Present knowledge about the internal structure is based largely on remote sensing techniques, landers and meteorites (SNC): shergottites, nakhlites, and chassigny (commonly believed as fragments of Martian crust ejected during large impacts – e.g.
Nyquist et al. 1979 , Jagoutz, 1991, Head et al. 2002). On this basis, it has been generally accepted that Mars, perhaps like other planets from the inner Solar System, have a differentiated interior where core, mantle and crust can be distinguished (Fig. 31). The segregation in layers gives a rise for thermally or/and chemically driven convection that eventually might lead to a global magnetic field generation (see 2.4).
2.3.1 Core
The mean density and moment of inertia that can be measured by probes from the orbit suggest large similarities to a scaled Earth model (Stevenson, 2001). Following this assumption, the core has been formed due to a gravitational differentiation of immiscible iron and silicates at the early stage of planet’s accretion. Based on 146Sm, 182Hf, 142Nd, 182W e.g. (Harper et al. 1995), (Lee and Halliday, 1997), (Kleine et al. 2004), (Foley et al. 2005), derived from SNC meteorites, the time for core-mantle separation was very short. The
Fig. 31) Model of Mars’ interior. A) Core, B) Mantle, C) Crust (Source: PSN National Science Center, Malaysia, www.psn.gov.my).
A B C
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presently accepted value is 10 - 12My years after the formation of the Solar System (Blichert-Toft et al. 1999), (Shih et al. 1999), (Kleine et al. 2004), but even a shorter period has been suggested (Yin et al. 2002), (Jacobsen, 2005). Heat needed for an extensive melting and possibly formation of magma ocean e.g. Tanton et al. 2005a), (Elkins-Tanton et al. 2005b), (Médard and Grove, 2006), could have been provided through the kinetic energy of impacting bodies e.g. (Wetherill, 1990), (Agnor et al. 1999). An influence of decaying radioactive elements, like 26Al, seems to be insignificant at this early stage of evolution (Elkins-Tanton et al. 2005a). Recently refined moment of inertia (Sohl et al.
2005) and tidal deformations data (Yoder et al. 2003) allow for the core radius estimation, with considerable precision, between 1520 and 1840km taking 3390km as a mean planetary radius (Stewart et al. 2007). Although, there are indications (Yoder et al. 2003) for a still relatively hot and at least partially molten core, it is still unclear if a solid inner core is developed. Evolutionary models propose several solutions e.g. (Stevenson, 2001), (Fei and Bertka, 2005), (Sohl et al. 2005), (Stewart et al. 2007) that are strongly correlated with core’s chemical composition. The most often invoked constituents in terrestrial planets are iron with small amount of nickel and some light element like sulfur or silicon (Gessmann et al. 2001), (Li and Agee, 2001), (Sohl et al. 2005), (Stewart et al. 2007).
2.3.2 Mantle
The mantle is even less understood. Additionally to a large variation in possible chemical composition (see review in Médard and Grove, 2006), difficulties arise also from the history of mantle formation. Possibly, the most widely accepted chemical composition model (Dreibus and Wänke, 1985), derived from SNC meteorites, seems to confirm a hypothesis of two distinct sources that took a part in terrestrial planets accretion (Ringwood, 1977), (Wänke, 1981). Proposed components are: A) chondrites with reduced chemical components (e.g. enstatite chondrites), B) volatile rich carbonaceous chondrites (Dreibus and Wänke, 1985). Quick, homogenous accretion from such elements would cause partial oxidation of more reduced components thus binding them before segregation could take a place. Fe, Cr and Mn enrichment in SNC meteorites seems to be in accordance with this scenario (Dreibus and Wänke, 1985). Moreover, high volatile content in component B would effectively support partial melting by lowering solidus temperature (Médard and Grove, 2006) that eventually lead to magma ocean formation and homogenization of upper 700-1500km of Mars (Righter et al. 1998), (Elkins-Tanton et al.
2005a), (Elkins-Tanton et al. 2005b). Even though, close analyses of SNC meteorites
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suggest that some parts of the mantle could retained heterogeneous domains (Foley et al. 2005) that were later on magma sources for shergottites, nakhlites, and chassigny. Although present Martian mantle is “dry”, it cannot be excluded that at the early stage of evolution it contained appreciable amount of water (Médard and Grove, 2006) that supports thermal convection by decreasing melt’s viscosity.
Presently, a minor amount of water (several ppm) from homogeneous accretion might be still bound in nominally anhydrous crystal yield similar mineralogical assemblages e.g. (Bertka and Fei, 1997). Upper Martian mantle is thought to be dominated by olivine, (ortho- and clino-) pyroxens and garnet (Fig. 32). At about 9GPa, orthopyroxene become unstable. Close to 13.5GPa, due to high Fe content, olivine transforms partially to ringwoodite marking the beginning of the Martian transition zone (in Earth’s upper mantle olivine transforms to wadsleyite and later to ringwoodite). A small increase in pressure, to 14GPa, causes transformation of olivine and ringwoodite into wadsleyite accompanied by clinopyroxene and high pressure garnet (majorite). From about 15GPa, ringwoodite is formed in expense of wadsleyite and clinopyroxene. The last one is completely consumed at about 17GPa. In Martian condiitons, where pressure increases much slower with depth than in Earth’s mantle, the phase transitions between (Fe,Mg)SiO4
polymorphs are likely to form diffusive boundaries. Close to the core/mantle boundary another transition that marks the beginning of the lower mantle is predicted. If the
Fig. 32) Mineralogical model of the Martian mantle.
Ol - olivine, Sp (γ) – ringwoodite, β-phase – wadsleyite, Opx – orthopyroxene, Cpx – clinopyroxene, Gt - garnet, Maj – majorite, Mw – magnesiowüstite, Mg-Pv – magnesium perovskite (Bertka and Fey, 1997).
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temperature is high enough and if current values for the core radius (Sohl et al. 2005) are close to the reality, an existence of the lower mantle becomes plausible. The assemblage of ringwoodite and partially majorite is replaced by magnesiowüstite, perovskite and majorite.
2.3.3 Crust
Unlike the core and mantle, Martian crust is much easier subject of investigation.
Gravimetric and topographic data (Fig. 33 A) defines the maximum crustal thickness on not greater than 150km (Sohl et al. 2005) with the average value between 30 and 80km (McGovern et al. 2002), (Wieczorek and Zuber, 2004). These results are consistent with geochemical data retrieved from the SNC meteorites (Norman, 1999) and in-situ measurements (McLennan, 2001). Generally, younger surfaces overlay regions with thinner crust. Older surfaces match closely to the regions with thicker crust (Fig. 33 A, B).
The lowest crustal thickness marks old, large impact basins (e.g. Hellas, Utopia, Isidis) (Fig. 33 A, C). A large volcanic province, Tharsis, with four large shield volcanoes (Olympus Mons, Arsia, Pavonis, Ascreaus) (see Appendix 1) marks another characteristic feature of the Martian surface where the thickest crust has been detected. The considerably large thickness on southern hemisphere rapidly decreases while moving to the northern hemisphere. The division is often referred as Martian dichotomy. Its origin is widely discussed (see review: Zuber, 2001, Solomon et al. 2005) but no conclusions have been reached yet. What is also intriguing, the dichotomy is not only in crustal thickness and age but also in hypsometry (Fig. 33 C) as well as in chemical composition. The bulk of crust is generally composed from basalts and possibly basaltic andesites. Old southern heavily cratered highlands are most likely composed from products of primitive magmas. Products of more evolved magma, andesitic rocks are suspected to build the relatively smooth younger northern plains. A presence of the basaltic andesites is still a subject of debate (Bandfield et al. 2000), (Wyatt and McSween, 2002), (Rogers and Christensen, 2003), (Chevrier and Mathé, 2007), (Karunatillake et al. 2007), since the formation process would have required an environment similar to the one in a terrestrial subduction zone. This analogy supports an idea of plate tectonics that perhaps was present at the early stage of
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Fig. 33) A) Martian crustal thickness, B) Approximate surface ages, C) Topography map with major regions marked (Solomon et al. 2005). More detailed shaded relief map in Appendix 1.
A
B
C
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Martian evolution (Sleep, 1994), (Fairén and Dohm, 2004), (Lenardic et al. 2004). This very much intriguing hypothesis neatly explains water transport into the mantle thus supporting partial melting and more evolved magmas generation. Plate tectonics is also a very effective interior cooling mechanism that stimulates mantle and core convection.
Although such short activity cannot be completely rejected in the early Martian history e.g.
(Nimmo and Stevenson, 2000), surface dating at least for southern hemisphere, based on a crater counting method (Hartmann and Neukum, 2001) seems to diminish its extent.
Retrieved ages are close to a value obtained from ALH84001 meteorite that is considered as a fragment of an ancient crust (4-4,5Gyr –Nyquist et al. 2001, Weiss et al. 2002).
Shergottites, that crystallized much later (0.15-1.3Gyr- Chen and Wasserburg 1986, Jagoutz, 1991), show only a little change in geochemical markers. This line of evidences suggests that the crustal recycling, if ever existed, was strongly limited. Presently Mars is assumed to be a single plate planet (Solomatov, 1995), (Grasset and Parmentier, 1998), (Spohn et al. 2001) covering a dry, low viscosity mantle where subduction is not possible.
Although, rifting related magmatism is very unlikely in the stagnant lid convection regime, it does not mean that there was no volcanic activity on Mars in later periods. The most prominent example is, as already mentioned, Tharsis related volcanism. Although, a mechanism that lead to such a focused magmatism, remains a mystery, it is reasonable to assume its plume affinities even if present models cannot reproduce this feature (see review: Solomon et al. 2005). Recurring throughout the time a volcanic activity emplaced an old crust (Arkani-Hamed, 2004), (Johnson and Phillips, 2005), with 20-30km thick lava layers covering about 25% of the total planet’s surface (Anderson et al. 2001).
The surface mineralogy, although dominated by olivine, pyroxenes and mafic feldspars, shows also products of alteration (Fig. 34). Initial determination by landers (Viking 1 and 2, Pathfinder) and orbital spectroscopy e.g. (TES-MGS, THEMIS-Mars Odyssey, OMEGA-Mars Express) has been significantly refined by the recent rover missions (Spirit, Opportunity) (see review: Chevrier and Mathé, 2007). In the investigated areas rovers found weathering products of mafic minerals (e.g. talc, iddingsite), clay minerals (e.g. nontronite, montmorillonite), sulfates (e.g. jarosite, gypsum, kiserite) and iron compounds (e.g. hematite, goethite). All those findings provide important clues for water activity and ancient climate reconstruction leading to the current state (see 2.5).
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