Geodynamics

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Geodynamics is a subfield of geophysics dealing with dynamics of the Earth. It applies physics, chemistry and mathematics to the understanding of how mantle convection leads to plate tectonics and geologic phenomena such as seafloor spreading, mountain building, volcanoes, earthquakes, faulting and so on. It also attempts to probe the internal activity by measuring magnetic fields, gravity, and seismic waves, as well as the mineralogy of rocks and their isotopic composition. Methods of geodynamics are also applied to exploration of other planets.[1]

Overview[edit]

Geodynamics is generally concerned with processes that move materials throughout the Earth. In the Earth’s interior, movement happens when rocks (or melts) deform and flow in response to a stress field.[2] This deformation may be brittle, elastic, or plastic, depending on the magnitude of the stress and the material’s physical properties, especially the stress relaxation time scale. Rocks and other geomaterials are structurally and compositionally heterogeneous and are subjected to variable stresses, so it is common to see different types of deformation in close spacial and temporal proximity.[3] When working with geological timescales and lengths, it is convenient to use the continuous medium approximation and equilibrium stress fields to consider the average response to average stress.[4]

Experts in geodynamics commonly use data from geodetic GPS, InSAR, and seismology, along with numerical models, to study the evolution of the Earth's lithosphere, mantle and core.

Work performed by geodynamicists may include:

Deformation of Rocks[edit]

This article is about deformation in geology. For a more rigorous treatment, see Deformation (mechanics).

Rocks and other geological materials experience strain according three distinct modes, elastic, plastic, and brittle depending on the properties of the material and the magnitude of the stress field. Stress is defined as the average force per unit area exerted on each part of the rock. Pressure is the part of stress that changes the volume of a solid; shear stress changes the shape. If there is no shear, the fluid is in hydrostatic equilibrium. Since, over long periods, rocks readily deform under pressure, the Earth is in hydrostatic equilibrium to a good approximation. The pressure on rock depends only on the weight of the rock above, and this depends on gravity and the density of the rock. In a body like the Moon, the density is almost constant, so a pressure profile is readily calculated. In the Earth, the compression of rocks with depth is significant, and an equation of state is needed to calculate changes in density of rock even when it is of uniform composition.[5]

Elastic Deformation[edit]

Elastic deformation is always reversible, which means that if the stress field associated with elastic deformation is removed, the material will return to its previous state. Materials only behave elastically when the relative arrangement of material components (e.g. atoms or crystals) remains unchanged. This means that the magnitude of the stress cannot exceed the yield strength of a material, and the time scale of the stress cannot approach the relaxation time of the material. If stress exceeds the yield strength of a material, bonds begin to break (and reform), which can lead to ductile or brittle deformation.[6]

Ductile deformation[edit]

Ductile or plastic deformation happens when the temperature of a system is high enough so that a significant fraction of the material microstates (figure 1) are unbound, which means that a large fraction of the chemical bonds are in the process of being broken and reformed. During ductile deformation, this process of atomic rearrangement redistributes stress and strain towards equilibrium faster than they can accumulate.[6] Examples include bending of the lithosphere under volcanic islands or sedimentary basins, and bending at oceanic trenches.[5]

Brittle Deformation[edit]

When strain localizes faster than these relaxation processes can redistribute it, brittle deformation occurs. The mechanism for brittle deformation involves a positive feedback between the accumulation of defects in areas of high strain, and the localization of strain along these dislocations and fractures. In other words, any fracture, however small, tends to focus strain at its leading edge, which causes the fracture to extend.[6]

Whichever mode of deformation ultimately occurs is the result of a competition between processes that tend to localize strain, such as fracture propagation, and relaxational processes, such as annealing, that tend to delocalize strain. In general, the mode of deformation is controlled not only by the amount of stress, but also by the distribution of strain and strain associated features.[6]

Deformation structures[edit]

Structural geologists study the results of deformation, using observations of rock, espcially the mode and geometry of deformation to reconstruct the stress field that affected the rock over time. Structural geology is an important complement to geodynamics because it provides the most direct source of data about the movements of the Earth. Different modes of deformation result in distinct geological structures. Figures 3 shows the appearance of brittle fracture in rocks while figure 4 shows the appearance of ductile folding.

Thermodynamics and Geodynamics[edit]

The physical characteristics of rocks that control the rate and mode of strain, such as yield strength or viscosity, depend on the thermodynamic state of the rock and composition. The most important thermodynamic variables in this case are temperature and pressure. Both of these increase with depth, so to a first approximation the mode of deformation can be understood in terms of depth. Within the upper lithosphere, brittle deformation is common because under low pressure rocks have relatively low brittle strength, while at the same time low temperature reduces the likelihood of ductile flow. After the brittle-ductile transition zone, ductile deformation becomes dominant.[2] Elastic deformation happens when the time scale of stress is shorter than the relaxation time for the material. Seismic waves are a common example of this type of deformation. At temperatures high enough to melt rocks, the ductile shear strength approaches zero, which is why shear mode elastic deformation (S-Waves) will not propagate through melts.[7]

The Dynamics of the Earth[edit]

The main motive force behind stress in the Earth is provided by thermal energy from radioisotope decay, friction, and residual heat.[8][9] Cooling at the surface and heat production within the Earth create a metastable thermal gradient from the hot core to the relatively cool lithosphere.[10] This thermal energy is converted into mechanical energy by thermal expansion. Deeper hotter and often have higher thermal expansion and lower density relative to overlying rocks. Conversely, rock that is cooled at the surface can become less buoyant that the rock below it. Eventually this can lead to a Rayleigh-Taylor instability (Figure 2), or interpenetration of rock on different sides of the buoyancy contrast.[2][11]

Figure 2 shows a Rayleigh-Taylor instability in 2D using the Shan-Chen model. The red fluid is initially located in a layer on top of the blue fluid, and is less buoyant than the blue fluid. After some time, a Rayleigh-Taylor instability occurs, and the red fluid penetrates the blue one.

Negative thermal buoyancy of the oceanic plates is the primary cause of subduction and plate tectonics,[12] while positive thermal buoyancy may lead to mantle plumes, which could explain intraplate volcanism.[13] The relative importance of heat production vs. heat loss for buoyant convection throughout the whole Earth remains uncertain and understanding the details of buoyant convection is a key focus of geodynamics.[2]

Methods in Geodynamics[edit]

Geodynamics is a broad field and combines many different types of observations. Close to the surface of the Earth, data includes field observations, geodesy, radiometric dating, petrology, mineralogy, drilling boreholes and remote sensing techniques, such as seismic tomography and magnetotellurics. Unfortunately, for anything deeper than a few kilometers, most direct observations become impractical. Geologists studying the mantle and core must rely almost entirely on seismic tomography, modeling and mineral physics experiments designed to recreate the conditions of the deep Earth. (see also Adams–Williamson equation).

See also[edit]

References[edit]

  1. ^ Ismail-Zadeh & Tackley 2010
  2. ^ a b c d Turcotte, D. L. and G. Schubert (2014). "Geodynamics."
  3. ^ Winters, J. D. (2001). "An introduction to igenous and metamorphic petrology."
  4. ^ Newman, W. I. (2012). "Continuum Mechanics in the Earth Sciences."
  5. ^ a b Turcotte & Schubert 2002
  6. ^ a b c d Karato, Shun-ichiro (2008). "Deformation of Earth Materials: An Introduction to the Rheology of Solid Earth."
  7. ^ Faul, U. H., J. D. F. Gerald and I. Jackson (2004). "Shear wave attenuation and dispersion in melt-bearing olivine
  8. ^ Hager, B. H. and R. W. Clayton (1989). "Constraints on the structure of mantle convection using seismic observations, flow models, and the geoid." Fluid Mechanics of Astrophysics and Geophysics 4.
  9. ^ Stein, C. (1995). "Heat flow of the Earth."
  10. ^ Dziewonski, A. M. and D. L. Anderson (1981). "Preliminary reference Earth model." Physics of the Earth and Planetary Interiors 25(4): 297-356.
  11. ^ Ribe, N. M. (1998). "Spouting and planform selection in the Rayleigh–Taylor instability of miscible viscous fluids." Journal of Fluid Mechanics 377: 27-45.
  12. ^ Conrad, C. P. and C. Lithgow-Bertelloni (2004). "The temporal evolution of plate driving forces: Importance of “slab suction” versus “slab pull” during the Cenozoic." Journal of Geophysical Research 109(B10): 2156-2202.
  13. ^ Bourdon, B., N. M. Ribe, A. Stracke, A. E. Saal and S. P. Turner (2006). "Insights into the dynamics of mantle plumes from uranium-series geochemistry." Nature 444(7): 713-716.

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