with the onset of analyses by electron microprobe, thousands of zeolite crystals have been analyzed showing a wide compositional range, and several new minerals with framework structures were discovered. This was achieved by the advent of automated single crystal X-ray diffractometers, resulting in much more detail concerning zeolite framework structures. With impeding nomenclature problems, the International Mineralogical Association’s Commission on New Minerals and Mineral Names assigned a subcommittee to review all minerals, and proposes a new definition and a system of nomenclature of zeolites.

The report of the International Mineralogical Association, Commission on New Minerals and Mineral Names, contains the following definition:

“A zeolite mineral is a crystalline substance with a structure characterized by a framework of linked tetrahedra, each consisting of four O atoms surrounding a cation. This framework contains open cavities in the form of

channels and cages. These are usually occupied by H2O molecules and extra- framework cations that are commonly exchangeable. The channels are large enough to allow the passage of guest species. In the hydrated phases, dehydration occurs at temperatures mostly below about 400°C and is largely reversible. The framework may be interrupted by (OH,F) groups; these occupy a tetrahedron apex, which is not shared with adjacent tetrahedra”

(COOMBS et al. 1998).

INTRODUCTION 6

Dehydration properties from Cronstedt’s original definition (1756) and a framework structure from Hey’s 1930 definition are retained. However, following the new definition (COOMBS et al. 1998), the framework needs not to be only aluminosilicate. Beryllosilicate, aluminophosphate, and a few similar compositions are allowed by definition.

1.4. ZEOLITE STRUCTURE AND CHEMISTRY

The basic feature of all zeolite structures is an aluminosilicate framework

(tectosilicate) composed of (Si,Al)O4 tetrahedra, each oxygen of which is shared between two tetrahedrons (Armbruster and Gunter 2001). The net negative charge on the tectosilicate framework is balanced by the incorporation of cations (extra- framework cations) in cages or channels. In most cases Ca2+, Na+ or K+ and less frequently Li+, Mg2+, Sr2+ and Ba2+ are situated in cavities within the framework structures. This feature can also be observed in feldspar and feldspathoid minerals. But in contrast to feldspar and feldspathoid minerals the zeolite aluminosilicate framework contains open cavities and open channels (i.e. they have lower densities) through which ions can be either extracted or introduced in the structure (Armbruster and Gunter 2001).

Their compositions are represented by the structural formula (1):

+z +3 (A )y/z(B )y(Si)xO2(x+y) · nH2O (1)

Where A represents extra-framework cations (such as Na+, K+, Ca2+, Ba2+, Sr2+, Mg2+ and Fe2+), B are tetrahedral coordinated trivalent cations in the zeolite framework (Al3+ and Fe3+), z is the charge of the extra-framework cations, n is the number of moles of extra-framework molecular water, and x and y are the stoichiometric coefficients for trivalent cations and Si4+ in tetrahedral sites, respectively. The quantities y/z and 2(x+y) represent the stoichiometries of the extra-framework cations and framework oxygens, respectively, necessary for maintaining charge balance in the tectosilicate lattices of zeolites (Armbruster and Gunter 2001). INTRODUCTION 7

An additional feature, which separates zeolites still further from the feldspar and feldspathoid minerals, is the presence of water molecules within the structural channels. These are relatively loosely bound to the framework and cations, and like the cations they can be removed and replaced without disrupting framework bonds (Deer et al. 2004).

Three types of solid solutions in zeolites are consistent with the stoichiometry of equation (1). These solutions are not strictly coupled and can occur independently from other substitutions as long as charge balance is maintained (Neuhoff et al. 2000). The first of these is the solid solution within the tetrahedral sites. Tetrahedral substitution of Si4+ and Al3+ observed in zeolites is highly variable, whereas the substitution Fe3+ for Si4+ or Al3+ is limited. Secondly, solid solutions among extra- framework cations are often quite extensive, as evidenced by the large ion-exchange capacities of some zeolites (e.g. Colella 1996). Total extra-framework ion charge is necessarily a function of Al3+ and Fe3+ content. Zeolites with high Si/Al ratios commonly are richer in monovalent extra-framework cations than are more aluminous samples of the same species. Twice as many monovalent ions as divalent ions are necessary to compensate for charge imbalances caused by Al3+ in the framework, and the additional monovalent ions often occlude H2O molecules present in isostructural zeolites with divalent extra-framework cations (e.g. natrolite and scolecite, Ca- heulandite and Na-heulandite). The third type of solid solution in zeolites is the variation in water content, which a consequence of the loose bounding nature of molecular water in zeolites, whereas the total water content is a sensitive function of temperature, total pressure and the partial hydrostatic pressure (Neuhoff et al. 2000).

1.5. ZEOLITE OCCURRENCES, ZEOLITE FACIES AND ZEOLITE ZONES

Zeolites are formed during reaction of aqueous fluids and rocks in several different geological environments: Most zeolite occurrences formed during diagenetic processes in sedimentary rocks (including volcanoclastic deposits) which can be grouped into several types of geological environments or hydrological systems (Hay 1966, 1977; Hay and Sheppard 1977; Surdam 1977; Gottardi 1989; Hay and Sheppard INTRODUCTION 8

2001), like (1) hydrologically open systems (e.g. Hay and Sheppard 2001), (2) hydrologically closed systems (e.g. Langella 2001), (3) soil and surficial deposits (e.g. Ming and Mumpton 1989), (4) deep marine sediments (e.g. Boles and Coombs 1977) and (5) marine sediments from arc-source terrains (e.g. Boles and Coombs 1977) (Fig. 1.1).

Fig. 1.1: Schematic diagrams showing patterns of zeolite zoning in silicic tephra deposits in various genetic environments (modified after Hay 1977 and Neuhoff et al. 2000). (a) Plan view and cross- section view of zeolites formed in closed hydrological systems (e.g. Playa lakes). The mineral distribution in these systems reflects an increase in salinity during fluid evaporation. (b) Cross-section view of zeolites formed in an open hydrological system. Mineral distribution is taken to reflect pH changes during progressive interaction with the host rock. (c) Cross-section view of mineral distribution in hydrothermal systems. Mineral distribution reflects a temperature gradient during alteration. (d) Cross-section view of mineral distribution under ongoing burial of a stratigraphic sequence. Mineral distribution reflects a temperature gradient during burial.

Almost every known zeolite occurs in cavities of volcanic lava flows (e.g. Tertiary lavas of Iceland, Deccan Plateau, India). These zeolites are formed either during burial metamorphism of the lava pile (e.g. Walker 1960; Neuhoff 1999), during INTRODUCTION 9 hydrothermal alteration of continental basalts (e.g. Walker 1963), or during diagenesis in areas of high heat flow caused by active geothermal systems (e.g. Kristmannsdóttir and Tómasson 1978; Weisenberger and Selbekk 2008).

Zeolites as products of hydrothermal crystallization are generally known from active geothermal systems associated with volcanic rocks. Very little work has been published on zeolite occurrences related to late stage crystallization of pegmatitic bodies (e.g. Orlandit and Scortecci 1985), in hydrothermal ore veins (Deer et al., 2004), as alteration along fault plains (e.g. Vincent and Ehlig 1988), and in hydrothermal fractures and veins in granites and gneisses (e.g. Borchardt et al. 1990; Borchardt and Emmermann 1993; Armbruster et al. 1996; Bish and Ming 2001; Freiberger et al. 2001; Fujimoto et al. 2001).

Fig. 1.2: Temperature-pressure diagram showing the metamorphic facies, including the field of zeolite facies which represents the lowest metamorphic facies (modified after Winter 2001).

The importance of zeolites as low temperature alteration phases in the Earth’s crust was noted early in the history of metamorphic petrology. However, Eskola (1939) rejected the concept of a metamorphic zeolite facies. The issue was revisited in the 1950s when Rengarten (1950) proposed a “geochemical zeolite facies” in which zeolite assemblages represent alteration of sediments in contact with aqueous solutions of unusual composition. The benchmark papers of Fyfe et al. (1958) and INTRODUCTION 10

Coombs et al. (1959) arise the present concept of zeolite facies as a necessary intermediate metamorphic grade between diagenesis and greenschist facies. Since then several zeolite facies distributions in various genetic environments were described. Nowadays the zeolite facies is accepted as an intermediate facies between the prehnite-pumpellyite facies and diagenesis (Fig. 1.2, 1.3)

The distribution of individual zeolite species within diagenetically altered or metamorphosed sediments and volcanic rocks is commonly characterized by isograds of first appearance of one or more zeolites bounding spatially restricted zones (e.g. Coombs et al. 1959; Hay 1977; Kristmannsdóttir and Tómasson 1978).

Fig. 1.3: Temperature-pressure ranges of zeolite-forming environments. (adapted from Deer et al. 2004). Solid curves are experimentally determined stability limits of selected zeolites: (1) epidote + quartz + H2O = laumontite + prehnite, at low-pressure end and epidote + chlorite + quartz + H2O = laumontite + pumpellyite at high-pressure end (2) laumontite + quartz + H2O = heulandite (Cho et al. 1987). Dashed line represents a retrograde PT-path in Alpine fissures, determined by fluid inclusions in fissure quartz (Mullis et al. 1994).

The golden spike for zeolite facies mineralization was done by the British geologist George Walker (1959, 1960, 1963). In his work, Walker made a careful study of Iceland´s amygdules and mapped the zeolite distribution in Tertiary lavas of Eastern Iceland. He recognized a systematic depth variation in the zeolite distribution of the lava sequences. Similar observations were done during the same period by Coombs et al. (1959) on low-grade meta-sediments in New Zealand. INTRODUCTION 11

With their pioneering work Walker (1959, 1960) and Coombs et al. (1959) identified regionally extensive mineral assemblages that define “depth” controlled “zeolite zones” formed during burial metamorphism that have been found in different environments during the last centuries (e.g. Bish and Ming 2001).

1.6. ZEOLITE STABILITY

The particular zeolite to form depends on any of five factors: (1) temperature, (2) pressure, (3) primary rock composition, (4) fluid composition, and (5) the water to rock ratio. The most important factor of all theses for a particular paragenesis is the composition of the material to be altered and the composition of the altering solution (e.g. Deer et al. 2004).

The effect of pressure on zeolite isograds usually cannot be assessed independently from that of temperature; however, Iijima (1988) has demonstrated that the temperatures corresponding to the isograds are essentially independent of pressure.

Several reactions involving Ca-zeolites have been studied experimentally (e.g. Liou 1971; Thompson 1971; Cho et al. 1987; Frey et al. 1991). In general the maximum temperature and pressure limits of zeolite stability are in agreement with observations on geothermal systems (Kristmannsdóttir and Tómasson 1978; Frey et al. 1991). Above 400°C anorthite is stable relative to wairakite, the Ca-zeolite stable at highest temperature (Frey et al. 1991).

There are large discrepancies between directly measured temperatures at the position of some zeolite boundaries in boreholes and temperature calculated from experimentally based thermodynamic data (Neuhoff et al. 2000). For example, in the North Tejon oil field in California, heulandite and laumontite coexist at around 90 °C (Noh and Boles 1993). This is a much lower temperature than the 240 ˚C derived from equilibrium phase calculation. Neuhoff (1999) has demonstrated that these discrepancies can be attributed for example to variations in structural order-disorder and in different chemical composition between natural and synthetic minerals.

INTRODUCTION 12

1.7. GEOLOGICAL SETTING

The Alps form a part of a Tertiary orogenic belt that stretches from southern Europe to Asia. The Alps formed as a result of the closure of Jurassic to Cretaceous Tethys ocean basins during convergence of the Apulian and European plates (e.g. Trümpy 1960; Frisch 1979; Schmid et al. 2004). An orogenic belt characterized by stacked nappes formed in the Tertiary when the Apulian and European plates collided. The collision caused a complicated tectonic structure and a regional metamorphic overprint. Deeply buried parts of the orogen were later exhumed and uplifted in the late Tertiary (Trümpy 1980) and finally reached the erosion surface.

Fig. 1.4: Simplified geological map of Switzerland (modified after Spicher 1980). The dashed line mark the trail of the Gotthard NEAT tunnel.

INTRODUCTION 13

Zeolites in fissures occur predominantly in rocks that belong to large basement windows exposed in the northern part of the Alps (Fig. 1.4). These so-called “external massifs” of the Alps belongs to the European plate (e.g. Trümpy 1980). The massifs represent parautochthonous units (Pfiffner 1986). Two major basement units are distinguished in the central Swiss Alps: the Aar Massif and the Gotthard Massif. They constitute of pre-Variscan basement, which is partly reworked by the Variscan and Alpine orogenesis. The massifs form a 115 km long and 23 to 40 km wide SW-NE trending outcrop. The large Aar Massif consists of pre-Variscan gneisses, pre- Variscan granitoids, migmatitic granites and gneisses, lower and upper Carboniferous intrusives and Carboniferous volcanics (Abrecht 1994). Many of the prominent high Alpine peaks and the largest glaciers of the Alps are located in the Aar massif.

The Gotthard Massif is located to the south of the Aar Massif and is followed further south by the north Penninic continental nappe stack. It consists of a poly- metamorphic continental basement with Variscan granites. It is separated from the Aar Massif in the north by the narrow Tavetsch Massif and the Mesozoic metasediments of the Urseren zone.

The rocks have been overprinted by Tertiary Alpine metamorphism. The metamorphic grade and Alpine peak metamorphism increases from nearly non- metamorphosed rock units in the north, over greenschist facies rocks in the Aar- and Gotthard Massif region up to amphibolite facies conditions in the Penninic nappes to the south (Labhart 1977; Frey et al. 1980; Frey and Mählmann 1999). In general, metamorphic grade increases within a specific tectonic unit from north (external position) to south (internal position), as well as from the top to the bottom (Frey and Mählmann 1999) (Fig. 1.5).

Deutsch and Steiger (1985) determined an age of 37 Ma for peak metamorphism in the Lepontine Alps, whereas peak metamorphism in the Aar Massif is dated around 25 Ma (Grimsel area, Dempster 1986).

During uplift and erosion brittle deformation structures became dominant once the rocks crossed the ductile-brittle transition zone. Semi-brittle shear zones related to backthrusting, normal faults and the opening of fissures and gashes are typical structures related to the late orogenic deformation phases including extension structures related to uplift. INTRODUCTION 14

The fissures are conductive to hot aqueous fluids, whereas the percolating hot fluids could react with the surrounding rocks along the fissure walls to form secondary fissure minerals.

Fig. 1.5: Schematic sketch of the T-t evolution of tectonic units in the Central Alps in relation to fissure formation and the timing of zeolite growth. (a) The T-t paths of individual tectonic units reflecting an increase of the Alpine peak metamorphism from north to south. The southern units have reached the ductile regime, northern units were deformed brittle. All units reached temperatures above the zeolite window (except for the parautochtonous cover rocks of the Aar massif). During uplift the units returned to the brittle deformation regime and extension fissures formed (b), subsequently zeolite- absent fissure assemblages developed (c) and finally the units entered the zeolite window (d).

Metamorphic conditions derived from fluid inclusion studies on Alpine fissure material represent conditions during various stages of exhumation, decompression and cooling. Mullis et al. (1994) determined minimum conditions for fissure mineral formation ranging from 400-430°C in temperature and fluid pressures from 240 to 380 MPa for the southern Aar Massif (Zinggenstock), for the Gotthard Massif and the northern Lepontine Alps.

An exhumations rate of 0.5 mm a-1 for the Reuss valley (northern Aar Massif) has been proposed by Michalski and Soom (1990) for the past 27 Ma from apatite and zircon fission track data. The corresponding cooling rate of 13°C Ma-1 agrees well with other apatite fission track data that suggest uplift rates of 0.3-0.6 mm a-1 during the last 6-10 Ma (Schaer et al. 1975). Using exhumation rates and trapping temperatures of early fluid inclusions (Mullis 1996) determined the time of the first INTRODUCTION 15 opening of fissures and precipitation of fissure minerals in the Aar Massif (Zinggenstock) and Gotthard Massif (La Fibbia) to 20 to 15 Ma b.p..

The deposition of fissure minerals occured along the cooling and decompression P-T-path after peak metamorphism (Fig. 1.5).

1.8. REFERENCES

Abrecht, J. (1994) Geological units of the Aar massif and their pre-Alpine rock associations: a critical review. Schweizerische Mineralogische und Petrographische Mitteilungen, 74, 5-27. Armbruster, T., and Gunter, M.E. (2001) Crystal structure of natural zeolites. In D.L. Bish, and D.W. Ming, Eds. Natural Zeolites: occurrence, properties, applications, 45, p. 1-68. Mineralogical Society of America, Geochemical Society. Armbruster, T., Kohler, T., Meisel, T., Nägler, T.F., Götzinger, M.A., and Stalder, H.A. (1996) The zeolite, fluorite, quartz assemblage of the fissure at Gibelsbach, Fiesch (Valais, Switzerland): crystal chemistry, REE patterns, and genetic speculations. Schweizerische Mineralogische und Petrographische Mitteilungen, 76, 131-146. Bish, D.L., and Ming, D.W. (2001) Natural Zeolites: occurrence, properties, applications, 45, 654 p. Mineralogical Society of America, Geochemical Society. Boles, J.R., and Coombs, D.S. (1977) Zeolite facies alteration of sandstones in the southland syncline, New Zealand. American Journal of Science, 277, 982-1012. Borchardt, R., and Emmermann, R. (1993) Vein minerals in KTB rocks. In R. Emmenmann, J. Lauterjung, and T. Umsonst, Eds. KTB Report 93-2, p. 481- 488. Project Management of the Continental Deep Drilling of the Federal Republic of Germany in the Geological Survey of Lower Saxony. Borchardt, R., Zulauf, G., Emmermann, R., Hoefs, J., and Simon, K. (1990) Abfolge und Bildungsbedingungen von Sekundärmineralen in der KTB-Vorbohrung. In R. Emmermann, and P. Giese, Eds. KTB Report 90-4, p. 76-88. Projektleitung Kontinentales Tiefbohrprogramm der Bundesrepublik Deutschland im Niedersächsischen Landesamt für Bodenforschung. Bucher, K., and Stober, I. (2000) Hydrochemistry of water in the crystalline basement. In I. Stober, and K. Bucher, Eds. Hydrogeology of Crystalline Rocks, p. 141-175. Kluwer Academic Publishers, Dordrecht. Cho, M., Maruyama, S., and Liou, J.G. (1987) An experimental investigation of heulandite-laumontite equilibrium at 1000 to 2000 bar Pfluid. Contributions to Mineralogy and Petrology, 97, 43-50. INTRODUCTION 16

Ciesielczuk, J., and Janeczek, J. (2004) Hydrothermal alteration of the Strzelin granite, SW Poland. Neues Jahrbuch für Mineralogie, Abhandlungen, 179, 239- 264. Colella, C. (1996) Ion exchange equilibria in zeolite minerals. Mineralium Deposita, 31, 554-562. Coombs, D.S., Alberti, A., Artioli, A., Armbruster, T., Colella, C., Galli, E., Grice, J.D., Liebau, F., Mandarino, J.A., Minato, H., Nickel, E.H., Passaglia, E., Peacor, D.R., Quartieri, S., Rinaldi, R., Ross, M., Sheppard, R.A., Tillmanns, E., and Vezzalini, G. (1998) Recommended nomenclature for zeolite minerals: report of the subcommittee on zeolites of the international mineralogical association, commission on new minerals and mineral names. Mineralogical Magazine, 62, 533-571. Coombs, D.S., Ellis, A.J., Fyfe, W.S., and Taylor, A.M. (1959) The zeolite facies, with comments on the interpretation of hydrothermal syntheses. Geochimica et Cosmochimica Acta, 17, 53-107. Cronstedt, A.F. (1756) Om en obekant barg art, som kallas Zeolites. Akad. Handl. Stockholm, 17, 120-123. Deer, W.A., Howie, R.A., Wise, W.S., and Zussman, J. (2004) Framework silicates: silica minerals, feldspathoids and the zeolites. 982 p. The Geological Society, London. Dempster, T.J. (1986) Isotope systematics in minerals: biotite rejuventation and exchange during Alpine metamorphism. Earth and Planetary Science Letters, 78, 355-367. Deutsch, A., and Steiger, R. (1985) Hornblende K-Ar ages and the climax of Tertiary metamorphism in the Lepontine Alps (South-Central Switzerland); an old problem reassessed. Earth and Planetary Science Letters, 72, 175-186. Eskola, P. (1939) Die Metamorphen Gesteine. In T. Barth, F.W. Correns, and P. Eskola, Eds. Die Entstehung der Gesteine, p. 263-407. Springer, Berlin. Freiberger, R., Hecht, L., Cuney, M., and Morteani, G. (2001) Secondary Ca-Al silicates in plutonic rocks: implications for their cooling history. Contributions to Mineralogy and Petrology, 141, 415-429. Frey, M., Bucher, K., Frank, E., and Mullis, J. (1980) Alpine metamorphism along the Geotraverse Basel-Chiasso - a review. Eclogae Geologicae Helvetiae, 73, 527- 546. Frey, M., de Capitani, C., and Liou, J.G. (1991) A new petrogenetic grid for low- grade metabasites. Journal of Metamorphic Geology, 9, 497-509. Frey, M., and Mählmann, R.F. (1999) Alpine metamorphism of the Central Alps. Schweizerische Mineralogische und Petrographische Mitteilungen, 79, 135-154. Frisch, W. (1979) Tectonic progradation and plate tectonic evolution of the Alps. Tectonophysics, 60, 121-139. Fujimoto, K., Tanaka, H., Higuchi, T., Tomida, N., Othani, T., and Ito, H. (2001) Alteration and mass transfer inferred from the Hirabayshi GSJ drill penetrating the Nojima Fault, Japan. Island Arc, 10, 401-410. INTRODUCTION 17

Fyfe, W.S., Turner, F.J., and J., V. (1958) Metamorphic reactions and metamorphic facies. 259 p. Gottardi, G. (1989) The genesis of zeolites. European Journal of Mineralogy, 1, 479- 487. Gottardi, G., and Galli, E. (1985) Natural Zeolites. 409 p. Springer, Berlin, Heidelberg, New York, Tokyo. Hauri, F. (2006) Natural zeolites from southern Germany: Applications in concrete. In R.S. Bowmann, and S.E. Delap, Eds. Zeolite'06 - 7th International Conference on the Occurrence, Properties, and Utilization of Natrural Zeolites, p. 130-131, Socorro, New Mexico, USA. Hay, R.L. (1966) Zeolites and zeolitic reactions in sedimentary rocks. Geological Society of America, Special Papers 85. Hay, R.L., and Sheppard, R.A. (1977) Zeolites in open hydrologic systems. In F.A. Mumpton, Ed. Mineralogy and geology of Natural Zeolites, 4, p. 93-102. Mineralogical Society of Amercia, Short Course Notes. Hay, R.L., and Sheppard, R.A. (2001) Occurrences of zeolites in sedimentary rocks. In D.L. Bish, and D.W. Ming, Eds. Natural Zeolites: occurrence, properties, applications, 45, p. 217-234. Mineralogical Society of America, Geochemical Society. Huber, H.M. (1943) Die Kluftminerallagerstätten im südöstlichen Gotthardmassiv. Schweizerische Mineralogische und Petrographische Mitteilungen, 23, 475-537. Iijima, A. (1988) Applications of zeolites to petroleum exploration. In D. Kalló, and H.S. Sherry, Eds. Occurrence, properties and utilization of natural zeolites, p. 29-37. Akadémiai Kiadó, Budapest. Kalló, D. (2001) Applications of natural zeolites in water and wastwater treatment. In D.L. Bish, and D.W. Ming, Eds. Natural Zeolites: occurrence, properties, applications, 45, p. 519-550. Mineralogical Society of America, Geochemical Society. Kenngott, A. (1866) Die Mineralien der Schweiz nach ihren Eigenschaften und Fundorten ausführlich beschrieben. 460 p. Verlag W. Engelmann, Leipzig. Kristmannsdóttir, H., and Tómasson, J. (1978) Zeolite zones in geothermal areas in Iceland. In L.B. Sand, and F.A. Mumpton, Eds. Natural Zeolites: Occurrence, Properties, Use, p. 277-284. Pergamon Press, New York. Labhart, T.P. (1977) Aarmassiv und Gotthardmassiv. 173 p. Borntraeger, Berlin. Langella, A., Cappelletti, P., and de'Gennaro, M. (2001) Zeolites in closed hydrologic systems. In D.L. Bish, and D.W. Ming, Eds. Natural Zeolites: occurrence, properties, applications, 45, p. 235-260. Mineralogical Society of America, Geochemical Society. Liou, J.G. (1971) P-T stabilities of Laumontite, Wairakite, Lawsonite, and related minerals in the system CaAl2Si2O8-SiO2-H2O. Journal of Petrology, 12, 379- 411. Michalski, I., and Soom, M. (1990) The Alpine thermo-tectonic evolution of the Aar and Gotthard massifs, Central Switzerland: Fission track ages on zircon and INTRODUCTION 18

apatite and K-Ar mica ages. Schweizerische Mineralogische und Petrographische Mitteilungen, 70, 373-387. Ming, D.W., and Allen, E.R. (2001) Use of natural zeolites in agronomy, horticulture, and environmental soil remediation. In D.L. Bish, and D.W. Ming, Eds. Natural Zeolites: occurrence, properties, applications, 45, p. 619-654. Mineralogical Society of America, Geochemical Society. Ming, D.W., and Mumpton, F.A. (1989) Zeolites in soils. In J.B. Dixon, and S.B. Weed, Eds. Minerals in Soil environments, 2nd ed., p. 873-911. Soil Science Society of America, Madison, Wisconsin. Mullis, J. (1995) Entstehung alpiner Kluftmineralien. Mitteilungen für Wissenschaft und Technik, Bd. XI, 54-64. Mullis, J. (1996) P-T-t path of quartz formation in extensional veins of the Central Alps. Schweizerische Mineralogische und Petrographische Mitteilungen, 76, 159-164. Mullis, J., Dubessy, J., Poty, B., and O'Neil, J. (1994) Fluid regimes during late stages of a continental collision: physical, chemical and stable isotope measurements of fluid inclusions in fissure quartz from a geotraverse through the Central Alps, Switzerland. Geochimica et Cosmochimica Acta, 58, 2239-2267. Murphy, C.B., Hrycyk, O., and Gleason, W.T. (1978) Natural zeolite: Novel uses and regeneration in wastewater treatment. In L.B. Sand, and F.A. Mumpton, Eds. Natural Zeolites: Occurrence, Properties, Use, p. 471-478. Pergamon Press, New York. Neuhoff, P.S., Fridriksson, T., and Arnórsson, S. (1999) Porosity evolution and mineral paragenesis during low-grade metamorphism of basaltic lavas at Teigarhorn, Eastern Iceland. American Journal of Science, 299, 467-501. Neuhoff, P.S., Fridriksson, T., and Bird, D.K. (2000) Zeolite parageneses in the North Atlantic Igneous Provinces: implications for geotectonics and groundwater quality of basaltic crust. International Geology Review, 42, 15-44. Niggli, E., Koenigsberger, J., and Parker, R.L. (1940) Die Mineralien der Schweizeralpen. 661 p. Wepf & Co, Basel. Noh, J.H., and Boles, J.R. (1993) Origin of zeolite cements in the Miocene sandstones, North Tejon oil fields, California. Journal of Sedimentary Petrology, 63, 248-260. Orlandi, P., and Scortecci, P.B. (1985) Minerals of the Elba pegmatites. Mineralogical Record, 16, 353-364. Parker, R.L. (1922) Über einige schweizerische Zeolithparagenesen. Schweizerische Mineralogische und Petrographische Mitteilungen, 2, 290-298. Pfiffner, O.A. (1986) Evolution of the north Alpine foreland basin in the Central Alps. Special Publication International Association of Sedimentology, 8, 219-228. Rengarten, N.V. (1950) Laumontite and analcime from Lower Jurrasic deposits in the North Caucasus. Dokl. Akda. Nauk SSSR, 70, 485-488. Schaer, J.P., Reimer, G.M., and Wagner, G.A. (1975) Actual and ancient uplift rate in the Gotthard region, Swiss Alps: a comparison between precise levelling and fission-track apatite age. Tectonophysics, 29, 293-300. INTRODUCTION 19

Schmid, S.M., Fügenschuh, B., Kissling, E., and Schuster, R. (2004) Tectonic map and overall architecture of the Alpine orogen. Eclogae Geologicae Helvetiae, 97, 93-117. Sigrist, F. (1947) Beiträge zur Kenntnis der Petrographie und der alpinen Zerrkluftlagerstätten das östlichen Aarmassives. Schweizerische Mineralogische und Petrographische Mitteilungen, 27, 39-182. Spicher, A. (1980) Tektonische Karte der Schweiz 1:500 000, 2nd edition. 71 p. Schweizerische Geologische Kommision, Bern. Stalder, H.A., Wagner, A., Graser, S., and Stuker, P. (1998) Mineralienlexikon der Schweiz. 579 p. Wepf Verlag, Basel. Stober, I., and Bucher, K. (1999) Deep groundwater in the crystalline basement of the Black Forest region. Applied Geochemistry, 14, 237-254. Surdam, R.C. (1977) Zeolites in closed hydrologic systems. In F.A. Mumpton, Ed. Mineralogy and geology of Natural Zeolites, 4, p. 65-92. Mineralogical Society of Amercia, Short Course Notes. Thompson, A.B. (1971) P CO2 in low-grade metamorphism; zeolite, carbonate, clay mineral, prehnite relations in the system CaO-Al2O3-SiO2-CO2-H2O. Contributions to Mineralogy and Petrology, 33, 145-161. Trümpy, R. (1960) Paleotectonic evolution of the Central and Western Alps. Bulletin of the Geological Society of America, 71, 843-908. Trümpy, R. (1980) Geology of Switzerland, a guide book, Part A: an outline of the geology of Switzerland. 104 p. Schweizerische Geologische Kommission, Basel. Tschernev, D.I. (2001) Natural zeolies in solar energy heating, cooling, and energy storage. In D.L. Bish, and D.W. Ming, Eds. Natural Zeolites: occurrence, properties, applications, 45, p. 589-617. Mineralogical Society of America, Geochemical Society. Vincent, M.W., and Ehlig, P.L. (1988) Laumontite mineralization in rocks exposed north of San Andreas Fault at Cajon Pass, southern California. Geophysical Research Letters, 15, 977-980. Walker, G.P.L. (1959) Geology of the Reydarfjördur area, Eastern Iceland. Quarterly Journal of the Geological Society of London, 114, 367-393. Walker, G.P.L. (1960) Zeolite zones and dike distribution in relation to the structure of the basalts of Eastern Iceland. Journal of Geology, 68, 515-528. Walker, G.P.L. (1963) The Breiddalur central volcano, Eastern Iceland. Quarterly Journal of the Geological ociety of London, 119, 29-63. Weisenberger, T., and Selbekk, R.S. (2008) Multi-stage zeolite facies mineralization in the Hvalfjördur area, Iceland. International Journal of Earth Sciences. DOI 10.1007/s00531-007-0296-6 Winter, J.D. (2001). An introduction to igneous and metamorphic petrology, 1, p. 697. Prentice Hall, New Jersey. ZEOLITES IN BASEMENT ROCKS 20

2. ZEOLITES IN FISSURES OF GRANITES AND GNEISSES OF THE CENTRAL ALPS

ZEOLITES IN BASEMENT ROCKS 21

2.1. ABSTRACT

Six different Ca-zeolites occur widespread in various assemblages in late fissures and fractures in granites and gneisses of the Swiss Alps. The zeolites form as a result of water-rock interaction at relatively low temperatures (<250 °C) in the upper continental crust. The low-grade fissure mineral assemblages are the key to the appreciation of water-rock interaction in hydrothermal and geothermal systems located in granites and gneisses of the crystalline basement. The zeolites typically overgrow earlier minerals of the fissure assemblages, but zeolites also occur as single stage fissure deposits in granite and gneiss. They represent the most recent fissure minerals formed during uplift and exhumation of the Alpine orogen. A systematic study of zeolite samples showed that the majority of finds originate from three regions particularity rich in zeolite-bearing fissures: (1) in the central and eastern part of the Aar- and Gotthard Massif, including the Gotthard road tunnel and the Gotthard- NEAT tunnel, (2) Gibelsbach/Fiesch, in a fissure breccia between Aar Massif and Permian sediments, and (3) in Penninic gneisses of the Simano nappe at Arvigo (Val Calanca).

The excavation of tunnels in the Aar- and Gotthard massif give an excellent overview of zeolite frequency in Alpine fissures, whereas 32 % (Gotthard NEAT tunnel, 12000-18555) and 18 % (Gotthard road tunnel) of all fissures are filled with zeolites. The number of different zeolites is limited to 6 species: laumontite, stilbite and scolecite are abundant and common, whereas heulandite, chabazite and epistilbite occur occasionally. Ca is the dominant extra-framework cations, with minor K and Na. Heulandite and chabazite additionally contain Sr up to 29 and 10 mole%, respectively. Na and K content of zeolites tends to increase during growth as a result of systematic changes in fluid composition and/or temperature. The K enrichment of stilbite found in surface outcrops compare to stilbite in the subsurface may indicate late cation exchange during interaction with surface water. Texture data, relative age sequences derived from fissure assemblages and equilibrium calculations shows that the Ca-dominated zeolites precipitated from fluid with decreasing temperature in the order (old to young = hot to cold): scolecite, laumontite, heulandite, chabazite and stilbite. ZEOLITES IN BASEMENT ROCKS 22

The components necessary for zeolite formation are derived from dissolving primary granite and gneiss minerals. The nature of these minerals depends on the metamorphic history of the host rock. Zeolites in the Aar Massif derived from the dissolution of epidote or calcite and albite that were originally formed during Alpine greenschist metamorphism. Whereas albitization of plagioclase in higher grade rocks releases the necessary components for zeolite formation, a process that is accompanied by a distinct porosity increase. Zeolite fissures occur in the zone where fluid inclusions in earlier formed quartz contain H2O dominated fluids. This is consistent with equilibrium calculations that predict a low CO2 tolerance of zeolite assemblages particularly at low temperature. Pressure decrease along the uplift and exhumation can increase zeolite stability. The major zeolite forming reaction consumes calcite and albite; it increases pH and the total of dissolved solids. The produced Na2CO3 waters are in accord with reported deep groundwater (thermal water) in the continental crust, which are typically oversaturated with respect to Ca- zeolites.

Keywords: zeolite, granite, water-rock interaction, laumontite, Swiss Alps

2.2. INTRODUCTION

The interaction of rocks with hot water circulating on the fractures of the continental crust produces fissure minerals in the open porosity of the fissures and leaches the original rock matrix. The chemical composition of the fluid thereby monitors the water-rock interaction process that controls the dissolution of primary minerals, as well as the precipitation of secondary minerals in the open spaces (e.g. Nordstrom et al., 1989; Stober & Bucher, 1999; Bucher et al., 2009). Veins and mineralized fractures are ubiquitous in regional metamorphic terrains. They bear considerable information on fluid movement, fluid-rock interaction and fluid sources (e.g. McCaig et al., 1990). Detailed mineralogical and petrological study of the low-grade fissure mineral assemblage provides quantitative access to fluid-rock interaction. From such data the evolution of porosity and permeability of the total system and the leached rock matrix can be deduced. These flow properties of fractured rocks are required for the understanding of geothermal systems and fluid migration in the upper continental ZEOLITES IN BASEMENT ROCKS 23 crust (e.g. Gianelli et al., 1998; Neuhoff et al., 1999; Weisenberger & Selbekk, 2008). The crystalline basement of the upper continent crust consists predominantly of granites and gneisses. A major mineral of both rock types is plagioclase. Hydrothermal alteration of basement rocks along fractures attacks predominantly plagioclase and replaces the mineral with secondary Ca-Al silicates such as epidote, prehnite and zeolites. Zeolites are the predominant secondary Ca-Al mineral at low temperature (< 250 °C) (e.g. Gottardi, 1989; Bish & Ming, 2001; Fig. 2.1).

Although zeolites are known from fissures and gashes of the crystalline basement from reports by mineral collectors for more than 150 years (e.g. Kenngott, 1866; Parker, 1922; Niggli et al., 1940; Huber, 1943; Sigrist, 1947; Stalder et al., 1998) (Table 2.1), they did not create much interest in the scientific research community. Previous publications on zeolite occurrences and their origin in the Central Swiss Alps and other areas of crystalline basement rocks are limited (Armbruster et al., 1996; Freiberger et al., 2001; Fujimoto et al., 2001; Ciesielezuk & Janeczek, 2004).

Fig. 2.1: Pressure-temperature ranges of environments of zeolite formation. (adapted from Deer et al., 2004). Solid curves are experimentally determined stability limits of selected zeolites: (1) epidote + quartz + H2O = laumontite + prehnite, at low-pressure end and epidote + chlorite + quartz + H2O = laumontite + pumpellyite at high-pressure end (2) laumontite + quartz + H2O = heulandite (Cho et al., 1987). Dashed line represents a retrograde PT-path in Alpine fissures, determined by fluid inclusions in fissure quartz (Mullis et al., 1994).

ZEOLITES IN BASEMENT ROCKS 24

In general, zeolites are among the most common products of chemical interaction between fluids and the crustal rocks during diagenesis and low-grade metamorphism (e.g. Bish & Ming, 2001). Zeolite minerals occur in low temperature (<250 °C), low pressure (<200 MPa), water saturated environments. The required constituents for the formation of zeolites are commonly derived from dissolution of volcanic glass (e.g. Sheppard & Hay, 2001) and zeolites are common in altered volcanic rocks. In granites and gneisses, the necessary components can be derived from the alteration of feldspars, particularly plagioclase, and other aluminous silicates (Engvik et al., 2008). Temperature and pressure largely control the kind of zeolite that will be formed. P and T is usually a function of burial depth or temperature changes during hydrothermal overprint. The composition of the altered material and the composition of the hydrothermal fluid are further controls on the product zeolite mineralogy (e.g. Bucher & Stober, 2001; Deer et al., 2004).

Zeolites as products of hydrothermal crystallization are generally know from active geothermal systems associated with volcanic rocks. Very little work has been published on zeolite occurrences related to late stage crystallization of pegmatitic bodies (e.g. Orlandit & Scortecci, 1985), in hydrothermal ore veins (Deer et al., 2004), as alteration along fault plains (e.g. Vincent & Ehlig, 1988), and in hydrothermal fractures and veins in granites and gneisses (e.g. Borchardt et al., 1990; Borchardt & Emmermann, 1993; Armbruster et al., 1996; Bish & Ming, 2001; Freiberger et al., 2001; Fujimoto et al., 2001).

The scarcity of reports on zeolite in fissures of granites and gneisses is astonishing because of the obvious very widespread occurrence of zeolites in granites and gneisses. For example in the latest special edition on natural zeolites (Bish & Ming, 2001), only two short notes were made on zeolites hosted in granites and gneisses. Nevertheless zeolite formation in fractures and cavities in granitic gneisses is a frequent feature in the continental crust.

In this paper, we present the “uncommon” zeolite occurrences in fractures and cavities in granites and granitic gneisses in the Central Swiss Alps. The presented data allow for a consistent model of zeolite formation in basement rocks (Table 2.2; Fig. 2.2).

ZEOLITES IN BASEMENT ROCKS 25

2.3. GEOLOGICAL SETTING

The Alps forms a part of a Tertiary orogenic belt that stretches from southern Europe to Asia. The Alps formed as a result of the closure of Jurassic to Cretaceous Tethys ocean basins during convergence of the Apulian and European plates (e.g. Trümpy, 1960; Frisch, 1979; Schmid et al., 2004). An orogenic belt characterized by stacked nappes formed in the Tertiary when the Apulian and European plates collided. The collision caused a complicated tectonic structure and a regional metamorphic overprint. Deeply buried parts of the orogen were later exhumed and uplifted in the late Tertiary (Trümpy, 1980) and finally reached the erosion surface.

Fig. 2.2: Map of Switzerland and the Central Swiss Alps. (a) Outline of Switzerland and the position of the external massifs (modified after Labhart, 1977). Numbered points marks zeolites localities and the numbers corresponds to Table 2.1. The positions of the Gotthard road tunnel and the Gotthard NEAT tunnel are marked central external massifs in gray. Black star marks the position of the zeolite locality Arvigo, Val Calanca/GR. (b) Simplified geological map with dashed line. (c) Spatial distribution of each zeolite, showing no preferred distribution.

Zeolites in fissures occur predominantly in rocks that belong to large basement windows exposed in the northern part of the Alps. These so-called “external massifs” of the Alps thus belong to the European plate (e.g. Trümpy, 1980). The massifs represent parautochthonous units (Pfiffner, 1986). Two major basement units are distinguished in the Central Swiss Alps: the Aar Massif and the Gotthard Massif. They constitute to the pre-Variscan basement, which is partly reworked by the ZEOLITES IN BASEMENT ROCKS 26

Variscan and Alpine orogenesis. The massifs form a 115 km long and 23 to 40 km wide SW-NE trending outcrop. The large Aar Massif consists of pre-Variscan gneisses, pre-Variscan granitoids, migmatitic granites and gneisses, lower and upper Carboniferous intrusives and Carboniferous volcanics (Abrecht, 1994). Many of the prominent high Alpine peaks and the largest glaciers of the Alps are located in the Aar massif.

The Gotthard Massif is located to the south of the Aar Massif and is followed further south by the north Penninic continental nappe stack. It consists of a poly- metamorphic continental basement with Variscan granites. It is separated from the Aar Massif in the north by the narrow Tavetsch Massif and the Mesozoic metasediments of the Urseren zone (Fig. 2.2).

2.3.1. Metamorphic conditions during Alpine orogenesis

All rocks have been overprinted by the Tertiary Alpine metamorphism. The metamorphic grade and Alpine peak metamorphism increases from nearly non- metamorphosed rock units in the north, over greenschist facies rocks in the Aar- and Gotthard Massif region up to amphibolite facies conditions in the Penninic nappes to the south (Labhart, 1977; Frey et al., 1980; Frey & Mählmann, 1999). In general, metamorphic grade increases within a specific tectonic unit from north (external position) to south (internal position), as well as from the top to the bottom (Frey & Mählmann, 1999).

Metamorphic conditions of the Central Alps exceed the zeolite facies with the exception of the parautochthonous units in the north. A continuous zone of very low- grade metamorphism (= anchizone) of up to 15 km width can be delineated along the southern Helvetic nappes and in the parautochthonous cover of the Aar Massif basement (Frey & Mählmann, 1999). Peak metamorphism determined from the Taveyanne greywacke (Glarus Alps) in parautochthonous units north of the Aar Massif range from zeolite (Lmt + Prh + Pmp + corrensite), to prehnite-pumpellyite (Prh + Pmp + Ep), pumpellyite-actinolite (Pmp + Act + Ep) and lower greenschist facies (Act + Ep) (Rahn et al., 1994). Metamorphic conditions in parautochthonous units north of the Aar Massif corresponds to a temperature range of 240-300 °C and 200-300 MPa (Frey & Mählmann, 1999; Rahn et al., 1994). ZEOLITES IN BASEMENT ROCKS 27

Fig. 2.3: Detailed geological map of the eastern Aar Massif (modified after Labhart, 1977). Numbered points marks zeolites localities and the numbers corresponds to Table 2.1. G = glacier

Tertiary greenschist facies metamorphism (= epizone) overprinted the old basement rocks of the Aar- and Gotthard Massif. The following isograds have been located from north to south: (1) first appearance of green biotite (Steck & Burri, 1971), (2) disappearance of stilpnomelane (Jäger et al., 1967), (3) the transformation isograd of microcline/sanidine (Bambauer & Bernotat, 1982; Bernotat & Bambauer, 1982; Frey & Mählmann, 1999). This corresponds to maximum temperatures in the Northern Aar Massif of about 270 °C. This temperature is above the zeolite window defined by data on illite crystallinity, vitrinite reflection and fluid inclusion measurements by Breitschmid (1982). The typical Alpine mineral assemblages of the Central Aar granite (Fig. 2) is Qtz + Ab + Kfs + Chl + Ms + Cal. Alpine green biotite occurs along a mappable isograd (Steck & Burri, 1971) in the Aar granite suggesting a minimum temperature at about 420 °C. The microcline/sanidine transformation isograd is located further south in the Aar granite, suggesting a further increase of metamorphic grade with minimum temperatures of 450 °C (Bambauer & Bernotat, 1982). The first appearance of oligoclase in granitic gneisses in the Gotthard Massif (Steck, 1976) still further south marks the beginning of amphibolite facies conditions (about 500 ˚C).

Metamorphic peak conditions regularly increase southward in the Lepontine Alps (the Penninic nappe stack). The stacked nappes involve staurolite-bearing micaschists, tremolite marbles and amphibolites all characteristic of amphibolite facies conditions ZEOLITES IN BASEMENT ROCKS 28

(Frey & Mählmann, 1999). Peak conditions during Eocene to Miocene metamorphism at the Arvigo zeolite locality (Fig. 2.2) in the Simano nappe (a member of the Penninic nappe stack range from 600-680 °C and 550 to 600 MPa (Engi et al.; 1995; Todd & Engi, 1997; Nagel et al., 2002). Deutsch & Steiger (1985) determined an age of 37 Ma for peak metamorphism in the Lepontine Alps, whereas peak metamorphism in the Aar Massif is dated around 25 Ma (Grimsel area, Dempster, 1986).

During uplift and erosion brittle deformation structures became dominant once the rocks crossed the ductile-brittle transition zone. Semi-brittle shear zones related to backthrusting, normal faults and the opening of fissures and gashes are typical structures related to the late orogenic deformation phases including extension structures related to uplift. The fissures can be divided into two characteristic groups based on geometry and morphology (Mullis et al., 1994; Mullis, 1995): (1) tension gashes and (2) interboudin gaps. Tension gashes generally develops parallel to the maximum stress (σ1) and perpendicular to the maximum elongation, at an angle of around 45° to the shear plane (Huber, 1948; Mullis et al., 1994). They are mostly arranged in en echelon fashion (Ramsay, 1967). The variation in length reach from <10 cm up to more than 10 meters. Interboudin gaps usually develop parallel to the direction of maximum extension in rocks of different viscosity. In contrast to tension gashes the interboudin gaps rarely exceed the size of 1 meter. The orientation of both types of fissures is normal to foliation or schistosity (Mullis et al., 1994).

Both structures are conductive to hot aqueous fluids. The percolating hot fluids could react with the surrounding rocks along the fissure walls to form secondary fissure minerals.

Metamorphic conditions derived from fluid inclusion studies on Alpine fissure material represent conditions during various stages of exhumation, decompression and cooling. Mullis et al. (1994) determined minimum conditions for fissure mineral formation ranging from 400-430 °C in temperature and fluid pressures from 240 to 380 MPa for the southern Aar Massif (Zinggenstock), for the Gotthard Massif and the northern Lepontine Alps.

An exhumations rate of 0.5 mm a-1 for the Reuss valley (northern Aar Massif) has been proposed by Michalski & Soom (1990) for the past 27 Ma from apatite and zircon fission track data. The corresponding cooling rate of 13 °C Ma-1 agrees well ZEOLITES IN BASEMENT ROCKS 29 with other apatite fission track data that suggest uplift rates of 0.3-0.6 mm a-1 during the last 6-10 Ma (Schaer et al., 1975). Using exhumation rates and trapping temperatures of early fluid inclusions (Mullis, 1996) determined the time of the first opening of fissures and precipitation of fissure minerals in the Aar Massif (Zinggenstock) and Gotthard Massif (La Fibbia) to 20 to 15 Ma b.p..

The deposition of fissure minerals, hence also the zeolite minerals described in this paper, occured along the cooling and decompression P-T-path after peak metamorphism.

2.4. SAMPLING AND ANALYTIC METHODS

Samples of fissure minerals from surface outcrops and from road and rail tunnels in the Central Swiss Alps provided assemblage data, relative age relationships and chemical composition data from zeolites and associated minerals. Samples were supplied from four sources (Table 2.2): (1) The Swiss Natural History Museums Bern (SNHMB) made the vast collection of Alpine minerals available for our systematic study of the spatial distribution of zeolites, the museum keeps samples of important zeolite localities, including the Gotthard road tunnel. (2) Excellent mineral samples from the Amsteg-Sedrun section of the new Gotthard rail base tunnel currently under construction were made available by the mineral representative Peter Amacher. He saved the specimens during excavation work during the last few years. (3) The Mineralogical Museum University Freiburg provided samples from classic localities. (4) During several field trips in the summer seasons 2006 and 2007 mineral samples were collected from the Aar- and Gotthard Massif, Gibelsbach and Arvigo/Val Calanca.

Quantitative zeolite analysis were performed at the Institute of Mineralogy and Geochemistry, University of Freiburg, using a CAMECA SX 100 electron microprobe equipped with five WD spectrometers and one ED detector with an internal PAP-correction program (Pouchou & Pichior, 1991). Major and minor elements for zeolites were determined at 15 kV accelerating voltage and 8 nA beam current with a defocused electron beam of 20 µm in diameter with counting time up to ZEOLITES IN BASEMENT ROCKS 30

20 s. Na and K were counted first to minimize the Na and K loosed during determination. Since zeolites lose water when heated, the crystals were mounted in epoxy resin to minimize loss of water due to the electron bombardment. Natural and synthetic standards were used for calibration. The standards employed were: albite (Na), periclase (Mg), wollastonite (Si), barite (Ba), hematite (Fe), celestine (Sr), orthoclase (K), anorthite (Ca), rhodonite (Mn) and rutile (Ti). Identification of various minerals was obtained by a BRUKER AXS D8 Advance X-ray powder diffractometer (XRD) and the DIFFRACplus v5.0 software for evaluation.

The content of zeolite water was determined by heating the samples that had been equilibrated with air of 50 % relative humidity to 873 K for 24 h and measuring the weight loss. For some samples zeolite water was determined by measuring mass loss between 298 and 1273 K by scanning-heating TGA (thermogravimetry analysis) at a heating rate of 10 K min-1 on a Netzsch STA 449C Jupiter simultaneous DSC-TGA (differential scanning calorimetry - thermogravimetry analysis) apparatus. The charge balance of zeolites formulas is a reliable measure for the quality of the analysis and which correlates with the difficulties related to the thermal instability of zeolites in microprobe analysis. A usefull error test investigates the charge balance between the non-framework cations and the amount of tetrahedral Al (Passaglia, 1970). Analyses are considered acceptable if the sum of the charge of the extra-framework cations (Ca2+, Sr2+, Na+, and K+) is within 10% of the framework charge (Al3+).

2.5. ZEOLITES IN THE CENTRAL ALPS

In contrast to other zeolite environments (e.g. basalts, alkaline lakes) only few different zeolite species occur in granites and gneisses. Laumontite, scolecite, heulandite, stilbite and chabazite mark the dominant species whereas epistilbite occurs only at three sites (Table 2.1; no 4, 25 & 57). Additionally thomsonite, phillipsite, natrolite, mesolite and analcime were found in the Central Alps associated with rocks of basic to ultramafic composition (e.g. amphibolites, metabasalts, serpentinites), e.g. within the Zermatt-Saas ophiolite (Table 2.1; no 67) or the Geisspfad (Binntal) serpentinite (Table 2.1; no 54) and the zeolite-bearing Taveyanne greywacke (Glarus Alps, Rahn et al., 1994). These zeolite occurrences will not be ZEOLITES IN BASEMENT ROCKS 31 discussed here. Faujasite was mentioned by Parker (1922) in the eastern part of the Central Aar granite body (Fig. 2.2) but the mineral has never been confirmed.

2.5.1. Spatial distribution

The spatial distribution of zeolites in Alpine fissures in crystalline basement rocks of the Central Alps is summarized in Figs 2.2 & 2.3 and Table 2.1. It follows from the data compilation that zeolites are not evenly distributed but rather have preferentially been reported from a relatively small number of localities. In a broad zone surrounding these localities zeolites are very common and widespread. These focus areas typically cover some km2. It is important to note that the patchy zeolite distribution does not result from sampling bias. Sampling bias can be excluded because the main target of the mineral hunters is rock crystal and smoky quartz, which is found in fissures over the entire area. However, quartz-bearing fissures are commonly devoid of zeolite minerals.

Fig. 2.4: Photograph and schematic illustration showing typical vein characteristics of Alpine fissures. (a) Photograph of a vein hosted in a biotite-rich gneiss (Arvigo/Val Calanca); hammer for scale. The vein is characterized by a 1 cm leaching zone trending in vertical direction, which appears to be lighter colored, due to the removal of mafic minerals. The open space of the fissure is filled with secondary minerals, mainly chlorite. (b) Schematic sketch of an Alpine fissure. The open fissures provide pathways for hot fluids. Leaching caused by fluid-rock interaction change primary mineralogy and composition of the host rock, visible as alteration zone along to the fissure wall. Secondary minerals precipitate in the open space. (c) Schematic sketch of a zeolite bearing Alpine fissure, exhibit euhedral mineral assemblages. Zeolites overgrow earlier formed minerals in the following order, as it observed in nature: Qtz → Ep → Prh → Sco. The alterations zones seem to be proportional to the aperture of fissures.

One of the prime sources of Alpine fissure quartzes is the Central Aar granite (Figs 2.2 & 2.3), which forms an over 100 km long and 8 to 10 km wide intrusive body ZEOLITES IN BASEMENT ROCKS 32 striking in SW-NE direction. It contains abundant fissure zeolites only in an eastern area centering around Piz Giuv (Figs 2.2 & 2.3). In the Giuv area zeolites are abundant in surface outcrops but also in fissures opened during major tunnel construction (e.g. Gotthard road tunnel, Stalder et al., 1980; new base rail tunnel Gotthard NEAT). In the tunnels zeolite fissures occur up to 1500 to 2000 m below the surface. Fissures in other parts of the Central Aar granite do not contain zeolites (Fig. 2.2). The Aar granite is rather uniform and homogeneous in mineralogical and chemical composition (Labhart, 1977).

Fissure zeolites are not restricted to a specific lithological unit (Figs 2.2 & 2.3). In the Giuv area, for example, zeolite-bearing fissures occur in most lithologies of the crystalline basement (Fig. 2.3). Zeolite occurrence is not controlled by changes in bulk composition of basement rocks. However, zeolites disappear south of the Aar Massif basement abruptly and are not present in the meta-sediments of the Tavetsch Massif (Fig. 2.3). In the Gotthard massif zeolites also occur in basement granites and gneisses but are absent in fissures in the meta-sediments of the cover units (e.g. Stalder et al., 1980).

The distribution of zeolites in vertical direction is accessible thanks to large numbers of zeolite samples recovered during tunnel construction operations (e.g. Gotthard road and rail tunnel, NEAT Gotthard base railway tunnel, NEAT Lötschberg base rail tunnel). These outstanding data show that laumontite is the dominant zeolite mineral in Alpine fissures. Scolecite, heulandite, chabazite and stilbite are present in distinctly smaller numbers of fissure assemblages (Stalder et al., 1980; Amacher, pers. com.). Laumontite is also abundant at the Arvigo locality (Fig. 2.2), where various zeolites are found in an active quarry with large blocks of basement gneisses. However, laumontite has been reported only sporadically, although from many localities, from surface outcrops where the fissures have been exposed to weathering conditions (Stalder et al., 1998) (Fig. 2.2). In contrast, scolecite, heulandite, chabazite and stilbite are the dominant zeolites species in fissures of surface outcrops (e.g. Niggli et al., 1940; Huber, 1943; Sigrist, 1947; Huber, 1948; Stalder et al., 1998).

The data show that no zonal distribution pattern can be recognized for any of the zeolite minerals. Both, a zonal regional distribution as well as a vertical zonal distribution is absent. This lack of mineral zone patterns in the distribution of Alpine fissure zeolites is in contrast to other environments (Langella et al., 2001; Sheppard & ZEOLITES IN BASEMENT ROCKS 33

Hay, 2001; Utada, 2001a, b) where zeolites are formed (e.g. deep marine sediments, hydrothermal alteration and burial metamorphism in volcanic rocks, saline high alkaline lakes).

Characteristic of Alpine zeolite fissures is that in many of the fissures a succession of zeolites minerals can be found as a paragenesis in a single fissure.

Table 2.1. Zeolite localities in the Central Swiss Alps (numbers corresponds to numbers shown in Figs 2.2 & 2.3). No Locality Ca Sco Lmt Stb Heu Cha AFMb Fc Referencesd 1 Bäregg, Oberaar, BE x x x Qtz, Kfs, x Stalder, 1964; Stalder et al., Grimsel Hem, Cal, 1998; SNHMB; ° Chl, Py 2 Lötschberg tunnel BE x Qtz, Cal x SNHMB 3 Bächistock GL x x x Niggli et al., 1940; SNHMB 4 Arvigo/Val GR x x x x Qtz, Kfs, xxx Stalder et al., 1998; Wagner, CalancaEpi Ttn, Act, 2000a, b; SNHMB; ° Ep, Prh, Chl, Ap 5 BergellPhil * GR x x x x x Qtz, Kfs, x Hirschi, 1925; Stalder et al., Grs, Prh, 1998; SNHMB Ttn, Cal 6 Cuolm da Vi GR x x Kenngott, 1866; Huber, 1948; SNHMB 7 Davos GR x x x Stalder et al., 1998; SNHMB 8 Drumtobel/Sedrun GR x x x x x Qtz, Kfs, xxx Stalder et al., 1998; SNHMB; Cal, Act, ° Ttn 9 Misox GR x Chl x SNHMB 10 OberalpstockTho GR Qtz x Niggli et al., 1940; Stalder et al., 1998 11 Sedrun GR x x x Kenngott, 1866; SNHMB 12 Sella/Gotthard GR x Qtz, Hem, x Kenngott, 1866; SNHMB; ° Chl 13 Stgegia, Medel GR x Qtz x SNHMB 14 Val Casatscha GR x x Cal x Stalder et al., 1998; Huber, 1943; ° 15 Val Cristallina GR x x Cal x SNHMB 16 Val Giuv GR x x x x Qtz, Kfs, xxx Kenngott, 1866; Sigrist, 1947; Ap Stalder et al., 1998; SNHMB; ° 17 Val Maighels* GR x Chl x Huber, 1943; Stalder et al., 1998; SNHMB 18 Val Medel GR x x Huber, 1943; SNHMB 19 Val Mila GR x Qtz x Niggli et al., 1940; Sigrist, 1947; SNHMB 20 Val Muretto/Bergell GR x x Stalder et al., 1998 21 Val Nalps GR x x Prh, Tur x SNHMB 22 Val Punteglias GR x Qtz, Chl x Huber, 1943; Sigrist, 1947; Stalder et al., 1998; SNHMB 23 Val Russein GR x x Qtz, Ep, Chl xx Huber, 1943; Stalder et al., 1998; SNHMB 24 Val Strem GR x x x x Qtz, Chl xxx Kenngott, 1866; Niggli et al., 1940; Huber, 1948; Stalder et al., 1998; SNHMB; ° ZEOLITES IN BASEMENT ROCKS 34 continue Table 2.1 25 BiascaEpi TI x Kfs, Ttn, x Stalder et al., 1998; SNHMB Ep. Chl 26 Camperio/Passo del TI x Qtz, Hem, x Wagner et al., 1972; SNHMB Lucomango Cal, Chl, Prh 27 Domodossola TI x Chl x SNHMB 28 La Fibbia, Gotthard TI x Qtz Chl, xx Kenngott, 1866; SNHMB Hem 29 Lodrino TI x Qtz, Chl x Stalder et al., 1998; SNHMB 30 Mt Ceneri TI x x Py x Toroni, 1984; Stalder et al., 1998 31 Pizzo Lucendro TI x x x Qtz Kfs, x Stalder et al., 1998; SNHMB Hem, Ms 32 Val Baveno TI x x x Stalder et al., 1998 33 Val Canaria* TI x x SNHMB 34 Val Maggia* TI x x x Kenngott, 1866; Stalder et al., 1998; SNHMB 35 Val Vergeletto, road TI x x x Simonetti, 1971; Stalder et al., tunnel 1998; SNHMB 36 Andermatt UR x x SNHMB 37 Brunnital UR x Cal x SNHMB 38 Chrützlistock UR x x Qtz xxx Kenngott, 1866; Niggli et al., 1940; Sigrist, 1947; Weibel, 1963; Stalder et al., 1998; SNHMB; ° 39 Etzlital UR x x x x xxx Kenngott, 1866; Niggli et al., 1940; Sigrist, 1947; Stalder et al., 1998; SNHMB; ° 40 Fedenstock UR x Qtz, Chl x SNHMB 41 Fellilücke UR x x Qtz, Flt xxx Niggli et al., 1940; Sigrist, 1947; SNHMB; ° 42 Fellital UR x x x Qtz, Chl xxx Niggli et al., 1940; Sigrist, 1947; SNHMB; ° 43 Göscheneralp UR x x x Qtz, Chl x Kenngott, 1866; Stalder et al., 1998; SNHMB 44 Gotthard road UR x x x x Qtz, Kfs, xxx Stalder et al., 1980, 1998; tunnel Cal, Ep, SNHMB; ° Prh, Chl, Py 45 Griesserental UR x Kfs x SNHMB 46 Maderanertal UR x Qtz, Chl, Py x SNHMB 47 NEAT, Amsteg - UR x x x x Qtz, Kfs, xxx ° Sedrun Cal, Hem, Act, Ep, Chl, Apo, Py, Anh 48 Piz Giuv UR x xxx Kenngott, 1866; Niggli et al., 1940; Sigrist, 1947; Huber, 1948; SNHMB; ° 49 Riental UR x x Qtz xxx Kenngott, 1866; Niggli et al., 1940; SNHMB; ° 50 Schattig Wichel & UR x x x x x Qtz, Kfs, xxx Kenngott, 1866; Niggli et al., Piz Giuv Ap, Ttn, 1940; Sigrist, 1947; Huber, Act, Chl, 1948; Stalder et al., 1998; Ep, Prh SNHMB; ° 51 Schijenstock UR x x Qtz, Flt, xx Niggli et al., 1940; Stalder et Hem al., 1998; ° 52 Tiefengletscher UR x Qtz, Gn x SNHMB ZEOLITES IN BASEMENT ROCKS 35 continue Table 2.1 53 Arolla VS x x Qtz x Stalder et al., 1998; SNHMB 54 Binntal Meso, Nat, Phil * VS x x x x x Qtz, Kfs, x Kenngott, 1866; Keusen & Hem, Chl, Bürki, 1969; Stalder et al., Ttn, 1998; SNHMB 55 Fieschergletscher VS x x Qtz, Kfs, x Kenngott, 1866; Niggli et al., Ep, Ap 1940; Stalder et al., 1998; SNHMB 56 Furka tunnel VS x Qtz, Kfs, x Wälti, 1984; SNHMB Chl, Ttn 57 Gibelsbach/FieschEpi VS x x x x Qtz, Kfs, Flt xxx Kenngott, 1866; Koenigsberger, 1917; Armbruster et al., 1996; Stalder et al., 1998; SNHMB 58 Gornergletscher VS x x Stalder et al., 1998 59 Gredetschtal/Brig VS x Qtz, Kfs, xx Niggli et al., 1940; Stalder et Chl, Ttn al., 1998; SNHMB 60 Grosses Sidelhorn VS x Qtz, Chl x SNHMB 61 Lax VS x x x Qtz, Chl x Kenngott, 1866; Stalder et al., 1998; SNHMB 62 Lötschental* VS x x x x Qtz, Kfs, xx Fellenberg von, 1893; Stalder Cal, Act et al., 1998; SNHMB 63 Martiny VS x Qtz x SNHMB 64 Massaschlucht VS x Kfs, Cal x SNHMB 65 Mättital VS x x Qtz, Cal x SNHMB 66 Nuffenenpass VS x Qtz x SNHMB 67 Pollux/ZermattNat VS xx Stalder et al., 1998 68 Simplontunnel VS x Chl x Stalder et al., 1998; SNHMB Mineral abbreviations used after Bucher & Frey (2002). Following abbreviations are used for zeolites: Sco = scolecite, Lmt = laumontite, Stb = stilbite, Heu = heulandite, Cha = chabazite. a Canton (Swiss districts, BE = Bern, GL = Glarus, GR = Grisons TI = Ticino, UR = Uri, VS = Valais). b major associated fissure minerals. c frequency of zeolites (x = sporadic zeolite occurrence, xx = cumulative zeolites occurrence, xxx = zeolites occur frequently). d SNHMB = collection of the Swiss Natural History Museum Bern.. Epi epistilbite. Meso mesolite. Nat natrolite. Phil phillipsite. Tho thomsonite. * = zeolites hosted in basic rocks (e.g. amphibolite). ° = localities from which samples were analyzed in this study

2.5.2. Field occurrences

2.5.2.1. General features

The zeolites reported in this paper occur exclusively in fissures, veins and gashes. Rock forming zeolite minerals such as in the Taveyanne greywacke (Rahn et al., 1994) are not considered here. The fissure zeolites cover and coat the walls of fractures in granites and gneisses but also occur as pore and cavity filling in leached host rocks. Zeolites typically overgrow earlier formed minerals of the fissure assemblage, but they also occur as single stage fissure deposits in granites and gneisses (e.g. Kenngott, 1866; Parker, 1922; Stalder et al., 1980; Armbruster et al., 1996). ZEOLITES IN BASEMENT ROCKS 36

Assemblage and distribution data of zeolite minerals are compiled in Table 2.1. Zeolites are present as transparent, white or brownish crystals. The crystal size is relatively small and usually does not exceed 1 cm. This is in contrast to other fissure minerals, which commonly reach crystal sizes of several cm to dm. Typically zeolites occur as very small (< 1 mm) inconspicuous whitish coatings on earlier formed minerals or direct on the fracture wall, which makes them difficult to recognize. Some zeolite crystals have a green color (Table 2.2; A8332), due to small inclusions of chlorite.

Fig. 2.5: Fissures in the Gotthard NEAT tunnel. (a) Tunnel head wall showing fissure, that strikes in S- W direction. (b) Sigmoidal fissure. (c) Photomicrograph of laumontite needles. (d) Laumontite covering earlier formed quartz as dense mats.

The fissures tend to be lens-shaped with large long- to short-axis ratios. Open fissure cavities range from cm to several tens of meters in length. The aperture of the fractures varies from mm to tens of cm (some open crystal caves of > 1m have been found). The fractured host rock is commonly chemically leached on both sides of the fissure (Fig. 2.4). The leached zone is usually lighter colored than the unaltered host rock, due to the lack of primary dark minerals such as biotite in the former (Fig. 2.4). The leached zone shows a higher porosity compared with the host rock (e.g. Parker, 1922; Sigrist, 1947; Mercolli et al., 1984; Ciesielezuk & Janeczek, 2004). The secondary porosity of the leached rock is locally filled with secondary minerals, forming a zone of impregnation (e.g. Huber, 1943; Mercolli et al., 1984). Leaching zones are not always present in the fissures. Leaching zones range from mm to m scale, but it seems that the width of leaching zone is related to the aperture of the fracture. Leaching always increases towards the central open space. Frequently much of the primary rock material has been removed and the remaining alteration products form a very porous crumbly disjointed mass (Table 2.2; DT, TW34). ZEOLITES IN BASEMENT ROCKS 37

The dominant minerals in Alpine fissures are quartz, adularia and chlorite. There is a remarkable absence of clay minerals (except chlorite) in all Alpine fissures. In the most common multi-mineral veins and fissures the volume fraction of zeolites is very low (Huber, 1943; Sigrist, 1947; Stalder et al., 1980). However, single-mineral zeolite fissures are common and widespread. In tunnel fissures where laumontite is the dominant zeolite covering fissure walls as dense mats (Figs 2.5 & 2.6), the zeolite may be modally abundant also in multi-mineral veins (Stalder et al., 1980). Because zeolites tend to be well preserved in fissures opened in the progress of tunnel constructions, tunnel data give a clue on the frequency and the abundance of zeolites in fractured granite and gneiss. During construction of the 16 km long Gotthard road tunnel 225 fissures were recorded during a systematic evaluation of fissures in the main- and security tunnel in the northern section (from north portal to 7 km to the south). 41 of the fissures contained zeolites (18 % of all fissures; Stalder et al., 1980). In some lithologies, for instance in the Southern gneisses of the Aar Massif, zeolites were found in 50 % of the fissures (Stalder et al., 1980). Similar data are available from the new Gotthard NEAT tunnel. In the section 12000 to 18555 meters (Fig. 2.7), 26 of 83 fissures (32 %) yield zeolite species (P. Amacher pers. com.).

Zeolites normally occur together with other mineral species in a vein. A general chronology of zeolite-bearing Alpine fissures is compiled in Table 2.3. The schematic sequence of successive minerals and mineral assemblages in Alpine fissures is based on observed textural (overgrowth) relationships providing relative age and sequence from a large number of fissures (Table 2.2).

Zeolites generally overgrow all earlier formed minerals and usually represent the latest mineral formed in Alpine fissure. In some veins euhedral crystals of apophyllite overgrow locally laumontite as (e.g. Gotthard NEAT, Table 2.2; KB868; Arvigo, Wagner et al., 2000a, b; Gotthard road tunnel, Stalder et al., 1980). All six Ca-zeolites (Table 2.3) never occur together, a maximum of three different successive zeolites may be found in a single fissure. Single zeolite veins are common but many veins contain at least two different zeolites. The structures of multi-zeolite veins show that zeolites never co-precipitate but rather are always diachronous.

Laumontite is the most common zeolite. It occurs in mono-zeolite veins and in fissures with stilbite or scolecite or both. Samples from the Gotthard NEAT tunnel (Table 2.2; TW01.2, TW01.3, TW02) consistently show that early laumontite is ZEOLITES IN BASEMENT ROCKS 38 overgrown by late stilbite. This observation is confirmed by data from the Gotthard road tunnel (Stalder et al., 1980). Samples from Arvigo (Table 2.2; A8) demonstrate that early scolecite is overgrown by laumontite (Fig. 2.6). However, one sample suggests an inverse growth relationship in which laumontite formed before scolecite (Table 2.2; Arvigo1).

Table 2.2. Zeolite assemblages in Alpine fissures analyzed in this study No- Mineral assemblages Sourcea, b, c Nod Description A3* Stb Fr 24 light brownish Stb crystals, Val Strem A4* Qtz-Kfs-Stb P.A 47 Stb covers Qtz and Kfs as dense mats, up to 3 mm long white euhedral needles, Gotthard NEAT A4190 Hm-Cc-Lmt SNHMB 26 Lmt on Cc scalenoeder, which overgrows Hm, Camperio A5203* Qtz-Heu SNHMB 1 euhedral Heu crystals up to 1cm in size growing on top of Qtz crystals in a fissure of altered sericite-gneiss, Oberaar/Grimsel A6* Qtz-Stb P.A. 47 Stb forming flat-topped crystals and fan-like crystal aggregates, grown on Qtz, Gotthard NEAT A8* Kfs-Prh-Ep-Chl-Sco-Lmt Fr 4 Sco needles and Lmt on top of Kfs, Ep, Prh and Chl in a fissure hosted in an orthogneiss, Arvigo/Val Calanca A8.1* Qtz-Kfs-Ep-Chl-Sco-Heu P.A. 4 Sco needles and Heu on top of Kfs, Ep and Chl in a fissure hosted in an orthogneiss, Arvigo/Val Calanca A8.2* Heu-Stb Fr 24 coffin shaped Heu crystals with a blocky habit, followed by Stb, as dense mats of crystals, up to 4 mm in size, Val Strem A8332* Kfs-Ap-Chl-Cha SNHMB 55 euhedral Cha crystals up to 5 mm, associated with Kfs and Ap; Cha appear to be green, because of Chl inclusions, Fieschergletscher Arvigo1* Ep-Prh-Sco-Lmt Fr 4 Sco tufts and Lmt cover Ep and Prh in a fissure hosted in an orthogneiss Arvigo/Val Calanca Arvigo12 Lmt Fr 4 anhedral Lmt crystals, fills up a highly porous zone in a leaching zone I* hosted in an orthogneiss, Arvigo/Val Calanca Arvigo13 Sco Fr 4 Sco needles up to 4 cm in length located in fissure, Arvigo/Val * Calanca Arvigo2* Lmt Fr 4 Lmt, which totally fill up a 4 cm wide and 10 cm long boudinage gash; crystal sizes up to 2 cm. Arvigo/Val Calanca B12* Sco-Heu-Stb Fr 24 euhedral crystals of Sco, Heu and Stb, Schattig Wichel, B981* Chl-Lmt SNHMB 44 green Lmt crystals up to 1 cm in length, green colour of Lmt appear because of Chl inclusions, Gotthard road tunnel B3140 Qtz-Kfs-Cc-Chl-Cc-Sco SNHMB 4 fissure mineralization in chloritized gneiss, Arvigo, Chaba1* Qtz-Cha Fr 51 rhombohedral, transparent Cha crystals, up to 2 mm in size, grown after Qtz in altered granite, Schijenstock Chaba2* Cha Fr 28 rhombohedral Cha crystals in a fissure of granite Stella/Gotthard DT* Heu-Stb Fr 11 light brownish Stb crystals, forming 5 cm long fanlike bow ties, which associated with early formed Heu crystals on top of a highly porous matrix, Drumtobel/Sedrun Fi1* Qtz-Heu-Stb Fr 57 euhedral Stb crystals, up to 1 cm in size associated with Heu and Qtz, Gibelsbach/Fiesch Fi2* Flt-Stb Fr 57 Stb on green fluorite, Gibelsbach KB868* Qtz-Kfs-Lmt-Apo P.A. 47 Apophyllite-(F) overgrows Lmt, Kfs and Qtz in a fissure of Gotthard NEAT K614* Qtz-Cha-Stb Fr 50 smoky Qtz, covered by Cha rhomboeder and fan-shape Stb aggregates, Val Val NL4b* Lmt P.S. 47 Lmt which fills up a 0.2 mm wide vein in a highly porous and altered gneiss, Gotthard NEAT R1* Qtz-Stb Fr 49 light brounish Stb, forming radial groups of 1 cm in diameter, hosted in quartz lenses in paragneiesses, Riental TW01* Lmt P.A. 47 Lmt on Qtz, Gotthard NEAT ZEOLITES IN BASEMENT ROCKS 39 continue Table 2.2 TW01.2* Qtz-Chl-Lmt-Stb P.A. 47 Stb associated with Lmt and Chl as early phases in a fissure of the Gotthard NEAT TW01.3* Lmt-Stb P.A. 47 Lmt associated with Stb as early phase in a fissure of the Gotthard NEAT TW02* Qtz-Lmt-Stb P.A. 47 Lmt on Qtz associated with Stb as early phase in a fissure of the Gotthard NEAT TW03* Cc-Lmt P.A. 47 Lmt associated with early formed Cc in a fissure of Gotthard NEAT TW11.1* Qtz-Chl-Lmt P.A. 47 Lmt associated with early formed Chl and Qtz in a fissure of Gotthard NEAT TW20* Sco-Heu-Stb Fr 50 euhedral crystals of Sco, Heu and Stb, Schattig Wichel TW34* Heu-Stb Fr 24 light brownish Stb crystals, forming 5 mm long bow ties, which associated with early formed Heu crystals, on a highly porous matrix, Val Strem 4172 Cc-Stb SNHMB 37 Stb grow on Cc, Brunnital 7844 Qtz-Chl-Cc-Sco SNHMB 50 Sco and Cc growing on Chl and Qtz, Schattig Wichel 7890 Cc-Heu SNHMB 46 Heu on Cc scalenoeder, Maderaneetal 30992 Qtz-Cc-Lmt SNHMB 2 Lötschberg exploring tunnel 35370* Stb SNHMB 44 light brownish Stb crystal, up to 3 mm, grown on top of fine Qtz crystals, fissure breccia 2830 meter after north portal, Gotthard road tunnel 35508* Kfs-Chl-Stb SNHMB 44 light pinkish Stb, southern granite gneisses, (3240m) Gotthard road tunnel 35795* Lmt SNHMB 44 Lmt from the Gotthard road tunnel 35843* Sco-Stb SNHMB 44 Stb on top of Sco from a fissure of the Gotthard road tunnel 36728* Qtz-Cc-Ms-Ttn-Stb SNHMB 44 Stb crystals up to 5 mm growing on Cc, Fibbia granite gneisses, Gotthard road tunnel s SNHMB = collection of the Natural History Museum Bern. b P.A. = sample provided by Peter Amacher. c Fr = sample collected during field work or owned by the Mineralogical Museum University Freiburg. d number corresponds to Table 2.1 and Figs 2.2 & 2.3. * = sample analysed during this study

Frequently heulandite succeeds scolecite (Table 2.2; A8.1, TW20; Fig. 2.6). Single heulandite crystals are penetrated by older scolecite tufts (Fig. 2.6e). In some veins additional stilbite forms the latest zeolite in this succession (Sco → Heu → Stb). Heulandite occurs either in the assemblage scolecite-heulandite or heulandite-stilbite or scolecite-heulandite-stilbite, whereas the assemblage laumontite-heulandite has not been found. Heulandite in mono-zeolite veins is rare and occurs occasionally grown on calcite as substrate mineral (Table 2.2; 7890).

Stilbite appears with all other zeolites except epistilbite, but it also appears frequently in mono-zeolite fissures (Table 2.2; A4, A6, R1, 35370). In samples from the NEAT tunnel stilbite associated with earlier formed laumontite. Overgrowth of stilbite on heulandite can be found at all zeolite localities (Table 2.2; DT, Fi1, TW34). Stilbite is always the last zeolite in the succession (Table 2.2). Chabazite occurs widespread but in small amounts and irregularly. It is subordinate to laumontite, stilbite, scolecite and heulandite (Fig. 2.2). It is associated with and overgrown by ZEOLITES IN BASEMENT ROCKS 40 later formed stilbite (Fig. 2.6; Table 2.2; K614). At some localities it also occurs in single zeolite veins (Table 2.2; A8332, Chaba1, Chaba2). Chabazite has never been observed together with heulandite, laumontite and scolecite. The rarest zeolite is epistilbite, which is only known from 3 localities (Table 2.1; no 4 25, 57). It is not associated with other zeolite species.

From these observations the following growth chronology can be deduced (from old to young): Sco → Lmt → Heu → Cha → Stb. Note that wairakite, a common Ca- zeolite elsewhere, has not been found in the Central Alps.

Most zeolites occur associated with earlier formed minerals in the veins. These minerals include silicates, oxides, sulfides, sulfates, carbonates, phosphates and halides (Table 2.1). Zeolites always overgrow all these minerals (except apophyllite). Although a large number of different minerals occur together with zeolites, in a single fissure not more than six different minerals are normally present. A full list of minerals occurring with zeolites in fissures from different localities is given in Table 2.1 and the data from the analysis of mineral assemblages in this study are given in Table 2.2. It is evident from the data compilation that zeolites do not occur preferentially in veins with a preferred mineral assemblage. The data presented in Table 2.2 demonstrate that zeolites occur unrelated to the type of minerals deposited in the vein prior to zeolite formation.

Table 2.3. Crystallization chronology of zeolite bearing Alpine fissures determined in this study. quartz adularia epidote prehnite chlorite zeolites apophyllite ± titanite calcite calcite - scolecite ± actinolite fluorite - laumontite ± ilmenite ± hematite - heulandite ± anhydrite - chabazite - stilbite

early late

Zeolites grow on a substrate, which is either the host rock surface or an earlier formed mineral. Preferential substrate minerals are quartz, adularia and calcite. Locally special and unique assemblages in fissures associated with zeolites occur. For instance, the assemblage quartz - adularia - green fluorite - zeolite occurs in a fissure breccia at Gibelsbach near Fiesch (Figs 2.2 & 2.6d; Table 2.1; no 57). A fissure at Tiefengletscher (Table 2.1; no 55) exhibits a unique paragenetic relation of laumontite overgrowth on galena and smoky quartz. Often titanium minerals, like titanite, ZEOLITES IN BASEMENT ROCKS 41 brookite, rutile and ilmenite appear in fissures as accessory phase in minor amounts. Remarkable is the sporadic presence of sulfide and sulfate minerals in veins from the same rock type. For instance in the Central Aar granite at tunnel level (Gotthard NEAT) the assemblages anhydrite-chlorite-laumontite-stilbite and quartz-chlorite- calcite-pyrite-laumontite are present (P. Amacher, pers. com.). Museum quality hematite roses from La Fibbia and Pizzo Lucendro (Table 2.1; no 28, 31), are associated with stilbite in the complete paragenesis of quartz-adularia-muscovite- hematite-rutile-stilbite (Stalder et al., 1998). Similar assemblages are known from the same Fibbia gneiss, approximately 1500 m beneath La Fibbia at the level of the Gotthard road tunnel (Stalder et al., 1980, 1998).

The most common mineral that occurs together with zeolites is quartz (Figs 2.5 & 2.6; Table 2.1 & 2.2). If quartz and zeolites are the only minerals in a fissure, the zeolite grows directly on the crystal surface of euhedral quartz (Table 2.2; A5203, A6, A8.1, Fi1). Substrate quartz does not show any sign of leaching and dissolution. Alpine fissure quartz crystals overgrow and include a wide range of different minerals such as chlorite, rutile, anhydrite and many others (Stalder et al., 1998). However, zeolite inclusions in quartz are unknown. This implies that quartz growth always precedes the onset of zeolite growth. Additional phases in fractures of gneiss and granite are often adularia and chlorite, which are together with quartz the most common minerals in Alpine fissures (Stalder et al., 1998). Sample KB868 (Table 2.2) provides a full growth chronology of these phases with the sequence quartz → adularia → laumontite → apophyllite. Similar to quartz-substrate, the contact surface between adularia and laumontite is planar and even without leaching or dissolution textures (Table 2.2; A8332, A8). Chlorite occurs in large quantities in Alpine fissures and in association with zeolites and marks the most frequent Mg-Fe silicate in Alpine fissures (Stalder et al., 1998). Chlorite very often occurs as unconsolidated chlorite sand in fissures forming vermicular grains, with grain sizes less than 2 mm. Because of the loose nature, chlorite is reworked and often fills up open spaces in the porous rock matrix and in open spaces between older minerals as well as between younger minerals. Inclusions of chlorite in later formed species including zeolites give them a green color typical of many localities (Table 2.2; A8332). The observation suggests that chlorite formed prior to zeolites. In addition to the frequent substrate minerals quartz and adularia, also calcite often serves as substrate for zeolite growth (Table 2.1 ZEOLITES IN BASEMENT ROCKS 42

& 2.2; TW03, 7844, A4190, 30992, 7890, 4172). Calcite may occur in different generations during vein mineralization (Table 2.3). Calcite formed before and after chlorite growth and it occurs with all zeolites except chabazite and epistilbite.

Fig. 2.6: Representative zeolite assemblages from the Central Swiss Alps. (a) Radial stilbite aggregates overgrows chabazite rhombohedra on quartz; Val Val/GR. (b) Fan-shape stilbite overgrows coffin shaped heulandite crystals; Val Strem/GR. (c) Laumontite fibrous; Gotthard NEAT/UR. (d) Stilbite associated with green fluorite; Gibelsbach/VS. (e) Coffin shaped heulandite grows on scolecite needles, which grow on quartz covering highly porous rock matrix; Schattig Wichel/UR. (f) Scolecite grows on Alpine fissure quartz; Gotthard NEAT tunnel/UR. (g) Scolecite on epidote, adularia and quartz; Arvigo/GR. (h) Laumontite and scolecite associated with calcite that shows dissolution Arvigo/GR.

ZEOLITES IN BASEMENT ROCKS 43

Calcite often occurs as scalenoeders on which zeolites developed (Table 2.2; 30992, 7890). It also occurs as paper spar that often shows evidence of dissolution and resorption (Table 2.2; TW03; Fig. 2.6).

Prehnite and epidote are Ca-Al silicate minerals that formed prior to the zeolites at some localities (Table 2.1 & 2.2). Actinolite that often forms asbestos fibers also occurs locally. Pumpellyite, a diagnostic mineral of the prehnite-pumpellyite and pumpellyite-actinolite metamorphic facies, has never been found in Alpine gneiss and granite fractures. However prehnite and epidote do occur in granite and gneisses (Table 2.1 & 2.2), but zeolites are often lacking (Stalder et al., 1998) in the succession. The Arvigo quarry, fissures around Piz Giuv and the Gotthard NEAT tunnel (Table 2.1 & 2.2) represent localities were prehnite and epidote are frequently found associated with zeolites. The mineral sequence deduced from these localities is: quartz-adularia-actinolite-epidote-prehnite-zeolite (Table 2; A8).

2.5.2.2. Aar Massif/Gotthard NEAT tunnel

During the ongoing excavation of the Gotthard NEAT tunnel (NEw Alp Transit, 57 km long rail base tunnel through the Central Alps, Figs 2.2 & 2.3) a large number of mineralized fissures and veins were opened, many of which contained a museum quality mineral specimens (Figs 2.5 & 2.6). Of particular interest for the present research was tunnel section under Chrüzlistock (Figs 2.3 & 2.7) and the surface outcrops above to the tunnel section, where zeolites have been frequently found in fissures (e.g. Kenngott, 1866; Parker, 1922; Niggli et al., 1940; Huber, 1943; Sigrist, 1947; Stalder et al., 1998). In other sections of the NEAT tunnel no zeolites have been found so far or the tunnel drilling is still under progress.

Figure 2.7 gives an overview of the main fissure mineral assemblages in relation to the lithologies in the tunnel section, where zeolites have been found. Zeolite occurrences in surface outcrops vertically above the tunnel correspond to those in the tunnel section. In surface outcrops zeolites are concentrated around the Giuv syenite (Fig. 2.3) a W-E trending igneous body of syenitic composition (Labhart, 1977). But the syenite-unit pinches out with depth and is not present on tunnel level (500 m) (Fig. 2.7). In total more than 25 different mineral species have been found in veins and fractures in the studied section of the NEAT tunnel (7950-8850 meter in distance ZEOLITES IN BASEMENT ROCKS 44 to the north portal). Adularia, albite, amianthus, calcite, chlorite, pyrite and quartz occur evenly distributed over the whole section. Anatase, anhydrite, ankerite, apatite, apophyllite, chalcopyrite, epidote, fluorite, galena, graphite, hematite, milarite, spahlerite, synchisite and titanite occur only sporadically or in traces (P. Amacher pers. com.). In surface outcrops the following additional minerals were found associated with zeolites (Huber, 1948): datholite, limonite, molybdenite and prehnite. To be mentioned is the well-known “Skolezitkehle” (Huber, 1948), located at Schattig Wichel, the north wall of Piz Giuv, where a numerous zeolite-bearing fissures were found.

Fig. 2.7: Main fissure mineral assemblages in the Gotthard NEAT tunnel. Profile shows the drilling section Amsteg-Sedrun, from tunnel meter 12000 to 18550 (distance from north portal at Erstfeld), where zeolite minerals were recovered. In the section between 7950 and 12000 meter no zeolites were found in fissures. Drilling in the three other tunnel sections is still in progress. Tunnel level is 500 meter above sea level and overburden ranges between 1290 and 2130 meter. The position of Chrüzlistock (Fig. 2.3) is marked with a black star; (data derived from P. Amacher, pers. com.).

In the tunnel section laumontite is the most abundant zeolite mineral, which is in sharp contrast to observations from surface outcrops (Table 2.1), where laumontite occurs only sporadically in small amounts (Huber, 1943; Sigrist, 1947; Stalder et al., 1998). Stilbite (stellerite) follows as the second most common zeolite in the tunnel, whereas scolecite, heulandite and chabazite occur only sporadically in the tunnel, ZEOLITES IN BASEMENT ROCKS 45 again in contrast to observations from surface outcrops where stilbite, scolecite and heulandite occur in similar proportions and chabazite is relatively rare (e.g. Huber, 1948).

Fissures that are filled with zeolites only indicate a late formation of the fracture. A systematic study of fissure chronology by Heijboer (2006) focused on quartz formation. It showed that zeolite species formed during the last mineral precipitation phase, during which the orientations of fissure and veins are generally SE-NW or NE- SW, respectively (Huber, 1946; Heijboer, 2006).

2.5.2.3. Gotthard massif/ Gotthard road tunnel

Excavation of the 16 km long Gotthard road tunnel in the seventies of the last century (Fig. 2.2) supplied considerable amounts of fissure minerals including zeolites from the Gotthard massif (Stalder et al., 1980). This amount of material is in distinct contrast to the small number of finds of zeolite veins from surface outcrops in the Gotthard massif (Fig. 2.2; Table 2.1). Thus zeolites are very frequent in both the Gotthard road tunnel and in the Gotthard NEAT tunnel.

The spatial distribution of zeolites in the Gotthard road tunnel can be related to the type of host rock of the fissure. The host rock controls the proportion of fissures filled with zeolites and the nature of the dominant zeolite present (Fig. 2.8). A detailed study of minerals in fissures of the Gotthard road tunnel during excavation by Stalder et al. (1980), has shown, that zeolites occur in different modal proportions in different lithologies but that they are not present in all rock types. Figure 2.8 gives an overview of fissure bearing lithologies and the relative amount of fissure that contain zeolites. The units along the tunnel sections are (from N to S): Central Aar granite (ZArg), Southern gneisses (SGn), Permian-Carboniferous (PC), Gurschen gneiss

(GGn), Fibbia granite gneiss (FGgn) and Tremola Series (TrS). Where CO2 fluids were found in fluid inclusions of quartz, e.g. in GGn fissures are zeolites devoid and in the PC unit only a small number of fissures (<10 % of all fissures) contain laumontite (Stalder et al., 1980). Laumontite with minor amounts of stilbite, heulandite scolecite and chabazite are typically found in the ZArg (17 % of fissure hold laumontite), SGn (45 %) and TrS (25 %) units. A distinctly different zeolite population pattern is observed in the FGgn rock unit. Stilbite (>50 %) represents the ZEOLITES IN BASEMENT ROCKS 46 major zeolite in the Fibbia gneisses, whereas laumontite, scolecite and chabazite were only found in less than 10 % of all fissures (Stalder et al., 1980). Similar to the Gotthard NEAT tunnel apophyllite is associated with laumontite and marks a younger fissure generation.

Fig. 2.8: Simplified geological cross section through the Gotthard road tunnel and incidence of zeolites (% of fissures with zeolites, n = number of fissures) in fissures for different lithological units (Stalder et al., 1980). (ZAgr: Central Aar granite, SGn: Southern gneisses, Me: Mesozoic units, PC: Permian- Carboniferous, GGn: Gurschen gneiss, GGr: Gamsboden granite gneisses, GGn: Guspis gneiss, FGgn: Fibbia granite gneiss, SGn: Sorescia gneiss, TrS: Tremola Series.

2.5.2.4. Gibelsbach/Fiesch

The mineral fissures from Gibelsbach (Fig. 2.2) have been described already by Kenngott (1866). The mineral assemblage of green octahedral fluorite, quartz, adularia, albite and various zeolite species (Fig. 2.6d) is unique to this locality. In addition to the five major zeolites (scolecite, laumontite, heulandite, stilbite/stellerite, and chabazite) Gibelsbach is one of three localities (Table 2.1; no 4, 25), where epistilbite is found. Single crystal studies have shown that the mineral previously identified as stilbite is in fact stellerite, the orthorhombic Ca-endmember equivalent of monoclinic stilbite (Armbruster et al., 1996).

The Gibelsbach mineral veins are hosted in a brecciated, highly porous and strongly foliated granite (Armbruster et al., 1996). The zeolite-bearing zone is bordered to the south by Permian sedimentary rocks occurring between Aar- and ZEOLITES IN BASEMENT ROCKS 47

Gotthard Massif and to the north by a coarse grained granite of the southern part of the Aar Massif (Zbinden, 1949) (Fig. 2.2). Important at this locality is that in contrast to most other zeolite localities chlorite or other iron-magnesium silicates have not been observed in the outcrop.

2.5.2.5. Arvigo/Val Calanca

Banded biotite gneiss and coarse-grained light colored two-mica gneiss are mined as building stones in Arvigo (Fig. 2.4a). The rocks belong to the upper Simano nappe of the crystalline Penninic basement. The Arvigo quarry became famous for a large number of Alpine fissure minerals, which occur in extensional fractures and cavities of the granitic gneisses. Arvigo is the prime zeolite locality within the Penninic nappes (Table 2.1; no 25, 30, 32, 35). The Arvigo fissures contain the assemblage quartz, adularia, epidote, prehnite, chlorite and all zeolite species known from Alpine fissures. Apophyllite is present as the latest mineral. In general, pale green sheaves of epidote are overgrown by prehnite, chlorite and zeolites marking the most common mineral assemblages in Arvigo (Fig. 2.6g). Scolecite and laumontite are the dominant zeolite species, whereas heulandite, chabazite, stilbite and epistilbite can only found sporadically. The occurrences of mesolite (Weiß & Forster, 1997) as overgrowth on scolecite could not be confirmed during this study. Comparing to other surface outcrops, the Arvigo quarry shows the highest abundance of laumontite in the Alps. This is related to the active quarrying, which reveals a steady supply of fresh unaltered material. More than 40 different minerals species are found in fissures of the Arvigo quarry, most of the minerals are rare (full list of minerals see Wagner et al., 2000a and b). Remarkable is the appearance of babingtonite, which usually is associated with zeolites in basic igneous rocks (Armbruster, 2000; Armbruster et al., 2000). The mineral assemblages vary with host rock composition and mineralogy, but zeolite species are always present as latest mineral species in the fractures. Fluid-rock interaction with the host rock, forms noticeable leaching zones (Fig. 2.4). These 6-7 cm wide leaching zones are usually lighter colored than the dark country rock, due to the removal of mafic minerals. The high abundance of zeolites in Arvigo is remarkable and unique for the crystalline Penninic nappes. There are a large number ZEOLITES IN BASEMENT ROCKS 48 of quarries in the area quarrying similar rocks from the same or the neighboring nappes but not in any of them comparable zeolites finds have been made.

2.5.3. Mineralogy and crystal chemistry of zeolites and associated minerals

2.5.3.1. Chabazite-Ca

Chabazite ((Ca0.5,Na,K)4(Al4Si8O24) •12 H2O, Passaglia & Sheppard, 2001) forms granular pseudorhombohedral, 1 to 4 mm large, transparent to translucent colorless or white crystals (average 2 mm; Fig. 2.6a). The triclinic crystals (Armbruster & Gunter, 2001) often form penetrating twins with corners projecting from the faces. It grows on fissure quartz crystals (rock crystals) and is often associated and overgrown by stilbite.

Chabazite in Alpine fissures can be classified as chabazite-Ca (Coombs et al., 1998), because the main extra-framework cation is Ca. Ca occupies on an average 60 mole% of the extra-framework cations (range from 56 to 66 mole%). Representative analyses are given in Table 2.4. K and Sr are present in high proportions in the channels structure of chabazite (Fig. 2.9), the chabazite-K content ranges from 24 to 35 mole%, with an average value of 30 mole% of all extra-framework cations. Sr occupies up to 10 mole% of the extra-framework sites, making Sr to an important zeolite cation similar to heulandite. Other extra-framework cations such as Na, Mg, Ba and Fe occur in traces only and can therefore be neglected. Fissure chabazite Si/(Si+Al) ratio of 0.70 to 0.71 (Fig. 2.10), which is higher than the mean value of amygdaloidal chabazite crystals (0.67; Passaglia & Sheppard, 2001) that can be related to the coupled substitution K+ + Si4+ = Ca2+ + Al3+. Zoned chabazite shows that the K content increases and the Ca decreases from core to the rim.

ZEOLITES IN BASEMENT ROCKS 49

Table 2.4. Representative zeolite analysis Sample no. A8332 Chaba1 Chaba2 Chaba2 A5203 TW20 TW20 A8.1 Fi1 TW34 Analysis no. 2 2 3 4 1 6 7 12 4 6 wt.% Cha Cha Cha Cha Heu Heu Heu Heu Heu Heu

SiO2 52.76 53.59 54.37 53.80 54.64 60.74 59.10 56.30 56.67 57.76

Al2O3 19.47 19.24 19.51 19.53 16.51 15.99 15.08 16.56 15.60 15.61 CaO 8.35 8.52 8.00 8.48 5.47 6.77 6.36 6.87 6.07 5.99 SrO 1.88 1.58 2.18 2.29 4.35 2.40 2.35 0.92 2.31 3.08 BaO -- 0.08 0.01 0.00 - - - - 0.00 0.28

Na2O 0.17 0.08 0.20 0.15 0.14 0.03 0.03 0.38 0.11 0.07

K2O 3.75 3.72 3.70 2.73 1.53 1.36 1.36 2.47 2.28 1.40

H2O 13.58* 13.14* 12.02* 13.00* 17.30* 12.67* 15.68* 16.47* 16.93* 15.75*

Totala 100.00 100.00 100.00 100.00 100.00 100.00 100.00 100.00 100.00 100.00

Si 8.300 8.364 8.385 8.360 26.573 27.446 27.636 26.713 27.139 27.276 Al 3.610 3.539 3.546 3.577 9.463 8.517 8.312 9.259 8.805 8.688 Ca 1.407 1.425 1.322 1.412 2.850 3.276 3.184 3.491 3.115 3.031 Sr 0.171 0.143 0.195 0.206 1.227 0.629 0.638 0.253 0.641 0.843 Ba - 0.005 0.001 0.000 - - - - 0.000 0.052 Na 0.052 0.024 0.060 0.045 0.132 0.027 0.024 0.350 0.102 0.064 K 0.753 0.741 0.728 0.541 0.949 0.786 0.811 1.496 1.393 0.843 O 24 24 24 24 72 72 72 72 72 72

H2O 7.125 6.840 6.183 6.737 28.074 19.088 24.452 26.073 27.041 24.807

E%b -8.89 -9.92 -7.23 -6.55 2.47 -1.50 -2.17 -0.81 -2.40 -1.60 Si/(Si+Al) 0.70 0.70 0.70 0.70 0.74 0.76 0.77 0.74 0.76 0.76 Ca/(Ca+Na+K+Sr) 0.59 0.61 0.57 0.64 0.55 0.69 0.68 0.62 0.59 0.63 Na/(Ca+Na+K+Sr) 0.02 0.01 0.03 0.02 0.03 0.01 0.01 0.06 0.02 0.01 K/(Ca+Na+K+Sr) 0.32 0.32 0.32 0.25 0.18 0.17 0.17 0.27 0.27 0.18 Sr/(Ca+Na+K+Sr) 0.07 0.06 0.08 0.09 0.24 0.13 0.14 0.05 0.12 0.18

Sample no. DT A8.2 A8 TW11.1 TW03 TW01.2 35795 NL4b Arviog12I Arvigo2 Analysis no. 5 3 2 6 1 8 3 9 10 7 wt.% Heu Heu Lmt Lmt Lmt Lmt Lmt Lmt Lmt Lmt

SiO2 57.53 57.66 52.23 53.00 53.18 52.33 51.49 51.95 52.06 52.43

Al2O3 15.08 15.96 21.89 20.87 20.64 21.46 21.63 20.64 21.48 21.15 CaO 5.96 5.98 12.20 11.35 11.28 11.97 12.02 11.06 11.65 11.82 SrO 1.95 3.72 0.03 0.12 0.03 0.05 0.24 0.27 0.17 0.00 BaO 0.41 0.50 0.00 - - - - 0.00 0.00 0.00

Na2O 0.26 0.08 0.04 0.12 0.08 0.04 0.08 0.16 0.00 0.09

K2O 1.97 1.51 0.05 0.33 0.33 0.05 0.05 0.46 0.32 0.12

H2O 16.81* 14.532* 15.24 14.17* 14.457* 14.08* 14.49* 15.44* 14.19* 14.37*

Totala 100.00 100.00 101.69 100.00 100.00 100.00 100.00 100.00 100.00 100.00

Si 27.442 27.062 16.036 16.362 16.443 16.149 16.013 16.313 16.124 16.228 Al 8.478 8.827 7.923 7.594 7.521 7.806 7.928 7.640 7.841 7.713 Ca 3.046 3.005 4.014 3.754 3.735 3.957 4.005 3.721 3.866 3.919 Sr 0.539 1.013 0.005 0.021 0.005 0.008 0.043 0.049 0.031 0.000 Ba 0.077 0.092 0.000 - - - - 0.000 0.000 0.000 Na 0.240 0.068 0.023 0.072 0.050 0.023 0.048 0.099 0.000 0.053 K 1.199 0.905 0.019 0.130 0.129 0.021 0.020 0.185 0.126 0.047 O 72 72 48 48 48 48 48 48 48 48

H2O 26.743 22.747 13.884 14.591 14.909 14.486 15.029 16.170 14.658 14.837 E%b -3.57 -4.80 -1.95 -2.06 -1.81 -2.11 -2.90 -2.36 -0.99 -2.85

Si/(Si+Al) 0.76 0.75 0.67 0.68 0.69 0.67 0.67 0.68 0.67 0.68 Ca/(Ca+Na+K+Sr) 0.61 0.60 0.99 0.94 0.95 0.99 0.97 0.92 0.96 0.98 Na/(Ca+Na+K+Sr) 0.05 0.01 0.01 0.02 0.01 0.01 0.01 0.02 0.00 0.01 K/(Ca+Na+K+Sr) 0.24 0.18 0.00 0.03 0.03 0.01 0.00 0.05 0.03 0.01 Sr/(Ca+Na+K+Sr) 0.11 0.20 0.00 0.01 0.00 0.00 0.01 0.01 0.01 0.00 ZEOLITES IN BASEMENT ROCKS 50 continue Table 2.4 Sample no. TW20 35843 A8 A4 Arvigo1 Arvigo13 Arvigo13 TW02 TW01.2 35370 Analysis no. 2 3 3 1 6 3 10 2 4 1 wt.% Sco Sco Sco Sco Sco Sco Sco Stb Stb Stell

SiO2 45.72 44.64 45.67 45.39 45.64 46.02 46.20 58.12 63.07 59.13

Al2O3 24.86 24.44 24.70 25.14 24.65 24.96 25.14 15.21 17.09 14.12 CaO 14.05 13.63 13.90 13.69 14.14 14.07 13.99 7.75 8.72 7.87 SrO 0.02 0.04 0.00 0.10 0.00 0.04 0.00 0.06 0.00 0.00 BaO - - - 0.00 0.00 0.00 0.01 - - -

Na2O 0.04 0.09 0.22 0.09 0.10 0.14 0.10 0.41 0.91 0.02

K2O 0.01 0.02 0.00 0.00 0.00 0.02 0.00 0.07 0.02 0.04

H2O 15.25* 17.09* 15.45* 14.01 15.38* 14.65* 14.55* 18.37* 10.18* 18.80*

Totala 100.00 100.00 100.00 98.43 100.00 100.00 100.00 100.00 100.00 100.00

Si 24.298 24.249 24.330 24.200 24.312 24.296 24.325 27.546 27.255 28.062 Al 15.569 15.647 15.507 15.797 15.476 15.532 15.599 8.495 8.705 7.898 Ca 8.000 7.933 7.932 7.819 8.070 7.961 7.890 3.938 4.035 4.002 Sr 0.006 0.013 0.000 0.032 0.000 0.013 0.000 0.018 0.000 0.000 Ba - - - 0.000 0.000 0.000 0.001 - - - Na 0.044 0.095 0.228 0.091 0.103 0.139 0.104 0.378 0.763 0.018 K 0.005 0.014 0.000 0.000 0.000 0.012 0.003 0.041 0.013 0.024 O 80 80 80 80 80 80 80 72 72 72

H2O 27.035 30.966 27.451 27.700 27.325 25.798 25.550 29.035 14.671 29.760

E%b -3.08 -2.30 -3.64 0.01 -4.73 -3.66 -1.82 2.00 -1.81 -1.84 Si/(Si+Al) 0.61 0.61 0.61 0.61 0.61 0.61 0.61 0.76 0.76 0.78 Ca/(Ca+Na+K+Sr) 0.99 0.98 0.97 0.98 0.99 0.98 0.99 0.90 0.84 0.99 Na/(Ca+Na+K+Sr) 0.01 0.01 0.03 0.01 0.01 0.02 0.01 0.09 0.16 0.00 K/(Ca+Na+K+Sr) 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.01 Sr/(Ca+Na+K+Sr) 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Sample no. 35370 35843 35843 R1 R1 Fi1 DT A4 A6 A6 Analysis no. 5 7 10 1 2 7 9 5 4 7 wt.% Stb Stb Stb Stb Stb Stb Stb Stb Stb Stb

SiO2 58.11 61.46 61.62 55.94 56.91 57.40 61.36 58.15 56.17 61.80

Al2O3 14.14 15.64 15.54 15.37 15.36 17.21 15.99 15.14 15.04 17.11 CaO 7.74 8.52 8.52 7.23 7.62 8.65 7.41 7.93 7.72 8.55 SrO 0.00 0.00 0.00 0.04 0.00 0.17 0.00 0.03 0.09 0.06 BaO - - - 0.00 0.02 0.00 0.03 0.06 0.00 0.08

Na2O 0.27 0.35 0.23 0.36 0.11 0.30 0.76 0.56 0.52 0.81

K2O 0.03 0.06 0.07 1.50 0.95 0.87 1.26 0.00 0.04 0.05

H2O 19.71* 13.90* 13.91* 20.64 20.64 15.34* 13.17* 18.06* 20.36* 11.46*

Totala 100.00 100.00 100.00 100.00 100.00 100.00 100.00 100.00 100.00 100.00

Si 27.932 27.627 27.691 27.185 27.342 26.579 27.527 27.509 27.353 27.133 Al 8.010 8.286 8.230 8.803 8.698 9.392 8.454 8.443 8.630 8.854 Ca 3.986 4.103 4.102 3.765 3.923 4.291 3.562 4.017 4.030 4.022 Sr 0.000 0.000 0.000 0.011 0.000 0.046 0.000 0.008 0.025 0.014 Ba - - - 0.000 0.004 0.000 0.005 0.011 0.001 0.013 Na 0.252 0.305 0.200 0.339 0.102 0.269 0.661 0.509 0.487 0.692 K 0.018 0.034 0.040 0.930 0.582 0.514 0.721 0.001 0.023 0.029 O 72 72 72 72 72 72 72 72 72 72

H2O 31.598 20.848 20.854 33.454 33.074 23.691 19.706 28.497 33.080 16.789

E%b -2.81 -3.05 -2.54 -0.52 1.88 -0.98 -0.88 -1.61 -0.14 0.39 Si/(Si+Al) 0.78 0.77 0.77 0.76 0.76 0.74 0.77 0.77 0.76 0.75 Ca/(Ca+Na+K+Sr) 0.94 0.92 0.94 0.75 0.85 0.84 0.72 0.89 0.88 0.85 Na/(Ca+Na+K+Sr) 0.06 0.07 0.05 0.07 0.02 0.05 0.13 0.11 0.11 0.15 K/(Ca+Na+K+Sr) 0.00 0.01 0.01 0.18 0.13 0.10 0.15 0.00 0.01 0.01 Sr/(Ca+Na+K+Sr) 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.01 0.00 * a b H2O calculated by difference. Totals include traces of Mg, Ti, Mn and Fe. E % = 100*((Al)- (Na+K)+2(Mg+Ca+Sr+Ba)/(Na+K)+2(Mg+Ca+Sr+Ba)), measure of charge balance

ZEOLITES IN BASEMENT ROCKS 51

2.5.3.2. Heulandite-Ca

The heulandite group of minerals is represented by heulandite (Na,K)Ca4(Al9Si27O72)

•24 H2O and clinoptilolite (Na,K)6(Al6Si30O72) •24 H2O (Armbruster & Gunter, 2001). In Alpine fissures only heulandite occurs as member of the heulandite group. Heulandite forms crystals up to 12 mm in size, but the average size of the crystals is 1 to 4 mm. The monoclinic crystals of space group 2/m occur in tabular habit parallel {010} and with its typical coffin-shaped appearance (Fig. 2.6b, e). Heulandite crystals are transparent to translucent and colorless or white in color and have a subconchoidal to uneven cleavage.

Fig. 2.9: Extra-framework cations (Ca-Na-K-Sr) in Alpine zeolites. All zeolites are dominated by Ca. Heulandite and chabazite incorporate a significant amount of Sr of up to 30 mole%.

Representative analyses of heulandite are given in Table 2.4. The average composition of heulandite, determined by 46 microprobe analysis from 9 different samples (Table 2.2), shows the composition of Ca3.09Na0.12K1.01Sr0.80(Al8.81Si27.13O72)

•24 H2O, which is very similar to the heulandite composition

(Ca3.37Na0.07K0.88Sr0.55(Al8.42Si27.49O72) •24 H2O) determined by Armbruster et al. (1996) from Gibelsbach in the western Aar Massif (Fig. 2.2; Table 2.2). Heulandite from Alpine fissures is heulandite-Ca (Coombs et al., 1998), with significant amounts ZEOLITES IN BASEMENT ROCKS 52 of Sr and K (Fig. 2.9; Table 2.4). Ca occupies on an average 62 mole% of all extra- framework sites, 16 mole% are occupied by Sr (maximum 29 mole%) and K is found on 20 mole% of the sites (maximum 31 mole%). Na is below 10 mole%, in most samples Na occurs, like Ba, Mg and Fe in traces only (Fig. 2.9).

Heulandite is often zoned which Ca decreasing and K and Sr increasing from core to rim. Samples (A8.1, A8.2, DT) from Arvigo and Drumtobel/Sedrun are low in Sr and enriched in Na compared with the other samples (Fig. 2.9). The Si/(Si+Al) content range between 0.74 and 0.77 (Fig. 2.10), which agrees with the definition of heulandite (Coombs et al., 1998), that can be distinguished by clinoptilolite, 0.8 < Si/(Si+Al).

2.5.3.3. Laumontite

Laumontite is a monoclinic (space group C2/m) zeolite. It forms thin, elongated fibers or prisms elongated along the c-axis with a squared cross-section. The common crystal form of laumontite is the {110} prism. Commonly twinning occurs on {100} to form “swallow tail” or “V” twins. It is normally white with a common length between <1 to 15 mm. The cleavage is perfect, with a slight pearly luster on the broad cleavage surface. Laumontite is the most widespread zeolite in Alpine fissures, exposed underground in tunnels sections or in active quarries. Because the mineral decomposes by dehydration at room temperature and decays to a powdery mass, laumontite occurs rarely in surface outcrops.

The composition of laumontite (79 analysis from 14 samples; Table 2.2 & 2.4) formed in Alpine fissures is close to endmember composition Ca4(Al8Si16O48) •18

H2O (Armbruster & Kohler, 1992). Ca is the dominant extra-framework cation (average value of 96 mole%), with Na and K typically below 5 mole% (Fig. 2.9; Table 2.4). Maximum values for K and Na are 6 and 7 mole%, respectively. Other elements occur only in traces. The extra-framework cations Na and K increase from core to rim of zoned crystals. The Si/(Si+Al) content varies only slightly between 0.67 and 0.69 (Fig. 2.10), with an average ratio of 0.68. Alkalis increase with increasing Si and decreasing Ca and Al during growth, related to the coupled substitution of Si4+ + (Na+, K+) = Al3+ + Ca2+. ZEOLITES IN BASEMENT ROCKS 53

Fig. 2.10: R2+ - R+ - Si compositional diagram of Alpine zeolites. Si/Al ratio increases in chronological order.

2.5.3.4. Scolecite

Fibrous scolecite occur as white crystals with vitreous or slightly silky luster, forming characteristic radiating sprays. The transparent to translucent crystals are thin prismatic with squared cross sections. Scolecite needles range between 1 and 20 mm in length, normally 3 to 6 mm. Scolecite (Ca8(Al16Si24O80) •24 H2O) has the same structural framework as natrolite (Na16(Al16Si24O80) •16 H2O) and mesolite

(Na16Ca16(Al48Si72O240) •64 H2O) (Armbruster & Gunter, 2001). Minerals of this 2+ + group are distinguished by a Ca + H2O ↔ 2Na substitution and by symmetry (Gottardi & Galli, 1985). The variation in composition of natrolite, mesolite and scolecite is very small. (e.g. Deer et al., 2004). 42-microprobe analyses from 7 different samples (Table 2.2) indicate no major chemical variation (Fig. 2.9). Ca occupies an average of 99 mole% of the extra-framework cation sites (Fig. 2.9). Minor amounts of Na up to 4 mole%, but in most analyses around 1 mole% can be observed. In contrast K and Sr are not significantly incorporated in the framework structure of scolecite and therefore occur only in traces, like Ba, Mg and Fe. The ZEOLITES IN BASEMENT ROCKS 54 small substitution range can also be recognized in the Si/(Si+Al) ratio, whereas the ratio varies in a small range between 0.60 and 0.61 (Fig. 2.10).

2.5.3.5. Stilbite/Stellerite

A complete solid solution exists between stellerite (Ca4(Al8Si28O72) •28 H2O) and stilbite (NaCa4(Al9Si27O72) •30 H2O). Although the tetrahedral framework of stellerite and stilbite is identical, with the symmetry Fmmm (Gottardi & Galli, 1985), the overall crystal symmetry is different due to different locations of the extra framework cation sites. Stoichiometric stellerite is orthorhombic, with only one extra framework cation site fully occupied by Ca. In contrast there is an additional extra framework cation site in stilbite, occupied by Na. The additional Na site in stilbite leads to a reduction of symmetry from orthorhombic Fmmm in stellerite to monoclinic C2/m in stilbite (Gottardi & Galli, 1985; Quartieri & Vezzalini, 1987).

The habit of stilbite and its chemical composition is very variable. They occur as thick tabular crystals flattened on {010} with pointed terminations, as sheaf like aggregates or globular aggregates, as characteristic “bow-tie” crystals or as clusters of elongated six-faced crystals. Stilbite is transparent to translucent, colorless to white in color, whereas stilbite from surface outcrops in the Riental valley hosted in quartz veins (Table 2.1; no 49) occurs as light reddish globular aggregates. Crystal sizes of stilbite/stellerite range from 1 to 12 mm, with a frequent size between 3-5 mm. Historically and in old museum collections stilbite and stellerite were summarized by the synonym “desmine”, which is not longer an official term for zeolites (Coombs et al., 1998). Single x-ray diffraction analysis of “stilbite” from Gibelsbach by Armbruster et al. (1996), has shown that zeolites described stilbite from Alpine fissure are stellerite. Between stilbite and stellerite exists a complete solid solution (Fridriksson et al., 2001) and it is not always possible to determine, whether it stellerite or stilbite is present. The chemical composition shows a large variation (Fig. 2.9).

112 microprobe analyses from 20 samples (Table 2.2 & 2.4) were measured to evaluate stilbite composition. Different groups of stilbite can distinguished on the basis of the K content (Fig. 2.9). One sample from the Gotthard road tunnel (35370) is pure Ca-endmember thus it can be classified as stellerite (Coombs et al., 1998). ZEOLITES IN BASEMENT ROCKS 55

Chemically stellerite and stilbite-Ca can be rather similar, so the formal mineral name stellerite is restricted to specimens of nearly stoichiometric Ca4(Al8Si28O72) •28 H2O.

It is distinguished from stilbite-Ca by containing low Na2O and high SiO2. The maximum amount of Na, K, Mg and Fe in stellerite is below 0.2 atoms per formula unit (apfu) (Coombs et al., 1998) and the sample 35370 has in average 0.18 apfu on the basis of 72(O), well within the stellerite field (Coombs et al., 1998). The Si/(Si+Al) content for stellerite is constant with a ration 0.78 (Fig. 2.10). All other samples show Na-Ca zoning, with Ca high in the core and low at the rim. Sodium concentration shows an inverse pattern. Using the K content, 2 different groups can be distinguished. Samples from Gotthard road tunnel and Gotthard NEAT tunnel (Table 2.2) have a K content of less than 3 mole% of the extra-framework cation side, whereas samples from surface outcrops (Table 2.2) can be distinguished by an elevated K content between 8 and 20 mole% of the extra-framework site (Fig. 2.9). A coupled substitution of Si4+ + K+ = Al3+ + Ca2+ can be seen in stilbite samples from tunnel sections, with a maximum Na value of 0.84 (apfu), whereas the Si/(Si+Al) ratio range between 0.75 and 0.78, with an average ratio of 0.77 (Fig. 2.10). Stilbite samples from the surface indicate a coupled substitution including Na and K. These samples are characterized by elevated K contents up to 0.93 apfu and slightly higher Na values. Si/(Si+Al) ratio ranges between 0.74 and 0.77, with an average ratio of 0.76 (Fig. 2.10).

2.6. DISCUSSION

The distribution patterns of the Alpine zeolites (Fig. 2.2c) show that the vein zeolites are not products of prograde zeolite facies metamorphism. They are rather related to the cooling and uplift history of the Alpine orogen. The zeolites formed in late fractures in areas that have experienced much higher metamorphic peak conditions (Fig. 2.11) and also underwent pervasive ductile deformation. Zeolites formed when the rocks entered the brittle deformation regime and cooled to temperature below about 250˚C (see below). The ductile-brittle transition is located at 350-400˚C and can be tied to the formation of fissure assemblages that indicate the highest temperature (about 400˚C, Mullis et al., 1994) corresponding to a depth of about 12 km. Consequently early fissures do not contain zeolites but rather a sequence of minerals ZEOLITES IN BASEMENT ROCKS 56 and assemblages that are summarized in Table 2.3. The onset of zeolite formation corresponds to the point where the local cooling path of the different areas entered the “zeolite window” (Fig. 2.11). This may have happened at different times in the different areas (Fig. 2.11; all cooling path merge eventually, which may not necessarily have been the case). Zeolites overgrowing earlier formed fissure minerals (e.g. quartz; Figs 2.4 & 2.5) suggest that the fracture formed before the cooling path entered the zeolite window. Monomineralic zeolite fissures indicate, in contrast, that the fracturing occurred in the zeolite window (Fig. 2.11).

Fig. 2.11: Schematic sketch of the T-t evolution of tectonic units in the Central Alps in relation to fissure formation and the timing of zeolite growth. (a) The T-t paths of individual tectonic units reflecting an increase of the Alpine peak metamorphism from north to south. The southern units have reached the ductile regime, northern units deformed brittle. All units reached temperatures above the zeolite window (except for the parautochtonous cover rocks of the Aar massif). During uplift the units returned to the brittle deformation regime and extension fissures formed (b), subsequently zeolite- absent fissure assemblages developed (c) and finally the units entered the zeolite window (d).

The presented data show that first scolecite, then laumontite, heulandite, chabazite and finally stilbite precipitated from the hot aqueous fluids with decreasing temperature along the cooling path. This means that scolecite formed when the cooling path first entered the zeolite window, stilbite is the low-temperature zeolite that may still form in deep groundwater environments in the crystalline basement of the Central Alps such as in 40˚C fissures of the Gotthard rail base tunnel (Seelig et al., 2007). Zeolite formation thus can be related to the T-interval from 250˚C to 50˚C with a characteristic mean temperature of about 200˚C where most of the dominant zeolite laumontite formed. ZEOLITES IN BASEMENT ROCKS 57

Heulandite and chabazite require Sr in the fluid to become stable together with a pure Ca zeolite. Sr-bearing heulandite and chabazite is associated with stilbite or scolecite (Fig. 2.6). From this it also follows that heulandite associated with scolecite (Fig. 2.6e) indicate a higher temperature regime compared to chabazite associated with stilbite (Fig. 2.6a) that point to lower temperature conditions.

Chemical zoning and compositional variation of laumontite, heulandite, chabazite and stilbite indicate changes in fluid composition or temperature during growth. Similar pattern were also observed in Icelandic geothermal system (Fridriksson et al., 2001), where a Na increase with decreasing temperature during formation of stilbite was confirmed. Remarkable are two distinct geochemical patterns comparing stilbite collected in tunnels and at surface outcrops. Because samples from surface outcrops formed earlier than samples from tunnel sections, the elevated values of K in surface samples suggests a late temperature dependent ion exchange: Na+ = K+ as a result of interaction with surface water. The zoning pattern in laumontite, heulandite and chabazite seems to be related to a decrease in temperature during formation. But late ion exchange with surface water after the formation of zeolite could also explain the observed compositional patterns in zeolite. Present day deep fluids in the Aar and

Gotthard basement are dominated by Na2SO4 and Na2CO3 (Seelig et al., 2007) and would thus support the presence of an ion exchange component.

A remarkable result of this study is the observation of significant differences of the observed rare occurrences of zeolites in surface outcrops and the very abundant occurrence of zeolites in tunnel fissures. This is particularly the case for the abundance of laumontite in fissures. The mineral is extremely rare at the surface and dominant in subsurface samples. The absence of laumontite in surface outcrops is a consequence of the instability of the zeolite in the presence of air. Laumontite exposed to low humidity air at low temperature partially dehydrates and decomposes (Blum, 1843; Armbruster & Kohler, 1992) and is easily eroded. We conclude from our data and observations that Ca-zeolites are the prime alteration products together with chlorite that form from the fundamental reaction (1):

granite + meteoric water = laumontite (Ca-zeolite) + chlorite + deep groundwater (1)

ZEOLITES IN BASEMENT ROCKS 58

It is important to recognize that granite alteration by water-rock interaction along fissures at temperatures of 150 ± 100˚C in a cooling orogen does not produce clay minerals with the exception of chlorite.

2.6.1. Reactions and processes of zeolite formation

The geologic overall context of the reported zeolites suggests that they precipitated from aqueous solutions circulating in open fissures. The zeolite crystals do not grow directly on the expense of primary minerals of unaltered granite. The structures and textures described imply that primary minerals of basement rocks dissolved along fractures into hot water (aqueous fluid). The hot aqueous liquid reached high degrees of super saturation with respect to a succession of Ca-zeolite minerals upon cooling, as observed and documented from many basement fluids at temperatures comparable to the imposed main temperature of zeolite formation of the Central Alps (Urach geothermal site, German continental deep drilling site KTB; Stober & Bucher, 2004, 2005). The process is a classic dissolution-precipitation process observed in many metamorphic environments (Carmichael, 1969).

The observed Ca-zeolites require the presence of Ca, Al and Si in the hot fluid to form. Source minerals of the rock matrix for necessary constituents are plagioclase, clinozoisite, quartz and earlier formed fissure calcite. The dissolution of plagioclase with an appreciable component of anorthite such as oligoclase in the samples from Arvigo (Fig. 2.12) provides a source for Ca2+.

2+ - plagioclase + H2O ⇒ albite + Ca + 2 AlO2 + 2 SiO2,aq + H2O (2)

2+ - Na4CaAl6Si14O40 + H2O ⇒ Na4Al4Si12O32 + Ca + 2 AlO2 + 2 SiO2,aq + H2O

Reaction (2) describes the dissolution of An-component of oligoclase, release of its component into the hot fluid and the production of a residual solid phase namely albite. Continuous albitization is accompanied by a porosity increase (Fig. 2.12), which provides the necessary permeability for water infiltration. The albitization reaction proceeds, as long fluid pathways are available to transport fluid undersaturated with plagioclase to the reaction interface of plagioclase and albite. In ZEOLITES IN BASEMENT ROCKS 59 the Arvigo samples plagioclase parent grains in the leaching zone are usually completely replaced by albite. Preserved plagioclase exists in the unaltered host rock, which could be an effect of the lack of porosity due to higher temperature transformation reaction (e.g. biotite-chlorite), or because fluid pathways get clogged by consumption of earlier formed porosity. The components released by plagioclase dissolution are continuously precipitated as zeolite in the open fissure by the model reaction (3):

2+ - Ca + 2 AlO2 + 4 SiO2,aq + 4 H2O ⇒ laumontite (3)

2+ - Ca + 2 AlO2 + 4 SiO2,aq + 4 H2O ⇒ CaAl2Si4O12 •4 H2O

Fig. 2.12: EMPA images showing products of albitization process (Pl + Qtz + H2O = Ab + Lmt). (a) BSE image showing pores in albite formed by albitization of primary plagioclase. Laumontite is easily recognized by its cleavage. A porosity of ~ 14 % was measured by image analyzing methods using IMAGEJ. (b) Ca distribution image showing irregular grain boundary of laumontite.

The additional silica necessary for the formation of zeolite may either derive locally from dissolution of primary quartz or from externally derived SiO2,aq and added by the fluid. Reactions (4) and (5) represent generic reactions describing plagioclase dissolution and albite and zeolite precipitation:

plagioclase + 4 H2O + 2 SiO2,aq ⇒ laumontite + albite (4)

Na4CaAl6Si14O40 + 4 H2O + 2 SiO2,aq ⇒ CaAl2Si4O12 •4 H2O + Na4Al4Si12O32 ZEOLITES IN BASEMENT ROCKS 60

plagioclase + 7 H2O + 5 SiO2,aq ⇒ stilbite + albite (5)

Na4CaAl6Si14O40 + 7 H2O + 5 SiO2,aq ⇒ CaAl2Si7O18 • 7H2O + Na4Al4Si12O32

Keep in mind that the reactions involve a transport step between dissolution and precipitation as the later two processes proceed at spatially different locations. The two reactions represent net balances of reaction (2) and (3) and are thus independent of pH.

During albitization of oligoclase in Arvigo Ca2+ and Al3+ were transported in solution from the rock matrix to the open fissure resulting in a porosity increase of 14 % in the matrix rock. This agrees well with the measured porosity in thin section (Fig. 2.12).

Zeolite formation in the Aar Massif requires another source for Ca2+ because plagioclase in the Aar granite and gneiss is pure albite. Alpine regional metamorphism reached only greenschist facies metamorphism in the Aar Massif (Fig. 2.11), resulting in a complete transformation of prealpine plagioclase to albite and a separate Ca-phase such as clinozoisite (epidote) or calcite. In the southern Gotthard massif plagioclase contains up to 18% anorthite component (Steck, 1976) as a result of increasing peak metamorphic grade toward the south (Fig. 2.11).

Thus Ca-zeolite may precipitate in fissures of the Aar massif from dissolved components that have been provided by the dissolution of clinozoisite and albite. Clinozoisite (epidote) is present in the host rock as result of Alpine greenschist facies metamorphism.

2 clinozoisite + 15 SiO2,aq + CO2 + 20 H2O ⇒ 3 stilbite + calcite (6)

2 Ca2Al3Si3O12(OH) + 15 SiO2,aq + CO2 + 20 H2O ⇒ 3 CaAl2Si7O18 •7 H2O + CaCO3

This plausible reaction mechanism co-precipitates zeolite and calcite. The assemblage is common in fissures (Table 2.2; TW03, 30992, 36728). If the fluid does not reach calcite saturation, the zeolite is the single product phase of the reaction and models the common zeolite (Lmt, Stb) fissures. ZEOLITES IN BASEMENT ROCKS 61

In some fissures, the dissolution of secondary calcite, originally formed from plagioclase breakdown during Alpine greenschist facies metamorphism, may have provided Ca2+ for zeolite formation during fluid-rock interaction:

+ 2- 2 albite + calcite + SiO2,aq + 7 H2O ⇒ stilbite + 2 Na + CO3 (7)

+ 2- 2 NaAlSi3O8 + CaCO3 + SiO2,aq + 7 H2O ⇒ CaAl2Si7O18 •7 H2O + 2 Na + CO3

Reaction (7) consumes calcite and forms stilbite that is accompanied by an increase in pH and the total of dissolved solids (TDS). The proposed reaction is supported by the + 2- high Na and CO3 concentrations, high pH, and high degrees of oversaturation with respect to stilbite in deep groundwater reported from the Gotthard rail base tunnel (Seelig et al., 2007).

2.6.2. Assemblage stability and phase relationships involving zeolites

Assemblage stability and phase relationships involving zeolites have been computed with the computer program Domino/Theriak (de Capitani & Brown, 1987) using the thermodynamic data by Bermann (1988), Evans (1990), Frey et al. (1991) and Maeder & Bermann (1991). For scolecite and chabazite we adopted the thermodynamic data of Johnson et al. (1983) and Ogorodova et al. (2002) and the heat capacity function was predicted from the equation [4Sco + 2Qtz = 3Lmt + An] and [7 Cha + 16 Qtz = 6 Stb + An], respectively.

The P-T model for the system CaAl2Si2O8–SiO2–H2O (Fig. 2.13) agrees well with field observations (Table 2.3). The predicted sequence of zeolites along a cooling path Lmt - Heu -Stb corresponds to the observed sequence in Alpine fissures. On Fig. 2.13 scolecite is restricted to low-pressures and would not form along the deduced cooling path in clear contrast to the field evidence. We conclude that the thermodynamic data for Sco needs improvement.

Fissure zeolites form at hydrostatic pressures below about 100 MPa corresponding to 10 km depth and temperatures below 400˚C. The effect of pressure on zeolite formation is thus assumed to be negligible. The absence of pumpellyite in Alpine fissures is consistent with low-pressure zeolite formation. In basaltic systems ZEOLITES IN BASEMENT ROCKS 62 with excess Qtz and Chl the laumontite-pumpellyite association is not stable at pressures below 100 MPa (Cho et al., 1986).

Fig. 2.13: Assemblage stability diagram in the Ca-Al-Si-O-H system. A retrograde cooling path adapted from Mullis et al. (1994) shows cooling during exhumation. Bulk: CaAl2Si2O8 + 5 SiO2 + 50 H2O; An + Qtz + water.

Zeolite forming reactions require dissolved silica in the hot water on the reactant side.

Consequently the reactions depend on the activity of SiO2 in the hot fluid (Fig. 2.14). The computed phase fields show the same sequence as observed in nature where the sequence also corresponds to a chronological order that can be related to a cooling path. The phase topology has been computed at 100 MPa and it shows a common boundary between Lmt and Stb. At lower pressure (10 MPa on Fig. 2.14) an inverted topology with a common boundary between Heu and Cha is predicted to be more stable than Lmt - Stb. The 100 MPa topology is consistent with the presented data, since laumontite-stilbite is a frequently observed zeolite assemblage (Table 2.2) and heulandite-chabazite has not been found. ZEOLITES IN BASEMENT ROCKS 63

During cooling and uplift (Fig. 2.11) zeolites start to form in fissures with the first appearance of scolecite and laumontite at a temperature of 280-300°C. Assuming a plausible crustal temperature gradient of 30˚C km-1 this would correspond to 10 km depth and a hydrostatic pressure of about 100 MPa.

Fig. 2.14: Equilibrium T- aSiO2 diagram at P = 10 MPa and 100 MPa, for the Ca-Al-Si-O-H system.

Standard state for aSiO2 is the pure stable SiO2 solid at P and T of interest. Thus quartz saturation is given at aSiO2 = 1. Note the topology inversion [Heu, Cha]⇔[Lmt, Stb] between 10 and 100 MPa pressure. Also note that pure chabazite-Ca is predicted to form from fluids with SIQtz < 0.

Temperature information can also be derived from chlorite geothermometry. As described above, chlorite forms earlier or with the zeolites and therefore provides maximum temperatures for zeolite formation. A temperature of 325°C ± 23°C (n = 170) has been calculated using the geothermometric calibration of Chatelineau (1988). The empirical calibration is based on the Al content on the tetrahedral sites. The temperature of 325˚C is interpreted as the maximum temperature for the zeolite window (Fig. 2.11). It probably marks the point when the crust has cooled along the path to start the first zeolites to form. ZEOLITES IN BASEMENT ROCKS 64

The low temperature limit of zeolite formation is not easy to establish. Stilbite is predicted to be thermodynamically stable in the presence of water at ambient P-T conditions. This is consistent with the observation that recent tunnel waters (40°C) from the Gotthard NEAT tunnel is strongly oversaturated with respect to stilbite, suggesting that stilbite formation is still under progress (Seelig et al., 2007).

Fig. 2.15: Equilibrium T-XCO2 diagram at P = 10 MPa, 50 MPa and 100 MPa, for the Ca-Al-Si-O-H-C system with excess quartz. (Bulk: CaAl2Si2O8 + 5 SiO2 + 50 H2O). Zeolites tolerate decreasing amounts of CO2 along cooling path.

2.6.3. Fluid composition

Zeolites in Alpine fissures are very irregularly distributed. For example zeolites in fissures of the central Aar granite (ZAgr) are frequent in some and rare in other regions. However, the bulk rock composition of ZAgr varies very little over the outcrop area of nearly 100 km (Labhart, 1977). Consequently, the different frequency of zeolite occurrence could be related to variations in externally derived fluids on the fractures. The consequence is that in such a case the rocks would be unable to buffer the fluid composition. This is in contrast to modern groundwaters in the basement, ZEOLITES IN BASEMENT ROCKS 65

which are controlled by the local lithology. It is obvious that H2O and CO2 will have a prime control on zeolite reactions (Zen, 1961). At relatively high pCO2, Ca-zeolites are replaced by clay minerals (Fig. 2.15).

It is evident that zeolites require low aCO2 conditions (Fig. 2.15). With increasing

CO2 in the fluid zeolites are replaced by kaolinite at low temperature and by margarite and calcite at high temperature. Stilbite does not tolerate much CO2; its presence indicates low CO2 fluids. CO2 tolerance increases with temperature in the order Stb <

Heu < Lmt < Sco. Zeolite CO2 tolerance decreases with increasing pressure and at pressure condition at about 100 MPa, zeolites decompose (or will not form) in fluids witch exceed 5 mole% CO2.

It can be concluded from the relationships that the CO2 content of the fluid controls the presence or absence of zeolites in a particular local area. The composition of fluid inclusions may thus give important information of fluid compositions during zeolite growth. Unfortunately, no fluid inclusions are preserved in zeolite minerals. Abundant fluid inclusions are present in Alpine fissure quartzes (Mullis, 1995). However, as documented above fissure quartz is clearly older than zeolites. Fluids trapped in earlier quartz are thus unrelated to zeolite forming fluids.

However, the fluids trapped in Alpine fissure quartzes show a very distinct N-S compositional trend (Frey et al., 1980). The Central Alps can be divided into four zones (from North to South: the higher hydrocarbon zone, 100-200 °C; the methane zone, 210-270 °C; the H2O zone, 240-430 °C; and the CO2 zone, 300-450 °C), whereas the Aar Massif and the northern part of the Gotthard Massif characterized by fluids dominated by H2O (>80 mole% H2O, <10 mole% CO2). Zeolites reported and discussed here formed in the CO2-poor H2O-dominated zone. The zeolite-rich zone of the eastern part of the Aar Massif and the northern part of the Gotthard Massif indicate H2O dominated fluids (Mullis et al., 1994). In the Penninic Alps to the South fluids tend to be dominated by CO2, explaining the absence of zeolite species in this region. The exceptional Arvigo zeolite locality in the Calanca valley within the CO2- rich region can be explained by locally H2O-rich fluids. Detailed studies have shown that the fluid composition changes from early CO2 dominated fluids (40 mole% CO2) to late stage CO2-absent fluids (Wagner et al., 2000 a, b). Mullis et al. (1994) observed that in general, every distinct fissure followed the same general retrograde ZEOLITES IN BASEMENT ROCKS 66

fluid evolution path, leading to a final water-rich fluid. Early H2O-poor fissure fluids have been unable to form zeolites and are thus responsible for the lack of the high- temperature Ca-zeolite wairakite.

Fig. 2.16: Activity diagram for the system Ca-Al-Si-O-H depicting mineral stability as a function of 2 2+ Log aSiO2,aq versus Log aCa /a H Standard state for aqueous species is a hypothetical one molal solution at infinite dilution. Quartz, apophyllite and wollastonite saturation and saturation of calcite as a function of pCO2marked with dashed lines. Deep waters from crystalline basement are shaded in gray: a = water samples from the Gotthard NEAT tunnel (Seelig et al., 2007), b = water samples from a granite located in Stripa/Sweden (Nordstrom et al., 1989), c = range of typical deep water from crystalline basement (Bucher & Stober, 2000), d = deep waters from the Urach drill site (Stober & Bucher, 2004).

Still, fluid inclusions in quartz crystals represent fluid compositions prior to zeolite formation. Regarding present-day fluid compositions in crystalline basement rocks of the continental crust, fluids are generally oversaturated in respect to zeolites. Deep continental fluids have the potential to form zeolites (Stober & Bucher, 2004, 2005). High-pH waters from the NEAT tunnel in the basement of the Aar Massif (Seelig et al., 2007), water from the basement at Stripa, Sweden, (Nordstrom et al., 1989), Bad Urach (Stober & Bucher, 2004) in the Black Forest basement (Bucher & Stober,

2000) are all oversaturated in respect to zeolites (Fig. 2.16). aSiO2,aq, pH and other fluid composition parameters may control the zeolite that will form in addition to temperature. Textural relationships presented above show that the fluids are not quartz-saturated during zeolite formation. Variations in aSiO2,aq control the zeolite assemblages along the cooling path. The following sequence of zeolites is present with decreasing aSiO2,aq: Lmt - Heu – Stb, Lmt – Stb and Lmt – Cha – Stb (Fig. 2.14). Thus the regional absence of zeolite and the observed different local sequences of ZEOLITES IN BASEMENT ROCKS 67

zeolites formed during cooling and uplift can be related to local variations in CO2 and

SiO2 content of the fissure fluid.

2.7. CONCLUSIONS

Zeolites are common minerals in late Alpine fissures of crystalline basement rocks of the Central Alps. The fissure zeolites precipitated from hot aqueous fluids that developed their load of dissolved solids from fluid-rock interaction. The zeolites form dense mats and crusts on fissure walls in surface outcrops and in rail and road tunnel up to 2000 meter below the surface. Three regions are particularly rich in zeolite- bearing fissures in granites and gneisses: (1) in the central and eastern part of the Aar- and Gotthard Massif, including the Gotthard road tunnel and the Gotthard-NEAT tunnel, (2) Gibelsbach/Fiesch with a large supply of zeolites in a fissure breccia between Aar Massif and Permian sediments, and (3) in Penninic gneisses of the

Simano nappe at Arvigo (Val Calanca). Ca-zeolites precipitated from the low-CO2 aqueous fluid with decreasing temperature in the following sequence: scolecite, laumontite, heulandite, chabazite and stilbite. Within this sequence an increase of the Si/Al ratio of the zeolites can be observed. Heulandite and chabazite incorporate significant amounts of extra components with respect to the Ca-Al-Si-O-H system. Stilbite samples from tunnel sections and from surface outcrops show distinctly different compositions that indicate a post-growth K-Ca exchange.

The components needed for zeolite formation were derived from the reaction of the hot fluid in the fissures with “primary” rock. The reaction dissolved plagioclase in rocks of the amphibolite facies (e.g. Arvigo) and precipitated albite and zeolite. In greenschist facies granites and gneisses of the Aar Massif, the reaction dissolved clinozoisite-epidote and/or calcite and precipitated zeolites that form the observed zeolite veins of this region. The dissolution process is accompanied with a porosity increase in the leaching zone. The remarkable and astonishing lack of zeolites in late fissures in many other regions with the potential for zeolite fissures is difficult to explain but could be related to pCO2 above critical threshold value that makes zeolite formation impossible.

ZEOLITES IN BASEMENT ROCKS 68

2.8. ACKNOWLEDGMENTS

We are grateful to Peter Amacher who provided high-quality mineral specimens from the Gotthard NEAT tunnel. Beda Hoffmann and Peter Vollenweider from the Swiss Natural History Museum in Bern, giving us the possibility to study their mineral collection. Special thanks go to the technicians and staff of the Mineralogical- Geochemical Institute, University of Freiburg: Isolde Schmidt for her help during sample preparation and for the XRD analyses; Melanie Katt for the careful preparation of fragile thin sections; Hiltrud Müller-Sigmund for her useful advise during EMP analyses and her patience with us at the electron microprobe. A special thanks deserved to the Friedrich Rinne foundation for the financial support

2.9. REFERENCES

Abrecht, J., 1994. Geological units of the Aar massif and their pre-Alpine rock associations: a critical review. Schweizerische Mineralogische und Petrographische Mitteilungen, 74, 5-27. Armbruster, T., 2000. Cation distribution in Mg, Mn-bearing babingtonite from Arvigo, Val Calanca, Grisons, Switzerland. Schweizerische Mineralogische und Petrographische Mitteilungen, 80, 279-284. Armbruster, T. & Gunter, M. E., 2001. Crystal structure of natural zeolites. In: Natural Zeolites: occurrence, properties; applications (eds Bish, D. L. & Ming, D. W.) Reviews in Mineralogy & Geochemistry, Vol. 45, pp. 1-68. Mineralogical Society of America, Washington, DC. Armbruster, T. & Kohler, T., 1992. Re- and dehydration of laumontite: a single- crystal X-ray study at 100 K. Neues Jahrbuch für Mineralogie, Monatsheft, 9, 385-397. Armbruster, T., Kohler, T., Meisel, T., Nägler, T. F., Götzinger, M. A. & Stalder, H. A., 1996. The zeolite, fluorite, quartz assemblage of the fissure at Gibelsbach, Fiesch (Valais, Switzerland): crystal chemistry, REE patterns, and genetic speculations. Schweizerische Mineralogische und Petrographische Mitteilungen, 76, 131-146. Armbruster, T., Stalder, H. A., Gnos, E., Hofmann, B. A. & Herwegh, M., 2000. Epitaxy of hedenbergite whiskers on babingtonite in Alpine fissures at Arvigo, Val Calanca, Grisons, Switzerland. Schweizerische Mineralogische und Petrographische Mitteilungen, 80, 285-290. Bambauer, H. U. & Bernotat, W. H., 1982. The microcline/sanidine transformation isograd in metamorphic regions. I. Composition and structural state of alkali feldspars from granitoid rocks of two N-S traverses across the Aar Massif and ZEOLITES IN BASEMENT ROCKS 69

Gotthard Massif. Schweizerische Mineralogische und Petrographische Mitteilungen, 62, 185-230. Berman, R. G., 1988. Internally-consistent thermodynamic data for minerals in the system Na2O-K2O-CaO-MgO-FeO-Fe2O3-Al2O3-SiO2-TiO2-H2O-CO2. Journal of Petrology, 29, 445-522. Bernotat, W. H. & Bambauer, H. U., 1982. The microcline/sanidine transformation isograd in metamorphic regions. II. The region of Lepontine metamorphism, Central Swiss Alps. Schweizerische Mineralogische und Petrographische Mitteilungen, 62, 231-244. Bish, D. L. & Ming, D. W. (eds), 2001. Natural Zeolites: occurrence, properties; applications. Reviews in Mineralogy & Geochemistry, Vol. 45, pp. 654. Mineralogical Society of America, Washington, DC. Blum, J. R., 1843. Leonhardit, ein neues Mineral. (Poggendorff's) Annalen der Physik und Chemie, 59, 336-339. Borchardt, R. & Emmermann, R., 1993. Vein minerals in KTB rocks. In: KTB Report 93-2 (eds Emmenmann, R., Lauterjung, J. & Umsonst, T.), pp. 481-488, Project Management of the Continental Deep Drilling of the Federal Republic of Germany in the Geological Survey of Lower Saxony. Borchardt, R., Zulauf, G., Emmermann, R., Hoefs, J. & Simon, K., 1990. Abfolge und Bildungsbedingungen von Sekundärmineralen in der KTB-Vorbohrung. In: KTB Report 90-4 (eds Emmermann, R. & Giese, P.), pp. 76-88, Projektleitung Kontinentales Tiefbohrprogramm der Bundesrepublik Deutschland im Niedersächsischen Landesamt für Bodenforschung. Breitschmid, A., 1982. Diagenese und schwache Metamorphose in den sedimentären Abfolgen der Zentralschweizer Alpen (Vierwaldstätter See, Urirotstock). Eclogae Geologicae Helvetiae, 75, 331-380. Bucher, K. & Frey, M., 2002. Petrogenesis of metamorphic rocks. 341 pp. Springer Verlag, Berlin. Bucher, K. & Stober, I., 2000. Hydrochemistry of water in the crystalline basement. In: Hydrogeology of Crystalline Rocks (eds Stober, I. & Bucher, K.), pp. 141- 175. Kluwer Academic Publishers, Dordrecht. Bucher, K. & Stober, I., 2001. Does plagioclase control the composition of groundwater in the crystalline basement? In: Water-Rock Interaction WRI-10, (ed Cidu, R.), pp. 149-152. A. A. Balkema Publishers, Lisse. Bucher, K., Zhang, L. & Stober, I., 2009. A hot spring in grnaite of the Western Tianshan, China. Applied Geochemistry, 24, 402-410. Carmichael, D. M., 1969. On the mechanism of prograde metamorphic reactions in quartz bearing pelitic rocks. Contributions to Mineralogy and Petrology, 20, 244-267. Cathelineau, M., 1988. Cation site occupancy in chlorites and illites as a function of temperature. Clay Minerals, 23, 471-485. Cho, M., Liou, J. G. & Maruyama, S., 1986. Transition from the zeolite to prehnite- pumpellyite facies in the Karmutsen Metabasites, Vancouver Island, British Columbia. Journal of Petrology, 27, 467-494. ZEOLITES IN BASEMENT ROCKS 70

Cho, M., Maruyama, S. & Liou, J. G., 1987. An experimental investigation of heulandite-laumontite equilibrium at 1000 to 2000 bar Pfluid. Contributions to Mineralogy and Petrology, 97, 43-50. Ciesielczuk, J. & Janeczek, J., 2004. Hydrothermal alteration of the Strzelin granite SW Poland. Neues Jahrbuch für Mineralogie, Abhandlungen, 179, 239-264. Coombs, D. S., Alberti, A., Artioli, A., Armbruster, T., Colella, C., Galli, E., Grice, J. D., Liebau, F., Mandarino, J. A., Minato, H., Nickel, E. H., Passaglia, E., Peacor, D. R., Quartieri, S., Rinaldi, R., Ross, M., Sheppard, R. A., Tillmanns, E. & Vezzalini, G., 1998. Recommended nomenclature for zeolite minerals: report of the subcommittee on zeolites of the international mineralogical association, commission on new minerals and mineral names. Mineralogical Magazine, 62, 533-571. de Capitani, C. & Brown, T. H., 1987. The computation of chemical equilibrium in complex systems containing non-ideal solutions. Geochimica et Cosmochimica Acta, 51, 2639-2652. Deer, W. A., Howie, R. A., Wise, W. S. & Zussman, J., 2004. Framework silicates: silica minerals, feldspathoids and the zeolites. 982 pp. The Geological Society, London. Dempster, T. J., 1986. Isotope systematics in minerals: biotite rejuventation and exchange during Alpine metamorphism. Earth and Planetary Science Letters, 78, 355-367. Deutsch, A. & Steiger, R., 1985. Hornblende K-Ar ages and the climax of Tertiary metamorphism in the Lepontine Alps (South-Central Switzerland); an old problem reassessed. Earth and Planetary Science Letters, 72, 175-186. Engi, M., Todd, C. S. & Schmatz, D., 1995. Tertiary metamorphic conditions in the eastern Lepontine Alps. Schweizerische Mineralogische und Petrographische Mitteilungen, 75, 347-369. Engvik, A., Putnis, A., Fitz Gerald, J. D. & Austrheim, H., 2008. Albitisation of granitoid: the mechanism of plagioclase replacement by albite. Canadian Mineralogist, 46, 1401-1415. Evans, B. W., 1990. Phase relations of epidote-blueschists. Lithos, 25, 3-23. Fellenberg von, E., 1893. Beschreibung des westl. Teils des Aamassivs BE/VS. Beiträge zur Geologischen Karte der Schweiz, 21. Freiberger, R., Hecht, L., Cuney, M. & Morteani, G., 2001. Secondary Ca-Al silicates in plutonic rocks: implications for their cooling history. Contributions to Mineralogy and Petrology, 141, 415-429. Frey, M., Bucher, K., Frank, E. & Mullis, J., 1980. Alpine metamorphism along the Geotraverse Basel-Chiasso - a review. Eclogae Geologicae Helvetiae, 73, 527- 546. Frey, M.. de Capitani, C. & Liou, J.G., 1991. A new petrogenetic grid for low-grade metabasites. Journal of Metamorphic Geology, 9, 497-509. Frey, M. & Mählmann, R. F., 1999. Alpine metamorphism of the Central Alps. Schweizerische Mineralogische und Petrographische Mitteilungen, 79, 135- 154. ZEOLITES IN BASEMENT ROCKS 71

Fridriksson, T., Neuhoff, P. S., Arnórsson, S. & Bird, D. K., 2001. Geological constraints on the thermodynamic properties of the stilbite-stellerite solid solution in low-grade metabasalts. Geochimica et Cosmochimica Acta, 65, 3993-4008. Frisch, W., 1979. Tectonic progradation and plate tectonic evolution of the Alps. Tectonophysics, 60, 121-139. Fujimoto, K., Tanaka, H., Higuchi, T., Tomida, N., Othani, T. & Ito, H., 2001. Alteration and mass transfer inferred from the Hirabayshi GSJ drill penetrating the Nojima Fault, Japan. Island Arc, 10, 401-410. Gianelli, G., Mekuria, N., Battaglia, S., Cheriscla, A., Garofalo, P., Ruggieri, G., Manganelli, M. & Gebregziabher, Z., 1998. Water-rock interaction and hydrothermal mineral equilibria in the Tendaho geothermal system. Journal of Volcanology and Geothermal Research, 86, 253-276. Gottardi, G., 1989. The genesis of zeolites. European Journal of Mineralogy, 1, 479- 487. Gottardi, G. & Galli, E., 1985. Natural Zeolites. 409 pp. Springer, Berlin, Heidelberg, New York, Tokyo. Heijboer, T. C., 2006. Origin and pathways of pro- and retrograde fluids, PTt paths and fluid-mineral equilibria from Alpine veins of the Central Alps: Case studies of the Fibbia and Amsteg areas. Unpublished PhD Thesis, University, Basel, Switzerland. Hirschi, H., 1925. Uranerz im tertiären Granit vom Bergell GR. Schweizerische Mineralogische und Petrographische Mitteilungen, 5, 429-430. Huber, H. M., 1943. Die Kluftminerallagerstätten im südöstlichen Gotthardmassiv. Schweizerische Mineralogische und Petrographische Mitteilungen, 23, 475- 537. Huber, W., 1946. Datolith vom Val Giuv, östl. Aarmassiv. Schweizerische Mineralogische und Petrographische Mitteilungen, 26, 79-84. Huber, W., 1948. Petrographisch-mineralogische Untersuchung im südöstl. Aarmassiv GR/UR. Schweizerische Mineralogische und Petrographische Mitteilungen, 28, 555-642. Johnson, G. K., Flotow, H. E. & O'Hare, P. A. G., 1983. Thermodynamic studies of zeolites: natrolite, mesolite and scolecite. American Mineralogist, 68, 1134- 1145. Jäger, E., Niggli, E. & Wenk, E., 1967. Rb-Sr-Alterbestimmungen an Glimmern der Zentralalpen. Beiträge zur Geologische Karte der Schweiz, 134. Kenngott, A., 1866. Die Mineralien der Schweiz nach ihren Eigenschaften und Fundorten ausführlich beschrieben. 460 pp.Verlag W. Engelmann, Leipzig. Keusen, H. R. & Bürki, H., 1969. Thomsonit und andere Zeolithe vom Geisspfad im Binntal VS. Schweizerische Mineralogische und Petrographische Mitteilungen, 49, 577-584. Koenigsberger, J., 1917. Über alpine Minerallagerstätten. Abhandlungen der Königlichen Bayrischen Akademie der Wissenschaften, 10, 1-108 ZEOLITES IN BASEMENT ROCKS 72

Labhart, T. P., 1977. Aarmassiv und Gotthardmassiv. Sammlung Geologischer Führer, 63. 173 pp. Borntraeger, Berlin. Langella, A., Cappelletti, P. & de'Gennaro, M., 2001. Zeolites in closed hydrologic systems. In: Natural Zeolites: occurrence, properties; applications (eds Bish, D. L. & Ming, D. W.) Reviews in Mineralogy & Geochemistry, Vol. 45, pp. 235-260, Mineralogical Society of America, Washington, DC. Maeder, U. K. & Berman, R. G., 1991. An equation of state for carbon dioxide to high pressure and temperature. American Mineralogist, 76, 1547-1559. McCaig, A. M., Wickham, S. M. & Taylor, H. P. J., 1990. Deep fluid circulation in alpine shear zones, Pyrenees, France: field and oxygen isotope studies. Contributions to Mineralogy and Petrology, 106, 41-60. Mercolli, I., Schenker, F. & Stalder, H. A., 1984. Geochemie der Veränderungen von Granit durch hydrothermale Lösungen. Schweizerische Mineralogische und Petrographische Mitteilungen, 64, 67-82. Michalski, I. & Soom, M., 1990. The Alpine thermo-tectonic evolution of the Aar and Gotthard massifs, Central Switzerland: Fission track ages on zircon and apatite and K-Ar mica ages. Schweizerische Mineralogische und Petrographische Mitteilungen, 70, 373-387. Mullis, J., 1995. Entstehung alpiner Kluftmineralien. Mitteilungen für Wissenschaft und Technik, 11, 54-64. Mullis, J., 1996. P-T-t path of quartz formation in extensional veins of the Central Alps. Schweizerische Mineralogische und Petrographische Mitteilungen, 76, 159-164. Mullis, J., Dubessy, J., Poty, B. & O'Neil, J., 1994. Fluid regimes during late stages of a continental collision: physical, chemical and stable isotope measurements of fluid inclusions in fissure quartz from a geotraverse through the Central Alps, Switzerland. Geochimica et Cosmochimica Acta, 58, 2239-2267. Nagel, T., de Capitani, C. & Frey, M., 2002. Isograds and P-T evolution in the eastern Lepontine Alps. Journal of Metamorphic Geology, 20, 309-324. Neuhoff, P. S., Fridriksson, T. & Arnórsson, S., 1999. Porosity evolution and mineral paragenesis during low-grade metamorphism of basaltic lavas at Teigarhorn, Eastern Iceland. American Journal of Science, 299, 467-501. Niggli, E., Koenigsberger, J. & Parker, R. L., 1940. Die Mineralien der Schweizeralpen. 661 pp. Wepf & Co, Basel. Nordstrom, D. K., Ball, J. W., Donahoe, R. J. & Whittemore, D., 1989. Groundwater chemistry and water-rock interactions at Stripa. Geochimica et Cosmochimica Acta, 53, 1727-1740. Ogorodova, L. P., Kiseleva, I. A., Melchakova, L. V. & Belitskiy, I. A., 2002. Thermodynamic properties of calcium and potassium chabazite. Geochemistry International, 40, 466-471. Orlandi, P. & Scortecci, P. B., 1985. Minerals of the Elba pegmatites. Mineralogical Record, 16, 353-364. Parker, R. L., 1922. Über einige schweizerische Zeolithparagenesen. Schweizerische Mineralogische und Petrographische Mitteilungen, 2, 290-298. ZEOLITES IN BASEMENT ROCKS 73

Passaglia, E., 1970. The crystal chemistry of chabazite. American Mineralogist, 55, 1278-1301. Passaglia, E. & Sheppard, R. A., 2001. Crystal chemistry of zeolites. In: Natural Zeolites: occurrence, properties; applications (eds Bish, D. L. & Ming, D. W.) Reviews in Mineralogy & Geochemistry, Vol. 45, pp. 69-116. Mineralogical Society of America, Washington, DC. Pfiffner, O. A., 1986. Evolution of the north Alpine foreland basin in the Central Alps. Special Publication International Association of Sedimentology, 8, 219- 228. Pouchou, G. & Pichior, F., 1991. Quantitative analysis of homogeneous or stratified microvolumes applying the model of "PAP". In: Electron probe quantitation (eds Heinrich, K. F. J. & Newbiry, D. E.), pp. 31-75, Plenum Press, New York. Quartieri, S. & Vezzalini, G., 1987. Crystal chemistry of stilbites: structure refinements of one normal and four chemically anomalous samples. Zeolites, 7, 163-170. Rahn, M., Mullis, J., Erdelbrock, K. & Frey, M., 1994. Very low-grade metamorphism of the Taveyanne greywacke, Glarus Alps. Switzerland. Journal of Metamorphic Geology, 12, 625-641. Ramsay, J. G., 1967. Folding and fracturing of rocks. 586 pp. McGraw-Hill, New York. Schaer, J. P., Reimer, G. M. & Wagner, G. A., 1975. Actual and ancient uplift rate in the Gotthard region, Swiss Alps: a comparison between precise levelling and fission-track apatite age. Tectonophysics, 29, 293-300. Schmid, S. M., Fügenschuh, B., Kissling, E. & Schuster, R., 2004. Tectonic map and overall architecture of the Alpine orogen. Eclogae Geologicae Helvetiae, 97, 93-117. Seelig, U., Stober, I. & Bucher, K., 2007. High-pH waters from the new, Gotthard rail base tunnel, Switzerland. In: Water-Rock Interaction WRI-12 (eds Bullen, T. & Wang, Y.), pp. 251-254. A. A. Balkema Publishers, Kunming. Sheppard, R. A. & Hay, R. L., 2001. Formation of zeolites in open hydrologic systems. In: Natural Zeolites: occurrence, properties; applications (eds Bish, D. L. & Ming, D. W.) Reviews in Mineralogy & Geochemistry, Vol. 45, pp. 235-260. Mineralogical Society of America, Washington, DC. Sigrist, F., 1947. Beiträge zur Kenntnis der Petrographie und der alpinen Zerrkluftlagerstätten das östlichen Aarmassives. Schweizerische Mineralogische und Petrographische Mitteilungen, 27, 39-182. Simonetti, A., 1971. Die Zeolithe und ihre Paragenesen in den Gesteinsklüften des Kt. Tessin, das Val Calanca und das Val Mesolcina (ital.). Bolletino Società ticinese di Scienze naturali, 62, 86-97. Stalder, H. A., 1964. Petrographische und mineralogische Untersuchung im Grimselgebiet. Unbublished PhD Thesis, University, Zürich, Switzerland. Stalder, H. A., Sicher, V. & Lussmann, L., 1980. Die Mineralien des Gotthardbahntunnels und des Gotthardstrassentunnels N2. 161 pp. Repof Verlag, Gurtnellen. ZEOLITES IN BASEMENT ROCKS 74

Stalder, H. A., Wagner, A., Graser, S. & Stuker, P., 1998. Mineralienlexikon der Schweiz. 579 pp. Wepf Verlag, Basel. Steck, A., 1976. Albite-Oligoklas-Mineralvergesellschaften der Peristeritlücke aus alpinmetamorphen Granitgneisen des Gotthardmassivs. Schweizerische Mineralogische und Petrographische Mitteilungen, 56, 269-292. Steck, A. & Burri, G., 1971. Chemismus und Paragenese von Granaten aus Granitgneisen der Günschiefer- und Amphibolitfazies der Zentralalpen. Schweizerische Mineralogische und Petrographische Mitteilungen, 51, 534- 538. Stober, I. & Bucher, K., 1999. Deep groundwater in the crystalline basement of the Black Forest region. Applied Geochemistry, 14, 237-254. Stober, I. & Bucher, K., 2004. Fluid sinks within the earth´s crust. Geofluids, 4, 143- 151. Stober, I. & Bucher, K., 2005. The upper continental crust, an aquifer and its fluid: Hydaulic and chemical data from 4 km depth in fractured crystalline basement rocks at the KTB test site. Geofluids, 5, 8-19. Todd, C. S. & Engi, M., 1997. Metamorphic field gradients in the Central Alps. Journal of Metamorphic Geology, 15, 513-530. Toroni, A., 1984. Chabasit-Fund im Tessin. Schweizer Strahler, 6, 447-457. Trümpy, R., 1960. Paleotectonic evolution of the Central and Western Alps. Bulletin of the Geological Society of America, 71, 843-908. Trümpy, R., 1980. Geology of Switzerland, a guide book, Part A: a outline of the geology of Switzerland. 104 pp. Schweizerische Geologische Kommission, Basel. Utada, M., 2001a. Zeolites in burial diagenesis and low-grade metamorphic rocks. In: Natural Zeolites: occurrence, properties; applications (eds Bish, D. L. & Ming, D. W.) Reviews in Mineralogy & Geochemistry, Vol. 45, pp. 277-304. Mineralogical Society of America, Washington, DC. Utada, M., 2001b. Zeolites in hydrothermal altered rocks. In: Natural Zeolites: occurrence, properties; applications (eds Bish, D. L. & Ming, D. W.) Reviews in Mineralogy & Geochemistry, Vol. 45, pp. 305-322. Mineralogical Society of America, Washington, DC. Vincent, M. W. & Ehlig, P. L., 1988. Laumontite mineralization in rocks exposed north of San Andreas Fault at Cajon Pass, southern California. Geophysical Research Letters, 15, 977-980. Wagner, A., Frey, M., Quadrio, F., Schwarzkopff, J. & Stalder, H. A., 1972. Die Mineralienfundstellen von Camperio und Campo Blenio, Kt. Tessin. Jahrbuch des Naturhistorischen Museum Bern, 4, 227-360. Wagner, A., Stalder, H. A., Stuker, P. & Offermann, E., 2000a. Arvigo - eine der bekanntesten Mineralfundstellen der Schweiz. Schweizer Strahler, 12, 41-70. Wagner, A., Stalder, H. A., Stuker, P. & Offermann, E., 2000b. Arvigo - eine der bekanntesten Mineralfundstellen der Schweiz. Schweizer Strahler, 12, 118-154. ZEOLITES IN BASEMENT ROCKS 75

Weibel, M., 1963. Chabasit vom Chrützlistock. Schweizerische Mineralogische und Petrographische Mitteilungen, 43, 361-366. Weisenberger, T. & Selbekk, R. S., 2008. Multi-stage zeolite facies mineralization in the Hvalfjördur area, Iceland. International Journal of Earth Sciences. DOI 10.1007/s00531-007-0296-6 Weiß, S. & Forster, O., 1997. Arvigo, Val Calanca: Kluftminerale aus dem Süden Graubündens. Lapis, 6, 13-42. Wälti, G., 1984. Mineralien aus dem Furka-Basis-Tunnel. Schweizer Strahler, 6, 559- 566. Zbinden, P., 1949. Geologisch-petrographische Untersuchung im Bereich südlicher Gneise des Aarmassivs (Oberwallis). Schweizerische Mineralogische und Petrographische Mitteilungen, 29, 221-356. Zen, E., 1961. The zeolite facies: an interpretation. American Journal of Science, 259, 401-409.

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 76

3. POROSITY EVOLUTION, MASS TRANSFER AND PETROLOGICAL EVOLUTION DURING LOW TEMPERATURE WATER-ROCK INTERACTION IN GNEISSES OF THE SIMANO NAPPE - ARVIGO, VAL CALANCA, GRISONS, SWITZERLAND

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 77

3.1. ABSTRACT

Low-grade mineral assemblages are the key to the appreciation of water-rock interaction in hydrothermal and geothermal systems located in granites and gneisses. Zeolite formation is an important process in rocks of the continental crust. It takes place at temperatures below 250°C under hydrothermal conditions. A detailed study of the mineralogical, chemical and petrological evolution of crystalline basement rocks in Arvigo was performed to assess information about the evolution of fluid-rock interaction during uplift of the Alpine orogen. The Arvigo fissures contain the assemblage epidote, prehnite, chlorite and various species of zeolites.

Fluid rock interaction takes place along a retrograde exhumation path which is characterized with decreasing temperature by: (1) coexisting prehnite/epidote, that reveals temperature conditions of 330 – 380 °C, (2) chlorite formation at temperature of 333 ± 32 °C and (3) formation of zeolites <250 °C. The formation of secondary minerals is related to the hydrothermal replacement reaction during albitization and chloritization that releases components for the formation of Ca-Al silicates and form a distinct reaction front. The fluid-rock interaction is associated with a depletion of

Al2O3, SiO2, CaO, Fe2O3 and K2O in the altered wall rock. The reaction is associated with an increase in porosity up to 14.2 ± 2.2 %, caused by the volume decrease during albitization and the removal of chlorite. The propagation of the sharp reaction front through the gneiss matrix occurred via a dissolution-reprecipitation mechanism. Zeolite formation is tied to the plagioclase alteration reaction in the rock matrix, which releases components for zeolite formation to a CO2-poor, alkaline aqueous fluid.

Keywords: water-rock interaction, laumontite, prehnite, epidote, albitization, Arvigo, Swiss Alps

3.2. INTRODUCTION

Fracture related fluid-rock interaction in gneisses and granitic rocks under hydrothermal conditions often causes changes in mineralogy and geochemistry and has been investigated by numerous authors (e.g. Ferry 1979; Mercolli et al. 1984; LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 78

Parneix and Petite 1991; Ciesielczuk and Janeczek 2004). Ca-Al-silicates, like epidote, prehnite and zeolites thereby formed in the open space and therefore gives information about temperature and pressure conditions as well as information about fluid composition during fluid-rock interaction and precipitation of secondary phases (e.g. Thompson 1971; Surdam 1973; Liou 1979; Liou 1985; Cho et al. 1986; Bevins et al. 1991; Young et al. 1991; Rose et al. 1992; Diegel and Ghent 1994; Gianelli et al. 1998; Faryad and Dianiska 2003). Detailed mineralogical and petrological study of the low-grade mineral assemblage can give an appreciation of the fluid-rock interaction and the porosity and permeability evolution, which is an important factor of geothermal systems and fluid migration in the upper continental crust (e.g. Gianelli et al. 1998; Neuhoff et al. 1999; Weisenberger and Selbekk 2008).

Ca-Al-silicate formation is a widespread and frequent feature in volcanic rocks of basaltic to acidic composition and in sedimentary environments, where elements necessary for the formation of secondary minerals mainly derive from dissolution of primary glass (e.g. Walker 1960, 1963; Hay 1966, 1977; Hay and Sheppard 1977; Surdam 1977; Gottardi 1989; Neuhoff et al. 1999; Hay and Sheppard 2001). However, infrequently similar Ca-Al-silicate phases have been reported in gneisses and granite, whereas elements most likely released during feldspar alteration (e.g. Freiberger et al. 2001; Faryad and Dianiska 2003; Weisenberger and Bucher 2009). A detailed study of the zeolite distribution in the Central Swiss Alps (Weisenberger and Bucher 2009) has shown that zeolites hosted in gneisses and granites occur frequently and has to be kept in mind during fluid-rock interaction in the upper continental crust.

The purpose of the paper is to report the results of an investigation of the hydrothermal alteration and metasomatic albitization of granitic gneisses of the Simano nappe. Samples were collected in an active quarry in Arvigo (Val Calanca, Switzerland; Fig. 3.1). Epidote, prehnite and zeolites display the main fissure minerals. In contrast only three other localities in the Lepontine Alps are known where zeolites occur (Biasca, Val Baveno, Val Vergeletto road tunnel; Stalder et al. 1998). Hydrothermal alteration in gneisses and granitic rocks is macroscopically seen as localized metasomatic bleached zone adjacent to fractures, which are filled by hydrothermal minerals. These expose the interface between generally unaltered and altered or albitized rock, respectively. This allows us to a detailed study of petrological changes, mass transfer and porosity evolution during metasomatism of LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 79 the upper continental crust, which contains a significant amount of plagioclase. The results provide an example of how observations of petrological changes, mass transfer and porosity evolution can be integrated in the geochemical interpretation of fluids in crystalline rocks of the upper continental crust.

3.3. GEOLOGICAL SETTING

The Arvigo locality is situated in the N-S striking Alpine valley Val Calanca, Grisons (Fig. 3.1). The rocks exposed in Arvigo belong to the Simano nappe, which is allocated to the lower Penninic basement nappes of the Central Alps and represents paleogeographically the southern passive margin of the European plate (Wenk 1955).

The Simano nappe is a metamorphic complex including several metagranitic bodies of Caledonian and Variscan age (Jenny et al. 1923; Keller 1968; Köppel and Grünenfelder 1975; Schaltegger et al. 2002), whereas the upper parts of the Simano nappe mainly consist of pre-Mesozoic gneisses and micaschists, intercalated with numerous amphibolite and calcsilicate lenses (Schaltegger et al. 2002; Rütti et al. 2005). The large Simano basement nappe is located between the underlying Leventina nappe and the Adula nappe as hanging wall (Berger et al. 2005), which is separated by thin Mesozoic metasediments (Fig. 3.1).

Barrovian-type Alpine metamorphism gradually increases from north to south from lower amphibolite facies condition in the southern Gotthard Massif to upper amphibolite facies conditions southwards to the Insubric Line (Frey et al. 1980; Frey and Mählmann 1999). Temperature determination by Engi et al. (1995) suggest a temperature range in the Simano nappe between 550°C in the north and 700°C conterminous to the Insubric Line. Corresponding pressure estimates in the Simano nappe achieve maximum pressures of about 650-700 MPa in a region approximately 20 km north of the Insubric Line. Peak metamorphic conditions for the region around Arvigo yield a temperature range from 600-680°C and pressure conditions of 550 to 600 MPa (Engi et al. 1995; Todd and Engi 1997; Nagel et al. 2002). LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 80

Fig. 3.1: Geological sketch maps: (a) Simplified tectonic map of the Simano nappe complex (modified after Spicher 1980); rectangle marks the section in part b. (b) Simplified geological map of the studied area along the NS trending Val Calanca valley (modified after Berger et al. 2005). (c) Outline of Switzerland, rectangle marks the location of the tectonic sketch in section a.

The rocks are mined as building stones in a quarry south of the town Arvigo (Fig. 3.1). The Arvigo gneisses general strike in NNW-SSE orientation with a dip of 20° to 30° in NE direction (146/30° NE to 158/31° E) that is parallel to the inclination of the valley. The Arvigo quarry became famous for a large number of Alpine fissure minerals, which occur in extensional fractures and cavities of the granitic gneiss (Ruppe 1966; Simonetti 1971; Wagner 1968, 1980, 1981, 1983; Weiß and Forster 1997; Wagner et al. 2000a, b). Fissures and gashes formed by brittle deformation were generated during exhumation and uplift of the Alpine orogen (Mullis et al. LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 81

1994). Fissures are generally perpendicular to the schistosity plane, whereas fissures are often brecciated due to late deformation stage.

Fig. 3.2: Field relationships and schematic sketches of fracture related hydrothermal alteration. Mineral abbreviation used after Bucher and Frey (2002). (a) Fissure mineral assemblages: Qtz-Kfs-Ep-Sco. (b) Assemblage: Cal-Sco-Lmt, whereas calcite shows corrosion. Coin for scale.(c) Vein photograph from a vein hosted in biotite-rich gneisses (Arvigo/Val Calanca). Hammer for scale. The vein is characterized by a ∼1 cm leaching zone trending in vertical direction, which appears to be lighter, due to the removal of biotite. The open space of the fissure is filled with secondary phases, which consists mainly on chlorite. (d) Schematic sketch of an Arvigo fissure. Extension forces lead to the opening of fissure, which describe pathways for fluids that increase the permeability and drive on fluid flow through the rock fissure. Leaching during fluid-rock interaction change primary mineralogy and geochemistry of the host rock, which is marked by an alteration zone, which grow perpendicular to the fissure orientation. Alteration zone shows an increase in porosity and appears to be lighter due to removal of primary minerals. Fluid, which gets saturated in respect to secondary minerals precipitate this in the open space, which affected a decrease of permeability and can hence stop the fluid movement. (e) Schematic sketch of a zeolite bearing Alpine fissure, exhibit euhedral mineral assemblages. Zeolite species overgrow earlier formed minerals in the following order, as it observed in nature: Qtz-Ep-Prh-

Sco.

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 82

The Arvigo fissures contain more than 40 different minerals (Weiß and Forster 1997; Armbruster 2000; Armbruster et al. 2000, Wagner et al. 2000a, b) hosted in fissures, whereas the main mineral assemblage is characterized by epidote, prehnite and zeolites (Fig. 3.2). The Arvigo gneisses are penetrated by various small (<10 cm) aplitic dykes as well as pegmatitic dykes, which consists of coarse-grained quartz, biotite and sulfides, like pyrrhotite or molybdenite.

3.4. PREVIOUS WORK

Due to the ongoing active mining, the Arvigo quarry exhibits a large suite of minerals and the quarry became famous for mineral collecting. Previous work is limited to fissure mineral descriptions due to the extensive collection possibility (Ruppe 1966; Simonetti 1971; Graeser and Stalder 1976; Brughera 1984; Wagner, 1968, 1980, 1981, 1983; Weiß and Forster 1997; Wagner et al. 2000a, b). Due to the presence of epidote, which known to crystallize with low CO2 contents and the knowledge about

CO2 rich fluids in the Lepontine Alps (e.g. Poty et al. 1974; Mullis et al. 1994) fluid inclusion studies on quartz crystals were made (Wagner et al. 2000a, b; Stalder 2007) to get information about the evolution of the fluid inclusions.

Fluid inclusion analysis (Wagner et al. 2000a, b; Stalder 2007) shows a distinct change of the fluid composition with time. Fluid inclusions in the core of quartz crystals are rich in CO2 (up to 40 vol. % CO2, Stalder 2007), whereas the fluid inclusions in the rim are CO2 free. Mineral inclusions in the core zone, which corresponds to the CO2 solution, are hornblende, ilmenite, biotite and carbonate (Wagner 2000a). Those minerals often found as precursor phases of Ca-Al-silicates of the mineral. The rim zone contains also mineral inclusions, which change from the inner part of the rim to the outer part, with the following sequence amianthus, epidote, chlorite and calcite, which occur over the whole rim zone (Wagner 2000a). Zeolites instead were not found as inclusions. Homogenization temperatures of fluid inclusions in the core are up to 365°C (Wagner et al. 2000a, b; Stalder 2007). In contrast fluid inclusions in the rim yield homogenization temperatures of 160 to LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 83

230°C. NaCl content increases from core to rim up to 5.9 wt. % NaCl (Wagner et al. 2000a, b; Stalder 2007).

3.5. SAMPLING AND ANALYTIC METHODS

A suite of representative samples of different veins and fissure was collected in the field in the years 2006, 2007 and 2008. A subset of the samples was selected for petrographic, bulk-rock and electron-microprobe studies.

Mineralogical analyses were carried out by point counting of more than 1600 evenly spaced points in each thin section using standard polarized microscope. Quantitative zeolite analysis were performed at the Institute of Geosciences (Mineralogy - Geochemistry), University of Freiburg, using a CAMECA SX 100 electron microprobe equipped with five WD spectrometers and one ED detector with an internal PAP-correction program (Pouchou and Pichior 1991).

Major and minor elements for zeolites were determined at 15 kV accelerating voltage and 10 nA beam current with a defocused electron beam of 20 µm in diameter with counting time up to 20 s. Na and K were counted first to minimize the Na and K loss during determination. Since the zeolite loses water when heated, the crystals were mounted in epoxy resin to minimize loss of water due to the electron bombardment. Natural and synthetic standards were used for calibration. The standards employed were: albite (Na), periclase (Mg), wollastonite (Si), barite (Ba), hematite (Fe), celestine (Sr), orthoclase (K), anorthite (Ca), rhodonite (Mn), fluorite (F) and rutile (Ti). The charge balance of zeolites formulas is a reliable measure for the quality of the analysis and which correlates with the difficulties related to the thermal instability of zeolites in microprobe analysis. A useful error test investigates the charge balance between the non-framework cations and the amount of tetrahedral Al (Passaglia 1970). Analyses are considered acceptable if the sum of the charge of the extra- framework cations (Ca2+, Sr2+, Na+, and K+) is within 10% of the framework charge (Al3+). LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 84

Identification of various minerals was obtained by a BRUKER AXS D8 Advance X-ray powder diffractometer (XRD) and the DIFFRACplus v5.0 software for evaluation. Whole rock analyses were performed by standard X-ray fluorescence (XRF) techniques at the Institute for Geosciences (Mineralogy - Geochemistry) at the University Freiburg/Germany, using a Philips PW 2404 spectrometer. Pressed powder and Li-borate fused glass discs were prepared to measure contents of trace and major elements, respectively. The raw data were processed with the standard XR-55 software of Philips. Relative standard deviations are < 1 % and < 4 % for major and trace elements, respectively. Loss on ignition was determined gravimetrically by calculated by heating at 1100 °C for 2 hours.

A slap, 7 mm in thickness, of sample Arvigo 12 was impregnated with fluorescent epoxy under high pressure conditions at the EMPA (Swiss Federal Institute for Materials Testing and Research) in Dübendorf, Switzerland to point out the porosity. Information of the percentage of porosity was conducted by digital analysis of photomicrographs. Images of the alteration profile Arvigo 12 were digitized using the software package ImageJ 1.38x (Wayne Rasband, National Institute of Health, USA) to calculate the area of porosity.

Assemblage stability and phase relationships involving zeolites have been computed with the computer program Domino/Theriak (de Capitani and Brown 1987) using the thermodynamic data by Bermann (1988), Evans (1990), Frey et al. (1991) and Maeder and Bermann (1991). For scolecite and chabazite we adopted the thermodynamic data of Johnson et al. (1983) and Ogorodova et al. (2002) and the heat capacity function was predicted from the equation [4Sco + 2Qtz = 3Lmt + An] and [7 Cha + 16 Qtz = 6 Stb + An], respectively.

3.6. RESULTS

3.6.1. Petrography

Fracture related hydrothermal alteration and metasomatism extends a few centimeters perpendicular out from discrete fractures (Fig. 3.2), Macroscopically the alteration zone is characterized by a light and porous leaching zone. The fracture or altered wall LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 85 rock respectively is associated with precipitation of hydrothermal mineral paragenesis.

Fig. 3.3: Sample Arvigo 12 that shows a complete section from unaltered host rock to the different alteration zone and fissure precipitation. (a) Hand specimen of the profile. Alteration grade increases in arrow direction. Schistosity is horizontal in respect to the sawed section. (b) Sawed slices showing the position of XRF analysis presented in Table 3.8. (c) Sketch of different alteration areas: unaltered gneiss (G) light altered zone (L), wherein biotite starts, the medium altered zone starts with the first occurrence of turbidity plagioclase, marked by the red dashed line. These zone graduals in the highly altered part (H), wherein chlorite is absent. Secondary phases are precipitated along the fissure wall (V). Additional a thin chlorite in vein in the light altered section is shown. (d) Thin section photographs along the profile. Red dashed line marks the sharp contact from the zone of albitization and the unaltered plagioclase. Albitization produce plagioclase turbidity due to the formation of porosity. Thin section numbers corresponds to Table 3.1.

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 86

Table 3.1: Modal mineral distribution through the alteration profile (see Fig. 3.3). mineral TS 12.1 TS 12.2 TS 12.3 plagioclase/albite 49.2 53.8 52.9 quartz 1.5 1.4 2.4 K-feldspar 17.3 13.5 14.0 biotite 0.0 8.7 10.3 muscovite 20.4 17.6 14.5 chlorite 10.4 3.6 4.5 others (apatite, calcite, epidote, titanite 1.3 1.5 1.5 ilmenite, laumontite, prehnite) counted points 2103 2179 1687

3.6.1.1. Unaltered rock

The unaltered gneisses (G, Fig. 3.2) in Arvigo are represented by a suite of biotite- gneisses and two-mica gneisses, showing distinct differences in their mineralogy. The dark to dark-colored rocks varying from anhedral/subhedral fine to medium grained biotite gneisses with an alternation of mafic and felsic layers, representing the schistosity in a scale of 2 to 4 mm to augen-gneisses with feldspar crystals up to 12 mm in size. The primary minerals are: plagioclase, quartz, K-feldspar, biotite, muscovite and ±hornblende with the accessories apatite, zircon, rutile, ilmenite and titanite. The modal composition varies in a wide range from biotite, biotite- hornblende, biotite-poor, and biotite-muscovite to muscovite dominated gneisses.

3.6.1.2.. Altered rock

The extent of fracture-related hydrothermal alteration ranges in difference stages, showing differences in altered mineralogy, fissure minerals and porosity. Generally the altered rocks are marked by an increase in lightness and porosity.

Slightly altered gneisses (L, Fig. 3.3, 3.4) differ from unaltered in the prevalence of biotite over chlorite. The alteration grade of biotite increases gradually in fissure direction (Fig. 3.3). In medium altered gneisses (M, Fig. 3.3), almost all biotite is replaced by chlorite. Microscopically observation shows the development of turbidity in plagioclase (Figs 3.3, 3.4). In the outer zone (H, Fig. 3.3) all chlorite is dissolved.

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 87

3.6.1.3. Fissure minerals

As mentioned above, more than 40 minerals are known, but most of them occur only infrequently. The Arvigo locality is characterized by the extensive occurrence of epidote, prehnite, chlorite and zeolites, beside the prevalent fissure minerals quartz, adularia and calcite (Fig. 3.2). Quartz, adularia represents earlier formed minerals in the fissure overgrown by Ca-Al silicates, chlorite and calcite, whereas chlorite appear in two temporal stages. The sequence of crystallization fissures can be determined by field and microscopy observations as followed: quartz, adularia, chlorite I, epidote, prehnite, chlorite II, calcite and zeolites. Scolecite and laumontite are by far the dominant zeolite species, whereas heulandite, chabazite, stilbite and epistilbite can only be found sporadically.

3.6.1.4. Changes in modal mineralogy

Changes in the modal mineral composition along the alteration profile (Fig. 3.3) are given in Table 3.1. Plagioclase is the dominant mineral in the gneisses and due to alteration the modal percentage decreases with alteration. Quartz decreases in the highly altered zone, suggesting a quartz consuming reaction during alteration. Considering the potassium bearing phases, an increase of the modal contents of K- feldspar and muscovite is accompanied by the biotite decrease up to the absence of biotite in the altered zone (M, Fig. 3.3). This decrease in biotite is again associated with the increase in chlorite in the rock matrix. Calcite, epidote, prehnite and laumontite occur as accessory minerals only in TS 12.1 (Fig. 3.3). Titanite and ilmenite are inversely correlated to the biotite decrease.

3.6.2. Mineralogy and mineral chemistry

To study the change of petrography and chemical composition of primary and secondary minerals to obtain changes in the mineralogy, mass change and porosity during fluid-rock interaction sample Arvigo 12 (Fig. 3.3) were taken. This sample was chosen, because the sample provides unaltered and altered zones in hand-specimen scale. LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 88

Fig. 3.4: Representative microphotographs of typical mineral assemblages in the fresh and altered rocks (Sample Arvigo 12, Fig. 3.3). (a) Unaltered plagioclase crystal showing albite twins. (b) Albitized plagioclase grains. Albite crystals appear to be turbidity due to the porosity increase during albitization. Red dashed line marks the grain boundary between two albite grains. The orientation of pores seems to be depended on the crystallographic orientation. (c) Alteration front between the medium altered and highly altered zone marked by the red dashed line. Plagioclase to the right is unaltered, whereas plagioclase left of the red dashed line are albitized. (d) Typical mineral assemblages not affect by hydrothermal alteration. (e) Biotite alteration to chlorite and K-feldspar in the light altered zone. K-feldspar is etched by H2F and stained by Na3Co(NO2)6 to make it distinguishable from plagioclase. (f) Complete alteration of biotite to chlorite, K-feldspar and ilmenite in the light altered zone.

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 89

Table 3.2: Plagioclase composition along the profile (see Fig. 3.3, 3.5, 3.6). x-direction* 783 1482 5402 5876 5702 11756 11766 12977 16722 y-direction* 8989 19263 20361 14628 5452 2314 11678 21050 19899

SiO2 68.72 68.78 69.25 68.86 68.18 68.91 67.98 69.03 68.81

Al2O3 19.50 19.51 19.79 19.69 19.61 19.71 19.43 19.69 19.87 BaO 0.00 0.04 0.00 0.02 0.03 0.00 0.00 0.00 0.00 SrO 0.00 0.00 0.00 0.00 0.00 0.03 0.00 0.04 0.09 CaO 0.14 0.23 0.12 0.24 0.12 0.30 0.12 0.13 0.27

Na2O 11.96 11.92 11.83 11.86 12.02 11.99 11.99 11.99 12.07

K2O 0.03 0.04 0.03 0.05 0.05 0.02 0.04 0.04 0.04 Total+ 100.48 100.60 101.03 100.82 100.04 101.02 99.65 100.98 101.18 Si 2.991 2.991 2.993 2.987 2.982 2.985 2.986 2.990 2.978 Al 1.000 1.000 1.008 1.007 1.011 1.006 1.005 1.005 1.013 Ba 0.000 0.001 0.000 0.000 0.001 0.000 0.000 0.000 0.000 Sr 0.000 0.000 0.000 0.000 0.000 0.001 0.000 0.001 0.002 Ca 0.007 0.010 0.005 0.011 0.006 0.014 0.006 0.006 0.012 Na 1.009 1.005 0.991 0.997 1.020 1.007 1.021 1.006 1.013 K 0.001 0.002 0.001 0.003 0.003 0.001 0.002 0.002 0.002 Sum 5.013 5.013 4.999 5.009 5.023 5.016 5.023 5.012 5.022 % An 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 % Ab 0.99 0.99 0.99 0.99 0.99 0.99 0.99 0.99 0.99 % Or 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

x-direction* 16780 22422 21787 23741 23543 25733 26417 27112 23778 y-direction* 15244 1917 18330 21711 15929 8818 3569 21789 21319

SiO2 67.93 68.71 68.40 66.67 68.07 68.62 67.34 64.53 66.34

Al2O3 19.78 19.68 19.57 21.28 19.02 19.37 19.53 22.59 20.85 BaO 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.01 0.00 SrO 0.00 0.04 0.15 0.00 0.00 0.14 0.00 0.16 0.11 CaO 0.47 0.51 0.18 2.20 0.22 0.21 0.53 3.64 1.76

Na2O 11.63 11.69 12.03 10.70 11.70 11.87 11.56 9.95 10.89

K2O 0.07 0.05 0.04 0.20 0.04 0.03 0.07 0.15 0.12 Total+ 99.93 100.74 100.46 101.07 99.22 100.24 99.15 101.09 100.14 Si 2.974 2.984 2.983 2.901 3.000 2.995 2.974 2.824 2.913 Al 1.021 1.007 1.006 1.091 0.988 0.996 1.016 1.165 1.079 Ba 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Sr 0.000 0.001 0.004 0.000 0.000 0.003 0.000 0.004 0.003 Ca 0.022 0.024 0.008 0.102 0.010 0.010 0.025 0.171 0.083 Na 0.987 0.984 1.017 0.903 1.000 1.004 0.990 0.844 0.927 K 0.004 0.003 0.002 0.011 0.002 0.001 0.004 0.008 0.007 Sum 5.010 5.005 5.023 5.009 5.007 5.010 5.014 5.019 5.014 % An 0.02 0.02 0.01 0.10 0.01 0.01 0.02 0.17 0.08 % Ab 0.97 0.97 0.99 0.89 0.99 0.99 0.97 0.83 0.91 % Or 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.01 0.01

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 90 continue Table 3.2 x-direction* 27500 29800 33149 32074 36294 36229 39015 46580 49981 y-direction* 16094 15749 3319 10519 18573 18572 3736 6781 19241

SiO2 65.72 65.62 64.28 64.35 4.72 63.43 63.38 63.36 64.12

Al2O3 21.63 21.64 22.65 22.20 3.96 22.56 22.93 22.75 22.54 BaO 0.05 0.02 0.00 0.05 0.01 0.06 0.05 0.00 0.01 SrO 0.00 0.17 0.12 0.07 0.00 0.00 0.00 0.06 0.04 CaO 2.31 2.64 3.48 3.39 0.62 3.85 3.98 3.83 3.29

Na2O 10.59 10.25 9.90 9.64 3.36 9.73 9.51 9.20 9.84

K2O 0.15 0.25 0.33 0.43 0.12 0.18 0.25 0.53 0.26 Total+ 100.47 100.66 100.83 100.19 12.82 99.81 100.20 99.78 100.14 Si 2.880 2.875 2.821 2.839 1.843 2.812 2.800 2.810 2.828 Al 1.117 1.117 1.171 1.154 1.821 1.179 1.194 1.189 1.172 Ba 0.001 0.000 0.000 0.001 0.001 0.001 0.001 0.000 0.000 Sr 0.000 0.004 0.003 0.002 0.000 0.000 0.000 0.001 0.001 Ca 0.108 0.124 0.164 0.160 0.259 0.183 0.188 0.182 0.155 Na 0.900 0.870 0.843 0.825 2.545 0.837 0.814 0.792 0.841 K 0.008 0.014 0.018 0.024 0.061 0.010 0.014 0.030 0.015 ∑ 5.015 5.008 5.023 5.008 6.542 5.021 5.016 5.006 5.014 % An 0.11 0.12 0.16 0.16 0.09 0.18 0.19 0.18 0.15 % Ab 0.89 0.86 0.82 0.82 0.89 0.81 0.80 0.79 0.83 % Or 0.01 0.01 0.02 0.02 0.02 0.01 0.01 0.03 0.01

x-direction* 53705 57320 64034 67184 69215 67096 67477 72561 76162 y-direction* 11861 3787 18791 9233 9202 7841 12718 15502 4900

SiO2 62.84 63.96 64.70 64.07 65.83 63.85 64.53 65.24 64.46

Al2O3 22.43 23.13 21.93 22.76 20.67 21.86 22.51 21.91 22.10 BaO 0.00 0.00 0.00 0.01 0.00 0.00 0.02 0.00 0.00 SrO 0.15 0.02 0.00 0.12 0.06 0.11 0.00 0.00 0.19 CaO 4.00 3.94 2.92 3.81 1.62 3.01 2.30 3.04 3.21

Na2O 9.53 9.44 10.11 9.84 10.78 10.20 10.46 10.03 9.88

K2O 0.40 0.26 0.31 0.26 0.24 0.21 0.63 0.29 0.29 Totat+ 99.39 100.78 100.02 100.95 99.23 99.25 100.53 100.57 100.20 Si 2.805 2.805 2.854 2.811 2.916 2.843 2.838 2.861 2.843 Al 1.180 1.196 1.140 1.177 1.079 1.147 1.166 1.132 1.149 Ba 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Sr 0.004 0.001 0.000 0.003 0.002 0.003 0.000 0.000 0.005 Ca 0.191 0.185 0.138 0.179 0.077 0.144 0.108 0.143 0.152 Na 0.824 0.802 0.865 0.837 0.926 0.880 0.891 0.853 0.845 K 0.023 0.014 0.018 0.015 0.013 0.012 0.036 0.016 0.016 Sum 5.028 5.005 5.017 5.026 5.014 5.030 5.042 5.007 5.013 % An 0.18 0.18 0.14 0.17 0.08 0.14 0.10 0.14 0.15 % Ab 0.79 0.80 0.85 0.81 0.91 0.85 0.86 0.84 0.83 % Or 0.02 0.01 0.02 0.01 0.01 0.01 0.03 0.02 0.02

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 91 continue Table 3.2 x-direction* 87090 93056 72561 105725 99345 y-direction* 4081 17791 15502 9528 17370

SiO2 64.87 63.99 63.51 64.33 63.96

Al2O3 22.10 22.32 22.91 22.22 22.58 BaO 0.00 0.01 0.02 0.02 0.00 SrO 0.09 0.15 0.11 0.10 0.08 CaO 3.08 3.40 3.57 3.13 3.74

Na2O 9.96 9.62 9.64 9.98 9.55

K2O 0.29 0.43 0.28 0.39 0.33 Totat+ 100.39 100.04 100.05 100.22 100.28 Si 2.851 2.830 2.808 2.838 2.821 Al 1.145 1.163 1.194 1.155 1.174 Ba 0.000 0.000 0.000 0.000 0.000 Sr 0.002 0.004 0.003 0.002 0.002 Ca 0.145 0.161 0.169 0.148 0.177 Na 0.848 0.825 0.826 0.854 0.816 K 0.016 0.024 0.016 0.022 0.019 Sum 5.009 5.013 5.016 5.022 5.010 % An 0.14 0.16 0.17 0.14 0.17 % Ab 0.84 0.82 0.82 0.83 0.81 % Or 0.02 0.02 0.02 0.02 0.02 + * in µm. totals include FeO, MgO, MnO, TiO2

Fig. 3.5: Plagioclase composition along the profile Arvigo 12 (Fig. 3.3). Vein mineral precipitation occurs on the left site of the diagram, whereas the host rock continuous to the right. The small spike at around 70000 µm corresponds to the chlorite vein (Fig. 3.3). Chemical analyses are given in Table 3.2.

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 92

3.6.2.1. Plagioclase and its alteration products

An important mineralogical difference between fresh and altered rock is the complete albitization of plagioclase (Table 3.2; Fig. 3.5, 3.6, 3.7). During albitization plagioclase of oligoclase composition (An15-19) has been replaced by albite (An0.5-2, Table 3.2; Fig. 3.5, 3.6). Figures 3.5 and 3.6 shows a one-dimensional and two- dimensional profile, respectively, of plagioclase composition through the alteration profile (Fig. 3.3). The profile shows a sharp decrease in anorthite component at about 25000 µm in distance to the fissure wall, which corresponds to the appearance of turbidity in plagioclase (Fig. 3.3, 3.4). The development of turbidity in albite grains in thin section (Fig. 3.4) is related to porosity increase (Fig. 3.7, 3.8).

Fig. 3.6: Modeled plagioclase composition along the profile Arvigo 12. X-direction is parallel to the schistosity and y-direction perpendicular to them.

The intra-granular pores have an angular to elongated shape, ranges from micrometer size up to 10 µm in length (Fig. 3.7, 3.8). The slight decrease of anorthite component at around 70000 µm from the fissure wall is geometrically related to the thin chlorite vein (Fig. 3.3). The albite component in plagioclase shows an inverse pattern of the anorthite pattern. Considering the orthoclase content of plagioclase along the profile a slight decrease in the altered zone is visible (Table 3.2; Fig. 3.3). Additionally a depletion of Sr in albite can be observed (Table 3.2). Other major changes of LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 93 plagioclase geochemistry are not detected. Saussuritization of plagioclase, a process that is characterized by the replacement of plagioclase by fine-grained sericite (e.g. Sandström et al. 2008) cannot be observed in our samples. Although the plagioclase is totally albitized, the original albite law twinning in plagioclase has been preserved in many of the altered grains (Fig. 3.4).

Fig. 3.7: EMPA images (TS 12.1) showing products of albitization process (Pl + Qtz + H2O = Ab + Lmt). (a) BSE-image showing the porosity (black dots) in albite during albitization of primary plagioclase. Laumontite is easy visible due to the perfect cleavage under an angel of ~90°. Using image analyzing methods a porosity of ~ 15% can be determined. (b) K distribution image. (c) Ca distribution image showing seriate - amoeboid grain boundary of laumontite. (d) Na distribution image. (e) Al distribution image. (f) Si distribution image.

Fig. 3.8: EMPA images (TS 12.1) showing a relict plagioclase grain, surrounding by albite. Considering the porosity enrichment around the plagioclase grain. (a) Ca distribution image. Plagioclase shows a zoning pattern, whereas the Ca content in the rims is higher, than in the core. (b) K distribution image showing seriate - amoeboid grain boundary of laumontite. (c) Na distribution image.

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 94

Table 3.3: Representative chlorite, biotite and muscovite composition and forming temperature for chlorite. sample A8.1 A8.1 A8.2 A10 A10 Arvigo Arvigo Arvigo Arvigo Arvigo Arvigo 4 8I 8I 8II 12.1 12.3 no. 8 9 1 21 24 1 13 19 36 14 14 mineral chlorite chlorite chlorite chlorite chlorite chlorite chlorite chlorite chlorite chlorite chlorite

SiO2 24.30 24.33 24.38 25.04 25.39 25.86 26.41 25.83 26.79 23.19 24.43

TiO2 0.00 0.00 0.09 0.02 0.02 0.01 0.00 0.04 0.07 0.02 0.09

Al2O3 21.98 21.89 21.48 20.23 20.28 19.52 18.36 19.17 17.82 20.73 20.60 BaO - - - - - 0.07 0.06 0.03 0.00 0.00 0.00 FeO 31.65 32.26 31.30 30.54 30.10 27.10 27.63 25.86 28.48 31.54 30.40 MnO 0.42 0.45 0.46 0.44 0.52 0.31 0.39 0.35 0.23 0.59 0.42 MgO 10.36 10.33 10.61 11.93 11.95 14.29 14.11 15.32 14.22 9.57 9.59 SrO 0.00 0.01 0.03 0.02 0.04 0.00 0.00 0.00 0.03 0.04 0.00 CaO 0.00 0.03 0.06 0.04 0.09 0.04 0.04 0.02 0.06 0.10 0.06

Na2O 0.00 0.02 0.01 0.00 0.00 0.03 0.01 0.00 0.03 0.06 0.00

K2O 0.01 0.00 0.01 0.00 0.01 0.01 0.01 0.01 0.03 0.01 0.00 Total 88.72 89.32 88.44 88.27 88.40 87.25 87.02 86.64 87.76 85.83 85.59 based on 20 oxygens Si 5.244 5.234 5.277 5.408 5.460 5.545 5.695 5.545 5.747 5.220 5.449 AlIV 2.756 2.766 2.723 2.592 2.540 2.455 2.305 2.455 2.253 2.780 2.551 AlVI 2.836 2.783 2.758 2.559 2.600 2.477 2.361 2.395 2.252 2.719 2.862 Ti 0.000 0.000 0.014 0.004 0.003 0.002 0.000 0.006 0.011 0.003 0.015 Fe2+ 5.713 5.803 5.666 5.517 5.413 4.859 4.983 4.642 5.109 5.937 5.670 Mg 3.333 3.311 3.422 3.842 3.830 4.568 4.536 4.904 4.547 3.211 3.187 Mn 0.077 0.082 0.084 0.081 0.094 0.056 0.071 0.064 0.042 0.112 0.080 Ba - - - - - 0.006 0.005 0.003 0.000 0.000 0.000 Sr 0.000 0.001 0.004 0.002 0.005 0.000 0.000 0.000 0.004 0.005 0.000 Ca 0.000 0.006 0.015 0.009 0.021 0.009 0.009 0.006 0.014 0.024 0.014 Na 0.000 0.010 0.005 0.000 0.000 0.012 0.004 0.000 0.012 0.026 0.001 K 0.002 0.000 0.003 0.001 0.003 0.003 0.003 0.002 0.008 0.003 0.001 T (°C)* 382 383 376 355 347 333 309 333 301 386 349 averagea: 333°C ± 32°C * chlorite temperature is calculated by the empirical calibration of Cathelineau (1988). a average of 39 analysis in 10 samples

sample Arvigo Arvigo Arvigo Arvigo Arvigo Arvigo Arvigo Arvigo Arvigo Arvigo 12.3 12.3 12.3 12.1 12.3 12.3 12.3 12.3 12.1 12.3 no. 5 6 13 10 7 5 6 13 10 7 musco- musco- musco- musco- mineral biotite biotite biotite vite vite biotite biotite biotite vite vite based on 22 oxygens

SiO2 35.31 34.91 34.63 49.47 47.70 Si 5.410 5.383 5.370 6.605 6.416 IV TiO2 3.06 2.98 2.43 0.52 0.83 Al 2.590 2.617 2.630 1.395 1.584 VI Al2O3 18.26 18.72 18.85 29.96 31.63 Al 0.707 0.786 0.815 3.319 3.431 BaO 0.11 0.15 0.10 0.22 0.22 Ti 0.353 0.345 0.283 0.052 0.084 FeO 23.10 22.41 23.36 2.70 2.41 Fe2+ 2.960 2.890 3.030 0.302 0.271 MnO 0.48 0.40 0.42 0.02 0.05 Mg 1.567 1.560 1.497 0.371 0.276 MgO 6.86 6.79 6.48 1.86 1.38 Mn 0.063 0.053 0.055 0.002 0.005 SrO 0.00 0.03 0.00 0.00 0.00 Ba 0.006 0.009 0.006 0.012 0.011 CaO 0.01 0.00 0.02 0.02 0.02 Sr 0.000 0.003 0.000 0.000 0.000

Na2O 0.12 0.12 0.13 0.25 0.34 Ca 0.002 0.000 0.003 0.003 0.003

K2O 9.32 9.19 9.22 10.42 10.02 Na 0.034 0.034 0.038 0.065 0.090 Total 96.63 95.70 95.63 95.45 94.59 K 1.821 1.807 1.823 1.775 1.719

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 95

3.6.2.2. Biotite-chlorite

Chloritization of biotite increases gradually from the fresh rock, where biotite is unaltered, into the light altered zone (Fig. 3.3, 3.4). In the medium and highly altered zone biotite is totally replaced by chlorite, which also occurs as unconsolidated and consolidated spherulitic aggregates in fissures. The zone of chloritizated biotite extends further away from the fracture than is indicated by the macroscopic leaching zone (Fig. 3.3, red-dashed line), wherein the albitization process is the dominated alteration feature.

The chloritization reaction of biotite is accompanied by the formation of K- feldspar and Ti-phases, like ilmenite and titanite (Fig. 3.4). Chloritization appears to take place preferentially along cleavage planes of biotite. Representative analysis for biotite and chlorite are given in Table 3.3. Biotite has a FeO content of 21.2 - 24.6 wt% and a MgO content of 6.3 - 7.0 wt%, reflecting values of XMg 0.21 - 0.24. Chlorite chemistry shows contents of MgO 9.4 - 16.6 wt% and FeO 23.2 - 32.3 wt% reflecting values of XMg 0.23 - 0.42 that is higher than in biotite. The most evident change occurs in the concentration of Ti, which is a major element in biotite with an average value of 3.12 wt% (Table 3.3) and in contrast incorporated only in traces in chlorite.

3.6.2.3. Muscovite

The muscovite content has been preserved during alteration. Muscovite composition varies in a limited range, but no systematic pattern over the whole range from unaltered to altered rock is obvious. The levels of MgO (0.9 - 2.1 wt%) and FeO (1.9 - 2.9 wt%) reflect a celadonite component of 7-10 mol%. The paragonite content is 4 - 7 mol% and significant higher than the margarite component (<1 mol%) (Table 3.3).

3.6.2.4. K-feldspar

The primary K-feldspar has generally been preserved during alteration. Additional K- feldspar is formed during chloritization of biotite. Chemical composition of rock forming K-feldspar ranges from Or85-94 and Ab15-06, respectively whereas the anorthite LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 96 component is less than 0.5 mol% (Table 3.4). In contrast secondarily formed adularia as euhedral crystals in fissures, succeeding fissure quartz show a higher Or component (Or93-99, Ab07-01, An00-01, Table 3.4).

Table 3.4: Representative chemical composition of K-feldspar and adularia. sample Arvigo Arvigo Arvigo Arvigo A8.1 A10 A4.1 12.1 12.1 12.2 12.3 no. 8 20 26 2 15 18 9 mineral k-feldspar k-feldspar k-feldspar k-feldspar adularia adularia adularia

SiO2 64.40 64.67 64.75 64.34 64.97 64.14 64.34

Al2O3 18.98 18.46 18.35 18.47 18.57 18.46 18.39 BaO 0.59 0.56 0.20 0.52 - - - SrO 0.00 0.06 0.07 0.00 0.00 0.00 0.00 CaO 0.05 0.02 0.04 0.05 0.01 0.00 0.01

Na2O 1.59 0.97 0.61 1.20 0.59 0.12 0.36

K2O 14.01 15.05 15.39 14.48 15.78 16.26 16.17 Total* 99.62 99.84 99.82 99.11 99.99 99.06 99.31 based on 8 oxygens Si 2.974 2.992 2.996 2.990 2.996 2.992 2.994 Al 1.033 1.006 1.000 1.011 1.009 1.015 1.009 Ba 0.011 0.010 0.004 0.009 0.000 0.000 0.000 Sr 0.000 0.002 0.002 0.000 0.000 0.000 0.000 Ca 0.002 0.001 0.002 0.002 0.001 0.000 0.000 Na 0.143 0.087 0.055 0.108 0.053 0.011 0.032 K 0.825 0.888 0.908 0.859 0.928 0.968 0.960 Or % 0.85 0.91 0.94 0.89 0.95 0.99 0.97 An % 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Ab % 0.15 0.09 0.06 0.11 0.05 0.01 0.03

* total include FeO, MgO, MnO, TiO2

3.6.2.5. Quartz

Quartz is a rock-forming mineral in the Arvigo gneisses. Considering the model mineral evolution during alteration, quartz is consumed during alteration (Table 3.1). However, fissure quartz as the first mineral precipitated in the fissure can give important information about fluid composition/evolution and mineral evolution during quartz growth by its fluid and solid inclusions (cf. Previous work).

3.6.2.6. Epidote

Secondarily formed epidote occurs as dominant phase in veins and fissures overgrowing quartz and adularia (Fig. 3.2). Epidote is overgrown by prehnite and zeolites. The textural relationship in sample Arvigo 1 (Table 3.5) suggests a co- genetic growth of epidote and prehnite. Epidote forms green to dark-green sheaf like LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 97 aggregates up to 20 mm in length, but common size ranges between 5 and 10 mm. 3+ The chemical composition of epidote (ideal composition: Ca2(Fe ,Al)3Si3O12(OH)) shows a pistacite component (Fe3+/(Fe3++Al)) that ranges from 15 to 30%. Except the main constitutes of epidote only MnO and SrO occurs in epidote up to 0.64 wt% and 0.30 wt%, respectively. No mineral zoning in epidote is visible.

Table 3.5: Analysis of coexisting epidote-prehnite (Sample Arvigo 1). sample Arvigo1 Arvigo1 Arvigo1 Arvigo1 Arvigo1 Arvigo1 Arvigo1 Arvigo1 no. 13 14 18 19 16 17 20 21 mineral epidote epidote epidote epidote prehnite prehnite prehnite prehnite

SiO2 37.63 37.45 37.35 37.77 42.52 42.44 42.22 42.54

TiO2 0.05 0.05 0.00 0.05 0.09 0.12 0.04 0.02

Al2O3 23.71 23.56 23.25 23.64 21.54 21.20 21.40 21.14

Fe2O3 12.72 12.24 12.64 10.84 3.81 3.98 4.05 3.66 FeO 0.10 - 0.38 - - - - - MnO 0.16 0.12 0.64 0.08 0.05 0.02 0.00 0.03 MgO 0.00 0.01 0.03 0.00 0.00 0.00 0.00 0.00 SrO 0.29 0.29 0.30 0.23 0.00 0.00 0.00 0.00 CaO 23.02 23.40 22.28 23.61 26.58 26.58 26.41 26.78

Na2O 0.02 0.02 0.00 0.00 0.04 0.01 0.00 0.02

K2O 0.00 0.00 0.00 0.01 0.01 0.00 0.00 0.02 Total* 97.70 97.19 96.90 96.24 94.74 94.35 94.12 94.21 based on 8 cations and 12.5 oxygens based on 7 cations and 11 oxygens Si 3.002 3.000 3.011 3.014 2.992 3.000 2.991 3.008 Al 2.229 2.224 2.209 2.223 1.786 1.766 1.787 1.762 Ti 0.003 0.003 0.000 0.003 0.005 0.006 0.002 0.001 Fe3+ 0.764 0.772 0.767 0.743 0.225 0.222 0.227 0.225 Mg 0.000 0.001 0.004 0.000 0.000 0.000 0.000 0.000 Fe2+ 0.007 - 0.025 - - - - - Mn 0.011 0.008 0.044 0.005 0.003 0.001 0.000 0.002 Sr 0.013 0.013 0.014 0.011 0.000 0.000 0.000 0.000 Ca 1.968 2.008 1.925 2.019 2.004 2.013 2.005 2.029 Na 0.003 0.003 0.000 0.000 0.005 0.001 0.000 0.003 K 0.000 0.000 0.000 0.001 0.001 0.000 0.000 0.002 * totals includes traces of BaO

3.6.2.7. Prehnite

Prehnite forms colorless to pale green fan-shaped radiating aggregates, so-called bow- tie structures (Phillips and Rickwood 1975), or sheaf like aggregates (Fig. 3.9). The aggregates are up to 10 cm in diameter overgrowing quartz and epidote. Prehnite used to be a substrate mineral for zeolites. As observed in epidote, prehnite composition is 3+ limited to their major elements (ideal composition: Ca2(Fe ,Al)2Si3O10(F,OH)2) and minor elements occur only in traces (Table 3.6). A small content of fluorine due to the OH ↔ F substitution is detectable (Table 3.6). Significant compositional variations within prehnite occur at the octahedral site in the Al2O3 - Fe2O3 ratio (Table 3.6; Fig. LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 98

3+ 3.9). The total iron content (as Fe2O3) varies from 0.0 to 9.5 wt. The content of Fe derby, controls the appearance of prehnite; colorless to pale green prehnite has low iron contents, whereas dark green prehnite is high in the iron content. Crystal aggregates often show a zoning pattern in Fe and Al (Table 3.6; Fig. 3.9), suggesting a Fe3+ ↔ Al substitution during growth, whereas the Fe content decreases to the rim or with time, respectively.

Table 3.6: Representative prehnite analysis. sample A 10 A 10 A 10 A 10 no. 2 3 4 5 rim core rim core

SiO2 43.75 42.95 43.78 43.21

Al2O3 23.33 19.12 23.92 21.20

Fe2O3 0.69 6.59 0.25 3.83 FeO 0.58 0.54 - 0.24 CaO 26.84 26.35 27.32 26.76

Na2O 0.01 0.01 0.02 0.00

K2O 0.01 0.02 0.01 0.00 F 0.05 0.11 0.05 0.01 Total 95.29 95.75 95.27 95.32 -O≡F 0.02 0.05 0.02 0.00 Total* 95.27 95.70 95.24 95.32 based on 7 cations and 11 oxygens Si 3.031 3.031 3.022 3.025 Al 1.905 1.590 1.946 1.749 Fe3+ 0.036 0.350 0.013 0.202 Fe2+ 0.034 0.032 - 0.014 Ca 1.992 1.992 2.021 2.007 Na 0.001 0.001 0.003 0.000 K 0.001 0.002 0.001 0.000 F 0.011 0.025 0.011 0.002 * totals includes traces of MnO, MgO, SrO BaO

Fig. 3.9: EMPA images showing element distribution in a prehnite aggregate indicating a Fe ↔ Al substitution during growth. (a) Fe distribution map showing an iron-enrichment in the core, which decrease to the rim. (b) Al distribution map showing an alumina-depletion in the core.

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 99

3.6.2.8. Zeolites

Scolecite and laumontite are by far the dominant zeolite species, whereas heulandite, chabazite, stilbite and epistilbite can only be found sporadically. Samples from Arvigo show the zeolite succession laumontite after scolecite (Fig. 3.2). However, one sample indicates an inverse growth pattern in which laumontite is the first zeolite that has formed. However from succession following chronology with increasing age can be compiled for Alpine fissures: scolecite, laumontite, heulandite, chabazite and stilbite (Weisenberger and Bucher 2009).

Laumontite (Ca4(Al8Si16O48) •18 H2O) often forms radiating aggregates. It forms thin, elongated fibers or prisms elongated along the c-axis with a square cross-section. Laumontite forms as {110} prism and commonly be twinned on {100} to form “swallow tail” or “V” twins. It is white with a length between <1 to 15 mm. The cleavage is perfect, with a slight pearly luster on the broad cleavage surface.

Considering the chemical composition calcium is the dominant extra-framework cation (average value of 97 %), with minor amounts of sodium and potassium (Fig. 3.10, Table 3.7), which consists less than 5 % of the extra-framework cations. Maximum value for K and Na are 7 and 1 %, respectively. Other elements occur only in traces. The extra-framework cation K increases during growth from core to the rim (Fig. 3.10). The Si/(Si+Al) content varies slightly between 0.67 and 0.69 (Fig 3.11). It can be observed that the content of alkalies increases with increasing Si and decreasing Ca and Al during growth, which can be expressed by the coupled substitution Si4+ + (Na+, K+) ↔ Al3+ + Ca2+.

Scolecite (Ca8(Al16Si24O80) •24 H2O) occur as white fibrous crystals with vitreous or slightly silky luster, forming characteristic radiating sprays (Fig. 3.2). The transparent to translucent crystals are slender prismatic with square cross section. Scolecite is between 1 and 20 mm in length, but the common length range is between 3 and 6 mm. 23 microprobe analyses from 3 different samples indicate no major chemical changes (Fig. 3.10, 3.11). Calcium is the dominant extra-framework cation (Fig 3.10, Table 3.7), which in average occupies 98 % of the extra-framework cation sites. Minor amounts of Na up to 4 %, but in average 2 % can be observed. In contrast K and Sr are not significantly incorporated in the framework structure of scolecite (Fig 3.10, 3.11) and therefore occur only in traces, like Ba, Mg and Fe (Table 3.7). LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 100

The small substitution range can also be recognized in the Si/(Si+Al) ratio, which shows a small range between 0.61 and 0.62 (Fig. 3.11).

Fig. 3.10: Triangular plots of extra-framework cation (Ca+Sr+Mg-Na-K) distribution of zeolites. Dashed framed area marks the chemical composition of zeolite found in granites and gneisses in the Swiss Alps (Weisenberger and Bucher 2009).

Heulandite ((Na,K)Ca4(Al9Si27O72) •24 H2O) forms crystals up to 12 mm in length, but the average size of the crystals is 1 to 4 mm. The monoclinic crystals occur in tabular habit parallel {010} and elongated in its typical coffin-shaped appearance. Crystals are transparent to translucent and colorless. Representative analyses for heulandite are given in Table 3.7. The average composition of heulandite from Arvigo is Ca3.27Na0.19K1.45Sr0.30(Al8.84Si27.16O72) •22 H2O, which is nearly identical to the heulandite composition (Ca3.37Na0.07K0.88Sr0.55(Al8.42Si27.49O72) •22 H2O) determined by Armbruster et al. (1996) from Gibelsbach in the western Aar Massif and from other Alpine fissures (Weisenberger and Bucher 2009). Heulandite can be classified by its geochemistry as heulandite-Ca (Coombs et al. 1998), with higher amounts of Sr and K (Fig. 3.10, Table 3.7). Calcium, with an average value of 63 % of all extra- framework cations, is the important extra-framework cation. Beside Ca, K marks a major element in heulandite, which yields an average value of 28 % (maximum to 31 %). Mentionable are the significant Sr content ranges between 5 and 7 %, which can be incorporated on the extra-framework cation sides and distinguish heulandite from the other zeolites found in Arvigo and in Alpine fissure, except chabazite, which comparably shows an enrichment of Sr (Weisenberger and Bucher 2009). The Si/(Si+Al) content ranges between 0.74 and 0.76 (Fig.11), which is in agreement with LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 101 the definition of heulandite (Coombs et al. 1998), that can be distinguished from clinoptilolite, 0.8 < Si/(Si+Al).

Fig. 3.11: R2+ - R+ - Si compositional diagram of zeolites. Si/Al ratio increases with chronologic order. Dashed framed area marks the chemical composition of zeolite found in granites and gneisses in the Swiss Alps (Weisenberger and Bucher 2009).

Table 3.7: Representative analysis of zeolite species. sample Arvigo12I Arvigo12I Arvigo2 Arvigo1 Arvigo 13 Arvigo 13 A8.1 A8.1 I no. 7 6 4 6 3 11 13 14 mineral laumontite laumontite laumontite scolecite scolecite scolecite heulandite heulandite

SiO2 52.50 51.43 51.52 45.64 45.64 45.82 57.84 59.46

Al2O3 21.88 19.97 21.24 24.65 24.72 25.33 16.44 15.79 CaO 11.42 10.90 11.86 14.14 13.55 13.97 6.41 6.29 SrO 0.11 0.00 0.04 0.00 0.00 0.04 1.40 1.14 BaO 0.01 0.00 0.02 0.00 0.00 0.01

Na2O 0.04 0.01 0.04 0.10 0.04 0.04 0.28 0.15

K2O 0.36 0.70 0.16 0.00 0.01 0.01 2.90 2.32

H2O 13.37 - - - - 13.69 - - Total 99.73 83.08 84.88 84.62 83.98 98.98 85.40 85.26 formula unit composition Si 16.124 16.415 16.112 24.312 24.413 24.187 26.921 27.449 Al 7.920 7.512 7.830 15.476 15.584 15.761 9.018 8.592 Ca 3.758 3.727 3.975 8.070 7.764 7.901 3.196 3.112 Sr 0.020 0.000 0.007 0.000 0.000 0.013 0.376 0.305 Ba 0.001 0.000 0.002 0.000 0.000 0.003 - - Na 0.000 0.006 0.024 0.103 0.039 0.039 0.256 0.134 K 0.141 0.285 0.064 0.000 0.004 0.005 1.723 1.366 O 48 48 48 80 80 80 72 72

H2O 13.696 - - - - 24.113 - -

TSi 0.669 0.686 0.673 0.611 0.610 0.605 0.749 0.762 E% 0.06 -3.14 -2.86 -4.73 0.00 -0.73 -1.58 3.09

* totals includes traces of FeO, MgO, MnOTiO2. TSi – Si/(Si+Al) E% - a measure of charge balance, = (100*((Al)-(Na+K)+2(Mg+Ca+Sr+Ba)/(Na+K)+2(Mg+Ca+Sr+Ba)) LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 102

In general all zeolite in Arvigo, as well as in fissure hosted in granites and gneisses in the Central Alps (Weisenberger and Bucher 2009), are Ca-dominated and the Si/Al ratio increases with decreasing temperature/time (Fig. 3.11).

3.6.3. Porosity

Porosity increase during fracture related alteration is a multiple process, by the volume increase changes during albitization process (Eq. 1, 2) and the removal of primary and secondary phases, like chlorite, which where formed during biotite dissolution. The porosity of gneiss varies between 1.0 and of 1.9 %. The porosity in the altered rock zone varies due to the removal of phases between 3.8 % at the alteration front of albitization and increases continuously up to 6.2 % in the medium altered rock (M, Fig. 3.3). The highly altered zone (H, Fig. 3.3), which is characterized by the removal of chlorite exhibit a porosity of up to 14.2 %. The porosity, which generated during albitization, is in submicroscopic scale and often not connected. This results in a non-considering of porosity during impregnation and digital analysis methods, which gives uncertainties of up to approximately 15 %. Considering the process of albitization a volume change of ~16 % (Eq. 1, 2) can be calculated in albite by using molar volume, which is in acceptable agreement with porosity determination using image analysis (Fig. 3.7) resulting in an value of 12.7 %.

2+ - 1 oligoclase + H2O => 4.19 albite + 0.81 Ca + 1.62 AlO2 + 1,62 SiO2,aq + H2O (1)

∆Vsolids = (4.19 VAb – 1 VOlg)/(1 VOlg) (2)

whereas following mineral composition were used: oligoclase = Na4.15Ca0.85Al5,85Si14.15O40 albite = Na0.99Ca0.01Al1.01Si2.99O8

Using the volume of plagioclase (Table 3.1) a porosity increase of 8.5 % can be related to the albitization process on whole rock scale. Figure 8 shows a relictic LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 103 plagioclase grain that did not completely replaced by albite and can give therefore an insight into the porosity evolution during replacement of a grain. Thereby the porosity in albite can be detected with 5 % whereas around the relictic plagioclase grain the porosity is enriched and increase up to 34 % (Fig. 3.8).

3.6.4. Whole rock geochemistry and mass changes

Changes of mass are common during hydrothermal alteration. A method for mass- balance analyses was described by Gresens (1967), which is based on the assumption that some elements are immobile and therefore conserved during the alteration. The ratio of mobile elements in the fresh and altered rock is then compared to the ratio of the immobile elements, in order to calculate the mass or volume change during alteration.

Table 3.8: Major and trace element composition through the alteration profile, including density (see Fig. 3.3). sample Arvigo 12 Arvigo 12 Arvigo 12 Arvigo 12 Arvigo 12 Arvigo 12 Arvigo 12 Arvigo 12 no. A I* A II A III A IV A V A VI A VII A VIII wt. % wt. % wt. % wt. % wt. % wt. % wt. % wt. %

SiO2 57.10 58.04 56.49 56.28 56.40 56.36 56.68 56.02

TiO2 0.66 0.62 0.57 0.57 0.57 0.55 0.54 0.58

Al2O3 21.65 21.29 22.67 22.94 22.94 22.42 22.90 22.96 tot Fe2O3 3.74 3.96 3.71 3.75 3.77 3.96 3.70 3.93 MnO 0.06 0.07 0.06 0.06 0.06 0.07 0.07 0.07 MgO 1.05 1.14 1.08 1.08 1.07 1.07 1.03 1.11 CaO 1.93 1.04 2.69 2.68 2.73 2.75 2.70 2.67

Na2O 5.37 6.22 5.68 5.75 5.84 5.93 5.91 5.82

K2O 4.71 4.43 4.13 4.25 4.05 3.83 3.98 4.26

P2O5 0.36 0.32 0.32 0.29 0.28 0.29 0.25 0.31 L.O.I. 2.44 1.66 1.45 1.20 1.27 1.34 1.31 1.06

Totals 99.21 98.95 99.01 99.00 99.12 98.72 99.21 98.94 ppm ppm ppm ppm ppm ppm ppm ppm V 59 53 51 50 52 52 53 57 Cr 31 25 26 24 25 27 25 27 Ni 18 15 15 19 13 15 15 17 Cu 10 < 5 < 5 8 < 5 1 5 15 Zn 54 55 54 55 53 54 52 58 Rb 144 126 122 141 132 124 129 152 Sr 105 118 291 298 304 313 307 302 Zr 323 323 280 265 275 266 264 276 Ba 801 773 701 665 641 610 633 638

(g/cm3) (g/cm3) (g/cm3) (g/cm3) (g/cm3) (g/cm3) (g/cm3) (g/cm3) density 2.884 2.797 2.849 2.848 2.850 2.825 2.832 2.874 *includes fissure minerals

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 104

Bulk rock chemistry and density measurements along the alteration profile (Fig. 3.3) are presented in Table 3.8. Using the simplified graphical method after Grant (1986) to solve Gresens’ (1967) equation, and the assumption that TiO2 is conservative during the hydrothermal alteration, Fig. 3.12 represents the isocon diagram. Anyway, titanite as Ti-bearing phase is rarely found as fissure minerals (Weiß and Forster 1997; Wagner et al. 2000a, b), suggesting a slight Ti mobility.

Fig. 3.12: Isocon diagram and histogram for the chemical loss and gain during alteration. Fresh rock is based on sample Arvigo 12 AVIII, whereas the altered rock composition is based on analyses Arvigo 12 AII (Fig. 3.3, Table 3.8). Isocon diagram showing constant mass (CM), constant volume (CV) and Isocon line. Elements below the lines are depleted in the altered rock relative to the fresh rock. Co = concentration of original element; Cf = concentration of transformed element. Histogram showing oxide mass changes compared to their respective mass in the fresh rock. Mfi = weight concentration of component i in transformed rock; Moi = weight concentration of component i in original rock; Mo = mass of the original rock; (Mfi-Moi)/Moi = mass change in relation to original element mass; (Mfi- Moi)/Mo = mass change in relation to original rock mass. Diagrams were constracted by using the program GEOISO (Coelho 2006)

Elements plotted above the isocon have been enriched relative to the fresh rock, whereas elements below the isocon line have been depleted during the alteration process. The slope of the obtained isocon is 1.068 (Fig. 3.12), equivalent to a mass loss of 6.8% (Grant 1986). Changes of the rock volume during alteration can be calculated using the mass ratio of immobile elements and the rock densities of the fresh and altered rock. Using the density and the mass ratio of immobile elements of 0.932 (inverted slope of the isocon), a volume loss of 3.9 % is obtained. Isocons representing constant mass and constant volume instead of constant TiO2 are included in Fig. 3.12 for comparison.

Considering gain and loss during alteration, CaO, Sr and Rb are the elements that shows the highest grade of depletion in respect to their element mass (Fig. 3.12) and LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 105

decreased by 64 %, 63 % and 25 %, respectively. K2O, SiO2, MgO, Fe2O3 MnO and

Al2O3 are depleted but in minor amounts, 3 %, 3 %, 4 %, 6 %, 7 % and 13 %, respectively. By defining TiO2 as immobile during the alteration process, Na2O behaves also conservative, without any changes (Fig. 3.12). Ba, Zr and P2O5 increase during alteration, 13 %, 10 % and 9 %, respectively. Nevertheless Al2O3, SiO2, CaO,

Fe2O3 and K2O are the significant elements, relative to the rock mass, whereas the other element changes are not significant, because they occur only in traces (Fig. 3.12). Given a mass loss of 6.8 % and the changes in major elements, the mass- balance equation for the hydrothermal alteration of the rock is (Eq. 3):

100 g fresh rock + fluid => 93.2 g of altered rock + 3.0 g Al2O3 + 1.7 g SiO2 + 1.7 g

CaO + 0.3 g Fe2O3 + 0.1 g K2O (3)

3.7. DISCUSSION

3.7.1. Mineral reactions

The most apparent mineralogical changes in the altered rock are the albitization (Fig. 3.3, 3.4, 3.7, 3.8) and the chloritization of biotite (Fig. 3.3, 3.4). Chloritization of biotite extents farther away from the fracture than is indicated by the brightening of the rock due to the albitization process (Fig. 3.3). This could be happens either because biotite is more easily altered than plagioclase, or due to fluids that are more easily transported along the connected sheet silicate clusters.

The proposed biotite chloritization reaction (Eq. 4, 5) is based on the assumption of conserved Al and Ti (Ferry 1979; Tulloch 1979; Parry and Downey 1982) and caused in observed mineral changes in samples from Arvigo, as well as on average biotite and chlorite composition. By reason that the Fe3+/Fetotal ration in biotite and chlorite is unknown, all Fe has been assumed to be Fe2+.

2+ 2+ + 1.8 biotite + 4.6 H2O + 1.3 Mg + 0.6 Ca + 0.4 H => 1.0 chlorite + 1.1 K-feldspar 2+ + + 0.6 titanite + 0.2 Fe + 0.2 SiO2 + 2.1 K (4) LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 106

2+ 2.05 biotite + 3.6 H2O + 0.9 Mg + 0.5 SiO2 => 1.0 chlorite + 2.0 K-feldspar + 0.8 ilmenite + 0.1 Fe2+ + 1.7 K+ + 0.6 H+ (5)

whereas following mineral composition were used: biotite = K1.8(Fe3.0Mg1.6)(Al0.8Ti0.4)(Si5.4Al2.6O20)(OH)4 chlorite = (Mg4.2Fe5.2Al2.6)(Si5.6Al2.4O20)(OH)16 adularia = KAlSi3O8 titanite = CaTiSiO5 ilmenite = FeTiO3

This solid-solid reaction is confirmed by petrographic observations (Fig. 3.4). However the reactions yields a volume increase of 6 % and 8 %, respectively, using the mineral molar volume of Eq. 4 and 5. Textural observation suggests that the chloritization process is volume conservative and therefore elements have to be transported away to achieve the fully pseudomorphic replacement (Fig. 3.4). Whether the chlorite or the K-feldspar component is dissolved cannot be achieved due to the fact that the stoichiometric coefficient of the products and reactants varies in the rock sample. Therefore the iso-volume reactions 6 and 7 display two endmember versions whereas chlorite and K-feldspar, respectively is dissolved to achieve iso-volume conditions.

1 biotite + 0.22 Mg2+ + 3.6 H+ => 0.43 chlorite + 0.98 K-feldspar + 0.39 ilmenite + 2+ + 3+ 0.35 Fe + 0.82 K + 0.04 SiO2 + 0.25 Al + 0.35 H2O (6)

2+ + 1 biotite + 0.11 H2O + 0.45 Mg + 3.58 H => 0.49 chlorite + 0.77 K-feldspar + 0.39 2+ + 3+ ilmenite + 0.07 Fe + 1.03 K + 0.36 SiO2 + 0.19 Al (7)

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 107

Considering the chloritization reaction of biotite (Eq. 6, 7), the reaction takes place in 2+ + 2+ 3+ the presence of fluid (H2O) and Mg and whereas SiO2, K , Fe and Al and are released.

The albitization process (Eq. 1) is a common equilibrium process under fluid saturated conditions over a wide PT range from diagenesis (Saigal et al. 1988, Lee et al. 2003) to greenschist (Leichmann et al. 2003) and even amphibolite facies metamorphism (Clark et al. 2005). During the albitization of plagioclase, dissolution 2+ 3+ of oligoclase occurs with coeval formation of albite. Ca , Al , SiO2, which were released during the process (Eq. 1), are transported in solution to the fracture or to open space in the adjacent rock (Fig. 3.7), where these elements precipitate as Ca-Al- silicates. Due to the limited solubility of Al3+, it seems likely that Al3+ does not migrate over significant distances and precipitates in proximate parts.

Oligoclase (An15-19) from Arvigo samples (Table 3.2, Fig. 3.5, 3.6) has been replaced by albite (An0.5-2). The product of the reaction that pseudomorphically replaced plagioclase crystals by albite produce porosity due to the differences in molar volume between the solid phases (Eq. 2). However, microporosity in plagioclase/albite increases with alteration as seen in Fig. 3.7 and 3.8, whereas the orientation of the pores are related to the crystallographic orientation (Fig. 3.4). Preserved albite twinning in albitized plagioclase can be seen in Fig. 3.4, which implies an epitactic overgrowth as previously described in altered plagioclase (Engvik et al. 2008) as well as in K-feldspar (Walker et al. 1995; Cole et al. 2004). This implies that the albitization process is controlled by dissolution-reprecipitation mechanism along a moving interface (e.g. Putnis and Putnis 2007, Engvik et al. 2008), resulting in porosity generation.

Primary muscovite and K-feldspar does not show any alteration texture without a signs of alteration and therefore they have not been regarded in the alteration scheme in Fig. 3.13.

3.7.2. Mass changes and element mobility

Mass changes and element mobility on whole rock scale is limited to few elements

(Fig. 3.12). Significant loss is marked by the major elements CaO, Al2O3, SiO2, LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 108

Fe2O3, MgO and K2O (Fig. 3.12). In contrast the loss of trace elements Rb and Sr and gain of Ba, Zr and P2O5 is relatively high, but their quantitative changes are evanescent in respect to the whole rock mass (Fig. 3.12). Mass changes can basically be linked to the two major alteration reactions of biotite chloritization and albitization of plagioclase (Eq. 1, 6, 7). Element mobility is summarized in Fig. 3.13 representing an alteration scheme of redistribution of elements during alteration between primary minerals, secondary minerals and hydrothermal fluid.

Fig. 3.13: Flow chart illustrating the exchange of ions and element mobility during hydrothermal alteration of gneisses from Arvigo. The diagram refers to alteration in rocks containing plagioclase altered to albite and chloritization of biotite. An external fluid is required to supply of H2O, CO2 and O2 for alteration. Primary Muscovite and K-feldspar regarded in the alteration scheme, due to the fact that they are not involved into the alteration. Polygons in the wall rock field represent primary minerals (Pl, Bt and Qtz) and their secondary alteration products (Ab, Kfs, Chl, Ttn and Ilm) that remain into the wall rock during alteration. Secondary minerals plotted into the fracture field, that are found as euhedral fissure minerals in the Arvigo fissures, but it is not excluded that these secondary minerals are not precipitated in open space in the wall rock. Black arrows represent migration of elements, which are based on alteration reaction and average mineral composition discussed in the text. Dashed black arrows represents element migration, that are limited onto framework conditions: chlorite and K- feldspar formation in the fissure will happen, if the volume increase during chloritization of biotite can not balanced by the precipitation in open spaces in the adjacent to the wall rock. SiO2 required for zeolite formation, may either derive from primary quartz, or is provided in solution in the hydrothermal fluid. Prehnite and epidote will incorporates iron, if Fe2+ is oxidized to Fe3+.

The decrease of Ca2+ is compatible with albitization of plagioclase (Eq. 1), during which Ca2+ is mobilized (Fig. 3.12, Table 3.8). The change in Ca2+ is strongly connected to Sr2+, due to similarity in size and charge that allows Sr2+ to substitute LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 109

Ca2+ in the plagioclase lattice (Sun et al. 1974). Therefore it is obvious that Sr2+ is also mobilized during plagioclase dissolution. Most of the Ca2+ content was transported out in solution of the wall rock (Fig. 3.12, Table 3.8) and precipitated in the fracture as Ca-Al-silicates (epidote, prehnite and zeolites) and calcite, depending on fluid composition (CO2-H2O) and temperature. Nevertheless around 1 wt% CaO is still stored in the altered rock. The remaining content of Ca2+, which is incorporated in albite only in traces (Table 3.2) is assumed to be hosted in titanite formed during chloritization (Eq. 6) and in secondary phases, which precipitated already in the open space of the wall rock (Fig. 3.7). Sr2+ is preferred to integrate into heulandite (Table 3.7).

The decrease of Al (Fig. 3.12) is noteably high in fact that Al is generally relatively immobile compared to other elements during fluid-rock interaction (Carmichael 1969; Ragnarsdottir and Walther 1985; Verdes et al. 1992). Al often has been assumed to be immobile and was therefore often used as a constant reference frame for mass balance calculations (e.g. Thompson 1975; Grant 1986). Nevertheless, field evidences (Fig. 3.3), mass changes (Fig. 3.12, Table 3.8) and mineral reaction (Eq. 1) suggest Al3+ mobility during fluid-rock alteration. Considering the albitization process, by which porosity is generated, Al3+ leached out to the fracture and precipitated as Ca-Al-silicate. The volume conservative chloritization reaction marks an additional source for the Al3+ deficiency and Al mobility. Therefore, Al3+ have to be transported away (Eq. 6, 7) to achieve the fully pseudomorphic replacement (Fig. 3.4), without volume expansion.

The decrease of SiO2 during alteration is linked with the leaching of silica during chloritization and albitization (Eq. 1, 6, 7).

Considering Eq. 6 and 7, Fe is released during chloritization and Mg has to be added to balance the chloritization reaction. Nevertheless mass balance calculation indicating a loss in both elements (Fig. 3.12). The Mg loss is related to the chlorite removal out of the wall rock into the fracture (Fig. 3.3, Table 3.1). Fe (Eq. 6, 7) migrates to the fracture, where it is oxidized and precipitates in epidote and prehnite (Table 3.5, 3.6).

The volume conservative chloritization (Eq. 6, 7) consume Mg2+. Mg could be added by an external fluid or due to migration of biotite from the adjacent rock. LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 110

However, previous research (Parry and Downey 1982; Parneix et al. 1985; Drake et al. 2008) has shown that the Mg content of chlorite decreases systematically as more biotite is replaced, which indicates that Mg is a very mobile element during biotite chloritization.

The decrease in K+ is consistent with the chloritization process (Eq. 6, 7) during which K+ leached out from the wall rock (Fig. 3.2). Primary K-feldspar and muscovite can be excluded as K source by textural evidences (Fig. 3.4).

The increase of volatiles (LOI, Table 3.8) is not astonishing, if we consider the hydration reaction of biotite and the formation of hydrous Ca-Al-silicates during albitization.

No changes are assessed in Na+ concentration (Fig. 3.12) during alteration. These agree with volume calculations during albitization and measured microporosity in altered plagioclase, regarding the chemical change in plagioclase.

Elements that were release during albitization and chloritization are transported out to the fissure and precipitates their in secondary minerals. In general solute transport in porous material is accomplished by three principal mechanism: advection, aqueous diffusion and hydrodynamic dispersion, whereas advection is the dominant mechanism (Steefel 2008). Therefore a temperature gradient between the fissure-fluid and the fluid that reacts with the minerals represents a reliable driving force that enhance advection of the solution into fissure direction.

3.7.3. Mineral stability and mineral equilibria

3.7.3.1. Prehnite and epidote

Prehnite and epidote are common minerals in low-grade metamorphic rocks (e.g. Fricke 1952; Kuniyoshi and Liou 1976; Tulloch 1979; Liou et al. 1983; Liou 1985; Cho et al. 1986; Rose and Bird 1987; Bevins et al. 1991; Freiberger et al. 2001).

Temperature and/or fO2 conditions, during which prehnite and epidote are formed, are reflected in the chemical composition. The occurrence of pumpellyite, which is often associated with prehnite and epidote, could not be confirmed and therefore can give an indication about pressure conditions (Kuniyoshi and Liou 1976). The quite LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 111

common inhomogeneity of the Al2O3-Fe2O3 ratio, which is common on thin section scale in prehnite, may be due to (1) partial re-equilibration during progressively changing P-T-fO2 conditions in process of Ca-Al-silicates formation, whereas the iron content in prehnite increases with decreasing temperature and increasing fO2, whereas the Al contents decrease (Kuniyoshi and Liou 1976; Liou et al. 1983), (2) successive discrete hydrothermal events (Freiberger et al. 2001), or (3) local chemical influence of host minerals. However elevated fO2 conditions necessary for the oxidizing process of Fe2+, released during biotite dissolution, seems like the cause for the Fe3+ enrichment in the core (Fig. 3.9).

Fig. 3.14: Predicted temperature of the formation of coexisting prehnite and epidote, using the Fe3+ - Al partitioning curves, determined by Rose and Bird (1987). Distribution of Fe3+ between coexisting epidote and prehnite is expressed as the mole fraction of Ca2FeAlSi3O10(OH)2 in prehnite and as pistacite component Ca2Fe3Si3O12(OH) in epidote. Dashed lines represents limits on the compositional range for coexisting prehnite and epidote (I: prehnite => zoisite + grossular + quartz; II: laumontite + prehnite => clinozoisite + quartz). Solid lines presents isotherms, based on constant log K by using thermodynamic properties of the reaction: Al-prehnite + epidote => Fe-prehnite + clinozoisite (adapted from Rose and Bird 1987).

Prehnite stability determined for metabasites reaches up to 400 °C and up to 300 MPa (Liou et al. 1985; Frey et al. 1991). However the absence of pumpellyite, which is stable between 100 and 800 MPa, suggests that pressure conditions during the LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 112 formation of Ca-Al-silicates were below 100 MPa (Kuniyoshi and Liou 1976; Liou et al. 1985; Frey et al. 1991).

Sample Arvigo 1 (Table 3.5) reveals the coexisting prehnite-epidote assemblage and therefore it can be used to get information about the formation temperature. Figure 3.14 shows the formation temperature of coexisting prehnite and epidote, using the approach from Rose and Bird (1987). Rose and Bird (1987) suggested that the iron partitioning of coexisting prehnite and epidote is a function of temperature.

Using the iron partitioning treatment after Rose and Bird (1987) a formation temperature between 330 and 380°C (Fig. 3.14) for coexisting prehnite and epidote can be diagnosed. Isotherms in Fig. 3.14 are based on thermodynamic calculations by iterative solutions for the composition of coexisting prehnite, epidote and grandite garnet (Rose and Bird 1987). The compositional limits on the stability of coexisting prehnite and epidote are represented by the two dashed lines (Fig. 3.14). Data from Arvigo (Fig. 3.14) requires that the coexisting prehnite and epidote pairs represent non-equilibrium Fe3+-Al partitioning and are metastable with respect to the reaction: Prh = Zo + Grs + Qtz. The evaluated temperature is consistent for prehnite stability in active geothermal systems (275-350°C; Bird et al. 1984) and in hydrothermal experiments (376°C, Liou et al. 1983).

Considering the determined temperature and formation temperature of chlorite, which are in the same range and therefore well agree with the textural appearances that suggests a contemporaneous growth.

3.7.3.2. Chlorite

Several chlorite thermometers, applying structural and chemical criteria are available from the literature (e.g. De Caritat et al. 1993). The empirical calibration based on AlIV content (Cathelineau and Nieva 1985; Cathelineau 1988) was tested for low- grade basic rocks within a regional metamorphic context (Bevins et al. 1991). However De Caritat et al. (1993) has shown that the content of AlIV is not dependent on the geochemical composition of the host rock and therefore the Cathelineau (1988) thermometer is frequently used including chlorite hosted in granites and gneisses (e.g. Rahn et al. 1994; Orvosová et al. 1998). Although the precision of such empirical LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 113 thermometers is difficult to assess, Vidal et al. (2001) showed that the variation of AlIV in chlorite with is temperature thermodynamically sound.

Chemical composition of chlorite in Arvigo was examined on 40 grains and the corresponding formation temperature varies from 27 to 380 °C in a wide range, with an average value of 333 °C (Table 3.3). Considering the temperature distribution (Fig. 3.15), two distinct groups of chlorite are evident. The first group shows a formation temperature around 310 °C and the higher temperature group varies from 330 to 380 °C. Chlorite, which pseudomorphic replaces biotite in the rock matrix, trends to higher formation temperatures, in contrast to spherulitic chlorite precipitated in the fissure and open space, respectively, which is related to lower temperatures.

Fig. 3.15: Distribution of calculated chlorite temperatures for Arvigo samples using the calibration after Cathelineau (1988).

3.7.3.3. Zeolites

Zeolites mark beside apophyllite, the youngest secondary mineral formed in the Alpine fissure in Arvigo. The general chronology of the Arvigo zeolite assemblages is: scolecite, laumontite, heulandite and stilbite and is comparable with recent evaluations of zeolite bearing fissures in the Alps and thermodynamic phase modeling in the system CaAl2Si2O8–SiO2–H2O (Weisenberger and Bucher 2009).

Reaction isograds for Ca-zeolites are well determined by experimental methods (e.g. Liou 1971; Thompson 1970; Cho et al. 1987; Frey et al. 1991). In general the maximum temperature and pressure limits of zeolite stability are in agreement with observations on geothermal systems (Kristmannsdóttir and Tómasson 1978; Frey et al. 1991). Nevertheless, there is a significant deviation between temperature noted at the position of a given zeolite isograd reaction and temperature resulted from phase LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 114 equilibrium calculations (Kristmannsdóttir and Tómasson 1978; Frey et al. 1991; Neuhoff et al. 2000), which arises difficulties in attempt to use experimental observations of phase equilibrium to assess thermobarometric conditions in zeolite- facies rocks that usually reflects higher temperatures as it observed in natural systems. According to this discrepancy various variables like pH, chemical composition of the water, pCO2, the presence of additional extra-framework cations like Sr, Na and K, the amount of H2O incorporated in the zeolite channel structure, order-disorder and fluid pressure, respectively, can affect the thermodynamic equilibrium conditions and consequently the reaction isograds (e.g. Thompson 1970; Liou 1971; Cho et al. 1987; Frey et al. 1991; Neuhoff et al. 1999).

Fig. 3.16: Temperature - fO2 phase-diagram (50 MPa) displaying the stable mineral assemblages for the bulk composition of host rock material of profile Arvigo 12. Thermodynamic calculation where done with iron-free prehnite.

The mineral evolution and the evolution of porosity were modeled by using computed assemblages stability diagrams with the Theriak/Domino software of de Capitani and

Brown (1987). Considering an increase in fO2 during alteration, which is implied by the iron zoning in prehnite (Fig. 3.9), Figure 3.16 represents a T-fO2 phase diagram. It LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 115

indicates that at constant temperature a change of fO2, which is externally controlled by the infiltrated fluid, can change the stability of zeolites.

The formation of zeolite occurs in a low-pressure regime (<200 MPa) (Bish and Ming 2001). The occurrences of scolecite and laumontite indicate maximum formation temperatures of 280 - 300 °C by using equilibrium phase modeling in the system CaO-Al2O3-SiO2-H2O-CO2 (Fig. 3.17). However, in situ temperature measurements in active hydrothermal systems, like in basaltic rocks on Iceland suggest lower temperatures (e.g. Kristmannsdóttir and Tómasson 1978).

An important factor, which controls the formation of zeolite, is the composition of the fluid from which the secondary minerals precipitated. Regarding the low frequency of zeolites in the Lepontine Alps, this lack can be related to CO2 dominated fluids (Poty et al. 1974; Mullis et al. 1994; Stalder 2007). Zen (1961) noticed that zeolite mineral assemblages could be obtained by the increase of the chemical potential of H2O relative to that of CO2, at constant temperature and pressure. At relatively low CO2 activities, calcium zeolites are destabilized relative to assemblages contain calcite, quartz and clay minerals (Zen 1961; Senderov 1973) that is supported by the fluid inclusion evolution to CO2 free fluids. However, the lack of zeolite inclusions in quartz, suggests that quartz growth was finished before zeolite formation starts and no information about fluid compositional at the time of zeolite formation is available. Nevertheless thermodynamic modeling points out that the fluid has to be low in CO2 (Fig. 3.17). The effect of CO2 bearing fluids on the stability of zeolites and other Ca-Al-silicates can be seen in the calculated thermodynamic phasediagrams for different Ca/Al ratios and pressure conditions in the system CaO-Al2O3-SiO2-

H2O-CO2 (Fig. 3.17). For calculation the ideal mixing model for H2O-CO2 was used.

Zeolite species are stable in fluids dominated by H2O with low CO2 concentrations.

With increasing CO2 activity zeolite species are replaced by kaolinite (e.g. Val Bedretto, Stader et al. 1998) at lower temperature and other Ca-Al silicates, calcite and quartz at higher temperature (Fig. 3.17). Stilbite is stable only at very restricted

XCO2 less than 0.04 at 10 MPa and low temperature, whereas heulandite, laumontite and scolecite are stable at higher XCO2 (Fig. 3.17). Scolecite stability is controlled by pressure conditions and occurs only at lower pressures. The Ca/Al ratio can also effect the stability variations of zeolites. However the occurrences of heulandite may also be controlled by additional extra-framework cations like Sr, Na and K, which LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 116 prefer to incorporate into heulandite in Alpine fissure (e.g. Weisenberger and Bucher 2009). Generally zeolite stability increases with decreasing pressure and at pressure condition at about 100 MPa, zeolites will destabilize at XCO2 lower than 0.05 (Fig. 3.17). However, in hydrothermal systems the fluid pressure is unlikely equal to the total pressure. The assumption of very low CO2 is supported by the absence of kaolinite, which would be present at lower temperatures with XCO2 > 0.04.

Fig. 3.17: Equilibrium T- XCO2 diagrams at P = 10 MPa, 50 MPa and 100 MPa, for the CaO-Al2O3- SiO2-H2O-CO2 system for different Ca/Al ratios.

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 117

3.7.4. Mineral evolution

Bulk rock composition of the unaltered rock of sample Arvigo 12 (Fig. 3.3, Table 3.8) has been recalculated to atomic proportions. To simplify the diagrams, Ti and Mn have been ignored. The presence of chlorite, epidote and prehnite indicate that Fe occurs in di- and trivalent states and that some provision for the redox state is need to be made. However the fact, of Fe zoning in prehnite, which is interpreted as change in oxygen fugacity, specification of the redox state is not possible. Thereby modeling with different redox state conditions was done with the result that the ratio 1/1 of Fe3+/Fe2+ reflects the best fit with observed secondary mineral inventory. The alteration is driven by hydrothermal process and therefore H2O was set in excess.

Figure 3.18 gives a PT diagram that is appropriate for hydrothermal alteration conditions. The corresponding predicted assemblage evolution is shown in Fig. 3.19. According to Fig. 3.18, epidote would start to form at temperature conditions of 450°C, which is higher than the temperature estimation by using iron distribution in epidote and prehnite. However, the CO2 rich fluid at higher temperature could be the reason for the delay and formation of epidote at lower temperature.

Fig. 3.18: Assemblage stability diagram for unaltered rock of sample Arvigo 12 (Fig. 3.3, Table 3.8). Note: note all assemblage fields are labeled.

LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 118

Chloritization occur at temperature between 350 and 325 °C, depending on pressure conditions and is in good agreement with the empirical calibrated chlorite thermometers by Cathelineau (1988) that yields an average chlorite formation temperature of 333 °C.

The phasediagram points out that K-feldspar is affected by a hydration reaction and forms muscovite. But in contrast to the computed diagram, the observed mineral inventory (Table 3.1), where K-feldspar occurs as rock-forming mineral, as well as fissure adularia, did not reflect the stable assemblage, and therefore K-feldspar behaves metastable in the hydrothermal Arvigo system. Considering the chloritization reaction (Eq. 6, 7), which releases Al2O3 during reaction, the phasediagram has to be regarded with care.

Fig. 3.19: Predicted assemblage evolution during hydrothermal alteration, calculated along a linear exhumation path diagonal through Figure 3.18.

Metasomatic reactions depend on the compositions of the fluid phases leaving and entering the system, and cannot be thermodynamically treated by merely considering the solid phases in the same way that isochemical metamorphic reactions are LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 119 traditionally used to determine PT conditions. Nevertheless the phase-diagram reflects a good approach to model the mineral evolution, because prehnite and the zeolite-in reaction reflect a plausible mineral evolution, if we consider field observations and compare them with other hydrothermal systems. In any case, which kind of zeolite is formed is also related to cation substitutions and might increase the stability in contrast to the pure Ca endmember, which lack in the thermodynamic data-base, like the incorporation of Sr into heulandite.

Fig. 3.20: Porosity evolution during hydrothermal alteration. Minerals that were precipitated during hydrothermal alteration were assumed to precipitated in the fissure, whereas elements necessary for formation were moved out from the host rock, producing porosity. One path reflects the calculation, that assumed the total removal of chlorite and one reflects the porosity path, whereas chlorite occurs in the rock matrix.

Using the predicted assemblage evolution along the cooling path (Fig. 3.18, 3.19) and assuming that the elements for the secondarily formed Ca-Al-silicates in the fissure were derived from the adjacent wall rock, the porosity evolution can be calculated by removing the molar volume portion of the secondary mineral, from the initial molar rock volume (Fig. 3.20). Figure 3.20 therefore shows the temporal porosity evolution along the PT path (Fig. 3.18). LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 120

One path is modeled by assuming that all chlorite remains in the wall rock, whereas the second path requires complete chlorite removal into the fissure space and resulted porosities of 9.2 and 11.3 vol. %. In contrast porosity calculation using modal mineral composition gives a porosity of ~17 %. The difference could be related to the 2-dimensional analysis of porosity of an anisotropic texture or to the meta-stability of K-feldspar in phase diagram calculation and therefore the non-consideration of K- feldspar into the porosity calculation.

3.7.5. Fluid accessibility and composition

The alteration process, forming hydrous Ca-Al-silicates requires a significant amount of H2O. Fluids have to infiltrate the wall rocks, where fractures act as fluid channels on outcrop scale (e.g. Austrheim 1987; Bons 2001) as well as on microscale (e.g. Fitz Gerald and Stünitz 1993; Oliver 1996).

Fracturing is caused by brittle deformation that is younger than the main Alpine deformation and related to the uplift of the Central Alps 10-20 Ma ago (Steck 1968; Purdy and Stalder 1973). Two distinct fracture directions can be observed in Arvigo, which differ in mineralogy and suggest a change in the stress field with time. The later formed fissures are mineralized with zeolites. Recent dating of the latest fissure minerals in the Central Alps (Weisenberger and Bucher 2008) suggests a younger age (∼2 Ma) of zeolite formation in the Central Alps.

Biotite as well as chlorite occur as connected cluster due to foliation and provide migration pathways for the fluid through the whole altered zone and increase into fissure direction. In contrast, the albitization process produces porosity that is constant over the whole sharp alteration front (Fig. 3.3). In general, the porosity of the albitization is the volume occupied by the fluid phase and is generated at the reaction interface where the volume of dissolved plagioclase is less than the volume of albite that reprecipitates. However, only the interconnected porosity provides permeability. Nevertheless, the porosity which is generated during volume loss of albitization is enriched at the grain boundaries (Fig. 3.8) and marks a prominent fluid channel, whereas intragranular porosity may not affect the fluid permeability due the poorly LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 121 connected pores. Both alteration reactions, chloritization of biotite and albitization, enhance the permeability.

If we include secondary minerals into the volume equation and neglecting that some elements are transported away or added from the fissure wall, a volume increase has to be assumed. This process of precipitation of secondary in the adjacent area of the wall rock can impregnate earlier formed porosity in the wall rocks, as well as the fracture. This decrease hereby the permeability to the point that the fluid low is interrupted and the alteration process is terminated due to the absent of fluid. This is seen at the chlorite vein in Fig. 3.3, where the adjacent plagioclase is depleted in Ca (Fig. 3.5, 3.6), but less depleted than the albite crystals in the appreciable alteration zone. These Ca remaining in plagioclase can be related to inaccessibility due to the impregnation of the fluid channel by secondary chlorite.

Early CO2 dominated fluids may derive from decarbonation processes or oxidizing of organic matter of Mesozoic metasedimentary rocks, which are integrated in the nappe-stack of the Lepontine Alps (e.g. Poty et al. 1974; Mullis et al. 1994). Magmatic waters seem to be an unlikely source regarding the geological setting and the absence of magmatic intrusions. If the CO2 was formed in lithological units below or above and migrated to the fissure remains as open question. The change in fluid composition could be caused due to infiltration of meteoric waters or by the lack of the sources for decarbonation, due to the erosion of the metasedimentary units above the Simano nappe. However, if we consider the element mobility (Fig. 3.13) the change in element concentration and mineralogy does not need infiltration of chemically exotic fluids.

3.8. CONCLUSION

Low-grade mineral assemblages are the key to the appreciation of water-rock interaction in hydrothermal and geothermal systems located in granites and gneisses. The Arvigo locality is a example for a crystalline basement unit consisting of granites and gneisses, which is significantly affected by fracture-related hydrothermal alteration. Fissures and gashes formed by semi-brittle deformation were generated LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 122 during exhumation and uplift of the Alpine orogen. These fractures and cavities were filled with fluids and new minerals crystallized in the open space.

The Arvigo fissures contain the assemblage epidote, prehnite, chlorite and various species of zeolites. In general epidote is overgrown by prehnite, chlorite and zeolites. Ca as extra-framework cation dominates all zeolites, whereas the specific zeolite formed in the fissures depends on the temperature. Following Ca-dominated zeolites precipitated from the low-CO2 aqueous fluid with decreasing temperature: scolecite, laumontite, heulandite, chabazite and stilbite.

The composition of coexisting prehnite/epidote reveals temperature conditions between 330 and 380 °C for the pre-zeolite assemblage using the Rose and Bird (1987) calibration. The iron zoning pattern in prehnite suggest elevated oxygen fugacity during early growth of prehnite. AlIV occupancy on the octahedral site in chlorite (Cathelineau 1988) suggests temperature conditions of 333 ± 32 °C. Zeolite formation takes place at temperatures below 250°C.

Fluid induced mineral reactions occurred during the hydrothermal alteration of rock-forming minerals in the wall rock. The reactions are marked by the albitization of plagioclase accompanied by chloritization of biotite, forming a reaction front propagating from central fractures into the gneiss matrix. A first replacement reaction changes biotite into chlorite within a 3 to 7 cm thick zone of the host rock. The plagioclase replacement reaction releases components for zeolite formation and forms a sharp reaction front in the gneiss at about 2 to 2.5 cm from the central fracture.

The albitization reaction is associated with a volume decrease for the solids. Thereby albite remains as daughter phase during in the wall rock and exhibit a porosity increase of ~16 %, whereas the anorthite component get dissolved. We conclude that much of the produced volume is transferred to the central extension fracture by laumontite precipitation in the open fracture. The porous product albite suggests that the propagation of the reaction front through the gneiss matrix occurred via a dissolution-precipitation mechanism. Chloritization is accompanied by the + 2+ 3+ release of K , Fe , Al and SiO2 to be volume conservative.

Temperature controlled advection can be assumed to control the transport of dissolved elements into fissure direction. The mineral evolution along an exhumation path is conforming to petrographic and mineralogical observations. LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 123

Although the calculations are isochemical the results are consistent with the observations including wall rock minerals and fissure minerals. Theses suggest that elements that were transferred out of the wall rock, precipitates in the fissure and did not transported away.

The remarkable and astonishing lack of zeolites in late fissures in the Lepontine

Alps, compare to the Arvigo locality could be related to pCO2 above critical threshold value that makes zeolite formation impossible. The calculated phase diagrams in the system Ca-Al-Si-O-C-H encourage the fluid evolution to CO2 poor fluids with time, which is observed in fluid inclusions in quartz, that were formed prior to the zeolite formation.

Mass balance calculations for the whole rock suggest a mass loss of 6.8 % and depletion in Al2O3, SiO2, CaO, Fe2O3 and K2O in the altered wall rock. These elements are subsequently found as major components in epidote, prehnite, calcite, adularia, chlorite and zeolites as fracture filling minerals. The mass transfer is associated with an increase in porosity, caused by the volume decrease during albitization and the removal of chlorite in the wall rock.

The mineral paragenesis in low-grade rocks, often result from fluid-rock interaction alteration. The mineralogical, geochemical and textural signatures, caused by low-grade metamorphism, can be interpreted by relatively simple paragenetic schemes, which can be linked to the tectonic and thermal history of the rocks as well on the fluid evolution.

3.9. ACKNOWLEDGMENTS

We are grateful to Giovanni Polti and Alfredo Polti SA for permission to do field work in the active quarry. Special thanks go to the technicians and staff of the Institute of Geosciences, Mineralogy – Geochemistry, University of Freiburg and particularly Hiltrud Müller-Sigmund for her useful advise during EMP analyses and her patience with us at the electron microprobe. Andreas Leemann from the Swiss Federal Laboratories for Materials Testing and Research for impregnation of rock LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 124 samples. A special thanks deserved to the Friedrich Rinne foundation for the financial support.

3.10. REFERENCES

Armbruster T (2000) Cation distribution in Mg, Mn-bearing babingtonite from Arvigo, Val Calanca, Grisons, Switzerland. Schweiz Mineral Petrogr Mitt 80: 279-284 Armbruster T, Kohler T, Meisel T, Nägler TF, Götzinger MA, Stalder HA (1996) The zeolite, fluorite, quartz assemblage of the fissure at Gibelsbach, Fiesch (Valais, Switzerland): crystal chemistry, REE patterns, and genetic speculations. Schweiz Mineral Petrogr Mitt 76: 131-146 Armbruster T, Stalder HA, Gnos E, Hofmann BA, Herwegh M (2000) Epitaxy of hedenbergite whiskers on babingtonite in Alpine fissures at Arvigo, Val Calanca, Grisons, Switzerland. Schweiz Mineral Petrogr Mitt 80: 285-290 Austrheim H (1987) Eclogitization of the lower crustal granulites by fluid migration through shear zones. Earth Planet Sci Lett 81: 221-232 Berger A, Mercolli I, Engi M (2005) Tectonic and petrographic map of the Central Lepontine Alps, 1:100’000. Schweiz Mineral Petrogr Mitt 85: 109-146 Berman RG (1988) Internally-consistent thermodynamic data for minerals in the system Na2O-K2O-CaO-MgO-FeO-Fe2O3-Al2O3-SiO2-TiO2-H2O-CO2. J Petrol 29: 445-522 Bevins RE, Rowbotham G, Robinson D (1991) Zeolite to prehnite-pumpellyite facies metamorphism of the late Proterozoic Zig-Zag Dal Basalt Formation, eastern North Greenland. Lithos 27: 155-165 Bird DK, Schiffman P, Elders WA, Williams AE, McDowell SD (1984) Calcsilicate mineralization in active geothermal systems. Econ Geol 79: 671-695 Bish DL, Ming DW (2001) Natural zeolites: occurrence, properties; applications. Reviews in Mineralogy & Geochemistry, Vol. 45, Mineralogical Society of America, Washington, DC, Bons PD (2001) The formation of large quartz veins by rapid ascent of fluid in mobile hydrofractures. Tectonophysics 336: 1-17 Brughera F (1984) Aquamarin aus dem Steinbruch von Arvigo (Calancatal). Schweizer Strahler 6: 498-501 Bucher K, Frey M (2002) Petrogenesis of metamorphic rocks. Springer, Berlin Carmichael DM (1969) On the mechanism of prograde metamorphic reactions in quartz bearing pelitic rocks. Contrib Mineral Petrol 20: 244-267 Cathelineau M (1988) Cation site occupancy in chlorites and illites as a function of temperature. Clay Minerals 23: 471-485 LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 125

Cathelineau M, Nieva D (1985) A chlorit solid solution geothermometer. The Los Azufres (Mexico) geothermal system. Contrib Mineral Petrol 91: 235-244 Cho M, Liou JG, Maruyama S (1986) Transition from the zeolite to prehnite- pumpellyite facies in the Karmutsen Metabasites, Vancouver Island, British Columbia. J Petrol 27: 467-494 Cho M, Maruyama S, Liou JG (1987) An experimental investigation of heulandite- laumontite equilibrium at 1000 to 2000 bar Pfluid. Contrib Mineral Petrol 97: 43- 50 Ciesielczuk J, Janeczek J (2004) Hydrothermal alteration of the Strzelin granite SW Poland. N Jb Miner Mh 179: 239-264 Clark C, Schmist Mumm A, Faure K (2005) Timing and nature of fluid flow and alteration during Mesoproterozoic shear zone formation, Olary Domain, South Australia. J Metamorph Geol 23: 147-164 Coelho J (2006) GEOISO - A WindowsTM program to calculate and plot mass balances and volume changes occurring in a wide variety of geologic processes. Comput Geosci 32: 1523-1528 Cole DR, Larson PB, Riciputi LR, Mora CI (2004) Oxygen isotope zoning profiles in hydrothermally altered feldspars; estimating the duration of water-rock interaction. Geology 32: 29-32 Coombs DS, Alberti A, Artioli A, Armbruster T, Colella C, Galli E, Grice JD, Liebau F, Mandarino JA, Minato H, Nickel EH, Passaglia E, Peacor DR, Quartieri S, Rinaldi R, Ross M, Sheppard RA, Tillmanns E, Vezzalini G (1998) Recommended nomenclature for zeolite minerals: report of the subcommittee on zeolites of the international mineralogical association, commission on new minerals and mineral names. Mineral Mag 62: 533-571 de Capitani C, Brown TH (1987) The computation of chemical equilibrium in complex systems containing non-ideal solutions. Geochim Cosmochim Acta 51: 2639-2652 De Caritat P, Hutcheon I, Walshe JL (1993) Chlorite Geothermometry: A Review. Clays and Clay Miner 41: 219-239 Diegel S, Ghent ED (1994) Fluid-mineral equilibria in prehnite-pumpellyite to greenschist facies metabasites near Flin Flon, Manitoba, Canada: implications for petrogenetic grids. J Metamorph Geol 12: 467-477 Drake H, Tullborg E-L, Annersten H (2008) Red-staining of the wall rocks and its influence on the reducing capacity around water conducting fractures. Appl Geochem 23: 1898-1920 Engi M, Todd CS, Schmatz D (1995) Tertiary metamorphic conditions in the eastern Lepontine Alps. Schweiz Mineral Petrogr Mitt 75: 347-369 Engvik A, Putnis A, Fitz Gerald JD, Austrheim H (2008) Albitisation of granitoid: the mechanism of plagioclase replacement by albite. Can Mineral 46: 1401-1415 Evans BW (1990) Phase relations of epidote-blueschists. Lithos 25: 3-23 Faryad SW, Dianiska I (2003) Ti-bearing andradite-prehnite-epidote assemblage from the Malá Fatra granodiorite and tonalite (Western Carpathians). Schweiz Mineral Petrogr Mitt 82: 47-56 LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 126

Ferry J, M. (1979) Reactions mechanism, physical conditions, and mass transfer during hydrothermal alteration of mica and feldspar in granitic rocks from South-central Maine, USA. Contrib Mineral Petrol 68: 125-139 Fitz Gerald JD, Stünitz H (1993) Deformation of granitoids at low metamorphic grade. I; reactions and grain size reduction. Tectonophysics 221: 269-297 Freiberger R, Hecht L, Cuney M, Morteani G (2001) Secondary Ca-Al silicates in plutonic rocks: implications for their cooling history. Contrib Mineral Petrol 141: 415-429 Frey M, Bucher K, Frank E, Mullis J (1980) Alpine metamorphism along the Geotraverse Basel-Chiasso - a review. Eclogae Geol Helv 73: 527-546 Frey M, de Capitani C, Liou JG (1991) A new petrogenetic grid for low-grade metabasites. J Metamorph Geol 9: 497-509 Frey M, Mählmann RF (1999) Alpine metamorphism of the Central Alps. Schweiz Mineral Petrogr Mitt 79: 135-154 Fricke G (1952) Ein Prehnit-Vorkommen im Schwarzwald. Aufschluss 3: 94 Gianelli G, Mekuria N, Battaglia S, Cheriscla A, Garofalo P, Ruggieri G, Manganelli M, Gebregziabher Z (1998) Water-rock interaction and hydrothermal mineral equilibria in the Tendaho geothermal system. J Volc Geothermal Res 86: 253- 276 Gottardi G (1989) The genesis of zeolites. Eur J Mineral 1: 479-487 Graeser S, Stalder HA (1976) Mineral-Neufunde aus der Schweiz und angrenzenden Gebieten. Schweizer Strahler 4: 158-171 Grant JA (1986) The isocon diagram - a simple solution the Gresens´equation for metasomatic alteration. Econ Geol 81: 1976-1982 Gresens RL (1967) Composition-volume relationships of metasomatism. Chem Geol 2: 47-65 Hay RL (1966) Zeolites and zeolitic reactions in sedimentary rocks. Geol Soc Amer Special Pap 85 Hay RL (1977) Geology of zeolites in sedimentary rocks. In: Mumpton FA (ed) Mineralogy and geology of natural zeolites, Mineralogical Society of America, Short Course Notes, pp 53-64 Hay RL, Sheppard RA (1977) Zeolites in open hydrologic systems. In: Mumpton FA (ed) Mineralogy and geology of natural zeolites, Mineralogical Society of America, Short Course Notes, pp 93-102 Hay RL, Sheppard RA (2001) Occurrences of zeolites in sedimentary rocks. In: Bish DL, Ming DW (eds) Natural zeolites: occurrence, properties; applications, Reviews in Mineralogy & Geochemistry, Vol. 45, Mineralogical Society of America, Washington, DC, pp 217-234 Jenny H, Frischknecht G, Knopp J (1923) Geologie der Adula. Beitr Geol Karte Schweiz, Schweizerische Geologische Kommision, Bern Johnson GK, Flotow HE, O'Hare PAG (1983) Thermodynamic studies of zeolites: natrolite, mesolite and scolecite. Am Mineral 68: 1134-1145 LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 127

Keller F (1968) Mineralparagenesen und Geologie der Campo Tencia-Pizzo Forno- Gebirgsgruppe. Beitr Geol Karte Schweiz, Schweizerische Geologische Kommision, Bern Kristmannsdóttir H, Tómasson J (1978) Zeolites zones in geothermal areas in Iceland. In: Sand LB, Mumpton FA (eds) Natural zeolites: occurrence, properties, use, Pergamon Press, New York, pp 277-284 Kuniyoshi S, Liou JG (1976) Contact metamorphism of the Karmutsen Volcanics, Vancouver Islands, British Columbia. J Petrol 17: 73-99 Köppel V, Grünenfelder M (1975) Concordant U-Pb ages of Monazite and xenotime from the Central Alps and the timing of the high temperature Alpine metamorphism, a preliminary report. Schweiz Mineral Petrogr Mitt 55: 129-132 Lee MR, Thompson P, Poeml P, Parsons L (2003) Peristeritic plagioclase in North Sea hydrocarbon reservoir rocks: Implications for diagenesis, provenance and stratigraphic correlation. Am Mineral 88: 866-875 Leichmann J, Broska I, Zachovalova K (2003) Low-grade metamorphic alteration of feldspar minerals: a CL study. Terra Nova 15: 104-108 Liou JG (1971) P-T stabilities of laumontite, wairakite, lawsonite, and related minerals in the system CaAl2Si2O8-SiO2-H2O. J Petrol 12: 379-411 Liou JG (1979) Zeolite facies metamorphism of basaltic rocks from the East Taiwan Ophiolite. Am Mineral 64: 1-14 Liou JG (1985) Phase equilibria and mineral parageneses of metabasites in low-grade metamorphism. Mineral Mag 49: 321-333 Liou JG, Kim HS, Maruyama S (1983) Prehnite-epidote equilibria and their petrologic applications. J Petrol 24: 321-342 Maeder UK, Berman RG (1991) An equation of state for carbon dioxide to high pressure and temperature. Am Mineral 76: 1547-1559 Mercolli I, Schenker F, Stalder HA (1984) Geochemie der Veränderungen von Granit durch hydrothermale Lösungen. Schweiz Mineral Petrogr Mitt 64: 67-82 Mullis J, Dubessy J, Poty B, O'Neil J (1994) Fluid regimes during late stages of a continental collision: physical, chemical and stabel isotope measurements of fluid inclusions in fissure quartz from a geotraverse through the Central Alps, Switzerland. Geochim Cosmochim Acta 58: 2239-2267 Nagel T, de Capitani C, Frey M (2002) Isograds and P-T evolution in the eastern Lepontine Alps. J Metamorph Geol 20: 309-324 Neuhoff PS, Fridriksson T, Arnórsson S (1999) Porosity evolution and mineral paragenesis during low-grade metamorphism of basaltic lavas at Teigarhorn, Eastern Iceland. Am J Sci 299: 467-501 Neuhoff PS, Fridriksson T, Bird DK (2000) Zeolite parageneses in the North Atlantic Igneous Provinces: implications for geotectonics and groundwater quality of basaltic crust. Int Geol Rev 42: 15-44 Ogorodova LP, Kiseleva IA, Melchakova LV, Belitskiy IA (2002) Thermodynamic properties of calcium and potassium chabazite. Goechem Int 40: 466-471 LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 128

Oliver NHS (1996) Review and classification of structural controls on fluid flow during regional metamorphism. J Metamorph Geol 14: 477-492 Orvosová M, Majzlan J, Chovan M (1998) Hydrothermal alteration of granitoid rocks and gneisses in the Sb-Au Dúbrava deposit, Western Carpathians. Geol Carp 49: 377-387 Parker RL (1973) Die Mineralfunde der Schweiz, Neubearbeitung. Verlag Wepf & Co, Basel Parneix JC, Beaufort D, Dudoignon P, Meunier A (1985) Biotite chloritization process on hydrothermal altered granites. Chem Geol 51: 89-101 Parneix JC, Petit JC (1991) Hydrothermal alteration of an old geothermal system in the Auriat Granite (Massif Central, France); petrological study and modelling. Chem Geol 89: 329-351 Parry WT, Downey LM (1982) Geochemistry of hydrothermal chlorite replacing igneous biotite. Clays and Clay Miner 30: 81-90 Passaglia E (1970) The crystal chemistry of chabazite. American Mineralogist 55: 1278-1301 Phillips ER, Rickwood PC (1973) The biotite-prehnite association. Lithos 8: 275-281 Poty BP, Stalder HA, Weisbrod AM (1974) Fluid inclusions studies in quartz from fissures of Western and Central Alps. Schweiz Mineral Petrogr Mitt 54: 717- 752 Pouchou G, Pichior F (1991) Quantitative analysis of homogeneous or stratified microvolumes applying the model of "PAP". In: Heinrich KFJ, Newbiry DE (eds) Electron probe quantitation, Plenum Press, New York, pp 31-75 Purdy JW, Stalder HA (1973) K-Ar ages of fissure minerals from the Swiss Alps. Schweiz Mineral Petrogr Mitt 53: 79-98 Putnis A, Putnis CV (2007) The mechanism of reequilibration of solids in the presence of a fluid phase. J Solid State Chem 180: 1783-1786 Ragnarsdottir KV, Walther JV (1985) Experimental determination of corundum solubilities in pure water between 400-700°C and 1-3 kbars. Geochim Cosmochim Acta 49: 2109-2115 Rahn M, Mullis J, Erdelbrock K, Frey M (1994) Very low-grade metamorphism of the Taveyanne greywacke, Glarus Alps. Switzerland. J Metamorph Geol 12: 625-641 Rose NM, Bird DK (1987) Prehnite-epidote phase relations in the Nordre Aputiteq and Kruuse Fjord Layered Gabbros, East Greenland. J Petrol 28: 1193-1218 Rose NM, Bird DK, Liou JG (1992) Experimental investigation of mass transfer - albite, Ca-Al-silicates, and aqueous solutions. Am J Sci 292: 21-57 Ruppe H (1966) Val Calanca - Graubünden. Aufschluss 17: 105-109 Rütti R, Maxelon M, Mancktelow NS (2005) Structure and kinematics of the northern Simano Nappe, Central Alps, Switzerland. Eclogae Geol Helv 98: 63-81 Saigal GC, Morad S, Bjørlykke K, Egeberg PK, Aagaard P (1988) Diagenetic albitization of detrital K-feldspar in Jurassic, Lower, Cretaceous, and Tertiary LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 129

clastic reservoir rocks from offshore Norway; I Textures and origin. J Sed Petr 58: 1003-1013 Sandström B, Annersten H, Tullborg E-L (2008) Fracture related hydrothermal alteration of metagranitic rock and associated changes in mineralogy, geochemistry and degree of oxidation: a case study at Forsmark, central Sweden. Int J Earth Sci. DOI 100.1007/s00531-008-0369-1 Schaltegger U, Gebauer D, von Quadt A (2002) The mafic and ultramafic rock association of Loderio-Biasca (lower Pennine nappes, Ticino, Switzerland); Cambrian oceanic magmatism and its bearing on early Paleozoic paleogeography. Chem Geol 186: 265-279

Senderov EE (1973) Effect of CO2 on the stability of laumontite. Geochem Int 10: 139-114 Simonetti A (1971) Le zeoliti a le loro paragenesi nelle fessure delle rocce del canton Ticino, della Val Calanca e della Val Mesolcina. Boll Soc Ticinese Sci Naturali 62: Spicher A (1980) Tektonische Karte der Schweiz 1:500’000, 2nd edition. Schweizerische Geologische Kommision, Bern Stalder HA (2007) Kluft-Mineralien aus dem Steinbruch von Arvigo im Calancatal. Schweizer Strahler 1: 5-7 Stalder HA, Wagner A, Graser S, Stuker P (1998) Mineralienlexikon der Schweiz. Wepf Verlag, Basel Steck A (1968) Junge Bruchsysteme in den Zentralalpen. Eclogae Geol Helv 61: 387- 393 Steefel CI Hay (2008) Geochemical kinetics and transport. In: Brantley SL, Kubicki JD, White AF (eds) Kinetics of water-rock interaction, Springer, New York, pp 545-589 Sun C-O, Williams RJ, Sun S-S (1974) Distribution coefficients of Eu and Sr for plagioclase-fluid and clinopyroxene-liquid equilibria in oceanic ridge basalt: an experimental study. Geochim Cosmochim Acta 38: 1415-1433 Surdam RC (1973) Low-grade metamophism of tuffaceous rocks in the Karmutsen Group, Vancouver Island, British Columbia. Geol Soc Amer Bull 84: 1911- 1922 Surdam RC (1977) Zeolites in closed hydrologic systems. In: Mumpton FA (ed) Mineralogy and geology of natural zeolites, Mineralogical Society of America, Short Course Notes, pp 65-92 Thompson AB (1970) Laumontite equilibria and the zeolite facies. Am J Sci 269: 267-275

Thompson AB (1971) PCO2 in low-grade metamorphism; zeolite, carbonate, clay mineral, prehnite relations in the system CaO-Al2O3-SiO2-CO2-H2O. Contrib Mineral Petrol 33: 145-161 Thompson AB (1975) Calc-silicate diffussion zones between marble and pelitic schist. J Petrol 16: 314-346 LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 130

Todd CS, Engi M (1997) Metamorphic field gradients in the Central Alps. J Metamorph Geol 15: 513-530 Tulloch AJ (1979) Secondary Ca-Al silicates as low-grade alteration products of granitoid biotite. Contrib Mineral Petrol 69: 105-117 Verdes G, Gout R, Castet S (1992) Thermodynamic properties of the aluminate ion and bayrite, boemite, diaspore and gibbsite. Eur J Mineral 4: 767-792 Vidal O, Parra T, Trotet F (2001) A thermodynamic model for Fe-Mg aluminous chlorite using data from phase equilibrium experiment and natural pelitic assemblages in the 100-600 °C, 1-5 kbar range. Am J Sci 6: 557-592 Wagner A (1968) Mineralien aus den Stenbrüchen von Arvigo. Schweizer Strahler 1: 128-131 Wagner A (1980) Die Mineralien aus den Gesteinsbrüchen von Arvigo im Bild (1. Teil). Mineralienfreund 18: 137-141 Wagner A (1981) Die Mineralien aus den Gesteinsbrüchen von Arvigo im Bild (2. Teil). Mineralienfreund 19: 56-64 Wagner A (1983) Die Mineralien aus dem Val Calanca und den Steinbrüchen von Arvigo. Schweizer Strahler 6: 336-355 Wagner A, Stalder HA, Stuker P, Offermann E (2000a) Arvigo - eine der bekanntesten Mineralfundstellen der Schweiz. Schweizer Strahler 12: 41-70 Wagner A, Stalder HA, Stuker P, Offermann E (2000b) Arvigo - eine der bekanntesten Mineralfundstellen der Schweiz. Schweizer Strahler 12: 118-154 Walker FDL, Lee MR, Parsons L (1995) Micropores and micropermeable texture in alkali feldspars; geochemical and geophysical implications. Mineral Mag 59: 505-534 Walker GPL (1960) Zeolite zones and dike distribution in relation to the structure of the basalts of Eastern Iceland. J Geol 68: 515-528 Walker GPL (1963) The Breiddalur central volcano, Eastern Iceland. Quart J Geol Soc Lond 119: 29-63 Weisenberger T, Bucher K (2008) Porosity evolution and mass transfer during low- grade metamorphism in crystalline rocks of the upper continental crust. In: 33rd IGC International Geological Congress, Oslo MPN03710L Weisenberger T, Bucher K (2009) Zeolite in fissure of granites and gneisses of the Central Alps. submitted to J Metamorph Geol Weisenberger T, Selbekk RS (2008) Multi-stage zeolite facies mineralization in the Hvalfjördur area, Iceland. Int J Earth Sci. DOI 10.1007/s00531-007-0296-6 Weiß S, Forster O (1997) Arvigo, Val Calanca: Kluftminerale aus dem Süden Graubündens. Lapis 6: 13-42 Wenk E (1955) Eine Strukturkarte der Tessineralpen. Schweiz Mineral Petrogr Mitt 35: 311-319 Young B, Dyer A, Hubbard N, Starkey RE (1991) Apophyllite and other zeolite-type minerals from the Whin Sill on the northern Pennines. Mineral Mag 55: 203- 207 LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 131

Zen E (1961) The zeolite facies: an interpretation. Am J Sci 259: 401-409

TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 132

4. TIMING AND MINERAL EVOLUTION DURING LOW-TEMPERATURE FLUID-ROCK INTERACTION ON UPPER CRUSTAL LEVEL: 40Ar/39Ar APOPHYLLITE-(KF) DATING AND APATITE FISSION TRACK ANALYSIS ON ALPINE FISSURES (CENTRAL ALPS/SWITZERLAND)

TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 133

4.1. ABSTRACT

The mineral assemblage quartz, laumontite and apophyllite-(KF) occur in a fissure within the Southern Aar granite located in the Aar Massif (Switzerland). They were formed during exhumation of the Alpine orogen and laumontite and apophyllite marks the latest fissure minerals in the Central Alps. A combined study of 40Ar/39Ar age dating, apatite fission track (FT) and chemical characterization of tunnel and surface samples are present to carry out the position of low-temperature water-rock interaction in respect to the Alpine history.

Apatite FT analysis yields an exhumation rate of 0.45 mm a-1, a cooling rate of 13 °C Ma-1 and a geothermal gradient of 28 °C km-1. Combining these with the 40Ar/39Ar plateau age for apophyllite of ∼2 Ma, a minimum formation temperature and depth of 70 °C and 2800 m, respectively can be assumed. Temperature-time evolution of fissures in the Aar Massif and thermodynamic mineral evolution indicate that laumontite were formed between 7 and 2 Ma before present at temperatures between 150 and 70 °C.

Elements for laumontite formation derived during dissolution of primary minerals. Changes of laumontite chemistry could be an effect of temperature drop or a change in fluid chemistry that would be supported by later apophyllite formation.

Keywords: laumontite, apophyllite-(KF), 40Ar/39Ar, apatite fission track, granite

4.2. INTRODUCTION

Fluid-rock interaction is an important process in the upper crust, with respect to porosity evolution, permeability and fluid migration. The fluid composition monitors the water-rock interaction and controls the dissolution of primary minerals and the re- precipitation of secondary minerals in open spaces (e.g. Nordstrom et al., 1989; Bucher and Stober, 2001). Thereby the formation of zeolites and apophyllite is widespread in basaltic rocks (e.g. Walker, 1959; Belsare, 1969; Sukheswala et al., TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 134

1974; Keith and Staples, 1985; Young et al., 1991; Neuhoff et al., 1997; Weisenberger and Selbekk, 2009) as well as in upper continental crust, particularly in hydrothermal fractures and veins in granites and gneisses (e.g. Borchardt et al., 1990; Borchardt and Emmermann, 1993; Armbruster et al., 1996; Freiberger et al., 2001; Fujimoto et al., 2001; Ciesielczuk and Janeczek, 2004, Weisenberger and Bucher,

2009). The formation of zeolites requires a H2O dominated fluid (Zen, 1963; Senderov, 1973; Weisenberger & Bucher, 2009) and is restricted to low temperature (<250 °C), low pressure (<200 MPa), water-saturated environments.

Considering deep continental fluids, they have the potential to form zeolites (Stober and Bucher, 2004). High-pH waters from the NEAT tunnel in the basement of the Swiss Aar Massif (Fig. 4.1)(Seelig et al., 2007), water from the crystalline basement at Stripa, Sweden, (Nordstrom et al., 1989), Bad Urach (Stober and Bucher, 2004) and from the Black Forest basement (Bucher and Stober, 2000) are all oversaturated in respect of zeolites. Therefore a detailed study of fissure minerals can give important information about the hydrogeochemical evolution in the upper continental crust.

The formation of zeolites and apophyllite-(KF) marks the last step of the long (20 Ma) history of fracture generation and mineralization as result of the uplift and exhumation of the Alpine orogen (e.g. Weisenberger & Bucher, 2009).

The Gotthard-NEAT tunnel is a good example to study subsurface samples, approximately 2000 m below the surface, which are usually not accessible. For instance laumontite is the most widespread zeolite in Alpine fissures, exposed underground in tunnels sections or in active quarries. Because the mineral decomposes by dehydration at room temperature and decays to a powdery mass, laumontite occurs rarely in surface outcrops (Weisenberger and Bucher, 2009). The timing of low-temperature water-rock interaction is commonly difficult to establish because of the paucity of suitable material for geochronology. The appearance of apophyllite-(KF) following laumontite during late stages of Alpine fissure mineralization gives the opportunity to get information on the age of this event. Apophyllite as tool of age dating by 40Ar/39Ar techniques is not used widely. However Fleming et al. (1999) and Molzahn et al. (1999) evaluated the feasibility of using apophyllite for geochronology by the 40Ar/39Ar method on secondary mineralization TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 135 in the Transantarctic Mountains and demonstrated that apophyllite dating can produce geologically meaningful ages.

The aim of this study is to assess information about the timing of low- temperature mineralization in the upper continental crust and carry out the relation to the temporal evolution of fissure mineralization during fluid-rock interaction at Alpine exhumation stage.

To assign these information geochemical characterization of fissure minerals are presented, as well as Ar/Ar age determination on apophyllite and apatite FT analysis to obtain the local exhumation rate and therefore estimate the formation depth of late stage fissure minerals.

4.3. GEOLOGICAL SETTING

The Aar Massif is one of the external massifs of the Central Swiss Alps (Frisch et al., 1990) situated in the Helvetic zone (Fig. 4.1). The Aar massif is formed by Hercynian intrusives, emplaced into a polymetamorphic basement (Fig. 4.1) of Paleozoic to late Proterozoic age (Grünenfelder et al., 1964; Gulson and Rutishauser, 1976). The Upper Carboniferous Central Aar granite is a lens-shaped batholith, consisting of granites and granodiorites, and is exposed over an area of about 550 km2. The Southern Aar granite is only exposed in the eastern part and can be traced over 20 km in W-E direction and 1-2 km in N-S direction (Fig. 4.1). The Southern Aar granite has earlier been considered as marginal southern facies of the Central Aar granite (Huber 1948), but modern age determination indicates that the Southern Aar granite is not genetically related to the Central Aar granite. Schaltegger and Corfu (1992) determined (U-Pb- method on zircon and allanite) the age of the emplacement of the Southern Aar granite took place at around 350 Ma. This age predates the late Hercynian Central Aar granite, which was emplaced in a short period of 2-4 Ma at around 298 Ma (Schaltegger and Corfu 1992).

The Aar Massif was subject to Cenozoic Alpine greenschist-facies metamorphism. The N-S transection is marked by different isograds with increasing metamorphic grade from the north to the south (Bambauer and Bernotat, 1982; Frey TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 136 et al., 1980; Frey and Mählmann, 1999): the first appearance of green biotite (Steck and Burri, 1971), disappearance of stilpnomelane (Jäger et al., 1967) and the transformation isograd of microcline/sanidine (Bambauer and Bernotat, 1982; Bernotat and Bambauer, 1982; Frey and Mählmann, 1999). Peak metamorphism exceeds zeolite facies all-over.

FIG. 4.1: (a) Detailed geological map of the Eastern Aar Massif (modified after Labhart, 1977) including sample locality and track of the Gotthard new railway base tunnel (NEAT). (b) Outline of Switzerland and the position of the central external massifs in gray. Rectangle mark the section a.

During late-orogenic exhumation the considerable increase of erosion rate and denudation forced the evolution of shear zones related to backthrusting, parallel normal faulting and the opening of fissures and gashes. These act as pathways for fluids which seep through and react with the surrounding rocks to finally form secondary fissure minerals (Berger et al., 2005; Mullis 1995, 1996). An exhumation TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 137 rate of 0.5 mm a-1 for the Reuss valley in the Central Aar Massif is proposed by Michalski and Soom (1990) for the past 27 Ma (apatite and zircon FT), with a cooling rate of 13 °C Ma-1, which are in good agreement with uplift rates of 0.3 - 0.6 mm a-1 during the last 6-10 Ma (Schaer et al., 1975). Using exhumation rates and trapping temperatures of early fluid inclusions (Mullis, 1996) the first opening of fissures and precipitation of fissure minerals in the Aar- (Zinggenstock) and Gotthard Massif (La Fibbia) is determined to around 20 Ma ago.

4.4. SAMPLES AND METHODS

4.4.1. Analytic

Whole rock analyses were performed by standard X-ray fluorescence (XRF) techniques at the Institute of Geosciences (Mineralogy and Geochemistry) at the University of Freiburg/Germany, using a Philips PW 2404 spectrometer. Pressed powder and Li-borate fused glass discs were prepared to measure contents of trace and major elements, respectively. The raw data were processed with the standard XR- 55 software of Philips. Relative standard deviations are < 1 % and < 4 % for major and trace elements, respectively.

Quantitative mineral analyses were performed at Institute of Geosciences (Mineralogy and Geochemistry), University of Freiburg, using a CAMECA SX 100 electron microprobe equipped with five WD spectrometers and one ED detector with an internal PAP-correction program (Pouchou and Pichior, 1991). Major and minor elements in zeolites were determined at 15 kV accelerating voltage and 10 nA beam current with a defocused electron beam of 20 µm in diameter with counting time up to 20 s. Na and K were counted first to minimize the Na and K loosed during determination. Since zeolites lose water when heated, the crystals were mounted in epoxy resin to minimize loss of water due to the electron bombardment. Natural and synthetic standards were used for calibration. The charge balance of laumontite formula is a reliable measure for the quality of the analyses and correlates with the difficulties related to the thermal instability of zeolites in microprobe analysis. A useful error test investigates the charge balance between the non-framework cations TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 138 and the amount of tetrahedral Al (Passaglia, 1970). Analyses are considered acceptable if the sum of the charges of the extra-framework cations (Ca2+, Sr2+, Ba2+, Na+, and K+) is within 10% of the framework charge (Al3+).

Isotopic dating was carried out at the 40Ar/39Ar laboratory of the Department of Mineralogy, University of Geneva, Switzerland. Crystals of apophyllite were crushed and clear, inclusion free chips were packed in copper foil. The samples were irradiated for 3 hours at 1 MW in the Oregon State University CLICIT facility, and J values were calculated via the analysis of Fish Canyon Tuff sanidines, which were spaced by <1cm throughout the columnar irradiation package. Stepwise degassing was performed using a 30W CO2-IR laser, and extracted gas was purified in a UHV extraction line equipped with SAES AP10 and GP50 getters, prior to analysis. Isotope ratios were measured with a GV instruments ARGUS multi-collector mass spectrometer, equipped with four high-gain (10-12 Ohms) Faraday collectors for the analysis of 39Ar, 38Ar, 37Ar and 36Ar and one single 10-11 Ohms Faraday collector for the analysis of 40Ar. Blanks were measured between every three degassing steps and before every new sample. Data reduction was performed using the program ArArCALC (Koppers, 2002) and corrections were applied for post-irradiation decay 37 39 of Ar (T0.5 = 35.1 days) and Ar (T0.5 = 269 years). The mass discrimination factor during analysis was 1.00436 based on ongoing measurements of 40Ar/36Ar ratios in quantitatively calibrated air shots from an air pipette. Correction factors for interfering Ca- and K- derived isotopes have been calculated from 10 analyses of two Ca-glass samples and 22 analyses of two pure K-glass samples, and are: 36Ar/37Ar(Ca)=2.603E-4 ± 2.373E-9, 39Ar/37Ar(Ca)=6.501E-4 ± 7.433E-9 and 40Ar/39Ar(K)=1.547E-2 ± 7.455E-7.

For the FT measurement, apatite grains were separated from sample 115900 and SueArGr (8-10 kg rock material) using standard crushing, magnetic and heavy liquid techniques. Separated apatites (fraction 63-300 µm) were mounted with epoxy on glass slides, polished, and etched for 20 s with 5N HNO3 to reveal the spontaneous fission tracks. Mounts were covered with U-free white mica sheets and sent to irradiation at the FRM-II reactor facility in Garching/Germany, together with other samples and top and bottom CN5 dosimeter glasses. After irradiation, the white micas were removed and etched for 45 minutes in 40 % HF to reveal the induced tracks.

Fission tracks were recorded and confined track lengths and Dpar measured using TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 139 transmitted and reflected light at 1600x magnification on a Zeiss Axioplan microscope, equipped with a computer-driven stage and a digitizing tablet at the Geosciences Department of Basel University. 40 grains were counted per sample in order to reduce the age error of the expectedly young ages. Central ages (Galbraith and Laslett, 1993) were calculated using the IUGS-recommended zeta calibration approach (Hurford and Green, 1983). Statistical χ2 tests were applied to search for internal variation assuming the single grain ages of mono-population samples to be Poissonian distributed. Failure of this test (P(χ2) < 5%) may indicate the presence of internal age variation due to partial annealing or chemical inhomogeneities.

FIG. 4.2: Photograph and schematic illustration showing fissure mineral assemblag at the 115935 (Table 4.1) locality in the NEAT tunnel from the sample KB868. (a) Photograph of fissure assemblages from the same fissure than sample KB868: clear apophyllite “cubs” overgrown by laumontite needles. (b) Schematic sketch on thin section scale of sample KB868 fissure mineral succession Qtz → Lmt → Apo. (c) Representative microphotographs under crossed polarized light.

TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 140

4.4.2. Samples

All sample sites are located in the Southern Aar granite (Fig. 4.1, Table 4.1). Sample KB868 is collected from a fissure during excavation of the Gotthard NEAT tunnel, 15935 m south of the north portal (Fig. 4.1, 4.2). The sample exhibits a fissure assemblage with the chronological order: quartz, adularia, laumontite and apophyllite (Fig. 4.2). Additionally chlorite and milarite are found in the same fissure (P. Amacher pers. com.). Euhedral fissure minerals, up to 12 mm in size (Fig. 4.2) grew on a thin leached matrix (< 1 cm in thickness). Leaching is indicated by higher porosity than in the fresh tunnel sample 115900 (Table 4.1) from the same lithological unit of the Southern Aar granite. A third sample (SueArGr) is collected from a surface outcrop of the Southern Aar granite (Fig. 4.1, Table 4.1) The vertical offset between the tunnel- (115900, KB868) and the surface specimen (SueArGr) is 1623 m (Fig. 1, Table 4.1).

TABLE 4.1. Sample description. Indicated x and y coordinates corresponds to the Swiss coordinate net (units: km). Sample No x y altitude [m] Description Assemblage* SueArGr 700 588 174 999 2 123 Rock sample of the Southern Aar granite Qtz, Kfs, Pl, collected on a surface outcrop. XRF- analysis, Ms, Chl, Ep, apatite FT Ap, Ttn, Rt, Py, (Bt) 115 900 700 062 173 863 ∼500 Drill core of the Southern Aar granite collected Qtz, Kfs, Pl, in the Gotthard-NEAT tunnel; 115900 m in Ms, Chl, Ep, distance to the north portal. XRF- analysis, Ap, Zrn, Rt apatite FT. KB868 700 079 173 832 ∼500 Fissure assemblages in the Southern Aar Apo, Lmt, (115 935) granite: quartz, adularia, laumontite and Kfs, Qtz, Chl, apophyllite; additionally chlorite and milarite Mil are found in the same fissure. Fissure: 40 x 60 x 30 cm. 15935 m in distance to the north portal in the Gotthard-NEAT tunnel *Abbreviations according to Bucher and Frey (2002); Mil = milarite, minerals in parentheses are metastable

4.5. RESULT

4.5.1. Petrography and geochemistry

The Southern Aar granite consists predominantly of albite, K-feldspar, quartz, muscovite and chlorite. The rock is deformed with K-feldspar megacrysts up to 1 cm TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 141 in length in a matrix of plagioclase, quartz, muscovite and chlorite as mafic components, usually not exceeding grain sizes of some mm (Fig. 4.3). Both rock samples (115900, SueArGr) exhibit the same mineralogy (Table 4.1), whereas the sample from the surface (SueArGr) is strongly be weathered.

FIG. 4.3: Representative microphotographs of mineral assemblages of rock sample 115900 and SueArGr. (a) Recrystallization quartz, saussuritizated plagioclase, chlorite and K-feldspar in sample 115900; plane polarized light. (b) Fracture between saussuritizated plagioclase filed with epidote in sample SueArGr; crossed polarized light. (c) Epidote flakes in between chlorite sheets in sample 115900; crossed polarized light.

TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 142

TABLE 4.2. Bulk rock geochemistry SueArGr 115 900 wt %

SiO2 68.86 67.50

TiO2 0.42 0.47

Al2O3 16.23 15.46 tot Fe2O3 2.62 3.16 MnO 0.06 0.06 MgO 1.19 1.29 CaO 2.10 2.66

Na2O 4.15 3.78

K2O 3.56 4.03

P2O5 0.22 0.23 LOI 1.05 0.91 Totals 100.67 99.80 ppm V 37 44 Cr 20 33 Ni 16 23 Cu 4 22 Zn 55 56 Rb 161 145 Sr 572 693 Zr 162 248 Ba 1071 1208

Plagioclase is altered to sericite, whereas the fine muscovite flakes are concentrated in the cores of the plagioclase grains. Saussuritization of plagioclase causes the formation of epidote that occurs as small inclusion therein, as well as interstitial filling in chlorite and in small (< 100 µm) veins (Fig. 4.3). Quartz crystals (Fig. 4.3) show a characteristic fabric of dynamic recrystallization by subgrain rotation that was caused during Alpine deformation. The assemblage K-feldspar and chlorite suggests that the mineralogy of the Southern Aar granite has been altered and retrogressed to temperatures below 400°C (Bucher and Frey, 2002). A few altered relictic biotite grains are present. Pseudomorphic replacements of chlorite after biotite often show preserved sagenitic intergrowth. The K-feldspar is microcline that shows a tartan plain pattern and often exhibits perthitic exsolution that indicates that the samples come from locations north of the microcline/sanidine transformation isograd (Bambauer and Bernotat, 1982; Bernotat and Bambauer, 1982; Frey and Mählmann, 1999). Accessory minerals include apatite, titanite, pyrite, allanite and zircon.

The peraluminous Southern Aar granite (Table 4.2) shows a slightly lower SiO2 content than the Central Aar granite (69.85 ± 3.60; Schaltegger, 1990) that is derived from calc-alkaline magmatism during the Hercynian orogenesis (Schaltegger and TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 143

Corfu, 1992). MgO, TiO2 and Fe2O3 contents are very low, whereas the K2O und

Na2O contents are elevated. Ca values are 2.10 and 2.66, respectively, slightly higher to the Central Aar granite of the Reuss valley (1,77 ± 0.78; Schaltegger, 1990) and significant higher than accordant units in the Grimsel area (Schaltegger, 1990). The trace elements are dominated by Ba and Sr. The bulk rock geochemistry of both samples (115900 and SueArGr) shows no major differences (Table 4.2).

4.5.2. Mineralogy and geochemistry

4.5.2.1. Laumontite

Laumontite is a monoclinic (space group C2/m) zeolite. It forms thin, elongated fibbers or prisms elongated along the c-axis with a squared cross-section (Fig. 4.2). Twinning occurs on {100} to form “swallow tail” or “V” twins. It is white with a length between <1 to 12 mm.

FIG. 4.4: Chemical variation in laumontite from sample KB868 as function of the Si/Al ration and extra-framework cations.

The composition of laumontite was obtained on sample KB868 (Table 4.3) and is close to endmember composition Ca4(Al8Si16O48) •18 H2O (Armbruster and Kohler, 1992). Ca is the dominant extra-framework cation (average value of 96 mole%), with TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 144

Na and K typically below 5 mole% (Table 4.3, Fig. 4.4). Additionally Ba is incorporated as extra-framework cation up to 3 mole%. Sr occurs only in traces (Fig. 4.4). The contents of Na, K and Ba increase from core to rim of zoned crystals (Fig. 4.4, 4.5). Often a second laumontite generation is observed (Fig. 4.5), enriched in K and Na. The Si/(Si+Al) ratio varies slightly between 0.67 and 0.69 (Fig. 4.4), with an average of 0.68. Extra-framework cations vary with the Si/(Si+Al) ratio (Fig. 4.4). With increasing Si/(Si+Al) ratio the Ca content decreases, while Na, K and Ba increase. This can be expressed by the coupled substitution of Si4+ + (Na+, K+, Ba2+) = Al3+ + Ca2+.

TABLE 4.3. Laumontite chemistry Sample no. KB 868.3 KB 868.3 KB 868.3 KB 868.2 KB 868.1 KB 868.1 Analysis no. 4 6 7 10 15 16 wt.%

SiO2 53.12 52.78 53.42 53.47 52.76 53.85

Al2O3 21.52 22.12 21.58 21.06 20.83 21.33 CaO 11.72 12.24 11.68 11.49 11.33 11.56 SrO 0.00 0.04 0.00 0.02 0.00 0.00 BaO 0.25 0.06 0.33 0.42 0.48 0.45

Na2O 0.09 0.03 0.12 0.16 0.18 0.16

K2O 0.15 0.07 0.00 0.15 0.02 0.19

Totala 86.85 87.36 87.20 86.78 85.62 87.55

Si 16.228 16.044 16.250 16.357 16.352 16.334 Al 7.748 7.925 7.737 7.593 7.609 7.625 Ca 3.836 3.986 3.807 3.766 3.762 3.757 Sr 0.000 0.007 0.000 0.004 0.000 0.000 Ba 0.030 0.007 0.039 0.050 0.058 0.053 Na 0.053 0.018 0.071 0.095 0.108 0.094 K 0.058 0.027 0.000 0.059 0.008 0.074 O 48 48 48 48 48 48

E%b -1.22 -1.51 -0.34 -2.57 -2.03 -2.09 Si/(Si+Al) 0.68 0.67 0.68 0.68 0.68 0.68 aTotals include traces of Mg, Mn and Fe. b E % = (100*((Al)- (Na+K)+2(Mg+Ca+Sr+Ba)/(Na+K)+2(Mg+Ca+Sr+Ba)), measure of charge balance

4.5.2.2. Apophyllite-(KF)

Apophyllite (KCa4Si8O20(F,OH) •8 H2O) occurs as overgrowth on laumontite in Alpine fissures (Fig. 4.2). It forms transparent tetragonal pseudo-cubes with truncated spikes of rhomboid faces that end in a pyramid. Sizes are up to 1.5 cm. The apex is truncated in which case the appearance is cubic (Fig. 4.2). TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 145

TABLE 4.4. Apophyllite chemistry Sample no. KB 868.1 KB 868.1 KB 868.2 KB 868.2 KB 868.3 KB 868.3 Analysis no. 2 3 10 11 16 18 wt %

SiO2 50.47 50.86 50.63 50.64 50.18 50.96

Al2O3 0.99 0.76 0.85 0.96 0.76 0.95 CaO 24.25 24.33 24.31 24.38 24.29 24.35

Na2O 0.27 0.11 0.11 0.14 0.10 0.25

K2O 4.32 4.63 4.68 4.67 4.57 4.18 F 1.95 2.06 2.02 2.06 2.02 2.06 -O≡F 0.82 0.87 0.85 0.87 0.85 0.87 Totala 81.43 81.94 81.77 82.03 81.11 81.95

Si 7.857 7.884 7.868 7.849 7.866 7.875 Al 0.182 0.139 0.156 0.175 0.140 0.173 Ca 4.045 4.041 4.048 4.048 4.080 4.032 Na 0.081 0.033 0.033 0.042 0.030 0.075 K 0.858 0.916 0.928 0.923 0.914 0.824 Total 13.022 13.020 13.035 13.046 13.032 12.985 O 20 20 20 20 20 20 F 0.960 1.010 0.993 1.010 1.001 1.007

K/Cab 0.178 0.190 0.193 0.192 0.188 0.172 aTotals include traces of Ba, Fem Mg, Mn and Sr. b ratio of weight

FIG. 4.5: K (a) and Na (b) element concentration map for laumontite in sample KB868. The electron microprobe concentration map clearly shows 2 distinct laumontite generations. Noticeable the sector zoning in the older generation. TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 146

The chemical composition of apophyllite from the NEAT tunnel is given in Table 4.4. It is near the stoichiometric composition of apophyllite-(KF) end-member. The individual crystals are relatively homogenous in terms of chemistry, with small differences in concentrations of minor components like Al and Na, with average values of 0.14 and 0.05 mole%, respectively and no other elements occur in apophyllite.

4.5.3. Ar/Ar age

The apophyllite-(KF) crystals (TW003-APO, ATW-APO) from sample KB868 were selected as age marker for secondary mineral formation. The 40Ar/39Ar total fusion and incremental-heating results for the samples are summarized in Table 4.5 and Fig. 4.6. The K/Ca ratio measured from nucleogenic Ar isotopes has the value of 0.19 and 0.18, respectively and is in good agreement with the values measured by electron microprobe 0.17-0.19 (Table 4.3).

For sample ATW-APO a 40Ar/39Ar total fusion age of 2.11 ± 0.06 Ma is obtained (Fig. 4.6). The error plateau age of the sample is 2.04 ± 0.14, by excluding the first three heating steps to improve the visual fit. These shows different ages, suggesting that these steps have either excess radiogenic argon, which they captured from hydrothermal fluids, impurities on surface of the crystal, or lost radiogenic argon from close to grain boundary sites.

For sample TW003-APO the weighted plateau age of 1.96 ± 0.08 Ma (Fig. 4.6) over 10 out of 11 steps and a 40Ar/39Ar total fusion age is 1.83 ± 0.05 Ma. The very low 36Ar content suggest that almost all of the 40Ar from the sample is radiogenic and not from fluid inclusions or argon which was trapped during crystallization.

Both samples overlap in age by using the plateau age, which gives the best statistical requirement and an apparent age of ∼2 Ma is reasonable.

40 39 TABLE 4.5. Ar/ Ar increment heating ages of apophyllite Sample no. Total fusion Plateau age MSWD 39Ar % of Isochron age 40Ar/36Ar J age (Ma) (Ma ±2σ) total (Ma ±2σ) intercept

ATW - APO 2.11 ± 0.06 2,04 ± 0.14 9.74 99.50 2.22 ± 0.41 225 ± 118 0.0008215 TW003 - APO 1.83 ± 0.05 1.96 ± 0.08 1.61 56.01 1.46 ± 0.81 407 ± 189 0.0008222

TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 147

40 39 FIG. 4.6: Plots of Ar/ Ar incremental-heating spectra for apophyllite.

4.5.4. Apatite fission track analysis

Two samples were collected for apatite FT analysis along a vertical section in the Southern Aar granite (Fig. 4.1, Table 4.1). The results are presented in Table 4.6 and Fig 4.7. Apatite FT ages are 9.6 Ma (SueArGr) and 6.1 Ma (115900), respectively, with 1σ age errors of less than 10 %. Both samples pass the χ2 test, indicating that the TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 148 single grain ages belong to one single age population. The confined length measurements give mean FT lengths of 14.1 and 13.7 µm that corresponds to an undisturbed steady cooling behavior.

TABLE 4.6: Apatite fission track data. For sample location see Fig. 4.1 and Table 4.1. Sample Mineral Spontaneous Induced Pχ2 Dosimeter Central D(par) Mean S.d. of No and No. rs ri rd* FT Age (Ma) (+/-S.d.) Track distribution Crystals (-2 /+2 ) Length (No. Tracks) (Ns) (Ni) (Nd) σ σ 115 900 apatite 0.007 0.315 87 % 16.08 6.1 2.33 13.70 1.21 (40) (133) (6018) (12518) (-1.0/+1.2) (± 0.29) (32) SueArGr apatite 0.011 0.316 92 % 01605 9.6 3.07 14.08 1.25 (40) (194) (5580) (12491) (-1.3/+1.5) (± 0.48) (100) (i) Track densities are (x107tr cm-2), *=(x105 tr cm-2) numbers of tracks counted (N) shown in brackets; (ii) analyses by external detector method using 0.5 for the 4π/2π geometry correction factor; (iii) apatite ages calculated using dosimeter glass CN5 with

ζCN5 =344 ± 5; (iv) P(χ2) is probability for obtaining χ2 value for v degrees of freedom, where v = no. crystals – 1; (v) track length and D(par) data are given in 10-6m, S.d. = 1σ standard deviation.

Dpar measurements (Table 4.6) revealed a significant difference in Cl content (and thus in Dpar) between the two samples suggesting that the closure temperature of the surface sample (SueArGr) is slightly higher than the one of the sample from the tunnel (Donelick et al., 2005). However, preliminary estimations revealed that the influence of such closure temperature variation on the estimation of the depth of apophyllite formation would be negligible in comparison to the variation introduced by the age errors.

FIG. 4.7: Apatite fission track length data.

TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 149

4.6. DISCUSSION

4.6.1. Mineral reaction

Secondary minerals formed during precipitation of oversaturated hot fluids in respect of the secondary minerals. Structural and textural evidences, like leaching zones and porosity increase in the altered wall rock (Weisenberger and Bucher, 2009) imply that primary minerals of the host rock are dissolved along the fractures and supply elements necessary for zeolite formation (Eq. 1).

2+ - Ca + 2 AlO2 + 4 SiO2,aq + 4 H2O ⇒ CaAl2Si4O12 •4 H2O (Lmt) (1)

Those hot aqueous fluids reach a high degree of super saturation with respect to zeolites as observed in the NEAT tunnel (Seelig et al., 2007) and in other deep continental fluids (Urach geothermal site, German continental deep drilling site KTB; Stober and Bucher, 2004).

Sources of Ca, Al and Si in the wall rock, which are necessary for the laumontite formation, are albite, clinozoisite, quartz and calcite (Eq. 2, 3). Clinozoisite (epidote), calcite and albite are present in the host rock as result of Alpine greenschist facies metamorphism due to the consumption of prealpine plagioclase. Whereas plagioclase is considered to be the source for elements that form zeolites in rocks of higher Alpine metamorphism like in Arvigo and in the Gotthard Massif (Weisenberger and Bucher, 2009).

2 Ca2Al3Si3O12(OH) (Czo) + 6 SiO2,aq + CO2 + 11 H2O ⇒ 3 CaAl2Si4O12 •4 H2O

(Lmt) + CaCO3 (Cc) (2)

Keep in mind that the reactions involve a transport step between dissolution and precipitation. The additional silica necessary for the formation of zeolites during clinozoisite dissolution may either be derived locally from dissolution of primary quartz or albite (Eq. 3) or from externally derived SiO2,aq. TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 150

This plausible reaction mechanism co-precipitates laumontite and calcite, which is a common assemblage in Alpine fissures (Weisenberger and Bucher 2009). However, calcite does not occur in our sample, which may indicates that calcite saturation was not obtained by the fluid.

+ Ca2Al3Si3O12(OH) (Czo) + NaAlSi3O8 (Ab) + 2 SiO2,aq + 7 H2O + H ⇒ + 2 CaAl2Si4O12 •4 H2O (Lmt) + Na (4)

This reaction consumes albite and clinozoisite and forms laumontite, and is accompanied by an increase in pH and the total of dissolved solids (TDS). The proposed reaction is supported by high Na+, high pH, and high degrees of over- saturation with respect to zeolites in deep groundwater reported from the NEAT Gotthard rail base tunnel (Seelig et al., 2007).

During laumontite growth an increase of Na, K and Ba can be observed (Fig. 4.4, 4.5). This change in chemistry can either be related to change in formation temperature, or to change in fluid chemistry during growth.

Since bulk Ca is low to very low in granites of the Aar Massif dissolution of widespread matrix and fissure fluorite (Stalder et al., 1998) provides some of the Ca necessary for zeolite development, and also F for late apophyllite growth:

+ 4 CaF2 (Flt) +8 SiO2,aq + 12 H2O + K ⇒ + - KCa4Si8O20(F) • 8 H2O (Apo) + 8 H + 7 F (5)

Reaction (5) consumes K+ and dissolves silica in addition to fluorite. The reaction releases F- to the water as a by-product of apophyllite (or laumontite) formation. The proposed reaction mechanism is supported by the presence of leached fissure fluorite in the Aar granite and by ultra-high fluoride concentrations in hot deep groundwater reported from the Gotthard rail base tunnel (Seelig et al., 2007). Seelig et al. (2009) reported pronounced fluoride concentrations in the tunnel waters raging from 5 to 29 mg L-1, whereas fluoride derived mostly from biotite alteration and fluorite leaching.

TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 151

4.6.2. Depth and temperature estimation

Although many deep continental fluids are oversaturated with respect to zeolites (e.g. Bucher and Stober, 2001; Seelig et al. 2007), the zeolite forming process could never been observed in situ. Therefore we used the approach to assign a minimum depth and temperature by comparing exhumation rate and the age of apophyllite.

The variation in the apatite FT ages reveals the vertical uplift of the two samples that differs in altitude. The ages gives the time which elapsed between passing the 120 °C isotherm and present time. Thus, a cooling rate and uplift rate can be calculated using the present rock temperature in the tunnel of 43 °C and a mean annual surface temperature of ∼0 °C. Therefore a cooling rates of 12.5 °C Ma-1 can be assessed, which is in agreement to already known cooling rates (Michalski and Soom, 1990).

Using the apatite FT age and the vertical height difference (Table 4.1) of the two sample localities an exhumation rate of 0.45 mm a-1 can be assessed agreeing with already known exhumation rates (Schaer et al., 1975; Michalski and Soom, 1990), observed along the Reuss valley ∼10 km west of the studied sample localities.

This implies that apophyllite with an Ar/Ar age of ∼2 Ma is formed ∼900 m below the NEAT tunnel level, or -400 m below sea level. This depth can be supposed as minimum depth of laumontite formation.

The geothermal gradient obtained by the product of cooling rate and the inverse of the exhumation rate results in value of 28 °C km-1 for the geothermal gradient in the studied area and coincidence with a achieved geothermal gradient in the Aar Massif by Reinecker et al. (2008), ranging from 25 - 30 °C km-1.

Taken the rock temperature of 43°C at the sample locality in the NEAT tunnel and the geothermal gradient the apophyllite Ar/Ar age reflects a minimum formation temperature for laumontite of ∼70 °C. However, whether the Ar age represents the formation temperature or a closure temperature below the apophyllite formation temperature is still unknown. Nevertheless, the appearance of apophyllite in basalts associated with zeolites (e.g. Betz, 1981; Keith and Staples, 1985) suggests formation temperatures in the same order.

Considering the overburden of ∼1900 m with respect to the apophyllite sample locality in the NEAT tunnel a total depth of 2800 m is presumed, which corresponds TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 152 to a hydrostatic pressure of 28 MPa. Regarding deep continental drill holes, like the German continental deep drilling site KTB, laumontite occurs down to depth of ∼5300 m (Borchardt and Emmermann, 1993). Stober and Bucher (2004) suggested the formation of laumontite at a depth of ∼3500 m, based on the over-saturation of fluids. However, the laumontite saturation index are calculated on the observed temperature of 150 °C which is twice as high as assumed minimum temperature in this study. Therefore an upper limit has to be assessed using thermodynamic modeling as well as the integration in the Alpine fissure history.

4.6.3. Thermodynamic approach

A thermodynamic approach was used to show the mineral evolution along a PT-path (Fig. 4.8). Computed assemblages stability diagrams were modeled with the Theriak/Domino software of de Capitani and Brown (1987) based on the thermodynamic data by Bermann (1988), Evans (1990), Frey et al. (1991).

Bulk rock composition of the unaltered rock sample 115900 (Table 4.2) has been recalculated to atomic proportions. To simplify the diagrams, Ti and Mn have been ignored. The presence of chlorite and epidote indicate that Fe occurs in di- and trivalent state and that some provision for the redox state is need to be made. Thereby modeling with different redox state conditions was done with the result that the ratio 1/1 of Fe3+/Fe2+ reflects the best fit with observed mineral inventory on observed thin section scale (Fig. 4.3). Due to the fact that alteration is driven by hydrothermal process H2O was set in excess. CO2 was excluded due to the fact that fluid inclusions in fissure-quartz of the studied region shows a very low CO2 content (Mullis et al.,

1994) and the knowledge that relative low CO2 activities are ample to destabilize zeolites (e.g. Senderov, 1973; Weisenberger and Bucher, 2008).

Figure 4.8 gives a predicted assemblage evolution along a PT path, whereas pressure conditions were chosen to be in between hydrostatic and lithostatic pressure. Calculations were done assuming closed system conditions, with the exception of the variable H2O content. Albite, quartz, K-feldspar, muscovite, biotite and epidote are the stabile phases at peak Alpine metamorphic conditions. Along the cooling path biotite and albite deceases whereas the epidote content increase. At a temperature of TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 153

330 °C biotite is replaced by chlorite, which is in good agreement with temperature estimates from chlorite in the NEAT tunnel using the empirical calibrated chlorite thermometers by Cathelineau (1988). The phase diagram points out that K-feldspar is affected by a hydration reaction and forms muscovite. But in contrast to the calculation, the observed mineral inventory with rock-forming K-feldspar does not reflect the stable assemblage, and therefore K-feldspar behaves metastable in the hydrothermal system. Laumontite launch to form at temperature of ∼150 °C under the consumption of epidote, albite and quartz that is postulated in Eq. 4.

FIG. 4.8: Predicted assemblage evolution during hydrothermal alteration, calculated along a linear exhumation path with H2O in excess.

The mineral reactions and the evolution of fissure minerals involve a transport step between dissolution and precipitation as the later processes proceed at a spatially different location in a fluid regime. Nevertheless calculations of the mineral evolution suggest that reactions took place in a fluid regime with no exotic chemical composition. Calculations by Dipple and Ferry (1992) showed that changes in major element concentrations in rocks, which is the case if we assume that elements necessary for laumontite formation are derived from the dissolution of primary minerals, do not need infiltration of chemically exotic fluids, but can instead be driven TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 154 by aqueous flow along normal temperature gradients under conditions close to local equilibrium.

4.6.4. Alpine history

Considering the Alpine history, fissure precipitates formed by fluid infiltration and the fluid rock interaction with primary minerals during uplift and exhumation of the orogenic units.

FIG. 4.9: Temperature-time path of fissure mineralization during Alpine exhumation in the Central Swiss Alps. Three time paths are given that coincidence with the path evaluated in this study. The shift in the higher temperature area can be explained by the south - north decline of peak metamorphism indicated by the solid triangles (1 = first appearance of oligoclase in the Gotthard Massif (Steck, 1976) 2 = transformation isograd of microcline/sanidine (Bambauer and Bernotat, 1982); 3 = first appearance of green biotite (Steck and Burri, 1971)). The Gibelsbach upper zeolite limit is defined by fluid inclusion measurement in fluorite proceeding the zeolite formation. Schematic illustrations showing the fissure formation in relation to the temperature-time path.

Figure 4.9 gives temperature-time paths during the phase of uplift in the Central Swiss Alps and shows the time interval at which formation of zeolites is to be likely. The formation of zeolites and apophyllite marks the last step in the Alpine fissure history.

The opening of fissures starts after Alpine peak metamorphic conditions by retrograde passage through the brittle-ductile transition (Fig. 4.9). Fissure quartz, TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 155 which is often a substrate mineral on which zeolites grow, formed during earlier fissure phases. By fluid inclusion analysis different growth generation in fissure quartzes can be attached to a temperature regime between 250 °C and 450 °C and pressure conditions between 180 MPa and 440 MPa (Mullis, 1995). Mullis (1995) linked the quartz population to the temperature-time path of the same area, generated by radiometric age data (apatite and zircon FT; Rb/Sr in biotite and muscovite). This yields the time for fissure quartz formation in the Eastern Aar Massif between 21 and 13 Ma before present, which is slightly retarded with respect to the southern Gotthard Massif (Fig. 4.9).

The textural evidence of no zeolite inclusions in fissure quartz suggests that the formation of zeolites initially starts after quartz formation was completed and therefore zeolite formation in the Eastern Aar Massif have to be assume formed later than 13 Ma before present.

Considering the zeolite locality at Gibelsbach/Fiesch (Valais, south-western Aar Massif) whereas the assemblages quartz, green fluorite and zeolites occur, fluid inclusions measurements in fluorite by Armbruster et al. (1996) yields formation temperatures above 200 °C by assuming a pressure conditions of ∼100 MPa. These implies that the zeolite formation start to form at temperatures below 200 °C and at later time (∼10 Ma; Fig. 4.9) than the formation of fluorite was finished.

From this follows that the formation of laumontite can be limited to a time range between ∼10 and 2 Ma and a temperature range between 200 and 70 °C. Nevertheless thermodynamic modeling (Fig. 4.8) indicate the formation of laumontite below 150 °C which would forward limited the formation range of the laumontite formation to a time range between 7 and 2 Ma before present.

4.7. CONCLUSION

Combining different methods, the study low-temperature fluid rock interaction leads to an evidence for zeolite (laumontite) and apophyllite generation in respect to the Alpine exhumation history: TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 156

(1) Zeolites (laumontite) followed by apophyllite formed as latest mineral in Alpine fissure. Age measurements indicate an Ar/Ar age of ∼2 Ma for apophyllite, reflecting the Pleistocene epoch.

(2) The exhumation rate of the studied area is 0.45 mm a-1 that is in the range of other known exhumation ranges in the Eastern Aar Massif. Apatite FT yield a cooling rate of 13 °C Ma-1 and a geothermal gradient of 28 °C km-1. Taken the exhumation rate and the 40Ar/39Ar age of apophyllite a minimum formation temperature and depth of laumontite of ∼70 °C and 2800 m, respectively can be determined.

(3) Considering a temperature-time path and the thermodynamic approach to estimate the formation of laumontite in the Southern Aar Granite (Aar Massif), laumontite are formed between 7 to 2 Ma before present in a temperature range of 150 to 70 °C.

(4) During growth of laumontite the Si/Al ratio, K, Na and Ba increases, whereas Ca decreases, which could be an effect of temperature drop. However the overgrowth of apophyllite do not exclude a chemical change in fluid composition during laumontite and apophyllite growth.

(5) Elements for the formation of laumontite derived during dissolution and transport of primary minerals (clinozoisite/epidote, albite and quartz) of the wall rock and no exotic fluid are necessary.

4.8. ACKNOWLEDGMENTS

We would like to thank Peter Amacher who provided high-quality mineral specimens from the Gotthard NEAT tunnel. We are grateful to Alptransit and the geologist Roger Rütti for sample supply from the tunnel. In addition a special thanks to the technicians of the Institute of Geosciences (Mineralogy – Geochemistry) University of Freiburg for the assistance in preparing and analysing samples. A special thanks deserved to the Friedrich Rinne foundation for the financial support.

TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 157

4.9. REFERENCES

Armbruster, T. and Kohler, T. (1992) Re- and dehydration of laumontite: a single- crystal X-ray study at 100 K. Neues Jahrbuch für Mineralogie, Monatsheft, 9, 385-397. Armbruster, T., Kohler, T., Meisel, T., Nägler, T.F., Götzinger, M.A. and Stalder, H.A. (1996) The zeolite, fluorite, quartz assemblage of the fissure at Gibelsbach, Fiesch (Valais, Switzerland): crystal chemistry, REE patterns, and genetic speculations. Schweizerische Mineralogische und Petrographische Mitteilungen, 76, 131-146. Bambauer, H.U. and Bernotat, W.H. (1982) The microcline/sanidine transformation isograd in metamorphic regions. I. Composition and structural state of alkali feldspars from granitoid rocks of two N-S traverses across the Aar massif and Gotthard massif. Schweizerische Mineralogische und Petrographische Mitteilungen, 62, 185-230. Belsare, M.R. (1969) A chemical study of apophyllite from Poona. Mineralogical Magazine, 37, 288-289. Berger, A., Mercolli, I. and Engi, M. (2005) Tectonic and petrographic map of the Central Lepontine Alps, 1:100 000. Schweizerische Mineralogische und Petrographische Mitteilungen, 85, 109-146. Berman, R.G. (1988) Internally-consistent thermodynamic data for minerals in the system Na2O-K2O-CaO-MgO-FeO-Fe2O3-Al2O3-SiO2-TiO2-H2O-CO2. Journal of Petrology, 29, 445-522. Bernotat, W.H. and Bambauer, H.U. (1982) The microcline/sanidine transformation isograd in metamorphic regions. II. The region of Lepontine metamorphism, Central Swiss Alps. Schweizerische Mineralogische und Petrographische Mitteilungen, 62, 231-244. Betz, V. (1981) Famous mineral localities: zeolites from Iceland and the Faeroes. The Mineralogical Record, 12, 5-26. Borchardt, R. and Emmermann, R. (1993) Vein minerals in KTB rocks. Pp 481-488 in: KTB Report 93-2 (R. Emmenmann, J. Lauterjung, and T. Umsonst, editor). Project Management of the Continental Deep Drilling of the Federal Republic of Germany in the Geological Survey of Lower Saxony. Borchardt, R., Zulauf, G., Emmermann, R., Hoefs, J. and Simon, K. (1990) Abfolge und Bildungsbedingungen von Sekundärmineralen in der KTB-Vorbohrung. Pp 76-88 in: KTB Report 90-4 (R. Emmermann, and P. Giese, editors). Projektleitung Kontinentales Tiefbohrprogramm der Bundesrepublik Deutschland im Niedersächsischen Landesamt für Bodenforschung. Bucher, K. and Frey, M. (2002) Petrogenesis of metamorphic rocks. Springer Verlag, Berlin, 341 pp. Bucher, K. and Stober, I. (2000) Hydrochemistry of water in the crystalline basement. Pp. 141-175 in: Hydrogeology of Crystalline Rocks (I. Stober, and K. Bucher, editors). Kluwer Academic Publishers, Dordrecht. TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 158

Bucher, K. and Stober, I. (2001) Does plagioclase control the composition of groundwater in the crystalline basement? Pp. 149-152 in: Water-Rock Interaction WRI-10 (R. Cidu, editor). A. A. Balkema Publishers, Lisse. Cathelineau, M. (1988) Cation site occupancy in chlorites and illites as a function of temperature. Clay Minerals, 23, 471-485. Ciesielczuk, J. and Janeczek, J. (2004) Hydrothermal alteration of the Strzelin granite, SW Poland. Neues Jahrbuch für Mineralogie, Abhandlungen, 179, 239-264. de Capitani, C. and Brown, T.H. (1987) The computation of chemical equilibrium in complex systems containing non-ideal solutions. Geochimica et Cosmochimica Acta, 51, 2639-2652. Dipple, G.M. and Ferry, J.M. (1992) Metasomatism and fluid flow in ductile fault zones. Contributions to Mineralogy and Petrology, 112, 149-164. Donelick, R.A., O'sullivan, P.B. and Ketcham, R.A. (2005) Apatite fission-track analysis. Pp 49-94 in: Low-temperature thermochronology: techniques, interpretations, and applications (P. W. Reiners, and T. A. Ehlers, editors). Reviews in Mineralogy and Geochemistry, 58, Minerlogical Society of America, Washington, D.C. Evans, B.W. (1990) Phase relations of epidote-blueschists. Lithos, 25, 3-23. Fleming, T.H., Foland, K.A. and Elliot, D.H. (1999) Apophyllite 40Ar/39Ar and Rb-Sr geochronology: Potential utility and application to the timing of secondary mineralization of the Kirkpatrick Basalt, Antarctica. Journal of Geophysical Research, 104, 20,081-020,095. Freiberger, R., Hecht, L., Cuney, M. and Morteani, G. (2001) Secondary Ca-Al silicates in plutonic rocks: implications for their cooling history. Contributions to Mineralogy and Petrology, 141, 415-429. Frey, M., Bucher, K., Frank, E. and Mullis, J. (1980) Alpine metamorphism along the Geotraverse Basel-Chiasso - a review. Eclogae Geologicae Helvetiae, 73, 527- 546. Frey, M., de Capitani, C. and Liou, J.G. (1991) A new petrogenetic grid for low-grade metabasites. Journal of Metamorphic Geology, 9, 497-509. Frey, M. and Mählmann, R.F. (1999) Alpine metamorphism of the Central Alps. Schweizerische Mineralogische und Petrographische Mitteilungen, 79, 135- 154. Frisch, W., Ménot, R.P., Neubauer, F. and von Raumer, J.F. (1990) Correlation and evolution of the Alpine basement. Schweizerische Mineralogische und Petrographische Mitteilungen, 70, 265-286. Fujimoto, K., Tanaka, H., Higuchi, T., Tomida, N., Othani, T. and Ito, H. (2001) Alteration and mass transfer inferred from the Hirabayshi GSJ drill penetrating the Nojima Fault, Japan. Island Arc, 10, 401-410. Galbraith, R.F. and Laslett, G.M. (1993) Statistical models for mixing fission track ages. Nuclear Tracks, 17, 207-214. Grünenfelder, M., Hofmänner, F. and Grögler, N. (1964) Heterogenität akzessorischer Zirkone und die petrogenetische Deutung ihrer Uran/Blei-Zerfallalter. II. TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 159

Präkambrische Zirkonbildung im Gotthardmassiv. Schweizerische Mineralogische und Petrographische Mitteilungen, 44, 543-558. Gulson, B.L. and Rutishauser, H. (1976) Granitization and U - Pb studies of zircons in the Lauterbrunnen Crystalline Complex. Geochemical Journal, 10, 13-23. Huber, W. (1948) Petrographisch-mineralogische Untersuchung im südöstl. Aarmassiv GR/UR. Schweizerische Mineralogische und Petrographische Mitteilungen, 28, 555-642. Hurford, A.J. and Green, P.F. (1983) The zeta age calibration of fission track dating. Isotope Geosciences, 1, 285-317. Jäger, E., Niggli, E. and Wenk, E. (1967) Rb-Sr-Alterbestimmungen an Glimmern der Zentralalpen. Beiträge zur Geologische Karte der Schweiz, 134, 77. Keith, T.E.C. and staples, L.W. (1985) Zeolites in Eocene basaltic pillow lavas of the Siletz river volcanics, central coast range, Oregon. Clays and Clay Minerals, 33, 135-144. Koppers, A.A.P. (2002) ArArCALC - software for 40Ar/39Ar age calculations. Computers and Geosciences, 28, 605-619. Labhart, T.P. (1977) Aarmassiv und Gotthardmassiv. Borntraeger, Berlin, 173 pp. Michalski, I. and Soom, M. (1990) The Alpine thermo-tectonic evolution of the Aar and Gotthard massifs, Central Switzerland: Fission track ages on zircon and apatite and K-Ar mica ages. Schweizerische Mineralogische und Petrographische Mitteilungen, 70, 373-387. Molzahn, M., Wörner, G., Henjes-Kunst, F. and Rocholl, A. (1999) Constraints on the Cretaceous thermal event in the Transantarctic Mountains from alteration process in Ferrar flood basalts. Global and Planetary Change, 23, 45-60. Mullis, J. (1995) Entstehung alpiner Kluftmineralien. Mitteilungen für Wissenschaft und Technik, 11, 54-64. Mullis, J. (1996) P-T-t path of quartz formation in extensional veins of the Central Alps. Schweizerische Mineralogische und Petrographische Mitteilungen, 76, 159-164. Mullis, J., Dubessy, J., Poty, B. and O'Neil, J. (1994) Fluid regimes during late stages of a continental collision: physical, chemical and stable isotope measurements of fluid inclusions in fissure quartz from a geotraverse through the Central Alps, Switzerland. Geochimica et Cosmochimica Acta, 58, 2239-2267. Neuhoff, P.S., Watt, W.S., Bird, D.K. and Pedersen, A.K. (1997) Timing and structural relations of regional zeolite zones in basalts of the East Greenland continental margin. Geology, 25, 803-806. Nordstrom, D.K., Ball, J.W., Donahoe, R.J. and Whittemore, D. (1989) Groundwater chemistry and water-rock interactions at Stripa. Geochimica et Cosmochimica Acta, 53, 1727-1740. Passaglia, E. (1970) The crystal chemistry of chabazite. American Mineralogist, 55, 1278-1301. Pouchou, G. and Pichior, F. (1991) Quantitative analysis of homogeneous or stratified microvolumes applying the model of "PAP". Pp. 31-75 in: Electron probe TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 160

quantitation (K.F.J. Heinrich, and D.E. Newbiry, editors) Plenum Press, New York. Reinecker, J., Danišík, M., Schmid, C., Glotzbach, C., Rahn, M., Frisch, W. and Spiegel, C. (2008) Tectonic control on the late stage exhumation of the Aar Massif (Switzerland): Constraints from apatite fission track and (U-Th)/He data. Tectonics, doi:10.1029/2007TC002247. Schaer, J.P., Reimer, G.M. and Wagner, G.A. (1975) Actual and ancient uplift rate in the Gotthard region, Swiss Alps: a comparison between precise levelling and fission-track apatite age. Tectonophysics, 29, 293-300. Schaltegger, U. (1990) Post-magmatic resetting of Rb-Sr whole rock ages - a study in the Central Aar Granite (Central Alps, Switzerland). Geologische Rundschau, 79, 709-724. Schaltegger, U. and Corfu, F. (1992) The age ans source of late Hercynian magmatism in the central Alps: evidence from precise U - Pb ages and initial Hf isotopes. Contributions to Mineralogy and Petrology, 111. Seelig, U., Stober, I. and Bucher, K. (2007) High-pH waters from the new, Gotthard rail base tunnel, Switzerland. Pp. 251-254 in: Water-Rock Interaction WRI-12, 1 (T. Bullen, and Y. Wang, editors). A. A. Balkema Publishers, Kunming. Seelig, U., Stober, I. and Bucher, K. (2009) Halogen geochemistry of waters from the new Gotthard rail base tunnel: a tale of water-rock interaction in a fractured basement aquifer. submitted to Geochimica et Cosmochimica Acta.

Senderov, E.E. (1973) Effect of CO2 on the stability of laumontite. Geochemistry International, 10, 139-114. Stalder, H.A., Wagner, A., Graser, S. and Stuker, P. (1998) Mineralienlexikon der Schweiz. Wepf Verlag, Basel, 579 pp. Steck, A. and Burri, G. (1971) Chemismus und Paragenese von Granaten aus Granitgneisen der Günschiefer- und Amphibolitfazies der Zentralalpen. Schweizerische Mineralogische und Petrographische Mitteilungen, 51, 534- 538. Stober, I. and Bucher, K. (2004) Fluid sinks within the earth´s crust. Geofluids, 4, 143-151. Sukheswala, R.N., Avasis, R.K. and Gangopadhyay, M. (1974) Zeolites and associated secondary minerals in the Deccan Traps of Western India. Mineralogical Magazine, 39, 658-671. Walker, G.P.L. (1959) Geology of the Reydarfjördur area, Eastern Iceland. Quarterly Journal of the Geological Society of London, 114, 367-393. Weisenberger, T. and Bucher, K. (2009) Zeolites in fissures of Granites and gneisses of the Central Alps. Submitted to Journal of Metamorphic Geology. Weisenberger, T. and Selbekk, R.S. (2008) Multi-stage zeolite facies mineralization in the Hvalfjördur area, Iceland. International Journal of Earth Sciences. DOI 10.1007/s00531-007-0296-6. Young, B., Dyer, A., Hubbard, N. and Starkey, R.E. (1991) Apophyllite and other zeolite-type minerals from the Whin Sill on the northern Pennines. Mineralogical Magazine, 55, 203-207. TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 161

Zen, E. (1961) The zeolite facies: an interpretation. American Journal of Science, 259, 401-409.

APPENDIX i

APPENDIX

I Own Contribution

Contribution of Tobias Weisenberger on paper manuscript processing presented in Chapter 2, 3 and 4.

Chapter 2:

1 1 Weisenberger T. and Bucher K. ZEOLITES IN FISSURES OF GRANITES AND GNEISSES OF

THE CENTRAL ALPS. submitted to “Journal of Metamorphic Geology” Own contribution: Idea: 50 %; Data generation: 100 %; Interpretation: 75 %; Manuscript writing: 60 %

Chapter 3:

1 1 Weisenberger T. and Bucher K. POROSITY EVOLUTION, MASS TRANSFER AND

PETROLOGICAL EVOLUTION DURING LOW TEMPERATURE WATER ROCK INTERACTION IN

GNEISSES OF THE SIMANO NAPPE – ARVIGO, VAL CALANCA, GRISONS, SWITZERLAND. shortly to be submitted to “Contributions to Mineralogy and Petrology” Own contribution: Idea: 90 %; Data generation: 100 %; Interpretation: 80 %; Manuscript writing: 90 %

Chapter 4:

1 1,2 3 3 1 Weisenberger T., Rahn M., van der Lelij R., Spikings R. and Bucher K. TIMING

AND MINERAL EVOLUTION DURING LOW-TEMPERATURE FLUID-ROCK INTERACTION ON 40 39 UPPER CRUSTAL LEVEL: AR/ AR APOPHYLLITE-(KF) DATING AND APATITE FISSION

TRACK ANALYSIS ON ALPINE FISSURES (CENTRAL ALPS/SWITZERLAND). shortly to be submitted to “Mineralogical Magazine” Own contribution: Idea: 90 %; Data generation: 40 %; Interpretation: 90 %; Manuscript writing: 90 %

1Institute of Geosciences, Mineralogy - Geochemistry, Albert-Ludwigs-University Freiburg, Albertstr. 23 b, 79104 Freiburg, Germany

2Eidgenössisches Nuklearsicherheitsinspektorat ENSI, 5232 Villigen, Switzerland

3Department of Mineralogy, Université de Genève, Rue des Maraîchers 13, 1205 Geneva, Switzerland APPENDIX ii

II Publications

Other related contrbutions by the author not included in the thesis.

Peer-Reviewed Papers

[1] SELBEKK R.S. AND WEISENBERGER T. (2005) Stellerite from the Hvalfjördur area, Iceland. Jökull 55, 49-52

[2] SPÜRGIN S., WEISENBERGER T. AND HÖRTH J. (2008) Das Leucitophyrvorkommen vom Strümpfekopf im Kaiserstuhl – eine historische und mineralogische Betrachtung. Berichte der Naturforschenden Gesellschaft zu Freiburg i. Br. 98, 221-244

[3] WEISENBERGER T. AND SELBEKK R.S. (2008) Multi-stage zeolite facies mineralization in the Hvalfjördur area, Iceland. International Journal of Earth Sciences, (DOI 10.1007/s00531-007-0296-6)

[4] WEISENBERGER T. AND SPÜRGIN S. (2009) Zeolites in alkaline rocks of the Kaiserstuhl volcanic complex, SW Germany - new micropobe investigation and their relationship to the host rock. Geolgica Belgica 12/1-2, 75-91

Talks und Poster Presentations on International Confereneces

[1] WEISENBERGER T. AND BUCHER K. (2006) Zeolites on fissures of crystalline basement rocks in the Swiss Alps. In Bowmann R.S. and Delap S.E. (eds). Zeolite `06 - 7th International Conference on the Occurrence, Properties, and Utilization of Natural Zeolites, Socorro, New Mecixo USA, 16-21 July 2006, p. 253 (Talk)

[2] SELBEKK R.S. AND WEISENBERGER T. (2006) Zeolite facies metamorphism in the Hvalfjördur area, Iceland. In Bowmann R.S. and Delap S.E. (eds). Zeolite `06 - 7th International Conference on the Occurrence, Properties, and Utilization of Natural Zeolites, Socorro, New Mecixo USA, 16-21 July 2006, p. 220 APPENDIX iii

[3] WEISENBERGER T. AND BUCHER K. (2006) Zeolites on fissures of alpine crystalline basement rocks in the Swiss Alps. Berichte der Deutschen Mineralogischen Gesellschaft, Beih. z. Eur. J. Mineral. Vol. 18, No. 1, p.153 (Talk)

[4] BUCHER K. AND WEISENBERGER T. (2006) Zeolites on fissures of crystalline basement rocks in the Swiss Alps. Geological Society of America Abstracts with Programs, Vol. 38, No. 7, p. 113

[5] SELBEKK R.S. AND WEISENBERGER T. (2007) Multi-stage zeolite facies metamorphism, Southwest Iceland. NGF Winterconference Stavanger, 8.-10. Januar 2007. NGF Abstracts and Proceedings of the Geological Society of Norway, no, 1, 90 (Poster)

[6] WEISENBERGER T. AND BUCHER K. (2007) Low-grade zeolite facies metamorphism in gneisses of the Simano nappe (Arvigo, Val Calanca, Grisons, Switzerland). Geochimica et Cosmochimica Acta 71(15) Supplement 1, A1100 (Talk)

[7] WEISENBERGER T. AND BUCHER K. (2007) Porosity increase during low- temperature metamorphism in gneisses of the Simano nappe (Arvigo, Val Calanca). 5th Swiss Geoscience Meeting, 17-18.11. 2007, Geneva, Switzerland, 104-105 (Poster)

[8] WEISENBERGER T. AND BUCHER K. (2008) Porosity evolution and mass transfer during low-grade metamorphism in crystalline rocks of the upper continental crust. 33rd IGC International Geological Congress, 06. - 14.08.2008 Oslo, MPN03710L (Talk)

[9] WEISENBERGER T. AND BUCHER K. (2008) Ca-Al Silicate Formation During Low- grade Metamorphism in the Upper Continental Crust. 86th Annual Meeting of the German Mineralogical Society – DMG 14th -17th September 2008 Berlin, S18T06 (Talk)

[10] SPÜRGIN S. AND WEISENBERGER T. (2008) Faujasite Growth During Palagonitisation of Mg-rich Sideromelane: an Example from the Kaiserstuhl Volcanic Complex, SW Germany. 86th Annual Meeting of the German Mineralogical Society – DMG 14th – 17th September 2008 Berlin, S18P07 (Poster) APPENDIX iv

Popular scientific papers

[1] WEISENBERGER T. (2005) Zeolithe in Island. Island; Zeitschrift der Deutsch- Isländischen Gesellschaft e.V. Köln und der Gesellschaft der Freunde Island

e.V. HAMBURG, ISSUE 1, APRIL 2005, 40-45

[2] WEISENBERGER T. AND SELBEKK R. S. (2006) Die Zeolith-Fundstelle Hvalfjordur, Island. Mineralien Welt 17, Heft 1, 50-56

[3] WEISENBERGER T., SPÜRGIN S. AND SELBEKK R. S. (2008) Die Fundstelle Helgustadir (Island): Geologie, Mineralogie und die bedeutende Geschichte des Isländischen Doppelspats für die Wissenschaft. Aufschluss 1, 53-63

Excursion Guide Books

[1] WEISENBERGER T. AND SPÜRGIN S. (2008) Secondary minerals in the limburgites. In Keller J.: Tertiary Rhinegraben volcanism: Kaiserstuhl and Hegau. 9th International Kimberlite Conference, Frankfurt/Main, field trip, 24-25

APPENDIX v

III Curriculum Vitae

Tobias Weisenberger

Date/ Place of birth: 04.12.1979 in Emmendingen/ Baden-Württemberg/ Germany since 2006 Dissertation at the Institute of Geosciences, Mineralogy - Geochemistry, Albert-Ludwigs-University of Freiburg. Subject of the thesis: “Zeolites in fissures of crystalline basement rocks”. Adviser Prof. Dr. Kurt Bucher and Prof. Dr. Reto Gieré

2000 – 2005 Geology (Diploma) study at the Albert-Ludwigs University of Freiburg. Subject of the thesis: "Zeolite facies mineralisation in the Hvalfjördur area, Iceland". Adviser Dr. Rune Selbekk

1990 – 1999 Secondary school, Gymnasium Kenzingen, Germany

1986 – 1990 Primary school Endingen, Germany