The Proceedings of the Eighth International Conference on Creationism (2018)

their morphology, and their stereology. When a solid (such as rock or metal) is subjected to stress, it first deforms elastically up to a certain point and then it deforms inelastically. Because the deformations that occur in the mantle typically are huge relative to the elastic limit (i.e., boundary between elasticity and inelasticity), we will be concerned primarily with inelastic deformation. In this study, we focus on the plastic deformation among the inelastic mechanisms. The plastic responses of mantle rocks depend in a fundamental way on several subscale heterogeneous structures and their interactions that can be coupled together on different spatial scales. In this section, some important multiscale features and their associated mechanisms that occur during plastic deformation, especially in the earth’s mantle, are briefly explained in order of length scale, from smallest to largest. A. Dislocations When polycrystalline materials like rocks deform plastically under stress, line defects at the atomic scale are generated within grains and at grain boundaries. These defects are known as “dislocations,” because the crystalline lattice is dislocated along a 1D line of atoms. Under certain thermomechanical conditions, these dislocations move in the grains to stabilize the material’s thermodynamic free energy and their motions lead to a reduction in the rock’s strength. The microstructural mechanisms related to these motions are referred to as “recovery.” Generally speaking, the number density of dislocations increases under deformation, and the interactions among dislocations as they move highly affect the material’s strength. In most cases, dislocation interaction increases the material’s strength, and this phenomenon is known as “work hardening.” During plastic deformation the material simultaneously undergoes both hardening due to dislocation interactions and weakening due to the recovery and annihilations of dislocations. The material’s strength is therefore determined by a competition between the two processes of hardening and recovery. Hence, understanding dislocation mobility, their density changes, and their interactions are crucial to correctly representing a material’s mechanical properties. B. Grains Polycrystalline rocks have stable crystal structures, and each individual crystal is known as a “grain.” Each grain in a rock has a specific shape, size, and orientation and shares polygonal boundaries with its neighboring grains. Grain boundaries can act as either a source or sink for dislocations. The diameter of a grain, called its grain size, affects the work hardening rate, because the boundaries act as barriers to inhibit dislocation motion. Hence, as the grain size decreases, the work hardening rate increases because the dislocations have less room to move (Hall 1951, 1954; Petch 1953). This Hall-Petch effect can be succinctly stated that as the grain size increases, the material weakens. Under certain thermomechanical conditions (e.g., pressure, temperature, and dislocation density), the grain boundaries can migrate to find energetically stable states. That is, the grain size changes until it finds a saturated state. This phenomenon is known as “grain boundary migration” or “grain growth.” Generally, under high temperature, the grain size increases as larger grains consume adjacent smaller grains. Under high pressure, grain growth is inhibited, however, because the thermodynamic activation energy barrier for the grain growth is raised. Hence, the resultant grain growth depends upon multiple thermomechanical quantities. In addition to grain growth, grain size reduction (or refinement) can also take place as a consequence of plastic deformation. This occurs when new small grains nucleate inside of the original grains or at the grain boundary. This phenomenon is known as “dynamic recrystallization.” For a geological example, dynamically recrystallized microstructures with very fine grains are commonly observed in highly deformed rocks like mylonite, a metamorphic rock found in zones of high shear such as folds or faults. Many laboratory studies have shown that dynamic recrystallization indeed occurs with dramatic grain size reduction (Hansen et al. 2012; Karato et al. 1980; Ohuchi et al. 2015; Van der Wal et al. 1993; Zhang et al. 2000). C. Crystallographic Preferred Orientation (CPO) or texture As discussed earlier, each grain of a polycrystalline material has its own orientation. However, during plastic deformation, the grains undergo rotations so that the grains display a distinctive average orientation. This distribution of grain orientations is known as “deformation-induced crystallographic preferred orientation (CPO)” (Karato 2012). In geophysical studies, researchers compare CPO observed in laboratory studies with the seismological wave data from the mantle to gain insight into the mantle’s plastic deformation history (Mainprice et al. 2000). A high-fidelity model able to predict the CPOs of mantle minerals for various types of deformation could potentially utilize seismic observations from today’s mantle to gain important insight into the mantle’s deformation that occurred during the Flood cataclysm. D. Mineralogical compositions of the earth’s mantle The earth’s mantle is thought to comprise at least eleven different minerals across its depth (Karato 2012; Ringwood 1991; Stixrude and Lithgow-Bertelloni 2011). As the depth increases, the dominant upper mantle minerals (olivine, pyroxenes, and garnet) undergo phase changes to other minerals with different crystalline structures due to increased pressure and temperature. These physical changes are known as “solid-solid phase transformations.” Most of these phase changes occur in the lower portion of the upper mantle between depths of 410 and 660 km, a region known as the mantle transition zone. The three main transition zone minerals are wadsleyite, ringwoodite, and majorite (Al-deficient garnet). Small amounts of clinopyroxene and Ca-perovskite can appear in the upper and lower portions of the transition zone, respectively (Irifune et al. 2008; Wood and Helffrich 2001). Since the dominant mineral phase in the upper mantle is olivine (approximately 60% in volume fraction), its various phases play significant roles throughout the mantle. At 410 km depth, where the transition zone begins, olivine first transforms to wadsleyite. Near 520 km depth, wadsleyite further transforms to ringwoodite, which has a spinel structure. Near 660 km depth, ringwoodite decomposes into Mg-perovskite and ferropericlase. This phase transition marks the boundary between the transition zone and the lower mantle. Also in the 600-700 km depth interval, Ca-perovskite exsolves from the majoritic garnet, and some portion of the garnet transforms to Mg- perovskite. Consequently, the lower mantle consists approximately of 70% Mg-perovskite, 20% ferropericlase, and 10% Ca- Cho et al. ◀ Strength-reducing mechanisms in mantle rock during the Flood ▶ 2018 ICC 708