Interiors of Terrestrial Planets and Major Moons

- August 2023 / 1 17 Chapter 5: Interiors of terrestrial planets and major moons • • • basic principles for modeling planetary interiors: internal heat; seismology; plate tectonics; magnetic fields comparisons of specific objects sample problems Basic principles of planetary interiors We have a few samples of the interiors of rocky / icy planetary bodies – many meteorites, for example, came our way courtesy of collisions that shattered their original parent asteroids – but by and large we rely on theoretical models supplemented by surface observations. We can observe seismic wave arrival times (at least on Earth), or the rate at which heat is emitted, or where and how frequently volcanoes erupt, or how the walls of craters sag, or how the object’s shape deviates from spherical or how the path of a passing spacecraft is affected – all these give us hints about the interior conditions. The most obvious piece of information about the interior of a planet is its mean density (mass / volume), because that’s related to its composition. Roughly speaking, ices have densities of ~1 g/cm3, surface rocks range from ~2.8 – 3.9 g/cm3, iron-rich minerals ~4½ – 5½ g/cm3, and nickel-iron meteorites, ~7 – 8 g/cm3. Mercury’s density is 5.43 g/cm3. Mercury can’t be just rock, but must have some fraction of denser material. Iron is a likely suspect, because it’s relatively cosmically abundant. Earth’s density is 5.52 g/cm3. . .but that doesn’t mean Earth has relatively more iron than Mercury. The reason is that Earth is massive enough to experience some amount of gravitational compression; our density is higher than the average of the densities of the planet’s constituent parts. Understanding the behavior of those constituent parts under the conditions of high pressure that exist at a planet’s core would help us disentangle the contributions of compression and composition to the overall density. We have evidence of grains of dust that pre-date the formation of the solar system; still, near the forming Sun the early solar system was warm and a lot of materials that we now think of as being solid would have been gas. The basic steps in the formation of the rocky / icy planetary bodies include initial condensation of gases into dust grains / snowflakes, accretion of those grains into larger objects, and, for objects that are large enough, differentiation. Refractory materials condense first, while the temperatures in the inner solar nebula are still quite high. At low pressures, elements such as tungsten and zirconium condense at temperatures of ~1,500 K. Oxides of calcium and aluminum soon follow, as evidenced by tiny droplets found in some primitive meteorites (the mineral perovskite, CaTiO3, is one example). Iron-nickel compounds and rocky silicates (e.g., enstatite, MgSiO3) are next, condensing at temperatures closer to 1,000 K. Below 500 K, which mostly means farther out in the solar system than the Earth, it cools enough for materials rich in carbon to condense and for water, ammonia, and methane to form ices. Reality is not quite so neat as these few examples suggest because chemistry keeps happening as dust grains are forming and the material in the protoplanetary disk cools. For instance, iron will react with the gas H2S at ~700 K, forming troilite (FeS), and will react with H2O at ~500 K to form iron oxide (FeO). The iron oxide will react with silicates to produce silicate minerals with relatively more iron; olivine, (Mg,Fe)2SiO4, and pyroxene, (Mg,Fe)SiO3, are two examples. Our most common ices may be water, ammonia (NH3) and methane (CH4) but at very cold temperatures, ~25 K, we find CO and N2 ices instead. Internal heat We expect that for smaller objects, e.g. asteroids or moons a few kilometers or a few tens of km in size, the average density and the density of the surface rocks should be more similar because small objects are not large enough to have differentiated. Small objects have less accretional heat to begin with and cool more rapidly because their ratio of surface area to volume is larger. Planets and major moons are large enough to have had enough internal heat to have become at least partially melted, enough for the denser 2 17 1 3 GM 2 E 0.3⋅ 6.67 ⋅10 −11 m 3 / (kg ⋅s 2 )⋅ 6 ⋅10 24 kg ΔT = ≈ 25 R = ≈ 20,000 K. M ⋅ cP M ⋅ cP 6.4 ⋅10 6 m ⋅ 900 J / (kg ⋅ K) 3 17 Figure 5.1: Folded rock. Chewuch river, northern Cascade range, Washington State. Photo by AKD. 4 17 Figure 5.2: Phase diagram, water 5 17 Figure 5.3: Isostatic compensation The icy surface of Europa shows much less vertical relief than our Moon does. Images of the surface of Pluto from the recent New Horizons flyby show icy mountain ranges, at least one of which is as high as the Rockies. It’s clear that Pluto’s crust is able to support taller mountains than Europa’s can and thus Pluto’s crust must be stiffer than Europa’s. Given how cold it is, that is no shock!; on the other hand, some models of Pluto’s surface features suggest that Pluto might also have a subsurface layer of liquid. That hypothesis is supported by the presence of ammonia near cracks on Pluto’s surface; ammonia should be destroyed by 6 17 Figure 5.4: Seismic waves 7 17 Figure 5.6: P waves from a shallow quake propagating through the Earth. (This source web site, as of August 2015, has a flash animation of waves propagating through the Earth.) 8 17 Figure 5.7: example seismometer recording Figure 5.8: Plate tectonics The new crust formed at spreading ridges is typically basalt, which is somewhat denser than rocks such as granites that make up much of our continental crust. It cools and thickens as it gets carried away from the warm spreading ridge. Near the ridge the lithosphere might be barely 10 km thick; across a large ocean, about to dive into a trench, the ocean lithosphere might be more than 100 km thick. Continental lithosphere, which includes the less dense continental crust, is thicker, ~200 km. Note that the Earth’s interior is not quite so neat as the sketches in figures 5.6 and 5.8. Seismic waves don’t change their speeds smoothly with depth in the lower mantle in the same manner at all locations on Earth. Most noticeable are two large irregularly shaped “blobs”, called large low-shear-velocity provinces, lying at great depth, one 9 17 Figure 5.9: Plate tectonics, a bit more detailed; Illustration by José F. Vigil in This Dynamic Planet , USGS / Smithsonian Inst. / US Naval Research Lab. Figure 5.10: Mid-Atlantic Ridge 10 17 Figure 5.11: Tectonic plates 11 17 Figure 5.12: Ring of Fire from This Dynamic Planet Smithsonian Institution, USGS, US Naval Research Laboratory 12 17 Figure 5.14: Dipole magnetic field 13 17 Figure 5.15: Magnetic field orientation record at a mid-ocean ridge http://pubs.usgs.gov/gip/dynamic/developing.html Between magnetic reversals there are also shorter, temporary geomagnetic excursions when the magnetic axis wanders much more than usual and may flip, and the field strength decreases, possibly not globally, over several hundreds of years before returning to the dominant field strength and orientation. The Laschamps event, ~ 41,400 years ago, is the best known excursion. A weakened magnetic field offers less protection from high-energy cosmic rays; there is some evidence suggesting that higher-than-average numbers of extinctions are associated with periods of weaker magnetic fields. Here too one might ask whether we see similar magnetic field patterns anywhere else in the solar system indicative of magnetic field reversals. Again, the answer is “maybe”. There are some hints of magnetic striping on Mars. It’s not recent, but suggests that the early Mars might have had incipient plate 14 17 Figure 5.17: Core sizes for rocky objects. Modeling the large icy outer solar system objects is more complicated. Water ice, as noted above, has many different phases. Water with additives such as salts and organic molecules has a lower freezing 15 17 Figure 5.18: Ganymede interior model NASA / JPL https://photojournal.jpl.nasa.gov/catalog/PIA18005 object average density (g/cm3) surface magnetic field (µT) Mercury 5.43 0.25 Venus 5.24 ~5 ⋅10−4 Earth 5.51 31 Moon 3.35 Mars field source / comments interior liquid layer? dynamo? partially molten core? induced by solar wind Fe-Ni core? dynamo Fe-Ni outer core ~0.03 localized crustal field thin outer core 3.93 ~10−3 loc. crustal field + solar wind Fe/FeS core? Io 3.53 ~1.3 interacts with Jupiter’s field thin outer mantle shell Europa 3.01 ~0.2 interacts with Jupiter’s field ocean Ganymede 1.94 ~0.75 intrinsic; dynamo? ocean; possibly two Callisto 1.83 ~0.04 Titan 1.88 ~0 interacts with Saturn’s field ocean Enceladus 1.61 ~0 interacts with Saturn’s field ocean Pluto 1.86 no? ocean? not completely differentiated ocean? 16 17 ΔP Δr = − ρ g. 17 17