Peridotite is widely considered to be the dominant component of the upper most mantle. Seismic velocities of a dry peridotite determined in the laboratory at high pressure and hypersolidus temperature show a rapid decrease with increasing temperature. The strong temperature dependence of velocities can be used to estimate temperature and melt fraction in the low‐velocity zone of the Earth. The laboratory results show that the pressure dependence of both velocity and melt fraction appears to be well accounted for by the pressure dependence of the solidus temperature of the peridotite, i.e., homologous temperature dependence. This observation allows us to extrapolate laboratory results to higher pressures (greater depths) with some confidence, requiring only a knowledge of the solidus as a function of pressure. Using the ratios of lithospheric to asthenospheric velocities, temperature and melt fraction in the asthenosphere (the low‐velocity zone) can be determined from laboratory velocity data. Examples of thermal structures of the upper mantle beneath oceanic plates are presented. We choose three locations of geophysical interest: Iceland Plateau, Pacific Ocean, and Philippine Sea, where reliable seismic velocities have been determined in the upper mantle from surface wave studies. A large amount of partial melt (≥5 vol %) and a higher temperature than the dry peridotite solidus are inferred in 0–5 m.y. asthenosphere under the Iceland Plateau and the active marginal basin of the east Philippine Sea. The partial melt zone appears to extend deeper (>100 km) and to greater ages (>20 m.y.) in the Pacific Ocean region than under the slowly spreading Iceland Plateau. Partial melting is not expected in the asthenosphere older than 5 m.y. under the Iceland Plateau. Even though the volume fraction of partial melt is not so large (≤3 vol %) beneath the Pacific plate, the melting zone may extend to a distance of above 2000 km from the ridge. Many seamounts observed in the Pacific Ocean may result from this vast region of melting under the plate. We suggest a linear relation between the width of melting zone and the plate velocity. The laboratory results are also applied to other low‐velocity zones. It is suggested that a velocity drop up to about 6% in the asthenosphere does not necessarily require the existence of partial melt but can be explained by high subsolidus temperature. This implies that partial melt may not exist generally but only in limited areas in the low‐velocity zone. Comparison of seismic results with theoretical models may overestimate the temperature and melt fraction, if the velocity drop at subsolidus temperature is neglected. Temperature and melt fraction obtained in this study are discussed together with results from heat flow and electrical studies. Almost the same temperatures as inferred from seismic velocity and a dry peridotite solidus are calculated from heat flow data, indicating that the upper mantle under mid‐ocean ridges may be fairly dry. Even a water‐ or CO2‐undersaturated solidus (0.1 wt ...
The anelastic properties of compressional waves in a peridotite have been determined in the laboratory at sufficiently high temperatures (to 1280°C) and pressures (to 0.73 GPa) to warrant comparison with seismic measurements of the Earth. A substantial decrease of Qp is observed at temperatures well below the onset of partial melting. Qp systematically increases with increasing pressure over the entire temperature range. Of major significance is the finding that Qp is dependent on the ratio of the temperatures to the melting (solidus) temperature; i.e., Qp depends on the homologous temperature. The pressure dependence of Qp appears through the pressure dependence of the solidus of the peridotite. Within the uncertainties of measurement of both Qp and the phase diagram, it appears that melting and high‐temperature anelastic properties have a common origin in peridotite. The homologous temperature dependence of Qp suggests that we may estimate the temperature and pressure dependence of Qp for peridotites of different compositions and possibly even for hydrous peridotites, if solidus temperatures are known as a function of pressure (a far easier measurement than elastic and anelastic properties). The activation volume of Qp is greatly reduced at high pressure, since the slope of solidus versus pressure rapidly decreases with increasing pressure. Pressure dependence of seismic velocity and melt fraction in peridotite also appears to be related to the homologous temperature. The Qp‐homologous temperature relation suggests a connection between Qp and the properties of the grain boundaries; that is, the major loss of seismic energy occurs at the grain boundaries. Grain boundary relaxation or high‐temperature background attenuation is a possible mechanism for the grain boundary damping. No frequency dependence of Qp is resolved (0<α<0.2 in Qp ∝ fα) over the pressure, temperature and frequency ranges of the measurement. The present results and the model of grain boundary relaxation suggest that an appropriate choice of grain size may give an ultrasonic Q that is applicable to the Earth. Experimentally determined anelastic properties of a peridotite are critical for modeling mechanical properties of the upper mantle. Implications of the results are as follows: (1) Seismic data commonly interpreted as indicating a partially molten asthenosphere may instead reflect a hot solid asthenosphere at 90–100% of the solidus temperature. (2) Partial melting may not produce any abrupt change of seismic velocity and Q; rather, elastic and anelastic properties of the upper mantle will change gradually at the boundary where the geotherm crosses the solidus. (3) There may be no sharp mechanical boundary between the lithosphere and the asthenosphere.
INTRODUCTIONRecent studies of igneous rocks have taken a refreshing new direction, mainly as the result of a greater awareness of the important role played by physical properties of magmas in determining the eruptive behavior and compositional variations of volcanic rocks. This rcview summarizes the present state of knowledge, along with some of the recent rheological studies having a direct bearing on interpretations of volcanic phenomena and processes of crystallization and differentiation of shallow magmatic intrusions.The rheological properties of magmas reflect the inherent structures of molten silicates and will not be thoroughly understood until more is known about the basic nature of silicate liquids. Recent progress in this field has been ably reviewed by Hess (1980) and needs no elaboration here aside from mention of the obvious need for better laboratory measurements in order to place greater constraints on interpretations of silicate melts. For petro genetic and volcanological processes, the most conspicuous need is for improved data on rheological properties, chiefly viscosity and yield strength, and for more precise equations for predicting these properties under a wide range of natural conditions. Viscosity, 1'/, is usually defined as the ratio of shear stress to strain rate and is expressed in units of poise (g cm -1 S -1) or Pascal seconds (10 poise = 1.0 Pa s). In more general terms, it is the coefficient for transfer of 337 0084--6597/84/0515-0337$02.00 Annu. Rev. Earth Planet. Sci. 1984.12:337-357. Downloaded from www.annualreviews.org Access provided by Northeastern University on 02/05/15. For personal use only. Quick links to online content Further ANNUAL REVIEWS 338 McBIRNEY & MURASE momentum in the equation • = .o+�(��r (1)in which r is the shear stress applied in the direction x parallel to the plane of fl o w, 'Lo is the minimum stress required to initiate permanent deformation, du/dy is the velocity gradient normal to the plane of shear, and n is a constant with a value of one or less. A fluid having no yield stress (.0 = 0) and a direct linear relation between shear stress and rate of strain (i.e. n = 1) is said to be Newtonian. Many complex fluids, suspensions, and emulsions are non-Newtonian in the sense that the relati on between strain rate and stress is nonlinear, and in some, viscous flow takes place only when the shear stress exceeds a finite value. Although, for convenience, magmas are commonly treated as Newtonian fluids, most are not Below their liquidus, particularly when they are charged with phenocrysts, their rate of shear is not directly proportional to stress, i.e. n <
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