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Abstract

In the last decade, the considerable amount of ultrahigh pressure terrains found around the world (Liou et al., 1998) and recent findings of mineral inclusions suggesting equilibration pressures up to more than 20 GPa (e.g. Stachel et al., 2000) have led to an increasing interest in petrological tools to be used to unravel the P-T evolution at extreme conditions. The expected resolution of such tools should provide a key for understanding tectonic processes responsible for the dynamics of UHP terrains.

Introduction

In the last decade, the considerable amount of ultrahigh pressure terrains found around the world (Liou et al., 1998) and recent findings of mineral inclusions suggesting equilibration pressures up to more than 20 GPa (e.g. Stachel et al., 2000) have led to an increasing interest in petrological tools to be used to unravel the P-T evolution at extreme conditions. The expected resolution of such tools should provide a key for understanding tectonic processes responsible for the dynamics of UHP terrains.

Even though there has been a substantial amount of experimental work in simple model systems, mainly devoted to define the maximum stability field of ultrahigh pressure phases or to retrieve thermodynamic data from phase equilibria constraints, information on phase relationships in more complex, Fe-bearing systems approaching natural compositions are still fragmentary. A brief list of references for a variety of model systems where ultrahigh pressure phases are found is given in Table 1. References cited are not intended to be exhaustive but provide a starting point for the relevant literature in some of the systems listed.

Table 1.

Model systems where ultrahigh pressure phases are found, with a brief list of references

SystemUHP phasesSource for literature
FMSolivine, pyroxene, wadsleyite, ringwoodite, ilmenite, majorite, perovskite, magnesiowustite, stishoviteAkaogi et al. (1998); Frost et al. (2001)
MASenstatite, pyrope, majorite, ilmenite, perovskite, “tetragonal phase”Akaogi et al. (2002); Heinemann et al. (1997)
FMASmajoriteO’Neill & Jeanloz (1994)
CMSdiopside, Ca-perovskite, walstromite, CaSi2O5Canil (1994); Gasparik et al. (1994);
Gasparik (1996a)
CFSskiagiteWoodland & O’Neill (1995)
CMAspinel, Ca-ferrite, “hexagonal Al-rich phase”Akaogi et al. (1999); Miyajima et al. (2001)
NCMASjadeite, Na Ca-ferriteGasparik (1996b)
NKASwadeite, hollandite, calcium ferrite(-structure)Yagi et al. (1994)
ASHtopaz-OH, phase Pi, phase eggDaniels & Wunder (1996); Schmidt et al. (1998)
MSHantigorite, clinohumite, phase A, phase B, superhydrous phase B (phase C), phase D (phases F and G), phase E, 10A phaseAngel et al. (2001); Ulmer & Trommsdorff (1999); Wunder (1998); Stalder & Ulmer (2001)
MASHchlorite, Mg-sursassite (MgMgAl pumpellyite), Mg-stauroliteBromiley & Pawley (2002); Fockenberg (1998)
CASHlawsonite, zoisitePoli & Schmidt (1998); Schmidt (1995)
CFASHclinozoisite, garnetBrunsmann et al. (2002)
KMASHphlogopite, wadeite, phase X, phengiteTronnes (2002); Massonne & Szpurka
KFASH(1997); Hermann (2002b); Schmidt et al. (2001)
CM-CO2dolomite, aragonite, magnesiteLuth (1995, 2001)
SystemUHP phasesSource for literature
FMSolivine, pyroxene, wadsleyite, ringwoodite, ilmenite, majorite, perovskite, magnesiowustite, stishoviteAkaogi et al. (1998); Frost et al. (2001)
MASenstatite, pyrope, majorite, ilmenite, perovskite, “tetragonal phase”Akaogi et al. (2002); Heinemann et al. (1997)
FMASmajoriteO’Neill & Jeanloz (1994)
CMSdiopside, Ca-perovskite, walstromite, CaSi2O5Canil (1994); Gasparik et al. (1994);
Gasparik (1996a)
CFSskiagiteWoodland & O’Neill (1995)
CMAspinel, Ca-ferrite, “hexagonal Al-rich phase”Akaogi et al. (1999); Miyajima et al. (2001)
NCMASjadeite, Na Ca-ferriteGasparik (1996b)
NKASwadeite, hollandite, calcium ferrite(-structure)Yagi et al. (1994)
ASHtopaz-OH, phase Pi, phase eggDaniels & Wunder (1996); Schmidt et al. (1998)
MSHantigorite, clinohumite, phase A, phase B, superhydrous phase B (phase C), phase D (phases F and G), phase E, 10A phaseAngel et al. (2001); Ulmer & Trommsdorff (1999); Wunder (1998); Stalder & Ulmer (2001)
MASHchlorite, Mg-sursassite (MgMgAl pumpellyite), Mg-stauroliteBromiley & Pawley (2002); Fockenberg (1998)
CASHlawsonite, zoisitePoli & Schmidt (1998); Schmidt (1995)
CFASHclinozoisite, garnetBrunsmann et al. (2002)
KMASHphlogopite, wadeite, phase X, phengiteTronnes (2002); Massonne & Szpurka
KFASH(1997); Hermann (2002b); Schmidt et al. (2001)
CM-CO2dolomite, aragonite, magnesiteLuth (1995, 2001)

Abbreviations (System column): A: Al2O3, C: CaO, F: FeO, H: H2O, K: K2O, M: MgO, N: Na2O, S: SiO2

Experimental data in simple model systems often provide fundamental constraints to basic mechanisms responsible for the appearance of ultrahigh pressure phases, but they are usually of limited help to resolve environmental conditions with sufficient accuracy in real rocks. Most natural systems are actually characterised by a large number of independent chemical components, therefore complex continuous reactions are expected to control most phase relationships.

A number of experimental difficulties are responsible for the relatively restricted number of experimental studies in complex systems at conditions reproducing cold geotherms (high dP/dT). Most researchers observe a progressive decrease in the “reactivity” of the experimental charges with increasing pressure and with the increasing complexity of solid solutions. As a result, most subsolidus charges show sluggish transformations at UHP conditions and poorly equilibrated textures, even in long duration experiments (compare Fig. 1a and 1b). Furthermore, conventional bracketing techniques cannot be applied in complex, natural systems. Despite these limitations, such experimental systems offer a unique perspective on phases and reactions of direct application to the Earth's interior.

Fig. 1.

Backscattered electron images of run products showing textural differences between experiments performed on similar bulk compositions and run durations, but at extremely different pressure and temperature conditions. (a) graphite-bearing, fluid-saturated model MORB at 2.7 GPa, 730 °C, 224 hours (courtesy Ada Crottini); (b) graphite-bearing, fluid-saturated natural MORB at 19 GPa, 1200 °C, 120 hours (courtesy Kazuaki Okamoto). gar: garnet; cpx: clinopyroxene; ky: kyanite; maj: majorite; ca-pvs: Ca-perovskite (bright spots); sti: stishovite (dark spots). Note that magnification in (a) is 1000× and in (b) is 1800×.

Fig. 1.

Backscattered electron images of run products showing textural differences between experiments performed on similar bulk compositions and run durations, but at extremely different pressure and temperature conditions. (a) graphite-bearing, fluid-saturated model MORB at 2.7 GPa, 730 °C, 224 hours (courtesy Ada Crottini); (b) graphite-bearing, fluid-saturated natural MORB at 19 GPa, 1200 °C, 120 hours (courtesy Kazuaki Okamoto). gar: garnet; cpx: clinopyroxene; ky: kyanite; maj: majorite; ca-pvs: Ca-perovskite (bright spots); sti: stishovite (dark spots). Note that magnification in (a) is 1000× and in (b) is 1800×.

This review will focus mainly on experimental studies both on complex model systems close to real rocks and on natural materials, in order to facilitate the approach to the world of ultrahigh pressure rocks for those who are not familiar with the variety of new crystalline phases found in the last decade.

“Fluid” phases at ultrahigh pressure conditions

A few definitions

Because of the global geochemical and geodynamic implications, there has been an increasing interest in the role and features of fluids at ultrahigh pressure conditions. Even though various classical textbooks have regarded the eclogite facies as a “dry” environment, there is an increasing consensus that ultrahigh pressures rather represent “wet” conditions, at least when a prograde pressure path is attained.

As a first step it is important to recall definitions which will be useful in the following discussion.

  • H2O: it is a chemical component, it represents the stoichiometry of a pure compound (water), but it often also indicates a chemical species in complex fluids. The definition of component H2O in a system does not imply any specific value for its chemical potential.

  • water: a fluid composed of H2O only. The chemical potential of H2O will be maximum, i.e. it implies H2O saturation and unit activity for H2O. Though unrealistic at geological conditions, water is the reference fluid in most thermodynamic calculations and in numerical simulations.

  • fluid: a volatile-rich phase, showing supercritical behaviour, characterised by a poorly polymerised structure. Fluid saturation implies the appearance of a fluid, i.e. it defines a vaporous surface, but it does not imply a univocally defined chemical potential of any of the species present in it, e.g. H2O, CO2 or O2.

    From the previous definitions, we obtain:

  • “dry” conditions (“water absent”, Type I in Robertson & Wyllie, 1971): a H2O-free system.

  • “wet” conditions: a fluid-bearing system, where the abundance of H2O is enough to saturate also a silicate liquid, e.g. on a wet solidus (“water excess”, Type IV in Robertson & Wyllie, 1971).

  • H2O-deficient and fluid-absent (“water deficient” in Yoder, 1952; “water deficient and vapor-absent”, Type II in Robertson & Wyllie, 1971): an assemblage of minerals including hydrates but not a free fluid. It is a condition of H2O undersaturation.

  • H2O-deficient and fluid-present (“water deficient and vapor-present”, Type III in Robertson & Wyllie, 1971): an assemblage of minerals plus a free fluid, but there is insufficient H2O (and fluid) present to saturate the liquid when the assemblage is completely melted (see Figure 2 in Robertson & Wyllie, 1971).

  • aqueous fluid: is a fluid mainly composed of H2O plus other volatile species and/or dissolved solids.

  • carbonic fluid: is a fluid mainly composed of C species plus other volatiles and/or dissolved solids.

  • mixed fluid: often refers to a fluid with mixed volatile species, usually mainly H2O and CO2.

Fig. 2.

Sequence of reactions involving hydrates and carbonates in the model systems MgO-SiO2-H2O (a-d) and MgO-CaO-CO2 projected from silica (e-f). Assemblages below the tie-lines brucite-antigorite-talc-quartz in (a) and (b) represent H2O-undersaturated fluid absent conditions; assemblages below the tie-lines magnesite-dolomite-aragonite are CO2-undersaturated. Points 1 and 2 stand for two hypothetical ultramafic bulk compositions. arag: aragonite; atg: antigorite; br: brucite; cc: calcite; di: diopside; dol: dolomite; en: enstatite; fo: forsterite; mag: magnesite; p: periclase; q: quartz; ta: talc; wo: wollastonite. See text for explanation.

Fig. 2.

Sequence of reactions involving hydrates and carbonates in the model systems MgO-SiO2-H2O (a-d) and MgO-CaO-CO2 projected from silica (e-f). Assemblages below the tie-lines brucite-antigorite-talc-quartz in (a) and (b) represent H2O-undersaturated fluid absent conditions; assemblages below the tie-lines magnesite-dolomite-aragonite are CO2-undersaturated. Points 1 and 2 stand for two hypothetical ultramafic bulk compositions. arag: aragonite; atg: antigorite; br: brucite; cc: calcite; di: diopside; dol: dolomite; en: enstatite; fo: forsterite; mag: magnesite; p: periclase; q: quartz; ta: talc; wo: wollastonite. See text for explanation.

How likely is fluid saturation at high pressure?

Whenever hydrates or carbonates are present in a system (as in the simple model chemographies depicted in Fig. 2), subsolidus compositional diagrams are separated into two regions, the first where assemblages are fluid-saturated, i.e. a fluid occurs together with solids, and the second one where hydrates and/or carbonates occur, but no fluid is present (H2O or CO2-deficient). In Figure 2a, the mineral assemblages brucite-antigorite-fluid, antigorite-talc-fluid and talc-quartz-fluid represent fluid-saturated assemblages, whereas assemblages below the tie lines brucite-antigorite, antigorite-talc, talc-quartz are H2O-deficient, fluid absent. Similarly, assemblages below the tie-lines magnesite-dolomite-calcite in Figure 2e are CO2-deficient and fluid absent. Such tie-lines define a “fluid-saturation surface” which separates fluid saturated from fluid absent assemblages. The fluid saturation surface is a line in simple ternary compositional diagrams while a plane in quaternary diagrams.

Fluid absent (H2O-conservative) reactions occur below such tie-lines, e.g. the reaction forsterite + talc = enstatite + antigorite (Fig. 2a to 2b) or the reaction enstatite + dolomite = magnesite + diopside (Fig. 2e to 2f). These reactions do not lead to dehydration or decarbonation because no H2O or CO2 components are liberated.

Because low-temperature, low-pressure metamorphism commonly promotes the formation of hydrates that contain large amounts of H2O (e.g. antigorite, talc, chlorite, clay minerals, lawsonite etc.), the likelihood of H2O-deficient assemblages is maximised at such conditions. Let us assume two partially serpentinised rocks from the ocean floor, a harzburgite and an orthopyroxenite (points 1 and 2 in Fig. 2a). For this example, they will store the same amount of H2O in the assemblage antigorite + talc + forsterite. With increasing pressure and temperature, the reaction forsterite + talc = enstatite + antigorite will first lead to two different hydrate-bearing assemblages (Fig. 2b). The reaction antigorite + talc = enstatite + H2O will cause fluid saturation in the orthopyroxenite and consequently dehydration, whereas the harzburgite will remain at H2O-deficient conditions (Fig. 2c). Further dehydration will be caused by the breakdown of antigorite to forsterite + enstatite + H2O. Only at this stage does a fluid appear in the harzburgite (Fig. 2d). This reaction sequence implies that the “fluid saturation surface” progressively moves away from the volatile component, as a function of the hydrates stable at the various P-T conditions of interest. Therefore there is no straightforward relationship between “amount of volatiles” in the rock and “amount of fluid”; it is the stable phase assemblage that governs how H2O is partitioned. From Figure 2 it is evident that:

  • at fixed maximum H2O contents, the attainment of fluid saturation, i.e. attainment of conditions where the assemblage is on or above the fluid saturation surface, will be promoted by an increase in pressure and temperature;

  • fluid removal at constant P and T (and therefore phase assemblage) will never lead to H2O deficiency, simply because there is no process able to remove a fluid below the fluid saturation surface, i.e. where the fluid does not exist.

However, because the appearance of a fluid is related to the thermodynamic definition of saturation, the two previous statements do not represent a paradigm. Let us first consider a schematic block diagram where the chemical potential of the volatile species is plotted on the z axis, and P and T on the other two axes (Fig. 3). The chemical potential for the volatile species, e.g. H2O, will be maximum when a fluid, composed of this species only, appears. Any phase assemblage, or reaction, occurring below the fluid saturation line previously discussed, necessarily plots in the (physically accessible) half-space below the surface which describes the equation of state for the pure fluid, e.g. water. Figure 3 shows that with increasing pressure and temperature, along a steep geothermal gradient (as it is the case for a number of UHP terrains), the surface representing the chemical potential of H2O buffered by antigorite + forsterite + enstatite progressively approaches the fluid saturation surface. The intersection between these two surfaces corresponds to the univariant breakdown reaction of antigorite and to the appearance of a fluid, in agreement with the chemographies shown in Figure 2.

Fig. 3.

Saturation surfaces for water and carbon dioxide compared to chemical potentials of H2O and CO2 buffered by antigorite-enstatite-forsterite and by magnesite-quartz-enstatite, respectively. Volatile saturation and therefore fluid release occurs if the P–T path drives μH2O or μCO2 toward the thermodynamic surfaces of H2O or CO2 fluid, respectively. Note that μi, = molar G (from Poli & Schmidt, 2002).

Fig. 3.

Saturation surfaces for water and carbon dioxide compared to chemical potentials of H2O and CO2 buffered by antigorite-enstatite-forsterite and by magnesite-quartz-enstatite, respectively. Volatile saturation and therefore fluid release occurs if the P–T path drives μH2O or μCO2 toward the thermodynamic surfaces of H2O or CO2 fluid, respectively. Note that μi, = molar G (from Poli & Schmidt, 2002).

As we consider CO2 fluids we observe that, similarly to the previous case, the appearance of a carbonic fluid from a CO2-deficient assemblage, e.g. magnesite + enstatite + quartz, is controlled by the intersection of the two surfaces defining the chemical potential of CO2 in the solid assemblage and in the fluid. However, in this example, as a result of the large dG/dP for a pure CO2 fluid and of the relative location of the surface for the assemblage magnesite + enstatite + quartz, a relatively “flat” dP/dT slope is obtained for the decarbonation reaction responsible for the generation of carbonic fluid. The same prograde P–T path will therefore promote devolatilisation of the solid assemblage as far as H2O is concerned, but conversely it will favour fixation of CO2 in the solid because decarbonation would be rather driven by a moderate pressure increase relative to the temperature increase.

When both H2O and CO2 are present, the free energy of the mixed fluid is not a single value at P and T, hence it cannot be plotted as a simple surface in the diagram of Figure 3. Nevertheless a similar line of reasoning can be applied in this more complex case.

Again, a chemographic analysis is useful to show some counterintuitive implications of mixed volatile systems. In the triangle of Figure 4 we represent the system CaO–Al2O3–SiO2–H2O–CO2 projected from kyanite and coesite and for each three-phase and two-phase assemblage the chemical potential of H2O is reported. Such calculation demonstrates two interesting points. First, at the same pressure and temperature conditions, the chemical potential of H2O can be identical in fluid-absent conditions and in the presence of a fluid. Second, the addition of CO2 to a fluid-absent mineral assemblage (point A in Fig. 4) may lead first to the crystallisation of more hydrous minerals (e.g. lawsonite and zoisite in B) and then to the appearance of an aqueous fluid in C. Nevertheless, it has been common practice to assume identity between the chemical potential of a volatile component and the chemical potential of that volatile in the fluid, when multi-equilibrium calculations (e.g. TWEEQU, Berman, 1991) are used to derive the composition of metamorphic fluids. However, the actual presence of a mixed fluid, rather than a condition of H2O deficiency and fluid absence, should be carefully demonstrated.

Fig. 4.

Composition diagram for the system CaO-Al2O3-SiO2-H2O-CO2 projected from kyanite and coesite (from Poli & Schmidt, 1998, modified) which shows occurrence of hydrates and carbonates at fluid-absent conditions, equality of chemical potential of H2O for selected fluid present and fluid absent conditions and possible appearance of an aqueous fluid by addition of CO2 to the system.

Fig. 4.

Composition diagram for the system CaO-Al2O3-SiO2-H2O-CO2 projected from kyanite and coesite (from Poli & Schmidt, 1998, modified) which shows occurrence of hydrates and carbonates at fluid-absent conditions, equality of chemical potential of H2O for selected fluid present and fluid absent conditions and possible appearance of an aqueous fluid by addition of CO2 to the system.

In conclusion, even in such very simple model systems, the prediction of fluid saturation is not an easy task. It obviously depends on the thermodynamic properties of the solid assemblages compared to the properties of the possible fluids in a system. The worst and most common condition is represented by rocks where components C, O and H are independent variables. This is the case, as an example, when hydrates, carbonates and other mineral phases coexist in a Fe-bearing system. It is beyond the goal of this contribution to review all of the complexities present in C-O-H bearing rocks (e.g. see Connolly, 1995). Nevertheless it is important to point out that the appearance of a fluid is related to a complex interplay of mass balance and thermodynamic constraints because variable amounts of C, H and O can be stored in carbonates, hydrates and ferrous/ferric minerals as well as in fluids.

Fluids, melts and the 2nd critical endpoints

Even though in most thermodynamic calculations performed to extract P-T paths, fluids are supposed to be composed of volatile species only, it has been known since a long time that fluids at geological conditions dissolve variable amounts of solids which are far from negligible.

It is widely accepted that temperature and, secondarily, pressure promote dissolution. However, there is still an open debate how and how much (Stalder et al., 2001; Mibe et al., 2002) solids are dissolved. Because of intrinsic experimental difficulties, there is still a very limited amount of systematic work devoted to determine the vaporous surface in the different chemical systems. Even though petrologists are used to think about the variability of the solidus and the liquidus surfaces as a function of the buffering assemblage in P-T-composition space (e.g. the solidus for a gabbro is different from the solidus for granite), there has been a significant underestimation of the possible variability of the vaporous surface, i.e. of the surface limiting the fluid phase field (Boettcher & Wyllie, 1969). As an example, Figure 5 represents the vaporous surface in the system MgO-SiO2-H2O determined by Ryabchykov et al. (1983). At fixed P-T conditions, the amount of silica dissolved in the fluid will be maximum in assemblages coexisting with coesite, but in an assemblage with orthopyroxene will depend on the bulk composition chosen.

Fig. 5.

The vaporous surface in the system MgO-SiO2-H2O at 3 GPa after Ryabchykov et al. (1983).

Fig. 5.

The vaporous surface in the system MgO-SiO2-H2O at 3 GPa after Ryabchykov et al. (1983).

It is therefore evident that in complex, high variance systems the amount of silica present in fluids and/or hydrous melts can be much different than in the reference system SiO2-H2O. It can be lower, as shown in Figure 5, or even higher, as in the granite-H2O system, where the hydrous liquid has a higher SiO2 content (e.g. at 0.4 GPa, 900 °C) than a liquid in the quartz-H2O system. Though it might seem a trivial statement, it should be remembered that a fluid is a solution and therefore its actual composition will depend upon the buffering phase assemblage.

Because of the increasing amount of dissolved solids present in fluids at increasing pressure and temperature and because of the mutually increasing solubility of H2O in melts, it has been found that the “miscibility gap” between fluids and melts shrinks with increasing pressure (Fig. 6, after Stalder et al., 2000). A number of experimental studies have shown that the critical point between fluids and melts can be attained under geological conditions (e.g. Bureau & Keppler, 1999), even though the amount of volatile components required to encounter such a phenomenon is expected to be unusual. However, shrinking of the solvus with increasing pressure rather implies a migration of the critical point toward lower temperature, as shown in Figure 6. When the critical curve intersects the reaction describing the wet solidus, the two volatile-bearing phases (fluid and melt) are replaced by one single “supercritical” phase. As a consequence, a singularity results: the wet solidus vanishes, the concept of melting looses its definition, and the breakdown of solid assemblages is replaced by a concept of continuous dissolution in a first volatile-rich, and then silicate-rich, “non-solid” phase. This singularity is called 2nd critical endpoint and it is of utmost importance in ultrahigh pressure environments.

Fig. 6.

Schematic presentation of a binary system solid + H2O (from Stalder et al., 2000, copyright Mineralogical Society of America). P2 > P1.

Fig. 6.

Schematic presentation of a binary system solid + H2O (from Stalder et al., 2000, copyright Mineralogical Society of America). P2 > P1.

As shown by Boettcher & Wyllie in their classical paper dated 1969, in a multi-component chemical system, there is more than one 2nd critical endpoint, and, again, the relative location of the vaporous, solidus and liquidus surfaces are related by the coexisting buffering assemblages (Fig. 7, from Boettcher & Wyllie, 1969). In the simple system SiO2-H2O the second critical endpoint is at only 1 GPa and 1100 °C (Kennedy et al., 1962), in albite-H2O and haplogranitic systems moves to ca. 1.5 GPa (Bureau & Keppler, 1999; Stalder et al., 2000), in CaO-SiO2-H2O-CO2 to 3.2 GPa and 500 °C (Boettcher & Wyllie, 1969), whereas in the ultramafic model system MgO–SiO2–H2O (MSH) the solidus terminates at ca. 12 GPa and 1100 °C (Stalder etal., 2001).

Fig. 7.

Isobaric, polythermal diagrams for the hypothetical ternary system MO-SiO2-H2O (from Boettcher & Wyllie, 1969). Dashed lines are isothermal field boundaries. Dotted lines connect temperature maxima on the liquidus and vaporous field boundaries with the crystalline phase with which the liquid and vapour are in equilibrium. Reprinted with permission from Elsevier.

Fig. 7.

Isobaric, polythermal diagrams for the hypothetical ternary system MO-SiO2-H2O (from Boettcher & Wyllie, 1969). Dashed lines are isothermal field boundaries. Dotted lines connect temperature maxima on the liquidus and vaporous field boundaries with the crystalline phase with which the liquid and vapour are in equilibrium. Reprinted with permission from Elsevier.

In the system quartz + H2O, silica dissolved in the fluid ranges from a few weight percent at 500–700 °C to a maximum of 12 wt% at 1 GPa and 900 °C. In MSH at 6 GPa, the amount of silicate dissolved in the fluid does not exceed ca. 10 wt% up to temperatures close to the solidus (ca. 1150 °C). Experiments performed by Schneider & Eggler (1986) suggest quantities of solute from 3 wt% in an amphibole peridotite at 1.5–2.0 GPa, 750–900 °C to 12–15 wt% in a phlogopite peridotite at 1.3–2.0 GPa, 1100°C. Addition of CO2 to the fluid strongly decreases the solubility of silicates, by ca. one order of magnitude.

As a conclusion, we emphasise that the inadequacy of current knowledge on thermodynamic properties of UHP fluids strongly suggests the use of fluid-absent reactions for geothermobarometric purposes.

Ultrahigh pressure rocks: The quartz-coesite transformation and the “internally consistent” thermodynamic databases

UHP conditions refer to rocks which underwent recrystallisation within (or above) the stability field of coesite. It is therefore of primary importance that the position of the reaction quartz-coesite is still debated, especially in the relatively low temperature region. The transformation quartz-coesite is of particular interest not only as a pressure reference in high pressure experimental studies or as a lower pressure limit for UHP terrains, but it also indirectly influences the thermodynamic properties of most phases tabulated in the so-called “internally consistent” databases. In fact the concept of “internal consistency” is related to the high correlation between all of the variables contained in a very large over-determined system of non-linear equations whose dimensionality is controlled by dimensionality of the chemical components introduced. Conversely, thermodynamic properties of quartz and coesite (and therefore the calculated location of quartz-coesite transformation) will be controlled not only by the experiments performed in the pure system SiO2, but also by all of the reactions which include either quartz or coesite or both, in a variety of assemblages belonging to a very large chemical system (up to more than 12 components in the latest release of the database by Holland & Powell, 1998).

Despite the rigorous approach adopted in the internally consistent databases, optimisation routines (least squares, mathematical programming, Bayes method) may often obscure the basic quality/accuracy/precision of the experimental data behind the recalculation scheme. The quartz to coesite transformation is a typical example of this problem.

Figure 8 shows the experimentally derived position of this reaction from 400 °C to 1400 °C compared to the calculated equilibrium according to the database of Holland & Powell (1998) and updates. There are still major discrepancies between the different experimental studies, which cannot be simply solved by a temporal selection. Two of the most recent studies (Bose & Ganguly, 1995 and Hemingway et al., 1998) actually show the largest deviation, both in the absolute position in P-T space and in dP/dT slope. Such a deviation, as inherited by experimental difficulties, is much larger at low temperature and as large as 0.2 GPa at 600 °C. Recently, Walter et al. (2002) obtained an in situ determination of quartz to coesite at only 3 GPa and 1350 °C. Such a position would increase uncertainty to more than 0.4 GPa, a huge number when translated to geodynamic applications.

Fig. 8.

Comparison between different experimental studies on the transformation quartz-coesite and the calculated location according to the database of Holland & Powell (1998 and recent updates). Determination by Bohlen & Boettcher (1982) and Mirwald & Massonne (1980) (reported by Hemingway et al., 1998) and Bose & Ganguly (1995) were all performed in a piston cylinder apparatus. The data by Walter et al. (2002) refer to an in situ estimate performed under a synchrotron source.

Fig. 8.

Comparison between different experimental studies on the transformation quartz-coesite and the calculated location according to the database of Holland & Powell (1998 and recent updates). Determination by Bohlen & Boettcher (1982) and Mirwald & Massonne (1980) (reported by Hemingway et al., 1998) and Bose & Ganguly (1995) were all performed in a piston cylinder apparatus. The data by Walter et al. (2002) refer to an in situ estimate performed under a synchrotron source.

The calculated curve mainly lies in the uppermost region of the experimental range, and below 600 °C it deviates significantly from currently available data. Though it is beyond the goal of this contribution to discuss the details of the experimental problems or of the limits of the optimisation routines, it should be noted that differences in the order of a few hundreds of MPa represent a major uncertainty in terms of the resolution of P-T paths from UHP terrains.

Ultrahigh pressure rocks: The graphite-diamond transformation

A somewhat different uncertainty affects the transformation of graphite to diamond. The crystallisation of diamond was proved in a variety of media, from molten metals (Sung & Tai, 1997), to sulfur (Sato & Katsura, 2001), carbonates (Pal’yanov et al., 2002), carbonatitic melts (Arima et al., 2002), and C-O-H fluids (Sokol et al., 2001; Yamaoka et al., 2002 and references therein). The lowest P, T conditions for diamond nucleation were found at 5.7 GPa and 1150 °C in an alkaline carbonate C-O-H fluid (Pal’yanov et al., 1999). The currently accepted equilibrium curve for the graphite-diamond transformation still relies on the reversal experiments by Kennedy & Kennedy (1976), performed in the range 1100-1600 °C, and on physical properties of graphite and diamond summarised in Bundy et al. (1996). However, it should be noted that most UHP diamond-bearing terrains attained temperature conditions far below the experimental conditions investigated to date, because of the kinetic constraints imposed by the sluggishness of this transformation. The possible role of fullerene structures (Sundqvist, 1999) at low temperatures in geologically relevant conditions has to be entirely explored.

Mafic systems

Bimineralic mafic eclogites and the relevance of minor and accessory phases

Garnet and clinopyroxene often constitute more than 90% of the mode of mafic eclogites at ultrahigh pressure conditions. Even though such abundance controls relevant physical properties, first of all density, minor and accessory phases are as important as major minerals because of their influence on a variety of geological processes.

As examples, we know that physical and chemical properties of rocks, including brittle failure and creep, initiation of melting, generation of volatile-rich magmas responsible for explosive volcanism, geochemical signatures etc., are mostly influenced by the presence of minor (typically hydrates and carbonates) and accessory phases (typically rutile, allanite, zircon, phosphates, ellenbergerite).

A first relevant, experimentally documented example concerns the stability field of allanite in mafic eclogites to ca. 4 GPa (Hermann, 2002a). More than 90% of LREE and Th is incorporated in allanite, and allanite is found to be a residual phase during melting. Therefore the geochemical signature of liquids eventually produced in a subducting slab will be strongly controlled by the presence of allanite and by reactions responsible for its breakdown. Furthermore, Hermann (2002a) shows that more than 95% of Ti, Nb and Ta is partitioned in rutile, 95% of Zr and Hf in zircon and P in apatite (Fig. 9, from Hermann, 2002a). Phengite, at a modal value of ca. 10%, incorporates more than 95% of the bulk rock Rb, Ba, and Cs. Such figures clearly demonstrate that careful geochemical analysis cannot leave aside such accessory phases.

Fig. 9.

Trace element distribution among the phases of eclogites (from Hermann, 2002a). Modal amounts are 36% omphacite, 35% garnet, 14% quartz, 12% phengite, 2.5% rutile, 0.52% apatite, 0.1% allanite, 0.032% zircon. Reprinted with permission from Elsevier.

Fig. 9.

Trace element distribution among the phases of eclogites (from Hermann, 2002a). Modal amounts are 36% omphacite, 35% garnet, 14% quartz, 12% phengite, 2.5% rutile, 0.52% apatite, 0.1% allanite, 0.032% zircon. Reprinted with permission from Elsevier.

Chopin et al. (1986) and Brunet et al. (1998) have shown that ellenbergerites and phosphates provide important complementary information to geothermobarometric reconstruction of eclogites. Schmidt (1996), Ono (1998), Domanik & Holloway (1996), Okamoto & Maruyama (1998) found that melting at UHP is controlled by phengite, which is expected to be the only hydrous phase on the solidus of both subducted sediments and mafic eclogites. A number of geophysical studies proposed that intermediate depth earthquakes can be ascribed to a process of dehydration embrittlement.

Despite their importance, most reactions involving minor and accessory minerals are still in a state of preliminary investigation because a clear gap exists between the chemical complexity of real rocks and model systems studied in the laboratory.

Phase relationships in H2O-bearing systems

Most experimental studies performed on mafic systems focussed on phase relationships in hydrated MOR basalt compositions. Because of the large number of chemical components present in these rocks, continuous reactions are the relevant mechanism controlling the appearance and the abundance of mineral phases. Hence, phase transitions observed in MORB + H2O cannot be extrapolated to the universe of gabbroic rocks found in crustal sections (continental or oceanic), often recorded in UHP terrains, due to the strong effects of variable bulk compositions.

Figure 10a displays current knowledge to pressures of about 6 GPa. Most experimental studies are substantially consistent and minor, though relevant, discrepancies can be ascribed to different experimental set-ups. All of the experimental studies agree that amphibole breakdown in the system MORB + H2O, both on a simplified chemical composition (Poli, 1993; Schmidt & Poli, 1998) and on natural starting materials (Pawley & Holloway, 1993; Liu et al., 1996) is located at approximately 2.4 GPa between 650 °C and 750 °C. At such conditions amphibole is barroisitic and because of the low reaction rates in the subsolidus region, there is still no information on the mutual relationships between sodic-calcic and sodic amphiboles. Though it has been shown that, in principle, the glaucophane stability field intersects the coesite field (Carman & Gilbert, 1983), there is no experimental demonstration that amphiboles in complex, Ca-bearing systems extend to more than 2.4 GPa.

Fig. 10.(a)

Review of the experimentally determined phase relationships in MORB eclogites to 5 GPa in the presence of an aqueous fluid. (b) Schematic displacement of invariant point “CaAl” in (a) as a function of the stable garnet composition, which, in turn, is a function of the bulk composition adopted.

Fig. 10.(a)

Review of the experimentally determined phase relationships in MORB eclogites to 5 GPa in the presence of an aqueous fluid. (b) Schematic displacement of invariant point “CaAl” in (a) as a function of the stable garnet composition, which, in turn, is a function of the bulk composition adopted.

The temperature region below 650 °C is essentially unexplored, and probably unexplorable, because, despite the presence of a free fluid, even experiments as long as one month do not show appreciable approach to equilibrium conditions. As a consequence we have only a very rough idea of reactions involving chlorite. Nevertheless, because most mafic compositions at high pressure are silica-saturated, as a result of the albite breakdown to form jadeite component and quartz, chlorite breakdown with pressure should be located at pressures comparable to amphibole breakdown and it is not expected to reach the UHP field. On the contrary, olivine-saturated compositions might show a completely different pattern for chlorite (see section devoted to ultramafic systems).

Lawsonite and epidote group minerals, which have been known as typical hydrous minerals for relatively shallow hydrothermal conditions, show some of the most interesting relationships in UHP metamorphism. Even though there has been some scepticism about the actual presence of lawsonite in ultrahigh pressure eclogites, recent findings of coesite in the famous lawsonite eclogite xenoliths from the Four-Corner Region (Colorado Plateau, Usui et al., 2003) demonstrate the relevance of these minerals even in tectonic settings that are not expected to be most suitable for such low-temperature assemblages. The importance of lawsonite is also related to the very high amount of H2O stored in its structure, which makes it of primary interest for geophysicists and geochemists modelling the geodynamics of subduction zone environments.

The presence of lawsonite in mafic rocks is controlled by a continuous reaction, with steep positive dP/dT slope, which consumes lawsonite (and clinopyroxene) to form either a relatively grossular-rich garnet (Poli & Schmidt, 1995; Okamoto & Maruyama, 1999), or zoisite (Poli & Schmidt, 1995) at pressures above or below the (pseudo-)invariant point located at ca. 3.2 GPa and 680 °C. This point, as demonstrated by the comparison between different experimental studies, is very sensitive to the bulk composition chosen (Fig. 10b). Lawsonite temperature stability and, secondarily, zoisite pressure stability apparently increase as bulk chemistry moves toward more aluminous compositions. As an example, lawsonite extends its stability to ca. 850 °C at 6 GPa in intermediate (“andesitic”) systems (Poli & Schmidt, 1995) and to more than 900 °C in peraluminous sedimentary compositions (Domanik & Holloway, 1996). It is worth to recall that the invariant point in pure CASH between lawsonite, zoisite, grossular, kyanite and silica lies at about 7 GPa and 1000 °C (Poli & Schmidt, 1998). As a consequence we might expect that the stability of lawsonite and epidote group minerals should be significantly extended in a number of troctolitic gabbros compared to normal MORB compositions. However it should be noted that Liu et al. (1996) did not find either epidote or zoisite in a large pressure range, though lawsonite was observed. Such observation does not seem to be consistent both with the other high pressure experimental studies (Pawley & Holloway, 1993; Poli, 1993) and with natural occurrences, because epidote is found in a variety of eclogite facies rocks. Absence of epidote in the experiments of Liu et al. (1996) might be related to conditions in oxygen fugacity but the inconsistency remains unresolved.

Current knowledge about the stability and the phase relations involving hydrous magnesian silicates such as talc, chloritoid and Mg-staurolite is still extremely vague. Even though Pawley & Holloway (1993) and Poli (1993) observed growth of Mg-chloritoid, Liu et al. (1996) and Forneris & Holloway (2001) proposed the metastability of this phase on the basis of long duration reversal experiments (up to almost 100 days long). Since Mg-chloritoid is occasionally found in metamorphosed high pressure Mg-gabbros, one possible explanation for this apparent discrepancy is the bulk composition chosen, which is much higher in normative olivine in Poli & Schmidt (1995) and Pawley & Holloway (1993). Another possible, additional reason, is a drift in reactive bulk composition with time. Because Mg-rich minerals are present in the starting materials used (e.g. tremolite in Poli, 1993), nucleation of Mg-chloritoid might be promoted on these localised chemical systems, similarly to natural rocks, but diffusion with time might lead to resorption of previously formed crystals. Further experiments on Mg-enriched compositions should be performed to demonstrate this hypothesis and to verify the stability of Mg-chloritoid in mafic rocks at UHP conditions.

Similarly, talc, though not ubiquitous in experimental products, is expected to play an important role in differentiated gabbros and, again, it is expected to be sensitive to the bulk Mg/(Mg + Fe) in the system. The fate of talc with pressure is closely related to the stability of the 10 Å phase chemically analogous to talc, which will be discussed in more detail in the section devoted to ultramafic systems.

Finally, we mention the occurrence of Mg-staurolite in various high pressure studies at conditions close to the wet solidus of mafic systems. Despite the apparent rarity of Mg-staurolite in natural rocks, its frequency is remarkable in quite different experimental configurations, which might suggest that this mineral has been overlooked in nature.

When enough potassium is present in the system (this can be the case for altered oceanic basalts or for impure mafic rocks derived from sedimentary sequences) phengite is a characteristic ultrahigh pressure mineral. Schmidt (1995) and Schmidt & Poli (1998) described the stability of phengite in both mafic and intermediate systems where the celadonite buffering assemblage is garnet + omphacite + silica. For such an assemblage a progressive increase in Si with pressure and decrease with temperature is observed. Celadonite end-member composition is achieved at ca. 9 GPa close to phengite s.s. breakdown to K-hollandite. Phengite is the hydrous mineral with the largest stability field observed in mafic systems and because it is also the only hydrate on the solidus above the epidote breakdown, phengite controls melting and geochemical signatures of first melts. Unfortunately, despite considerable effort performed both in model (e.g. Massonne & Szpurka, 1997; Hermann, 2002b) and in more complex systems (Schmidt & Poli, 1998; Schmidt, 1996; Okamoto & Maruyama, 1998), there is still no complete formulation of a phengite geothermobarometer available for mafic eclogites at ultrahigh pressure conditions.

Phase relationships in a C-O-H-bearing system

The effect of the presence of a C-O-H mixed fluid on phase relationships in mafic systems at high pressure subsolidus conditions are substantially unknown to date. Only a few reconnaissance works on simple systems (see Luth, 1999, and references therein) or on more complex systems approaching natural rocks (Yaxley & Green, 1994; Molina & Poli, 2000) are currently available.

Experimental difficulties arise from the extremely complex interplay of mass balance constraints relating the solid assemblage (including hydrates and carbonates), the amount of fluid present, the speciation of the fluid, the iron oxidation state, the physical state of carbon (graphite/diamond vs. carbon-bearing volatiles) etc.

Phase relations at UHP conditions are further complicated by the enlargement of the stability field of graphite/diamond with increasing pressure and/or decreasing temperature in a diagram log fO2 vs. P or T. The stippled field in Figure 11 shows the potential coexistence of graphite and carbonates at 680 °C and 1300 °C. This field is realistically bounded at high oxygen fugacities by the equilibrium graphite + O2 = CO2 (GCO), whereas at low oxygen fugacity by the ultimate decomposition of carbonates to form graphite + periclase + silicates. If oxygen fugacities represented by the equilibrium NNO (Ni–NiO) or FMQ (fayalite-magnetite-quartz) are assumed to be a middle course of conditions at high pressure, calculations in Figure 11 suggest that phase assemblages of UHP terrains are expected to present either graphite/diamond only or graphite/ diamond + carbonates or carbonates only as a function of the reactive bulk composition, i.e. of the relative availability of carbon, hydrogen and oxygen.

Fig. 11.

The stability of graphite (gph) with carbonates (stippled field) at high and low temperature conditions. FMQ: fayalite-magnetite-quartz; FFsM: fayalite-ferrosilite-magnetite; NNO: nickel-bunsenite; GCO: graphite-CO2; EMOG: enstatite-magnesite-olivine-graphite; DSDG: dolomite-quartz/coesite-diopside-graphite. See text for explanation.

Fig. 11.

The stability of graphite (gph) with carbonates (stippled field) at high and low temperature conditions. FMQ: fayalite-magnetite-quartz; FFsM: fayalite-ferrosilite-magnetite; NNO: nickel-bunsenite; GCO: graphite-CO2; EMOG: enstatite-magnesite-olivine-graphite; DSDG: dolomite-quartz/coesite-diopside-graphite. See text for explanation.

Molina & Poli (2000) performed a series of piston cylinder experiments to 2.0 GPa and temperatures to 730 °C on a tholeiite composition in the presence of a mixed fluid produced by the decomposition of oxalic acid dihydrate, at oxygen fugacities externally buffered by NNO. Amphibole was found to coexist with calcite at P ≤ 1.4 GPa, with dolomite at 1.4 ≤ P ≤ 1.8 GPa, and with dolomite + magnesite at pressures higher than 1.8 GPa (Fig. 12). Garnet, paragonite, kyanite and epidote participate to complex reactions with carbonates buffering the fluid. Estimates of the coexisting fluid compositions, on the basis of mass balance and thermodynamic calculations, demonstrate a continuous H2O enrichment with increasing pressure and decreasing temperature. An almost purely aqueous fluid is obtained at 2 GPa and 665 °C. This implies that carbon tends to fractionate into the solid with increasing pressure. However, the nature of such a solid assemblage, as previously stated, cannot be straightforward predicted.

Fig. 12.

Review of the experimentally determined phase relationships in MORB eclogites to 6 GPa in the presence of a C-O-H mixed fluid at variable oxygen fugacities and bulk C-O-H contents.

Fig. 12.

Review of the experimentally determined phase relationships in MORB eclogites to 6 GPa in the presence of a C-O-H mixed fluid at variable oxygen fugacities and bulk C-O-H contents.

For this purpose it is useful to compare the results obtained by Yaxley & Green (1994) with assemblages found by Molina & Poli (2000) and by Crottini et al. (2002) (Fig. 12). Even though the stability of hydrous phases demonstrates a substantial consistency between these different works, carbon-bearing phases are extremely variable. Amphibole breakdown in C-O-H bearing MORB is found at 2.5-2.6 GPa and the fluid saturated solidus is located at ca. 730 °C at 2.2 GPa. These observations unequivocally indicate that the coexisting fluids have to be enriched in H2O component, despite the variable amounts of carbon introduced in the system for the variable experimental strategies adopted in these studies. On the contrary, Yaxley & Green (1994) observes the appearance of dolomite only above 2.5 GPa whereas Crottini et al. (2002) in the uppermost pressure region obtained graphite only, aragonite + dolomite + graphite or aragonite + graphite as a function of both the oxygen buffer and of the amount of C-O-H added to the system. Given currently available experiments it is therefore a challenging task to predict carbonate and hydrate-bearing assemblages at UHP conditions unless specific experiments are performed. It is worth noting that theoretical calculations are of limited help in such a case, because most petrologists assume that C-O-H mixed fluids can be approximated along the join H2O–CO2 (Kerrick & Connolly, 2001) which is evidently not the case, as graphite or diamond are widely recognised around the world in ultrahigh pressure terrains.

A few experimental studies in very simple model systems offer a useful reference frame, even though application to natural complex systems should be cautious. As an example the breakdown of dolomite with pressure (Luth, 2001) can be applied to a variety of bulk compositions, from ultramafic to mafic and intermediate eclogites. However, as shown in Figure 12 (compare dolomite-in in CMS and in experiments of Yaxley & Green, 1994), such model reactions may differ enormously from natural systems in P–T location because carbonate species are controlled by a variety of mass balance and thermodynamic constraints, first fluid speciation.

Ultramafics

Peridotite compositions

Although peridotites are typically characterised by a rather simple phase assemblage constituted of olivine, orthopyroxene, clinopyroxene and an Al-phase, i.e. plagioclase, spinel or garnet depending on pressure (see Ulmer & Trommsdorff, 1999 for a review on mantle mineralogy in simple systems), mantle phases may partake to complex solid solutions by considering both compositions approaching natural peridotites, i.e. complex systems and natural materials. As a result phase equilibria are mainly governed by continuous reactions and, again, strongly depend on bulk compositions.

Model natural peridotite compositions are mainly derived by variably enriched Hawaiian xenoliths (spinel and garnet lherzolites from Hawaiian xenoliths; Mysen & Boettcher, 1975), fertile spinel lherzolites (KLB-1; Hirose & Kushiro, 1993) and peridotites which are depleted in incompatible elements, but not depleted in Ca and Al such as the Tinaquillo lherzolite (Robinson & Wood, 1998). Furthermore Green (1973) calculated a model theoretical peridotite composition, the MORB pyrolite, combining a primitive basalt with a harzburgite residue and used such a pyrolitic composition as starting material to investigate the amphibole stability at subsolidus and near-solidus conditions (Niida & Green, 1999). The MORB pyrolite is an intermediate member between fertile lherzolites (Hawaiian xenoliths) and the Tinaquillo lherzolite.

In order to investigate the effect of potassium in mantle petrology, i.e. the mantle metasomatism due to the interaction with alkali rich fluids, natural bulk compositions have been doped with phlogopite or K-amphibole (Northern Depression Hessian peridotite + 1.5% phlogopite, Mengel & Green, 1989; Mont Briancon, Massif Central, France, spinel lherzolite plus 5 wt% phlogopite or plus 10 wt% K-amphibole, Konzett & Ulmer, 1999; KLB-1 lherzolite plus 10 wt% K-amphibole, Konzett & Fei, 2000).

The spinel to garnet transition

Among phase equilibria occurring at increasing pressure in peridotite systems, the spinel to garnet transition was one of the first tools to recognise UHP metamorphism in many ultramafic xenoliths and orogenic peridotites. Available data on complex ultramafic compositions are represented in Figure 13 and compared with most recent results on the spinel to garnet transition in the simple system CaO–MgO–Al2O3–SiO2 – CMAS (Klemme & O’Neill, 2000; Walter etal., 2002). Niida & Green (1999) located the spinel to garnet transition in a MORB pyrolite at 2.0 GPa and 1050 °C, which is in perfect agreement with the higher temperature, near-solidus data obtained in a dry MORB pyrolite by Robinson & Wood (1998).

In the simple model system CaO–MgO–Al2O3–SiO2 the spinel–garnet transition is governed by the univariant subsolidus equilibrium 

formula
Walter et al. (2002) performed in situ X-ray experiments in order to investigate and solve many discrepancies for the high temperature position (T > 1200 °C) of the spinel–garnet boundary reported in previous results. Although Walter et al. (2002) focussed on a simple model system and on high temperature regions, the innovative approach of in situ reconnaissance of subsolidus phase assemblages is worth of mention as reference for the CMAS system.

In order to investigate the effect of composition and kinetic metastability of phases, experiments were performed starting from mixtures of clinopyroxene, garnet and forsterite and orthopyroxene, clinopyroxene, garnet and forsterite, and the phase boundary was approached from both the high and low-pressure sides. Walter et al. (2002) drew two alternative curves: curve 1 and the shaded area in Figure 13 represent the phase boundary and its error while curve 2 takes into account a suspicious occurrence of spinel + orthopyroxene at about 1000 °C not solved yet in terms of possible metastability. The slope of curve 1, which is in agreement with reversals of Klemme & O’Neill (2000), should therefore represent a maximum.

Fig. 13.

The spinel to garnet transition in the simple system CaO–MgO–Al2O3–SiO2 – CMAS (Klemme & O’Neill, 2000; Walter et al., 2002) and in model peridotites (MORB-pyrolite and Tinaquillo lherzolites; Niida & Green, 1999; Robinson & Wood, 1998). Curves 1 and 2 represent two alternative phase boundaries proposed by Walter et al., 2002 (see text for further details).

Fig. 13.

The spinel to garnet transition in the simple system CaO–MgO–Al2O3–SiO2 – CMAS (Klemme & O’Neill, 2000; Walter et al., 2002) and in model peridotites (MORB-pyrolite and Tinaquillo lherzolites; Niida & Green, 1999; Robinson & Wood, 1998). Curves 1 and 2 represent two alternative phase boundaries proposed by Walter et al., 2002 (see text for further details).

Adding further elements to the simple system, such as Fe2+, Cr and Na, i.e. considering a complex model peridotite system, substantially influences the position of the spinel to garnet transition. Since Fe2+ is preferentially partitioned into garnet, its stability field is expanded towards lower pressure. For a composition with a typical mantle Mg value (XMg = Mg/(Mg + Fe) = 0.9), O’Neill (1981) found that Fe2+ tends to lower the transition of about 0.2–0.3 GPa. Chromium has the opposite effect: it expands the stability of spinel-bearing assemblages because of the preferential partitioning of Cr into spinel. As a result the spinel to garnet transition is not only shifted towards higher pressure as the ratio Cr/Cr + Al in the bulk rock increases, but the pressure and temperature field over which the reaction takes place is greatly broadened, i.e. the transition is a continuous reaction. For a ratio Cr/Cr + Al of about 0.1, Webb & Wood (1986) estimated that, if spinel and garnet are the only Cr-bearing phases, the transition should spread out over a pressure of about 1 GPa. However in natural lherzolite or pyrolite compositions (Na, Cr, and Fe-bearing systems) the occurrence of clinopyroxene may substantially change this prediction. At low mole fractions of Cr (XCr = Cr/(Cr + Al) of about 0.1) the NaCrSi2O6 component in clinopyroxene buffers the Cr content of spinel, sharpening the transition and reducing to only 0.2 GPa the pressure range over which spinel and garnet may coexist (Webb & Wood, 1986). In more refractory compositions, i.e. at higher XCr values, the chromium partitioning between clinopyroxene and spinel is such that spinel composition is not affected by clinopyroxene and the Cr-spinel + garnet assemblage may survive at higher pressure. The Bardane peridotites in Western Norway testify such an ultrahigh pressure phase assemblage (Cr-spinel, garnet and diamond).

Although the width of the continuous reaction has been quantitatively evaluated both in the absence and presence of clinopyroxene, the entity of shifting towards higher pressure is still poorly constrained. However, experiments performed both in natural peridotite compositions (Tinaquillo lherzolite; Robinson & Wood, 1998) and complex model peridotites (MORB pyrolite; Robinson & Wood, 1998; Niida & Green, 1999), i.e. in Fe, Cr, Na-bearing systems, suggest that the effect of chromium predominates over the effect of Fe2+, locating the spinel to garnet transition at pressures substantially higher than those observed in the simple system (CMAS).

The peridotite + H2O system

The hydration of peridotites along mid-ocean ridges has widely been accepted. Many efforts have therefore been devoted to unravelling phase relations in hydrated ultramafic systems, where hydrous minerals such as antigorite, chlorite, amphibole, and a variety of Dense Hydrous Magnesium Silicates may persist under subsolidus conditions at high and ultrahigh pressure (Kawamoto et al., 1996; Ulmer & Trommsdorff, 1999; Fumagalli, 2000). Phase relations in hydrous peridotites are of primary importance in depicting the dehydration evolution of a subducting slab, the metamorphic and igneous history of the mantle wedge at convergent margins and the petrology of deep mantle rocks.

It is worth noting that, although many experimental studies have been devoted to investigate near-solidus and supersolidus phase relations in natural systems approaching ultramafic compositions, subsolidus equilibria are still fragmentary. Figure 14 shows a summary of H2O-saturated phase relations in ultramafic rocks. Antigorite dominates the low temperature field up to a pressure of 6.0 GPa (Ulmer & Tromsdorff, 1999). Aluminium leads to a wide occurrence of chlorite from greenschists to eclogite facies. Amphibole is stable in a wide range of temperature conditions although it shows a temperature dependent mineral chemistry: a calcic amphibole occurs at low temperatures, a relatively Al-poor, tremolitic amphibole coexists with chlorite and, as temperature further increases, above the thermal stability of chlorite (750–800 °C), a Na-Al-rich pargasitic amphibole is stable up to the solidus (Niida & Green, 1999). Amphibole pressure stability is dependent on bulk compositions, mainly on the Na and Ca bulk content, i.e. on fertility. Mysen & Boettcher (1975) found that at 1000 °C amphibole breaks down at 2.2 GPa in harzburgites, at 2.5–2.8 GPa in lherzolites and at 2.8–3.0 GPa in pyrolites. Experiments performed at H2O-undersaturated conditions on MORB pyrolite by Niida & Green (1999) corroborated the hypothesis of the effect of bulk rock alkali content on the amphibole stability: in MORB pyrolites the maximum stability of pargasite lies at 2.8 GPa and 1000 °C, intermediate conditions between those of the Tinaquillo lherzolite (2.6 GPa, 1000 °C; Wallace & Green, 1991) and those of the Hawaiian pyrolite and the Northern Hessian Depression (2.9–3.0 GPa, 1000 °C and 2.8–2.9 GPa, 1000 °C, respectively). Niida & Green (1999) also showed that the maximum temperature stability expands from 1025 to 1150 °C with increasing bulk alkali contents. The higher temperature is related to the higher potassium content (as in the Northern Hessian Depression peridotite modified with 1.5% phlogopite, K2O = 0.43 wt%) which would stabilise amphibole to higher temperatures.

Fig. 14.

Experimentally determined phase diagram for hydrous peridotites.

Fig. 14.

Experimentally determined phase diagram for hydrous peridotites.

At subsolidus conditions pargasitic amphibole and clinopyroxene are the only Na-bearing phases. Niida & Green (1999) investigated systematically the composition of amphibole and coexisting phases in a constant bulk composition, the MORB pyrolite, from 0.4 to 3.0 GPa and between 925 and 1100 °C and found that amphibole breaks down through complex continuous reactions. Based on mass balance calculations and mineral chemistry they suggested that the ultimate breakdown of amphibole is related to the destabilisation of the richteritic component. The ultimate breakdown of amphibole at high temperature occurs as a dehydration reaction, providing therefore a release of fluid. This is not the case at temperatures lower than 800 °C. Fumagalli (2000) performed experiments at H2O saturated conditions on model lherzolite compositions (Na2O–CaO–FeO–MgO–Al2O3–SiO2–H2O, NCFMASH system) and found that, while at temperatures higher than 800 °C a dehydration reaction governs the amphibole breakdown, at temperatures lower than 800 °C, its pressure stability is controlled by the degenerate H2O conserving reaction: 

formula
Chlorite thermal stability, related to the reaction chlorite + enstatite = garnet + olivine + water, results depressed towards lower temperatures in the peridotite model system as compared to the MASH system due to the preferential partitioning of iron into garnet (Fumagalli, 2000).

The pressure stability of chlorite in complex ultramafic systems is still poorly explored and defined. Available experiments performed in the model peridotite system NCFMASH have shown that the pressure stability of chlorite is controlled by the appearance, at temperatures lower than 700 °C, of a 10 Å phase structure. The 10 A phase, first synthesised in the MgO–SiO2–H2O system by Sclar et al. (1965), is a phyllosilicate chemically analogous to talc but with excess water: the chemical formula may be written as Mg3Si4O10(OH)2-nH2O. However, in the complex ultramafic system a 10 A phase structure, identified on the basis of X-ray diffractometry, shows, at the electron microprobe scale, a homogeneous aluminous composition (Fumagalli & Poli, 1999; Fumagalli et al., 2001). The surprising coincidence in composition between this phase and the mixed-layered mineral kulkeite (clinochlore : talc = 1 : 1; Schreyer et al., 1982), may suggest a structural rearrangement of chlorite, intercalating the aluminium-free 10 Å phase through a continuous reaction. The thermal stability of the 10 Å phase structure is governed by the reaction at T > 700 °C: 10 Å phase + clinopyroxene = garnet + orthopyroxene + H2O.

The K-peridotite system

The relevance of metasomatic processes in subduction zones, affecting peridotites of the mantle wedge, has been recognised since a long time. Even though such effects have mainly been addressed on mantle xenoliths incorporated in magmatic suites, peridotites “intruded” in subducted continental crust provide a unique petrological window on processes undergoing at the slab–mantle interface. Buoyancy forces are responsible for peridotite emplacement in the subducted continental crust; the devolatilisation of felsic rocks and mass transfer toward peridotite bodies are strongly enhanced. A variety of volatile-bearing phases develop in ultramafic bodies and recent findings of Ti-phlogopite and diamond as inclusions in garnet peridotites from Bardane (Norway) testify to the intensity of mass exchange at depth (van Roermund et al., 2002).

Although data exist on the stability of pure phases such as phlogopite and K-amphibole (Trønnes, 2002, and references therein), available experimental data on phase relations in potassium doped model peridotites are still limited to just a few studies (Konzett & Ulmer, 1999; Konzett & Fei, 2000). Phase relations (summarised in Fig. 15) are governed by the occurrence of three potassic phases, which are, in the order of pressure stability: phlogopite, a potassic amphibole of richteritic composition, and a K-rich hydrous silicate, termed phase X (K2–xMg2Si2O7HI, x = 0–1, Luth, 1997). K-amphibole represents the breakdown product of phlogopite-bearing assemblages and phase Xrepresents the breakdown product of K-amphibole.

Fig. 15.

Summary of available experimental data in K-doped peridotites. KU99: Konzett & Ulmer, 1999; ST90: Sudo & Tatsumi, 1990; KF00: Konzett & Fei, 2000; L97: Luth, 1997; I98: Inoue et al., 1998. Phl: phlogopite; K-amp: K-amphibole.

Fig. 15.

Summary of available experimental data in K-doped peridotites. KU99: Konzett & Ulmer, 1999; ST90: Sudo & Tatsumi, 1990; KF00: Konzett & Fei, 2000; L97: Luth, 1997; I98: Inoue et al., 1998. Phl: phlogopite; K-amp: K-amphibole.

Sudo & Tatsumi (1990), investigating a phlogopite + diopside ± enstatite composition in the KCMASH system, suggested that phlogopite stability is governed by the reaction: 

formula
In the absence of orthopyroxene the reaction is divariant in the KCMASH system and therefore a field exists where phlogopite and K-amphibole coexist: the lower boundary represents the K-amphibole-in reaction (K-amp-in ST90) while the upper limit is the phlogopite out reaction (Phl-out ST90). In the presence of orthopyroxene, however, the reaction becomes univariant in the KCMASH system and phlogopite is not stable at pressures above 6 GPa (Sudo & Tatsumi, 1990). The potassic amphibole, which forms at the expense of phlogopite, is an Al-poor potassic amphibole termed KK-richterite, KKCaMg5Si8O22(OH)2. In the Na-bearing system sodium will contribute to the formation of K-richterite, KNaCaMg5Si8O22(OH)2, as a breakdown product of phlogopite. Konzett & Ulmer (1999) investigated the KNCMASH system using bulk compositions with an excess of phlogopite over orthopyroxene, in order to determine the maximum pressure stability of phlogopite in K-richterite bearing assemblages. Their results show that K-amphibole appears between 6 and 6.5 GPa at 800 °C and between 6.5–7.0 GPa at 1100 °C (Kr-in KU99).

Konzett & Ulmer (1999) underlined a discrepancy in the slope of the K-amphibole-in reaction as compared with what Sudo & Tatsumi (1990) determined. Whether the slope is positive or negative is still under debate.

Below the K-amphibole-in reaction the stable assemblage is constituted of phlogopite, clinopyroxene, orthopyroxene, garnet and fluid, and close to the reaction an olivine-bearing assemblage occurs. Above the K-amphibole-in reaction, phlogopite and K-amphibole coexist due to the excess of phlogopite respect to orthopyroxene up to the ultimate breakdown of phlogopite at pressures between 8 and 9 GPa, 1100 °C (Phl-out KU99). At pressures above 13–14 GPa (1100 °C) phase X replaces K-amphibole (K-amphibole-out reaction). As the complexity of the system investigated increases, the breakdown of K-amphibole occurs at lower pressure. The K-amphibole to phase X transition involves continuous changes in garnet compositions through Ca–Mg exchange and the limited solubility of MgSiO3 to form majoritic component (Konzett & Fei, 2000).

The thermal stability of K-amphibole is determined by the appearance of the anhydrous assemblage garnet, olivine, orthopyroxene and clinopyroxene (K-amp-out KU99). In the KNCMASH system it occurs between 1300 and 1400 °C at 8.0 GPa and shows a positive slope.

In an Fe-bearing system and lherzolitic composition, the effect of iron has to be taken into account. Konzett & Ulmer (1999) investigated a K-doped lherzolite modifying the composition of the Mont Briancon lherzolite (Massif Central, France) by adding phlogopite or K-richterite (see Table 1).

In the lherzolitic system the K-amphibole-in reaction slightly shifts towards lower pressure (between 6.0 and 6.5 GPa at 1100 °C) due to the preferential Fe2+ partitioning into garnet, which is a product of phlogopite breakdown. The coexistence of phlogopite and K-amphibole is, however, reduced to less than 1 GPa. It is worth noting that care should be taken in dealing with Fe-bearing systems due to the not easily predictable influence of the Fe3+/Fe2+ ratio on phase relations.

The K-amphibole thermal stability in the Fe-bearing lherzolitic system seems influenced again by oxidation reactions: the slope suggested by experimental data from Konzett & Ulmer (1999) is indeed opposite compared to what it was predicted in the Fe-free system. An explanation given by the authors refer to the pressure and temperature dependent oxidation of graphite inner capsule which, producing CO2, would lower the H2O activity and destabilise K-amphibole.

The pressure stability of K-amphibole in Fe-bearing lherzolitic systems was investigated by Konzett & Fei (2000): they run experiments by using a K-amphibole doped peridotite KLB1 from 12 to 14 GPa and 1200 °C. Potassic phases, either K-amphibole or phase X, coexist with garnet, low-Ca clinopyroxene, high-Ca clinopyroxene and Mg2SiO4. In the Fe-bearing system the K-amphibole to phase X transition, occurring between 12 and 13 GPa at 1200 °C, is shifted by about 1.0 GPa towards lower pressure as compared with what was found in the Fe-free system.

Metasedimentary rocks

Despite the fact that most UHP terrains have been unequivocally identified also in metasedimentary compositions, and that metapelites and metagreywackes give a unique geochemical signature to orogenic magmas, high-pressure experiments are still in a state of a preliminary investigation. Most published systematic work within the coesite stability field has been performed in the model systems KMASH, KFASH (Massonne & Szpurka, 1997) and KCMASH (Hermann, 2002b). Domanik & Holloway (1996) and Ono (1998) conversely performed experiments above 6 GPa, but on natural sedimentary compositions approaching metapelites. Ferri et al. (2000) and Poli & Schmidt (2002) presented data on experiments in KCFMASH to 2.7 GPa for bulk compositions representative of metapelites. Therefore the general picture available for the interpretation of natural cases is still fragmentary. The assemblage garnet-phengite-kyanite-coesite-clinopyroxene is predominant in most pelitic compositions, as biotite is completely consumed by the reaction biotite + kyanite + SiO2 = garnet + phengite (Hermann, 2002b; Poli & Schmidt, 2002), unless unusually high XMg bulk compositions are considered. Appearance of the join talc–garnet at ca. 2.4–2.5 GPa in the model system KCFMASH should be responsible for the widespread occurrence of talc in most natural bulk compositions (Ferri et al., 2000; Poli & Schmidt, 2002), but the fate of talc with pressure and its relationships with chloritoid are still unknown.

Most experimental studies have been performed at fluid saturation, and at such conditions K-feldspar was never found to be stable either in metapelites or in metagreywackes. It is therefore only at fluid-undersaturated conditions that K-feldspar is expected to be a possible UHP phase in metasedimentary material. However, as illustrated above, attainment of fluid-undersaturation during a prograde, subduction related P–T path for metasedimentary (initially H2O-rich) material is unlikely, and therefore the relevance of K-feldspar is expected to be very limited.

In conclusion a number of questions remains unresolved:

  • What is the pressure–temperature stability of the join garnet–talc?

  • What is the pressure stability of biotite solid solution?

  • What is the fate of magnesium-rich chloritoid?

  • Which reactions control the behaviour of clinopyroxene in common metapelites and metagreywackes?

  • Are indeed some unusual high-pressure phases found in experimental studies (Mg-sursassite, topaz-OH, phase Pi) relevant for UHP rocks and which reactions govern their appearance?

Okamoto & Maruyama (1998) and Schmidt & Poli (1998) revealed that potassium enters in clinopyroxene and sodium in garnet with increasing pressure; Hermann (2002b) demonstrated significant deviation of phlogopite toward talc and of clinopyroxene toward Ca-Eskola component. Such features are responsible for the peculiar intragranular textures found in UHP rocks, such as K-feldspar lamellae in clinopyroxene or silica rods in omphacite (Liou et al., 1998).

Given all of the complexities present in these bulk compositions and their importance for geodynamic processes, a particular effort has to be done in the near future to provide a reliable tool for the interpretation of metasedimentary rocks.

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Copyright permission

Figure 7 (from Geochimica et Cosmochimica Acta, Vol. 33, Boettcher, A.L. & Wyllie, P.J.: The system CaO-SiO2-CO2-H2O – III. Second critical end-point on the melting curve, pp. 611-632, copyright 1969) and Figure 9 (from Chemical Geology, Vol. 192, Hermann, J.: Allanite: thorium and light rare earth element carrier in subducted crust, pp. 289-306, copyright 2002) in this chapter are reprinted with permission from Elsevier.

Figures & Tables

Fig. 1.

Backscattered electron images of run products showing textural differences between experiments performed on similar bulk compositions and run durations, but at extremely different pressure and temperature conditions. (a) graphite-bearing, fluid-saturated model MORB at 2.7 GPa, 730 °C, 224 hours (courtesy Ada Crottini); (b) graphite-bearing, fluid-saturated natural MORB at 19 GPa, 1200 °C, 120 hours (courtesy Kazuaki Okamoto). gar: garnet; cpx: clinopyroxene; ky: kyanite; maj: majorite; ca-pvs: Ca-perovskite (bright spots); sti: stishovite (dark spots). Note that magnification in (a) is 1000× and in (b) is 1800×.

Fig. 1.

Backscattered electron images of run products showing textural differences between experiments performed on similar bulk compositions and run durations, but at extremely different pressure and temperature conditions. (a) graphite-bearing, fluid-saturated model MORB at 2.7 GPa, 730 °C, 224 hours (courtesy Ada Crottini); (b) graphite-bearing, fluid-saturated natural MORB at 19 GPa, 1200 °C, 120 hours (courtesy Kazuaki Okamoto). gar: garnet; cpx: clinopyroxene; ky: kyanite; maj: majorite; ca-pvs: Ca-perovskite (bright spots); sti: stishovite (dark spots). Note that magnification in (a) is 1000× and in (b) is 1800×.

Fig. 2.

Sequence of reactions involving hydrates and carbonates in the model systems MgO-SiO2-H2O (a-d) and MgO-CaO-CO2 projected from silica (e-f). Assemblages below the tie-lines brucite-antigorite-talc-quartz in (a) and (b) represent H2O-undersaturated fluid absent conditions; assemblages below the tie-lines magnesite-dolomite-aragonite are CO2-undersaturated. Points 1 and 2 stand for two hypothetical ultramafic bulk compositions. arag: aragonite; atg: antigorite; br: brucite; cc: calcite; di: diopside; dol: dolomite; en: enstatite; fo: forsterite; mag: magnesite; p: periclase; q: quartz; ta: talc; wo: wollastonite. See text for explanation.

Fig. 2.

Sequence of reactions involving hydrates and carbonates in the model systems MgO-SiO2-H2O (a-d) and MgO-CaO-CO2 projected from silica (e-f). Assemblages below the tie-lines brucite-antigorite-talc-quartz in (a) and (b) represent H2O-undersaturated fluid absent conditions; assemblages below the tie-lines magnesite-dolomite-aragonite are CO2-undersaturated. Points 1 and 2 stand for two hypothetical ultramafic bulk compositions. arag: aragonite; atg: antigorite; br: brucite; cc: calcite; di: diopside; dol: dolomite; en: enstatite; fo: forsterite; mag: magnesite; p: periclase; q: quartz; ta: talc; wo: wollastonite. See text for explanation.

Fig. 3.

Saturation surfaces for water and carbon dioxide compared to chemical potentials of H2O and CO2 buffered by antigorite-enstatite-forsterite and by magnesite-quartz-enstatite, respectively. Volatile saturation and therefore fluid release occurs if the P–T path drives μH2O or μCO2 toward the thermodynamic surfaces of H2O or CO2 fluid, respectively. Note that μi, = molar G (from Poli & Schmidt, 2002).

Fig. 3.

Saturation surfaces for water and carbon dioxide compared to chemical potentials of H2O and CO2 buffered by antigorite-enstatite-forsterite and by magnesite-quartz-enstatite, respectively. Volatile saturation and therefore fluid release occurs if the P–T path drives μH2O or μCO2 toward the thermodynamic surfaces of H2O or CO2 fluid, respectively. Note that μi, = molar G (from Poli & Schmidt, 2002).

Fig. 4.

Composition diagram for the system CaO-Al2O3-SiO2-H2O-CO2 projected from kyanite and coesite (from Poli & Schmidt, 1998, modified) which shows occurrence of hydrates and carbonates at fluid-absent conditions, equality of chemical potential of H2O for selected fluid present and fluid absent conditions and possible appearance of an aqueous fluid by addition of CO2 to the system.

Fig. 4.

Composition diagram for the system CaO-Al2O3-SiO2-H2O-CO2 projected from kyanite and coesite (from Poli & Schmidt, 1998, modified) which shows occurrence of hydrates and carbonates at fluid-absent conditions, equality of chemical potential of H2O for selected fluid present and fluid absent conditions and possible appearance of an aqueous fluid by addition of CO2 to the system.

Fig. 5.

The vaporous surface in the system MgO-SiO2-H2O at 3 GPa after Ryabchykov et al. (1983).

Fig. 5.

The vaporous surface in the system MgO-SiO2-H2O at 3 GPa after Ryabchykov et al. (1983).

Fig. 6.

Schematic presentation of a binary system solid + H2O (from Stalder et al., 2000, copyright Mineralogical Society of America). P2 > P1.

Fig. 6.

Schematic presentation of a binary system solid + H2O (from Stalder et al., 2000, copyright Mineralogical Society of America). P2 > P1.

Fig. 7.

Isobaric, polythermal diagrams for the hypothetical ternary system MO-SiO2-H2O (from Boettcher & Wyllie, 1969). Dashed lines are isothermal field boundaries. Dotted lines connect temperature maxima on the liquidus and vaporous field boundaries with the crystalline phase with which the liquid and vapour are in equilibrium. Reprinted with permission from Elsevier.

Fig. 7.

Isobaric, polythermal diagrams for the hypothetical ternary system MO-SiO2-H2O (from Boettcher & Wyllie, 1969). Dashed lines are isothermal field boundaries. Dotted lines connect temperature maxima on the liquidus and vaporous field boundaries with the crystalline phase with which the liquid and vapour are in equilibrium. Reprinted with permission from Elsevier.

Fig. 8.

Comparison between different experimental studies on the transformation quartz-coesite and the calculated location according to the database of Holland & Powell (1998 and recent updates). Determination by Bohlen & Boettcher (1982) and Mirwald & Massonne (1980) (reported by Hemingway et al., 1998) and Bose & Ganguly (1995) were all performed in a piston cylinder apparatus. The data by Walter et al. (2002) refer to an in situ estimate performed under a synchrotron source.

Fig. 8.

Comparison between different experimental studies on the transformation quartz-coesite and the calculated location according to the database of Holland & Powell (1998 and recent updates). Determination by Bohlen & Boettcher (1982) and Mirwald & Massonne (1980) (reported by Hemingway et al., 1998) and Bose & Ganguly (1995) were all performed in a piston cylinder apparatus. The data by Walter et al. (2002) refer to an in situ estimate performed under a synchrotron source.

Fig. 9.

Trace element distribution among the phases of eclogites (from Hermann, 2002a). Modal amounts are 36% omphacite, 35% garnet, 14% quartz, 12% phengite, 2.5% rutile, 0.52% apatite, 0.1% allanite, 0.032% zircon. Reprinted with permission from Elsevier.

Fig. 9.

Trace element distribution among the phases of eclogites (from Hermann, 2002a). Modal amounts are 36% omphacite, 35% garnet, 14% quartz, 12% phengite, 2.5% rutile, 0.52% apatite, 0.1% allanite, 0.032% zircon. Reprinted with permission from Elsevier.

Fig. 10.(a)

Review of the experimentally determined phase relationships in MORB eclogites to 5 GPa in the presence of an aqueous fluid. (b) Schematic displacement of invariant point “CaAl” in (a) as a function of the stable garnet composition, which, in turn, is a function of the bulk composition adopted.

Fig. 10.(a)

Review of the experimentally determined phase relationships in MORB eclogites to 5 GPa in the presence of an aqueous fluid. (b) Schematic displacement of invariant point “CaAl” in (a) as a function of the stable garnet composition, which, in turn, is a function of the bulk composition adopted.

Fig. 11.

The stability of graphite (gph) with carbonates (stippled field) at high and low temperature conditions. FMQ: fayalite-magnetite-quartz; FFsM: fayalite-ferrosilite-magnetite; NNO: nickel-bunsenite; GCO: graphite-CO2; EMOG: enstatite-magnesite-olivine-graphite; DSDG: dolomite-quartz/coesite-diopside-graphite. See text for explanation.

Fig. 11.

The stability of graphite (gph) with carbonates (stippled field) at high and low temperature conditions. FMQ: fayalite-magnetite-quartz; FFsM: fayalite-ferrosilite-magnetite; NNO: nickel-bunsenite; GCO: graphite-CO2; EMOG: enstatite-magnesite-olivine-graphite; DSDG: dolomite-quartz/coesite-diopside-graphite. See text for explanation.

Fig. 12.

Review of the experimentally determined phase relationships in MORB eclogites to 6 GPa in the presence of a C-O-H mixed fluid at variable oxygen fugacities and bulk C-O-H contents.

Fig. 12.

Review of the experimentally determined phase relationships in MORB eclogites to 6 GPa in the presence of a C-O-H mixed fluid at variable oxygen fugacities and bulk C-O-H contents.

Fig. 13.

The spinel to garnet transition in the simple system CaO–MgO–Al2O3–SiO2 – CMAS (Klemme & O’Neill, 2000; Walter et al., 2002) and in model peridotites (MORB-pyrolite and Tinaquillo lherzolites; Niida & Green, 1999; Robinson & Wood, 1998). Curves 1 and 2 represent two alternative phase boundaries proposed by Walter et al., 2002 (see text for further details).

Fig. 13.

The spinel to garnet transition in the simple system CaO–MgO–Al2O3–SiO2 – CMAS (Klemme & O’Neill, 2000; Walter et al., 2002) and in model peridotites (MORB-pyrolite and Tinaquillo lherzolites; Niida & Green, 1999; Robinson & Wood, 1998). Curves 1 and 2 represent two alternative phase boundaries proposed by Walter et al., 2002 (see text for further details).

Fig. 14.

Experimentally determined phase diagram for hydrous peridotites.

Fig. 14.

Experimentally determined phase diagram for hydrous peridotites.

Fig. 15.

Summary of available experimental data in K-doped peridotites. KU99: Konzett & Ulmer, 1999; ST90: Sudo & Tatsumi, 1990; KF00: Konzett & Fei, 2000; L97: Luth, 1997; I98: Inoue et al., 1998. Phl: phlogopite; K-amp: K-amphibole.

Fig. 15.

Summary of available experimental data in K-doped peridotites. KU99: Konzett & Ulmer, 1999; ST90: Sudo & Tatsumi, 1990; KF00: Konzett & Fei, 2000; L97: Luth, 1997; I98: Inoue et al., 1998. Phl: phlogopite; K-amp: K-amphibole.

Table 1.

Model systems where ultrahigh pressure phases are found, with a brief list of references

SystemUHP phasesSource for literature
FMSolivine, pyroxene, wadsleyite, ringwoodite, ilmenite, majorite, perovskite, magnesiowustite, stishoviteAkaogi et al. (1998); Frost et al. (2001)
MASenstatite, pyrope, majorite, ilmenite, perovskite, “tetragonal phase”Akaogi et al. (2002); Heinemann et al. (1997)
FMASmajoriteO’Neill & Jeanloz (1994)
CMSdiopside, Ca-perovskite, walstromite, CaSi2O5Canil (1994); Gasparik et al. (1994);
Gasparik (1996a)
CFSskiagiteWoodland & O’Neill (1995)
CMAspinel, Ca-ferrite, “hexagonal Al-rich phase”Akaogi et al. (1999); Miyajima et al. (2001)
NCMASjadeite, Na Ca-ferriteGasparik (1996b)
NKASwadeite, hollandite, calcium ferrite(-structure)Yagi et al. (1994)
ASHtopaz-OH, phase Pi, phase eggDaniels & Wunder (1996); Schmidt et al. (1998)
MSHantigorite, clinohumite, phase A, phase B, superhydrous phase B (phase C), phase D (phases F and G), phase E, 10A phaseAngel et al. (2001); Ulmer & Trommsdorff (1999); Wunder (1998); Stalder & Ulmer (2001)
MASHchlorite, Mg-sursassite (MgMgAl pumpellyite), Mg-stauroliteBromiley & Pawley (2002); Fockenberg (1998)
CASHlawsonite, zoisitePoli & Schmidt (1998); Schmidt (1995)
CFASHclinozoisite, garnetBrunsmann et al. (2002)
KMASHphlogopite, wadeite, phase X, phengiteTronnes (2002); Massonne & Szpurka
KFASH(1997); Hermann (2002b); Schmidt et al. (2001)
CM-CO2dolomite, aragonite, magnesiteLuth (1995, 2001)
SystemUHP phasesSource for literature
FMSolivine, pyroxene, wadsleyite, ringwoodite, ilmenite, majorite, perovskite, magnesiowustite, stishoviteAkaogi et al. (1998); Frost et al. (2001)
MASenstatite, pyrope, majorite, ilmenite, perovskite, “tetragonal phase”Akaogi et al. (2002); Heinemann et al. (1997)
FMASmajoriteO’Neill & Jeanloz (1994)
CMSdiopside, Ca-perovskite, walstromite, CaSi2O5Canil (1994); Gasparik et al. (1994);
Gasparik (1996a)
CFSskiagiteWoodland & O’Neill (1995)
CMAspinel, Ca-ferrite, “hexagonal Al-rich phase”Akaogi et al. (1999); Miyajima et al. (2001)
NCMASjadeite, Na Ca-ferriteGasparik (1996b)
NKASwadeite, hollandite, calcium ferrite(-structure)Yagi et al. (1994)
ASHtopaz-OH, phase Pi, phase eggDaniels & Wunder (1996); Schmidt et al. (1998)
MSHantigorite, clinohumite, phase A, phase B, superhydrous phase B (phase C), phase D (phases F and G), phase E, 10A phaseAngel et al. (2001); Ulmer & Trommsdorff (1999); Wunder (1998); Stalder & Ulmer (2001)
MASHchlorite, Mg-sursassite (MgMgAl pumpellyite), Mg-stauroliteBromiley & Pawley (2002); Fockenberg (1998)
CASHlawsonite, zoisitePoli & Schmidt (1998); Schmidt (1995)
CFASHclinozoisite, garnetBrunsmann et al. (2002)
KMASHphlogopite, wadeite, phase X, phengiteTronnes (2002); Massonne & Szpurka
KFASH(1997); Hermann (2002b); Schmidt et al. (2001)
CM-CO2dolomite, aragonite, magnesiteLuth (1995, 2001)

Abbreviations (System column): A: Al2O3, C: CaO, F: FeO, H: H2O, K: K2O, M: MgO, N: Na2O, S: SiO2

Contents

GeoRef

References

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