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45th Lunar and Planetary Science Conference (2014) 2652.pdf PLATE BOUNDARY LOCALIZATION: WHAT PROCESSES ACTIVE ON EARTH DO NO APPLY TO OTHER PLANETARY LITHOSPHERES? L.G.J. Montesi1, 1University of Maryland, Department of Geology, College Park MD 20742, [email protected]. Introduction: Few problems in Earth Science are as fundamental as the origin of plate tectonics. Our planet has adopted a global tectonic regime that appears unique in the solar system [1]. Deformation is localized in narrow deformation zones, some of which do no contribute to removing heat from the interior of the planet [2]. Because of the higher strain rate inside narrow deformation zones, energy dissipation is higher in a localized shear zone than in a distributed one. Therefore, plate boundary localization must involve a weakening process to make deformation in a localized more favored over distributed deformation [3]. Here, I summarize localization processes proposed to be active on Earth, their effect, and whether or not they may be active in other terrestrial bodies. Localization processes: Rocks deform following brittle processes at relatively shallow pressure and temperature and plastic mechanisms are greater depth. Geological observations on Earth show that deformation can localize under both conditions [4], even though plastic rheologies, being strain-rate hardening, are fundamentally stable in the laboratory [3, 5]. Localization can be understood in either case if a state variable enables rock strength to weaken as strain accumulates. Brittle processes involve the formation of, and sliding along, cracks between and inside grains that form a rock [e.g., 6]. It is therefore not surprising that brittle failure should result in a localized shear zone. Additional processes that weaken brittle shear zones include nucleation of weak, often hydrated minerals like serpentine, graphite, and mica [7] and the development of a foliated cataclastic zone [8]. All these phenomena contribute to decrease the coefficient of friction from 0.6 to 0.1. A key to achieve such weakness is that fault gouge adopts a fabric that allows the weak phase to form a through-going layer [9, 10]. Localization in the ductile regime is harder to understand, because of the fundamental strain-rate hardening behavior of plastic deformation process [11]. Several studies have appealed to shear heating to weaken ductile rocks [12, 13], but there only scant evidence of this process in the field [14,15]. It is most efficient at relatively low temperature [5]. Ductile shear zone rocks often display marked fabric differences compared to the protolith and wall rocks. A reduced grain size is often observed, which leads to the many shear zones being classified as mylonites [16–19]. Reducing grain size is possible only when dislocations are active inside the rocks [e.g. 20]. However, it leads to weakening only if grain boundary sliding and diffusion are dominant. Therefore, grain size reduction is a viable localization process only in the dislocation-accommoated Grain Boundary sliding regime (dis-BGS, [21, 22]). The mantle can localization in this regime at temperatures less than less than 800°C [18]. Ice also deform by disGBS is the superplastic regime [23]. Ductile shear zones also display a layered fabric of anastomosing high-strain bands at many scales [4, 24]. The high-strain zone are often rich in phyllosilcates, sometimes reaching the stage of being a phyllonite [25, 26]. Development of a layered fabric enables the weakest phase to control the rheology of the deforming rock. If that phase is weak or highly nonlinear, like micas, pronounced weakening is associated with the appearance of a layered fabric [5]. In summary, geological observations imply that fabric development is the dominant weakening process that enables shear zone localization. This process is active in both brittle and ductile regimes. Montési [5] shows that the potential for localization associated with this mechanism far exceeds that of other processes. Grain size evolution often accompanies shear zone development but participates to weakening only if disGBS creep is active. This mechanism may be important in relatively low temperature mantle and in ice. Localization in the continental lithosphere. Gueydan et al. [27] have recently proposed a model of strain-dependent lithospheric strength. At low strain, the upper mantle and upper crust are the strongest regions of the lithosphere. However, fabric evolution reduces strength in the upper crust (brittle fabric), middle crust (fabric in presence of mica) and upper mantle (grain size reduction). The reduced strength of the lithosphere as a whole leads to increased strain rate if the loading stress, controlled by large-scale tectonics, does not change. The lower crust, which does not weaken, becomes the load-bearing layer at high strain. This model is able to reconcile different views of the lithospheric strength profile by assocating the classic “jelly sandwich” model to low strain areas and the “crème brûlée” model to high strain areas [28]. 45th Lunar and Planetary Science Conference (2014) Localization on Venus. Applying these concepts to Venus requires taking into consideration the higher temperature of the lithosphere compared to the Earth. Most phyllosilicates are unstable just a few km below the Venusian surface. Moreover, phyllosilicates are often hydrated. The low water content expected in the Venusian lithosphere makes it unlikely that these phases are present. Therefore, ductile layering is unlikely an important player on Venus. DisGBS is limited to ~800°C in the mantle. Assuming a geothermal gradient of 10 K/km, such temperature are present everywhere the crust is less than 34 km, that is, everywhere except under the highest crustal plateaus [29]. Brittle localization is less efficient than on Earth because connection of a hydrated weak phase is unlikely on Venus. Localization on Mars. Mars presents the same lithological possibilities as the Earth. However, the crust is dominantly basaltic, with only limited evidence for felspathic rocks [30]. Many micas in terrestrial shear zones are formed by the breakdown of Feldspar in presence of water [27]. As this is unlikely to take place on Mars, crustal-level localization is probably limited to the brittle field. The Martian geotherm has probably chaged dramatically over time. For a 30 km thick crust (lowlands), the temperature necessary for disGBS would have been exceeded if the geotherm exceeded 30 K/km. These conditions would only have have been exceeded since the early Noachian, making localization in the shallow mantle a possibility for most of the geological history of Mars. Localization on Icy Satellites. Icy shells are essentially monomineralic, which limits localization to grain size reduction in presence of disGBS. This mechanism is thought to be dominante in icy shells [23]. Therefore it is not surprising that icy satellite tectonics display clear evidence for localized deformation. Indeed, it seems difficult for icy shells to retain the rigid interiors that are necessary to define plates [2]. Taking stock: These theoretical considerations reveal that localization in the brittle upper crust and in the upper mantle are expected in every planet. However, the Earth is unique in the possibility of having localization in the middle crust thanks to its unusual mineralogy. The combination of felsic minerals and water makes it possible to form micas, which are associated with a strong potential for localization dues to their highly non-linear rheological behavior. These minerals are likely to be rare in the crust of Mars and Venus because of the absence of source minerals and water, respectively. Therefore, the Venusian and Mar- 2652.pdf tian lithosphere have a structure where deep and shallow localizing levels are probably separated by a thick layer of non-localizing materials. This makes it difficult to achieve wholesale failure of these lithosphere which is necessary for plate tectonics [31]. The lithosphere remains strong overall, favoring a stagnant lid regime. The icy satellites lithosphere, like the terrestrial oceanic crust, can be regarded as a single layer, where localization is possible thanks to grain size reduction and the involvement of disGBS. This behavior favors localization. The icy satellite ice layer essentially shatters because the subjacent water layer opposes very little resistance, unlike the oceanic asthenosphere. References: [1] O’Neill C. et al. (2007) EPSL, 261, 20–32. [2] Bercovici, D. et al. (2000) Geophysical Monograph 121, 1–46. [3] Montési, L.G.J., and Zuber, M.T. (2002) JGR, 102, doi:10.1029/2001JB000465. [4] Vauchez A. et al. (2012) Tectonophysics 558-559, 1-27. [5] Montési, L.G.J. (2013) JSG 50, 254–266. [6] Paterson, M.S., Wong, T.-f., (2005) Experimental rock deformation – the brittle field Springer. [7] Lockner, D.A., et al. (2011) Nature 472, 82-85. [8] Collettini, C. (2009) Nature 462, 907-911. [9] Niemeijer, A. GRL 37, doi:10.1029/2009GL041689. [10] Rutter, E.H. et al. (2013) JSG 51 doi:10.1016/j.jsg.2013.03.008 [11] Paterson, M.S. (2007) Tectonophysics 445, 273–280. [12] Regenauer-Lieb, K., and D. Yuen (1998) GRL 25, 2737– 2740. [13] Kaus, B.J.P., and Podladchikov, Y.Y. (2006) JGR 111, doi:10.1029/2005JB003652. [14] Brun J.-P., and P.R. Cobbold (1980) JSG 2, 149– 158. [15] Camacho A. et al. (2001). JSG 23, 1007– 1013. [16] Drury, M.R.et al. (1991), Pure Appl. Geophys. 137, 439– 460 [17] Jin, D. et al. (1998) JSG 20, 195– 209. [18] Precigout, J. and Gueydan, F. (2009) Geology 37, 147-150 [19] Platt J.P., and Behr, W.M. Behr (2011) JSG 33, 537– 550. [20] Montési, L.G.J., Hirth, G. (2003) EPSL 211, 97–110 [21] Hirth, G., and Kohlstedt, D. (2003) Geophysical Monograph 138, 83–105. [22] Hansen L.N. (2011) JGR 116, doi: 10.1029/2011JB008220. [23] Golsdby, D.L. and Kohlstedt, D.L. (2001) JGR 106, 11017–11030. [24] Gerbi, C.,N. et al. (2010) JSG 32, 107– 117. [25] Jefferies, S.P.,et al. (2006) JSG 28, 220–235. [26] Gueydan, F., et al. (2003) JGR 108, 2064 doi:10.1029/2001JB000611. [27] Gueydan, F. et al (2014) Tectonophysics, under revision [28] Jackson, J., (2002) GSA Today, September, 1-9. [26] Jefferies, S.P.,et al. (2006) JSG 28, 220–235. [29] James, P.B. et al. (2013) JGR 118, doi:10.1029/2012JE004237. [30] Wray, J.J. et al. (2013) Nature Geoscience, 6, 1013– 1017. [31] Regenauer-Lieb K. et al. (2001), Science 294, 578-580