EARTH AS AN EVOLVING PLANETARY SYSTEM Part 9 ppsx

before the terminal large impact event about 3.9 Ga. Continued fracturing and volcanism on Mars extended to at least 1000 Ma and perhaps 100 Ma. Venus Comparison with the Earth. Unlike the other terrestrial planets, Venus is similar to the Earth both in size and mean density (5.24 and 5.52 g/cm 3 , respectively) (Table 10.1). After correcting for pressure differences, the uncompressed density of Venus is within 1% of that of the Earth, indicating that both planets are similar in composition, with Venus having a somewhat smaller core/mantle ratio. Although both planets have similar amounts of N 2 and CO 2 , most of the Earth’s CO 2 is not in the atmosphere but in carbon- ates. Venus also differs from the Earth by the near absence of water and the high density and temperature of its atmosphere. As described later, Venus may at one time have lost massive amounts of water by the loss of hydrogen from the upper atmosphere. Unlike the Earth, Venus lacks a satellite, has a slow retrograde rotation (244 Earth days for one rota- tion), and does not have a measurable magnetic field. Because Venus orbits the Sun in only 225 days, the day on Venus (244 Earth days) is longer than the year. The absence of a magnetic field in Venus may be caused by the absence of a solid inner core because, as described in Chapter 5, crystallization of an inner core may be required for a dynamo to operate in the outer core of a planet. Of the total Venusian surface, 84% is flat rolling plains, some of which are more than 1 km above the average plain elevation. Only 8% of 362 Comparative Planetary Evolution N. Kasei Ayres Gibraltar Bosporus Missoula Amazon Mississippi 10 km 0 0 500 m 5 Altai Figure 10.4 Comparison of channel cross-sections for cata- clysmic flood channels on Mars (upper two) with straits (Gibraltar and Bosporus) and river chan- nels on the Earth. Modified from Baker (2001). the surface is true highlands; the remainder (16%) lies below the average radius, forming broad, shallow basins. This is unlike the topographic distribution on the Earth, which is bimodal because of plate tectonics (Fig. 10.5). The unimodal distribution of elevation on Venus does not support the existence of plate tectonics on Venus today. The spectacular Magellan imagery indicates that unlike the Earth, deformation on Venus is distributed over thousands of kilometers rather than occurring in narrow orogenic belts (Solomon et al., 1992). There are numerous examples of compressional tectonic features on Venus, such as Maxwell Montes deformational belt in the western part of Ishtar Terra (Fig. 10.6). Ishtar Terra is a highland about 3 km above the mean plan- etary radius surrounded by compressional features suggestive of tectonic convergence resulting in crustal thickening. Maxwell Montes stands 11 km above the surrounding plains and shows a wrinkle-like pattern suggestive of compressional deformation (Kreep and Hansen, 1994). The deformation in this belt appears to have occurred passively in response to horizontal stresses from below. Coronae are large circular features (60–2600 km in diameter with most 100–300 km) with a great diversity of morphologies (Stefan et al., 2001). Almost all coronae occur between 80° N and 80° S latitude and show a high concentration in equatorial areas. Venus is the only planet known to have coronae. An approximately inverse correlation between crater and corona density suggests that the volcano–tectonic process that forms coronae may be the same process that destroys craters (Stefanick and Jurdy, 1996). The most widely accepted models for the origin of coronae are those involving mantle plumes. A rising plume creates a region of uplift accompanied by radial deformation and dyke emplacement (Copp et al., 1998). Volcanism may also accompany this stage. As the plume head spreads at the base of the lithosphere, it elevates the surface, producing annuli in some coronae. This is generally followed by collapse as the plume head cools. Members of the Solar System 363 Venus Earth Surface area per kilometer interval (%) 65 60 55 50 45 40 35 30 25 20 15 10 0 −6 −4 −2 Altitude (km) 0246 5 Figure 10.5 Comparison of relief on Venus and the Earth. Surface height is plotted in 1-km intervals as a function of surface area. Height is measured from the sphere of average planetary radius for Venus and from the sea level for the Earth. Modified from Pettengill et al. (1980). Another unique and peculiar feature of the Venusian surface is the closely packed sets of grooves and ridges known as tesserae, which appear to result from compression. A combination of structural, mechanical, topographic, and geologic evidence suggests that tesserae record interaction of deep mantle plumes with an ancient, globally thin litho- sphere, resulting in regions of thickened crust (Hansen et al., 1999). Perhaps the most important data from the Magellan mission are those related to impact craters (Kaula, 1995). Unlike the Moon, Mars, and Mercury, Venus does not preserve a record of heavy bombardment from the early history of the solar system (Price and Suppe, 1994). Crater size–age distribution shows an average age of the Venusian surface of only 600 to 400 Ma, indicating extensive resurfacing of the planetary surface at this time. Most of this resurfacing is with low-viscosity lavas, presumably mostly basalts as inferred from the Venera geochemical data. Crater distribution also indicates a rapid decline in the resurfacing rate within the last tens of millions of years. However, results suggest that some large volcanoes (72 Ma), some basalt flows (128 Ma), some rifts (130 Ma), and 364 Comparative Planetary Evolution Figure 10.6 Magellan image of Maxwell Montes, the highest mountain range on Venus, which stands 11 km above the average diameter of the planet. The complex pattern of intersecting ridges and valleys reflects intense folding and shearing of the crust. Courtesy ofU. S. Geological Survey. many coronae (120 Ma), are much younger than the average age the resurfaced plains and probably represent ongoing volcanic and tectonic activity (Price et al., 1996). The differences between Venus and the Earth, with the lower bulk density of Venus, affect the nature and rates of surface processes (weathering, erosion, and deposition), tectonic processes, and volcanic processes. Because a planet’s thermal and tectonic history depends on its size and the area/mass ratio as described later, Venus and the Earth are expected to have similar histories. However, the surface features of Venus are quite different from those of the Earth, raising questions about how Venus transfers heat to the surface and whether plate tectonics has ever been active. The chief differences between the Earth and Venus appear to have two underlying causes: (1) small differences in plan- etary mass leading to different cooling, degassing, and tectonic histories, and (2) differ- ences in distance from the Sun, resulting in different atmospheric histories. Surface Composition. Much has been learned about the surface of Venus from scientific missions by the United States and Russia. The Russian Venera landings on the Venusian surface have provided a large amount of data on the structure and composition of the crust. Results suggest that most of the Venusian surface is composed of blocky bedrock surfaces and that less than one-fourth contains porous, soil-like material (McGill et al., 1983). The Venera Landers have also revealed the presence of abundant volcanic features, complex tectonic deformation, and unusual ovoid features of probable volcanic–tectonic origin. Reflectance studies of the Venusian surface suggest that iron oxides may be important components. Partial chemical analyses made by the Venera Landers indicate that basalt is the most important rock type. The high K 2 O recorded by Venera 8 and 13 is suggestive of alkali basalt, and the results from the other Venera landings clearly indicate tholeiitic basalt, perhaps with geochemical affinities to terrestrial ocean-ridge tholeiites (Fig. 10.2). A Venusian crust composed chiefly of basalt is consistent with the presence of thousands of small shield volcanoes that occur on the volcanic plains, typi- cally 1 to 10 km in diameter and with slopes of about 5 degrees. The size and distribu- tion of these volcanoes resembles terrestrial oceanic-island and seamount volcanoes. Venusian Core. Venus has no global magnetic field, although it likely has a molten outer core with or without an inner core (Stevenson, 2003). The absence of a dynamo in the outer core probably reflects the lack of convection caused either because an inner core is absent or because the outer core is not cooling. If the inside of Venus is hotter than the corresponding depth in the Earth, which seems likely, an inner core is not expected. Alternatively, or in addition, the Venusian core may not be cooling at present because it is still recovering from heat loss associated with a resurfacing event some 500 Ma. Cooling and Tectonics. To understand the tectonic and volcanic processes on Venus, it is first necessary to understand how heat is lost from the mantle. Four sources of information are important in this regard: the amount of 40 Ar in the Venusian atmosphere, lithosphere thickness, topography, and gravity anomalies. The amount of 40 Ar in planetary atmospheres can be used as a rough index of past tectonic and volcanic activity because Members of the Solar System 365 it is produced in planetary interiors by radioactive decay and requires tectonic–volcanic processes to escape. Venus has about one-third as much 40 Ar in its atmosphere as does the Earth, which implies less tectonic and volcanic activity for comparable 40 K contents. In contrast to the Earth, where at least 90% of the heat is lost by the production and subduction of oceanic lithosphere, there is no evidence for plate tectonics on Venus. The difficulty of initiating and sustaining subduction on Venus is probably because of a combination of high mantle viscosity; high fault strength; and thick, relatively buoyant basaltic crust. Thus, it would appear that Venus, like the Moon, Mercury, and Mars, must lose its heat through conduction from the lithosphere, perhaps transmitted upward chiefly by mantle plumes. The base of the thermal lithosphere in terrestrial ocean basins is about 150 km deep, where the average geotherm intersects the wet mantle solidus. On Venus, however, where the mantle is likely dry, an average geotherm does not intersect the dry solidus, indicating the absence of a distinct boundary between the lithosphere and the mantle (Fig. 10.7). The base of the elastic lithosphere in ocean basins is at the 500° C isotherm, or about 50 km deep. Because 500° C is near the average surface temperature of Venus, there is no elastic lithosphere on Venus. Another important difference between Venus and Earth is the strong positive correlation between gravity and topography on Venus, implying compensation depths in the Venusian mantle of 100 to 1000 km. This requires strong coupling of the mantle and lithosphere, and hence the absence of an asthenosphere, agreeing with the thermal arguments presented previously. This situation may have arisen from a lack of water in Venus. One of the important consequences of a stiff mantle is the inability to recycle lithosphere through the mantle, again showing that plate tectonics cannot occur on Venus. The deformed plateaus and the lack of features characteristic of brittle deformation, such as long faults, suggest the Venusian lithosphere behaves more like a viscous fluid than a brittle solid. The steep-sided high-elevation plateaus on Venus, however, attest to the strength of the Venusian lithosphere. 366 Comparative Planetary Evolution Dry Mantle Solidus Wet Mantle Solidus OCEAN BASIN VENUS Base of Elastic Lithosphere 150 100 50 Depth (km) Temperature (°C) 0 0 400 800 1200 1600 Figure 10.7 Comparison of geotherms from an average terrestrial ocean basin and Venus. Conduction is assumed to be the only mode of litho- spheric heat transfer on Venus. Also shown are wet (0.1% water) and dry mantle solidi. Two thermal-tectonic models have been proposed for Venus: the conduction and the mantle-plume models (Bindschadler et al., 1992). In the conduction model, Venus loses heat by simple conduction through the lithosphere, and tectonics is a result of compres- sion and tension in the lithosphere in response to the changing thermal state of the planet. It is likely that such a model describes the Moon, Mercury, and Mars at present. Not favoring a conduction model for Venus, however, is the implication that the topography is young because it cannot be supported for long with warm, thin lithosphere. In the mantle-plume model, which is preferred by most investigators, Venus is assumed to lose heat from large mantle plumes coupled with delamination and sinking of the lith- osphere (Turcotte, 1995; Turcotte, 1996) (Fig. 10.8). Also consistent with plumes are the deep levels of isostatic gravity compensation beneath large topographic features, suggest- ing the existence of plumes beneath these features. Unlike the Earth, most of the topo- graphic and structural features on Venus can be accounted for by mantle plumes, with compressional forces over mantle downwellings responsible for the compressional fea- tures on the surface. On the whole, the geophysical observations from Venus support the idea that mantle downwelling is the dominant driving force for deformation of the surface of Venus. The return flow in the mantle would also occur in downwellings and would undoubtedly involve delamination and sinking of significant volumes of the lithosphere. This is a striking contrast to the way the Earth cools, as shown in Figure 10.8. Although the age of the Venusian surface is likely variable, studies of crater distribu- tions indicate that more than 80% of the surface had all of its craters removed in a short period, probably between 10 and 100 My (Nimmo and McKenzie, 1998). It is still debated Members of the Solar System 367 Lithosphere Lithosphere Subduction 90% Delamination 70% Plumes 10% Plumes 30% EARTH VENUS Figure 10.8 Comparison of the Earth’s cooling mechanisms with possible cooling mecha- nisms on Venus. Estimates of magnitudes given in per- centages. Modified from Turcotte (1995). whether this resurfacing of the planet from 600 to 400 Ma was caused by a catastrophic planetwide mantle-plume event, a short-lived plate tectonic event in the waning stages today, or simply ongoing resurfacing whose mean age is 600 to 400 Ma (Strom et al., 1994; Nimmo and McKenzie, 1998). Giant Planets Jupiter and Saturn, the two largest planets, have densities indicating that they are com- posed chiefly of hydrogen and helium (Table 10.1). In the outer parts of the planets, these elements occur as ices and gases and at greater depths as fluids. The cores of the giant planets include a mixture of high-density ices and silicates. Relative to the Sun, the giant planets are enriched in elements heavier than He. Magnetic fields of these planets vary significantly in orientation or magnitude, and the origin of these fields is poorly understood. They are not, however, produced by dynamo action in a liquid Fe core, as is the case in the terrestrial planets. Unlike Jupiter and Saturn, the densities of Uranus and Neptune require a greater silicate fraction in their interiors. Models for Uranus, for instance, suggest a silicate core and icy inner mantle composed chiefly of water, CH 4 , and NH 3 and a gaseous and icy outer mantle composed chiefly of H 2 and He. Neptune must have an even greater proportion of silicate and ice. Except for Jupiter, with a 3-degree inclination to the ecliptic, the outer planets are highly tilted in their orbits (Saturn 26.7 degrees, Uranus 98 degrees, and Neptune 29 degrees). Such large tilts prob- ably result from collisions with other planets early in the history of the solar system. Whatever hit Uranus to knock it completely over must have had a mass similar to that of the Earth. Satellites and Planetary Rings General Features There are about 60 satellites in the solar system. Although there is great diversity in the satellites and no two are alike, three general classes of planetary satellites are recognized: 1. Regular satellites, which include most of the larger satellites and many of the smaller satellites, are those that revolve in or near the plane of the planetary equator and revolve in the direction in which the parent planet moves about the Sun. 2. Irregular satellites have highly inclined, often retrograde, and eccentric orbits, and many are far from the planet. Many of Jupiter’s satellites belong to this category as do the outermost satellites of Saturn and Neptune (Phoebe and Nereid, respec- tively). Most, if not all, of these satellites were captured by the parent planet. 3. Collisional shards are small, often irregular-shaped satellites that appear to have been continually eroded by ongoing collisions with smaller bodies. Many of the satellites of Saturn and Uranus are of this type. Phobos and Deimos, the tiny satel- lites of Mars, may be captured asteroids. 368 Comparative Planetary Evolution There are regularities in satellite systems that are important in constraining satellite origin. For instance, the large, regular satellites of Jupiter, Saturn, and Uranus have low inclination, prograde orbits indicative of formation from an equatorial disk (Stevenson, 1986). Although regular satellites extend to 20 to 50 planet radii, they do not form a scale model of the solar system. Although the large satellites are mostly rocky or rock–ice mix- tures, small satellites tend to be more ice rich, suggesting that some of the larger satel- lites may have lost ice or accreted rock. Volatile ices, such as CH 4 and N 2 , appear only on satellites distant from both the Sun and the parent planet, reflecting the cold temper- atures necessary for their formation. One thing that emerges from an attempt to classify satellites is that no general theory of satellite formation is possible. Planetary Rings Since the Voyager photos of planetary rings in the outer planets, the origin of planetary rings has taken on new significance. Some have suggested that the rings of Saturn can be used as an analogue for the solar nebula from which the solar system formed. Although Jupiter, Saturn, Uranus, and Neptune are now all known to have ring systems, they are all different, and no common theory can explain all of them. Although the rings of Saturn are large in diameter, the thickness of the rings is probably less than 50 m. The average particle size in the rings is only a few meters, and single particles orbit the planet in about 1 day. Three models have attracted most attention for the origin of planetary rings. In the first two models, rings are formed with the parent planet as remnants of an accretionary disk or of broken pieces of satellites. Neither of these origins is likely, however, because rings formed in such a manner should not have survived beyond a few million years. Alternatively, the rings may be debris from the disruption of captured comets such as Chiron. In this model, the small particles become rings and the larger fragments may become satellites. If the rings around the giant planets are the remains of captured comets, they are latecomers to the solar system because, as you shall see later in this chapter, comets are among the youngest members of the solar system. The Moon As a planetary satellite, there are many unique features about the Moon. Among the more important are the following, all of which must be accommodated by any acceptable model for lunar origin: 1. The orbit of the Moon about the Earth is neither in the equatorial plane of the Earth nor in the ecliptic; it is inclined 6.7 degrees to the ecliptic (Fig. 10.9). 2. Except for the Pluto–Charon pair, the Moon has the largest mass of any satellite– planetary system. 3. The Moon has a low density compared with that of the terrestrial planets, implying a relatively low iron content. 4. The Moon is strongly depleted in volatile elements and enriched in some refractory elements such as Ti, Al, and U. Members of the Solar System 369 5. The angular momentum of the Earth–Moon system is anomalously high compared to other planet–satellite systems. 6. The Moon rotates in the same direction as does the Earth. A great deal has been learned about the geochemistry and geophysics of the Moon from the Apollo landings (Taylor, 1982; Taylor, 1992). Although average lunar density is much less than the average Earth density (Table 10.1), its uncompressed density is about the same as the Earth’s mantle, implying that the Moon is composed largely of Fe and Mg silicates. Unlike most other satellites, which are mixtures of silicates and water ice, the Moon must have formed in the inner part of the solar system. From seismometers placed on the Moon by astronauts, scientists can deduce the broad structure of the lunar interior. The Moon has a thick crust (60–100 km) comprising about 12% of the lunar volume, and it appears to have formed soon after planetary accretion, about 4.5 to 4.4 Ga (Taylor, 1992). From the limited sampling of the lunar crust by the astronauts, scientists have learned that it is composed chiefly of anorthosites and gabbroic anorthosites as represented by exposures in the lunar highlands. These rocks typically have cumulus igneous textures, although they have been modified by impact brecciation. Sm-Nd isotopic dating indicates that this plagioclase-rich crust formed about 4.45 Ga. As you shall see later, it appears to have formed by crystallization of an extensive magma ocean. Also characteristic of the lunar surface are the mare basins, large impact basins formed before 3.9 Ga, covering about 17% of the lunar surface (Fig. 10.10). These basins are flooded with basalt flows only 1 to 2 km thick and probably erupted chiefly from fissures. Isotopically dated mare basalts range from 3.9 to 2.5 Ga. The impacts that formed the mare basins did not initiate the melting that produced the basalts, which were erupted up to hundreds of millions of years later and thus represent a secondary crust on the Moon. The youngest basaltic eruptions may be as young as 1 Ga. The lunar crust overlies a mantle composed of two layers. The upper layer or lithosphere extends to a depth of 400 to 500 km and is probably composed of cumulate ultramafic rocks. The second layer extends to about 1100 km, where a sharp break in seismic velocity occurs. Although evidence is still not definitive, it appears that the Moon has a small metallic core (300–500 km in diameter), comprising 2 to 5% of the lunar volume (Fig. 10.1). Although the Moon does not have a magnetic field, remnant magnetization in lunar rocks suggests a lunarwide magnetic field at least between 3.9 and 3.6 Ga (Fuller and Cisowski, 1987). The maximum strength of this field was probably only about one-half that of the present Earth’s field. It is likely this field was generated by fluid motions in the lunar core, much like the present Earth’s field is produced. A steady decrease in the 370 Comparative Planetary Evolution Moon Earth ECLIPTIC E q uator 6.7 ° 23.4 ° Figure 10.9 Orbital rela- tions of the Earth–Moon system. magnetic field after 3.9 Ga reflects cooling and complete solidification of the lunar core by no later than 3 Ga. The most popular model for lunar evolution involves the production of an ultramafic magma ocean that covered the entire Moon to a depth of 500 km or more and crystallized Members of the Solar System 371 Figure 10.10 Oblique view of the southern part of the Imbrium basin, one of the large mare basins on the Moon. Courtesy of the Lunar and Planetary Institute. [...]... wobble can be measured by tracking its position Planets can also be detected as they pass in front of the star The most unexpected aspect of the new planets is their peculiar properties (Taylor, 199 9) Most extrasolar planets have masses of 0.75 to 3 Jovian masses, but masses less than 0.5 of a Jovian mass are difficult to detect Unlike planets in the solar system, most extrasolar planets are much closer... planet is hot and viscosity is low, chaotic mantle convection rapidly cools the planet and crystallizes magma oceans (Drake, 2000) As mantle viscosity increases, beginning 100 My after accretion, convection should cool planets at reduced rates Accretion of the Earth With astrophysical models, the chemical composition of planets provides an important constraint on planetary accretion (Greenberg, 198 9;... of planetary masses? Was there ever a single, small planet in the asteroid belt, and if not, why? Cooling rate data from iron meteorites that come from the asteroids, as well as an estimate of the tidal forces of Jupiter, indicate that a single planet never existed in the asteroid belt The tidal forces of Jupiter would fragment the planet before it grew to planetary size Hence, it appears that the asteroids... the solar system accreted This means that there was not widespread mixing of the material in the inner solar system 396 Comparative Planetary Evolution during accretion and that planets growing in the inner solar system received most of their masses from narrow rings around the Sun (Drake and Righter, 2002) Extrasolar Planets As the ability of scientists to detect planets around stars increases with... the solar system would appear to extend from about 0 .95 to 1.5 AU (Kasting et al., 199 3b) This observation is important in that the continuously habitable zone is wide enough that habitable planets may exist in other planetary systems Although in any planetary system most life, and all higher forms of life, would be limited to the continuously habitable zone, some microbes that can withstand extreme... and accrete into the planets (Fig 10.21) Small planetesimals form within the cloud and spiral toward the ecliptic plane, where they begin to collide with one another and grow into planetary embryos,” which range in size from that of the Moon to that of Mercury It is these embryos that collide and grow into the terrestrial planets and the silicate–ice cores of the giant outer planets Most of the gas... mean atomic weight for the Earth of about 27 (mantle = 22.4 and core = 47.0) and show that it is composed chiefly of iron, silicon, magnesium, and oxygen When meteorite classes are mixed to give the correct core/mantle mass ratio (32:68) and mean atomic weight of the Earth, results indicate that iron and oxygen are the most abundant elements followed by silicon and magnesium (Table 10.3) Almost 94 %... the Sun in an orbit between Mars and Jupiter (Lebofsky et al., 198 9; Taylor, 199 2) Of the 10,000 or so known asteroids, most occur between 2 and 3 AU from the Sun (Fig 10.12) The total mass of the asteroid belt is only about 5% of that of the Moon Only a few large asteroids are recognized, the largest of which is Ceres with a diameter of 93 3 km Most asteroids are less than 100 km in diameter, and there... groups of asteroids are recognized: (1) the near -Earth asteroids (Apollo, Aten, and Amor classes), some of which have orbits that cross that of the Earth; (2) the main belt asteroids; and (3) the Trojans revolving in the orbit of Jupiter (Fig 10.12) Most meteorites arriving on the Earth are coming from the Apollo asteroids The orbital gaps in which no asteroids occur in the asteroid belt (for instance,... planets from volatile-poor planetesimals that were probably already differentiated into metallic cores and silicate mantles in a gas-free inner part of the nebula (Taylor, 199 9) During the time of supercollisions, the Earth was probably struck at a high velocity many times by bodies ranging up to Mercury in size and at least once by a Mars-size body Much of the kinetic energy associated with these collisions . in planetary atmospheres can be used as a rough index of past tectonic and volcanic activity because Members of the Solar System 365 it is produced in planetary interiors by radioactive decay and. members of the solar system. Asteroids Asteroids are small planetary bodies, most of which revolve about the Sun in an orbit between Mars and Jupiter (Lebofsky et al., 198 9; Taylor, 199 2). Of the 10,000. interpolation of planetary masses? Was there ever a single, small planet in the asteroid belt, and if not, why? Cooling rate data from iron meteorites that come from the asteroids, as well as an estimate