Download Complementi di Petrografia N.O Scienze Geologiche, Lezione n. 3

Complementi di Petrografia N.O Scienze Geologiche, Lezione n. 3
Definizione, limiti, tipi e fattori del metamorfismo
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PRESSURE, TEMPERATURE, FLUID PRESSURE CONDITIONS OF METAMORPHISM
Marco Scambelluri, Dipartimento per lo Studio del Territorio e delle sue Risorse, University of
Genova, Corso Europa 26, 16132 Genova, Italy
Email: [email protected]
Keywords: metamorphism, temperature, pressure, fluid
Contents
1. Introduction
2. General Features of Metamorphism
2.1. Pressure-Temperature Conditions of Metamorphism
2.2. Types of Metamorphism
2.3. Kinetics of Metamorphic Reactions
3. Temperature
4. Pressure
5. Variations of Metamorphic Mineral Assemblages in Dependence of Pressure and
Temperature
6. Role of the Fluid Phase During Metamorphism
6.1
The Catalysis of Mineral Reactions by Fluids
6.2
Incorporation of Fluids in Metamorphic Rocks and Associated Compositional
Variations
6.3
Fluid Release in Metamorphic Rocks: Dehydration and Decarbonation Reactions
6.4. Influence of Fluids on the Behaviour of Metamorphic Rocks
7. Trends in Metamorphic Petrology
Glossary
Metamorphism: the mineralogical and structural changes of solid rock in response to
environmental conditions.
Temperature: it is the most important variable in metamorphic processes, since many metamorphic
reactions are driven by changes in temperature. It is a measure of how hot the rocks are; it is
ususally referred to in degree Celsius (°C) or in Kelvins (K; K = °C + 273,15).
Pressure: it is the second most important variable and many metamorphic reactions are pressuredependent. It is a measure of the force per unit area to which a rock is subjected, and depends
on the density of the overlying rock column and on the depth. It is measured in bar ( 1 bar =
0.987 atmospheres), kilobar (1 kbar = 1000 bar) or in gigapascal (1 Gpa = 10 kbar).
Fluid: most metamorphic rocks contain chemical species such as H2O, CO2, H2, O2, CH4 and S
either as components of metamorphic mineral (e.g. hydrated silicates, carbonates, sulfiddes), or
as a fluid phase in the pores of rocks. Significant volumes of H2O and CO2 can be evolved
from rocks during metamorphic dehydration and/or decarbonation reactions, thereby producing
a free fluid phase in metamorphic rock systems.
Summary
This paper presents an insight on the main aspects related to the metamorphism of rocks.
Metamorphism is a process of mineralogical and compositional change that has affected
considerable volumes of rocks that are presently exposed on the Earth’s surface, and that constitute
large parts of old cratonic areas, mountain buildings, oceanic basins. These rocks still preserve at
surface environments relict, metastable, associations of metamorphic minerals and/or metamorphic
mineral inclusions that indicate their past crystallization at geological environments such as the base
of the continental crust, the boundaries between colliding plates, the deep mantle. The study and
analysis of these rocks has thus given important answers on the modes of accretion and
consumption of the Earth’s crust, on the evolution of mountain chains and continents and on a
number processes which have driven the evolution our planet through geological timescales.
Metamorphic transformations in rocks normally occur in the solidus state and in absence of
significant production of partial melts and magmas. They are driven by changes, which normally
result from tectonic processes, in some of the following variables: the pressure and temperature
conditions of a given rock body, its bulk chemical composition, the availability and access of a C-OH bearing fluid phases, the presence and/or absence of rock deformation. Pressure, temperature,
bulk rock and fluid compositions are the principal factors controlling the type of the metamorphic
mineral assemblages produced, which are indicative of the dominant thermal regime during the
transformation and of the geodynamic environment. An increasing amount of research work and
evidence has been brought in the last decade to recognize the fundamental role played in
metamorphic processes by the fluid phases, particularly in the kinetics and enhancement of
metamorphic reactions, in the diffusion and transport of elements at variable scales in metamorphic
environments, in the development of large scale phenomena such as partial melting of crust and
mantle.
1. Introduction
Metamorphism of a given rock body occurs in response to significant changes of intensive
variables, such as pressure, temperature and composition, which disturb the pre existing equilibrium
conditions and force the rock to reach a new state of more stable equilibrium.
The IUGS subcommission on the systematics of metamorphic rocks, has defined
metamorphism as the process causing substantial changes in the mineralogy, structure and/or bulk
chemical composition of a given rock volume. Changes are due to physico-chemical conditions
different from the ones attained in sedimentary and diagenetic environments, and may include
partial melting as long as most of the rock volume remains in a solid state. If significant change in
the bulk rock composition is the dominant process due to open system behaviour of a given rock
body, the term metasomatism is applied.
In metamorphic terranes, preservation of large volumes of rocks recording various events of
recrystallization and of reconstitutive phase changes, provide evidence on the evolution of the
Earth’s system through time. Metamorphism is generally associated with large-scale tectonic
processes that cyclically occurred during the Earth’s history. On a global scale, metamorphic rocks
are dominant terrestrial materials and are among the oldest rock dated so far at 3.7 billion years:
their study is therefore a key to deciphering the main events and large-scale mass movements which
dominated the Earth’s history from her infancy to present.
This article aims at reviewing the main driving forces of metamorphism, grossly
individuated with temperature, pressure and role of the fluid phases involved. It will not consider
the important effects of deformation on rock recrystallization, on the catalysis of mineral reactions
and on reaction kinetics. This article bases on several textbooks on the petrography and petrology of
metamorphic rocks (Powell, 1978; Best, 1982; Yardley, 1989; Bucher and Frey, 1994, Spear, 1995),
these texts represent the most relevant sources of basic information on the process of
metamorphism.
2. General Features of Metamorphism
2.1 Pressure-Temperature Conditions of Metamorphism
The physico-chemical conditions at which metamorphic transformations begin depend on
the type, texture and composition of the rock material involved. Organic matter for instance, records
metamorphic changes at significantly lower temperatures than silicate and carbonate rocks. In many
rock systems, the boundary between diagenesis and metamorphism is faint and arbitrary, and
metamorphic phase transitions appear to develop at temperatures as low as 200 °C (Figure 1A). Key
minerals such as carpholite, stilpnomelane, paragonite and zeolites are indicators for the beginning
of metamorphism. The high temperature limit of metamorphism corresponds to the onset of partial
melting of a given rock system: metamorphic rocks recording high temperature recrystallization
associated with partial melting, display granitic melt layers and/or pockets, closely associated with
restitic rock volumes. The latter appear depleted in fusible components which were uptaken by the
melt phase. The partial melting temperatures of rocks depend on pressure, rock composition, as well
as on the amount and on the composition of fluid present. At a constant pressure and in presence of
water-rich fluids, partial melting of granitic rocks starts at lower temperature (about 650 °C) than
melting of basaltic rocks (about 700 °C). In absence of fluid, partial melting temperatures are as
high as 1000 °C for granitic systems and 1100 °C for basaltic ones (Fig. 1A).
8
7
Pressure Gpa
6
5
4
Pressure Gpa
4
3
5
Not realized
on Earth
3
METAMORPHISM
2
1
2
METAMORPHISM
1
MAGMATISM
MAGMATISM
0
200 400 600 800 1000 1200 1400
400
600
800 1000 1200 1400 1600
Temperature °C
Figure 1. A: pressure versus temperature diagram showing the main domain of
metamorphism in crustal rocks, comprised between the field of diagenesis at low
temperature and pressure and the wet solidus of granites and of basaltic systems (drawn
after Bucher and Frey, 1994). B: pressure - temperature diagram showing the field of
metamorphism in mantle peridotites, delimitated at high temperatures by the dry peridotite
solidus and, at lower temperatures, by the peridotite solidus in presence of water and of
CO2.
The above temperature ranges thus represent the upper temperature limits for crustal
metamorphism. At mantle conditions the temperature limit of metamorphism rises significantly,
solid state changes in this environments being attained at much higher temperatures. Figure 1B
shows the solidus curves of peridotites in absence of fluid (dry solidus) and in presence of carbonic
(solidus + CO2) and aqueous solutions (solidus + H2O): it appears that in absence of fluid phases,
subsolidus (metamorphic) phase transitions in mantle rocks can be attained at temperatures as high
as 1500 - 1600 °C and at relevant depths for a range of geothermal gradients. Great part of upper
mantle peridotites thus record subsolidus changes and thereby behave as metamorphic rocks.
Concerning the pressure limits of metamorphism, low pressure transformations occur in most
contact aureoles formed during magma ascent and emplacement at shallow levels in the Earth’s
crust. The high pressure limit of metamorphism is yet unknown. In the recent past, it was thought
that the maximum pressures attained by metamorphic rocks buried at convergent plate margins and
recrystallized under high-pressure conditions (eclogite-facies rocks) did not exceed 10 kilobar
pressure, roughly corresponding to 30 kilometers depth. The more recent discovery that coesite is
the stable SiO2 form in eclogite-facies rocks of several orogenic belts (as for instance the Italian
Alps, the Caledonides of Norway, the Dabie-Shan mountains of China) has set the high pressure
limit of metamorphism above 30 kilobars (90-100 km; ultra-high pressure metamorphism) (Chopin,
1984; Smith, 1984; Wang et al., 1989). Finding of diamonds in some ultra-high pressure rocks has
set the possible limit at much deeper levels (Sobolev and Shatsky, 1990; Dobrzhinetskaya et al.,
1995). An extraodinary record of ultradeep provenance of mantle rocks has been recently proposed
for the metamorphic garnet lherzolites of Alpe Arami in the Swiss Alps (Dobrzhinetskaya et al.,
1996; Bozhilov et al., 1997) and of Western Norway (Van Roermund and Drury, 1998). Although a
debate is going on to assess the exact origin and depth of provenance of these rocks, particularly the
Arami garnet peridotite (Hacker et al., 1997; Green et al., 1997; Trommsdorff et al., 2000), they are
thought to retain phase transformations of deep mantle olivine (wadsleyte) and of majoritic garnet
(forming at depths close to the transition zone in the upper mantle,) into mineral assemblages stable
at shallower depths of about 80-100 km in the upper mantle. These recent discoveries derive from
the application of advanced techniques to the current analysis of metamorphic rocks and shift the
pressure boundary of metamorphism recorded by rocks presently exposed on surface to extreme
depths into the upper mantle, thereby enabling to considerably deepen our knowledge on the
behaviour of the Earth’s interior.
2.2 Types of Metamorphism
Metamorphism can be manifest over large regions such as orogenic chains at convergent
plate margins, cratonic areas, oceanic basins, and extensional environments where deep crustal
and/or mantle rocks are slowly exhumed to the surface. On the other hand, metamorphism can be
the result of local-scale processes such as development of kilometer-large contact aureoles around
plutons intruded at high crustal levels in cool country rocks, or frictional heating along major faults.
Orogenic metamorphism dominates in mountain buildings, where considerable volumes of
rocks with different paleogeographic and lithosheric provenance are tectonically stacked together as
the result of large scale movements during plate convergence and collision. These orogenic cycles
bring surficial crustal rocks to mantle depths and then return them to the Earth surface, thus causing
superposition of several metamorphic events coupled with permanent ductile deformations. Main
features of metamorphic rocks exposed in orogenic belts are their strongly deformed structures,
developed as the result of stress and deformations developed at variable temperatures and pressures
during their orogenic pathways. However, metamorphism and deformation are extremely
heterogeneous in these rocks, and the records of the starting materials are systematically preserved
in several undeformed rock domains, thus enabling to reconstruct the whole history of rock
materials involved in the orogenic cycle.
Oceanic basins are diffusely floored by mafic and ultramafic rocks metamorphosed at
variable temperatures and moderate pressures in the presence of seawater-derived solutions. A great
amount of petrographic, petrologic and geochemical works have been performed on oceanic
metamorphic rocks in the course of Ocean Drilling Projects aimed at defining the dynamics of
present-day oceans. Diffuse features of oceanic gabbros, basalts and peridotites are hydration
reactions at variable conditions which affect these rocks in the vicinity of oceanic ridges and during
lateral spreading off the oceanic centres. Close to mid oceanic ridges, the lithosphere is cut by
hydrothermal systems where deep seawater penetration occurs (down to about 3 km and more)
accompanied by complex fluid/rock interactions which determine an exchange of components with
the surrounding rocks. These processes locally bring to metasomatism of rocks, consisting of Mg-,
Ca, Na-enrichments and diffuse Si-depletion. The above features significantly control the element
and volatile budgets in oceans and enrich the oceanic lithosphere in exogenic and crustal
components, which become recycled into the mantle once the oceanic lithosphere is deeply buried
along subduction zones.
Contact metamorphism is the most diffuse type of local metamorphism affecting rocks at the
contact with intrusive and extrusive igneous bodies. Metamorphic and metasomatic changes in these
country rocks are determined by heat flow and by infiltration of late-stage igneous fluids emanating
from the magma chamber. As a consequence, aureoles of contact metamorphic rocks generally
develop around plutons. The extension and width of aureoles depend on several factors such as
volume and composition of intruded magmas, on the depth of emplacement as well as on the
properties of country rocks. The volume of an intrusive body is important because large plutons
bring more heat than smaller ones. Also, the compositions of the dominant intrusive rock type is
another parameter controlling the overall temperature, since granitic melts form at much lower
temperatures than basaltic ones. The intrusion depth determines the thermal gradient and heat flow
between hot plutons and country rocks. The highest temperature differences are attained at surface
crustal levels, since at deep crustal environments the temperature differences between country rocks
and magmas are much lower.
2.3 Kinetics of Metamorphic Reactions
The process of metamorphism does not affect homogeneously a given rock body; from
kilometric to microscopic scales the records of earlier evolutionary stages can survive a
metamorphic event. Metamorphic terrains therefore contain rock domains which equilibrate at
certain (dominant) metamorphic conditions; these domains are spatially associated with minor
volumes of rocks that still record previous geologic events, manifest as relict minerals and as
preserved rock textures. The basement rocks of Western Norway represent a well known example of
such an association of rock volumes recording metamorphic imprints acquired at different geologic
ages. Here mafic rocks with Precambrian granulite - facies metamorphism outcrop as relict bodies
inside major volumes of mafic rocks reequilibrated at eclogite - facies conditions during the
Caledonian orogenesis. Granulitic and eclogitic domains were never separated during the whole
time span from Precambrian to Caledonian (Griffin and Carswell, 1985), and their spatial
association has been shown to result from unfavourable reaction kinetics during transformation of
granulites into eclogites. Detailed structural and petrologic studies (Austrheim, 1987) demonstrated
that this transition mainly took place in the rock volumes more intensely affected by plastic
deformation and by infiltration of aqueous fluid solutions.
Similarly, widespread exposure of rocks formed in the deep crust and mantle, indicates
unfavourable reaction kinetics during their exhumation pathways to the Earth’s surface. Deep
metamorphic mineral assemblages survive outside their stability field due to a combined effect of
exhumation rates faster than the ones of metamorphic mineral reactions, and lack of fluid
infiltration. Presence of fluids would cause hydration reactions and re-equilibration of high grade
minerals into hydrous mineral assemblages stable at shallow crustal environments. Despite mineral
reactions are expected on the base of thermodynamic predictions, high activation energies can be
required for the transformation to develop. Until these energy barriers are not trespassed,
metamorphic reactions and, most importantly, nucleation and growth of new metamorphic minerals,
do not occur. Activation energies are lowered by plastic rock deformation, by significant
temperature overstepping of a mineral reaction boundary, and by presence of fluids and deformation
which control the diffusions of cations. Field and petrographic experience indicate that rock
volumes less affected by fluid and deformation activity are the ones less intensely transformed and
better preserving the pre-metamorphic features.
The general preservation of deep rocks at the Earth’s surface, as well as diffuse survival in
metamorphic terrains of pre-metamorphic features, indicate that slow reactions kinetics occurred
systematically, thereby preventing the full recrystallization of a given rock volume during a certain
metamorphic event. When catalytic factors do not assist mineral reactions, rocks do not transform
during a given metamorphic event and the nature of pre-metamorphic materials metastably survives
the transformation.
3. Temperature
Surface heat flow W/m
2
Temperature is a driving force of metamorphism. Temperature increases with depth in the
Earth and the rate at which it changes with depth defines a geothermal gradient. Geothermal
gradients in the average continental crust are usually around 30 °C/Km, but large variations occur
depending on the thermal structure of the various terrestrial environments. The variability of
geothermal gradients is reflected by variations of heat flow measured at the surface, and the amount
of heat at a given geological setting depends on the following sources: 1) heat rising from the
mantle; 2) heat produced by radiogenic decay, particularly effective in acid crustal rocks; 3) heat
associated with flow and intrusions of large magmatic bodies; 4) heat brought to the surface during
fast uprise of hot, deep seated rocks either in extensional settings, or in areas of rapid uplift of
mountain chains.
Variations of surface heat flow shown in Figure 2 reflect the geothermal gradients at several
geodynamic environments (after Yardley, 1989).
Mid-oceanic ridges
0.4
0.3
oceanic crust of
increasing age
0.2
Subduction zone
trenches
0.1
Volcanic
arcs
Back-arc
basins
0
500 km
Figure 2. Variability of surface heat flux at several geodynamic settings (drawn after
Yardley, 1989).
The highest surface heat flows are attained at mid-oceanic ridges, where uprise of
asthenospheric mantle to superficial levels is accompanied by partial melting and widespread
magmatic activity. Here the heat flux from the mantle and from large magmatic chambers are the
major components of such high geothermal gradients. The surface heat flow progressively decreases
away from mid-oceanic ridges, since lateral spreading off oceanic centres is accompanied by aging,
hydration and cooling of the oceanic lithosphere. The lowest geothermal gradients are attained in
Figure 2 at subduction zones, where burial to great depths of cool crustal material screens the heat
flux from the mantle and strongly depresses the geotherms. In crustal and in supra-subduction
environments the geothermal gradients rise, due to a combined effect of increased heat flux from the
mantle, radiogenic decay and uprise of magmatic bodies. The latter activity is particularly
significant at compressive margins.
The different values of surface heat flow therefore reflect various thermal regimes attained in
the different tectonic and geodynamic settings. The rates at which temperature changes with depth,
i.e. the geothermal gradient, in the various tectonic and geodynamic environments are shown in the
pressure-temperature diagram of Figure 3. The diagram shows the range of metamorphic conditions
attained in various settings and also reports for reference the melting curves for dry and wet granite
and basaltic systems. High-pressure low temperature metamorphism develops in subduction zones
and in geologic regions were rapid overthrusting of large continental masses has occurred. Low
pressure high temperature metamorphism characterizes the environments of high heat flow, such as
the oceanic ridges, the island arcs and the contact aureoles. Finally, the pressure temperature region
of intermediate gradients is generally achieved in orogenic belts were collision of continental slices
has occurred.
2
1.8
60
1.6
g es
e a n rid
c
o
id
a rc s, m oles
Isla nd
t a ure
c onta c
0.2
so lid
us
g ra n
20
a sa lt
l
10
0
0
200
400
600
800
Temperature °C
1000
1200
Depth km
30
d ry b
0.4
a
nt
e
in
nt
o
C
c
ni
ts
d ry
zo
n
uc
tio
n
0.6
e
og
r
o
l
be
ite s
o lid
us
es
d on
Ea rth
Su
bd
0.8
Not r
e
1
40
wet b asalt solidus
a lize
1.2
50
lid us
wet gra nire so
Pressure GPa
1.4
Figure 3. Pressure versus temperature diagram reporting the ranges of geothermal
gradients (shaded areas) attained at different geodynamic settings (drawn after Spear,
1995).
A(O
H) +
C+
B
D+ 2
HO
Pressure
A general process showing metamorphic and textural variations in rocks during an overall
process of temperature increase during Burial of a metamorphic rock along a given geotherm is
shown in Figure 4.
P2
P1
1
T1
B
1
C
A(OH)
2
E
C
E
2
C
T2 Temperature
Figure 4. Schematic pressure - temperature diagram reporting a metamorphic path (bold
line) crosscutting two mineral reaction boundaries (1 and 2) producing new mineral
assemblages at the expense of reactant phases. The two drawings at the right side of the
diagram show the reaction microstructures and the possible textural relations between
reactant and product phases formed as the result of reactions 1 and 2 (after Bard, ).
As a result of increasing temperature and pressure the rock will experience heating and will
undergo prograde metamorphism. In Figure 4, a rock recording prograde metamorphism will
trespass the reaction curve where minerals A + B will react to produce a new metamorphic
assemblage made of B + C + a free aqueous fluid phase. Subsequently, at higher temperature and
pressure a new mineral phase E will form at the expense of mineral C through the reaction C = E.
The nucleation and growth of product phases C, D, E will take place at either at the grain
boundaries, or above the reactant phases A, B (or C in case of reaction 2 in Fig. ). The newly formed
minerals will surround the reacting materials and the latter will appear as corroded, relict precursor
mineral phases. Along the prograde metamorphic path of Figure 4, a new paragenesis will be stable
at the increased temperature conditions and a new rock fabric will be therefore produced.
Conversely, if a volume of rock is involved in a cooling process, it may undergo retrograde
metamorphism.
The reactions shown in Figure 4 all involve formation nucleation and growth of new mineral
assemblages during a heating process. Crystallization of the new mineral assemblages will lower the
free energy of the whole system. However, even when formation of the new mineral assemblage is
energetically favoured, minerals will not necessarily nucleate and grow. This largely depends on the
activation energies required for the transformation to occur: these represent energetic barriers to
nucleation and growth, that must be overcome by means of catalysts and through significant
temperature increase beyond the reaction boundary. The excess heating required is called
overstepping and the amount of overstepping is related to the nature of the mineral reactions and on
A
Nuc leii of B
B
A
+
C
∆T
Overstepping
Temperature
Gibbs Free Energy
Gibbs Free Energy
their temperature dependence (Fig. 5). Figure 5A shows a reaction with large temperature
dependence, such as many dehydration and decarbonation reactions: this implies that the reaction
boundary must be overstepped of a small amount to nucleate and grow the reaction products. The
reaction shown in Figure 5B has a small temperature dependence which may result in a larger
overstepping. The overstepping effect if minerals have an energy budget stored in the form of strain.
B
Nuc leii of E
A+
C
E
∆T
Temperature
Overstepping
Figure 5. Gibbs energy versus temperature diagram showing the temperature dependence of
two mineral reactions (A + B = C and A + B = E) and the amount of overstepping required to
actually achieve nucleation of the product phases B and E.
4. Pressure
The changes in pressure are another dominant cause of metamorphism: they are related to
variations in the depth reached by metamorphic rocks during their history. The prevailing pressure
at a given depth is defined by the average density of the overlying rock column. It can be calculated
from P = ρgh, where g is the acceleration due to gravity, ρ is the density, h is the depth. Pressure is
normally measured either in kilobar (1000 bar; 1 bar = 0.987 atmospheres) or in gigapascal (10
Kbar = 1GPa). The total pressure acting at a point in the crust or mantle is given by the weight of
overlaying rocks and is called lithostatic pressure. This pressure is assumed to be isotropic, i.e. is
equal in all directions. Lithostatic pressure does not cause rock deformation, and many rocks
undergoing metamorphism at high and very high pressures may not display significant distortion
and may still preserve pre-metamorphic textures. Non isotropic (oriented) pressures often occur in
metamorphic environments as the result of tectonic forces causing permanent rock deformation.
Non-isotropic pressure is one major factor that causes permanent ductile and brittle deformations
and that controls the textural features of rocks (metamorphic foliations, compositional layerings,
vein systems). Ductile deformation causes plastic flow of rock materials along high strain horizons
(shear zones), diffusion and redistribution of chemical components in metamorphic rocks. In
metamorphic terrains deformation is heterogeneous and rocks showing highly deformed fabrics
(tectonites and mylonites) outcrop close to rocks with the same bulk composition and metamorphic
mineral assemblages, but showing undeformed textures (massive and coronitic rocks). The
similarity of metamorphic assemblages indicates the these texturally different rock types underwent
metamorphic recrystallizations at the same pressure-temperature conditions, but under different
deformation regimes. Heterogeneity in the intensity of deformation in closely associated rock
volumes is a general feature of metamorphic rocks and depends on a process known as deformation
partitioning, which causes the spatial coexistence of high-strain and of low-strain domains. The
latter still preserve textural and mineralogical records of evolutionary stages predating a given
metamorphic and deformation event.
In all geologic environments changes in pressure- and depth-related changes are induced by
tectonic processes. Pressure changes recorded by rocks presently exposed in orogenic belts indicate
that during subduction at convergent plate margins large portions of surface materials have been
tectonically buried into the mantle to depths exceeding 100 kilometers, thereby developing
assemblages made of high density metamorphic minerals (including garnet and pyroxene). These
tetctonic units made of high and ultrahigh pressure metamorphic rocks were then returned to the
surface during tectonic exhumation processes, causing the retrograde metamorphic imprint of the
high pressure minerals by less dense, hydrated mineral assemblages. Variations in depth associated
with pressure changes also occur in subsiding sedimentary basins during overall lithospere
extension and thinning. This type of metamorphism experienced by deeply buried sedimentary rocks
is called burial metamorphism.
5. Variations of Metamorphic Mineral Assemblages in Dependence of Pressure and
Temperature
Barrow (1893) firstly mapped in the crystalline basement rocks of the Scottish Highlands a
sequence of zones characterized by the first appearance in pelitic schists of metamorphic index
minerals (such as chlorite, biotite, garnet, staurolite, kyanite and sillimanite), which were related to
an increasing grade of metamorphism, i.e. to a progressive increase of temperature and pressure
conditions. The individual minerals are systematically distributed in distinct regional mineral zones
(metamorphic zones), and adjacent zones are separated by isograds, and/or reaction isograds. These
correspond to surfaces of equal metamorphic grade separating domains metamorphosed at different
pressure and temperature conditions, and therefore represent paleogeotherms. Since Barrow’s work,
studies and maps have been performed on metamorphic crystalline basements in order to understand
the distribution of the different mineral assemblages in and to relate them to a regional variability in
the main parameters governing metamorphism, such as pressure, temperature and the bulk
compositions of rocks. The concept of index mineral then evolved to mineral paragenesis: an
assemblage of metamorphic minerals coexisting in textural and chemical equilibrium in a rock of a
given bulk composition, and formed at a certain temperature and pressure. In a metamorphic terrain
rocks metamorphosed at the same pressure-temperature conditions develop different parageneses in
dependence of their different bulk composition. Therefore the concept of metamorphic facies
(Eskola, 1915) was proposed to define the whole set of mineral parageneses occurring in spatially
associated rock types of diverse bulk chemical compositions and metamorphosed under the same
broad pressure – temperature conditions. Figure 6 shows the ranges of pressure and temperature
conditions at which the metamorphic facies are interpreted to form.
The boundaries between the different facies are not sharp because most characterizing
mineral assemblages form by continuous, rather than discrete (discontinuous), reactions. This
depends on the fact that most minerals form complex solid solutions and that metamorphic fluids
correspond to compositionally complex solutions: these features determine a shift of mineral
reactions over pressure temperature regions and make the transition among different facies to along
pressure temperature intervals. Figure 6 also reports the gradients typical of various geodynamic
settings and shows that the transition from one facies to the other, as mapped in several
metamorphic terrains, depends on the thermal gradients that affect rocks during their history. As a
consequence Miyashiro (1973) emphasized that certain metamorphic facies are commonly
associated to the exclusion of others in different orogenic belts; a metamorphic facies series in a
given terrain is therefore diagnostic of the geothermal gradient attained and of the metamorphic
paths followed by rocks during their evolution.
2
1.8
1.6
1.4
d on
Ea rth
Eclogite
a lize
Greenschist
0.6
Prehnited
Pumpellyitete rm e
P/T
ia te
In
0.4
Amp
hib
P/T
Hig
h
0.8
o lite
Bluesc hist
Ep id
a mp ote
hib o
lite
1
No t
re
Pressure GPa
1.2
Granulite
Zeolite
0.2
Lo w P/T
0
0
200
400
600
800
1000
1200
Temperature °C
Figure 6. Pressure temperature diagram ilustrating the stability field of the various
metamorphic facies and reporting the major types of metamorphic facies series attained
at different tectonic envirronments (drawn after Spear, 1995). High P/T metamorphism
is characteristic of subduction zones, intermediate P/T is characteristic of continental
collision zones and low P/T characterizes ocean ridges and island arcs. The ranges of
geothermal gradients of the above geodynamic environments are shown in Figure 3.
6. Role of the Fluid Phase During Metamorphism
It has been stressed in the previous chapters that metamorphic processes are extremely
heterogeneous and that development of metamorphic reactions and mineral growth strongly depends
on the availability of catalysers. In this scenario, presence of fluid phases at the reaction sites and
the composition of these fluids play an important role. As outlined in section 3.2, in the basement
rocks of Western Norway fluid infiltration controlled the eclogitization of deep continental
granulitic crust. The metamorphic recrystallization of these rocks was not simply a function of
pressure, temperature and rock composition, but strongly depended on the presence of metamorphic
fluids. Fluids also accomplish mass transfer and element mobility at various scales. Dehydration
and/or dehydration-melting of crustal rocks during orogenic processes can be associated to large
scale mass transfer and can produce significant compositional variations in the deep continental
crust (Clemens and Droop, 1998). Upwards migration of metamorphic fluids released during
subduction of oceanic plates causes hydration and metasomatism of the mantle wedges overlaying
the subduction zones: this lowers the mantle solidus and favours its partial melting with consequent
formation of the arc and calcalkaline magmas typically produced at convergent plate margins
(Tatsumi, 1989). These geologic examples illustrate that fluids play a major control on the catalysis
of metamorphic reactions, on mass transfer and on the mechanical properties of rocks and minerals.
6.1 The Catalysis of Mineral Reactions by Fluids
The efficiency of fluids as catalysers of mineral reactions is qualitatively well known and has
led several authors to suggest that water availability at the reaction sites represents the most
important factor for the development of relevant mineral reactions, such as for instance the gabbro
to eclogite transition at subduction zones (Ahrens and Schubert, 1975; Rubie, 1990). The presence
of water during transformation of aragonite to calcite lowers the activation energy required for the
reaction to occur (about half the energy necessary at dry, water absent, conditions), and rises the
reaction rates of several orders of magnitude (Fig. 7; Brown et al., 1962).
50
100 150 200 250 300 400 T °C
ARAGONITE
CALCITE
100 Ma
DRY
W
ET
103 a
time
1 Ma
1a
1month
1 d ay
1 hr
3.0
2.5
2.0
3
1.5 10 /T K
Figure 7. Aragonite - calcite transformation at dry and wet conditions shown as a
function of time and temperature (Brown et al., 1962).
Fluids are generally incorporated in hydrous minerals and are released in the course of
devolatilization reactions during heating and burial of rocks. The presence of metamorphic fluids
can thus be limited to short periods of time coincident with mineral devolatilization episodes and
with infiltration of fluids deriving from external sources. Fluids are heterogeneously distributed in
rocks and preferentially locate along intergranular surfaces and within pore spaces. The interganular
domains of a rock can be: i) dry; ii) hydrated, but undersaturated in fluid; iii) saturated in fluid.
Below a given water/rock ratio, water molecules are incorporated in crystalline lattices, particularly
in the regions close to grain boundaries. The water dissolved in such domains facilitates diffusion
(grain boundary diffusion), probably substituting strong Si-O bonds with weaker Si-OH bonds,
thereby effectively reducing the activation energy required for metamorphic transformations. Above
a critical water/rock ratio, grain boundaries become water saturated and a free fluid phase is present
in the system, and this fluid phase can be interconnected along grain edges. In presence of a free
fluid phase the diffusion mechanism changes from diffusion in a solid crystalline net (grain
boundary diffusion) to diffusion within a fluid medium along interfaces between mineral grains
(interface diffusion).
6.2. Incorporation of Fluids in Metamorphic Rocks and Associated Compositional Variations.
Sedimentary rocks contain large amounts of fluid fixed in the structure of hydrous minerals
and present as a free phase in pore spaces; before diagenesis and prograde orogenic and burial
metamorphism the rocks are at their maximum state of hydration. Differently, igneous rocks have
low initial fluid and water contents and incorporate fluid as the result of secondary alteration driven
either by hydrothermal, or by meteoric waters. In this respect, alteration and hydration of the oceanic
crust and lithosphere is particularly relevant, since in oceanic basins rocks of igneous and/or
metamorphic origin having very low initial water contents (basalts, gabbros, mantle peridotites)
become significantly enriched in water, carbon dioxide and other volatile components as the result
of interaction with seawater-derived solutions. Water/rock interaction during the alteration of
oceanic basalts induce relevant compositional variations of several rock components (MgO, CaO,
Na2O, K2O, Cl, H2O, CO2, stable C-O-H isotopes) as a function of fluid composition and chemistry,
of the alteration temperature and of water rock ratios during the alteration process (Mottl, 1983;
Gregory and Taylor, 1981). The progressive enrichment in MgO and depletion in CaO documented
in many oceanic basalts (Fig. 8) is the result of alteration by seawater at increasing water/rock ratios.
Compared to fresh (unaltered basalts), the compositions of altered basalts sampled in the Mid
Atlantic Ridge (MAR), in the East Pacific Rise (EPR) and in the Indian Ocean (Fig. 8) display a
general decrease in CaO coupled with an increase in MgO. Comparable compositional trends have
been obtained during laboratory experiments by reacting seawater with basalts at increasing water
rock ratios (W/R 10 to 50, 62 and 150).
MAR basalts
EPR and Indian
Ocean basalts
14
Fresh
basalts
12
W/R 10
CaO
W/R 50, 62, 125
10
W/R 10
8
6
4
W/R 50,
62, 125
2
A
0
0
5
10
15
20
25
MgO
30
35
40
Figure 8. MgO versus CaO diagram showing the compositional variations in oceanic
basalts as the result of interaction with seawater. The diagram reports the compositional
field of unaltered basalts; the compositions of variably altered oceanic basalts from the
Mid-Atlantic Ridge (MAR), from the East Pacific Rise (EPR) and from the Indian
ocean. Also reported are the results of seawater-basalts experiments at increasing
water/rock ratios (W/R 10, W/R 50, 62, 125) (Mottl, 1983).
The data indicate that alteration of oceanic rocks is a complex process that can imply,
besides enrichment in the fluid and volatile contents of rocks, significant gains and losses of major
and trace element components, and can even cause metasomatic compositional changes. The
metamorphic reactions between fluids and rocks can produce significant variations of the starting
rock compositions; they will control the type of mineral assemblages and the composition of fluids
which are going to be produced if these rocks become involved in a new metamorphic cycle (for
example prograde metamorphism during subduction and/or plate collision).
6.3 Fluid Release in Metamorphic Rocks: Dehydration and Decarbonation Reactions.
= 1
= 0
.8
1
O2 =
XC
O2 =
Wollastonite
+ fluid
2O
1
Calcite
+ quartz
+ fluid
B
aH
= 0
.6
= 0
.4
O
aH2
2O
2O
aH
P kbar
En +
qtz+
H2O
tlc
aH
Pressure
2
XC
XCO2 = 0.13
A
3
0.5
Fluid release during tectonic and metamorphic processes is both related to compaction of
rock material with consequent expulsion of free fluid phases present in pore spaces and fluid
inclusions, and to devolatilization reactions. The latter occur in the general form hydrated mineral
assemblage = less hydrated (or dry) assemblage + fluid at increasing temperature, and control the
depths of fluid release and the amounts of fluids produced during a metamorphic cycle. The
reactions involving breakdown of OH-bearing minerals and liberation of a water-rich phase are
most common. Figure 9A shows as an example the breakdown reaction of talc to enstatite + quartz
+ fluid.
0
Temperature
200
400
600
800
T°C
Figure 9. A: pressure - temperature diagram showing the dehydration curve talc = enstatite + quartz
+ water for different fluid compositions (after Bucher and Frey, 1994). B: pressure - temperature
diagram reporting the reaction calcite + quartz = wollastonite + CO2 for different CO2
concentrations in the fluid phase (after Best, 1982).
In a pressure-temperature diagram, the dewatering reactions generally display a change in the slope
of the reaction curve from positive at low pressure, to negative at high pressure. This fact is due to
significant variations in the molar volume of the fluid produced (high molar volumes at low
pressure and low at high pressures), which in turns determines the ∆V of the reaction and, hence,
the slope of the reaction curve. The shape of the reaction curve in Figure 9A indicates that talc
dehydration (as well as dehydration of other hydrous mineral phases) can be caused by increasing
pressures and temperatures (heating and burial), and by decrease in pressure. The temperature at
which a dewatering reaction takes place depends on the fluid composition. Breakdown of talc, for
instance, occurs at progressively lower temperatures if the associated fluid is not pure water and
dissolves other components (such as CO2, CH4, N2, Cl) and the water activity in the fluid becomes
lower than 1 (Fig. 9A). Presence of complex fluids the compositions of which is significantly
different from pure water has been largely described in the geological and petrological literature
(Andersen et al., 2001): one main consequence of the presence of such fluids is a significant
restriction of the pressure-temperature the stability fields of reactant hydrous phases. Similarly, the
stability fields of carbonatic phases reduce strongly if the associated fluid does not correspond to
pure CO2 and contains extra components, such as water, methane and nitrogen. This relationship is
shown in Figure 9B, where the reaction curve calcite + quartz = wollastonite + CO2 is shown for
variable concentrations of CO2 in the fluid. The CO2/H2O ratios in metamorphic fluids can
significantly change from one rock domain to the other during a given metamorphic event, as a
function of the amount of fluid and of its mobility, of the type of mineral assemblage buffering the
fluid composition, and of the permeability of rocks involved.
6.4 Influence of Fluids on the Behaviour of Metamorphic Rocks.
The distribution of water in the Earth’s crust ad mantle plays a control on the rheological
properties and on the ductility of rocks. The effects of water on the behaviour of rocks have been
quantified by means of laboratory deformation experiments involving dry and wet rocks and
minerals as starting materials. Water has a softening effect on the minerals and rocks undergoing
ductile deformation: this effect is known since a long time as hydrolitic weakening and has been
quantitatively well studied in quartzites, granites and dunites, i.e. the rock materials representative
of crust and mantle (Ranalli, 1995). The stress versus strain diagrams of Figure 10 indicate that
quartz and olivine crystals containing trace amounts of dissolved water are much more ductile than
dry crystals.
1400
a tu ra
l
A
d ry n
stress MPa
1200
B
1000
800
600
400
wet syntetic
200
0
1
2
3
4
5
strain
Figure 10. Stress versus strain diagrams for dry and wet olivine and quartz.
The weakening effect of water on olivine indifferently affects coarse and fine grained
crystals (Figure 10B). Comparable relationships have been shown to occur in granitic rocks and in
materials made of polycrystalline albite aggregates: in these materials the temperature at which
occurs the transition between fragile and ductile regimes lowers of 150-200 °C for water contents of
about 0.2 weight percent (Tullis and Yund, 1980). The distribution of metamorphic aqueous fluids
in the crust and mantle therefore has relevant implications on their rheological properties and
strongly enhances the localization of important deformation processes along given horizons
corresponding to water-enriched domains. The latter therefore represent candidates to become
ductile shear zones, where significant plastic flow and mass transfer occur during main large scale
tectonic events which accomplish the main changes ruling the geodynamic history of the Earth.
The partial pressure of the free fluid phase in the pore spaces of rocks and its volume
expansion during heating represent another important variable during metamorphism. The pressure
of fluid (Pfluid) counteracts the lithostatic pressure (Plitho), and locally the fluid pressure can
overcome the lithostatic one. In such cases rocks undergo hydraulic fracturing: fluid and rock
pressures re-equilibrate and fluids flow throught the rock along fracture systems. Norris and Henley
(1976) have shown that compaction of rocks during regional metamorphism and tectonic is
accompanied with dehydration reactions and temperature increase, causing heating and volume
expansion of the produced fluids. These processes cause an increase in fluid pressure, thereby
leading to hydrofracturing of rocks. Diffuse development vein systems during the entire tectonic and
metamorphic history of regional metamorphic rocks indicates that the one envisaged by Norris and
Henley can be an ubiquitous process. Figure 11 illustrates the molar volumes of water as a function
of pressure and temperature, together with a typical clockwise pressure temperature path of
metamorphic rocks undergoing a prograde burial followed by retrograde and decompressional
exhumation, as diffusely documented in most orogenic regional metamorphic terrains.
9
3
VH2O cm mole
16
17
-1
18
19
7
20
Pressure kbar
22
25
5
30
3
40
50
1
100
300
500
700
900
Temperature °C
Figure 11. Pressure - temperature diagram showing the molar volumes of water and a
typical clockwise metamorphic path (Norris and Henley, 1976).
Once a free fluid phase is present in the rock pores, either as entrapped connate water, or as a
metamorphic phase produced during dewatering reactions, its molar volume will continuously
increase along the burial path, as well as along the exhumation path. The continuous volume
expansion of fluids in the metamorphic pile episodically brings to fluid pressures higher than the
lithostatic ones and hence to hydrofracturing of the metamorphic rocks. The approach of Norris and
Henley (1976) explains the diffuse finding of multiple generations of fracture systems in
metamorphic rocks and the continuous spatial redistribution of fluids, mass transfer and fluid
channelling along fracture systems.
The above features indicate that ductile deformation appear extremely favoured in presence
of small amounts of dissolved fluids, whereas high partial pressures of free pore fluids favour the
brittle fracturing of rocks. Actually, little is known on the exact state and location of fluid during
active deformation. By means of experiments and microstructural studies at the transmission
electron microscope, Tullis and Yund (1996) have shown that free intergranular fluids become
incorporated in minerals during active deformation, thus significantly favouring plastic deformation
and mass and cation diffusion. During the post-deformation annealing of rocks materials
investigated, the fluids previously incorporated in minerals during active deformation and wetting
crystal defects, slip planes and microfractures, become re precipitated at the grain edges under the
form of pore fluids.
6. Trends in Metamorphic Petrology
In the recent past, research in metamorphic petrology has been devoted to understand
processes controlling the development of mineral parageneses in the various rock types, to develop
thermodynamic modelling of metamorphic processes and mineral reactions, and to the assessment
of the pressure and temperature conditions of metamorphism through the application of
geothermobarometers. All of these studies, joined with textural investigations to recognize the
protoliths and the reaction steps of metamorphic rocks, and with radiogenic isotope geochemistry to
define the actual ages of metamorphic events, have been largely used to reconstruct the thermal tectonic regimes during metamorphism and the pressure - temperature - time paths followed by
metamorphic rocks during their histories.
More recently, the interest has been focused on the role of fluids during metamorphism, their
control on reaction processes, on rock deformation, as well as on mass and heat transfer at several
crustal and mantle environments. Modelling of fluid behaviour in rocks, coupled with experiments
on element partitioning between fluids and the coexisting solids at variable pressures and
temperatures, is a fertile branch of metamorphic petrology and geochemistry enabling to quantify
the mass transfer via fluid phases.
Another important aspect of metamorphic research concerns the continuous finding of
ultrahigh pressure mineral records that witness the extraordinary depths reached by some
metamorphic rocks. This research line has been favoured by the advancements in experimental
petrology, and by the application of technologies enabling to investigate at nanoscales the structures
of rock forming minerals.
These research trends will deepen our knowledge on the processes acting at the deepest,
inaccessible, levels of our planet and governing the dynamics of the Earth’s litosphere and mantle.
Bibliography
Anderson B.D.O. and Moore J.B. (1979). Optimal Filtering, Englewood Cliffs: Prentice Hall. [This
is a book that presents all the essential aspects of filtering in dynamical systems]
Andersen T., Frezzotti M.L., Burke E.A.J. (2001). Fluid inclusions: phase relationships – methods –
applications. Lithos 55, 320 pg. [This is a special volume that presents all the essential aspects of
fluid –phase petrology through the analysis of fluid inclusions in rocks]
Ahrens T.J. & Schubert G. (1975). Gabbro-eclogite transition and its geophysical significance.
Reviews in Geophisics and Space Physics 13( 2), 383-400.
Austrheim H. (1987). Eclogitization of lower crustal granulites by fluid migration through shear zones.
Earth and Planetary Science Letters 82, 221-232. [This is a scientific article showing the
interdependence of fluid infiltration and metamorphic reactions in deep crustal rocks that
underwent continental collision]
Bard J. P. (1980). Microtextures of igneous and metamorphic rocks. D. Riedel Publishing
Company, Dordrecht, Boston, Lancaster, 264 pp. [This is a book that presents all the essential
aspects of microstructures in rocks]
Barrow G. (1893) On an intrusion of muscovite biotite gneiss in the S.E. Highlands of Scotland and
its accompanying metamorphism. Quarterly Journal of the Geological Society of London 49, 330-358.
[This is the first description and mapping of metamorphic zones in metamorphic basement rocks]
Best M. G. (1982). Igneous and metamorphic petrology. W.H. Freeman and Company, New York San
Francisco, 630 pp. [This is a book that discusses many basical aspect of petrology and presents a
very good link between field observations and theoretical systems]
Bozhilov K. Green H.W., Dobrzhinetskaya L.F. (1999). Clinoenstatite in Alpe Arami peridotite:
Additional evidence of very high pressure. Science 284, 128-132.
Bucher K and Frey M. (1994). Petrogenesis of metamorphic rocks. Springer-Verlag, Berlin
Heidelberg New York, 318 pg. [This is a book that presents the essential aspects of metamorphism
in all rock systems]
Chopin C. (1984). Coesite and pure pyrope in high grade pelitic blueschists of the Western Alps: a
first record and some consequences. Contribus to Mineralalogy and Petrology 86, 107-118. [This is
a first and very well known example of coesite in rocks of crustal origin, demonstrating their
subduction down to about 100 km depth]
Clemens J.D. and Droop G.T.R. (1998). Fluids, P-T paths and the fates of melts in the Earth’s crust.
Lithos 44, 21-36.
Dobrzhinetskaya L.F, Eide E.A., Larsen R.B., Sturt B.A., Tronnes R.G., Smith D.C., W.R. Taylor,
Posukhova T. (1995). Microdiamond in high-grade metamorphic rocks of the Western Gneis region,
Norway. Geology 23, 597-600.
Dobrzhinetskaya L.F, Green H.W., Wang S. (1996). Alpe Arami: A peridotite Massif from depths
more than 300 km. Science 271, 1841-1845. [This is a scientific article providing evidence for the
possible origin of Swiss Alpine ultramafic rocks at extraordinary depths in the upper mantle]
Eskola P. (1915) On the relations between the chemical chemical and mineralogical composition in
the metamorphic rocks of the Orijarvi region. Bulletin de la Commission Gelogique de Finlande 44.
Green H.W., Dobrzhinetskaya L.F, Bozhilov K. (1997). Determining the origin of ultrahigh-pressure
lherzolites. (Reply). Science 27, 704-707.
Gregory R.T. and Taylor H.P.J. (1981). An oxygen isotope profile in a section of oceanic crust,
Samail ophiolite, Oman: evidence of d18O buffering of the oceans by deep (> 5 km) seawter
hydrothermal circulation at mid ocean ridges. Jourbal of Geophysical Research 86, 2737-2755.
[This is a research paper providing a comprehensive picture of the stable isotope composition of
rocks in a section of oceanic lithosphere]
Griffin W.L. and Carswell D.A. (1985). In situ metamorphism of Norwegian eclogites: An example.
In Gee D.G. and Sturt B.A. The Caledonide orogen – Scandinavia and related areas, John Wiley, 813822.
Hacker B.R., Sharp T., Zhang R.J., Liou J.G., Hervig R.L. (1997). Determining the origin of ultrahighpressure lherzolites. Science 278, 702-704.
Mottl M.J. (1983) Metabasalts, axial hot springs, and the structure of hydrothermal systems at midocean ridges. Geological Society of America Bulletin 94, 161-180.
Miyashiro A. (1973). Metamorphism and Metamorphic Belts. George Allen and Unwin, London
492 pp.
Norris R.J and Henley R.W. (1976). Dewatering of a metamorphic pile. Geology 4, 333-336.
Powell R. (1978). Equilibrium thermodynamics in petrology. Harper and Row Publishers, 284 pp.
[This is a book that presents the basic aspects of thermodynamic modeling of rock systems]
Ranalli G. (1995) Rheology of the Earth. Deformation and flow processes in Geophysics and
Geodynamics. London: Allen and Unwin Eds., 366 pp.
Rubie D.C. (1986). The catalysis of mineral reactions by water and restrictions to the presence of
aqueous fluid during metamorphism. Mineralogical Magazine 50, 399-415.
Rubie D.C. (1990). Role of kinetics in the formation and preservation of eclogites. In: Carswell DA
(ed): Eclogite Facies Rocks, Blackie, 111-141.
Smith D.C. (1984). Coesite in clinopyroxene in the Caledonides and its implications for
geodynamics. Nature 310, 641-644.
Sobolev N.V. and Shatsky V.S. (1990). Diamond inclusions in garnets from metamorphic rocks.
Nature 343, 742-746. [This is a research paper that documents formation of diamond in
metamorphic crustally-derived rocks as the result of subduction at great depths]
Spear F. S. (1995). Metamorphic phase equilibria and Pressure - Temperature - time paths.
Mineralogical Society of Americs, Monograph, Washington DC, 799 pp. [This is a book that
thoroughly discusses many aspects of metamorphic petrology and their thermodynamic analysis]
Tatsumi Y., 1989. Migration of fluid phases and genesis of basalt magmas in subduction zones.
Journal of Geophysical Research 94 4697-4707.
Trommsdorff V., Hemann J., Muentener O., Pfiffner M., Risold A.C. (2000). Geodynamic cycles of
subcontinental lithosphere in the Central Alps and the Arami enigma. Journal of Geodynamics 30,
77-92.
Tullis J. and Yund R.A. (1980). Hydrolitic weakening of experimentally deformed Westerly granite
and Hale albite rock. Journal of Structural Geology 2, 439-451.
Tullis J., Yund R.A., Farver J. (1996). Deformation enhanced fluid distribution in feldspar
aggregates and implications for ductile shear zones. Geology 24, 63-66.
Van Roermund H.L.M. and Drury M. (1998). Ultra-high pressure (P > 6 Gpa) garnet peridotites in
Western Norway: exhumation of mantle rocks from > 185 km depth. Terra Nova 10, 295-301.
Wang X., Liou J.G., Mao H.K. (1989). Coesite-bearing eclogites in the Dabie Mountains in Central
China. Geology 17, 1085-1088.
Yardley B.W.D. (1989). An introduction to metamorphic petrology. Longman, Earth Sciences
Series, 248 pp.