Seismology of Earth s interior. Structure of other terrestrial bodies (6.3) Structure of the giant planets (6.4)

Transcription

1 Planetary interiors

2 Topics to be covered Models of interior structure (6.1) Seismology of Earth s interior Seismology (6.2.1) Density profile of the Earth ( ) Structure of other terrestrial bodies (6.3) Structure of the giant planets (6.4) 2/76

3 Bulk properties of the planets 3/76 Increasing distance from Sun

4 Primary materials Why are planets made of these? Because they (or at least their constituent atoms) are abundantly produced by stars (see chapter on planet formation) The primary building blocks of the planets are metals (ρ ~ 8g/cm 3 ), silicates (ρ ~ 3g/cm 3 ), ices (ρ ~ 1g/cm 3 ) and H & He (ρ low, but varies with pressure) Because materials with higher condensation temperatures will be able to condense closer to the Sun when the planets were forming, we can use the bulk densities of the planets to infer that Mercury, Venus, Earth and Mars (terrestrial planets) are rock/metal Jupiter and Saturn (gas giants) are primarily H/He (but probably have an ice/rock core) Uranus and Neptune (ice giants) contain a lot of ices 4/76

5 Hydrostatic equilibrium If the material in the interior of the planet is not moving, then there must be zero net force everywhere inside (note: this ignores possible convection inside the planet) The weight of material above any point must be balanced by an equal upward pressure. (Hydrostatic equilibrium) The weight depends on: Density of material a above Composition of material Temperature (and phase) Pressure above (relating to density) Thus the accurate determination of a planet s internal structure, from the outside, can be difficult. 5/76

6 Melting point The melting point of most materials is also a function of pressure The increase in melting point with pressure for many (not all) common materials means that a body could have an outer liquid layer on top of a hotter solid lower layer. 6/76

7 Equation of state To determine the phase of a particular portion of the planet, we need the phase diagram for the material at the temperature T and also the pressure P in question The information in a phase diagram is equivalent to knowing the equation of state of the material, relating ρ, T and P Note that mixtures may be complicated, involving the coexistence of different minerals as liquids and solids together. Even pure materials may exist in a variety of solid and liquid states, each with different crystal structures, densities, etc. An example of a phase diagram 7/76

8 Equation of state For an ideal gas, it s simply P = nkt The equation of states for highly compressed gases, liquid and solids may be much more complicated For a solid or liquid, one usually tries to fit an empirical relation to whatever laboratory data is at hand. At higher temperatures or pressures, this may involve lasers and/or diamond anvil cells (to generate high temperatures and pressures) or studying as best one can the effects of shocks (e.g. explosions) on samples. Some progress can alos be made sometimes from theoretical quantum mechanical calculations, but the equation of state is often 8/76 largely unknown for planetary interiors

9 Giant planet T ranges from K in atmosphere to 10,000-20,000 K (hotter than surface of the Sun) at their centres. Pressure ranges up to Mbar at their center That their composition is primarily H and He can be guessed from their bulk densities (1.33 and 0.69 g/cm 3 for J &S resp) Giant planets 9/76

10 Hydrogen in Jupiter and Saturn Below 1 Mbar, molecular H 2 is dominant Beyond 1 Mbar liquid H 2 begins to dissociate to atomic H, a process complete at ~3 Mbar At higher temps, H may become degenerate or a plasma; however, the nature of the transition and its location is controversial. Dashed lines = adiabats for the giant planets, the expected T-P profiles for convective interiors 10/76

11 Jupiter In Jupiter, (86.4% hydrogen 12.6% helium by number) the transition from molecular hydrogen to atomic occurs between and d08r 0.8 J Owing to the close packing of molecules/atoms at these pressures, their electron orbitals overlap, allowing electrons to hop easily from one atom to the next (fluid metallic hydrogen, either molecular or atomic) This highly conductive fluid allows a strong magnetic field to be produced A core of rock/ice, around which the gas originally accumulated through gravity (perhaps Earth masses) remains at the core 11/76

12 Hydrogen and helium Helium in the core does not reach a metallic state at pressures seen in the giant planets He and H are not fully mixed at all points in the giant planets Their miscibility (ability to dissolve each other) varies with temperature and pressure At lower temperatures and pressures, H and He don t mix as well Separate Mixed 12/76

13 Atmosphere is observed to contain 96.7% hydrogen and 3.3% 3% helium by number Missing helium (compare to Jupiter) Because temperature at a given pressure is lower than Jupiter, immiscibility is more of an issue It is thought that droplets of helium condense out of the H/He mixture and rain down into the core (releasing grav. potential energy, an internal heat source) Saturn Molecular hydrogen Metallic hydrogen Rock/ice core 13/76

14 Understanding Jovian planet sizes and densities Stacking pillows increases the height, but eventually compression occurs. 14

15 Water is a major constituent t of the planets, particularly the outer planets and their satellites Solid water has at least 10 different crystalline forms varying in density from 0.92/cm 3 for ordinary ice (ice Ih, h=hexagonal, Ic=cubic is lower T form) to 1.66g/cm 3 for ice VII Interiors of the icy satellites are expected to vary from T=50-100K and P=0 at the surface to several hundred K and a few tens of kbars at their centers Water in the cores of the giant planets is likely supercritical (neither a liquid nor a gas) Ices Phase diagram for water. Dashed lines indicate adiabats for the giant planets 15/76

16 Uranus and Neptune Uranus and Neptune are often called ice giants because they are expected to be largely l water slush? H 2 /He/CH 4 /NH 3 gas Pure water is naturally slightly ionized into H 3 O + and OH - at 1 part in 10 million at room temperature; this is enhanced at higher T rocky core A mix of ice, non-conducting NH 3 and isopropanol ( synthetic Uranus ) also ionizes slightly at high pressure and thus becomes conducting. This conductivity is sufficient to account for the magnetic fields of Uranus and Neptune, Possible structure of Uranus and Neptune as the standard dynamo effect requires circulating conductive fluids. 16/76

17 Rocks and metals A great variety of minerals can be produced by the handful of rocky/metal elements that are most common on the terrestrial planets The particular mineral depends on the temperature, pressure and relative abundance of the constituents. 17/76

18 The Earth s mantle Earth s mantle begins below the 5-70 km thick crust The next layer, the mantle, is 2900 km thick and comprises 80% of the Earth s volume. 18/76

19 The Earth s mantle Olivine Composed primarily of olivine (=peridotite) (Mg,Fe) 2 SiO 4 olivine is rare in the crust) and pyroxene (Mg,Fe)SiO 3 Pyroxene 19/76

20 From the different speeds of propagation of seismic waves (later) we know that olivine changes to a (more densely- packed) spinel crystal structure at about 400 km This breaks down to periclase (MgO) and enstatite (MgSiO 3 ) at about 660 km MgSiO 3 is very stable at high h pressures and may be the dominant mineral in the Earth s interior i Note: the book calls MgSiO 3 perovskite, but that s CaTiO 3. MgSiO 3 is enstatite but has the same structure as perovskite and geologists often refer to it that way 20/76

21 Enstatite Periclase (green) on matrix Perovskite 21/76

22 The Earth s core Liquid Phase diagram (Fe) The Earth s core is likely primarily iron from its density Iron is well-studied up to 200 kbar but has not especially well-studied at high T and P The liquid can be solidified by high P It is known to have at least 5 solid phases (allotropes) with different structures and densities. One allotrope, α-fe is normal iron, β-fe is a nonmagnetic allotrope (though the lattice remains unchanged!) that s stable between about 760 and 900 C 22/76

23 Iron and sulfur The Earth s core has a density 5-10% lower than pure iron which indicates a mixture with S,O or H FeO is expected to combine with crustal minerals, but FeS is expected to settle in the core FeS shows eutectic behaviour: an alloy of FeS + Fe melts at temperatures below that of either Fe or FeS At 1 bar, Fe melts at 1808K, FeS at 1469K and a 27% S + 73% Fe mixture melts at 1262K At higher pressures, the melting temperature can be depressed by as much as 1000K, allowing the material to remain liquid at lower temperature. Pyrrhotite (FeS + S alloy) 23/76

24 Earth s core The Earth s core is about 7,100 kilometers in diameter, slightly larger than half the diameter of Earth and about the size of Mars. The outermost 2,250 kilometers of the core are liquid. Currents flowing in the core are thought to generate Earth's magnetic field. The innermost part of the core, about 2,600 kilometers in diameter, is made of a similar material to the outer core, but it is solid and perhaps even a giant single crystal of iron. 24/76

25 Earth s core temperature Earth gets hotter toward the center. At the bottom of the continental crust, the temperature is about 1000 C. The temperature increases about 1 C per kilometer below the crust. Geologists believe the temperature t of Earth's outer core is about 3700 to 4300 C. The inner core may be as hot as 7000 C -- hotter than the surface of the Sun. But, because it is under great pressure, the center of Earth remains solid. Depth-temperature p profile of the Earth 25/76

26 Earth s crystal core Iron crystal configurations. Bodycentered cubic (bcc), face- centered cubic (fcc) and hexagonal close-packed (hcp). Bcc seems to be ruled out. Tromp, J. (1993) Nature 366, Seismic waves from earthquakes pass through the earth's core faster when they travel parallel to the Earth's axis than when they travel in the plane of the equator. The transit time difference is 2-4 seconds. Apparently, the Earth's core is not perfectly spherical or its properties are different in different directions. The natural vibration or "ringing" frequencies of the Earth are "split," that is, instead of a series of single "tones" we detect a series of closely paired frequencies. This is symptomatic of a core that is anisotropic; that is, its properties are different in different directions. These could be explained if the Earth s inner core is a giant crystal aligned with its spin axis. 26/76

27 Gravity field If a planet were perfectly spherical, it would produce a gravitational field outside it equivalent to that of a point source. Φ g = GM p r In reality planets are not perfectly spherical. The equipotential surface is called the geoid. A glass of water will have a surface parallel to the geoid at any spot; a plumb bob will hang perpendicular to it. Earth s geoid (exaggerated) 27/76

28 Gravity field If the planet is axisymmetric, its gravitational potential can be approximated by a series expansion n GM R Φ = e r, θ, 1 g ( φ) J np n(cos θ ) r n=2 r where the angle θ is the colatitude the angle φ is the longitude R e the planet s equatorial radius, P n (cosθ) are Legendre polynomials J n are the gravitational moments of the planet, which are measures of their non-sphericity. 28/76

29 Gravity field and J n J 1 = 0 by our choice of coordinates (C of M at origin) For rotating fluid bodies, J n =0 where n=odd (e.g. giant planets) ) Φ = GM r 2 4 Re Re 1 J P2 (cosθ ) J 4 P4 r r 2 θ (cosθ ) L The coefficients J 2 2, J 4 etc. can be determined from spacecraft observations (the deviation of their orbits from that for a spherical planet) We can relate J 2,J 4... to the internal structure of the planet For the terrestrial planets, non-zero odd moments, as well as non- axisymmetric components of the geoid have been measured. 29/76

30 Gravitational moments of the planets 30/76

31 Isostatic equilibrium Consider a planet with a solid crust on a liquid interior layer Floating objects displace their own weight in the substance on which they float A large low-density mass (mountain) in isostatic equilibrium is compensated by a ʺdeficiencyʺ of mass underneath, with total mass of displaced matter (liquid) equal to the mass of the mountain A small dense mass (ocean crust, impact basin with thin crusts) in isostatic equilibrium has ʺextraʺ mass underneath it, so it sits lower than a low-density mass 31/76

32 Low-density wood blocks float high and have deep roots, whereas high-density blocks float low and have shallow roots. 32/76

33 33

34 Isostatic rebound In areas formerly covered by ice sheets (around the Baltic Sea and Hudson Bay, for example), sea cliffs and beach ridges are now found nearly 300 m above sea level! 14 C ages on marine shells and driftwood show that these features are postglacial (less than 14,000 years old). They were formed at sea level and have risen from isostatic rebound. Isostatic rebound usually occurs at an exponentially declining rate. The half-recovery time is commonly several thousand years, thus recovery is still continuing around the Baltic Sea and Hudson Bay, albeit much more slowly than it did immediately following deglaciation. 34/76

35 Earth s upper layers Crust: ocean crust (as little as 5km thick) and continental crust (up to 70km thick) Lithosphere: crust + rigid upper layers of mantle (extends down to about 80 km deep) Asthenosphere: liquid, highly viscous layer of mantle upon which the lithosphere floats. The base of the lithosphere is conventionally defined as the 1300 C isotherm since mantle rocks below this temperature t are sufficiently i cool to behave in a rigid manner. 35/76

36 Thickness of Earth s crust 36/76

37 Geothermal Gradients in the Earth The geothermal gradient varies widely with geography from 5 o C/km to 75 o C/km on Earth. 37/76

38 Heat sources If we observe planets and compare their luminosities (emitted heat) with the amount of sunlight they receive, several show signs of an excess, pointing to an internal heat source eg. Earth emits 75 erg cm -2 s -1 (0.075 W m -2 ) or 31 TW total Heat sources Gravitational Accretion Differentiation Radioactive decay Tidal dissipation Heat losses Conduction Radiation Convection Plate tectonics, volcanism Map of Earth s heat flow 38/76

39 Accretional heating Planets (or at least their cores) are thought to be accumulations of smaller bodies (planetesimals) brought together by gravity. The potential and/or kinetic energy of these bodies is converted largely into heat upon impact. 39/76

40 Accretional heating Assuming the material coming in is very cold, the GM energy input will be mainly due to the gravitational energy of the infalling body The protoplanet will radiate energy as a blackbody and cool. E& GMm& / in E& σt 4 out If accretion is slow, heat will be radiated into space faster than it is delivered and the body will not heat up If accretion is fast, heat will be buried and the temperature of the body will increase. r A 40/76

41 Differentiation The potential energy of a homogeneous planet is reduced if the densest components sink to the centre differentiation is energetically favoured Differentiation is opposed by the rigidity of body, so differentiation is generally favoured at higher temperatures Small/cold bodies may not be able to differentiate 41/76

42 Differentiation Differentiation releases gravitational energy, driving further differentiation positive feedback, potential runaway Heat released may generate thermal expansion and form a source of stress Sinking materials may undergo phase changes leading to volume changes and either expansion or contraction (Mercury) Differentiation may be ongoing in the giant planets as denser helium settles to the core (Saturn in particular) Helium rain-out on Saturn 42/76

43 Scarps on Mercury As we saw last week, Mercury is thought to have shrunk by a few km as it cooled, creating scarps 43/76

44 Energy budget of the outer planets Incident solar radiation much less than that at Earth, so surface (cloud top) temperatures are lower We can compare the amount of solar energy absorbed with that emitted. It turns out that there is usually an excess (not for Uranus though). Why? emitted After Hubbard, All units in W/m 2 in New Solar System (1999) 48 reflected incident Jupiter Saturn Uranus Neptune 44/76 0.3

45 Sources of Energy One major one is contraction gravitational energy converts to thermal energy. Helium sinking is another. Gravitational energy of a uniform sphere is 2 E g = 3GM / 5R So the rate of energy release during contraction is 2 de g dt 3GM = 2 5R dr dt Jupiter is radiating 3.5x10 17 W in excess of incident solar radiation. This implies it is contracting at a rate of 0.4 km / million years Another possibility is tidal dissipation in the interior. This turns out to be small for the giant planets. Radioactive decay is also a minor contributor for the outer planets. 45/76

46 Puzzles 1. Saturn has an excess heat loss that cannot be explained by primordial (accretional) heating, ongoing contraction, radioactivity it or tidal heating. However, it is consistent with gravitational energy released by differentiation as denser helium comes out of solution with hydrogen and sinks to the planet s core. 2. Why is Uranus heat budget so different (no internal heat source)? Perhaps due to compositional density differences inhibiting convection at levels deeper than ~0.6R p. May explain different abundances in HCN,CO between Uranus and Neptune atmospheres. These compositional differences may have been caused by an impact with a large protoplanetary body at early times, which may have also caused Uranus s s large obliquity (tilt). 46/76

47 Differentiation of the Earth For the accretion time of the Earth (~100 million year time scale), the temperature is not expected to have risen much directly from accretional heating (some debate). Yet the Earth is differentiated, which requires a rather large temperature Where could the extra heat have come from? 47/76

48 Radioactive decay Early on, short-lived radionuclides ( 26 Al -> 26 Mg in particular, 0.74 Myr half-life) would have contributed much more than longer-lived radionuclides do now Current radionuclides of geological interest About 50% of the Earth s current excess heat is thought to be due to radioactivity 235 U, 238 U, 232 Th, 40 K with half-lives of 0.71, 4.5, 13.9 and 1.4 Gyr respectively are now the major contributors 48/76

49 Tidal heating P b planet M R θ a m satellite Tidal potential at P Cosine rule b = a 1 2 m V G b = (recall acceleration = - V ) R a cosθ + R a 2 1/ 2 (R/a)<<1, so expand square root 2 m R R 1 = L 2 V G ( ) 1 + cosθ + 3cos θ + a a a 2 1 Constant Mean gravitational ti Tide-raising part of => No acceleration acceleration (Gm/a 2 ) the potential 49/76

50 Tidal bulges We can rewrite the tide-raising part of the potential as m G R ( 3cos θ 1 ) = HgP 2(cosθ ) 3 a 2 Where P 2 (cos θ) is a Legendre polynomial, g is the surface gravity of the planet, and H is the equilibrium tide g = GM 2 R H = m R M R a 3 a = semimajor axis of satellite orbit R = planet radius m/m = mass ratio For a uniform fluid planet with no elastic strength, the amplitude of the tidal bulge is 5H/2 An ice shell decoupled from the interior by an ocean will have a tidal bulge similar to that of the ocean For a rigid body, the tide may be reduced due to the elasticity of the planet 50/76

51 Rigidity or shear modulus We can write a dimensionless number which tells us how important the rigidity or shear modulus μ is compared with gravity: ~ 19 μ μ = (g is surface gravity, ρ is density) 2 ρgr For Earth, μ ~10 11 Pa, so ~ μ ~3 (gravity and rigidity are comparable) For a small icy satellite, μ ~10 10 Pa, so ~ μ ~ 10 2 (rigidity dominates) We can describe the response of the tidal bulge and tidal potential of an elastic body by the Love numbers h 2 and k 2, respectively h 2 5/ 2 ~ 3/ 2 = k2 = 1+ μ 1+ ~ μ ~ μ 51/76

52 Tidal dissipation For a rigid body, the tidal bulge amplitude is given by h 2 H The quantity k 2 is important in determining the tidal torque The rate of tidal dissipation, i the rate of energy being pumped into the satellite being flexed by the larger parental body, like the Earth or Jupiter is (homogeneous moon, synchronous rotation) E& = 21 k R GM 2 2 Q a 5 a 2 p ne where n is the mean motion of the satellite, a is the semimajor axis, R is the radius of the satellite or moon, e is the eccentricity cty of the orbit, bt, Q is sthe eseismic s quality factor or dissipation factor of the satellite and is (inversely) 52/76 related to the viscosity. 2

53 Tidal dissipation Tidal heating involves both viscoelasticity through Q and μ as well as orbital mechanics, and is time-dependent. Lower Q and lower μ = greater amount of tidal dissipation. This can also lead to positive feedback from temperature dependence of μ (higher T = lower rigidity) and Q (though in most simple fluids Q increases with T as the liquid becomes less viscous at larger T). Tidal dissipation is responsible for volcanism observed in Io (predictions made before the arrival of Pioneer in 1979 by Pat Cassen at NASA Ames) and other icy planets. Tidal heating may also be important In Ganymede and also in Europa (subsurface ocean?). Some values of Q Probably ~2x10 5 5x10 5 for Jupiter ~17-43 for Io, for Europa ~100 for terrestrial planets ( for most rocks and metals) Europa 53/76

54 Ohmic heating If a conducting object moves through a varying magnetic field, eddy currents will be set up inside it which will dissipate heat Ohmic heating may be important t for Io, Jupiter s innermost large moon. Ohmic heating has been invoked to melt planetesimals (in the early Sun s strong magnetic field) and thus to explain some properties of meteorites Ohmic heating is being touted by Agriculture Canada for industrial sausage production(!) 54/76

55 Seismology The study of the mechanical properties of the Earth, primarily by tracing the passage of elastic waves through it. These waves may be generated by earthquakes, volcanoes, meteorite impacts or artificially. Two types Body waves (S & P) through the bulk of the Earth Slower surface waves travel in the near-surface 55/76 layers

56 Surface waves Larger amplitude and longer duration that body waves, but slower to arrive. They are usually most destructive. Rayleigh waves move the surface in vertical ellipses, just like water waves (v~70% of S wave vel) Love waves move horizontally, across the direction of motion (v~90% of S wave vel) Rayleigh waves Love waves 56/76

57 Body waves Travel through the interior of the Earth Are reflected/refracted at interfaces in density/composition or slowly refracted by gradients in Earth s internal properties. Two types P (Primary, Push or Pressure) waves Longitudinal waves, compression and rarefaction along direction of motion Similar to sound waves Can travel through liquid S (Secondary, Shake or Shear) waves Shear waves, oscillations transverse to propagation Similar to EM waves, or waves in a rope Cannot travel through h liquidid 57/76

58 Body waves 58/76

59 P and S wave propagation Increasing sound speed with depth means body waves curve typically upwards PCP waves PKP waves P waves can travel through liquid but S waves cannot pass through liquid are reflected (PCP) and refracted This led to the discovery of a liquid (PKP) layer in the Earth s core, as there is a There is also a P-wave shadow zone S-wave shadow zone opposite where refraction prevents the arrival seismic sources where no direct s- of direct p-waves waves are received 59/76

60 Seismology and the interior of the Earth The Moho or Mohorovicic i discontinuity between the crust and mantle Varying depth: 5-10 km below ocean floor, 35 km below continents, <60 km below mountain ranges m thick P wave velocity increases Discontinuities in upper mantle Depths below 670 km Stepwise increasing P and S wave velocities Phase and density changes of material, e.g., olivine spinel enstatite 60/76

61 Seismology and the interior of the Earth The interface between the solid mantle and liquid outer core 3000 km depth P wave velocity decreases, S waves disappear The interface between the liquid outer core and the solid inner core 5200 km depth P wave velocity increases S waves reappear 61/76

62 Interior of the Moon crust depth <5 km under maria, up to >100 km under highlands on average thicker on the far side, center-of-mass to center-of-geometry displacement of 1.7 km towards the Earth in the Earth- Moon direction fractured at <25 km depth Back side of Moon 62/76

63 mantle Interior of the Moon upper mantle at depths <500 km, olivine middle mantle at depths <1000 km, olivine and pyroxene lower mantle at depths <1400 km, S wave velocity decreases, partially molten? core P wave velocity increases, solid radius < km iron-rich seismically active zones near surface (meteors) at depths of km (tidal) 63/76

64 Mercury Lithosphere depth 200 km scarps (1-3km high), possibly due to contraction after cooling Mantle depth 600 km rocky, silicate material abnormally thin overal Fe content of Mercury is 2x chondritic Core size 75% of the planetary radius (high bulk density) iron, or FeS outer liquid core, probably convective (magnetic field) large core may be primordial, or due to a large late impact that vaporized much of the mantle 64/76

65 Venus Similar interior structure to Earth, as suggested from size and mean density Surface heat flow probably less than on Earth interior may be heating up with radioactive heating most effective in mantle, the core may be relatively cooler Lithosphere: e high surface temperature suggests it is relatively thin, depth about 20 km heat loss less than on Earth: hot-spot and volcanic activity possibly more important absence of planet-wide tectonic plate activity (which is lubricated by H 2 O on Earth) crust may be too buoyant to subduct may lead to catastrophic crust renewal every few ~10 8 years Computer generated surface view of Venus Eistla Regio (from the northeast). 65/76

66 Mantle Venus Lower-viscosity asthenosphere may be lacking due to volatile (particularly water) depletion Core probably less FeS than Earth s core owing to its formation closer to the Sun (hotter = less volatile sulfur) which implies higher freezing T Venus has no magnetic field, which implies that there is no conductive core possible state: frozen solid liquid but without liquid to solid phase change that drives convection in Earth s core 66/76

67 Mars Surface and mantle high in FeO relative to Earth, as well,, and as in volatiles ((Na,P,K Rb) Lithosphere Low surface temperature p suggests relatively thick lithosphere elevation difference 5 km between S and N hemispheres absence of plate tectonics significant early heat loss from large shield volcanos (eg. Ol mp s Mons) Olympus topographic features are not fully isostatically compensated, g consistent with thick rigid lithosphere Topographic map of Mars (Mars Global Surveyor) 67/76

68 Mantle Mars silicate composition probably be enriched in FeO volcanic magma sources at a depth of 200 km (based on viscosity) Core solid Fe-Ni, or liquid Fe-FeS with larger radius weak magnetic field implies no liquid layer but models do not reproduce this 68/76

69 Interior structure of the terrestrial planets 69

70 Internal structure is modelled, but constrained by mass, radius, rotation, oblateness, internal heat sources, J n, etc Core Probably relatively small, 5-10 Earth masses Probably an inner rocky/iron and an outer ice-rich component, but could also be homogeneous Much mass was present after solidbody accretion, but some may have been added later by gravitational settling Jupiter Adiabatic (fully convective) model 70/76

71 Core envelope High-Z (Z>2) elements constitute Earth masses distributed within core and its surrounding envelope Enriched in high-z Mantle material by a factor 3-5 compared to chondritic composition Large amount of ices of H 2 O, NH 3, CH 4 and S- bearing materials Jupiter Most internally generated heat is attributed to gravitational contraction and accretion in the 71/76 past

72 Saturn Similar interior structure to Jupiter Like Jupiter, the total mass of high-z elements is about Earth masses As Saturn is smaller and less massive than Jupiter, the relative abundance of high-z elements is thereby greater consistent with atmospheric C, N, and S which h appears enhanced by a factor 2-3 relative to Jupiter 72/76

73 Saturn Like Jupiter, the magnetic field is generated in a metallic hydrogen region The field is weaker than Jupiter s due to a metallic envelope of smaller extent Internal excess heat is generated about equally by Gravitational contraction/accretion ti ti Release of gravitational energy due to He rainout onto the core 73/76

74 Uranus and Neptune Neptune e is 3% smaller and 15% more massive than Uranus Models much less constrained than J&S H and He constitute only a few Earth masses of material High-Z element mass is similar to Jupiter and Saturn Consistent with atmospheric C and S abundances enhanced about 50 times relative to Jupiter and Saturn Rock core Possibly about 1 Earth mass, but may be absent Possibly differentiated: iron in solid phase, rock may be liquidid 74/76

75 Uranus and Neptune Mantle Constitutes 80% of the mass Icy composition: hot, dense, ionic oceans of H 2 O and minor NH 3, CH 4, N 2, H 2 S Atmosphere Outer 5-15% of the radius H- and He-rich Internal magnetic fields indicate electrically conductive and convective icy interiors (P too low for metallic H) generated at radii from the core 75/76

76 Interior structure of the giant planets 76