zeolites in fissures of crystalline basement rocks

Transcription

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 extraframework 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
40
Ar/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.
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
2.
1
7
1.6.
ZEOLITE STABILITY
11
1.7.
GEOLOGICAL SETTING
12
1.8.
REFERENCES
15
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
2.5.3.5.
2.6.
3.
VII
Stilbite/Stellerite
54
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
POROSITY
EVOLUTION,
PETROLOGICAL
TEMPERATURE
MASS
EVOLUTION
WATER-ROCK
TRANSFER
AND
DURING
LOW
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
3.7.1.
Mineral reactions
105
105
TABLE OF CONTENTS
4.
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
TIMING AND MINERAL EVOLUTION DURING LOWTEMPERATURE FLUID-ROCK INTERACTION ON UPPER
CRUSTAL LEVEL:
40
Ar/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
UPPER CRUSTAL LEVEL:
40
AR/39AR 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
granites and gneisses of the Central
Alps
Chapter 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
Chapter 4: Timing and mineral
evolution during low-temperature
fluid-rock interaction on upper
40
39
crustal level: Ar/ Ar apophyllite(KF) dating and apatite fission track
analysis on Alpine fissures (Central
Alps/Switzerland)
• general overview and compilation
of all zeolite localities in the
Central Swiss Alps
• detailed study of a zeolite locality
to assess the formation of Ca-Al
silicates during uplift
• detailed study of a zeolite locality
to assess the formation of Ca-Al
silicates during exhumation
• what zeolite species occur in
crystalline basement rocks?
• temporal evolution of secondary
minerals
• timing of zeolite formation
• spatial distribution of zeolites
• mineral reactions during waterrock interaction
• chemical change during formation
of secondary minerals
• zeolite forming reactions
• mass transfer during water-rock
interaction
• relation of zeolite formation with
respect to the Alpine history
• what factors control the zeolite
formation?
• porosity evolution during waterrock interaction
• estimation of laumontite forming
temperature
• thermodynamic phase-diagram
calculation approaching physical
and chemical conditions
• fluid-composition in relation to
the formation of zeolites
• chemical characterization of
zeolites
• depth of zeolite formation
• 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 extraframework 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 (extraframework 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):
(A+z)y/z(B+3)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 extraframework 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, Caheulandite 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 crosssection 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, preVariscan 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 polymetamorphic 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 nonmetamorphosed 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 zeoliteabsent 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).
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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
2. ZEOLITES IN FISSURES OF GRANITES AND
GNEISSES OF THE CENTRAL ALPS
20
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 GotthardNEAT 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.
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 Cazeolites.
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
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).
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).
25
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
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 polymetamorphic 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 nonmetamorphosed 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 lowgrade 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).
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
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
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
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
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
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 &
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
1
Locality
Bäregg, Oberaar,
Ca
Sco
BE
x
Lmt
Stb
Heu
x
x
Cha
Grimsel
AFMb
Fc
Qtz, Kfs,
x
Hem, Cal,
Referencesd
Stalder, 1964; Stalder et al.,
1998; SNHMB; °
Chl, Py
2
Lötschberg tunnel
BE
3
4
Bächistock
Arvigo/Val
GL
GR
x
x
x
x
x
Qtz, Cal
x
Qtz, Kfs,
x
xxx
x
x
CalancaEpi
Ttn, Act,
SNHMB
Niggli et al., 1940; SNHMB
Stalder et al., 1998; Wagner,
2000a, b; SNHMB; °
Ep, Prh,
Chl, Ap
5
BergellPhil *
GR
6
Cuolm da Vi
GR
7
Davos
GR
8
Drumtobel/Sedrun
GR
9
Misox
GR
10
OberalpstockTho
GR
11
12
Sedrun
Sella/Gotthard
GR
GR
13
Stgegia, Medel
GR
14
Val Casatscha
GR
15
16
Val Cristallina
Val Giuv
GR
GR
x
x
x
x
x
Qtz, Kfs,
Grs, Prh,
x
Hirschi, 1925; Stalder et al.,
1998; SNHMB
x
x
Kenngott, 1866; Huber, 1948;
x
x
Stalder et al., 1998; SNHMB
xxx
Stalder et al., 1998; SNHMB;
°
Ttn, Cal
SNHMB
x
x
x
x
x
x
Qtz, Kfs,
Cal, Act,
Ttn
x
Chl
x
SNHMB
Qtz
x
Niggli et al., 1940; Stalder et
Qtz, Hem,
x
x
Kenngott, 1866; SNHMB
Kenngott, 1866; SNHMB; °
al., 1998
x
x
x
Chl
x
Qtz
x
SNHMB
x
x
Cal
x
Stalder et al., 1998; Huber,
x
x
x
Cal
Qtz, Kfs,
x
xxx
1943; °
x
x
x
Ap
SNHMB
Kenngott, 1866; Sigrist, 1947;
Stalder et al., 1998; SNHMB;
°
17
Val Maighels*
GR
18
19
Val Medel
Val Mila
GR
GR
20
Val Muretto/Bergell
GR
21
Val Nalps
GR
22
Val Punteglias
GR
23
Val Russein
GR
24
Val Strem
GR
x
Chl
x
Qtz
x
x
Huber, 1943; Stalder et al.,
1998; SNHMB
x
x
Huber, 1943; SNHMB
Niggli et al., 1940; Sigrist,
1947; SNHMB
x
x
x
Stalder et al., 1998
x
Prh, Tur
x
SNHMB
x
Qtz, Chl
x
Huber, 1943; Sigrist, 1947;
x
x
Qtz, Ep, Chl
xx
Huber, 1943; Stalder et al.,
x
x
Qtz, Chl
xxx
1998; SNHMB
x
x
Kenngott, 1866; Niggli et al.,
1940; Huber, 1948; Stalder et
al., 1998; SNHMB; °
34
continue Table 2.1
25
BiascaEpi
TI
x
Kfs, Ttn,
x
x
Wagner et al., 1972; SNHMB
Prh
Chl
x
SNHMB
Qtz Chl,
xx
Kenngott, 1866; SNHMB
Ep. Chl
26
Camperio/Passo del
TI
x
Qtz, Hem,
Lucomango
Cal, Chl,
27
Domodossola
TI
28
La Fibbia, Gotthard
TI
x
x
Hem
29
Lodrino
TI
x
30
Mt Ceneri
TI
x
31
Pizzo Lucendro
TI
32
Val Baveno
TI
33
Val Canaria*
TI
34
Val Maggia*
TI
35
Val Vergeletto, road
TI
x
x
Qtz, Chl
x
x
Py
x
Toroni, 1984; Stalder et al.,
x
Qtz Kfs,
x
1998
Hem, Ms
x
x
x
x
x
x
x
SNHMB
x
x
Kenngott, 1866; Stalder et al.,
x
x
1998; SNHMB
Simonetti, 1971; Stalder et al.,
tunnel
1998; SNHMB
36
Andermatt
UR
x
37
Brunnital
UR
x
38
Chrützlistock
UR
x
x
x
SNHMB
Cal
x
SNHMB
Qtz
xxx
1940; Sigrist, 1947; Weibel,
1963; Stalder et al., 1998;
SNHMB; °
39
Etzlital
UR
x
x
x
x
xxx
1940; Sigrist, 1947; Stalder et
al., 1998; SNHMB; °
40
41
Fedenstock
Fellilücke
UR
UR
x
x
x
Qtz, Chl
Qtz, Flt
x
xxx
42
Fellital
UR
x
x
x
Qtz, Chl
xxx
43
Göscheneralp
UR
x
x
x
Qtz, Chl
x
44
Gotthard road
UR
x
x
x
Qtz, Kfs,
xxx
SNHMB
1947; SNHMB; °
1947; SNHMB; °
x
tunnel
Cal, Ep,
1998; SNHMB
Stalder et al., 1980, 1998;
SNHMB; °
Prh, Chl, Py
45
Griesserental
UR
46
47
Maderanertal
NEAT, Amsteg -
UR
UR
x
x
x
x
x
x
Sedrun
Kfs
x
SNHMB
Qtz, Chl, Py
Qtz, Kfs,
x
xxx
SNHMB
°
xxx
1940; Sigrist, 1947; Huber,
Cal, Hem,
Act, Ep,
Chl, Apo,
Py, Anh
48
Piz Giuv
UR
x
49
Riental
UR
x
x
50
Schattig Wichel &
UR
x
x
1948; SNHMB; °
Qtz
xxx
Qtz, Kfs,
xxx
1940; SNHMB; °
x
x
Piz Giuv
x
Ap, Ttn,
Act, Chl,
1940; Sigrist, 1947; Huber,
Ep, Prh
51
Schijenstock
UR
x
x
Qtz, Flt,
SNHMB; °
xx
Hem
52
Tiefengletscher
UR
x
Qtz, Gn
al., 1998; °
x
SNHMB
35
continue Table 2.1
53
Arolla
VS
x
54
Binntal Meso, Nat, Phil *
VS
x
x
x
x
x
x
Qtz
x
Qtz, Kfs,
x
Kenngott, 1866; Keusen &
Hem, Chl,
Bürki, 1969; Stalder et al.,
Ttn,
55
Fieschergletscher
VS
56
Furka tunnel
VS
x
1998; SNHMB
x
Qtz, Kfs,
Ep, Ap
x
x
Qtz, Kfs,
x
SNHMB
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;
58
Gornergletscher
VS
59
Gredetschtal/Brig
VS
x
x
Qtz, Kfs,
x
xx
Chl, Ttn
60
61
Grosses Sidelhorn
Lax
VS
VS
62
Lötschental*
VS
x
x
x
x
x
al., 1998; SNHMB
x
Qtz, Chl
Qtz, Chl
x
x
x
Qtz, Kfs,
xx
SNHMB
1998; SNHMB
x
Cal, Act
63
Martiny
VS
64
65
Massaschlucht
Mättital
VS
VS
66
Nuffenenpass
VS
67
Pollux/ZermattNat
VS
68
Simplontunnel
VS
x
et al., 1998; SNHMB
Qtz
x
SNHMB
x
Kfs, Cal
Qtz, Cal
x
x
SNHMB
SNHMB
x
Qtz
x
x
Fellenberg von, 1893; Stalder
x
Chl
x
SNHMB
xx
x
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).
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 SW 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).
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
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
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
A8*
Kfs-Prh-Ep-Chl-Sco-Lmt
Fr
4
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
A8332*
Kfs-Ap-Chl-Cha
needles, Gotthard NEAT
on Qtz, Gotthard NEAT
Sco needles and Lmt on top of Kfs, Ep, Prh and Chl in a fissure hosted
in an orthogneiss, Arvigo/Val Calanca
dense mats of crystals, up to 4 mm in size, Val Strem
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
Fr
4
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
Fr
24
euhedral crystals of Sco, Heu and Stb, Schattig Wichel,
I*
hosted in an orthogneiss, Arvigo/Val Calanca
Arvigo13 Sco
*
Sco needles up to 4 cm in length located in fissure, Arvigo/Val
Calanca
Arvigo2* Lmt
B12*
Sco-Heu-Stb
B981*
Chl-Lmt
SNHMB 44
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
green Lmt crystals up to 1 cm in length, green colour of Lmt appear
because of Chl inclusions, Gotthard road tunnel
associated with early formed Heu crystals on top of a highly porous
matrix, Drumtobel/Sedrun
euhedral Stb crystals, up to 1 cm in size associated with Heu and Qtz,
Fi1*
Qtz-Heu-Stb
Fr
57
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
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
R1*
Qtz-Stb
Fr
49
TW01*
Lmt
P.A.
47
Gibelsbach/Fiesch
NEAT
gneiss, Gotthard NEAT
light brounish Stb, forming radial groups of 1 cm in diameter, hosted
in quartz lenses in paragneiesses, Riental
Lmt on Qtz, Gotthard NEAT
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
TW01.3* Lmt-Stb
P.A.
47
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
P.A.
47
Lmt associated with early formed Chl and Qtz in a fissure of Gotthard
Gotthard NEAT
Lmt associated with Stb as early phase in a fissure of the Gotthard
NEAT
TW11.1* Qtz-Chl-Lmt
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,
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
35370*
Qtz-Cc-Lmt
Stb
SNHMB 2
SNHMB 44
Lötschberg exploring tunnel
light brownish Stb crystal, up to 3 mm, grown on top of fine Qtz
Val Strem
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
35795*
35843*
Lmt
Sco-Stb
SNHMB 44
SNHMB 44
Lmt from the Gotthard road tunnel
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,
tunnel
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
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 Cazeolite 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
± titanite
epidote
calcite
prehnite
chlorite
calcite
zeolites
- scolecite
± actinolite
fluorite
- laumontite
± ilmenite
± hematite
- heulandite
± anhydrite
apophyllite
- 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,
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-chloritecalcite-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-muscovitehematite-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
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.
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
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,
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 NESW, 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
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: PermianCarboniferous, 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
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
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.
49
Table 2.4. Representative zeolite analysis
Sample no.
Analysis no.
wt.%
A8332
2
Cha
Chaba1
2
Cha
Chaba2
3
Cha
Chaba2
4
Cha
A5203
1
Heu
TW20
6
Heu
TW20
7
Heu
A8.1
12
Heu
Fi1
4
Heu
TW34
6
Heu
SiO2
Al2O3
CaO
SrO
BaO
Na2O
K2O
H2O
52.76
19.47
8.35
1.88
-0.17
3.75
13.58*
53.59
19.24
8.52
1.58
0.08
0.08
3.72
13.14*
54.37
19.51
8.00
2.18
0.01
0.20
3.70
12.02*
53.80
19.53
8.48
2.29
0.00
0.15
2.73
13.00*
54.64
16.51
5.47
4.35
0.14
1.53
17.30*
60.74
15.99
6.77
2.40
0.03
1.36
12.67*
59.10
15.08
6.36
2.35
0.03
1.36
15.68*
56.30
16.56
6.87
0.92
0.38
2.47
16.47*
56.67
15.60
6.07
2.31
0.00
0.11
2.28
16.93*
57.76
15.61
5.99
3.08
0.28
0.07
1.40
15.75*
Totala
100.00
100.00
100.00
100.00
100.00
100.00
100.00
100.00
100.00
100.00
Si
Al
Ca
Sr
Ba
Na
K
O
H2O
8.300
3.610
1.407
0.171
0.052
0.753
24
7.125
8.364
3.539
1.425
0.143
0.005
0.024
0.741
24
6.840
8.385
3.546
1.322
0.195
0.001
0.060
0.728
24
6.183
8.360
3.577
1.412
0.206
0.000
0.045
0.541
24
6.737
26.573
9.463
2.850
1.227
0.132
0.949
72
28.074
27.446
8.517
3.276
0.629
0.027
0.786
72
19.088
27.636
8.312
3.184
0.638
0.024
0.811
72
24.452
26.713
9.259
3.491
0.253
0.350
1.496
72
26.073
27.139
8.805
3.115
0.641
0.000
0.102
1.393
72
27.041
27.276
8.688
3.031
0.843
0.052
0.064
0.843
72
24.807
E%b
Si/(Si+Al)
Ca/(Ca+Na+K+Sr)
Na/(Ca+Na+K+Sr)
K/(Ca+Na+K+Sr)
Sr/(Ca+Na+K+Sr)
-8.89
0.70
0.59
0.02
0.32
0.07
-9.92
0.70
0.61
0.01
0.32
0.06
-7.23
0.70
0.57
0.03
0.32
0.08
-6.55
0.70
0.64
0.02
0.25
0.09
2.47
0.74
0.55
0.03
0.18
0.24
-1.50
0.76
0.69
0.01
0.17
0.13
-2.17
0.77
0.68
0.01
0.17
0.14
-0.81
0.74
0.62
0.06
0.27
0.05
-2.40
0.76
0.59
0.02
0.27
0.12
-1.60
0.76
0.63
0.01
0.18
0.18
DT
5
Heu
57.53
15.08
5.96
1.95
0.41
0.26
1.97
16.81*
A8.2
3
Heu
57.66
15.96
5.98
3.72
0.50
0.08
1.51
14.532*
A8
2
Lmt
52.23
21.89
12.20
0.03
0.00
0.04
0.05
15.24
TW11.1
6
Lmt
53.00
20.87
11.35
0.12
0.12
0.33
14.17*
TW03
1
Lmt
53.18
20.64
11.28
0.03
0.08
0.33
14.457*
TW01.2
8
Lmt
52.33
21.46
11.97
0.05
0.04
0.05
14.08*
35795
3
Lmt
51.49
21.63
12.02
0.24
0.08
0.05
14.49*
NL4b
9
Lmt
51.95
20.64
11.06
0.27
0.00
0.16
0.46
15.44*
Arviog12I
10
Lmt
52.06
21.48
11.65
0.17
0.00
0.00
0.32
14.19*
Arvigo2
7
Lmt
52.43
21.15
11.82
0.00
0.00
0.09
0.12
14.37*
Totala
100.00
100.00
101.69
100.00
100.00
100.00
100.00
100.00
100.00
100.00
Si
Al
Ca
Sr
Ba
Na
K
O
H2O
E%b
27.442
8.478
3.046
0.539
0.077
0.240
1.199
72
26.743
-3.57
27.062
8.827
3.005
1.013
0.092
0.068
0.905
72
22.747
-4.80
16.036
7.923
4.014
0.005
0.000
0.023
0.019
48
13.884
-1.95
16.362
7.594
3.754
0.021
0.072
0.130
48
14.591
-2.06
16.443
7.521
3.735
0.005
0.050
0.129
48
14.909
-1.81
16.149
7.806
3.957
0.008
0.023
0.021
48
14.486
-2.11
16.013
7.928
4.005
0.043
0.048
0.020
48
15.029
-2.90
16.313
7.640
3.721
0.049
0.000
0.099
0.185
48
16.170
-2.36
16.124
7.841
3.866
0.031
0.000
0.000
0.126
48
14.658
-0.99
16.228
7.713
3.919
0.000
0.000
0.053
0.047
48
14.837
-2.85
0.76
0.61
0.05
0.24
0.11
0.75
0.60
0.01
0.18
0.20
0.67
0.99
0.01
0.00
0.00
0.68
0.94
0.02
0.03
0.01
0.69
0.95
0.01
0.03
0.00
0.67
0.99
0.01
0.01
0.00
0.67
0.97
0.01
0.00
0.01
0.68
0.92
0.02
0.05
0.01
0.67
0.96
0.00
0.03
0.01
0.68
0.98
0.01
0.01
0.00
Sample no.
Analysis no.
wt.%
SiO2
Al2O3
CaO
SrO
BaO
Na2O
K2O
H2O
Si/(Si+Al)
Ca/(Ca+Na+K+Sr)
Na/(Ca+Na+K+Sr)
K/(Ca+Na+K+Sr)
Sr/(Ca+Na+K+Sr)
50
continue Table 2.4
Sample no.
Analysis no.
wt.%
SiO2
Al2O3
CaO
SrO
BaO
Na2O
K2O
H2O
TW20
2
Sco
45.72
24.86
14.05
0.02
0.04
0.01
15.25*
35843
3
Sco
44.64
24.44
13.63
0.04
0.09
0.02
17.09*
A8
3
Sco
45.67
24.70
13.90
0.00
0.22
0.00
15.45*
A4
1
Sco
45.39
25.14
13.69
0.10
0.00
0.09
0.00
14.01
Arvigo1
6
Sco
45.64
24.65
14.14
0.00
0.00
0.10
0.00
15.38*
Arvigo13
3
Sco
46.02
24.96
14.07
0.04
0.00
0.14
0.02
14.65*
Arvigo13
10
Sco
46.20
25.14
13.99
0.00
0.01
0.10
0.00
14.55*
TW02
2
Stb
58.12
15.21
7.75
0.06
0.41
0.07
18.37*
TW01.2
4
Stb
63.07
17.09
8.72
0.00
0.91
0.02
10.18*
35370
1
Stell
59.13
14.12
7.87
0.00
0.02
0.04
18.80*
Totala
100.00
100.00
100.00
98.43
100.00
100.00
100.00
100.00
100.00
100.00
Si
Al
Ca
Sr
Ba
Na
K
O
H2O
24.298
15.569
8.000
0.006
0.044
0.005
80
27.035
24.249
15.647
7.933
0.013
0.095
0.014
80
30.966
24.330
15.507
7.932
0.000
0.228
0.000
80
27.451
24.200
15.797
7.819
0.032
0.000
0.091
0.000
80
27.700
24.312
15.476
8.070
0.000
0.000
0.103
0.000
80
27.325
24.296
15.532
7.961
0.013
0.000
0.139
0.012
80
25.798
24.325
15.599
7.890
0.000
0.001
0.104
0.003
80
25.550
27.546
8.495
3.938
0.018
0.378
0.041
72
29.035
27.255
8.705
4.035
0.000
0.763
0.013
72
14.671
28.062
7.898
4.002
0.000
0.018
0.024
72
29.760
-3.08
0.61
0.99
0.01
0.00
0.00
-2.30
0.61
0.98
0.01
0.00
0.00
-3.64
0.61
0.97
0.03
0.00
0.00
0.01
0.61
0.98
0.01
0.00
0.00
-4.73
0.61
0.99
0.01
0.00
0.00
-3.66
0.61
0.98
0.02
0.00
0.00
-1.82
0.61
0.99
0.01
0.00
0.00
2.00
0.76
0.90
0.09
0.01
0.00
-1.81
0.76
0.84
0.16
0.00
0.00
-1.84
0.78
0.99
0.00
0.01
0.00
35370
5
Stb
58.11
14.14
7.74
0.00
0.27
0.03
19.71*
35843
7
Stb
61.46
15.64
8.52
0.00
0.35
0.06
13.90*
35843
10
Stb
61.62
15.54
8.52
0.00
0.23
0.07
13.91*
R1
1
Stb
55.94
15.37
7.23
0.04
0.00
0.36
1.50
20.64
R1
2
Stb
56.91
15.36
7.62
0.00
0.02
0.11
0.95
20.64
57.40
17.21
8.65
0.17
0.00
0.30
0.87
15.34*
61.36
15.99
7.41
0.00
0.03
0.76
1.26
13.17*
A4
5
Stb
58.15
15.14
7.93
0.03
0.06
0.56
0.00
18.06*
56.17
15.04
7.72
0.09
0.00
0.52
0.04
20.36*
A6
7
Stb
61.80
17.11
8.55
0.06
0.08
0.81
0.05
11.46*
Totala
100.00
100.00
100.00
100.00
100.00
100.00
100.00
100.00
100.00
100.00
Si
Al
Ca
Sr
Ba
Na
K
O
H2O
27.932
8.010
3.986
0.000
0.252
0.018
72
31.598
27.627
8.286
4.103
0.000
0.305
0.034
72
20.848
27.691
8.230
4.102
0.000
0.200
0.040
72
20.854
27.185
8.803
3.765
0.011
0.000
0.339
0.930
72
33.454
27.342
8.698
3.923
0.000
0.004
0.102
0.582
72
33.074
26.579
9.392
4.291
0.046
0.000
0.269
0.514
72
23.691
27.527
8.454
3.562
0.000
0.005
0.661
0.721
72
19.706
27.509
8.443
4.017
0.008
0.011
0.509
0.001
72
28.497
27.353
8.630
4.030
0.025
0.001
0.487
0.023
72
33.080
27.133
8.854
4.022
0.014
0.013
0.692
0.029
72
16.789
-2.81
0.78
0.94
0.06
0.00
0.00
-3.05
0.77
0.92
0.07
0.01
0.00
-2.54
0.77
0.94
0.05
0.01
0.00
-0.52
0.76
0.75
0.07
0.18
0.00
1.88
0.76
0.85
0.02
0.13
0.00
-0.98
0.74
0.84
0.05
0.10
0.01
-0.88
0.77
0.72
0.13
0.15
0.00
-1.61
0.77
0.89
0.11
0.00
0.00
-0.14
0.76
0.88
0.11
0.01
0.01
0.39
0.75
0.85
0.15
0.01
0.00
E%b
Si/(Si+Al)
Ca/(Ca+Na+K+Sr)
Na/(Ca+Na+K+Sr)
K/(Ca+Na+K+Sr)
Sr/(Ca+Na+K+Sr)
Sample no.
Analysis no.
wt.%
SiO2
Al2O3
CaO
SrO
BaO
Na2O
K2O
H2O
E%b
Si/(Si+Al)
Ca/(Ca+Na+K+Sr)
Na/(Ca+Na+K+Sr)
K/(Ca+Na+K+Sr)
Sr/(Ca+Na+K+Sr)
Fi1
7
Stb
DT
9
Stb
*
H2O calculated by difference. aTotals include traces of Mg, Ti, 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
A6
4
Stb
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
52
of Sr and K (Fig. 2.9; Table 2.4). Ca occupies on an average 62 mole% of all extraframework 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+.
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
group are distinguished by a Ca2+ + 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
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).
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
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 zeoliteabsent 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.
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)
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+.
plagioclase + H2O ⇒ albite + Ca2+ + 2 AlO2- + 2 SiO2,aq + H2O
(2)
Na4CaAl6Si14O40 + H2O ⇒ Na4Al4Si12O32 + Ca2+ + 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
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):
Ca2+ + 2 AlO2- + 4 SiO2,aq + 4 H2O ⇒ laumontite
(3)
Ca2+ + 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
Na4CaAl6Si14O40 + 4 H2O + 2 SiO2,aq ⇒ CaAl2Si4O12 •4 H2O + Na4Al4Si12O32
(4)
plagioclase + 7 H2O + 5 SiO2,aq ⇒ stilbite + albite
60
(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.
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 albite + calcite + SiO2,aq + 7 H2O ⇒ stilbite + 2 Na+ + CO32-
(7)
2 NaAlSi3O8 + CaCO3 + SiO2,aq + 7 H2O ⇒ CaAl2Si7O18 •7 H2O + 2 Na+ + CO32-
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
high Na+ and CO32- 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
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.
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.
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,
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 CO2rich 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
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 hightemperature Ca-zeolite wairakite.
Fig. 2.16: Activity diagram for the system Ca-Al-Si-O-H depicting mineral stability as a function of
Log aSiO2,aq versus Log aCa2+/a2H 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
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 zeolitebearing fissures in granites and gneisses: (1) in the central and eastern part of the Aarand 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.
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 MineralogicalGeochemical 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
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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.
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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
Steck, A. & Burri, G., 1971. Chemismus und Paragenese von Granaten aus
Granitgneisen der Günschiefer- und Amphibolitfazies der Zentralalpen.
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, 143151.
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.
Utada, M., 2001b. Zeolites in hydrothermal altered rocks. In: Natural Zeolites:
occurrence, properties; applications (eds Bish, D. L. & Ming, D. W.) Reviews
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Weibel, M., 1963. Chabasit vom Chrützlistock. Schweizerische Mineralogische und
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401-409.
LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO
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
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;
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
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.
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).
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.
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-PrhSco.
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
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 extraframework cations (Ca2+, Sr2+, Na+, and K+) is within 10% of the framework charge
(Al3+).
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
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.
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
49.2
53.8
52.9
1.5
1.4
2.4
K-feldspar
biotite
17.3
0.0
13.5
8.7
14.0
10.3
muscovite
20.4
17.6
14.5
chlorite
10.4
3.6
4.5
1.3
1.5
1.5
2179
1687
plagioclase/albite
quartz
others
(apatite, calcite, epidote, titanite
ilmenite, laumontite, prehnite)
counted points
2103
3.6.1.1. Unaltered rock
The unaltered gneisses (G, Fig. 3.2) in Arvigo are represented by a suite of biotitegneisses 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, biotitehornblende, 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.
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 Kfeldspar 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.
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.
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
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
+
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.
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 twodimensional 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
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.
94
Table 3.3: Representative chlorite, biotite and muscovite composition and forming temperature for
chlorite.
sample
no.
A8.1
A8.1
A8.2
A10
A10
Arvigo
Arvigo
Arvigo
Arvigo
Arvigo
Arvigo
4
8I
8I
8II
12.1
12.3
8
9
1
21
24
1
13
19
36
14
14
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
mineral
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
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
382
383
376
355
347
333
309
333
301
386
349
based on 20 oxygens
T (°C)*
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
no.
mineral
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
5
6
13
10
7
5
6
13
10
7
musco-
musco-
musco-
musco-
vite
vite
vite
vite
biotite
biotite
biotite
biotite
biotite
biotite
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
TiO2
3.06
2.98
2.43
0.52
0.83
AlIV
2.590
2.617
2.630
1.395
1.584
Al2O3
18.26
18.72
18.85
29.96
31.63
AlVI
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
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 Kfeldspar 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 Kfeldspar is formed during chloritization of biotite. Chemical composition of rock
forming K-feldspar ranges from Or85-94 and Ab15-06, respectively whereas the anorthite
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
no.
Arvigo
12.1
Arvigo
12.1
Arvigo
12.2
Arvigo
12.3
A8.1
A10
A4.1
8
20
26
2
15
18
9
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
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
mineral
based on 8 oxygens
* 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 cogenetic growth of epidote and prehnite. Epidote forms green to dark-green sheaf like
97
aggregates up to 20 mm in length, but common size ranges between 5 and 10 mm.
The chemical composition of epidote (ideal composition: Ca2(Fe3+,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
no.
mineral
Arvigo1
Arvigo1
Arvigo1
Arvigo1
Arvigo1
Arvigo1
Arvigo1
13
14
18
19
16
17
20
Arvigo1
21
epidote
epidote
epidote
epidote
prehnite
prehnite
prehnite
prehnite
42.54
SiO2
37.63
37.45
37.35
37.77
42.52
42.44
42.22
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
3.66
12.72
12.24
12.64
10.84
3.81
3.98
4.05
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
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
based on 8 cations and 12.5 oxygens
based on 7 cations and 11 oxygens
* totals includes traces of BaO
3.6.2.7. Prehnite
Prehnite forms colorless to pale green fan-shaped radiating aggregates, so-called bowtie 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
limited to their major elements (ideal composition: Ca2(Fe3+,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.
98
3.9). The total iron content (as Fe2O3) varies from 0.0 to 9.5 wt. The content of Fe3+
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
no.
A 10
A 10
A 10
2
3
4
A 10
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
Total*
0.02
95.27
0.05
95.70
0.02
95.24
0.00
95.32
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
based on 7 cations and 11 oxygens
* 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.
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).
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 extraframework 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
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
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
Al2O 3
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
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
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
mineral
formula unit composition
H2O
* totals includes traces of FeO, MgO, MnOTiO 2. 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))
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 %.
1 oligoclase + H2O => 4.19 albite + 0.81 Ca2+ + 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
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 massbalance 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
no.
Arvigo 12
Arvigo 12
Arvigo 12
Arvigo 12
Arvigo 12
Arvigo 12
Arvigo 12
Arvigo 12
A I*
A II
A III
A IV
AV
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
TiO2
0.66
0.62
0.57
0.57
0.57
0.55
0.54
0.58
Al2O 3
21.65
21.29
22.67
22.94
22.94
22.42
22.90
22.96
Fe2O 3
3.74
3.96
3.71
3.75
3.77
3.96
3.70
3.93
MnO
MgO
0.06
1.05
0.07
1.14
0.06
1.08
0.06
1.08
0.06
1.07
0.07
1.07
0.07
1.03
0.07
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
tot
ppm
ppm
ppm
ppm
ppm
ppm
ppm
56.02
ppm
V
59
53
51
50
52
52
53
57
Cr
31
25
26
24
25
27
25
27
Ni
Cu
18
10
15
<5
15
<5
19
8
13
<5
15
1
15
5
17
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
3
(g/cm )
density
2.884
*includes fissure minerals
3
(g/cm )
2.797
3
(g/cm )
2.849
3
(g/cm )
2.848
3
(g/cm )
2.850
3
(g/cm )
2.825
3
(g/cm )
2.832
3
(g/cm )
2.874
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; (MfiMoi)/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
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 massbalance 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+.
1.8 biotite + 4.6 H2O + 1.3 Mg2+ + 0.6 Ca2+ + 0.4 H+ => 1.0 chlorite + 1.1 K-feldspar
+ 0.6 titanite + 0.2 Fe2+ + 0.2 SiO2 + 2.1 K+
(4)
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2.05 biotite + 3.6 H2O + 0.9 Mg2+ + 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 +
0.35 Fe2+ + 0.82 K+ + 0.04 SiO2 + 0.25 Al3+ + 0.35 H2O
(6)
1 biotite + 0.11 H2O + 0.45 Mg2+ + 3.58 H+ => 0.49 chlorite + 0.77 K-feldspar + 0.39
ilmenite + 0.07 Fe2+ + 1.03 K+ + 0.36 SiO2 + 0.19 Al3+
(7)
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Considering the chloritization reaction of biotite (Eq. 6, 7), the reaction takes place in
the presence of fluid (H2O) and Mg2+ and whereas SiO2, K+, Fe2+ and Al3+ 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
of oligoclase occurs with coeval formation of albite. Ca2+, Al3+, 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-Alsilicates. 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,
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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 Kfeldspar 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
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.
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
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
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 lowgrade 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
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
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 zeolitefacies 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
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-SiO2H2O-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
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-Al2O3SiO2-H2O-CO2 system for different Ca/Al ratios.
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.
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
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).
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 Kfeldspar 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
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
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
release of K+, Fe2+, Al3+ 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.
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.
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
124
samples. A special thanks deserved to the Friedrich Rinne foundation for the financial
support.
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minerals from the Whin Sill on the northern Pennines. Mineral Mag 55: 203207
Zen E (1961) The zeolite facies: an interpretation. Am J Sci 259: 401-409
131
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)
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 reprecipitation 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.,
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
40
Ar/39Ar method on secondary mineralization
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 lowtemperature 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.
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
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
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 XR55 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
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
40
Ar/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
40
Ar. 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
of 37Ar (T0.5 = 35.1 days) and 39Ar (T0.5 = 269 years). The mass discrimination factor
during analysis was 1.00436 based on ongoing measurements of
40
Ar/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:
36
Ar/37Ar(Ca)=2.603E-4 ± 2.373E-9,
40
Ar/39Ar(K)=1.547E-2 ± 7.455E-7.
39
Ar/37Ar(Ca)=6.501E-4 ± 7.433E-9 and
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
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.
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
SueArGr
x
y
altitude [m]
700 588
174 999
2 123
Description
Assemblage*
Rock sample of the Southern Aar granite
Qtz, Kfs, Pl,
collected on a surface outcrop. XRF- analysis,
apatite FT
Ms, Chl, Ep,
Ap, Ttn, Rt,
Py, (Bt)
115 900
KB868
700 062
700 079
(115 935)
173 863
173 832
∼500
∼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.
Fissure assemblages in the Southern Aar
Apo, Lmt,
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
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.
142
TABLE 4.2. Bulk rock geochemistry
SueArGr
115 900
wt %
SiO2
68.86
TiO2
0.42
0.47
Al2O3
16.23
15.46
Fe2O3tot
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
100.67
99.80
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
Totals
67.50
ppm
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
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
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
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
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
-1.22
-1.51
-0.34
-2.57
-2.03
-2.09
0.68
0.67
0.68
0.68
0.68
0.68
wt.%
Totala
E%b
Si/(Si+Al)
a
Totals 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).
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
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
13.022
13.020
13.035
13.046
13.032
12.985
Total
a
50.96
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
b
Totals include traces of Ba, Fem Mg, Mn and Sr. 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.
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
40
Ar/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.
TABLE 4.5. 40Ar/39Ar increment heating ages of apophyllite
Sample no.
Total fusion
Plateau age
age (Ma)
(Ma ±2σ)
ATW - APO
2.11 ± 0.06
2,04 ± 0.14
TW003 - APO
1.83 ± 0.05
1.96 ± 0.08
MSWD
39
40
Ar/36Ar
Ar % of
Isochron age
total
(Ma ±2σ)
intercept
J
9.74
99.50
2.22 ± 0.41
225 ± 118
0.0008215
1.61
56.01
1.46 ± 0.81
407 ± 189
0.0008222
147
FIG. 4.6: Plots of 40Ar/39Ar 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
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
No
and No.
Crystals
115 900
SueArGr
Spontaneous Induced
Dosimeter
Central
D(par)
Mean
S.d. of
rs
ri
rd*
(+/-S.d.)
(Ns)
(Ni)
(Nd)
FT Age (Ma)
(-2σ/+2σ)
Track
Length
distribution
(No. Tracks)
13.70
1.21
14.08
1.25
apatite
0.007
0.315
(40)
(133)
(6018)
apatite
0.011
0.316
(40)
(194)
(5580)
Pχ2
87 %
92 %
16.08
6.1
2.33
(12518)
(-1.0/+1.2)
(± 0.29)
01605
9.6
3.07
(12491)
(-1.3/+1.5)
(± 0.48)
(32)
(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.
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).
Ca2+ + 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.
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 oversaturation 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.
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
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
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
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,
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:
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.
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.
157
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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.
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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.
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.
161
Zen, E. (1961) The zeolite facies: an interpretation. American Journal of Science, 259,
401-409.
APPENDIX
APPENDIX
I Own Contribution
Contribution of Tobias Weisenberger on paper manuscript processing presented in
Chapter 2, 3 and 4.
Chapter 2:
Weisenberger1 T. and Bucher1 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:
Weisenberger1 T. and Bucher1 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”
Chapter 4:
Weisenberger1 T., Rahn1,2 M., van der Lelij3 R., Spikings3 R. and Bucher1 K. TIMING
AND MINERAL EVOLUTION DURING LOW-TEMPERATURE FLUID-ROCK INTERACTION ON
UPPER CRUSTAL LEVEL:
TRACK ANALYSIS ON
40
AR/39AR
ALPINE
APOPHYLLITE-(KF) DATING AND APATITE FISSION
FISSURES
(CENTRAL ALPS/SWITZERLAND). shortly to be
submitted to “Mineralogical Magazine”
1
Institute of Geosciences, Mineralogy - Geochemistry, Albert-Ludwigs-University Freiburg, Albertstr.
23 b, 79104 Freiburg, Germany
2
3
Eidgenössisches Nuklearsicherheitsinspektorat ENSI, 5232 Villigen, Switzerland
Department of Mineralogy, Université de Genève, Rue des Maraîchers 13, 1205 Geneva, Switzerland
i
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 lowtemperature 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 Lowgrade 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 DeutschIsländischen 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

zeolites in fissures of crystalline basement rocks

Transcription

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