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Information about Ge11cDIfferentiation

Published on February 20, 2008

Author: Haggrid


Terrestrial Planetary Differentiation:  Terrestrial Planetary Differentiation Guest lecture for Ge/Ay 11c/103 Paul Asimow April 25, 2005 Planetary Evolution:  Planetary Evolution Giant Molecular Cloud Solar Nebula Undifferentiated Planetesimals Differentiated Planets This weeks topic(s): How do you get from here to here? Planetary Differentiation:  Planetary Differentiation Defined as the family of processes subsequent to or accompanying accretion that cause large-scale segregation of material into different compositional reservoirs These are generally physical processes (melting, impacts, gravity-driven flows) but we recognize their occurrence mostly through their chemical consequences (separation of the elements) There is a competing family of mixing/ homogenization/recycling processes, but evidence shows that, on the whole, differentiation wins We know the average composition of solar system matter:  We know the average composition of solar system matter Nearly undifferentiated objects do exist:  Nearly undifferentiated objects do exist Carbonaceous chondrites are excellent models of bulk solar system composition except for most volatile elements. C class asteroids appear to be spectrally similar to such meteorites. These objects imply that some solid bodies formed without separation into chemically distinct reservoirs and maintain that homogeneity to the present. Zr Planetary Differentiation:  Planetary Differentiation However, all objects larger than some critical size (somewhere ≤1000 km radius) are differentiated Established, beyond dispute, by direct observation of surface composition and/or by geophysical inference of deep structure Challenge for models: Explain differentiation of 4 Vesta but not 1 Ceres or 2 Pallas Vesta Moon Mars Planetary Differentiation:  Planetary Differentiation Why? There is a motive Layers of different chemical composition can have different density, and gravity provides a driving force whereby planets can lower their potential energy by sorting the denser material towards the center There is a means Solids are hard to sort, but liquids are easily separated. Partial or complete melting allows large-scale differentiation There was an opportunity Heating beyond the melting point of at least some components of undifferentiated solar material during planet formation is inevitable for bodies above a certain size that formed early enough or fast enough Hot planet formation :  Assembling a planet means dumping lots of matter into a gravitational field, leading to large negative potential energy and large positive kinetic energy which is dissipated as heat by the time the matter comes to rest. Total gravitational binding energy of a uniform-density planet increases with the fifth power of radius. For the Earth, Of course, slow accretion could allow this heat to escape, but models as well as timing constraints from short-lived isotopes show that accretion was much too fast for large bodies to stay cold. Differentiation is driven by further lowering the gravitational binding energy. The actual density structure of the Earth has 10% lower potential energy than the uniform sphere. Hence, once core formation begins, it is catastrophic and self-sustaining. If Earth formed homogenous and then a core separated, enough energy would have been liberated to heat everything 3000 K and melt it completely. Hot planet formation Hot planet formation :  As an example (small compared to whole planet formation), consider the moon-forming impact between an ~80% proto-Earth and a ~10% Earth-mass object… Hot planet formation Robin Canup Hot planet formation :  But that’s not all. There are other sources of energy for heating early planets. Hot planet formation Short-lived radioisotopes For example, 26Al (half-life 0.7 Ma) Wasserburg and co-workers showed that CAI’s formed with 26Al/27Al ~ 5 x 10-5 Decay of all the 26Al in material of CI composition at CAI time liberates 6 kJ/g (!), with initial heating of 10–2 K/yr Lee, Papanasatassiou and Wasserburg Hot planet formation :  A model of thermal evolution of asteroids due to 26Al heating; assumes more rapid growth at smaller semi-major axis to melt Vesta but not Ceres Hot planet formation Ghosh, Weidenschilling and McSween Hot planet formation :  Another possibility is electromagnetic induction heating by T-Tauri solar magnetic field; with suitable tweaking this model can explain radial dependence of asteroid differentiation Hot planet formation Herbert and Sonett Differentiation: Major Reservoirs :  All the terrestrial planets have (or had) four major reservoirs, resulting from differentiation: Core, mantle, crust, atmosphere Core formation is the most energetic and most permanent segregation event. Completed early in planetary evolution. Probably no significant re-mixing Crust-mantle separation takes longer, may still be going on, and may involve some amount of recycling/mixing (certainly on Earth) Atmosphere formation (“degassing”) mostly occurred early, but continues as long as igneous activity continues; only a few atmospheric components are recycled (CO2, H2O, Xe? On Earth) Differentiation: Major Reservoirs Core formation: chemical evidence:  Core formation: chemical evidence Relative to volatility trend, some elements are grossly depleted in silicate portion of the earth (but N.B. the most depleted elements are in chondritic relative proportions) …if our understanding of accretion is right there is a big hidden reservoir. What do the depleted elements have in common? Geochemical Affinity:  Geochemical Affinity • In the classification scheme of Goldschmidt, elements are divided according to how they partition between coexisting silicate liquid, sulfide liquid, metallic liquid, and gas phase…defined by examining ore smelting slags and meteorites Silicate Liquid Sulfide Liquid Metallic Liquid Gas Phase Siderophile Chalcophile Lithophile Atmophile H, He, N, Noble gases Alkalis, Alkaline Earths, Halogens, B, O, Al, Si, Sc, Ti, V, Cr, Mn, Y, Zr, Nb, Lanthanides, Hf, Ta, Th, U Cu, Zn, Ga, Ag, Cd, In, Hg, Tl, As, S, Sb, Se, Pb, Bi, Te Fe, Co, Ni, Ru, Rh, Pd, Os, Ir, Pt, Mo, Re, Au, C, P, Ge, Sn • To first order, the distribution of elements between core and mantle resembles equilibrium partitioning between metal liquid and silicates…confirmed by iron and achondrite meteorites (but at high P, no separate sulfide phase) • Melting a chondrite gives 3 immiscible liquids plus vapor: Geochemical Affinity and Electronic Chemistry:  Geochemical Affinity and Electronic Chemistry • OK, but what makes an element siderophile or lithophile? Notably, the Goldschmidt categories are well-grouped in the periodic table of the elements: • Siderophiles - intermediate electronegativity, like metallic bonding; Lithophiles - prefer ionic bonds; Chalcophile - covalent bonding; Atmophile - noble Core Formation: When? :  Core Formation: When? We can distinguish whether (a) impact and short-lived nuclides or (b) long-lived radionuclides raised T to melting and allowed core formation by determining how quickly it occurred Moon postdates core formation and age of moon is no more than ~60 Ma after formation of meteorites; moon formation is part of earth accretion 182Hf-182W (extinct siderophile-lithophile pair): Earth and moon are not chondritic, so core formation ≤ 30 Ma after iron meteorite formation Xe isotopes requires that accretion completed 50-70 Ma after meteorites Pb segregation into core or by volatile loss altered U/Pb ratio of mantle affecting subsequent evolution of Pb isotopes; implies t < 100 Ma Conclusion: Core formation before the end of accretion, too late for short-lived nuclide heating, too fast for long-lived nuclide heating…impact driven 4.55 Ga 4.50 4.45 formation of chondrites formation of irons and achondrites end of earth accretion age of moon permissible range of core formation times Core Formation: more How?:  Core Formation: more How? Early differentiation in Moon-sized bodies collision EMULSIFICATION DURING IMPACT (Hf-W timescale ~ planet formation timescale if emulsification is sufficiently small scale Early differentiation in Moon-sized bodies collision CORE MERGING EVENT (Hf-W timescale ≠ planet formation timescale Evidence for cores in terrestrial planets:  Evidence for cores in terrestrial planets Chemistry Earth, Moon, Mars, Basaltic achondrites are all depleted in siderophile elements Iron meteorites and M-class asteroids are fragments of cores Magnetic Fields Earth, Mercury, early Mars and Moon Moment of Inertia ratio I/MR2 Uniform sphere, 0.4; Moon, 0.3932 Earth, 0.33; Mars, 0.365; Mercury, 0.34; Venus not measured, thought to be ~0.33 Seismology For the moment, Earth only Crust Formation:  Crust Formation All the differentiated terrestrial planets also have a crust, and all these crusts are broadly basaltic Exceptions/special cases: lunar highlands, Earth continents Evidence: Earth and Lunar maria by direct sampling Mars from SNC meteorites and some surface chemistry data from Viking, Pathfinder, and MER Venus from trace element data from Venera Mercury from limited spectroscopy data Crust: Mercury:  Crust: Mercury Mercury from limited spectroscopy data Mostly very-low FeO, similar to lunar highlands Possibly some maria-like features in smoothest, youngest terrains Crust: Venus:  Crust: Venus Major elements by XRF at 3 sites, K-U-Th from gamma spectroscopy at 5 sites Mostly basalt or alkali basalt Venera 8 is possibly a more evolved plutonic rock Sparse pancake domes suggest some evolved magmas Crust: Moon, Mars, and basaltic achondite parents:  Crust: Moon, Mars, and basaltic achondite parents Postpone these to Wednesday and Friday’s lectures Earth: mantle:  Earth: mantle The average composition of the Earth’s upper mantle is: SiO2 TiO2 Al2O3 FeO MgO CaO Na2O 46% 0.2% 4% 7.5% 38% 3.2% 0.3% Nothing else amounts to more than 0.5%. H2O content ~100 ppm. (Mg+Fe+Ca)/(Si+Al) is between 1 and 2, so the upper mantle is dominated by olivines (isolated tetrahedra structure) and pyroxenes (chain structure). olivine (Mg,Fe)2SiO4 [Mg/(Mg+Fe)~0.9] orthopyroxene (Mg,Fe)2SiO6 clinopyroxene Ca(Mg,Fe)Si2O6 Plus an aluminous mineral that depends on pressure: 0-1 GPa, feldspar (plagioclase) CaAl2Si2O8-NaAlSi3O8 [Ca/(Ca+Na) ~0.9] 1-3 GPa, spinel MgAl2O4 >3 GPa, garnet (Fe,Mg,Ca)3Al2Si3O12 A rock with this mineralogy is a peridotite. Earth: oceanic crust:  Earth: oceanic crust The average composition of the Earth’s oceanic crust is: SiO2 TiO2 Al2O3 FeO MgO CaO Na2O K2O 50.5% 1.6% 15% 10.5% 7.6% 11.3% 2.7% 0.1% Nothing else amounts to more than 0.5%. H2O content is ~1000 ppm. Note large enrichments over mantle of TiO2, Al2O3, CaO, Na2O, and K2O; small enrichments of SiO2 and FeO; massive depletion of MgO. (Fe+Mg+Ca)/(Si+Ti+Al) ~ 1, so basalts are dominated by pyroxenes, with alkalis in feldspar Clinopyroxene Ca(Mg,Fe)Si2O6 Feldspar (plagioclase) CaAl2Si2O8-NaAlSi3O8 [Ca/(Ca+Na) ~0.4-0.7] plus olivine, orthopyroxene, and perhaps a bit of quartz. H2O lives in Amphibole (hornblende) Ca2(Mg,Fe)4Al2Si7O22(OH)2 A volcanic rock with this mineralogy is a basalt. A plutonic rock with this mineralogy is a gabbro. Earth: continental crust:  Earth: continental crust The average composition of the Earth’s continental crust is: SiO2 TiO2 Al2O3 FeO MgO CaO Na2O K2O 57% 0.9% 16% 9% 5% 7.4% 3.1% 1.0% Nothing else amounts to more than 0.5%. H2O content is variable, up to several percent. Note even larger enrichments over mantle of SiO2, K2O. There are few octahedral cations, so lots of framework silicates (quartz and feldspars to take alkalis). H2O gives micas & amphiboles before alteration Feldspars (plagioclase) CaAl2Si2O8-NaAlSi3O8 [Ca/(Ca+Na) ~0.1-0.6] Feldspar (Alkali feldspar) NaAlSi3O8-KAlSi3O8 Quartz SiO2 Mica: biotite KMg3(AlSi3)O10(OH)2 Mica: Muscovite KAl2(AlSi3)O10(OH)2 Volcanic rocks with this composition range from andesite to rhyolite. Plutonic rocks range from diorite to granite. Igneous Differentiation: Density Relations:  Igneous Differentiation: Density Relations The structure of planetary crusts and mantles results from a fascinating interplay between chemistry (melting relations, mineralogical structures, partition coefficients) and physics (liquid and solid densities) There are several interesting twists because chemistry depends on ionic radius and charge, whereas density also depends on atomic mass Minerals:  Minerals Most minerals are best thought of as periodic structures constructed by packing of ions (either single-atom ions like Na+ or compound ions like carbonate CO32-). The ionic radii and charge balance are the critical factors determining mineral structure. Since anions (negatively charged ions) are generally big and cations (positively charged ions) generally small, the volume of a mineral is usually dominated by the anions, with cations occupying interstitial spaces. Radius determines whether a cation is likely to be coordinated by 4 (tetrahedral), 6 (octahedral), 8, or 12 oxygen ions. Melting, mineralogy, and differentiation:  Melting, mineralogy, and differentiation Why does partial melting of mantle yield enrichment in partial melt (which goes to form crust) of SiO2, Al2O3, FeO, CaO, Na2O, K2O; leaving a residue enriched in MgO? We can gain insight into this with a few essential phase diagrams. The olivine binary phase loop: an example of continuous solid solution. Mg has a higher ionic potential than Fe, so by Goldschmidt’s rules Mg end-members have higher melting points than Fe end-members. The phase diagram shows that this translates into Mg being more compatible than Fe…the liquid is always enriched in Fe/Mg relative to the residue. Melting, mineralogy, and differentiation:  Melting, mineralogy, and differentiation The Mg2SiO4-SiO2 binary: an example with negligible solid solution and an intermediate phase. The first liquid that appears on melting of a rock consisting of forsterite (olivine) plus enstatite (orthopyroxene) is more SiO2-rich than enstatite. If we turn around and crystallize it, it will make enstatite plus quartz (a model basalt, not a peridotite!). This shows how oceanic crust is enriched in SiO2. We can make similar arguments for CaO, Na2O, and K2O, but they require ternary phase diagrams... Melting, mineralogy, and differentiation:  Melting, mineralogy, and differentiation The Forsterite-anorthite-silica ternary: shows large enrichment of CaO and Al2O3 in partial melts Melting, mineralogy, and differentiation:  Melting, mineralogy, and differentiation The result of the chemical differences between terrestrial mantles and their partial melts: At pressure <8 GPa, partial melts are buoyant and segregate upwards At pressure <1 GPa, basalts crystallize lots of plagioclase and are the solidified rocks are buoyant relative to peridotite But basalts are rich in Fe and, at high pressure, the plagioclase recrystallizes to garnet, which makes eclogite - denser than peridotite! In magma oceans, partial crystallization has its own set of density relations: Plagioclase floats in most melts, hence anorthositic crust can develop on Moon and mercury Perovskite sinks; olivine may sink or float; in some models this creates upper/lower mantle distinction on Earth Partition coefficients and Earth differentiation:  Partition coefficients and Earth differentiation Partition coefficients can be measured experimentally at particular conditions, or inferred from natural samples. The partition coefficients that obtained during melting of the primitive mantle to form the continents can be obtained (on the assumption of batch melting) from the bulk composition of the continental crust: Continental crust Mid-ocean ridge basalt Here elements are ordered by enrichment in the continental crust over bulk silicate earth, a sort of qualitative partition coefficient. If we assume DRb=0, then F=1.6% and we may assign D to all the other elements. Partition coefficients and Earth differentiation:  Partition coefficients and Earth differentiation The humped pattern of mid-ocean ridge basalts in these figures can be modeled as resulting from 8% melting of the source previously depleted of incompatible elements by 1.6% melting to form the continental crust. This demonstrates that the upper mantle is the complementary depleted reservoir to the continents. Partition coefficients and Earth differentiation:  Partition coefficients and Earth differentiation D=Cliquid/Cresidue F~Cresidue/Cliquid @D~0 Differentiation and radiogenic heat production:  Differentiation and radiogenic heat production All the significant long-lived radioactive nuclei (40K, 87Rb, 147Sm,187Re, 232Th, 235U, 238U) are lithophile and incompatible Hence they all partition into mantle, leaving core with only initial heat (possible exception: 40K) Also, they all concentrate strongly into crust, leaving depleted mantle with substantially less heat production Differentiation and radiogenic heat production:  Differentiation and radiogenic heat production This is thought to be an issue for Mars, since the Nakhlite source is 2x more depleted than the MORB source and the shergottite source is 6x more depleted - yet shergottite igneous activity continued until <200 Ma. Jones (2003) proposes a layered Martian mantle with a deep nakhlite source, a thermal boundary layer, and a shallow shergottite source Where did the atmosphere come from?:  Where did the atmosphere come from? A primordial atmosphere was formed during accretion by direct outgassing from the magma ocean and from volatile-rich impactors. It was rich in reducing species (in equilibrium with H2, no O2) and noble gases. This atmosphere was lost by catastrophic impacts; the Earth did not become closed to Xe loss until ~4.4 Ga. In the early Archean, a second atmosphere formed by outgassing from the Earth, now almost a closed system. Whether this was a strongly reducing atmosphere (H2, CH4, NH3) or a weakly reducing atmosphere (H2O, CO2, N2) depends on the relative timing of core formation, which oxidized the mantle by withdrawing metallic Fe -> probably weakly reducing. There was no O2, and so no O3, so solar UV reached surface Earth probably looked like Venus, with enough CO2 and H2O in the atmosphere to make a very hot surface and keep H2O in form of steam. But at 1 A.U. the equilibrium temperature would have settled at ~85°C, just cool enough to condense the ocean 38 Early atmosphere:  Early atmosphere A reducing atmosphere or a high solar UV flux favors loss of H2 from the exosphere. How do we trace how much H a planet has lost? 1. Stable isotopes of hydrogen, which is directly fractionated by preferential loss of 1H relative to 2D Earth: D/H = 1.5 x 10-4 Mars: D/H = 9 x 10–4 Venus: D/H = 0.022 2. Stable isotopes of other noble gases, which are fractionated during hydrodynamic escape driven by a large flux of escaping hydrogen: All the terrestrial planets have 20Ne/36Ar ~ 0.01x Solar Since H is lost as molecular or atomic hydrogen, the effect is irreversible oxidation of the planet Together with core formation, leads to a weakly oxidizing early atmosphere 39 The Ne 3-isotope diagram:  The Ne 3-isotope diagram Nucleogenic production moves degassed mantle to the right; undegassed mantle moves less because of low U/22Ne. The Ne 3-isotope diagram:  The Ne 3-isotope diagram Note that atmospheric composition and solar composition lie (approximately) on a mass fractionation line. Any physical or chemical process, thermodynamic or kinetic, fractionates isotopes by mass and will change 20Ne/22Ne twice as much as 21Ne/22Ne. But if Earth started out with Solar Ne isotope composition, how did the atmosphere get to be heavier in isotopic composition? Preferential degassing of light isotopes would make atmosphere lighter than mantle. Best story is hydrodynamic escape: earliest earth atmosphere was dominated by H2, whose escape flux under influence of early solar wind was so big it could carry away other atoms along with it. Light Ne was preferentially lost from the earth, leaving an isotopically heavy atmosphere. Terrestrial Xenology:  Terrestrial Xenology Fundamental observation: MORB data differ from atmosphere, and show mixing between air and a component with excesses of both 129Xe (from 129I decay) and fissiogenic Xe isotopes (from Pu and/or U decay) Continental samples (granites, etc.) show only fissiogenic Xe from U decay. CO2 well gases from continents show mixing between continental and mantle components. No OIB has ever shown Xe isotopes different from air -- either totally contaminated or lower mantle has atmospheric composition, we do not know which! Terrestrial Xenology:  Terrestrial Xenology Here is a simple degassing and gas-loss model: total loss of anything degassed until closure, total retention of everything since. With suitable choice of initial Xe composition, 129I and 244Pu abundances, you can make both I and Pu clocks give same age, 90 Ma after origin of solar system Terrestrial Xenology:  Terrestrial Xenology Unlike Ne and Ar, the mantle composition for the nonradiogenic Xe isotopes is similar to atmospheric, not to solar. Hence either (1) whatever fractionated the atmosphere in Ne and Ar was unable to separate Xe isotopes, presumably because Xe is too heavy to escape, or (2) the upper mantle Xe is mostly recycled atmosphere because Xe is retentive enough to be subducted. Summary of Earth Differentiation:  Summary of Earth Differentiation (refractories) (volatiles) (siderophile & chalcophile) (lithophile) (atmophile) (lost due to impacts) (late veneer) Solar Nebula Condensation and Accretion Core Silicate Earth Primitive Atmosphere Inner Core Outer Core Primitive Mantle Moon Lower Mantle Upper Mantle Continental Crust Oceanic Crust Modern Ocean & Atmosphere (continuing cometary flux?) (partial melting; liquid-crystal partitioning) (plate tectonics: partial melting, recycling) (nucleosynthesis, mixing) (gas-solid equilibria) (melting; gravity and geochemical affinity) (freezing) (catastrophic impact) (hotspot plumes) degassing degassing

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