Chapter 11 Silicate melts and rocks

793

Chapter 11 Silicate Melts and Rocks

11.1. Introduction Determination of the chemical structure of silicate melts and aqueous fluids are important because these pertain to processes involving heat and mass transfer in the terrestrial planets. During cooling of the silicate planetary bodies, materials constantly recycled throughout the crust and mantle result in chemically diverse rock compositions. Since melts served as the main agents in recycling, it is necessary to understand the chemical structure of both silicate melts involving alkali and alkaline earths, aluminium and also hydrothermal fluids. The fields of major rock types in the system SiO2-A1203M,,O,,-2 (where M = K, Na, Ca, Mg and Fe 2+) are shown in Fig. 11.1. Volatiles such as H20 and CO2 significantly influence melting behaviour. Melts generated below a critical depth would remain neutrally buoyant and thus not rise to the surface. With pressure, the melting curve rises continuously and the melting line may (in general) terminate in a critical point. The melting curve may rise to a limiting or asymptotic value. Possibly, at sufficiently high pressure, the distinction between a solid and a liquid would disappear and the material would persist in an amorphous state. Interactions between silicate melt and water influence the fundamental physical and chemical processes within the Earth. At high pressure, water is released from silicate melts, altering the surrounding volume of fluid, which may govern the explosive nature of many volcanic processes. Evidently, a relationship exists between the mixing behaviour of solutions in silicate melts (with or without silicate-beating aqueous solutions) and the transport properties or rheology of fluids and melts. These fluids may also act as solvents for silicates and serve as the principal transporting agent within the Earth. This transport may cause enrichment of elements to cause deposition of important ore bodies. Magmatic liquids are regarded as the main agents for heat and transfer within the Earth and other terrestrial planetary bodies. In magmatic melts positive correlation exists between melt viscosity or polymerization and silica content. Positive correlation also exists between electrical conductivity and metal/silicon ratio in the melt. Conductivity is related to diffusivity. Nernst-Einstein equation relates conductivity of a component, or/, to diffusivity, Di,

~

F2ziDi kT

794

Chapter 11

$i0 2 rhyolitic1 ~ andesitic~ basaltic~

itllllll

90

o)%.

,o

ooT/,K Figure 11.1. Approximate fields of major rock types expressed in terms of the pseudo-ternary system MnOn/2A1203-SIO2, where M = K, Na, Ca, Mg, and Fe2+ (after Mysen, 1990).

where F is Faraday's constant, Zi the electric charge of a particle i, k the Boltzmann's constant and T the absolute temperature. Diffusivity (Di) is related to viscosity (7/) using the Stokes-Einstein equation r

---

KT 67rriDi

where ri is the radius of the moving particle. For silicate melts the equation stated below is found more reliable than the StokesEinstein equation rl =

KT aiDi

is the jump distance. Bond strength activation energy. For melts in system Na20-SiO2 activation energy decreases rapidly from a value of pure SiO2 that is similar to the S i - O bond energy (---600 kJ/mol) to a nearly constant value for Na20 contents above ---15 mol% (Mysen, 2003). The structural control of viscous flow may be breakage of bridging S i - O - S i bonds. However, viscous flow does not depend solely on the energetic of bond breakage. Arrhenius expression for viscosity and activation energy is where

o~i

lnrt = lnr/o + (En/Rt) where ~7 is the viscosity, E,7 the activation energy of viscous flow, T the temperature in Kelvin and R the gas constant. Viscosity of most silicate systems including melt and glass do not follow Arrhenius expression. Rather viscosity displays a distinct curvature in In ~/vs 1/T space. For describing viscosity of melt and glass the most successful model involves the theory of configurational entropy, first described by Adam and Gibbs (1965).

Silicate Melts and Rocks

795

For describing viscosity Richet (1984) used the relation Inn

-- A e + Be/ZSconf

where Sconf is the configurational entropy, Ae a constant and Be the molar free hindrance energy, which is the energy required for changing from one configuration to another in melt. This energy is positively correlated with the silica content.

11.1.1. Magmatic melt under pressure Under pressure, the polymerized silicate melts behave anomalously: their viscosities decrease with increasing pressure. Viscosities can be estimated from the diffusivities of network-forming ions using the Eyring relation. Poe et al. (1995) have determined oxygen self-diffusion coefficients up to 15 GPa for three melts with varying degrees of polymerization. They found that, for sodium tetrasilicate (Na2Si409), which nominally has an average of 0.5 non-bridging oxygens per tetrahedral cation (NBO/T), oxygen diffusivities continue to increase with pressure up to 15 GPa. Results from molecular dynamics simulations of Na2Si409 liquid have indicated that diffusivities pass through a maximum near 20 GPa but, for the slightly more polymerized Na3A1SiTO17 (NBO/T = 0 . 2 5 ) , the experimentally determined diffusivities pass through a maximum near 5 GPa. These results would thus suggest that viscosities pass through a minimum with increasing pressure, and that the minimum is polymerization dependent. Solid-state NMR studies have shown that glasses quenched from high pressures contain increasing amounts of high-coordinate A1 (in alumino-silicates) and Si (in binary silicates) with increasing pressure. This is interpreted as a densification mechanism which initially enhances oxygen diffusivities. NBOs are likely to participate in the formation of high-coordinate A1 and Si species. However, at higher pressures, after all NBOs have been consumed by this reaction, bridging oxygens are required to continue the increase in the coordination of A1 and Si (thus forming three-coordinate oxygen). This marks the diffusivity maximum (and viscosity minimum) and accounts for its polymerization dependence. Such a diffusivity maximum is observed for the fully polymerized albite melt. This possibly indicates that oxygen self-diffusion is sensitive to trace amounts of water. However, there is little variation in water content as a function of pressure. The estimated concentrations (< 0.1 wt% H20) would reflect only a minor extent of depolymerization by incorporation of H20. Finally, increasing oxygen diffusivity with pressure up to 5 GPa is consistent with decreasing viscosity up to 3 GPa. Experimental studies on rock melts provided varying melt products. The salient observation on peridotite study may be outlined as below. In the low-pressure range (< 5 GPa), the chemistry of the melt is dependent on the exerted pressure and the relative proportions of the volatiles (H20 and CO2) present. With increasing pressure, the products of peridotite melt range from basalt through picrite and

Chapter 11

796

komatiite to lherzolite, accompanying a change in the chemical content of MgO and Si02 as: Oxidewt%

Picrite

~

Komatiite

--,

Lherzolite

MgO

> 10

> 20

> 35

SiO2

45-50

35-45

40

Thermodynamic modelling of mantle melting shows that garnet-spinel and spinel-plagioclase transitions are regions of low freezing but not of enhanced melting. The implications of this for problems in MORB petrogenesis have been discussed by Ashimow (1995). Kimberlites could be derived from relatively low-temperature melting of an H20rich mantle at depths of 150-300 km. Thus, the diamond-producing deposits could be associated with hydrous melting processes. Abundant evidence from xenoliths indicate that some regions of the sub-continental upper mantle may be significantly hydrated.

11.2. A l u m i n o - s i l i c a t e m e l t s

A knowledge of the effect of pressure on alumino-silicate glass and liquid structure is critically important to understand the magma flow up from depth and the melting of the igneous rocks. A pressure-induced coordination change is likely to be a very important mechanism in influencing the properties of natural magmas in the Earth's upper mantle ( < 300 km depth). The anionic structure of most magmatic liquids will consist of units that have average values of NBO/T equal to 2, 1 and 0. Alumino-silicate melt shows a change of A13+ coordination from fourfold at room pressure to sixfold coordination at upper mantle pressures. 11.2.1. C a O - A I 2 0 3 - S i O 2 m e l t s : c o m p r e s s i b i l i t i e s The melts in the CaO-A1203-SiO2 system were studied by Webb and Courtial (1996) for their compressibilities over the range 1,350-1,600~ For these melts, the bulk modulus increases with the addition of CaO and decreases with the addition of A1203 and SIO2. The compressibilities of CaO-SiO2 and CaO-A1203-SiO2 melts are lower than those of Na20-A1203-SiO2 and binary alkali-silicate melts in general. The mechanism of compression of CaO-A1203-SiO2 melts is different from that of Na20-A1203-SiO2 melts. The compressibility of the former is a complex function of the structural geometry variations arising from the presence of CaO, the C a O - O - S i bond interactions and the packing of cations around the anions in a melt, having clusters of preferred A1-Si-A1 bonding. The compressibility of the latter melts appears to be mainly a function of the A1-O interactions.

Silicate Melts and Rocks

797

11.2.2. Na20-Al203-SiO2 melts: AbsoNTS50 Yargar et al. (1995) investigated a model composition between albite (Na A1Si3Os) and sodium tetrasilicate (NTS, Na2Si409). a melt of 5 0 : 5 0 of these produced a glass of composition Na3A1Si7017 (AbsoNTSso). The 27A1 MAS N M R spectra (Fig. 11.2) for the glasses prepared at high pressure showed the resonance lines ascribed to as A1TM, A1v and A1vI species (Sato et al., 1991). The relative abundance of A1vI markedly increases with pressure while the amount of A1v goes through a m a x i m u m (---28%) around 8 GPa (Table 11.1). In Table 11.1, the abundance of A1v and A1vI in quenched glasses was estimated from the asymmetric peak areas of the - 1 / 2 - - - . 1/2 (central transition) 27A1 N M R resonance. The 6iso was estimated near the base of the asymmetric peaks. A convolution minimization fit to all three resonances in the spectra, allowing the amplitude and width to vary, gave a reasonable representation of the original spectra. Relative errors for the quantification of A1 species were ___6%. Figure 11.2 shows that, at high pressure, the line of the 23Na resonance (in AbsoNTSso) becomes narrower. This is caused by a decrease in quadrupole coupling

oPo__JA<___ 120 BO t,O

0-/+0

150 50 -50 -150

Parts per million Figure 11.2. 23Naand 27A1MAS NMR spectra for glasses under pressure showing the resonance liner for A1Iv, A1v, and A1vI spices (Sato et al., 1991).

TABLE 11.1 The abundance of A1w, A1v and A1vI at varying pressures in AbsoNTSso Species

A104 (A1Iv) A105 (A1v) A106 (A1vI)

6iso (ppm)

77 40 8

Integral (%) 6 GPa

8 GPa

10 GPa

12 GPa

80 12 8

49 28 23

42 20 38

35 17 48

798

Chapter 11

parameters or by a reduction in the range of chemical environment, i.e., lesser distortion (more symmetrical average coordinate of Na) at high pressure.

11.3. Viscosity: controlling factors Viscosity and density are related to the number of NBOs and also to the difference between total positive and total negative electrical charge. The property of interest in most high-pressure experimental investigations of melt rheology is the Newtonian viscosity (see Webb and Dingwell, 1995).

11.3.1. Diffusivity: Stokes-Einstein equation The Stokes-Einstein equation was formulated to explain the inverse proportionality between diffusivities of specific components and viscosity of the liquid: r i O = kBT/A

(ll-1)

where ~ is the viscosity, D the diffusivity, kB the Boltzmann's constant and A an effective length scale or jump distance for diffusion, which is also represented as 2 7ra, where a is the diameter of a Brownian particle (in liquid) (see also Sections 15.14.1 and 15.14.3). The determination of diffusivities can be accomplished by the detection of very short concentration profiles within a binary diffusion couple. Silicate melts posses a viscoelastic behaviour, which can be usefully approximated by a Maxwell body with a distribution of relaxation times. The relaxation mode controlling viscous flow in silicate melts has been linked quantitatively in temperature-time-space to the exchange of Si and O atoms in the melt, as determined by motional averaging in spectroscopic experiments. Many homogeneous equilibria involve reactions in the silicate melt phase whose kinetics are closely reflected in the relaxational time scale of resultant properties such as enthalpy, volume and shear stress. Hence, a spectroscopic investigation of relaxation in melts can be used to derive the viscosity data. Time-domain longitudinal dilatometric experiments reveal an onset of nonNewtonian rheology at approx. 2.5 log units of strain below the relaxation strain rate derived from the Maxwell equation: "r = rl/G = 1let

(11-2)

where r/is the shear viscosity, G the shear modulus, T the shear relaxation time and et is the relaxation strain rate. (Note: Experimental strain rate < 1,000 strain rate (relaxation).)

11.3.2. Temperature dependence: Arrhenian approximation Over the restricted temperature ranges defined by the phase equilibria of melting within the Earth, the Arrhenian approximation log'r/= a + b / T

(11-3)

Silicate Melts and Rocks

799

holds in a limited way but the temperature dependence of liquid viscosities over the range of viscosities encountered from liquidus temperatures (as low as 10 Pa s) to the glass transition (--~10la Pa s at typical dilatometric experimental timescales) reveals enormous departures from the Arrhenian approximation (see Fig. 3 of Dingwell, 1998). In the case of the diffusion of silicon, the Arrheninan relationship such as the Tamann-Vogel-Fulcher equation: l o g r / = a + b / ( T - c)

(11-4)

stands valid over a wide range of viscosities (Dingwell, 1990). Relaxational spectroscopic data also reveal oxygen-diffusivity change with viscous flow at high viscosities. 11.3.3. Alkali oxides

The viscosity can be seen to decrease strongly and non-linearly with addition of alkali oxide to silica melt. The non-Arrhenian temperature dependence is typical of very basic melts of geological relevance. In the high-viscosity range, the apparent activation energy increases with addition of alkali oxide to silica. 11.3.4. Water effect

Traces of water, typically in the range of tens of hundreds of ppm, can result in a reduction in melt viscosity of several orders of magnitude. However, the effect of water is more pronounced in a high-viscosity range rather than in a low-viscosity range (e.g., Hess and Dingwell, 1996). 11.3.5. Pressure effects on viscosity

The effect of pressure on the viscosity of silicate melts has been determined by using the Stokesian falling sphere law: V = 2r2Apg/2TI

(11-5)

where V is the settling viscosity, r the sphere radius, Ap the density contrast between sphere and liquid and g the acceleration. The influence of pressure on melt structure and properties has been discussed by Wolf and McMillan (1995). In the CaO-AlaO3-SiOa system, there is a transition to a positive pressure dependence of viscosity as the SiO2 content is reduced below 50 mol%. The degree of polymerization of the melt determined from the NBO/T (Mysen et al., 1982) content shows a trend of less viscosity decrease with pressure as SiO2 decreases. Melt compositions approaching ultra-basic chemistry show little influence of pressure on the viscosity. The general nature of the transition from negative to positive pressure dependence of viscosity in more depolymerized melts was investigated by Brearley et al. (1986). In magmatic processes under pressure, the viscosity of many silicate melts shows a decrease. Again, pressure may enhance the fragility of silicate melt. In alumino-silicate

800

Chapter 11

melts, the type of charge-compensating cation affects the pressure dependence of the viscosity as well as the compressibility of the melt (Kushiro, 1981).

11.3.6. Silicate polymerization A melt of SiO2 shows a high viscosity because of the presence of a fully polymerized tetrahedra. Viscous flow in pure SiO2 requires breaking strong S i - O bonds (452 kJ/mol), which form a high-viscosity network. When a network modifier like Na20 is added to it, the S i - O - S i linkages break to form S i - O - . . . N a +. This constitutes what are called NBOs and causes a lowering of the viscosity. Thus, adding alkali creates NBOs, which bar the formation of high-coordinate network cations. This in turn reduces the viscosity but, if A1203 is added to the melt (say, an alkali silicate) to such an extent that Na20:A1203--> 1, the NBOs are removed, the activation energy is increased and the tetrahedral network is reconstructed with increased viscosity. Under pressure, the formation of high-coordinated species is favoured with consequent lowering of the activation energy for oxygen exchange. High-pressure lowering of the viscosity can be related to the increase in the number of isoergic configuration states in the melt, which results in an increase in the number of channels for stress relaxation (Dickinson et al., 1990). By incorporating these conditions in the Adam-Gibbs expression for cooperative relaxation in polymer melts, one obtains the expression for shear viscosity: rl = r l o e x p ( - C / T S c )

(11-6)

where C relates the activation energy barrier for relaxation and Sc is the configurational entropy of the melt. NMR studies have revealed the presence of substantial amounts of highcoordinated silicon species (Si v and Si vI) in partially depolymerized alkaline silicate glasses quenched from high-pressure melts (Xue et al., 1991). In most natural melts, the ratio of metal oxide to A1203 is higher than 1 and A1 enters the network in fourfold coordination. Any increase in A1 coordination with pressure is likely to lead to a decrease in the high-pressure viscosity of these alumino-silicate melts. Alumino-silicates, covering the chemistry of andesitic to basaltic magmas, form tetrahedral networks with extremely high viscosities. Under pressure, the viscosity decreases and for this reason, at depth, the mobility of the melt can be higher by several orders of magnitude. This viscosity decrease is related to the pressure-induced increase in the A1 coordination and the resultant weakening of the A1-O bond strength by a change in the bond angles in the alumino-silicate tetrahedral network. The viscous flow of alkali-alumino-silicate melts is constrained by oxygen exchange between polymeric units (Poe et al., 1992). In this flow, structural relaxation may occur through the formation of transient five-coordinated (Si or A1) species. At T close to their glass transitions, liquid SiO2 and SiO2-Na20 melts have configurational entropies attributed to their bridging and non-bridging O atoms. The decrease in the specific volume with the decrease in polymerizaton causes the P-sensitivity of the viscosity of silicate melts. Using some assumptions, Bottinga and

Silicate Melts and Rocks

801

Ricket (1995) showed that the Kushiro discovery (1976, and subsequent papers) can be explained with the Adam and Gibbs theory (J. Chem. Phys., 43, 139-146, 1965) for the P-sensitivity of polymerization and the change in appropriate configurational entropy. For modelling the physical chemistry of the melts and for parametrizing melt viscosities for modelling of igneous processes, complete viscosity-temperature relationships are needed. This is all the more so because of the non-Arrhenian temperature dependence of the viscosity of most melts. This non-Arrhenian temperature dependence remains one of the main hindrances in obtaining a fully generalized model of melt viscosities for petrological calculations.

11.3.7. Density and viscosity determination Silicate melt densities at high pressure are usually estimated using molecular dynamic simulations and elastic property measurements at low pressure. Dynamic determinations using shock-wave techniques and static determinations using the "falling sphere" technique (density + viscosity) and the "sink"/"float" method (density) have been carried out on silicate melts to determine the variation of density as a function of pressure. In the falling sphere technique, the density is determined by measuring the distance of sinking (or floating) of a crystalline or metal sphere as a function of time and is calculated using Stoke's Law. In such studies, spheres of such minerals as ruby, forsterite, diamond, graphite, platinum, etc., have been used. A centrifuge can also be employed to accelerate the Stokesian settling velocity by up to a factor of 1,000-1,500. This accelerated falling sphere technique is used to provide data in the viscosity range of 105-108 Pa s. The experiments by Zhen-Ming Jun (1994) suggest an effective mantle viscosity to be around 1012 Pa s (the ratio of the stress about 107 Pa to the strain rate is --~ 10 -5 s-l), which is some six orders of magnitude lower than that commonly assumed for the Earth' s asthenosphere. A drastic lowering of the viscosity of silicate rock by the presence of water has important geodynamic consequences. A strong partitioning of water in melts migrating upwards leaves behind a largely dehydrated residual solid. This dehydration strengthening may be the reason for the narrowness of the upwelling zone beneath ridges. Dehydrated viscous melt is sufficiently viscous to keep lateral comer-flow gradients capable of focusing melt into a narrow ridge axis (Hirth and Kohlsted, 1996; Braun et al., 2000).

11.3.8. Melt percolation In the upper mantle, small amounts of melt may distribute so as to form an inter-connected framework of tubules (such as along the triple junction where olivine grains meet). However, to a certain extent, the melt equilibrates chemically with the solid through which it migrates. The melt migrates only a short distance by percolation before finding its way into broader channels where it can move

802

Chapter 11

rapidly upwards (Iwamori, 1993). However, when, under stress, the sample is stretched a little, the melt redistributes, spreading to grain boundaries. This melt redistribution greatly reduces the strength of the matrix, aiding the escape of the melt. A partial melt channel with locally high strain rate develops a greatly increased ability for upwelling of the material. Partial melting and separation of melt through melt percolation led to thermal expansion, contraction and volcanism, all of which are linked to plate tectonism. As the less dense hot material rises and displaces the dense material downwards, the total gravitational energy is lowered. Gravitational energy dominates in large-scale processes within the Earth and other planets. The effects of solid-solid phase transitions on melt productivity during isentropic pressure-release melting were investigated by Ashimow (1995). The principles governing these effects were developed by graphical construction in one- or two-component systems. These constructions show that solid-solid transitions with positive Clapeyron slopes diminish rather than enhance melt productivity.

11.3.9. Crystal- melt phase equilibria Crystal-melt phase equilibria are also affected by pressure. For example, enstatite (MgSiO3) melts incongruously to forsterite and liquid at --<0.3 GPa pressure but, at higher pressure, it melts incongruently. Olivine (in basalt) is typically a liquidus phase at crustal pressures but, at upper mantle pressure, pyroxene becomes a liquidus phase. Other pressure effects include cation diffusion, trace-element partitioning in crystal-liquid and redox ratios of iron (Mao et al., 1981). The inter-tetrahedral angles in the three-dimensional network decrease by several percent under pressure. A significant reduction in cavity volume will most likely affect the viscosity diffusion property, etc., in addition to the density.

11.3.9.1. fo2, fn20 and amo Oxygen fugacities, fo2, of mantle-derived melts and mineral assemblages are commonly considered to lie within approximately _+2 - 3 log units of NNO buffer curve but fugacities up to that of the H - M buffer have been estimated in some cases (e.g., Righter and Carmichael, 1993). There is little quantitative documentation of aH20 for derived melts and mineral assemblages. Bell (1994), however, has determined log aFi~o in the range - 2 to - 3 based on the H contents of a site of olivines from xenoliths (entrained in alkali basalts) from a number of localities. The low H20 activities may represent low wt% H20 contents in the co-existing magmas under vapour-absent conditions. Under vapour-present conditions, low ai-i~o would result if H20 represents only a small mole fraction of all species in a C - O - H fluid (Poppet al., 1995).

803

Silicate Melts and Rocks

11.4. H20 in silicate melts

H20 dissolves both as OH and as molecular H20 in silicate melts (e.g., Stolper, 1982). Speciation of H20 in melts and glasses poses a problem. Dingwell and Webb (1990) suggested that the H20 speciation in a glass represents the equilibrium at the bulk Tg of the glass. Shen and Keppler (1995) have reported direct FTIR measurements of H20 speciation in hydrous silicate melts to 1,000~ and 1 GPa in a DAC designed for FTIR studies. A peraluminious sodium silicate melt with 3 0 - 4 0 wt% HzO shows the presence of both OH and H20. The standard enthalpy of the speciation reaction H20 -k- O -- 2OH is very different for the melt and glass phases. The temperature dependence of the H20 speciation equilibrium can be described by two equations: (a) for glass phase: AH -- 1.6kJ/mol, and lnK = -2.65 - 1.91 • 102 KT -1 (b) for melt phase: AH = 30.3 kJ/mol, and lnK = 3.04 - 3.64 X 103 KT -1 The intersection of these two equilibrium curves defines a glass-transformation temperature of 335~ The water contents of glasses from MOR in close proximity to the Azores and Iceland hotspots are enriched by factors of 2 - 4 relative to the surrounding mantle (Kingsley and Schitting, 1995). Hence, these two areas are referred to as mantle "wetspots". 11.4.1. K 2 0 - S i O 2 - H 2 0 system The solubility and solubility mechanisms of H20 have been determined by Mysen (1998) for melts in the system K z O - S i O 2 - H 2 0 from 0.8 to 2.0 GPa in the 700-1,100~ temperature range. For the most potassic composition studied, KzSiOs, nearly complete miscibility between hydrous melt and silicate-saturated aqueous fluids were observed at 2.0 GPa. The H20 solubility and silicate solubility are strongly non-linear but show positive functions of pressure. The temperature dependence becomes more pronounced with increasing pressure and with increasing K/Si. The experimental study by Mysen (1998) reveals that the in situ high-T/high-P Raman spectroscopic information for potassium silicate melts is consistent with a solution mechanism for H20 in potassium silicate melts, which can be schematically summarized as: KzO.nSiO2 + H20 ~ Q~

+ QZ(K) + Q3 (K) + K...OH

(11-7)

In this expression, Q~ denotes a unit with no bridging oxygens and where all the NBOs are associated with H + (thus in effect producing a H4SiO4 species), the QZ(K) and Q3(K) denote structural units with two and three bridging oxygens, respectively, where the NBOs are associated with K +. The silicate-solution mechanism in aqueous fluid can possibly be described by a relation like that presented in equation (11-7). The phase relations along the joins KzSiO4-H20 and K2SiO9-H20 under pressure (Mysen,1998) show an immiscibility between potassium silicate melts and fluids. But the

804

Chapter 11

immiscibility gap shrinks rapidly with increasing pressure and with increasing K/Si ratio. Strangely, however, the gap continues to exist at 2 GPa.

11.5. REE patterns The depth constraints on the formation of tonalitic magmas in the continental crust are provided by REE patterns of the synthetic melts calculated from the REE abundances in metagabbro and metabasalt. The REE pattern of tonalites from active continental margins and Archean TTG associations show low values in REE, with LaN (chondritic normalized) of 10-30 and YbN of 1-2. These values are reproduced at 1-1.25 GPa pressure from metagabbro, which displays a slightly low REE-enriched pattern with LaN = 8 and YbN = 3 (see Springer and Seck, 1997). 11.5.1. Fe 3+ in glass The content of Fe 3+ in melts and glasses is a strong function of T, fo2 and composition. At high values off%, Fe 3+ is in tetrahedral coordination in many silicate melts. It has been demonstrated by Virgo and Mysen (1985) that Fe 3+ transforms gradually to octahedral coordination, either when oxygen fugacity decreases at constant temperature or when temperature increases at constant fo2 if, in the melt, the Fe3+/EFe is less than 0.5. Fe3+/EFe decreases initially with increasing pressure in glasses quenched from N a 2 0 - A 1 2 0 3 - S i O 2 - F e - O melt (Fig. 11.3). In alumina-bearing systems, however, the melt shows greater oxidation at higher pressures. The quenched melts (glasses) were analysed using 57Fe Mrssbauer spectroscopy by Brearly (1990), who observed pressure changes of the spectral envelopes. The changes may result from an increased distortion of Fe 3+ (IV).

11.5.2. Partition coefficient in melt/solid The partition behaviour of elements under pressure is controlled not only by P and T but also by composition. The effect of temperature on the partition (or rather, distribution) coefficient, D, is estimated by thermodynamic considerations (e.g., Murthy, 1992). By assuming Nernst partitioning and neglecting the non-ideality of the system, the equilibrium condition is expressed as:

txL + RT In C L = / z s + RT In CSx

(11-8)

where R is gas constant, T temperature and /xL and ~s represent standard chemical potentials of component X of the liquid phase and solid phase, respectively. Then the partition coefficient is expressed as:

l n D - (IxL -/xSx)RT

(11-9)

This equation predicts that the partition coefficient approaches 1 with increase in temperature. It is therefore expected that the relative difference of the elements in D will

Silicate Melts and Rocks

805

Fe 3+

7__Fe ~.0

lt+50 o C

0.8

0.6 Fe

0.4

- O. 9 ~_Fe

-0.7 -0.5

Fe3+ ~Fe 0.7

.25

-0.3 0.5

x-0.15

0.3

0.1

1 0

A 10

20

I

30

/,13

Pressure~k l:~r

Figure 11.3. Fe3+/'ZFe of quenched glasses as a function of pressure and composition of starting material (X -- mole fraction of acmite component of the starting material) at 1,450~ (Mysen and Virgo, 1978).

decrease with increase in temperature. Since the liquidus temperature increases with increase in pressure, all the values should be approaching to 1 under high temperature if the effects of temperature alone are considered. The partitioning of selected trace elements between peridotite minerals (such as olivines, orthopyroxene, clinopyroxene and phlogopite) and carbonatite melt at 1.84.6 GPa pressure was investigated by Sweeney et al. (1995). They determined the partition coefficients for trace-element partitioning between olivine-orthopyroxene-carbonatite melt. Earlier (Bell et al., 1979; Heinz and Jeanloz, 1987), it had been observed that Fe partitions into melt and the oxide structure relative to perovskite structure. Between H20-rich fluids and natural peridotite (olivine + orthopyroxene + clinopyroxene + accessory minerals), the partition coefficients of Rb, Nb, La, Sr, Sm, Zr, Tm and Y were measured by Ayers et al. (1997) at 2.0-3.0 GPa and 900-1,100~ The stable sub-solidus assemblages observed were: at 2.0G Pa ----, spinel-lherzolite (+_rutile) at 3.0G Pa ----,garnet-peridotite + zircon

806

Chapter 11

The high solubility of TiO2 is seen in spinel-lherzolite at high temperature. Therefore, rutile should hardly be present at the source region of island-arc basalts during melting. Peridotite/fluid bulk-partition coefficients at 2.0 GPa range from --~0.1 for Rb to --- 100 for Tm. Fluid/melt partition coefficients (D fluid/melt) have been reported by Ayers and Eggler (1995) for Ca, Mg, Na, K, Rb, Sr, La, Sm, Y and Tm at 1.5 and 2.0 GPa at 1,250~ They reported that coefficient D fluid/melt ranges from 0.43 to 1.31 for these elements (and Ti), which suggests a near-congruent dissolution of melt in fluid.

11.6. Rocks under pressure 11.6.1. Transformation under shock: pseudotachylites Impacting of meteorites on planetary bodies creates enormous transient pressure on rocks (or ices) which causes phase transition, deformation, melting and vaporization. Such impacted regions are characterized by shattered cones and pseudotachylites (pulverized rocks melted by impact or frictional melting). At Sudbury, impact-structure pseudotachylites (B type) are seen. Coesite and stishovite have been reported in the pseudotachylites in quartzites from the Vredefort impact structure in South Africa (Martini, 1978). In laboratory hyper-velocity impact experiments of Fiske et al. (1995), quartz was shock-loaded from 42 to 56 GPa by using the 6.5 m two-stage light gas gun at Lawrence Livermore National Laboratory. In this experiment, the melting of SiO2 provides a lower limit on the temperature of the pseudotachylite (2,000 K). These temperatures are similar to those required to reduce SiO2 in soils struck by lightning (Essene and Fisher, 1986). During a meteorite impact, the compression and development of a hemi-spherical transient crater cavity involves bulk plastic deformation for a time of 10 - 3 - 1 0 ~ s. Plastic flow and pseudotachylite formation are important processes in meteorite impacts. The central part of large impact structures may have a substantial static thermal metamorphic overprint. This explains the increase in crater melt with crater size. Such thermal structures have, however, been attributed (wrongly?) by some to endogenic processes.

11.6.2. Terrigenous and pelagic sediments under subduction Subduction of the continental crust along with the terrigeneous and pelagic sedimentary plays a major role in the evolution of the crust-mantle geochemical system (see also Section 2.10.2). The evidence of subducted sediments has been well observed in many calc-alkaline magmas that erupted in island arcs and continental margins from source regions at depths of 100-150 km. The isotopic signatures of Pb, Sr and Nd in many intra-plate magmas indicate that these were derived from subducted terrigeneous and pelagic sediments (e.g., Tera et al., 1986).

Silicate M e l t s a n d R o c k s

807

About 1.6 km3/yr of terrigenous and pelagic sediments are subducted into the mantle (Von Huene and Scholl, 1991). Some are returned to the crust via underplating and calcalkaline magmatism, while some go to greater depths to be incorporated into the mantle. Subduction of terrigeneous materials may also occur along the zones of continental collision. Recent discoveries of coesite in high-grade metamorphic rocks indicate that large, coherent volumes of quartzose-felspathic continental crust may be subducted to depths exceeding 100 km (Schreyer et al., 1987). Also, the reports of inclusions of microdiamonds in garnets from these metamorphosed lithologies suggest that the subduction in such cases reached depths of---120 km (e.g., Sobolev and Shatsky, 1990). Evidently, buoyancy forces at these depths become sufficient to inhibit the transportation of these sediments to greater depths via entrainment by subducting slabs. The depth of subduction is, therefore, delimited by the buoyancy of these materials. At shallow depths (<150 km), terrigeneous material is buoyant relative to the surrounding mantle and opposes entrainment by the subducting slab. In isolated cases, the entrainment continues and is reflected by the occurrences of coesite and diamond in some materials which were later pushed up to the surface (Irifune et al., 1994) (Fig. 11.4, from Liou et al., 1998; see inset). The subducting slab includes crustal-hosted fragments and mantle wedge blocks. The VHP and HP slabs retumed to shallow depths after recrystallization within coesite or diamond stability fields at depths > 100 km due to slab breakoff and the buoyancy of a low-density continental sheet. However, small rafts of fragmented continental blocks on the oceanic plates may be dragged deeper along the subduction zones. This process was very significant in the early past. The compositions of the argillaceous and siliceous facies of anhydrous pelagic sediments (Table 11.2) are sufficiently close to that of the upper crust. Irifune et al. (1994) experimented on a material of continental crust composition, similar to that of the average upper continental crust. During partial melting of continental crust lithologies at relatively low pressures (5-10 GPa), the early formed orthoclase, wadeite and K-hollandite are eliminated near the solidus, whereas the stability field of Na-clinopyroxene extends to temperatures well above the solidus. Liquidus in this pressure interval shows high K/Na ratios and high SiO2 contents. At higher pressures (16-24 GPa), the stability fields of K-hollandite and stishovite extend towards the liquidus. The resulting partial melts show lower K/Na ratios and Si02 contents.

11.6.2.1. Density change and buoyancy At depths with pressures lower than 6 GPa, the major phases in the crustal compositions are orthoclase, coesite and clinopyroxene. Near to 6.5 GPa, orthoclase transforms to wadeite + kyanite + coesite and the density of the assemblage increases to 3.2-3.3 g/cm 3. Around 9 GPa, K-hollandite and stishovite become major stable phases. At 10 GPa, the density of the assemblage becomes --~3.7-3.9 g/cm 3 (compared with pyrolite at --~3.4 g/cm3). The density of crustal compositions remains substantially higher that that of pyrolite up to --~24 GPa, i.e., near the bottom of the transition zone (--~660 km). Below this depth, however, the stated density relationship is reversed.

Chapter 11

808

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Figure 11.4. A tectonic model and P - T time path for subduction and exhumation of crustals fragments from continental lithosphere using for example the Dabie-Sulu collision zone. The required time between subduction and exhumation is about 10 million years (modified after Liou et al., 1998, 9 1998 Mineralogical Society of America). In m a n y cases, b u o y a n c y c a u s e s s u b d u c t e d s e d i m e n t s along w i t h the slab materials to r e t u r n to the n e a r - s u r f a c e crustal e n v i r o n m e n t 9 E v i d e n c e o f such a h i s t o r y is offered by the p r e s e n c e o f c o e s i t e and d i a m o n d . In o t h e r cases, the g e o c h e m i s t r y o f s o m e oceanic basalts has o f f e r e d e v i d e n c e for the c o n t i n e n t a l l y d e r i v e d r o c k s and s e d i m e n t s to have b e e n

TABLE 11.2 Chemical composition of continental crust and pelagic sediments

SiO2 TiOe AleO3 FeO a MgO CaO Na20 K20

Continental crust (Taylor and McLennan, 1985)

Pelagic sediments Argillaceous clay facies (Chester, 1990)

Siliceous facies (Chester, 1990)

66.0 (66.2) 0.5 (0.6) 15.2 (15.9) 4.5 (4.6) 2.2 (1.9) 4.2 (4.8) 3.9 (2.7) 3.4 (3.3)

60.50 0.92 19.50 8.17 1.50 4.18 1.67 3.56

70.80 0.72 14.73 6.38 1.51 2.71 1.04 2.10

Values in parentheses represent the results of electron microprobe analyses of the starting material actually used in the present study (FeO was replaced by COO). aAll iron has been calculated as FeO.

Silicate Melts and Rocks

809

transported down to depths of ---200 km. Once the critical depth of ---200 km is reached, the subduction of the continentally derived material may attain the density (and loss of buoyancy) for it to sink deep into the mantle. If significant amounts of associated water are subducted to depths greater than 150 km, partial melting of the subducted slab would occur. The melt would react and hybridize with the surrounding mantle, modifying the properties of the magmas which are derived from the mantle. Several lines of evidence suggest that significant amounts of water may be subducted to depths of 150-160 km (e.g., Thompson, 1992). Serpentinite underlying oceanic crust may provide the most plausible sources of water during deep subduction. H20 is thereby released and is added to that formed by the process of successive dehydration of dense hydrated magnesium silicates (DHMS). This would ascend and cause partial melting of the subducted oceanic crust along its overlying continentally derived lithologies. 11.6.2.2. Potassium mobility in subduction pressures The upper continental crustal rocks at low pressures (--~6-8 GPa) and near to solidus are seen to eliminate orthoclase and wadeite (Irifune et al., 1994). Thus, the partial melts are expected to be K-rich but, because the partition coefficient of Na in crystal/liquid is high (>--~ 2.5), Na-rich clinopyroxene exists throughout much of the melting interval. Accordingly, melts formed by a modest degree of partial melting in the low-pressure region have high K20 contents (up to 6%) and high K20]Na20 ratios (--~2-3). These results suggest that partial melts of subducted terrigenous lithologies between depths of --~ 150-250 km could be partially responsible for potassium-enrichment processes, which are believed to be widespread in the sub-continental lithosphere. They may also contribute to the systematic increase in K20 contents of calc-alkaline magmas. This is observed to occur as the distances from their eruption centres to the underlying Benioff-Wadati zones increase. Small degrees of partial melting with increasing depth cause enrichment of K20 in the melts (Irifune et al., 1994). However, at higher pressures between 10 and 21 GPa, the K-hollandite field expands during partial melting while that of Na-clinopyroxene contracts. At 21 GPa (/1,900~ the partial melt contains only 1.4% K20, compared with 7.1% Na20, and hollandite as the refractory residue continues to enlarge. The SiO2 content in the melt also reduces (--~55% SiO2) through expansion of the stishovite field. 11.6.2.3. Lead paradox Potassium-hollandite strongly hosts Pb (and La) but not U or REE (Kesson and White, 1986). Consequently, the liquid after partial melting will be richer in the U/Pb ratio when K-hollandite is present and lower in LREE/HREE, K/U and La/Nb ratios compared with the initial compositions (Irifune et al., 1994). The capacity for retaining Pb by residual K-hollandite during partial melting of subducted continentally derived lithologies at high pressures is significant. In the early history of the Earth (4.0-2.0 b.y. ago), the subduction of continentally derived lithologies might have occurred on a larger scale than exists at present and these were carried to greater depths ( 3 0 0 - 6 0 0 k m ) where they

810

Chapter 11

experienced partial melting. The K-hollandite phase as residium or the megaliths would retain Pb in its structure and concentrate at the boundary layer overlying the 660 km discontinuity. These would ultimately have sunk into the lower mantle (Allbgre and Turcotte, 1985). The extraction of these lithologies from the crust to the upper mantle may have caused the over-abundance of U in them compared with Pb. This is known as the 'lead paradox'. Indeed, Class et al. (1993) found that magmas that erupted along the hot-spot trail of the Ninety-EaSt Ridge in the Indian Ocean were characterized by an exceptionally high U/Pb ratio (/x --~ 30-60). The lead paradox with reference to calc-alkali magmatism has been discussed in Section 2.8.5.2.

11.6.2.4. Subducting slabs In subducted slabs where the temperatures are lower than the average mantle temperatures, volatile-bearing phases could be present. At depths of temperatures < 1,500 ~ C, the content of H20 in [3-phase (wadsleyite) is suggested to be nearly as high as that of nominally hydrous phases. And, since under hydrous conditions the stability of [3-phase + stishovite is much increased, they are likely to be the major constituents of the slabs sinking through the transition zone. During subduction, as the slab warms up to sufficiently high temperatures, hydrous [3-phase + stishovite would melt incongruently to garnet + melt. This melt may serve as a lubricant between the slab and the enclosing mantle. A sudden movement of the slab through lubrication by the volatile-rich melt from the mantle rock would cause a deep-focus earthquake. Such melting can also occur through the introduction of volatiles to the transition zone from the lower mantle or from the subducted slabs. Oceanic lithosphere may show a trend of increasing olivine at depth, corresponding to a change from basalt at the upper part to residual dunite at the base. This subducting lithosphere would show a gradual change with depth from garnet to [3-phase/spinel. Such a change would account for the observed global presence of a high-velocity gradient in the transition zone. Hence, the transition zone may serve as the shallow graveyard (initial to the D" zone) for the subducting lithosphere (see also Section 2.10).

11.6.3. Ultra-high-pressure metamorphism Metamorphism occurring at pressures greater than ---2.5 GPa (---80-90 km depth) is designated as an ultra-high-pressure or very high-pressure (UHP or VHP) process (see also Section 1.2.3.2).

11.6.3.1. Coesite-diamond The rocks formed at a UHP of 3 - 4 GPa ( ~ 100-120 km depths) and showing the occurrence of coesite and diamond are believed to be the crustal material which has been transported back to the surface (Xu et al., 1992; Okay et al., 1993).

Silicate Melts and Rocks

811

An appropriate chemistry through UHP at great depths can produce coesite plus micro-diamond. In garnet-peridotite, the assemblage is of magnesite + diopside _+Ticlinohumite and in eclogite talc is formed. UHP regimes offer important clues towards an understanding of subduction and continental collision. These help to bridge the gap in our knowledge about the upper mantle and crustal processes. The VHP rocks reveal a complete record of geodynamic pathways and offer the constraints for the mechanism of subduction and tectonic exhumation. However, coesite and diamond, as UHP minerals of mantle origin, have long been recognized in kimberlite pipes and in meteorite craters. The discovery of coesite and microdiamonds in UHP crustal rocks has revolutionized our understanding of continental collision zones and mantle dynamics attending subduction. In large UHP terranes, widespread (but minor) occurrences of garnet peridotites are seen which may represent fragments of mantle wedge overlying the subduction zones in convergent-plate boundaries. 11.6.3.2. Crustal metamorphic regimes For HP and UHP metamorphism of the crustal rocks, the maximum temperatures noted are in the range 750-800~ Above these temperatures, granitic melt + migmatite may be generated when H20 is available. At successive P and T, the assemblages that appear are: granulite ~ amphibolite and epidote amphibolite ~ green-schist. While the green-schist assemblage is stable at low pressure, blue-schist facies occur between 0.5 and 1.6 GPa and T --~ 400-450~ The metamorphic regimes along with geotherms of 5~ (extremely high P/T) and 20~ (ancient cratons) are shown in Fig. 11.5 (source Liou et al., 1998). Stabilities of diamond (Bundy 1980), coesite (Hemingway et al., 1998), glaucophane (Holland, 1988), jadeite + quartz (Holland, 1980), A12SiO5 (Bohlen et al., 1991), paragonite (Holland, 1979) and aragonite (Hacker et al., 1992) and the minimum melting of granitic and tonalite solidus (Huang and Wyllie, 1975) are also shown. 11.6.3.3. Hot and cold eclogites: coUision/subduction zones The VHP terranes can be classed as (a) 'hot eclogite' formed at P greater than coesite ---. quartz transition and (b)' cold eclogite' formed at P lower than coesite ---, quartz transition. HP blue-schist and 'cold eclogite' belts have been seen to occur in tectonic contact with the coesite-bearing VHP belt. These segments were evidently formed from a subduction zone during, or even prior to continent-continent collision (Liou et al., 1998). The VHP terranes, manifesting solid-state crystallization in high P and moderate T tectonic settings, are seen with major continental collision belts in Eurasia (and Africa) and are confined to Alpine-type orogens. These are mostly exhumed units, show lithologies mainly of eclogites and garnet peridotites (included in pods and slabs) and bear geochemical characteristics which are sub-continental in nature. The occurrences are commonly associated with late-stage granitic plutons. The VHP phases like microdiamonds are seen to occur within garnet and zircon while coesite occurs in garnet and omphacite.

Sulawesi Indonesia subduction zone. Recently, coesite inclusions in zircon have been reported from eclogitic rocks from Sulawesi Indonesia (Parkinson et al., 1998).

812

Chapter 11

Figure 11.5. P - T regimes corresponding to various metamorphic types: (1) Very high-P (VHP), (2) High-P, and (3) Low-P. Geotherm of 5~ and 20~ are indicated. P - T boundaries of various metamorphic facies are from Spear (1993) and subdivision of the eclogite field into amphibole eclogite, epidote eclogite, lawsonite eclogite and dry eclogite are from Okamoto and Maruyama (published in 1999) (Source: Liou et al., 1998, 9 1998 Mineralogical Society of America).

The terranes show metamorphic phases which appear at depths of--~ 100 km (e.g., Coleman and Wang, 1995). Phase-equilibrium constraints and thermobarometric calculations reveal that the peak temperatures were ---700-900~ at confining pressures greater than 2.8-4.0 GPa. These conditions may reflect a low geotherm < 7 - 8 ~ (present in subduction-zone environments).

11.6.3.4. D a b i e - S u l u collision zone

The most recognized VHP terrane in the world occurs in the Dabie-Sulu collision zone (210-240 Ma), in east central China. The UHP terranes, with areas of up to

Silicate Melts and Rocks

813

20,000 km a, consist mostly of crustal rocks. In these rocks, inclusions of high-pressure minerals such as coesite and/or micro-diamond bear an imprint of pressure to over 3 GPa during aborted subduction to depths as great as 135 km. This has been thoroughly investigated by Liou and his co-workers. Their tectonic model and the P - T time path for the terrane are illustrated in Fig. 11.7 (see caption). In the Dabie-Sulu terrane, the pre-Cambrian protoliths (granitic, pelitic, psammitic, carbonate and minor mafic-ultramafic rocks) were subjected to VHP metamorphism (> 2.5 GPa) at mantle depths through subduction prior to and during the Triassic collision of the Yangtze and Sino-Korean cratons. In this region, the rocks in the Dabie Mountains show the presence of coesite and micro-diamond (?) in eclogites and other metamorphosed crustal rocks (e.g., Liou et al., 1996). The model for the origins of Dabie-Sulu peridotites in the collision between the Sino-Korean and Yangtze plates and the cold subducting plate reaching UHP is schematically shown in Fig. 11.4 (read the caption). The Dabie-Sulu VHP rocks are unique in the occurrence of (Liou et al., 1998): (1) abundant coesite and hydrous phases (such as talc, zoisite/epidote, nyb6ite and phengite) in eclogite (Zhang et al., 1995a,c), (2) the world's lowest ~lSo values (ruffle shows - 1 5 per mil) for mineral separates from eclogites and meta-sediments (Zhang et al., 1998; Rumble and Yui, 1998), (3) abundant garnet peridotites of mantle origin (Zhang and Liou, 1998) and (4) abundant exsolution textures in VHP minerals from garnet peridotite and eclogite (e.g., Zhang and Liou, 1997). For such VHP complexes, a descent of ancient cold sialic crust overlying a subducting lithosphere is indicated (Ernst and Peacosk, 1996). Through buoyancy, the crust may decouple from the downgoing slab and experience adiabatic decompression while traversing up through the P - T regime of granulite--~ amphibolite facies (600800~ 0.3-1.0 GPa). However, some abnormally high-P peridotites may be transported by mantle convection from great depths to the subduction zone and be incorporated into the subducting continental crust. The North Dabie Complex (NDC) consists of granitic to monzonitic plutons of Cretaceous age intruding into amphibolite-facies (ortho) gneisses. The mineral assemblages and compositions point to an early eclogite-facies metamorphism (--~800- 820~ The C a - N a clinopyroxene present shows oriented quartz needles (---20-20 Ixm wide, --~5-200 lxm long) within it. This implies the prior existence of a non-stoichiometric "supersilicic" omphacite stabilized at UHP (> 2.5 GPa) conditions, although no evidence for coesite is noted. The SiO2 needles are interpreted by Tsai and Liou (2000) to be an exsolution from a precursor, non-stoichiometric omphacitic clinopyroxene, which contained excess silica at the peak metamorphic condition. The implication of excess silica in clinopyroxene has been discussed in the Section "Supersilicic clinopyroxene" in Chapter 6.

11.6.3.5. Alpe Arami UHP lherzolite The garnet lherzolite from the Alpe Arami peridotite massif (400 • 1,100 m 2) of the Central Alps may have been metamorphosed at much higher pressures of 10-15 GPa

814

Chapter 11

( ~ 300-450 km depths) (Dobrzhinetskaya et al., 1996). This implies that pieces of mantletransition zone (10-13 GPa) could be transported to the Earth's surface. The basis for considering this lherzolite to be of transition-zone origin lies in the discovery of (i) a lattice preferred orientation (LPO) of olivine, (ii) a few ixm-size FeTiO3 rods, topotactic with the host olivine, and (iii) high TiO2 (inferred) contents of olivine (Dobrzhinetskaya et al., 1996). FeTiO3 rods in crystallographic structures are ilmenite and the structures are intermediate between ilmenite and perovskite. These rods are hypothesized to have exsolved (along [010] direction of host olivine) at 10-15 GPa pressure ( ~ 300-450 km) (Mehta et al., 1994), which is high enough for olivine to transform from [3- to ~/-phase and manifest a LPO of these phases (wadsleyite or ringwoodite). The great abundance of FeTiO3 and spinel precipitates in first-generation olivine in Alpe Arami have no known counterpart in any other peridotite massif or xenolith. However, commenting on the work of Dobrzhinetskaya et al. (1996), Hacker et al. (1997) opined that the Chinese or Alpine lherzolite experienced maximum pressures only > 4 - 5 GPa (i.e., not 10-15 GPa). They suggested that the stu0v needs some additional (see Fig. 11.6, top center) evidence as is also required after the discovery of some such high-pressure polymorphs as: TiO2 with e~-PbO2 structure (stable at 5 - 7 GPa), high-pressure C2/c clinopyroxene (6GPa) and majorite garnet (7 GPa) (Angel et al., 1992).

Temperature (~ 400 I

600 I

800 I

1000 1

1200 ~

1400 i

o

n 9

,

lOO

I

& Q. t9 0,,

200

~.~

o. olivine

%,

c~ r

10

soo ~::T

r

t,,=

ul

3

o,, t... 0..15-

400

SO0 20600

1200

1600

Temperature ( K ) Figure 11.6. Pressure-temperature stability fields for Mgt.sFeo.2SiO4 (Mehta et al., 1994) and FeTiO3 (Helfrich et al., 1989). Conditions within the upper mantle range from the coldest subducting lithosphere to the subcontinental upper mantle, as shown by dotted lines. Box shows equilibrium pressure and temperature of silicate minerals in Chinese garnet lherzolite (Hacker et al., 1997).

Silicate Melts and Rocks

815

11.6.3.6. Exsolutions in VHP minerals Exsolution textures are common in VHP minerals obtained from both eclogitic and ultramafic rocks. These are: (i) ilmenite exsolution: such as rods, (ii) magnetite exsolution: such as plates in olivine and clinohumite, (iii) quartz exsolution: such as rods in clinopyroxene omphacite, (iv) monazite exsolution: such as lamellae in apatite, and (v) rutile exsolution: such as needles in garnet. Exsolution of lamellae in VHP minerals may have taken place during nearly isothermal decompression, in contrast to exsolution with falling temperature in primary igneous minerals (Liou et al., 1998). llmenite rods in olivine

Alpe Armi massif in Central Alps. Micron-sized ilmenite rods in olivine have recently been identified in garnet peridotites from several VHP terranes (Fig. 11.6, from Hacker et al., 1997). TEM study of Alpe Arami garnet lherzolite revealed that the FeTiO3 rods in olivine are topotactic with the host, parallel to its [010] direction and presumed to have exsolved at 10-15 GPa (300-450 km) as perovskite, followed by variable conversion to ilmenite. The unique preferred orientation of Alpe Arami olivine may have formed during recrystallization of a wadsleyite- or ringwoodite-bearing protolith (Green and Dobrzhinetskaya, 1997). Thus, it can be argued that the Alpe Arami complex was possibly derived from the transition zone in the mantle. Sulu garnet peridotites in Chijiadian (China). The olivine grains in the Chijiadian lherzolite show remarkably homogeneous distribution of rod-like inclusions of titanates (Feo.82Mg0.15Mno.o3TiO3), which are more magnesian than the Alpe Arami iron titanates (Feo.94Mgo.o6TiO3). The long axes of the inclusions are parallel to [010] of the host olivine crystals. TiO2 contents of olivine from the Alpe Arami and Sulu terrane are seen to be of values < 300 ppm, which implies P - T conditions --~4 - 6 GPa and 780-820~ (Nakajima and Ogasawara, 1997). Magnetite exsolutions. From some mafic-ultra-mafic rocks, magnetite rods/plates have been reported to have exsolved in olivine and clinohumite. One such case is Dabie harzburgite and garnet pyroxenite. The topotaxial intergrowth between olivine and oriented magnetite lamellae is as follows: [220]magll[220]ol [111]magll[33 i]ol [11 i]magll[331]ol [242]magII[220]ol This offers evidence for an exsolution phenomenon involving the binary system Fe304- (Fe,Mg)2SiO4. If the original (Fe,Mg)2SiO4 phase was actually [3-phase (wadsleyite) with distorted spinel structure, it would accommodate Fe304 as spinel solid solution along the binary join

816

Chapter 11

Fe304-(Fe,Mg)2SiO4. During decompression, the transformation of wadsleyite ([3-phase) to olivine (a-phase), exsolution of excess Fe304 (magnetite) takes place bearing the topotaxial relationship. Under pressure, a significant amount of Fe 3+ can be incorporated into a wadsleyite structure through substitution of 2Fe 3+ -- Fe 2+ + Si 4+. The difference in stability of Fe 3+ wadsleyite (5 .0- 6.0 GPa at 1,100-1,200~ is adequate to stabilize Fe3+-enriched (Mg,Fe)2SiO4 to lower pressure (shallower depth) compared with a Fe3+-poor system. Evidently, such a relationship would alter the depth of the '410 km discontinuity' in the mantle (Woodland and Angel, 1998).

Silica rods. The nature of silica rods in clinopyroxenes has already been discussed in the Section "Supersilicic clinopyroxene" in Chapter 6. Monazite lamellae. Monazite lamellae in apatite have been discussed elsewhere. 11.6.4. Basalts and eclogites The melting T of eclogite increases with P whereas potassic basalt shows this characteristic only at 1.5-2.5 GPa and at > 3.0 GPa. However, between 2.5 and 3.0 GPa, the melting T decreases with P. If the whole deeper mantle is composed of eclogite (Anderson, 1979), then a sharp discontinuity would appear at 400 km depth (~ 13 GPa). The composition of the rocks in the lithosphere and the types of hydrous mineral and their stable P - T conditions are important factors controlling the melt behaviour of rocks. This may explain the partial melting of rocks and the origin of the low-velocity zone in the deep lithosphere. Eclogite with a jadeite-rich (Na-Al-rich) pyroxene shows a density of---3.33.8 g/cm 3. This metamorphic equivalent of basalt is composed primarily of garnet and omphacite. Eclogites such as from the Fransiscan tectonic block, experiencing retrograde and prograde metamorphism to eclogite facies, consist of garnet, glaucophane, phengite, albite, quartz and rutile, which record peak metamorphism at pressures ---1 GPa and temperatures --~300~ (Oh and Liou, 1990). The other distinctive rock types associated with these are blue-schists and garnet amphibolite. A typical P - T path for this complex is shown in Fig. 11.7 (Source: Mysen et al., 1998) (see caption). Dissolution of omphacite in the co-existing garnet is minimal at 13 GPa and is complete at 17 GPa when eclogite transforms to garnetite. The formation of the pyroxenerich garnet could occur isobarically by a univariant transformation from pyrope-rich garnet and omphacite at pressures between 13 and 17 GPa (Gasparik, 1996). In a pyrolite mantle, the dissolution of pyroxene in garnet occurs over a wide range of pressures below 17 GPa with pyroxene fully dissolved at 17 GPa. If the deeper part of the upper mantle is composed of eclogite, discontinuity in seismic velocities at 400 km depth is expected to be sharp. The stability field of garnet expands with increasing pressure from 13 to 17 GPa by the dissolution of pyroxene in garnet, while the field contracts between 17 and 25 GPa by the exsolution of CaSiO3 perovskite and MgSiO3 ilmenite. The compositions of the first

Silicate Melts and Rocks

817

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Figure 11.7. Pressure-temperature path of in situ eclogite facies rocks from Jenner, Fransiscan Complex. Modified after Oh and Liou (1990) with albite breakdown curve from Newton and Smith (1967) (Source Mysen et al., 1998a, 9 1998 Mineralogical Society of America).

garnet at 13 GPa and the garnet co-existing with CaSiO3 perovskite and MgSiO3 ilmenite at 22 GPa are coincidentally identical and close to Ens1Di9Jd40. Compositional gaps of variable widths could exist between the ternary and more pyrope-rich garnets, between the Na-poor and Na-rich garnets and between the Ca-poor and Ca-rich garnets. (Note: The stability field of garnet is seen to expand substantially by addition of even a small amount of Na.) Such immiscibility could produce a sharp discontinuity in an eclogite mantle at 400 km depth. However, a sharp 500 km discontinuity, corresponding to the breakdown of the diopside omphacite to garnet and CaSiO3-perovskite, can only be present in a Ca-rich mantle (Gasparik, 1996). Partial fusion experiments with basic granulites were performed at 0.5-1.5 GPa by Springer and Seck (1997). The melt compositions were seen to change as: trondhjemitic ---, tonalitic ---, choritic, with an increasing degree of partial melting. At 0.5 GPa, the crystalline residua with plagioclase and pyroxene are dominant. At 1.5 GPa, garnet/pyroxene dominate. Melts from granulites match the major element compositions of natural trondhjemites and tonalites. At 0.5 GPa, their A1203 content is relatively low, similar to tonalites; at 1.5 GPa, A1203 is high due to the near absence of plagioclase in the crystalline residua.

11.6.4.1. Dehydration melting of metabasalt at 0.8-3.2 GPa Partial melting experiments on amphibolites representing metamorphosed Archean tholeiite (greenstone), high-alumina basalt, low-K tholeiite and alkali-basalt were carried out by Rapp and Watson (1995). Silicic to intermediate liquids result from --~20-40% melting between 1,050 and 1,100~ leaving a granulite residue at 0.8 GPa and garnet granulite to eclogite residues at 1.12- 3.2 GPa.

818

Chapter 11

The experimental data suggest that the Archean TTG (tonalite-trondhjemitegranodiorite) suite of rocks can be generated by 10-40% melting of partially dehydrated metabasalt at P > the garnet phase boundary (>- 1.2 GPa) and T of 1,000-1,100~ 11.6.5. MORB

Because MORB has high Si, A1, Fe and Na content, the minerals developed at mantle pressures are substantially complicated in nature compared with those obtained from transformed mantle peridotite, viz. perovskite, CaSiO3-perovskite and Mg-wtistite (Kesson et al., 1998). The high A1203 content of MORB also results in higher majorite-perovskite transition pressure than required for peridotite transition. The zeropressure density of basaltic crust with perovskitic lithology is 4.23 g/cm 3. Hirose et al. (1999) calculated the zero-pressure densities at 24 and 26 GPa as 3.87 and 3.92 cm 3, respectively. These values are consistent with those obtained much earlier by Irifune and Ringwood (1993). Up to 27 GPa, the sub-solidus phase relations are determined by Hirose et al. (1999). At 24 GPa (/2,023 K), the mineral assemblage is composed of majorite + stishovite + CaSiO3-perovskite. Along with this, an aluminous phase with Ca-ferrite structure is noted. This observation is consistent with that recorded in earlier experiments by Irifune and Ringwood (1993). At 26 GPa (/> 2,473 K) a new A1-Ca phase is found. This is akin to the CAS phase described by Irifune et al. (1994). A majorite-perovskite transformation is seen to start at this pressure of 26 GPa ( ~ 720 km depth) and 2,000 K (Hirose et al., 1999). The lithology changes from garnetite to perovskitite ---26 GPa. Under the 660 km discontinuity, if the slabs of basaltic lithosphere accumulate to form a megalith of 60 km thickness, its transformation to denser perovskite lithology would cause it to penetrate deep inside the mantle. This transition boundary has a positive P - T slope, whereas the transition boundary in the underlying harzburgite mass has a negative P - T slope (Irifune and Ringwood, 1987). However, no other major phase transformations have been reported at higher pressures up to 100 GPa (Kesson and Ringwood, 1994). This suggests that these phases remain stable in the deep mantle except for majorite, which is fully transformed to perovskite at P > 27 GPa. At this pressure (---4,000 K), the partial melt generated should have a compositional enrichment in MgO, FeO (with depletion in SiO2 + A1203). Thus, the solid residue in MORB composition at lower mantle pressure would become denser because of its higher iron content. At a depth of 1,500 km ( ~ 6 4 GPa), the melting temperature of basalt is about 250 K lower than that of mantle peridotite (Zerr et al., 1998). Extrapolation to 135 GPa yields a melting temperature of MORB ---4,000 K, i.e., at the CMB. The melting curve of MORB extrapolated to the CMB is presented in Fig. 11.8. Thus, if the temperature locally reaches 4,000 K in the D" region, which may be a graveyard for subducted lithosphere, the crustal material of the basaltic component would partially melt. This melt at the base of the mantle can account for the recent observations of the seismic anisotropy (Kendall and Silver, 1996) and anomalously slow P-wave velocities

819

Silicate Melts and Rocks 5 OO0

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Pressure (GPa) Figure 11.8. Melting curve of MORB extrapolated to the core-mantle boundary using the melting relationships of Simon (S) and Kraut and Kennedy (KK). Open and solid circles represent melting temperatures of MORB and MgSiO3, respectively (determined in a laser-heated diamond cell). Open squares represent melting temperatures of MORB (determined in the multi-anvil apparatus). Melting curves of St, Ca-pv and Mg-pv are substantially higher than that of MORB (from Hirose et al., 1999).

(Williams et al., 1996; Revenaugh and Mayer, 1997). However, under such a scenario, the temperature of the outer core must be higher than the 4,000 K required for the melting of MORB perovskite (Fig. 11.8; Hirose et al., 1999). The temperature difference over the thermal boundary between core and mantle may reach 1,500 K and hot mantle plumes, including partially molten slab materials, are likely to arise from this depth (Hirose et al., 1999).

11.6.6. Komatiite, picrite and lherzolite: CaO-MgO(FeO)-SiO2 systems The geochemistry of komatiites holds considerable clues to the scale of mass transfer in the Earth (cf. Herzberg, 1995). High-pressure liquid-phase diagrams obtained from multi-anvil experiments on komatiite and peridotite (e.g., Herzberg et al., 1990; Herzberg and Zhang, 1996) in the 5 - 2 5 GPa pressure range offered important information. The C a O - M g O - F e O - F e 2 0 3 - Fe~ (CMFS and C M F A S _ Fe ~ system represents komatiitic and picritic rock compositions. Komatiites record multiple saturation at --- 3 - 1 0 GPa in the upper mantle (Herzberg and Zhang, 1996). The 3,500-Myr Barberton komatiite has a range of CaO:AleO3 that indicates multiple saturation pressures of 8-11 (+2) GPa.

820

Chapter 11

Herzberg and Zhang (1997) reported the results of multi-anvil melting experiments on a wide range of komatiite analogue mixed with compositions in the system C a O - M g O - F e O - F e 2 0 3 + Fe~ at 5 GPa. The liquidus crystallization fields for olivine (ol), orthopyroxene (opx), clinopyroxene (cpx) and garnet were mapped out, as were their intersections at various cotectic and invariant points. The effect of FeO is to expand the liquidus crystallization phases: pyroxenes at the expense of olivine and clinopyroxene at the expense of orthopyroxene (see also TrCnnes et al., 1992). Herzberg and Zhang (1998) experimented on a range of compositions in the system CaO-MgO-A1203-SiO2. The liquidus crystallization fields for forsterite, orthopyroxene, clinopyroxene and garnet have been mapped out at 10 GPa, as have been at their intersections at various cotectics. It is observed that pressure reduces the content of A1203 and increases MgO and SiO2 contents in magmas formed by the melting of garnet lherzolite with increasing pressure (Herzberg, 1992). The viscous flow of alkali-alumino-silicate melts is constrained by oxygen exchange between polymeric units (Poe et al., 1992). In this flow, structural relaxation may occur through the formation of transient five-coordinated (Si or A1) species. Progressive melting must take place because in the ascending plume the adiabatic gradient has a smaller dT/dP than the solidus. The chemical composition of the upper mantle is considered to be lherzolitic. High-pressure experimentation on garnet lherzolite (liquid + olivine + orthopyroxene + clinopyroxe + garnet) showed that, during melting on decompression, all clinopyroxene and garnet can be dissolved (e.g., Herzberg, 1995). The liquids that erupt solidify to picrites and komatites. Experiments in both CaO-MgO-A1203-SiO2 and MgO-SiO2 demonstrate that there is a maximum normative olivine content to liquids formed by the initial or advanced melting of peridotite in the upper mantle and this occurs at 7 - 8 GPa. During the ascent of a peridotite in a plume, clinopyroxene and garnet are the early phases to melt out during decompression. Advanced anhydrous melting will yield liquids with a residual harzburgite mineralogy (L + ol + opx). Komatiites and picrites of Cretaceous age are thought to have melted in a plume that gave rise to the Carribean plateau (Storey et al., 1991). The parental komatiite composition was dominated by harzburgite (L + opx + ol) (e.g., Herzberg, 1995). A residual harzburgite signature for most komatites with Cretaceous and late-Archean ages has been interpreted to have been formed by about 25-60% anhydrous melting of mantle peridotites and in plumes with potential temperatures that were 200-400~ higher than those of present-day oceanic ridges (e.g., Walter, 1998; Herzberg and O'Hara, unpublished work).

11.6.7. Garnet peridotites: "forbidden zone" The P - T diagram shows the positions for garnet peridotite xenoliths from Lesotho kimberlites (Carswell and Gibb, 1980). The estimated P and T for Lesotho garnet peridotite xenoliths range from 870 to 1,450~ and from 26 to 56 kbar, respectively. A small portion of the terranes in SW China is made up of garnet peridotites, studies on

Silicate Melts and Rocks

821

which have led to some exciting findings. The garnet peridotites of the Dabie-Sulu terrane of east central China may have crystallized in the 'forbidden zone' at depths of 185250 km. These peridotites originate in the mantle and their mineralogical characteristics imply that they have experienced recrystallization at high temperatures (> 1,000~ In some cases, their crystallization ages are considerably older than their host country rocks (Krogh and Carswell, 1995). Amongst all Eurasian garnet peridotites, P - T conditions are seen to be highest in Sulu-Debie UHP terrane, indicating processing conditions of 750-950 ~ and 4.0-6.7 GPa. Some of these P - T conditions lie within the forbidden zone (Liou and Zhang, 1998). 11.6.7.1. Exsolutions Garnets in peridotites from the Western Gneiss Region contain pyroxene needles, indicating that exsolution (unmixing) has occurred after initial formation of a supersilicic garnet at depths > 185 km (e.g., Terry et al., 1999). The Alpe Arami garnet peridotite of northern Italy contains exsolution lamellae of clinoenstatite within diopside, with crystallographic evidence implying initial formation of depths > 250 km (Bozhilov et al., 1999). In Chinese UHP rocks, the peridotites contain unusual FeTiO3 rods observed in olivine (Debrzhinetskaya et al., 1996).

11.6. 7.2. Emplacement of garnet peridotites The mechanism for emplacement of deep, mantle-bome garnet peridotite at the Earth's surface may involve transportation by asthenopheric upwelling from depths of about 135 km, tectonically inserted into subducted continental crust and then exhumed. These forbidden-zone garnet peridotites of Dabie-Sulu may have been placed in the crust before subduction and subjected to in situ UHP metamorphism together with the subducted slab. These bodies offer the first evidence of continental rocks being subducted to depths of 200 km or more. The discovery of garnet peridotites from the forbidden zone now provides a revolutionary new window into the subduction of continental margins, the thermal structure of the subduction zones and the recycling of volatiles into the mantle (Liou et al., 2000). Beneath northeast Japan, a 130-million-year-old Pacific Plate is foundering at the rate of 9.0 cm/year (Peacock and Wang, 1999). This could be a locale for the crust to be subducted down to 200 km depth without being heated beyond 900~ If the subduction to 2 km depth takes --~5 m.y., the rocks in the middle zone of the subducted slab will undergo little heating as the diffusion distance for this time scale is 10-15 km. These cold subduction zones are clearly the sites of major recycling of H20 into the mantle.