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Origin and Evolution of Earth: Research Questions for a Changing Planet
2
Earth’s Interior
As planets age they slowly evolve as the heat trapped and generated in the interior is transported to the surface. The internal planetary processes that move this heat—including volcanism and convection—have a huge influence on the nature of planetary surfaces. Yet the vast interior is inaccessible to direct study and must be understood with geophysical observations, experimental studies of materials under deep-Earth conditions, and theoretical models. For over a century, seismic wave, geomagnetic, and gravity measurements made at the surface have been improving our characterization of Earth’s internal structure. Experimental and theoretical determinations of material properties at high temperatures and pressures and numerical modeling of mantle and core heat transport and convection over very long timescales also play key roles in studies of internal dynamics. However, despite continuing advances, we still cannot uniquely describe Earth’s mantle structure or explain in any detail how the core and mantle work, why Earth differs from other planets, or how it may change in the future.
The three questions included in this chapter describe scientific challenges for understanding Earth’s evolution and internal dynamics. Question 4 addresses deep-Earth dynamical processes, from the inner metallic core at the center of Earth to the convecting mantle to the volcanoes at the surface. Question 5 focuses on the near-surface features of Earth—old continents, young ocean basins, and plate tectonics—that make Earth unique among Solar System planets and that also seem inextricably linked to the presence of water and the preservation of life-sustaining conditions. Question 6 deals with Earth materials properties, which control many of the internal processes discussed in this chapter.
QUESTION 4:
HOW DOES EARTH’S INTERIOR WORK, AND HOW DOES IT AFFECT THE SURFACE?
The previous chapter discussed evidence that Earth and the Moon, and by extension the other terrestrial planets, started out with high internal temperatures about 4.5 billion years ago. Once the planetary accretion process tails off, the planets cool, first through a period of active geological processes and ultimately to a state of geological quiescence. When the planet is geologically active, evidence of that activity is reflected in the nature of its surface and atmosphere and perhaps the existence of a magnetic field. After the interior cools and its viscosity increases sufficiently, geological activity grinds to a halt, and the planet’s surface stops regenerating. Thereafter, only external processes, such as bombardment with asteroids, further modify the surface.
Some planetary bodies, like the Moon, cooled quickly and have been geologically inactive for billions of years. Despite rapid cooling after the Moon-forming impact (Questions 1 and 2), Earth produced and retained enough heat to power geological activity until the present, and it is likely to do so for several billion more years. However, both the amount of Earth’s cooling and resulting changes in the internal dynamics and surface environment are still poorly
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Origin and Evolution of Earth: Research Questions for a Changing Planet
known. Although we know that heat is transported by mantle convection, we do not yet have the capability to exactly describe these convective patterns, calculate with confidence how different they were in the past, or predict how they will change in the future. Resolving the critical questions about planetary evolution will require much more advanced knowledge of planetary materials and how they affect convection (Question 6), better constraints from seismology on the present configuration of mantle flow at both large and small scales, and significant advances in mathematical modeling of convection that is driven by both temperature and chemical variations.
Convection and Heat Flow
About 43 TW (1012 J/s) of heat flows from Earth’s interior through its surface at present, based on global heat flow measurements and thermal models for cooling oceanic lithosphere. Sources of this surface heat flow include the slow cooling of the mantle and core over the history of the planet; heating produced by radioactive decay of U, Th, and K; and minor sources such as tidal heating. The exact contribution of each to the planet’s heat flow is uncertain. For example, we do not know how much U, Th, and K are contained inside Earth and how these elements are distributed (McDonough, 2007). These elements are more effective at keeping Earth hot if they are located deep within the mantle, or even to some degree in the core, rather than near the surface. As a result of these uncertainties, we cannot yet answer the simple question: How fast is Earth cooling?
The primary mechanism for transporting heat within Earth’s interior is convection. It was once believed that mantle convection was impossible because the mantle was demonstrably solid. But much like a glacier, the mantle can behave like both a brittle solid and a liquid: it fractures when deformed rapidly but flows on long timescales. We now know that both the mantle and the outer core circulate in a complex pattern of large- and small-scale flows. In the molten outer core, which has very low viscosity (some estimates suggest a value similar to that of liquid mercury), convection is rapid. Hot liquid metal circulates up to the top of the core where it loses heat to the base of the mantle and then sinks again in a turbulent pattern that is affected by rotation and the magnetic field the flow generates. By contrast, mantle motions are ponderous. Typical velocities are about 5 cm/yr (based on geodetic, magnetic, seismic, and geological measurements), and at this rate the nominal “round-trip” journey of a mantle wide convection cell—across the surface for 5,000 km, down 2,900 km to the bottom of the mantle, and back to the surface again—would take about 300 million years. This rate of travel is consistent with simple thermal convection models that treat the mantle as if it were a liquid with a viscosity (estimated from postglacial rebound rates) of about 1021 Pa-s. The configuration of convection in Earth’s mantle provides the primary control on how Earth cools, mainly because the mantle makes up roughly two-thirds of Earth’s mass and 85 percent of its volume (Figure 2.1).
Mantle motions carry hot material from deep inside Earth toward the surface, where heat is lost to the atmosphere and ultimately to space, and also carry cold surface rocks down to great depths. Unresolved issues concerning mantle convection arise from uncertainties about material properties at high pressures and temperatures. Experiments and field evidence show that mantle rock becomes soft enough to flow over geological time periods at depths of just 30 to 60 km, where the temperature surpasses 700°C and pressure reaches several thousand atmospheres. At higher temperature—above 1200°C—the viscosity of mantle rock is low enough that it behaves much like a thick liquid; almost all of the mantle is hotter than 1200°C. Mantle viscosity exerts the primary control on the form of convection and the efficiency at which heat is moved toward Earth’s surface. However, other factors also are important. For example, viscous dissipation associated with deformation of stiff lithospheric plates at subduction zones strongly affects the form of convection and the relationship between convective vigor and surface heat flow. The largest uncertainties are for the lower mantle. Seismological data suggest that the flow pattern there is complex. Other observations suggest that viscosity increases in the lower mantle, and numerical models indicate that flow velocities in the lower mantle may be much slower than plate velocities such that the overturn time is a billion years or more (Kellogg et al., 1999; Ren et al., 2007).
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Origin and Evolution of Earth: Research Questions for a Changing Planet
FIGURE 2.1 Cutaway view of Earth’s interior showing major layers (oceanic and continental crust, upper mantle, lower mantle, outer core, inner core) and features (mantle plumes, subduction zones, midocean ridges, convection currents, magnetic field). SOURCE: Lamb and Sington (1998). Copyright 1998 Princeton University Press. Reprinted with permission.
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Origin and Evolution of Earth: Research Questions for a Changing Planet
How Are Mantle Convection and Earth’s Thermal Evolution Related?
We know that mantle convection is driven by the heat of Earth’s interior, but what controls Earth’s temperature? The current understanding is that the mantle itself acts as Earth’s primary “temperature regulator,” and its actions depend on the atomic-scale properties of mantle minerals that determine viscosity. The effective viscosity of the mantle depends on the rate at which the mineral grains can deform in response to an applied stress, which in turn is strongly dependent on temperature. Laboratory data indicate that for a given stress a 100°C temperature increase lowers the viscosity by about a factor of 10. Consequently, if Earth were to heat up, it would convect more vigorously and lose heat faster. As heat is lost, temperature drops and convection slows, decreasing the rate of heat loss. This temperature-viscosity feedback should keep Earth’s internal temperature well regulated. The temperature at which the thermostat is most likely to be set is just below the melting point of mantle rock because there is an even faster decrease in viscosity with temperature once the mantle begins to melt.
The temperature-viscosity feedback model is useful but it implies a steady system that undergoes only slow changes over long periods of time. This implication is at odds with much of what we know and suspect about mantle materials and geological history. For example, the continents, which are an end product of Earth’s evolution, show evidence of rapid growth spurts (Question 5), which may or may not be associated with accelerated plate tectonics (Hoffman and Bowring, 1984). The seafloor of the western Pacific Ocean contains enormous volcanic mountain ranges, which suggests that the Cretaceous Period (65 to 150 Ma) was a time of exceptionally intense volcanic activity and possibly also fast seafloor spreading (Engebretson et al., 1992). We also know that the Cretaceous was a period of exceptional global warmth and high sea level (Question 7) and stability of Earth’s magnetic field. These observations as well as theoretical considerations raise the question of whether Earth’s thermal evolution and internal processes are adequately described by our (quasi-) steady state models or whether the evolution has been unsteady and punctuated by catastrophic reconfigurations. Thus, even though we understand the most basic features of mantle convection, our level of understanding is insufficient to explain many of the most important geological and geochemical features of our planet.
We are even further from understanding the internal evolution of other rocky bodies of our Solar System, where we have fewer data, and interactions between thermal evolution and orbital evolution provide additional complications (see Box 2.1). Earth (and possibly Venus) has apparently maintained a high enough internal temperature to ensure continued geological activity. However, on smaller planetary bodies, geological surface activity has either long since stopped (Moon) or slowed greatly (Mars). It is believed that the mantles of other terrestrial planets should function in the same way as Earth’s, unless there are different amounts of radioactive elements or different amounts of water dissolved in the mantle minerals. The addition of tiny amounts of water to mantle minerals would lower both the viscosity of the mantle and the melting temperature (Question 6) and may prolong a planet’s geologically active life.
What Do Mantle Plumes Tell Us About Convection and Heat Transport?
The viscosity of Earth’s mantle is sufficiently low and sensitive to temperature that convection can include complex small-scale currents. Evidence of this small-scale convection is provided by hot spots—large clusters of volcanoes, the most active of which are in Hawaii, Iceland, the Galapagos Islands, Yellowstone, and Réunion (Indian Ocean). Hot spots are usually explained as the surface outpourings of magma formed in mantle plumes, which are cylindrical upwellings of hot (and hence low viscosity) rock that are thought to form near the base of the mantle and rise to the surface at rates much faster than plate velocities (Figure 2.2). Mantle plumes should form as a consequence of heat entering the bottom of the mantle from the much hotter outer core.
Mantle plumes may also be responsible for large igneous provinces, which are vast basalt lava plateaus on continents and the ocean floor. The best current explanation is that they form when the bulbous top of a new plume approaches Earth’s surface (Figure 2.2), then spreads out and causes widespread melting (Ernst et al., 2005). These large, rapid lava outpourings may have caused major perturbations to Earth’s climate (Question 7)
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BOX 2.1
Planetary Comparisons
Our understanding of our planetary neighbors has advanced substantially over the last several decades through spacecraft exploration and analysis of lunar samples and meteorites from Mars and the Moon. The other terrestrial planetary bodies (Venus, Mars, Mercury, and the Moon) formed by the same processes as Earth (see Question 1) and are governed by the same physical and chemical laws and processes. Nevertheless, each has taken a distinct evolutionary track, deepening the questions we pose for how Earth works the way it does.
Venus, at 0.8 Earth masses, is sometimes called Earth’s “sister planet,” but its massive carbon dioxide atmosphere (90-bar surface pressure) and global cloud cover have led to a runaway greenhouse, a surface temperature of 470°C, and the loss of most water from the atmosphere. Venus also lacks Earth-like plate tectonics, but the planet has been subjected to resurfacing—probably by some form of lithospheric recycling not understood—at least once and perhaps multiple times. The density of impact craters indicates that the surface has an average age between several hundred million years and 1 billion years. There are mountain belts and pervasively deformed plateaus, both of which are stratigraphically older than the widespread volcanic plains, known to be basaltic from spacecraft lander measurements. Unlike Earth, Venus has no detectable internal magnetic field. A strong correlation of long-wavelength gravity and topography in the plains is the signature of ongoing mantle convection. Rifting and volcanism have occurred more recently than the average surface age, and the planet is likely to be volcanically and tectonically active at present.
Mars, at 0.1 Earth masses, evolved more rapidly than Earth or Venus. Isotopic evidence from Martian meteorites indicates that Mars formed its core, mantle, and most of its crust within a few tens of millions of years after the beginning of Solar System formation, probably without any plate tectonics era. Large segments of the most ancient crust on Mars are strongly magnetized, relics of a core dynamo that began early in Martian history but probably died out after several hundred million years. The Martian surface has seen a mix of plains volcanism and more focused magmatism in regional centers, dominated by the Tharsis volcanic province, largely constructed before 4 Ga (billion years ago). Fluvial landforms, widespread chemical alteration, and sedimentary deposits visited by surface rovers all indicate that water was an important agent of geological change early in Martian history. At about 4 Ga, Mars lost its global magnetic field, its carbon dioxide atmosphere was substantially thinned by solar wind stripping, the climate cooled, and water lost its dominant role in surface change. Martian volcanism continued at generally declining rates, and the planet may still be active at low levels today.
The Moon and Mercury, at 0.01 and 0.05 Earth masses, respectively, have heavily cratered surfaces and only extremely tenuous atmospheres, but their similarities end there. The Moon began largely molten, presumably the result of the accumulation of hot ejecta from a giant impact on the early Earth. Cooling and solidification of the resulting magma ocean led to formation of the crust and the mantle source regions of later volcanic lavas. Those lavas erupted to partially fill the lunar maria, mostly on the lunar nearside at 3 to 4 Ga, but there are also isolated younger volcanic deposits. The Moon may have a small iron-rich core, but if so it is no more than a few percent by mass. Lunar rocks from 3 to 4 Ga are magnetized, but whether the magnetizing field was a central core dynamo or transient field generated during surface impacts is an open question. The Moon is seismically active at low levels today. Shallow moonquakes are probably the signature of interior cooling, whereas deep moonquakes occur in clusters and appear to be triggered by tidal stresses.
Mercury, in contrast, has such a high bulk density that its iron-rich core comprises at least 60 percent of the planet’s mass. Mercury has a global magnetic field, dipolar like that of Earth, and the outer core is known to be molten on the basis of the amplitude of the planet’s libration forced by solar torques as Mercury progresses along its elliptical orbit. The planet has an ancient, heavily cratered crust, as well as somewhat younger plains units that may be volcanic in origin. The surface composition is poorly known, but Earth-based measurements indicate that surface silicate materials have little or no ferrous iron. The dominant tectonic landforms on Mercury are high-relief lobate scarps, the surface expressions of large-offset thrust faults. Because of the extensive distribution and apparently random orientation of these features, the lobate scarps have been interpreted to record an extended period of global contraction, the result of some combination of interior cooling and solidification of an inner core.
and perhaps even major extinctions (Question 8). Other indications of plumes include broad bulges in the ocean floor, such as those around Hawaii, and the tremendously excessive amount of lava produced at Iceland in comparison to other places along the Mid-Atlantic ridge.
Although there is good geological evidence that mantle plumes exist, seismological evidence for the existence of narrow, hot, cylindrical upwellings in the lower mantle is only equivocal. Some cylindrical regions of low velocity appear to extend downward to 200 to 600 km, while others seem to extend almost to the core-mantle boundary. However, there is abundant evidence for much larger, domical or irregularly shaped low-velocity features in the lower mantle that are sometimes called superplumes (Figure 2.3). Does this mean that thermal plumes do not exist in the lower mantle or that the seismic resolution is still too low to make them out? Seismic data suggest that the large low-seismic-velocity regions near the base of the mantle are anomalously dense, which is contrary to expectations for buoyant thermal upwellings (Ishii and
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FIGURE 2.2 Sketch of mantle convection and structure based on inferences from fluid mechanics and seismological data. SOURCE: Courtesy of Geoff Davies, Australian National University. After Davies (1999). Copyright 1999 by Elsevier Science and Technology Journals. Reproduced with permission.
Tromp, 1999). However, it is becoming better appreciated that temperature variations may not be the only source of buoyant upwellings in the mantle. Chemical variations may be large enough to affect large-scale mantle flow, and mantle plumes can have both thermal and chemical components to their buoyancy (Davaille, 1999; Farnetani and Samuel, 2003).
Does Convection Occur Through the Whole Mantle or in Layers?
A key question about the modern form of mantle flow is whether convection occurs through the whole mantle
FIGURE 2.3 Representation of large-scale seismic velocity structure of the mantle. Red zones have relatively slow P-wave velocity and blue zones are relatively fast. Slow velocities are thought to represent hotter parts of the mantle. SOURCE: <http://www.seismology.harvard.edu/Projects.html>. See also Su et al. (1994). Used with permission.
or in layers. Models, geochemical analyses of mantle magmas erupted on the surface, and interpretations of seismic waves that have passed through Earth have all yielded different answers. In general, mantle models based on geochemistry suggest that mantle convection occurs in two layers, whereas most geophysical evidence and numerical models strongly support whole-mantle convection. Reconciling these differences is important for understanding Earth’s volcanic and thermal evolution.
Geochemical analysis. Interactions of the mantle with the core and surface chemically alter the upper and lower boundary regions of the mantle (discussed below). Convection then stirs this altered material back into the main volume of the mantle. The chemical composition of lavas derived from the mantle provides clues about the extent to which these heterogeneities persist in time and hence about the nature of mantle convection (Van Keken et al., 2002). Lavas (and most other rocks) contain every one of the 90 naturally occurring elements in the periodic table, although about 75 are present in small abundances. With new techniques the concentration of each of the 90 elements and the relative amounts of isotopes of about half of the elements can be measured precisely. The isotopes formed by radioactive decay (206Pb, 207Pb, 208Pb, 87Sr, 143Nd, 230Th, 226Ra, and others) provide detailed information about mantle evolution as well as the processes that produce and transport magma.
Low-abundance trace elements and isotopes of Pb, Sr, Nd, Hf, He, and Os show large, nonrandom variations among volcanic rocks. Basalt lavas erupted at midocean ridges differ systematically from those
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erupted at hot spots. Midocean ridge lavas also vary from ridge to ridge and along individual ridges (Figure 2.4). The chemical differences between hot spots and midocean ridges have long been considered evidence that the lower mantle (whence mantle plumes presumably come) is different, and convects separately, from the upper mantle (Hofmann, 1997).
Nevertheless, there are complications in the isotopic data. For example, 3He data suggest that parts of the mantle are relatively unaltered, or at least less degassed, but other isotopes (Nd, Sr, Pb, Hf) tell a different story (Moreira et al., 2001). The mantle overall does not seem to have an Nd or Hf isotopic composition that properly complements that of the continental crust. Many such chemical clues must be sorted out before we can develop a model for mantle convection and the mantle-crust system that agrees with models for Earth’s bulk composition derived from meteorites (Question 1) and with the distribution of heterogeneities at depth.
Seismic interpretation. The most direct observations available for inferring the present-day configuration of mantle convection are provided by seismicity in subduction zones and three-dimensional seismic tomography models of the interior. Seismic velocity variations are caused by changes in pressure, temperature, composition, and mineral alignment, so interpretation of the models requires information from mineral physics (Question 6) and geodynamics. High-seismic-velocity features corresponding to cold sinking oceanic lithosphere are clearly observed in regional and global seismic tomography models (Figure 2.5). Low-velocity features (presumably signifying
FIGURE 2.4 (Left) Bathymetry of the Mid-Atlantic ridge and topography of adjacent continents. SOURCE: <http://www.ngdc.noaa.gov/mgg/image/2minrelief.html>. (Right) Variations of neodymium isotopic composition in basalt lavas from along the Mid-Atlantic ridge plotted against latitude. Zero on the epsilon scale corresponds to the bulk Earth value, which assumes Earth has the same Sm/ Nd ratio as chondritic meteorites. The high degree of heterogeneity indicates that diverse materials are generated in the mantle by melting and subduction and that these heterogeneities are not homogenized by convection. SOURCE: Data from the online database PetDB, averaged by ridge segment by Su (2002).
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FIGURE 2.5 Seismic tomography data indicating that in some areas subducted slabs extend through the 660-km discontinuity and well down into the lower mantle. Blue shading indicates higher seismic body wave (P) and shear wave (S) velocity, both of which should correlate with lower temperature. The thickness of the cold slab, however, is only about 50 to 100 km, whereas the thickness of the high-velocity (blue) zone is close to 500 km in the lower mantle. The greater thickness in the lower mantle could be due to deformation of the slab or to decreased spatial resolution of the image at greater depth. SOURCE: After Trampert and van der Hilst (2005). Copyright 2005 American Geophysical Union. Reproduced with permission.
relatively high temperature) underlie ocean ridges, back arc basins, and tectonically active areas of continents. Continental cratonic areas are underlain by high-seismic-velocity regions extending 250 to 350 km deep, indicating fundamental differences between oceanic and continental plates (Question 5). Deeper seismic-velocity structures are less easily related to surface tectonics, with very large scale structures tending to dominate in the transition zone from 410 to 660 km deep, and in the lowermost mantle above the core-mantle boundary. For several decades the resolution of seismic tomography models has been improving, and this is guiding numerical modeling of mantle flow processes.
Seismic evidence shows a large velocity discontinuity 660 km below the surface, which is thought to involve mineral phase transformations that tend to impede flow through the transition depth. However, seismological data also show some subducted slabs extending to depths greater than 1,000 km (Figure 2.5), clearly penetrating the 660-km boundary. The variable depth of lithospheric slab subduction is not easily understood in the context of simple thermal convection and is the primary observation driving consideration of more complex convection models involving both thermal and chemical effects.
Models. Mantle convection models have progressed from simplified two-dimensional models to complex three-dimensional simulations, in concert with increasing computing power and improving knowledge of mantle material properties (Figure 2.6; see also Cohen, 2005). Comparison with seismological models allows some parameters in convection models to be tested, but many issues are still unresolved. Among the challenges of simulating mantle convection are the strong dependence of viscosity on temperature and composition, mineralogical heterogeneity in the mantle on both large and small scales, departures from simple fluid behavior,
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FIGURE 2.6 Computer simulation of mantle convection in two dimensions. Red-green-blue color scale depicts temperatures from 4000°C to 0°C. Fine-scale features, which arise from extreme variations in material properties at small length scales, are not well represented in this simulation, but hot upwellings from the core-mantle boundary region, and cold downwellings (analogous to subduction) from the cold surface boundary layer, are prominent features. SOURCE: Butler et al. (2005). Copyright 2005 by Elsevier Science and Technology Journals. Reproduced with permission.
and the effects of melting and phase changes on material properties. Although the simulations are guided by observations and experimental measurements, these are often indirect or subject to varying interpretations as a result of the difficulty in specifying material properties at conditions of high pressure and temperature. For example, the uppermost mantle is mostly made of three minerals: olivine, orthopyroxene, and clinopyroxene. In the lower mantle these minerals are transformed by pressure into higher density forms, and the size and even the composition of the mineral grains are poorly known. It is the deformation characteristics of these mineral aggregates that determine the nature of mantle convection. Because the grain size and other properties of the deep mantle have yet to be determined, our ideas about convection in the lower mantle involve large extrapolations of the properties we can determine for Earth materials at lower pressure and temperature (see Question 6).
Numerical simulations of mantle convection show that even with phase transitions inhibiting flow and a viscosity increase in the lower mantle, it is plausible that large-scale transport of material between the upper and lower mantle does occur. All in all, current seismological and geodynamic results tend to favor an intermediate model of mantle convection that is neither strictly layered nor simple whole-mantle convection.
When Did Earth’s Inner Core Form?
Earth’s thermal evolution is reflected in and strongly influenced by the temperature of the liquid outer core. The fact that Earth’s outer core is liquid rather than solid is evidence for the hot origin of Earth, and the fact that the core has not completely solidified over Earth’s 4.5-billion-year history means that it has been prevented from losing heat too quickly. Laboratory experiments suggest that the top of the core is about 1500°C hotter than the deep mantle (Figure 2.7). Therefore, heat must be flowing from the outer core into the lower mantle, and the core must be cooling. The core must also be close to its solidification temperature because the inner core is solid. As the core cools, it solidifies from the bottom up, so we deduce that the solid inner core is growing and the liquid outer core is shrinking.
The inner core–outer core boundary must have a temperature exactly equal to the melting temperature of
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FIGURE 2.7 Schematic representation of average temperature in Earth’s interior versus depth. Viscosity estimates are also shown. Temperature is highly uncertain below about 500-km depth. The average mantle temperature (red line) is based on an adiabatic gradient and a temperature of 1350°C at a pressure of 1 bar. Higher and lower temperatures for plumes and subducted slabs are approximate but close to estimates. Temperature in the core and the large temperature drop at the base of the mantle are poorly constrained. See van der Hilst et al. (2007) for a recent estimate of temperature at the core-mantle boundary and Bunge (2005) for a discussion of nonadiabatic temperature structure in the mantle.
the core at the corresponding pressure. The core melting temperature is uncertain because the core contains minor elements other than Fe, and it is not known exactly which elements and how much of them. Hence, the melting temperature of the core is likely to be a complex function of both composition and material properties at high temperature and pressure. Ongoing research is examining the possibility that heat-producing elements (e.g., potassium) may be present in the core and may contribute to a slowing of core cooling.
How long the inner core has existed, its rate of growth, and why the core has not fully solidified are fundamental unresolved issues (Butler et al., 2005). Part of the answer seems to be that the core has been kept in a molten state by the mantle, which because of its much higher viscosity does not remove heat fast enough. Also, once crystallization of the inner core started, it would slow cooling of the core because crystallization releases heat. It has recently been inferred from convection models that the inner core may be relatively young; it may have begun forming about 1.5 billion to 2 billion years ago (Labrosse et al., 2001). This idea, however, is inconsistent with theoretical models that suggest the presence of a solid inner core may be important for the strength of the magnetic dipole field and for the occurrence of reversals. Moreover, there is evidence that Earth’s magnetic field is older than 2 billion years (Tarduno et al., 2007). This apparent conundrum may be partly a consequence of our still poor understanding of the characteristics and processes near the core-mantle boundary, including the values of the temperature contrast and the amount of heat flowing across the boundary (e.g., Bunge, 2005).
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How Has Earth’s Magnetic Field Evolved Through Time?
It has long been recognized that the main part of the geomagnetic field is sustained by fluid motions in Earth’s electrically conducting outer core. These motions cause the magnetic field to change over many timescales, from diurnal to annual to geological timescales. However, a unified picture of how the geodynamo and the core fit into the Earth system has not yet emerged. Important questions about the internal operation of the geodynamo and the relationship between the geodynamo and other Earth processes remain unanswered. These include: How do the geodynamo, mantle convection, and plate tectonics interact? What role did the geodynamo play in Earth’s early history? The age of the magnetic field is of interest because the magnetosphere may help keep Earth habitable. For example, the magnetosphere may have been necessary to help Earth retain its atmosphere against the eroding powers of the solar wind, and it partly shields Earth’s surface against radiation from space. How important the latter is in preserving life or in modulating the rates of evolution is not agreed upon. New insights on these questions will come from continued satellite and ground-based observations of the geomagnetic field and the paleomagnetic field, dynamical interpretations of the core’s seismic structure, and sophisticated numerical dynamos (e.g., Figure 2.8) and models of core evolution.
What Are the Chemical Consequences of Mantle Convection?
The mantle interacts with Earth’s surface environment through volcanism, heat and mass exchange at midocean ridges, and subduction. The mantle may also exchange material with Earth’s outer core. Overall, the mantle mediates a grand-scale circulation of materials that may extend from the core-mantle boundary to the surface and back again. The nature of this mantle circulation and the processes that produce the interactions at the mantle boundaries are critical to understanding how Earth’s chemistry is continually modified. For example, volcanism builds oceanic and continental crust (Question 5) and releases to the atmosphere water,
FIGURE 2.8 A snapshot of a three-dimensional computer simulation of the geodynamo. The magnetic field is illustrated using lines of force; blue lines represent the inward directed field and yellow lines represent the outward directed field. The field is intense and complicated in the model’s fluid iron core, where it is generated by fluid motions. Like Earth’s field, the simulated field has a dominantly dipolar structure outside the core. SOURCE: Courtesy of Gary Glatzmaier, University of California, Santa Cruz; and Paul Roberts, University of California, Los Angeles.
carbon dioxide, and other gases, continually renewing the oceans and atmosphere. Mountain building, erosion, and subduction, which also reflect the effects of mantle convection, remove these same materials and tend to recycle them into the deepest parts of the mantle. At the core-mantle boundary we infer there is mainly heat exchange, but there is tantalizing evidence of chemical interaction as well (Brandon et al., 1999). Still unknown are whether the processes that mediate these exchanges were different in the past. An interesting possibility is that the nature of continents and oceans that support a habitable surface environment today reflect only a particular phase of Earth’s cooling and hence might have been absent or much different in the past and might also be much different in the future.
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quantitatively predict the dependence of erosion rate on other variables and the strength and deformation properties of rocks in the lower continental crust and the upper mantle.
QUESTION 6:
HOW ARE EARTH PROCESSES CONTROLLED BY MATERIAL PROPERTIES?
Geology is founded on the central insight that rocks can be read as a record of Earth’s history. Rocks and minerals are produced and altered by geological processes—melting, eruption, weathering, erosion, deformation, and metamorphism. Therefore, deciphering the secrets of the rock record begins with an understanding of large-scale geological processes. The keys to understanding these processes are the basic physics and chemistry of the materials that make up the planet. Scientists now recognize that macroscale behaviors—plate tectonics, volcanism, and so on—arise from the microscale composition of Earth materials and indeed from the smallest details of their atomic structures. Understanding materials at this microscale is essential for comprehending Earth’s history (NRC, 1987) and making reasonable predictions about how things may change in the future.
The high pressures and temperatures of Earth’s interior, the enormous size of Earth and its structures, the long expanse of geological time, and the vast diversity of materials and properties present challenges to investigation. Moreover, minerals are complicated solids that generally contain not only their essential chemical constituents but also trace amounts of almost every element known in nature. Although we can learn much about Earth from the study of pure compounds that approximate real minerals, we also know that even minute amounts of other chemical elements can radically change a mineral’s behavior.
Fortunately, the surge of interest in understanding Earth materials at the atomic level has been accompanied by rapid development of new tools, including new synchrotron sources that bring the ability to probe the atomic structure of minerals and liquids (Figure 2.16); high-pressure devices to simulate the distortion of atomic arrangements under huge pressures; and advanced quantum mechanical theory, which promises major advances in our understanding of physics and chemistry at the extreme conditions of planetary interiors and at the smallest scales of mineral surfaces and nanoparticles. Advances at the other end of the spectrum, when the scale is extremely large and/or the processes are extremely slow, will require advances in experiment, theory, computation, and observation. Only the combination of all four is likely to bring progress.
What Minerals Comprise Planetary Interiors?
As noted in Questions 4 and 5, the nature of the convection and deformation that affect Earth’s mantle and crust, and hence models for plate tectonics and Earth’s temperature history, depends directly on the material properties of rocks and minerals at the high temperatures and pressures of planetary interiors. The pressure is 136 GPa (1.36 million atmospheres) at the base of the mantle and 364 GPa at Earth’s center, while the temperature reaches 4000 K at the base of the mantle and 6000 K at Earth’s center (similar to the temperature at the surface of the Sun; Figure 2.17).
Phase transformations. The pressure in Earth’s interior is so enormous that it alters the fundamental properties of elements; for example, it can convert insulators to metals and cause magnetism to collapse (Figure 2.18). Such changes occur because pressure compresses and distorts the electron orbitals, thereby changing the most basic properties of the materials. Changing pressures bring about many kinds of phase transformations. The most familiar of these are melting and freezing, but many more complex phase transformations have been identified. Structural phase transitions are also common. The transition from graphite to diamond is well known, but more important for Earth processes is how mantle olivine and pyroxenes change at high pressure.
High-pressure mineral transformations, and their dependence on temperature, allow us to estimate the temperature of the deep Earth and provide constraints on how mantle convection works. Temperatures inside Earth can be estimated by comparing the pressure and temperature conditions at which mineral transformations occur in the laboratory to the depths at which sudden changes in the physical properties of the mantle and core occur (Figure 2.19). We know, for example, that the boundary between the liquid outer core and the
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FIGURE 2.16 (Top) Aerial view of the storage ring at the Advanced Photon source. Such third-generation synchrotron sources have revolutionized the study of Earth materials by dramatically increasing spatial and temporal resolution of experimental measurements and allowing for the study of much smaller samples than had been possible. A similar qualitative advance is expected when the first fourth-generation synchrotron sources (X-ray-free electron lasers) come online in 2009. SOURCE: <www.aps.anl.gov/About/APS_Overview/index.html>. Courtesy of Argonne National Laboratory, managed and operated by the University of Chicago, Argonne, LLC, for the U.S. Department of Energy. (Bottom) Results of a quantum mechanical computation based on density function theory, showing the predicted structure and distribution of electrons in SiO2 at high pressure. Such computational methods can provide estimates of material properties over the vast range of pressures and temperatures encountered in planetary interiors. SOURCE: Oganov et al. (2005). Reprinted with permission. Copyright 2005 by the American Physical Society.
solid inner core must be at the melting temperature of the core (Question 4), although the temperature is not known precisely due to uncertainty in the composition of the core and the difficulties of exploring these high temperatures in the laboratory. The temperature of the most important changes of seismic wave velocity in the mantle, which happen at depths of about 400 and 660 km, is well constrained by laboratory studies of the conversion of olivine and pyroxene to higher density minerals. These phase transformations are so drastic that they can influence mantle convection; a phase transformation that causes a large change in density can work either for or against the thermal buoyancy that drives convection.
Although the effects of phase transitions on mantle convection are generally appreciated, we still do not know how the natural system actually works—for example, the extent to which the phase transitions impede or enhance the sinking of subducted slabs or change the size and shape of mantle plumes as they rise. A previously unknown phase transformation was recently discovered at pressures well beyond those previously probed (Murakami et al., 2004). The new transformation, from perovskite, the main mineral structure of the deep mantle, to a higher pressure postperovskite form, occurs at the top of the D″ region, an anomalous zone above the core-mantle boundary (corresponding to some 100-GPa pressure) that exhibits intriguing and highly variable seismological features (see Question 4), some of which may be caused by the transformations.
What is the melting temperature of rocks under pressure? Much of what we know about how Earth’s interior works is based on knowledge of the melting temperature of rock and metal, and how this temperature changes with pressure (Question 4). To expand this knowledge,
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FIGURE 2.17 Diamond-anvil apparatus (top). The sample is placed between two opposed diamond anvils, the tips of which range from 0.01 to 1 mm across, depending on the pressure range of interest. The vertically oriented strip is a metal gasket that prevents the sample from extruding. Diamond is ideal for high-pressure studies because is it strong, chemically inert, and transparent to most light. SOURCE: <http://www.physics.missouristate.edu/Faculty/Mayanovic/mayanovic_research_webpage.htm>. Used with permission. A shock wave experiment can be carried out using a gun (right), magnetic drive, or laser. The projectile can produce pressures and temperatures that exceed those at Earth’s center (like a diamond anvil cell) but for very short periods of time (in contrast to static anvil experiments). New methods combine both static and dynamic approaches to reach pressure-temperature domains (Jeanloz et al., 2007). SOURCE: <www.gps.caltech.edu/~sue/TJA_LindhurstLabWebsite/index.html>. Used with permission.
FIGURE 2.18 Influence of pressure on the iron atom. Shown is the predicted charge density of the doubly charged iron cation (Fe) in the mineral ferropericlase (Mg,Fe)O, in which it is surrounded by six oxygens (O). (Left) At low pressure the spins of the d electrons are maximally aligned, producing a net magnetic moment on each iron atom (called the high-spin or HS state) and the magnetic properties that we are familiar with, such as the tendency of magnetic minerals to align with the magnetic north pole. (Right) At high pressures characteristic of Earth’s deep mantle the spins pair (called the low-spin or LS state), the atomic magnetic moments vanish, and iron-bearing minerals are nonmagnetic. The figures show that the size and shape of the iron cation also change across the high-spin to low-spin transition: iron is smaller (by about 10 percent in volume) and less spherical in the low-spin state, which should produce a change in density and other physical properties of iron-bearing minerals. SOURCE: Tsuchiya et al. (2006). Reprinted with permission. Copyright 2006 by the American Physical Society.
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FIGURE 2.19 Photograph looking into a diamond cell at a 100-micron blue single crystal of hydrous ringwoodite (ideally Mg2SiO4 composition) held in situ at 30 GPa, corresponding to a depth of 800 km in Earth. The brown spots indicate where the sample has been heated with a laser to a few thousand degrees, causing a phase transformation to the assemblage MgSiO3 perovskite + MgO periclase that is thought to comprise most of Earth’s mantle below a depth of 660 km. SOURCE: Courtesy of Steven Jacobsen, Northwestern University. Used with permission.
we need to understand the processes that control the melting and freezing of rocks and minerals in the planetary interior. Melting of rocks involves complex chemistry, because rocks are typically composed of four or more mineral phases, none of which are pure. As rock melts, the composition and density of the liquid portion are different from those of the solid, and thus, with the help of gravity, one can segregate from the other. For example, the lava that erupts from volcanoes is both less dense and compositionally different from the parent mantle rock. Over Earth’s long history the repeated processes of melting, melt ascent due to buoyancy, and eruption onto the surface have completely rearranged many of its chemical elements. This process of planetary differentiation, making chemically distinct domains out of a homogeneous starting material, is one of the most fundamental features of planetary evolution (Questions 2, 4, and 5).
One of the more intriguing questions about melting is whether, under some conditions, magma may be denser than the surrounding solid mantle. Magma is highly compressible, so its density must increase rapidly with increasing pressure. The density of solids also increases with pressure but more slowly. Although there is so far only scant experimental and theoretical evidence, it suggests that magma can be denser than mantle rock deep inside Earth (Figure 2.20; Miller et al., 1991). The consequences of this for Earth’s evolution would be profound. If silicate melt sinks instead of rising toward the surface, it could be stored at depth for long periods, where it would be kept hot. The geochemical consequences of this inverted gravitational separation could also be important, but little is known about the distribution of trace elements between solids and liquids at high pressures. Iron-rich liquid would likely exist as a separate, denser phase than Earth’s silicate fraction and sink to the center, forming the core (Question 2). The timescale of this descent and the partitioning of elements between the iron-rich and silicate portions during core formation are still uncertain and have profound implications for the chemical composition of the core and the origin of the geomagnetic field (Question 4).
There is confirming evidence that liquid may be present in the deep mantle, especially near the coremantle boundary. Seismologists have identified thin layers of extremely low shear wave velocity at the base of the mantle, a characteristic of liquid. It has been suggested that this region could be made of dense, partially solidified magma and that it could even be a remnant of the Hadean planetary magma ocean (Williams and Garnero, 1996; see Question 2). If U, Th, and K are concentrated in this deep liquid, it could mean that the base of the mantle produces extra heat from radioactivity, which would affect how we think about the core dynamo and about the overall chemical composition of the mantle. If mantle liquid is in contact with the liquid outer core, it would also mean that chemical exchange across the boundary would be much more effective than if the mantle is solid; this would change the way we think about the origin of chemical heterogeneity in the mantle (Question 4). To resolve these issues we need to know much more about the properties of silicate liquids and solids at very high pressures and temperatures. Recent experimental advances, including measurements of liquid structure in situ at high pressure (Shen et al., 2004), will work hand in hand with theoretical and computer modeling. Modeling of high-pressure properties (Figure 2.20), using the principles of quantum mechanics, shows promise, although at present only a
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FIGURE 2.20 Predicted atomic-scale structure of a model magma (MgSiO3 composition) showing that the large compressibility of liquids is caused by rearrangement of the structure from an open configuration near zero pressure (left) to a much more compact and highly coordinated structure at the pressure of the core-mantle boundary (right). Silicon-oxygen coordination polyhedra are shown in blue and magnesium ions in yellow. SOURCE: Stixrude and Karki (2005). Reprinted with permission from the American Association for the Advancement of Science (AAAS).
small number of atoms can be modeled, which means that it is not yet possible to use this approach to explore how trace elements behave.
Can seismic waves be used to uniquely determine mantle mineralogy? Material properties and seismology are interdependent in a fundamental way. Seismologists can measure the speed at which seismic waves traverse the mantle and use this information to construct pictures of the deep Earth in a process analogous to a medical CAT scan. At the same time, pictures of the deep mantle cannot be interpreted without information about mantle minerals and rocks, just as radiologists need to know how bone and other types of tissue transmit X-rays. The changes in seismic wave velocity through different structures in the deep Earth are small—about 1 percent—so the elastic properties of the minerals need to be known precisely to interpret the changes. As these properties become better known, geologists hope to use seismic images to map the temperature and composition variations in the mantle and perhaps even the pattern of convection. The latter is possible because seismic wave velocity is dependent on direction, or anisotropy, and can be related to flow patterns if there is sufficient knowledge of the elasticity of minerals and the mechanisms by which they deform (Karato, 1998). A striking example of anisotropy inside Earth may be seen in the inner core, where longitudinal seismic waves travel 3 percent faster along the rotational axis than in the equatorial plane. This difference may be due to alignment of iron crystals in the core, although the mechanism for producing the alignment is still uncertain (Stixrude and Brown, 1998). Understanding the origin of this alignment is likely to tell us a great deal about the dynamics at Earth’s center, the history of the core, and the origin of the geomagnetic field.
How Much Water Is in the Solid Earth?
Earth is unique in the Solar System for its abundant surface water, and most models for the early Earth suggest that the source of this water was the mantle via volcanic eruptions. Based on recent research, it seems likely that the interior continues to be a major reservoir of both water and carbon dioxide (Williams and Hemley, 2001). Earth is so massive that if the mantle is only 0.03 percent water, it would hold the equivalent of all the water in the modern oceans. Upwelling mantle material at midocean ridges appears to contain about this much water, so at present Earth’s interior has at least one ocean’s worth of water. How much more it might have and how this amount has changed over Earth’s history are outstanding questions.
We do not know whether Earth has always had the present amount of water at its surface, but the answer has implications for a variety of processes. To reach the answer, we need a deeper understanding of where water and carbon dioxide are stored in the mantle. We know of two potential reservoirs of water: hydrous phases, such as clays that contain predictable amounts of water within their crystal structures, and nominally anhydrous phases, such as olivine (the most abundant mineral in the upper mantle), which include hydrogen as defects (Figure 2.21). Knowing more about these reservoirs may frame our view of the long-term evolution of the hydrosphere, including formation of the oceans (Question 2). Understanding the evolution of the deep hydrosphere is also central to our view of mantle dynamics, since even small amounts of hydrogen can change the viscosity of the mantle by orders of magnitude and the melting temperature of rocks by hundreds of degrees (Question 4). For example, if the mantle has more water, it might convect faster and produce more volcanism, by which it loses water to the
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FIGURE 2.21 An example of how water might be stored in Earth’s interior. Shown is the predicted structure of a nominally anhydrous mantle mineral (stishovite, ideally SiO2) with trace amounts of hydrogen incorporated via the replacement: Si4+ = Al3+ + H+. Dark- and light-blue polyhedra are SiO6 and AlO6 coordination environments, respectively; red spheres are oxygen atoms; and the green sphere is a hydrogen atom. The solubility of water in this mineral reaches a few percent at conditions typical of the shallow lower mantle. SOURCE: Courtesy of Lars Stixrude, University of Michigan.
surface. If the mantle loses too much water, volcanism might slow down until enough water is returned to the mantle by subduction. This type of feedback may help regulate Earth’s surface environment and the water content of the mantle (see also Question 7).
How Do Minerals and Fluids React?
Chemical reactions between minerals and water enable the oceans and atmosphere to exchange chemicals with the rocks of the crust and mantle. These chemical reactions control the mineral weathering that accompanies erosion and ultimately affect the composition of seawater, the bioavailability of nutrients and toxins in the environment, and the amount of carbon dioxide in the atmosphere (Question 7). All of this chemistry occurs in the microenvironment at the surfaces of minerals. New data about natural materials, especially about the microstructure of mineral surfaces, are changing ideas about how minerals and fluids react. Recent studies have shown that reactivity is exquisitely sensitive to the finest details of surface structure. For example, the rate of exchange with water for oxygen atoms on distinct but structurally similar sites on an aluminum hydroxide surface may vary by seven orders of magnitude (Phillips et al., 2000).
In addition, a major new realization is that most of the mineral surface area in the environment may be in the form of nanophases: extremely small mineral particles, 1 to 100 nm in size, orders of magnitude too small to see with the naked eye. These very small mineral grains have dramatically different physical and chemical properties than larger ones (Banfield and Zhang, 2001). The surface energy of nanophases is so important that it can stabilize structures that do not exist in bulk material (Navrotsky, 2004). These structures may have unique reactive sites, adsorptive properties, and reaction kinetics. The structures of nanophases also vary depending on whether they are surrounded by water, air, or organic ligands. Nanophases are important for their role as a unique reactive surface area, and they also help us understand how minerals form, since all minerals start out as nanophases in the form of small nucleation centers (Figure 2.22).
At and near Earth’s surface, the formation and dissolution of minerals take place in the presence of microorganisms, and there is a growing awareness that biology plays a significant role in mediating chemical reactions at mineral surfaces (Question 8). In addition, many minerals are formed entirely by living organisms, both large and small. Limestone, for example, is almost entirely formed as calcium carbonate shell material by small marine organisms. Much of the modern study of mineral formation lies at the interface of biology, chemistry, and geology. With new analytical techniques it is becoming possible to study how minerals are made by organisms and to compare biological and inorganic processes. For example, it is possible that an organism can produce a microenvironment that causes calcite to be precipitated essentially by inorganic processes. By altering the microenvironment, the organism can control the particular form, and hence trace element composition, of the mineral that is precipitated (Bentov and Erez, 2006). We may have much to learn about how minerals form by carefully watching how organisms make them (Figures 2.22 and 2.23).
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FIGURE 2.22 Necklace of titania nanocrystals that have aggregated spontaneously by oriented attachment. In this mineral growth pathway, crystals no more than a few nanometers in diameter aggregate and rotate so that adjacent surfaces share the same crystallographic orientation. The pair of adjacent interfaces is eliminated and the pair of nanoparticles is converted to a larger single crystal. Individual atoms are visible in the lower view. SOURCE: (Top) Penn and Banfield (1999). Copyright 1999 by Elsevier Science and Technology Journals. Reproduced with permission. (Bottom) Courtesy of Lee Penn, University of Minnesota, and Jillian Banfield, University of California, Berkeley. Used with permission.
Can Large-Domain, Multiscale, and Extremely Slow Earth Processes Be Predicted?
Many properties and processes depend on length scale and timescale in ways that are difficult to predict. The general idea of scaling, or inferring the behavior of materials at one scale from knowledge of those materials at another scale, underlies much of our thinking about Earth. For example, our understanding of mantle convection is founded on our ability to relate planet-scale (large) and laboratory-scale (small) flows that have the same ratio of buoyancy forces to viscous resisting forces (the Rayleigh number). Laboratory analogs are likely to be accurate for some aspects of mantle convection, but they have limits. For example, we know that the crust and uppermost mantle exhibit nonfluid behavior, or there would be no plate tectonics (Question 5). We also know that most of the surface deformation caused by plate tectonics takes place in narrow zones at the edges of the plates. The localization of deformation probably has an origin in complex failure processes that are dependent on both size and timescale. Rocks and even magmas can exhibit a behavior called strain softening, which means that as the amount or rate of deformation increases, the resistance to deformation decreases, which increases the amount and rate of deformation further. Consequently, deformation is most likely to continue wherever it has already started and to be concentrated in narrow zones rather than being widely distributed. Other feedbacks of this sort include thermal weakening and damage weakening (Bercovici and Karato, 2002). In the latter, deformation either reduces grain size or increases crack density, making the material easier to deform. There are many ways that rocks can behave when stressed; these different deformation processes affect one another; and the larger the rock body under consideration, the more processes that can come into play. Hence, predicting what will happen at a large scale from information about what happens at a small scale is a major challenge.
The behavior of faults raises many scale-related fundamental questions: How are earthquakes (large scale) generated and can we predict them using small-scale models (Question 9)? What localized (small-scale) process and set of conditions trigger a (large) fault to rupture a particular distance on a particular day? How much of continental deformation (large) is caused by slip on faults (small)? Some of the most influential predictors of fault movement have been laboratory measurements of rock strength: squeeze a rock in one direction and eventually it will break or slide along preexisting faults, once friction is overcome. However, rock at the scale of a great earthquake rupture is much weaker than rock in the laboratory. One possible explanation is that water is pervasive in the crust and weakens fault planes by acting as an easily sheared but
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FIGURE 2.23 Orange, polymer-laden ferric iron oxyhydroxides from a submerged mine. The slime consists of colloidal aggregates of nanoparticles, mineralized cell products, and cells (left) of two bacteria. The twisted stalks are characteristic of iron-oxidizing bacteria belonging to the Gallionella genus, while sheathed elongate cells are typical of bacteria belonging to the iron-oxidizing Lepthothrix genus. The contrast is due to iron oxyhydroxide nanoparticles. (Right) A closeup of the nanoparticle aggregates reveals that while the individual particles are separated (white regions), they have been bio-assembled so that they are crystallographically oriented in the same direction. SOURCE: Banfield et al. (2000). Reprinted with permission from AAAS.
incompressible lubricant that dramatically reduces the friction between the two rock surfaces (Figure 2.24). But as noted above, there are many other possible ways to cause Earth’s crust to appear weak in comparison to rocks in the laboratory.
Another reason scaling is challenging is that Earth is heterogeneous: material properties, including viscosity, electrical and thermal conductivity, chemical diffusivity, and elasticity, may vary spatially by orders of magnitude on scales ranging from nanometers to kilometers. Heterogeneity may dramatically influence dynamics. Cappuccino drinkers are familiar with the fluid dynamical oddities of composites, seen in the relative stiffness of milk foam as compared with its constituents, air and milk. Analogous phenomena are common in nature. For example, as magma forms by melting inside Earth, it juxtaposes relatively fluid magma with mineral crystals that are essentially rigid. The viscosity of crystal mush, which largely determines how fast it rises (or sinks), depends strongly and nonlinearly on the amount of suspended solid crystals it contains. Deformation and/or dissolution of the solid matrix through which magma moves can also organize solid and liquid fractions so that the liquid becomes channelized, dramatically increasing the rate of liquid-solid segregation, with important implications for magma migration in the mantle and formation of the core (Questions 2 and 4; Holtzman et al., 2003). The mantle is made of solid minerals with varying strengths. Just as in the case of magma channelization, mantle convection may organize weaker and stronger minerals into layers (foliation), dramatically influencing the viscosity as well as the seismic signal and our interpretation of it in terms of composition, temperature, and flow pattern. Chemical reaction of fluids and melts with surrounding solids can also produce channels, which can significantly influence the composition of the magma and our inferences about its origin (Spiegelman and Kelemen, 2003; Figure 2.25).
The importance of time. The solid-like or fluid-like behavior of the mantle illustrates the importance of time in the material properties of large domains. The boundary between fluid-like and solid-like behavior is set by the Maxwell relaxation time—the ratio of viscosity to shear modulus—which is on the order of 1,000 years for the mantle. This means that we can only determine the viscosity of mantle materials in the laboratory at extremely slow rates of deformation or at unrealistically high temperatures to bring the Maxwell relaxation time within the window achievable by experiment. Just as solids behave like fluids on long
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FIGURE 2.24 Photograph (upper left) and thin section (upper right) of the Punch Bowl fault in southern California. The principal slip surface (pss) is thought to have accommodated several kilometers of slip. The slip is localized to a 1-mm (white) region, including a microshear zone with more intense shearing (dark) occurring within a few hundred microns. SOURCE: (Upper left) Chester and Chester (1998). Copyright 1998 Elsevier, reprinted with permission. (Upper right) Courtesy of Judith Chester, Texas A&M University. (Bottom) Results of experiments on fault slip in natural rocks showing that the friction coefficient depends on slip velocity and nearly vanishes for slip velocities similar to those of earthquakes (1 m/s). SOURCE: Di Toro et al. (2004). Reprinted by permission from Macmillan Publishers Ltd.: Nature, copyright 2004.
timescales, fluids behave like solids and rupture on short timescales. When magma is deformed very rapidly—for example, during an eruption—it may fracture. Understanding this behavior is helping us sort out the dynamics of volcanic eruptions (Question 9) and how these depend on features such as magma composition (e.g., Gonnermann and Manga, 2003).
Summary
Understanding how Earth works depends on knowledge of the properties of rocks and minerals. After a period of steady progress, breakthroughs are now at hand because of new analytical tools provided by advanced radiation sources (e.g., synchrotron, neutron, and laser facilities) and advanced computing. Much of
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FIGURE 2.25 Simulation of the distribution of melt (as measured by porosity) in a deforming, reacting matrix. The melt organizes itself into channels that vary in width, position, and melt content with time. SOURCE: Courtesy of Marc Spiegelman, Columbia University. Used with permission. See also Spiegelman et al. (2001).
the essential physics and chemistry of Earth materials arises from structures and processes that occur at the atomic level. The new tools allow these small scales to be studied directly as well as simulated, bridging the gap between quantum mechanics and microscopy and paving the way for a new level of understanding of planetary processes at longer length scales.
Earth materials present a challenge to understanding because of their complex chemical composition and the high pressures and temperatures of planetary interiors. The long timescales of geological processes also create difficulties because some of the critical processes that affect planetary evolution take place so slowly that they cannot be simulated in the laboratory and because they may be caused by mechanisms that are not important or even perceptible at laboratory timescales. The physics of large domains, long timescales, and multiple interacting scales remains a major challenge in Earth science and one that will advance only with interdisciplinary effort.
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