(lecture07)= # Planetary Surfaces: Geology, Geomorphology, & Geophysics ```{highlights} **Learning objectives:** By the end of this lecture, you will be able to describe the major surface processes shaping planetary bodies, derive the crater scaling law from dimensional analysis and apply it to estimate crater sizes, use crater counting to constrain surface ages, and compare tectonic, volcanic, and erosional styles across the solar system. ``` ```{seealso} **Slides:** [Download Lecture 7 (PDF)](../_static/slides/lecture07.pdf) ``` ## Surface processes A planet's surface is its geological record: the integrated result of billions of years of competing processes that create, modify, and destroy landforms. By reading this record, we can reconstruct a body's geological history even without visiting it in person {cite:p}`Melosh2011`. Surface processes fall into two broad categories: - **Endogenic processes** are driven by a body's internal heat ({ref}`Lecture 3 `): volcanism builds new terrain by erupting molten material, and tectonics deforms the crust through faulting and folding. Bodies with more internal heat (larger size, more radioactive elements, stronger tidal heating) show more vigorous endogenic activity. - **Exogenic processes** are driven by external forces: impact cratering reshapes surfaces through hypervelocity collisions, while erosion (by wind, water, ice, or radiation) gradually wears them down. These processes dominate on bodies with little internal heat. The relative importance of these processes varies dramatically across the solar system, depending on a body's size, composition, atmosphere, and distance from the Sun. {numref}`fig:surface-processes` uses Mars as a single-body anchor for the four-way taxonomy (volcanism, tectonics, impacts, erosion) developed in this lecture. | Body | Dominant surface process | Key evidence | |------|:------------------------:|:------------:| | Moon | Impact cratering | Saturated highland craters, preserved for $>$4 Gyr | | Earth | Plate tectonics + erosion | Youngest ocean floors $<$200 Myr; rapid weathering | | Mars | Volcanism + impacts | Olympus Mons, cratered southern highlands | | Venus | Volcanism | Volcanic plains cover $\sim$80% of surface | | Io | Tidal volcanism | $\sim$300–400 active volcanoes, surface age $<$1 Myr | | Europa | Ice tectonics | Lineae, chaos terrain, very few craters | ```{figure} figures/surface_processes_montage.avif :name: fig:surface-processes :width: 500px :align: center Viking Orbiter global colour mosaic of Mars (PIA00407), an orthographic projection of the Valles Marineris hemisphere centred near 20$^\circ$N, 60$^\circ$W. The image is used here as an anchor for the four-way taxonomy of surface processes. Endogenic processes are evident in Valles Marineris, the dark extensional canyon system running west-to-east across the lower middle of the disk, and in the Tharsis shield volcanoes near the western limb (visible as dark patches at left, foreshortened by the limb). Exogenic processes are recorded in the heavily cratered southern highlands and in the bright water-ice north polar cap (top). Mars preserves all four major process classes (volcanism, tectonics, impact cratering, erosion) on a single body and is therefore a useful template for the rest of this lecture. Credit: NASA/JPL/USGS, public domain. ``` ## Impact cratering Impact cratering is the most universal geological process in the solar system. Every solid body bears the scars of hypervelocity collisions, from Mercury's heavily cratered surface to Pluto's ancient terrains. On bodies without atmospheres or active geology (like the Moon), impact craters are preserved for billions of years, making them invaluable tools for dating surfaces {cite:p}`Melosh2011`. ### Impact mechanics When a projectile strikes a planetary surface at typical solar system velocities of 10–70 km s$^{-1}$ (set by orbital mechanics; {ref}`Lecture 2 `), the collision releases enormous energy in a fraction of a second. The impact process unfolds in three stages: 1. **Contact and compression:** The projectile makes contact and a shock wave propagates into both the target and the impactor, compressing them to pressures of $\sim$100 GPa (comparable to pressures at Earth's core-mantle boundary). The shock wave converts kinetic energy into internal energy, vaporising and melting material near the impact point. 2. **Excavation:** The shock wave expands hemispherically outward and reflects off the free surface, generating a **rarefaction wave** (a decompression pulse that follows the initial shock) that accelerates material upward and outward. A transient cavity forms, with a depth-to-diameter ratio of roughly 1:3. Material is ejected in a cone-shaped curtain, creating the **ejecta blanket**, the layer of debris draped over the terrain surrounding the crater. 3. **Modification:** The transient cavity is gravitationally unstable. In small craters, the walls simply slump inward, producing a bowl-shaped **simple crater**. In larger craters, the floor rebounds upward (forming a **central peak**) and the walls collapse in terraces, producing a **complex crater**. The kinetic energy of an impactor is: $$ E_k = \frac{1}{2}mv^2 $$ For a 1 km diameter rocky asteroid ($\rho \approx 3000$ kg m$^{-3}$, $m \approx 1.6 \times 10^{12}$ kg) hitting at 20 km s$^{-1}$: $$ E_k = \frac{1}{2} \times 1.6 \times 10^{12} \times (2 \times 10^4)^2 \approx 3 \times 10^{20} \text{ J} $$ This is roughly $10^3$ times the energy of the largest nuclear weapon ever detonated (Tsar Bomba, $\sim$50 Mt $\approx 2 \times 10^{17}$ J), released in less than a second and concentrated at a single point. {numref}`fig:impact-stages` sketches the three stages in cross-section. ```{figure} figures/impact_stages.avif :name: fig:impact-stages :width: 100% :align: center The three stages of impact crater formation in schematic cross-section. (a) Contact and compression: the projectile arrives at $10$ to $70$ km s$^{-1}$, and shock waves (dashed red) propagate into both the target and the projectile, compressing the material to pressures of order $100$ GPa and vaporising and melting the contact zone (orange). (b) Excavation: the shock front expands hemispherically, and the rarefaction wave behind it (dotted blue) accelerates material up and out along ballistic paths, opening a transient cavity with a depth of about one third of its diameter and depositing the ejecta blanket around the rim. (c) Modification: in a small crater the walls slump inward and leave a simple bowl; in a large crater the floor rebounds into a central peak and the walls collapse in terraces around a flat floor, producing a complex crater. Not to scale. Course-original figure. ``` ## Blackboard derivation: Crater scaling law ````{admonition} Blackboard derivation: Crater scaling law from dimensional analysis :class: tip **Goal:** Use dimensional analysis to derive how crater diameter $D$ depends on impact energy $E$, target density $\rho$, and surface gravity $g$, then apply the result to estimate the crater produced by a 1 km asteroid impact on the Moon. **Setup.** We want to find the diameter $D$ of the crater (a length) produced by an impact with kinetic energy $E$ into a target with density $\rho$ under surface gravity $g$. In the **gravity regime** (where crater size is limited by gravity rather than material strength, valid for craters larger than $\sim$100 m), dimensional analysis requires: $$ D = C \, E^a \, \rho^b \, g^c $$ where $C$ is a dimensionless constant of order unity and $a$, $b$, $c$ are exponents to be determined. {numref}`fig:board-sketch-l07` shows the quantities that enter the scaling: the impact energy $E$ arrives from above, the target has density $\rho$ and gravity $g$, and the crater has diameter $D$. ```{figure} figures/board_sketch_l07.avif :name: fig:board-sketch-l07 :width: 45% :align: center Setup sketch of the board derivation of the crater scaling law: an impactor with kinetic energy $E$ strikes a target of density $\rho$ under surface gravity $g$ and excavates a bowl of diameter $D$. Course-original figure. ``` **Dimensional analysis.** Writing the dimensions of each quantity in terms of mass $M$, length $L$, and time $T$: - $[D] = L$ - $[E] = M L^2 T^{-2}$ - $[\rho] = M L^{-3}$ - $[g] = L T^{-2}$ Requiring $[E^a \rho^b g^c] = L$: $$ [E^a \rho^b g^c] = (M L^2 T^{-2})^a \, (M L^{-3})^b \, (L T^{-2})^c = M^{a+b} \, L^{2a-3b+c} \, T^{-2a-2c} $$ This must equal $[D] = M^0 L^1 T^0$, so each exponent must match: **Mass:** $a + b = 0 \implies b = -a$ **Time:** $-2a - 2c = 0 \implies c = -a$ **Length:** $2a - 3b + c = 1$ Substituting $b = -a$ and $c = -a$: $$ 2a - 3(-a) + (-a) = 1 \implies 2a + 3a - a = 1 \implies 4a = 1 \implies a = \frac{1}{4} $$ Therefore $a = 1/4$, $b = -1/4$, $c = -1/4$, and: $$ \boxed{D \sim \left(\frac{E}{\rho g}\right)^{1/4}} $$ (eq:crater-scaling) **Dimensional check:** $[E/\rho g] = [M L^2 T^{-2}/(M L^{-3} \cdot L T^{-2})] = [L^4]$. Taking the fourth root gives $[L]$, the dimensions of length. $\checkmark$ This is the **crater scaling law** in the gravity regime {cite:p}`Holsapple1993`. It tells us that crater diameter scales as the fourth root of impact energy: doubling the energy increases the crater diameter by only a factor of $2^{1/4} \approx 1.19$ (about 19%). This weak dependence on energy explains why craters have a relatively narrow size range even though impactor energies span many orders of magnitude. **Worked example: 1 km asteroid on the Moon.** For our 1 km asteroid ($E \approx 3 \times 10^{20}$ J) impacting the Moon's regolith, the loose, fragmented layer of surface debris ($\rho \approx 2500$ kg m$^{-3}$, $g = 1.62$ m s$^{-2}$): $$ D \sim \left(\frac{3 \times 10^{20}}{2500 \times 1.62}\right)^{1/4} = \left(7.4 \times 10^{16}\right)^{1/4} \approx 1.65 \times 10^{4} \text{ m} \approx 16.5 \text{ km} $$ This is consistent with the observed sizes of lunar craters formed by $\sim$1 km impactors. The same law also fixes how crater size grows with impactor size: at fixed impact velocity and target density, $E \propto L^3$ for an impactor of diameter $L$, so $D \propto L^{3/4}$. Applied to the 85 km crater Tycho, this gives an impactor $(85/16.5)^{4/3} \approx 9$ times wider than our 1 km asteroid, so roughly 9 km across. Read that as an order of magnitude, not a measurement: the law describes the transient cavity in the gravity regime, whereas Tycho is a complex crater whose final rim was widened by collapse, and the dimensionless prefactor $C$ is not fixed by dimensional analysis (it depends on target material properties and is of order unity for rocky surfaces). The more complete **pi-scaling framework** of {cite:p}`Holsapple1993` parameterises the transition between the gravity regime and the strength regime (where material cohesion, not gravity, limits crater growth) and accounts for target porosity and impactor properties. The two regimes are illustrated in {numref}`fig:holsapple-piscaling`. ```` ```{figure} figures/holsapple1993_strength_gravity_regimes.avif :name: fig:holsapple-piscaling :width: 700px :align: center The regimes of cratering for a material with strength: cratering efficiency $\pi_V = \rho V/m$ as a function of gravity-scaled size $\pi_2 = g a / U^2$. The three curves correspond to three impact velocities ($U = 2.5$, 10, and 40 km s$^{-1}$). In the **strength regime** (small craters, left), $\pi_V$ depends on the impact velocity $U$ but is essentially independent of $\pi_2$, so the curves separate. For increasing size at fixed velocity, the system transitions to the **gravity regime** (large craters, right) where $\pi_V \propto \pi_2^{-\alpha}$ and the curves converge to a common law. Most laboratory experiments in geological materials are necessarily in the strength regime; planetary-scale craters are firmly in the gravity regime. Reproduced from {cite:p}`Holsapple1993`, Fig. 3. ``` ## Crater morphology and chronology ### Crater morphology Craters come in three morphological classes, determined primarily by their diameter relative to the **transition diameter** $D_t$ {cite:p}`Melosh2011`: - **Simple craters** ($D \lesssim D_t$): Bowl-shaped depressions with smooth walls and a depth-to-diameter ratio of $\sim$1:5. On the Moon, simple craters have $D \lesssim 15$ km. - **Complex craters** ($D_t \lesssim D \lesssim 300$ km on the Moon): Feature central peaks (formed by rebound of the crater floor), terraced walls (from gravitational collapse), and flat floors. The transition from simple to complex occurs because gravity overcomes the strength of the crater walls. - **Multi-ring basins** ($D \gtrsim 300$ km): The largest impacts produce concentric ring structures. Examples include the Orientale basin on the Moon (930 km), Caloris on Mercury (1550 km, {numref}`fig:caloris-basin`), and the Hellas basin on Mars (2300 km). ```{figure} figures/mercury_caloris_basin.avif :name: fig:caloris-basin :width: 450px :align: center The Caloris basin on Mercury, $\sim$1550 km in diameter, imaged in enhanced colour by NASA's MESSENGER spacecraft. Caloris is one of the largest and best-preserved multi-ring impact basins in the solar system; the orange interior plains are smooth volcanic deposits emplaced after the impact, while the surrounding annulus shows ejecta and concentric ring structures. The basin's antipode on Mercury contains chaotic "weird terrain" thought to have formed from the focused seismic shock of the same event. Credit: NASA/Johns Hopkins APL/Carnegie Institution of Washington, public domain. ``` The transition diameter scales inversely with surface gravity, $D_t \propto 1/g$. The Moon only fixes the constant, because its transition diameter is well measured: $$ D_t \approx D_{t,\text{Moon}} \cdot \frac{g_{\text{Moon}}}{g} $$ (eq:transition-diameter) where $D_{t,\text{Moon}} \approx 15$ km and $g_{\text{Moon}} = 1.62$ m s$^{-2}$ {cite:p}`Melosh2011`. On Earth ($g = 9.81$ m s$^{-2}$), this gravity-only scaling gives $D_t \approx 15 \times 1.62/9.81 \approx 2.5$ km; target rock strength shifts the transition between $\sim 2$ km in crystalline rock and $\sim 4$ km in sedimentary targets, but virtually all terrestrial craters larger than a few kilometres are complex. The central peak that gives a complex crater its morphological signature is shown in {numref}`fig:crater-morphology`. ```{figure} figures/crater_morphology.avif :name: fig:crater-morphology :width: 600px :align: center Oblique view of the central peak of Tycho Crater (~85 km diameter), imaged by NASA's Lunar Reconnaissance Orbiter Camera (LROC) in low Sun illumination. The prominent ~2 km tall central peak in the foreground formed by gravitational rebound of the crater floor immediately after the impact; the rough terrain behind it is the hummocky interior of the crater, and the low Sun casts the peak's shadow across it. The transition from simple bowl-shaped craters to complex craters with central peaks and terraced walls occurs near $D_t \approx 15$ km on the Moon and scales inversely with surface gravity (Eq. {eq}`eq:transition-diameter`). Credit: NASA/GSFC/Arizona State University, public domain. ``` ### Crater counting and surface ages The key principle of **crater chronology** is simple: older surfaces have had more time to accumulate impact craters, so they have higher crater densities. By counting craters as a function of diameter, we can determine relative ages, and with calibration, absolute ages {cite:p}`Neukum2001`. The **crater size-frequency distribution (SFD)** follows a power law: $$ N(>D) = a \, D^{-b} $$ (eq:crater-sfd) where $N(>D)$ is the cumulative number of craters per unit area with diameter greater than $D$, and $a$ and $b$ are constants. For the production population (before saturation), $b \approx 2$ to 3, depending on the size range and the impactor population. The **lunar chronology** provides the critical calibration. Samples returned by the Apollo and Luna missions give radiometric ages for specific lunar surfaces. By counting craters on those same surfaces and measuring their crater density, we establish the relationship between crater density and absolute age. This calibrated lunar chronology can then be extrapolated (with corrections for different impact rates and gravity) to other solar system bodies. Key results from crater counting include: - The lunar highlands are **saturated** with craters (crater density has reached an equilibrium where new craters destroy old ones) and are $>$4 Gyr old. - The lunar maria have ages of 3.9–3.1 Gyr, confirmed by Apollo sample dating. - Mars's southern highlands are heavily cratered ($\sim$4 Gyr), while the northern lowlands are much younger. - Venus has a remarkably **uniform** crater density across its entire surface, implying a young mean surface age of a few hundred Myr and a global resurfacing episode (see {ref}`Lecture 9 `). The heavily cratered lunar farside highlands are shown in {numref}`fig:crater-counting`; the calibrated lunar-chronology curve relating cumulative crater density to surface age is shown in {numref}`fig:neukum-chronology`. ```{figure} figures/crater_counting.avif :name: fig:crater-counting :width: 450px :align: center The lunar farside as imaged by NASA's Lunar Reconnaissance Orbiter, showing a heavily cratered surface. Almost the whole hemisphere is ancient highland crust ($>$4 Gyr); the few dark patches are small farside maria, and the large nearside maria (3.1–3.9 Gyr) lie on the opposite hemisphere and are not visible here. The difference in crater density between the two hemispheres is what crater counting turns into relative and absolute surface ages. Forward references: crater chronology is applied to Mars in {ref}`Lecture 10 `. Credit: NASA/GSFC/Arizona State University, public domain. ``` ```{figure} figures/neukum_chronology.avif :name: fig:neukum-chronology :width: 600px :align: center The lunar crater chronology of {cite:t}`Neukum2001`, plotted from the published coefficients $a = 5.44 \times 10^{-14}$ km$^{-2}$, $\lambda = 6.93$ Gyr$^{-1}$, and $b = 8.38 \times 10^{-4}$ km$^{-2}$ Gyr$^{-1}$. The black curve gives the cumulative density of craters with $D \geq 1$ km expected on a surface of model age $T$. The function has two regimes: a roughly linear segment for $T \lesssim 3$ Gyr, where the linear term $bT$ describes a steady impact flux, and a steep upturn beyond $T \approx 3.6$ Gyr, where the exponential term overtakes the linear one and reflects the much higher impact rate during the early bombardment of the inner solar system. The red construction shows how the relation is used in practice: a measured crater density of $10^{-2}$ km$^{-2}$ inverts to a surface age of about 3.7 Gyr. The shaded intervals mark where the calibration is weakest: no returned sample dates a lunar unit between about 1 and 3 Gyr, and none dates a unit older than 3.92 Gyr, so the curve is an interpolation over the first interval and an extrapolation beyond the second {cite:p}`Robbins2014`. Ages on other bodies follow from the same inversion, with corrections for the local impact flux and gravity. Course-original figure. ``` ## Volcanism Volcanism is the primary mechanism by which a planet's internal heat reaches the surface ({ref}`Lecture 3 `). The style of volcanic activity (whether gentle lava flows or explosive eruptions) depends on the magma composition, volatile content, and the body's gravity and atmospheric pressure {cite:p}`Melosh2011`. ### Effusive vs. explosive volcanism The key variable is **magma viscosity**, which is controlled primarily by the $\mathrm{SiO_2}$ (silica) content: - **Low-viscosity (basaltic) magma** ($\sim$50% $\mathrm{SiO_2}$): Flows easily, producing broad, flat **shield volcanoes** and extensive **flood basalt** plains, broad sheets of low-viscosity lava erupted from fissures rather than a single central vent. Dissolved volatiles ($\mathrm{H_2O}$, $\mathrm{CO_2}$) escape gradually from the fluid magma, so eruptions are typically **effusive** (gentle lava flows). This is the dominant style on the Moon, Mars, and Io. - **High-viscosity (silicic) magma** ($>$65% $\mathrm{SiO_2}$): Traps dissolved volatiles until the pressure exceeds the magma's strength, producing violent **explosive** eruptions (e.g., Mount St. Helens, Krakatoa). These build steep **stratovolcanoes** (cone-shaped edifices built from alternating layers of lava and ash) and deposit widespread ash layers. Explosive volcanism requires both high-silica magma *and* significant volatile content, conditions met primarily on Earth. {numref}`fig:effusive-explosive` contrasts the two eruption styles. ```{figure} figures/effusive_explosive.avif :name: fig:effusive-explosive :width: 100% :align: center Effusive and explosive volcanism in schematic cross-section. (a) Low-viscosity basaltic magma (about 50% SiO$_2$) flows easily: dissolved H$_2$O and CO$_2$ escape gradually, eruptions are effusive, and the products are broad shield volcanoes and thin lava sheets fed from fissures, the dominant style on the Moon, Mars and Io. (b) High-viscosity silicic magma (more than 65% SiO$_2$) traps its volatiles under a viscous plug until the pressure exceeds the strength of the magma, and the eruption is explosive, as at Mount St. Helens and Krakatoa; repeated eruptions build steep stratovolcanoes from alternating layers of lava and ash. Course-original figure. ``` ### Volcanic landforms across the solar system **Olympus Mons (Mars)** is the largest volcano in the solar system: a shield volcano with a base diameter of $\sim$600 km and a summit elevation of $\sim$21.2 km above the Mars datum (2.4 times the height of Mount Everest above sea level). It grows so large because Mars lacks plate tectonics: the volcanic hotspot remains stationary beneath the **lithosphere**, the rigid outer shell of crust and uppermost mantle, for billions of years, piling up lava in one location. On Earth, plate motion carries the crust over the hotspot, creating chains of smaller volcanoes (e.g., the Hawaiian Islands) rather than a single massive edifice {cite:p}`dePaterLissauer2010`. Surface gravity sets a second, independent limit on how tall a volcano can grow: the maximum height a crust of a given strength can support scales inversely with $g$, and with $g_{\mathrm{Mars}} \approx 0.38\,g_\oplus$ the Martian crust can hold an edifice roughly 2.6 times taller than an equally strong crust on Earth. The two effects act together: the fixed plume supplies one edifice for billions of years, and the low gravity raises the ceiling that edifice can reach. **Lunar maria** are vast flood basalt plains that fill ancient impact basins on the Moon's nearside ({numref}`fig:lunar-maria`). Radiometric dating of Apollo samples shows they erupted between 3.9 and 3.1 Ga, during a period of residual internal heating. The maria cover $\sim$16% of the lunar surface but are visible from Earth as the dark patches that form the "face" of the Moon. ```{figure} figures/moon_nearside_annotated.avif :name: fig:lunar-maria :width: 500px :align: center The lunar nearside, dominated by the dark **maria**: flood-basalt plains that erupted into the great impact basins between 3.9 and 3.1 Ga and now form the pattern visible from Earth. The bright, heavily cratered ground is the older anorthositic highland crust; the smooth, low-albedo maria are younger basaltic lava that flooded Mare Imbrium, Serenitatis, Tranquillitatis, and the other nearside basins, covering $\sim$16% of the lunar surface. Their concentration on the nearside, and near-absence on the farside, reflects the thinner nearside crust that the basaltic melt could reach. Lunar Reconnaissance Orbiter Camera wide-angle mosaic (PIA14011). Credit: NASA/GSFC/Arizona State University, public domain; annotations added. ``` **Venus** and **Io** each get a lecture of their own. Lava plains cover $\sim$80% of Venus, and its uniform crater population points to global resurfacing in a short interval a few hundred Myr ago ({ref}`Lecture 9 `). Tidal heating from the Laplace resonance makes Io the most volcanically active body in the solar system, with $\sim$300–400 active centres and a mean surface age below 1 Myr ({numref}`fig:io-volcanism-surface`; {ref}`Lecture 11 `) {cite:p}`dePaterLissauer2010,Davies2024PSJ`. ```{figure} figures/io_volcanism.avif :name: fig:io-volcanism-surface :width: 500px :align: center Full-disk view of Jupiter's moon Io from NASA's Galileo spacecraft (PIA00583), revealing a surface dominated by sulfur and silicate volcanism. Yellow, white, and pale-green regions are sulfur and sulfur-dioxide deposits; darker patches mark recent silicate lava flows and active volcanic centres (a few of which have visible halos of fresh pyroclastic deposits). Credit: NASA/JPL-Caltech/University of Arizona, public domain. ``` | Body | Volcanic style | Driving mechanism | Example landforms | |------|:--------------:|:-----------------:|:-----------------:| | Earth | Effusive + explosive | Internal heat + plate recycling | Hawaii, Yellowstone, mid-ocean ridges | | Moon | Flood basalts (extinct) | Residual heat (3.9–3.1 Ga) | Maria (Imbrium, Serenitatis) | | Mars | Shield volcanoes (extinct?) | Mantle plumes, no plate motion | Olympus Mons, Tharsis Montes | | Venus | Lava plains + shield volcanoes | Internal heat, stagnant lid | Maat Mons, pancake domes | | Io | Ultramafic lava flows (active) | Tidal heating | Loki Patera, Pele | The size and morphology of Olympus Mons are shown in {numref}`fig:olympus-mons`, with a scale comparison to Earth's largest mountains in {numref}`fig:olympus-comparison`. ```{figure} figures/olympus_mons.avif :name: fig:olympus-mons :width: 450px :align: center Olympus Mons on Mars, the largest volcano in the solar system, seen from above in this Viking Orbiter colour mosaic. The shield volcano has a base diameter of $\sim$600 km and a summit elevation of $\sim$21.2 km above the Mars datum. The caldera complex at the summit (centre) contains multiple nested collapse craters formed by episodic magma withdrawal. The steep basal escarpment (up to 6 km high) is visible as the sharp boundary encircling the edifice. Olympus Mons grew to this immense size because Mars lacks plate tectonics: the volcanic source remained fixed beneath the lithosphere for billions of years. Credit: NASA/JPL/USGS, public domain. ``` ```{figure} figures/olympus_mons_size_comparison.avif :name: fig:olympus-comparison :width: 600px :align: center Height comparison of Olympus Mons (Mars), Mauna Kea (Hawaii) and Mount Everest (Earth). Each edifice rises from the base that its quoted height is measured against: the Mars datum for Olympus Mons at 21.2 km, the ocean floor for Mauna Kea at 10.2 km of total relief, and sea level for Mount Everest at 8.8 km. The three heights share one linear scale, so they can be read directly against each other; Olympus Mons is 2.4 times the height of Everest and 2.1 times the height of Mauna Kea. The drawn widths are schematic and strongly compressed, since Olympus Mons is $\sim$600 km across at its base, many times the width of either terrestrial mountain, so no single horizontal scale can show all three. The compression keeps Olympus Mons the flattest of the three profiles, a broad and gently sloping shield beside the steep Himalayan peak. The height contrast is set by the absence of plate tectonics on Mars, where a stationary mantle plume has continuously fed the same volcanic edifice for billions of years (whereas the Pacific plate carries Hawaiian volcanoes off the underlying hotspot in $\sim$10$^6$ yr), and by the lower Martian surface gravity, which lets a crust of given strength support an edifice roughly 2.6 times taller than on Earth. Course-original figure. ``` ## Tectonics Tectonics encompasses the large-scale deformation of a planet's crust and lithosphere, driven by forces arising from internal convection, thermal contraction, tidal stresses, and (on Earth) the motion of lithospheric plates. ### Plate tectonics: Earth's unique regime Earth is the **only body** in the solar system with active plate tectonics: the lithosphere is divided into $\sim$15 rigid plates that move at velocities of 1–10 cm yr$^{-1}$, driven by mantle convection. Three types of plate boundary exist: - **Divergent boundaries** (mid-ocean ridges): Plates move apart; new crust is created by seafloor spreading. - **Convergent boundaries** (subduction zones): One plate descends beneath another into the mantle; associated with volcanism and mountain building. - **Transform boundaries** (e.g., San Andreas Fault): Plates slide horizontally past each other. Plate tectonics enables the **carbonate-silicate cycle** ({ref}`Lecture 6 `) by recycling carbon through subduction and volcanic outgassing, a critical component of Earth's long-term climate regulation. The global plate-boundary network is shown in {numref}`fig:plate-tectonics`. Why Earth has plate tectonics while other terrestrial bodies do not remains one of the central open questions in geophysics. Contributing factors likely include Earth's size, water content (which weakens the lithosphere), and the specific rheology of its mantle {cite:p}`Stern2018`. ```{figure} figures/plate_tectonics.avif :name: fig:plate-tectonics :width: 600px :align: center Earth's major tectonic plates, each drawn in its own colour, with the plate boundaries as dark outlines. The red arrows give the sense of relative motion at each boundary: arrows pointing away from one another mark **divergent boundaries** (mid-ocean ridges, where new crust is created), arrows pointing towards one another mark **convergent boundaries** (subduction zones, where one plate descends beneath another), and offset parallel arrows mark **transform boundaries** (where plates slide past each other). Earth is the only known body with active plate tectonics; all other terrestrial planets and moons operate in the stagnant-lid regime. Credit: Wikimedia Commons, [CC BY-SA 3.0](https://creativecommons.org/licenses/by-sa/3.0/). ``` ### Stagnant-lid convection All other terrestrial bodies in the solar system operate in the **stagnant-lid regime**: the mantle convects beneath a single, rigid, immobile lithospheric lid. Heat escapes primarily by conduction through the lid and by occasional volcanic eruptions that breach it. The stagnant lid grows thicker over time as the interior cools, eventually shutting down surface volcanism {cite:p}`Stern2018`. The stagnant-lid regime is the *default* outcome of mantle convection with strongly temperature-dependent viscosity. At the base of the lithosphere, the viscosity contrast between the cold lid and the hot interior is large enough (many orders of magnitude) that the lid effectively decouples from the convecting mantle below. Earth's plate tectonics requires a mechanism to *break* the lid, likely involving water weakening and self-sustained damage along plate boundaries. {numref}`fig:l07-lid-regimes` places the stagnant lid among the tectonic modes a rocky planet can adopt: the lid can move, creep, switch between states, or stay put. ```{figure} figures/lenardic2018_tectonic_modes.avif :name: fig:l07-lid-regimes :width: 700px :align: center Tectonic modes of terrestrial planets, beyond the simple mobile-versus-stagnant binary. An active lid moves as plate tectonics or as distributed deformation; a sluggish lid shows ridge-only spreading or regional to global mobility driven by sub-lithospheric convection; episodic (transient) modes switch between states globally or locally, leaving non-recyclable "continental" regions; a stagnant lid is cold (thermal stress only) or hot (heat pipes with vertical cycling, or plumes with lower-crust recycling). Reproduced from {cite:p}`Lenardic2018`, Fig. 2. ``` ### Tectonic features across the solar system **Mars** displays dramatic tectonic features despite lacking plate tectonics. The **Tharsis bulge** (a volcanic plateau $\sim$5000 km across and $\sim$10 km high) may have been uplifted by a **mantle plume**, a persistent upwelling of hot, buoyant rock rising through the mantle. **Valles Marineris**, the solar system's largest canyon system ($\sim$4000 km long, up to 7 km deep, and 200 km wide), formed by extensional rifting associated with the Tharsis bulge {cite:p}`dePaterLissauer2010`. **Venus** shows a puzzling tectonic style. The uniform crater density across the surface implies a mean age of a few hundred Myr, leading to the **episodic resurfacing hypothesis**: Venus may experience periodic catastrophic overturns where the stagnant lid founders and the entire surface is volcanically resurfaced in a geologically short interval. Between these episodes, the surface remains tectonically quiet. This idea remains debated; we will discuss Venus's geology in detail in {ref}`Lecture 9 `. **Mercury** has undergone significant **global contraction** as its large iron core cooled and solidified over time ({ref}`Lecture 4 `), shrinking the planet's radius by up to $\sim$7 km (with method-dependent estimates spanning $\sim$4–7 km) {cite:p}`Byrne2014`. This contraction compressed the crust, producing **lobate scarps** ({numref}`fig:mercury-scarp`), thrust faults up to several hundred kilometres long and 1–3 km high, discovered by Mariner 10 and mapped extensively by MESSENGER. We will discuss Mercury's surface in {ref}`Lecture 10 `. The Tharsis-related rifting on Mars is recorded in Valles Marineris ({numref}`fig:valles-marineris`). ```{figure} figures/mercury_lobate_scarp.avif :name: fig:mercury-scarp :width: 400px :align: center A lobate scarp near Pourquoi-Pas crater on Mercury, imaged by the MESSENGER spacecraft. The scarp runs diagonally across the field of view as a sinuous step in the surface, picked out by the shadow along its steep front. It is the surface expression of a thrust fault formed during global contraction as Mercury's interior cooled and the radius shrank by up to $\sim$7 km. Because such a fault breaks the surface it cuts, the contraction that produced it must post-date the cratered terrain it deforms {cite:p}`Byrne2014`. Credit: NASA/Johns Hopkins APL/Carnegie Institution of Washington, public domain. ``` ```{figure} figures/valles_marineris.avif :name: fig:valles-marineris :width: 650px :align: center Valles Marineris, the solar system's largest canyon system, stretching $\sim$4000 km across the Martian surface (roughly the distance from Lisbon to Moscow). The canyon is up to 7 km deep and 200 km wide, dwarfing Earth's Grand Canyon. It formed primarily through extensional rifting associated with the Tharsis volcanic bulge to the west, with subsequent widening by mass wasting and possibly fluvial erosion. This Viking Orbiter mosaic is a horizontal strip showing the canyon along its full $\sim$4000 km length, from the fractured terrain of Noctis Labyrinthus on the west (left) through the main Melas, Candor, and Coprates chasmata, to the chaotic terrain at the canyon's eastern outflow into Chryse Planitia (right); the Tharsis Montes lie off the left edge of the frame. Credit: NASA/JPL/USGS, public domain. ``` ## Erosion and weathering Erosion is the removal and transport of surface material by wind, water, ice, or chemical reactions. On bodies with thick atmospheres, erosion can be the dominant surface-modifying process; Earth's surface is almost entirely shaped by erosion on timescales of millions of years. On airless bodies, the only "erosion" comes from impact gardening and space weathering (see below). ### Aeolian (wind) processes Wind-driven erosion and deposition require an atmosphere dense enough to mobilise surface particles, so the surface pressures compared in {ref}`Lecture 5 ` set which bodies can have aeolian landforms at all. Active aeolian processes shape the surfaces of {cite:p}`Catling2017`: - **Mars:** Extensive dune fields in craters and polar regions. Global dust storms can loft particles to 30 km altitude. Sand dunes in Gale Crater show active migration observed by the Curiosity rover. - **Titan:** Vast equatorial dune fields composed of organic particles (tholins) produced by atmospheric photochemistry. The dunes are longitudinal, up to 150 m tall and hundreds of kilometres long. - **Venus:** Despite its dense atmosphere, surface winds are only $\sim$1 m s$^{-1}$ due to the sluggish near-surface dynamics, limiting aeolian activity. However, the thick atmosphere allows even these slow winds to mobilise fine particles. Repeat HiRISE imaging of Martian dune fields from the Mars Reconnaissance Orbiter has shown that these dunes migrate at sand fluxes comparable to those of terrestrial dunes despite Mars's $\sim$170$\times$ thinner atmosphere (Mars surface pressure $\sim$600 Pa vs Earth's $\sim$10$^5$ Pa), and that entire dunes advance rather than only the ripples on their surfaces {cite:p}`Bridges2012`. {numref}`fig:mars-dunes` shows the barchan dune field of Nili Patera; Titan's massive equatorial dune fields are visible in Cassini SAR imagery ({numref}`fig:titan-dunes`). ```{figure} figures/mars_dunes_bridges.avif :name: fig:mars-dunes :width: 600px :align: center Barchan and barchanoid dunes on the floor of Nili Patera, a volcanic caldera on Mars, imaged by HiRISE on the Mars Reconnaissance Orbiter in March 2014 (PIA18244, observation ESP_035603_1890). The horns of each crescent extend downwind, the steep lee face is the concave slope between them, and small ripples cover the gentler upwind side. Comparison with an image of the same field from November 2012 shows ripple movement, shifting dune margins, and fresh landslides on the lee face of the central dune. An earlier study of this same dune field measured sand fluxes comparable to terrestrial values and showed that entire dunes advance, not only the ripples on them, despite the thin atmosphere {cite:p}`Bridges2012`. Credit: NASA/JPL-Caltech/University of Arizona, public domain. ``` ```{figure} figures/titan_dunes.avif :name: fig:titan-dunes :width: 600px :align: center Cassini Synthetic Aperture Radar (SAR) image of longitudinal dune fields in the Shangri-La region near Titan's equator (PIA20710). The dark, parallel ridges are dunes of organic particles (tholins) that settled out of Titan's atmospheric photochemistry; they reach up to $\sim$150 m in height and stretch hundreds of kilometres along the prevailing winds. Bright patches are topographic high-standing islands of water-ice bedrock that the dune-forming sand flows around. Credit: NASA/JPL-Caltech/ASI, public domain. ``` ### Fluvial (water) processes Liquid water is the most powerful erosive agent on Earth, and evidence for past fluvial activity on Mars is one of the most important discoveries in planetary science: - **Earth:** Rivers, glacial meltwater, and coastal waves continuously reshape the surface; the Colorado River cut the 1.6 km deep Grand Canyon in $\sim$5–6 Myr. - **Mars:** **Valley networks** on the Noachian-aged southern highlands (the **Noachian** is the oldest of the three main Martian geological periods, spanning roughly 4.1 to 3.7 Ga) resemble terrestrial river drainage systems, implying sustained liquid water flow during the late Noachian to early Hesperian, with localised reactivation continuing into the Amazonian for individual systems such as Warrego Valles {cite:p}`Ansan2006`. **Outflow channels** (e.g., Ares Vallis, Kasei Valles) are enormous flood features, hundreds of kilometres long and tens of kilometres wide, carved by catastrophic releases of groundwater. These features are discussed further in {ref}`Lecture 10 `. - **Titan:** Saturn's largest moon has **methane rivers** that carve channels into its icy surface; the Huygens probe landed among rounded ice pebbles in a dry riverbed in 2005. Titan's methane hydrological cycle is the only known active fluvial system beyond Earth ({ref}`Lecture 11 `). The two kinds of Martian water-carved feature, catastrophic outflow channels and sustained-discharge dendritic valley networks, are shown in {numref}`fig:mars-outflow` and {numref}`fig:mars-valley-networks`. {numref}`fig:grand-canyon` shows the terrestrial type example of sustained fluvial erosion. ```{figure} figures/grand_canyon_colorado.avif :name: fig:grand-canyon :width: 550px :align: center The Colorado River winding through the Grand Canyon, Arizona. The river has incised roughly 1.6 km through nearly flat-lying Paleozoic sedimentary rock over the past 5 to 6 Myr, the type example of sustained fluvial erosion on Earth. The stepped canyon walls record differential erosion: resistant limestone and sandstone form cliffs, and weaker shale erodes to slopes. Credit: Alex Demas/USGS, public domain. ``` ```{figure} figures/mars_outflow_aram.avif :name: fig:mars-outflow :width: 320px :align: center A Martian outflow channel cutting through Aram Chaos. The braided streamlined islands and broad scoured trough are the geomorphological signature of catastrophic flood discharges, plausibly fed by sudden release of subsurface ice or groundwater on early Mars. Such channels are tens of kilometres wide and hundreds of kilometres long, dwarfing any terrestrial flood feature, and are distinct from the dendritic valley networks which formed under more sustained, lower-discharge conditions. Credit: NASA/JPL-Caltech/MSSS, public domain. ``` {numref}`fig:titan-huygens` shows the rounded pebbles of the dry riverbed at the Huygens landing site. ```{figure} figures/titan_huygens_surface.avif :name: fig:titan-huygens :width: 400px :align: center Scale composite of three progressively enlarged crops from the post-landing surface image returned by ESA's Huygens probe on 14 January 2005, the only in-situ image ever taken from the surface of an outer-solar-system body. The rounded, decimetre-scale "rocks" in the foreground (bottom panel, shown at their actual apparent size) are blocks of water-ice rounded by methane fluvial transport; the surface is a damp dark plain of methane-soaked organic sediment in a dried-out riverbed. Together with the descent imagery showing dendritic drainage networks, this image confirmed that liquid methane actively shapes Titan's surface today. Credit: ESA/NASA/JPL-Caltech/University of Arizona, public domain. ``` ```{figure} figures/mars_valley_networks_viking.avif :name: fig:mars-valley-networks :width: 600px :align: center Fine dendritic channel networks dissecting the heavily cratered Martian highlands, imaged by the Viking Orbiters (detail from PIA00413, a composite of high-resolution monochrome frames with lower-resolution colour). The branching tributary pattern closely resembles terrestrial fluvial drainage and is among the strongest geomorphological evidence for sustained surface runoff (and hence precipitation) during the late Noachian to Hesperian on early Mars. Such networks are concentrated in the ancient southern highlands and are central to the case that early Mars had a warmer and wetter climate than today. Credit: NASA/JPL/USGS, public domain. ``` ### Glacial processes Ice can flow slowly under its own weight, carving valleys and transporting debris: - **Earth:** Glaciers and ice sheets have profoundly shaped mid- and high-latitude landscapes ({numref}`fig:fedchenko-glacier`). During the Last Glacial Maximum ($\sim$20 ka), ice sheets covered $\sim$30% of Earth's land area. - **Mars:** Polar ice caps of $\mathrm{CO_2}$ and $\mathrm{H_2O}$ ice. Mid-latitude features including lobate debris aprons and lineated valley fill strongly resemble terrestrial rock glaciers, suggesting subsurface ice flow. ```{figure} figures/fedchenko_glacier.avif :name: fig:fedchenko-glacier :width: 550px :align: center The Fedchenko Glacier in the Pamir Mountains of Tajikistan, photographed from the International Space Station. At roughly 77 km long it is the longest glacier outside the polar regions. Tributary ice streams join the main trunk from side valleys, and the dark medial moraines that trace each junction downstream mark the rock debris carried along the flow. Valley glaciers erode U-shaped troughs and transport sediment over long distances, a dominant landscape-shaping process at mid and high latitudes on Earth. Credit: NASA/ISS Crew Earth Observations Facility, Johnson Space Center, public domain. ``` ### Chemical weathering Chemical reactions between surface rocks and atmospheric or liquid agents alter mineral compositions: - **Earth:** Silicate weathering by carbonic acid is the critical carbon sink in the carbonate-silicate cycle ({ref}`Lecture 6 `), regulating climate over geological time. - **Mars:** Orbital spectroscopy (OMEGA on Mars Express, CRISM on MRO) has detected hydrated minerals (phyllosilicates or clays, sulfates, and carbonates) formed by aqueous alteration of basaltic rock, providing mineralogical evidence for past liquid water ({numref}`fig:mars-clay-unit`). - **Venus:** The high surface temperature ($\sim$737 K) and dense $\mathrm{CO_2}$ atmosphere drive surface-atmosphere chemical reactions that may alter rock compositions on relatively short timescales. ```{figure} figures/mars_clay_unit_glen_torridon.avif :name: fig:mars-clay-unit :width: 500px :align: center Layered bedrock of the clay-bearing unit at Glen Torridon on the lower slopes of Mount Sharp, Gale crater, imaged by the Curiosity rover's Mastcam (the rover's nameplate is visible in the foreground). Orbital and in-situ spectroscopy show this unit is rich in phyllosilicate clay minerals, which form when liquid water chemically alters basaltic rock. Clay deposits like this are among the strongest mineralogical evidence that early Mars sustained surface or near-surface water. Credit: NASA/JPL-Caltech/MSSS, public domain. ``` ## Remote sensing of surfaces Most of what we know about planetary surfaces comes from remote sensing: observing from orbit or from Earth. Different wavelengths and measurement techniques reveal different properties of the surface {cite:p}`dePaterLissauer2010`. ### Reflectance spectroscopy Every mineral has a characteristic pattern of absorption features in the visible and near-infrared (VIS/NIR, 0.3–5 $\mu$m) caused by electronic transitions and molecular vibrations. By measuring the reflected sunlight spectrum from orbit, we can identify minerals remotely: - **CRISM** (Compact Reconnaissance Imaging Spectrometer for Mars, on MRO) and **OMEGA** (on Mars Express) have mapped the distribution of phyllosilicates (clays), sulfates, and other hydrated minerals across the Martian surface, providing definitive evidence that liquid water chemically altered the rocks during Mars's early history ({numref}`fig:crism-mineral-map`) {cite:p}`Ehlmann2008`. - Similar instruments on lunar orbiters have mapped the distribution of pyroxene, olivine, and plagioclase across the Moon's surface. ```{figure} figures/crism_mineral_map.avif :name: fig:crism-mineral-map :width: 600px :align: center Two complementary views of carbonate-bearing terrain near Nili Fossae on Mars (PIA19816). Left (labelled THEMIS TI): thermal-inertia map derived from thermal-infrared observations by the THEMIS instrument on Mars Odyssey, showing surface morphology and the contrast between dusty and rocky ground. Right (labelled CRISM): spectral classification strips from the CRISM instrument on the Mars Reconnaissance Orbiter, laid over a greyscale basemap, where colour codes the dominant mineralogy: Mg-carbonates (green), olivine-bearing sands (yellow to brown), and basaltic terrain (blue). Each panel carries its own 20 km scale bar. The carbonates form by aqueous alteration of basaltic crust and provide direct mineralogical evidence for sustained liquid water on early Mars; nearby Mg/Fe-phyllosilicates (clays) detected in the same Nili Fossae region by CRISM further constrain the alteration history {cite:p}`Ehlmann2008`. Credit: NASA/JPL-Caltech/ASU/JHU APL, public domain. ``` ### Radar imaging and laser altimetry Synthetic aperture radar (SAR) transmits microwave pulses and images a surface regardless of illumination or cloud cover. Magellan (1990–1994) mapped 98% of Venus's surface at $\sim$100 m resolution through the permanent cloud deck ({numref}`fig:venus-magellan`), and almost all of Venus's surface geology still rests on that one dataset ({ref}`Lecture 9 `); Cassini's radar found Titan's methane lakes and seas through its organic haze. ```{figure} figures/venus_magellan.avif :name: fig:venus-magellan :width: 450px :align: center Hemispheric view of Venus from NASA's Magellan radar mission (1990–1994), with colour from Soviet Venera lander surface measurements. Magellan used synthetic aperture radar to penetrate Venus's permanent cloud deck and map 98% of the surface at $\sim$100 m resolution, revealing more than 1600 volcanic centres, extensive lava plains covering $\sim$80% of the surface, and a remarkably uniform crater population implying a global mean surface age of only 300–700 Myr. Venus's surface geology is discussed in detail in {ref}`Lecture 9 `. Credit: NASA/JPL, public domain. ``` Laser altimeters time the round trip of a laser pulse to measure elevation with metre-scale precision. MOLA on Mars Global Surveyor produced the definitive topographic map of Mars ({numref}`fig:mars-topography`), including the $\sim$6 km hemispheric dichotomy, the Tharsis bulge, and the full extent of Valles Marineris {cite:p}`Smith2001`; LOLA on the Lunar Reconnaissance Orbiter did the same for the Moon and revealed the permanently shadowed polar craters. ### Gravity field mapping Precise tracking of spacecraft orbits reveals variations in a body's gravitational field, which reflect lateral density variations in the interior: - **GRAIL** (Gravity Recovery and Interior Laboratory, 2012) mapped the Moon's gravity field to extraordinary precision, revealing the crustal thickness variations ({numref}`fig:grail-crust`) and the structure of impact basins and mascons (mass concentrations). - **GRACE** did the same for Earth and tracks ice-sheet mass loss and groundwater depletion. We will discuss gravity field measurements and their interpretation further in {ref}`Lecture 8 `. ```{figure} figures/grail_crustal_thickness.avif :name: fig:grail-crust :width: 600px :align: center Lunar crustal thickness derived from the GRAIL mission, shown for the nearside (left) and farside (right) hemispheres {cite:p}`Wieczorek2013`. GRAIL measured the Moon's gravity field at unprecedented precision by tracking the inter-spacecraft distance between two co-orbiting satellites; combined with topography, the gravity solution yields the depth to the crust-mantle interface. The map reveals an average crustal thickness of $\sim$34-43 km {cite:p}`Wieczorek2013`, with strong thinning beneath the largest impact basins (deep blue), a thicker farside crust, and the Procellarum KREEP Terrane on the nearside (cooler colours within the Procellarum outline). Credit: NASA/MIT/GSFC/JPL-Caltech, public domain. ``` ```{figure} figures/mars_topography.avif :name: fig:mars-topography :width: 650px :align: center Global cylindrical topographic map of Mars from the Mars Orbiter Laser Altimeter (MOLA) on Mars Global Surveyor {cite:p}`Smith2001`. Elevations span $\sim$30 km, from the summit of Olympus Mons ($\sim$21 km above datum, white, left) to the floor of the Hellas basin ($\sim$8 km below datum, deep blue, lower right). The Tharsis bulge and its four shield volcanoes dominate the western hemisphere; Valles Marineris stretches eastward across the equator. The $\sim$6 km elevation difference between the cratered southern highlands and the smooth northern lowlands (the **hemispheric dichotomy**) remains one of the major unsolved problems in Martian geology. Credit: NASA/JPL/GSFC/MOLA Science Team, public domain. ``` ## Regolith formation and space weathering On airless bodies (the Moon, Mercury, asteroids), the surface is not bedrock but a loose, fragmented layer called **regolith**, produced by billions of years of impact processing and radiation exposure {cite:p}`Hapke2001`. ### Regolith The lunar regolith is a layer of unconsolidated debris (rock fragments, mineral grains, and glass beads) produced by the cumulative effect of impacts at all scales, from micrometeorite bombardment to basin-forming events. This process is called **impact gardening**: each impact excavates material, mixes the surface layer, and breaks rocks into progressively finer particles. Regolith depth therefore tracks how long a surface has been exposed. A global survey of crater morphologies in Lunar Reconnaissance Orbiter images gives median depths of typically 2–4 m on the young mare basalts and 6–8 m on the farside and non-mare nearside, so the ancient highlands hold roughly twice as much regolith as the maria {cite:p}`Bart2011` ({numref}`fig:lunar-regolith`). The Hayabusa2 and OSIRIS-REx sample returns showed that even sub-kilometre **rubble-pile** asteroids, with negligible gravity, retain a regolith rather than losing their ejecta to space; the returned samples themselves are discussed in {ref}`Lecture 12 `. {numref}`fig:bennu-osiris-rex` shows Bennu: boulders at every scale, no smooth ponded plains, and a bulk density that implies a porous rubble pile. ```{figure} figures/bennu_osiris_rex.avif :name: fig:bennu-osiris-rex :width: 450px :align: center The near-Earth asteroid Bennu, a mosaic of images taken by NASA's OSIRIS-REx spacecraft in December 2018 from a range of 24 km. Boulders crowd the $\sim$490 m body at every scale, with no smooth ponded plains, and its very low bulk density implies high internal porosity: a rubble pile of collisional fragments reaccumulated under self-gravity. OSIRIS-REx returned 121.6 g of this material to Earth in September 2023; the samples carry hydrated clay minerals, carbonates, amino acids, and nucleobases {cite:p}`Lauretta2024`. Credit: NASA/Goddard/University of Arizona, public domain. ``` ```{figure} figures/lunar_regolith.avif :name: fig:lunar-regolith :width: 400px :align: center Apollo 11 photograph AS11-40-5878 of an astronaut bootprint in the lunar regolith, taken by Buzz Aldrin during the first crewed lunar EVA on 20 July 1969 as part of an in-situ soil mechanics experiment. The cohesive, fine-grained regolith preserves the boot impression in sharp relief, demonstrating both its low bearing strength and the absence of erosion (no wind, water, or atmosphere). Such bootprints will remain visible on the lunar surface for millions of years. Credit: NASA/Apollo 11, public domain. ``` ### Space weathering The surfaces of airless bodies are continuously exposed to the space environment: solar wind ions (primarily protons and $\mathrm{He}^{2+}$), micrometeorite impacts, and galactic cosmic rays. These agents collectively produce **space weathering**, a set of physical and chemical changes that modify the optical properties of the surface over time {cite:p}`Hapke2001`: - **Solar wind implantation:** Energetic ions are implanted into the top $\sim$100 nm of mineral grains, creating crystal damage and amorphous coatings. - **Micrometeorite melting:** Tiny impacts at 10–70 km s$^{-1}$ melt and vaporise surface material, leaving **nanophase iron** (np-Fe$^0$), metallic iron droplets of a few to several hundred nm embedded in glassy rims on mineral grains. - **Cosmic ray damage:** High-energy particles create lattice defects in crystal structures. The net effect is that space-weathered surfaces become **darker and redder** over time. This is why fresh impact craters (e.g., Tycho on the Moon, with its bright ray system) stand out as brighter features against the darker, mature regolith surrounding them. Space weathering complicates the spectroscopic identification of surface minerals, since the absorption features are weakened and shifted, a significant challenge for remote sensing. {numref}`fig:space-weathering` summarises the agents of space weathering and their spectral result. ```{figure} figures/space_weathering.avif :name: fig:space-weathering :width: 100% :align: center Space weathering of an airless surface. (a) Schematic cross-section of a regolith grain: implanted solar wind ions (H$^+$, He$^{2+}$) and micrometeorites arriving at $10$ to $70$ km s$^{-1}$ convert the top $\sim 100$ nm of the grain into an amorphous rim that holds nanophase metallic iron blobs of a few to several hundred nm; micrometeorite impacts also produce melt splash and vapour, and galactic cosmic rays penetrate deeper into the grain. (b) Qualitative reflectance spectra: a mature, space-weathered surface is darker and redder than a fresh surface and shows a shallower $1$ $\mu$m absorption band. No numerical reflectance values are implied. Course-original figure. ``` ## Cryovolcanism on icy bodies In the outer solar system, where surface temperatures are far below the freezing point of water, volcanic processes take a different form. **Cryovolcanism** involves the eruption of volatile-rich "magma" (liquid water, ammonia-water mixtures, or methane) rather than silicate melts. The energy source is typically tidal heating ({ref}`Lecture 3 `), which can maintain subsurface oceans beneath icy shells. ### Enceladus, Europa, and Triton Saturn's small moon **Enceladus** ($R \approx 252$ km) is the most dramatic case. Cassini found geysers of water vapour and ice erupting from four parallel fractures, the **tiger stripes**, near the south pole ({numref}`fig:enceladus-cryovolcanism`) {cite:p}`Porco2006`. The plumes come from a global subsurface ocean in contact with the rocky core, kept liquid by tidal heating; they carry molecular hydrogen, silica nanoparticles, and organic molecules, consistent with hydrothermal vents on the ocean floor {cite:p}`NimmoPappalardo2016`, and the tiger stripes radiate $\sim$15.8 GW, far more than radiogenic heating can supply {cite:p}`Howett2011`. ```{figure} figures/enceladus_cryovolcanism.avif :name: fig:enceladus-cryovolcanism :width: 500px :align: center Plumes of water vapour and ice particles erupting from the south polar region of Saturn's moon Enceladus, imaged by NASA's *Cassini* spacecraft. The geysers originate from four parallel fractures called "tiger stripes" and are sourced from a global subsurface ocean in contact with the rocky core. The plumes contain molecular hydrogen, silica nanoparticles, and complex organic molecules, ingredients consistent with active hydrothermal chemistry on the ocean floor ({ref}`Lecture 14 `). Credit: NASA/JPL-Caltech/SSI, public domain. ``` Jupiter's moon **Europa** ($R \approx 1561$ km) holds a global ocean $\sim$100 km deep beneath an ice shell $\sim$15–25 km thick, maintained by tidal heating in the Laplace resonance with Io and Ganymede {cite:p}`NimmoPappalardo2016`. Its surface is crossed by **lineae**, cracks filled from below, and broken into **chaos terrain** where the shell fractured, rotated, and refroze ({numref}`fig:europa-chaos`); the near-absence of craters gives a surface age of only $\sim$40–90 Myr. ```{figure} figures/europa_chaos.avif :name: fig:europa-chaos :width: 450px :align: center Galileo SSI image of Europa's surface near Pwyll crater (the bright ray system in the lower centre, with the dark crater floor at its hub), showing the network of dark double ridges and lineae that crisscross the moon's icy crust. The criss-crossing lineae record successive episodes of fracturing and resurfacing as tidal stresses flexed the brittle ice shell over a $\sim$100 km deep subsurface ocean {cite:p}`NimmoPappalardo2016`. Europa's crater density implies a mean surface age of only 40–90 Myr, making it among the youngest surfaces in the solar system. Credit: NASA/JPL-Caltech/University of Arizona/University of Colorado, public domain. ``` Neptune's moon **Triton** vents nitrogen geysers $\sim$8 km high, seen by Voyager 2 in 1989 ({numref}`fig:triton-surface`), shows a dimpled "cantaloupe terrain" with no analogue elsewhere, and may hold a subsurface ocean of its own. The three bodies, and the Europa Clipper and JUICE missions now on their way to the Jupiter system, return in {ref}`Lecture 11 `. ```{figure} figures/triton_voyager.avif :name: fig:triton-surface :width: 450px :align: center Voyager 2 colour mosaic of Neptune's moon Triton (PIA00317), captured during the 1989 flyby. The pinkish southern polar cap (lower half) is a thin $\mathrm{N_2}$/$\mathrm{CH_4}$ frost; the dark streaks pointing northeast across the cap are wind-deposited dust from active nitrogen geysers, the only confirmed cryovolcanic plumes outside Enceladus. North of the cap, the "cantaloupe terrain" of dimples and ridges is unique in the solar system and remains poorly understood. Triton's retrograde orbit suggests it is a captured Kuiper-belt object, making it a probable analogue for Pluto-class dwarf planets. Credit: NASA/JPL/Voyager 2, public domain. ``` ## Recent advances Io is the one body beyond Earth where volcanic resurfacing can be watched from month to month. JunoCam, on the Juno extended mission, imaged Io on a series of close flybys in 2023 and 2024, and two frames of the Nusku region taken two months apart in early 2024 bracket a single eruption: a bright red ring of fresh sulfur-rich pyroclastic deposits appears around the vent in April and is absent in February ({numref}`fig:io-nusku`). Together with ground-based adaptive-optics monitoring, this gives the first sustained spatially resolved spacecraft record of Io's volcanism since Galileo and constrains where the tidal heat comes out ({ref}`Lecture 11 `). ```{figure} figures/io_nusku_change.avif :name: fig:io-nusku :width: 600px :align: center JunoCam imagery of the Nusku volcanic region on Io taken two months apart in early 2024 (PIA26488). A new bright red ring of fresh sulfur-rich pyroclastic deposits appears around the central vent in the April image but is absent in February, recording a single eruptive event observed in real time. JunoCam's repeated close flybys in 2023 and 2024 provide the first sustained spatial monitoring of Io's volcanism since Galileo. Credit: NASA/JPL-Caltech/SwRI/MSSS/Jason Perry, public domain. ``` The New Horizons flyby in July 2015 showed that even a small, distant dwarf planet can be geologically active today ({numref}`fig:pluto-true-color`) {cite:p}`Stern2015`. Sputnik Planitia, a $\sim$1000 km wide plain of nitrogen ice, is divided into polygonal cells tens of kilometres across that record active solid-state convection in the ice, and its complete lack of impact craters implies a surface age below $\sim$10 Myr {cite:p}`McKinnon2016`. Water-ice mountains up to $\sim$3 km tall stand along its margins, and the deep interior may hold a liquid water ocean. ```{figure} figures/pluto_true_color.avif :name: fig:pluto-true-color :width: 450px :align: center Pluto in enhanced true colour, from images taken by New Horizons on 14 July 2015. The bright western lobe of the heart-shaped Tombaugh Regio is Sputnik Planitia, a vast crater-free plain of convecting nitrogen ice bordered by water-ice mountain ranges; the dark reddish equatorial terrains are coated in tholins, complex organics produced by radiation chemistry. Credit: NASA/JHUAPL/SwRI, public domain. ``` ## References ```{bibliography} :filter: docname in docnames ```