Volkanoloji Volcanoes 3C11 Planetary Geology Volcanism VOLCANISM (4 lectures) 1 Introduction Volcanoes and lava fields are found on all the terrestrial planets, and many smaller satellites of the outer planets. Some volcanic provinces have been extinct for billions of years, while others are much younger, and at least three planetary bodies are currently active. The prominent role that volcanism plays in the surface history of most planets is due to the fact that planets are born hot - and continue to produce heat - throughout their lifetime. It is through the processes of magmatism and volcanism that planets can cool off, by transporting heat to the surface. 2 Getting rid of heat The energy budget of a planet is a balance between heat production, and the escape of heat through the surface. The heat production is usually proportional to the amount of radioactive material locked within the body, which to the first order is in proportion to the planet’ s volume. Conversely, the escape of heat is governed by the surface area of the planet from which heat can escape. Planetary bodies of a limited radius have a surface area which can maintain an equilibrium in the internal heat budget by simple conduction of the heat flow through a solid lithosphere up to the surface; their advantageous area-to-volume ratio ensures that the bulk of their heat is evacuated as fast or faster than it is generated at depth. Larger planets, on the other hand, suffer from an unbalanced area-to-volume ratio, so that they heat up faster than they can cool through conduction alone. Temperatures become so high that the planet’ s interior begins to melt. When a planet begins to melt, it does so at very specific levels, dictated by the temperature (activating melting) and pressure (inhibiting melting) gradients, as well as mineral and volatile constituents. 2.1 Partial melting Melting can occur only where there is a change in temperature, pressure, or bulk chemistry. When such a variation occurs, it is not the entire rock that melts uniformly, but only a fraction. The various minerals within a rock have different melting temperatures, and the most fusible crystals will liquefy first. As the conditions for melting become more favourable, more refractory minerals will break down in turn, joining the melt. This gradual process, called partial melting, strongly controls the chemical composition of any melted rock (magma) being produced, being highly dependant upon the fraction of the source rock that melts. However, the total volume of any rock within a planet will rarely be dominated by partial melting. Geophysical studies of the Earth’ s lithosphere suggest that partial melt constitutes only 1-3% of the upper mantle region. FIGURE 1: Partial melting If planetary interiors were immobile, then whatever the changes which brought about melting would be essentially random. However, as we shall see, volcanism is usually localised to specific regions on the planetary surface, reflecting variations in the extent of partial melting controlled by localised variations in temperature, pressure and composition. 2.2 Ascent of magma With only a few per cent partial melting, any magmatic liquid that is generated will exist tenuously along grain boundaries, held within the solid matrix. As melting proceeds, a greater volume of liquid accumulates, and pockets of melt can join. Once the melt can ooze out of its rocky matrix - along fine fractures in the rock - magma collects in larger fissures and starts to ascend to the surface. The liquid magma is less dense that the rock matrix from which it derives, and so its buoyancy causes it to rise through any cracks it can find or create on the way. The greater the difference in density, the faster the ascent velocity. The force from the melt’ s buoyancy is counteracted by resistance in the enclosing material, and can only rise once it overcomes the total of all resistive forces, so that it will preferentially follow previously produced faults. Magma begins to rise as bulges, and though the smaller bulges cool and lose their density contrast, the larger ones slowly grow upwards, their rate of progress gradually accelerating until they form into steep-sides diapirs or plumes.3C11 Planetary Geology Volcanism FIGURE 2 (Figure 4.1 Cattermole p. 45): The generation of diapirs In order to reach the surface directly, a magma needs to rise quickly - in the range of 1-2 metres per second - because if the rise is too slow, the magma has time to cool through contact with the surrounding rock, eventually solidifying and coming to a halt. However, the rising body of hot magma leaves behind it a heated and volumetrically expanded zone, resulting in preferential inclusions of subsequent batches of melt. With the repeated influx of successive batches of newly generated magma, the surrounding rocks would gradually be heated to melting, allowing the magma to continue to rise. This may partially explain the persistence of volcanism at given locations over prolonged periods. Solidified magma which never reaches the surface is termed plutonic rock, in contrast to volcanic rocks at the surface. Magma which infiltrates horizontal and vertical faults tends to open up the cracks, resulting in extensive sheets of magma, called sills when horizontal, and dykes when vertical. The ascent of magma is thus a complex process. Some melts never reach the surface, others rise directly to the surface without stopping on the way, while some batches ascend in a stepwise fashion, stagnating at various levels before a change in the local conditions - a surge in temperature, a drop in pressure, or the wrenching open of new cracks - allow the magma to resume its ascent. 2.3 Formation of magma reservoirs Reservoirs of magma can form beneath the surface of a planet, storing large amounts of magma until conditions change in such a way as to cause an eruption, or, if there is no replenishing source, until the magma solidifies. In order for a reservoir to form, the magma’ s ascent must be curtailed. Deep magma reservoirs can be produced when the magma encounters a rheological barrier, and shallow magma reservoirs from when the magma reaches a point of neutral buoyancy beneath the surface. 2.3.1 Deep magma reservoirs Deep magma reservoirs can form in a wide range of mantle environmental settings (Figure 3), where they can no longer rise buoyantly upon reaching a rheological barrier in the form of a viscous lid, commonly the lithosphere. Figure 3a - the diapiric rise of magma from depth through the mantle can result in the diapir head solidifying as the temperature drops below the melting-point, leading to the diapir flattening and spreading laterally against a viscous lid. This produces a large, partially molten magma source region serving as a deep magma reservoir. Figure 3b - plumes under a stationary overlying lithosphere will produce a flattened plume head and a large, partially molten zone underlying the region of uplift, providing abundant magma to the surface. As the plume evolves, the head mixes and FIGURE 3 (Figure 1, Head and Wilson, p. 3878): Generalised environments of mantle upwelling, formation of deep magma reservoir and pressure-release melting3C11 Planetary Geology Volcanism cools, though the plume tail continues to supply magma at a lower rate, resulting in a smaller and more centralised reservoir. Figure 3c - plumes intersecting with a moving viscous lid initially produce a flattened diapir resulting in voluminous eruptions and crustal thickening. This plateau is moved from the plume, leaving a small plume tail which produces a sequence of individual volcanic edifices. Figure 3d - phase changes or other discontinuities in the mantle can cause a plume head to become detached, behaving more like an independent diapir as it encounters the viscous lid. Figure 3e & 3f - relatively deep reservoirs can also occur with rift zones, as a result of regional extension, leading to crustal thinning, localised mantle upwelling and pressure-release melting. They can also be associated with deeper patterns of mantle upwelling. Provided magma is still buoyant, the rheological barrier doesn’ t necessarily lead to its migration being halted. Dykes may propagate from the stalled diapir and magma body, allowing continued melt ascent. Any dyke more than a few hundred metres in vertical extent containing buoyant magma is likely to continue to fracture its host rocks indefinitely once it has begun to propagate, and is thus not greatly influenced by discontinuities in the mechanical properties of the host rock. 2.3.2 Shallow magma reservoirs FIGURE 4: The development of a shallow magma reservoir Magma will continue to rise within a dyke until it reaches a level of neutral buoyancy, becoming trapped at its upper and lower edges, unless its vertical extent is greater than several kilometres. Regions of neutral buoyancy are produced by a crossover in the densities of the rising magma and the surrounding countryrock. Beneath this horizon, the magma will buoyantly ascend, whereas above the horizon, it will descend under the influence of negative buoyancy forces. The trapping of a large number of magma batches within a small region in this way can then lead to the accumulation of a substantial shallow magma reservoir, as the magma extends laterally along the neutral buoyancy zone. Normally, given that the magma is hotter than the surrounding countryrock, it should also be less dense. The density of the surrounding rock can reduce below that of the magma nearing the surface in one of two ways. If the composition of the crustal rocks is significantly different to that of the magma, the density can decrease sharply at a boundary, and the magma is trapped beneath. FIGURE 5: The decrease in density through contractancy A second source of decreasing countryrock pressure near the surface is the progressive opening up of pore space as the pressure within the rock decreases. The progressive reduction of macroscopic and microscopic pore space as a function of increasing depth and pressure is called contractancy. This process is able to significantly change the density of the uppermost fractured veneer of rock above an active magma reservoir, and in doing so, produce the horizons of neutral buoyancy which result in magma stabilisation. The pressure in the countryrock at any particular depth z can be calculated:3C11 Planetary Geology Volcanism (Eqn 1) Where, r cr = density of crust at depth z, r = density of compressed country rock, g = gravitational force, V 0 = surface void space and is a constant, given by Head and Wilson (1992) as 1.18 x 10 -8 Pa -1 , on the basis of fitting data from Earth 2.4 Magma chambers When a source of magma is feeding into a neutral buoyancy zone, and the magma becomes trapped vertically, a build-up in pressure occurs. This forces the magma to spread laterally along the neutral buoyancy zone, opening up cracks and melting down whole panels of the host rock, forming into a magma chamber. Chambers tend to regulate the pace of eruptions to the surface, since they store and distribute the rising magma. They can exist for extended periods of time, with the magma being stored within them for centuries, even millennia, before cooling to stone, or encountering new conditions which allow them to grow new shoots upwards. Chambers also play an important chemical role, because in times of storage, the magma cools and crystals form. As the most refractory minerals settle out of the liquid, withdrawing certain atoms from the melt preferentially, the chemical proportions in the remaining melt change respectively. This chemical evolution of magma is called crystal fractionation, and is a highly complex process. Denser crystals sink to the bottom of the chamber, while lower density minerals rise and coat the sides and top of the chamber. In addition, convection within the chamber, eruptions, new injections of magma from beneath, and erosion of previously settled crystals all complicate the process significantly. The changing chemical composition with age can be followed on Earth, from one eruption to the next, with the succession of changing lava being called a lava suite. The progression within such a suite is driven by the fractional crystallisation of the more refractory minerals in the melt, leading to the depletion of the atoms which make up these minerals in the melt. In silicate rocks (the most common rock type on all the terrestrial planets, which are dominated by silica), this results in the gradual elimination of magnesium, iron and calcium (predominately the minerals olivine and pyroxene) from the melt. This in turn results in a melt that is enriched with sodium and potassium (which will result in minerals called orthoclase and albite), and silica itself. If cooling proceeds and more crystals settle out of an ever-shrinking magma pocket, ever more radical chemistries are reached. If there are any volatiles within the melt (principally carbon dioxide and water for the terrestrial planets), they will not be affected by this cooling process unless they pass the saturation point at the given pressure of the system. The composition of an erupting magma will thus reflect the history of the magma from which it was produced, including variations in the partial melting process at the source, and evolution during ascent, storage, and eruption. FIGURE 6: The differentiation of magma through crystal fractionation 2.5 Eruption processes Magma trapped under pressure at either a rheological boundary or a neutral buoyancy zone will remain in place until the magma solidifies, or the conditions change to allow an eruption to occur. One of the commonest causes of change in condition is through the opening of cracks by pressurised volatiles. As the volatile content of the magma reaches supersaturation (when the magma cannot host its compliment of volatiles), either through a decrease in pressure, or through the increasing concentration of volatiles in the fractionation process, these volatiles begin to exsolve. The pressure in the magma chamber changes gradually due to this exsolution of volatiles, and the stresses produced by this increase will tend to open up cracks to the surface, initiate the propagation of dykes, and form a volcanic conduit. Once a conduit from the magma chamber to the surface is opened, magma will tend to flow upwards under the influence of a negative pressure gradient. However, the rise of magma from a reservoir at depth is primarily controlled by its volatile content.3C11 Planetary Geology Volcanism As a given batch of magma passes through the conduit, it moves from higher to lower pressures, causing the saturation point of the volatiles to drop. This leads to the exsolution of volatiles from the melt, causing bubbles to form. Any bubbles that are already exsolved will begin to expand, as the pressure decreases. This expansion causes momentum and energy to be passed into the magmatic flow, thus increasing the velocity along the conduit. FIGURE 7: Conduit processes during an eruption If the volatile content is high enough, as the pressure continues to drop, the magma becomes foam-like, and when total gas volume fraction reaches ~75%, the bubbles begin to coalesce to form a continuous phase, fragmenting the magma. If the bubbles are large enough to decouple from the surrounding melt - and the magma has a low enough viscosity - this will result in the volatiles flowing up the conduit separate from the surrounding magma. The eruptive behaviour of a volcano is thus closely related to the magma’ s volatile content, as well as the size and shape of the eruptive conduit. If a rising magma is gas-free, the magma rises relatively slowly to the surface, passes through the vent, and flows away as lava, in a relatively quiescent fashion. If the melt contains a gas component, at some depth beneath the vent the magma begins to vesiculate, so that upon reaching the surface it will spray out under decompression, throwing out molten clots which fall back to the ground and coalesce into a flow. If the gas has formed into a separate phase, however, upon reaching the vent, the magma disrupts into pyroclasts, which are ejected explosively as an ascending cloud of gas and solids. FIGURE 8: The influence of volatile content on the eruption process 3 Volcanic Materials The term volcanism implies the activity of molten materials (magmas), which eventually crystallise to form solid rocks or volcanic glasses. General usage of the term covers the activities of silicate magmas, but the definition of volcanism can be widened to include materials found in other conditions: surface units deposited by sulphur (Io), and by ice and ammonia compounds (other moons of the outer planets). 3.1 Silicate rocks Silicate rocks (SiO 2 -based) are made of combinations of seven mineral families (with a few exceptions, such as carbonitites, which are rocks which have carbon, rather than silica, as their main constituent). These families are: olivines, pyroxenes, amphiboles, micas, feldspars, quartz and oxides. All but oxides are silicate minerals; oxides are mostly iron and titanium.3C11 Planetary Geology Volcanism Within each family, the chemical composition may vary considerably within certain limits. You will see many names given to volcanic rocks depending upon their composition; they can be considered to vary in two ways - their silica content and their alkali content. The alkali content of a rock is expressed in terms of its (Na 2 O + K 2 O) content. Within each trend (say, basalt, andesite and rhyolite), there will be a variation in alkali content, (which are most easily dealt with by naming the rocks alkali basalts or alkali rhyolites, etc.) FIGURE 9: A composition- classification cube for silica rocks 3.2 Carbonate rocks Carbonate rocks (CO 3 -based) are far less common than silicates. They are most common on Earth, where life produces limestone rocks which are recycled underground to produce carbonate magmas. Carbonates can, however, be produced through non- organic processes, under the right conditions. They do not form through crystallisation, as silicates do; they form from an immiscible fluid that separates from the main melt, so that the formation of predominantly carbonate melts requires significant processing of the rocks. Some carbonates have a melting point as low as ~800K, and a viscosity similar to that of water. Carbonates may be more common on Venus, where the thick CO 2 atmosphere provides a rich source of the volatile. 3.3 Sulphur eruptions The surface of Io contains a wide range of colours, from black, through red and orange, to yellow. These colours are best explained by the presence of sulphur on the surface, along with sulphur dioxide frost, and sulphurous salts of sodium and potassium. Recent analysis of images from Galileo shows that Io’ s volcanism is not consistent with a simple sulphur cooling model, but the possibly remains that a more complex behaviour involving sulphur flows is at work. Equally, sulphur might form only a thin veneer at and near the surface, or simply affect the chemistry of predominantly silicate magmas. FIGURE 10: The viscosity of molten sulphur as a function of temperature Sulphur is an unusual substance. It boils at higher than about 550K, and can remain molten down to 392K. Thus, magmas on Io can be generated at much lower temperatures and lava flows can remain molten at much lower temperatures. The variation of viscosity with temperature is also unusual (Figure 10), so that lava flows can actually become more fluid as they cool. 3.4 Water eruptions Water ice is a very special ’ mineral’ , in that it increases in density when it melts. Because of this reversed density behaviour, melt water would not be expected to rise through its icy matrix to the surface. However, the ices in the outer Solar System are impure, mixed with other frozen volatiles (namely ammonia) and silicate rock: these dirty ices behave differently than pure ice, experiencing ’ normal’ phase changes with expansion during melting, and becoming buoyant in the liquid phase. These are the ices that fuel cryovolcanism, causing flows and icy sprays at the surface.3C11 Planetary Geology Volcanism 4 Types of Eruption There are two principle styles of eruptive process: there are eruptions which emanate from a central vent (central eruptions) and those which are associated with an extended fracture vent (fissure eruptions). Central eruptions are related to the rise of magma up a single conduit, erupting material in a roughly radial pattern around the vent. Fissure eruptions occur in regions suffering crustal extension, where magma is able to force its way to the surface along one or more pathways, forming a series of dykes along the region of fractures, so that lava can erupt from open fissures. As we’ ve seen, different volatile contents can vary an eruption between effusive activity (passive emission of lava) and explosive activity, through everything in-between. On small scales, magma may erupt either effusively or explosively. On large scales, however, viscosity (controlled by composition) becomes important; large basaltic eruptions are almost exclusively effusive, large silicic eruptions almost exclusively explosive. The viscosity of a fluid is also an important parameter when considering volcanic processes. Viscosity may be defined as the internal resistance to flow by a substance when a shear stress is applied. The behaviour of different fluids with various viscosities are given specific terms. FIGURE 11: Stress vs. Strain As we have seen with impact crater processes, fluids such as water are described as Newtonian, flowing at the slightest stress. Bingham fluids do not flow until a critical stress has been applied. This initial shear stress is termed its yield strength. Once moving, however, Bingham fluids flow at a rate directly proportional to the applied stress, like Newtonian substances. Intermediate between Bingham and Newtonian are the pseudo- plastics which show a non-linear relationship between stress and strain, and have no definite yield strength. Most lavas behave in a pseudo-plastic way, although in simple terms they are often regarded as Binghams. Several factors affect the viscosity of a fluid, including temperature, composition, volatile content and crystallinity. Temperature has a rather dramatic effect on magmas, increasing its viscosity by up to several orders of magnitude for a decrease of several hundreds of degrees. All types of lava show this trend, but the amount by which viscosity increases for a particular temperature drop depends on composition. FIGURE 12: Effect of volatiles and temperature on viscosity To reduce the viscosity of a silicate melt, its silicon-oxygen bonds must be broken. One effective way of doing this is by adding water, which breaks the bonds by forming OH - ions. Water will reduce the viscosity of both silicic and mafic melts, but the effect is more dramatic in silicic ones. The presence of solid crystals in the melt will tend to increase its viscosity, but the exact effect is difficult to quantify due to the variation in shape and size of the crystal.3C11 Planetary Geology Volcanism 4.1 Flood lavas If a magma is totally devoid of volatiles, it will flow peacefully out of the vent. Here it will form a lava flow, moving akin to water under the influence of gravity, finding topographic lows. Unlike water, however, a moving lava flow will cease to move when the rheological properties of the lava prevent its further progress. 4.1.1 Lava Flows FIGURE 13: Thickness of a flow The thickness, t, a lava flow will attain before it begins to flow is given by: a r t tan g t = (Eqn. 2) where t = yield strength, r = density, a = slope and g = acceleration due to gravity. Thus, a more viscous lava will have a larger yield strength and hence greater thickness. The velocity of a flow is then given by: a h r sin v 2 B gt = (Eqn. 3) where B = constant (~3 for a flat surface) and h = viscosity. Here the more viscous lavas will travel slower than the less viscous ones. How far a lava flow will travel is difficult to predict. With all else being equal, the effusion rate is one of the most important factors. In general, basaltic melts have far higher effusion rates than silicic ones and hence travel farther. The lava flows made of basaltic material are likely to be the ones we refer to most when looking at bodies such as the Moon and others. Two forms of basaltic lava are familiar to geologists: a’a and pahoehoe. Both types of lava may be erupted from the same vent; chemically a’a and pahoehoe may be the same, it is just their physical structure that is different. The mass eruption rate of lava , M, can be given as d u wL M r = (Eqn. 4) where w = width of the fissure, L = length of the fissure, r = density of the magma and u d = rise velocity of the magma. Thus, longer and wider fissures produce much higher eruption rates. These are typically basaltic in composition, low viscosity, high velocity flows. It is a common form of volcanism on other planets, where flows extend for hundreds of kilometres. 4.2 Hawaiian eruptions One of the mildest forms of eruption is that of Hawaiian activity, characterised by relatively quiet effusion of low viscosity (hot, fluid, and basaltic) lava, and brought about because the high-pressure gases within the rising magma are able to escape both readily and steadily, preventing a larger-scale coalescence of gas bubbles within the magma. The vigorous decompression of these volatiles as they rise to the surface blows the surrounding magmatic froth to shreds, spraying ash and magma droplets into the air, and producing a fire fountain at the vent head. 4.3 Strombolian activity Strombolian Activity also involves basaltic magma, but with a higher viscosity and yield strength than in Hawaiian activity. It is also more explosive, being characterised by intermittent bursts of activity, throwing pyroclastic material tens or hundreds of metres into the air. The magma in a Strombolian eruption rises slowly in the vent, allowing bubbles to coalesce and separate from the melt, as they rise to the top of the magma column. Upon reaching the surface, the gas pocket disrupts the magma, throwing it out ballistically. The whole cycle repeats, with individual blasts separated by anything from a fraction of a second to hours.3C11 Planetary Geology Volcanism FIGURE 14: Strombolian eruption 4.4 Vulcanian activity A more explosive form of volcanism is known as Vulcanian. These eruptions are usually small (<1km 3 ), but material can be erupted to far greater heights than in Strombolian eruptions. A higher viscosity magma is required for Vulcanian activity and andesitic magma is usually involved. Separate bursts of activity take place at intervals of minutes to hours. The ejected material is often found to be the remains of a shattered lava plug which formed in the vent. The exsolution of gas from great depths causes a build-up of pressure which eventually gets so great it breaks the overlying plug, ejecting the fragments at high velocity (200-400ms -1 ). FIGURE 15: Vulcanian eruption 4.5 Plinian eruptions Plinian eruptions are another form of explosive volcanism, but are relatively rare. They differ from previous examples in that their eruptions consist of sustained jets lasting for minutes or even hours. In Plinian eruptions, exsolution may start at deep levels. Crucially, the bubbles do not rise through the magma. Instead the magma and exsolved gases rise up at the same velocity. This can happen in high viscosity magmas (i.e. rhyolites) although basaltic Plinian eruptions are not unknown. In these cases the magma would have to have a high velocity through the vent such that the bubbles did not have much time for relative movement. Material is ejected at several hundred metres per second, with huge amounts of material rising convectively in a column above the vent. Once the material in the column is finally deposited, the consequences can be devastating. Pompeii is an example of what can happen when a Plinian eruption (in this case Vesuvius) occurs.3C11 Planetary Geology Volcanism FIGURE 16: Plinian eruptions 4.6 PelØan eruptions PelØan eruptions begin in the same way as Plinian eruptions, with huge amounts of material being ejected at high velocity, rising convectively in a column above the vent. However, a high discharge rate means that the clouds of matter erupted are particularly dense, and the pressure of material being pushed up from beneath is not enough to carry the weight of the overlying material. This results in the gravitational collapse of the entire vertical eruption column back onto the vent region, and spreading as a base surge over the flanks of the volcano. The mix of gas and lava flows out over the surrounding regions in the form of pyroclastic flows (also known as nuØes ardentes). FIGURE 17: PelØan eruptions 4.6.1 Pyroclastic flows Pyroclastic flows are high-density mixtures of hot, dry rock fragments and hot gases that move away from their source vents at high speeds. They may result from the collapse of vertical eruption columns of ash and larger rock fragments, or from the explosive eruption of molten or solid rock fragments. They can also result from a laterally directed explosion, or the fall of hot rock debris from a dome or thick lava flow. Most pyroclastic flows consist of two parts: a basal flow of coarse fragments that moves along the ground, and a turbulent cloud of finer particles (ash cloud) that rises above the basal flow. Pyroclastic flows are extremely hazardous because of their high speeds and temperatures. Objects and structures in their paths are generally destroyed or swept away by the impact of debris or by accompanying hurricane-force winds. Wood and other combustible materials are commonly burned by the basal flow; people and animals may also be burned or killed beyond the margins of a pyroclastic flow by inhalation of hot ash and gases. Pyroclastic flows generally follow valleys or other depressions, but can have enough momentum to overtop hills or ridges in their paths. Pyroclastic flows are one kind of sediment gravity flow. Sediment gravity flows are unique fluids because properties such as density and viscosity can change as they3C11 Planetary Geology Volcanism move, unlike fluids such as water or air within which density and viscosity change very little, if at all, during movement. FIGURE 17: Pyroclastic flows 4.7 Phreatic Eruptions Phreatic eruptions are driven by the explosive expansion of volatiles coming into contact with hot rock or magma. The distinguishing feature of phreatic explosions is that they only blast out the gaseous volatiles and fragments of pre-existing solid rock from the volcanic conduit; no new magma is erupted. On Earth and Mars, this process will involve cold groundwater or ice coming into contact with hot magma, resulting in pressurised steam decompressing at the surface, forming an energetic geyser at the surface. It is predicted that at ~1000m beneath to surface of Io, SO 2 could circulate as a phreatic fluid in the crust. In the vicinity of molten magma, the compound would vaporise and blow out onto the surface as a long period geyser. An even more powerful, but shorter term geyser, is produced if hot silicic rock (>1200K) comes into contact with a solid sulphur deposit, vaporising it to high temperatures and venting it out from the surface. Geysers also occur on Triton, but are driven in a slightly different way. The change in temperature due to solar exposure causes nitrogen ice beneath the surface to vaporise, and this then, upon reaching a vent to the surface, decompresses out as a plume. FIGURE 18: Phreatic eruptions - steam geyser Reference List Wilson L., Head J.W., 1983, “A comparison of volcanic eruption processes on Earth, Moon, Mars, Io and Venus”, Nature, 302, 663 Head and Wilson, 1992, "Magma Reservoirs and Neutral Buoyancy Zones on Venus: Implications for the Formation and Evolution of Volcanic Landforms", J. Geophys. Res, 97, 3877-3903 Francis P., 1996, “Volcanoes - A Planetary Perspective”, Oxford University Press Frankel C., 1996, “Volcanoes of the Solar System”, Cambridge University Press Cattermole, P., 1996, “Planetary Volcanism”, Wiley Praxis Publishing3C11 Planetary Geology Volcanism VOLCANISM (4 lectures) 5 Volcanic Landforms 5.1 Flood volcanism Flood volcanism covers large areas of all the terrestrial planets, and are amongst the most important manifestations of basaltic volcanicity. Flood basalts emerge in large volumes, so that they cover wide regions, because they are erupted through fissures, spreading out as extensive flows that pile up on top of one another, forming giant lava fields. The flood lavas are very low viscosity, consisting of low silicate basalts which are erupted very hot. With flood volcanism, volcanic constructs such as shields do not build up, because flood lavas stay so hot and fluid that they follow gentle slopes and spread thinly over large areas. As a result, the source vents of flood lavas are covered, both by their own products and by flows from other sources. Flood volcanism is related to extensional tectonics within the locality, where the stretching and thinning of the lithosphere allows regions of massive melting and the expansion of cracks. This allows easy access to the surface for underlying magma, erupting out through extensive dykes. On Earth, flood volcanism results in phenomena called traps (the Swedish for steps, formed by the multiple layers of lava), which occur on a timescale of once every 20-30 million years. An average trap represents one million years of eruptions within a particular province. However, within this period there are long stretches of quiescence. Some individual trap flows reach 50-100km in length, and several thousand cubic kilometres in volume, making them massive events, that dwarf the largest flows seen on Earth in recorded history by a factor of at least one hundred. Similar types of flow cover extensive regions on Mars, Venus, and especially the Moon. . On the Moon, the mare floods must have had high temperatures, high eruption rates and low viscosities in order to travel such long distances before cooling. Flows in the Imbrium basin, for instance, appear to travel for hundreds of kilometres, the largest extending 1200km into the basin. We know these flows were of low viscosity because the thickness of the flows average just 35m in thickness. Lava flows on the Moon did not occur all at the same time. In some cases they were separated by sufficient time to have significantly different crater densities, implying different exposure ages to space. The most prominent basalt eruptions appear to have occurred between 3.9 and 3.1 billion years ago, as determined from dating of lunar samples. A small amount of volcanism probably occurred as little as 2 billion years ago on the basis of crater density studies, and it is likely that volcanism may have started a lot earlier than 3.9 billion years ago, with the lavas from those eruptions having been covered over by subsequent events. On Venus, over 200 flows of a similar scale to terrestrial flood basalts have been located, larger than 50,000km 2 . No correlation between the varying flow morphology and the total area has been found. 5.1.1 Lava channels and tubes A common occurrence in basaltic lava flows are tunnels and tubes. These form when the surface of a flow crusts over whilst lava continues to flow underneath. When the source is cut off, the lava drains away downslope leaving behind an empty tube. Lava tubes allow a flow to reach much farther than on the surface by dramatically cutting down heat loss. Lava may flow ten times farther through a tunnel than by flowing on the surface. Sinuous rilles are basically meandering channels formed by the action of flowing lava moving downslope. They vary in size from tens of metres to 3km wide, by a few to 300 kilometres long, but their sides remain remarkably parallel throughout their entire length, apart from a slight narrowing further from the vent. Lunar sinuous rilles have often been compared to terrestrial lava channels, representing collapsed lava tubes or open lava channels on a massive scale, often showing much smoother morphologies. As with lava tubes, their origin is within basaltic lava flows, and they are commonly found within basaltic flood lavas, though they often have a higher concentration near the edges of these regions. The width of the rilles implies that high effusion rates were necessary to form the sinuous rilles. Turbulent lava flows are more likely to be capable of causing these features, which in turn would require fairly fluid lavas with low yield strength and low viscosity. Hulme (1982) reviews the possible formation processes of sinuous rilles, as shown in Figure 19. Sinuous rilles often begin in depressed source areas, looking very much like craters. These source depressions may been caused by a lava fountain erupting lava, which builds up around a vent to the point at which the lava begins to flow. The vent vicinity is where the lava is at its hottest and most erosive, and it is this that probably produce the depressions at the source.3C11 Planetary Geology Volcanism FIGURE 19: The formation of sinuous rilles (from Hulme, 1982, Geophys. Sur., 5, 245) On typical slopes at the edge of the Moon’ s mare (~0.004), with the eruption rates necessary for turbulent flow (~10 5 m 3 s -1 , Hulme 1982), the lava flows once it reaches a critical depth (Hulme quotes 20m for his model). With thicknesses and effusion rates such as these, the width of the flow will be several kilometres, and Hulme suggests that a channel within the flow will have a width of 1500m. In his model, the lava in the channel will be turbulent, whilst the rest of the flow is not (Fig 19a). The rille may begin to meander as a result of this turbulence, which can be considered as a series of eddies superimposed on the mean flow (Fig 19b), though of course local topography will also have some influence. While the meandering of the rille begins, the ground has been heating up beneath the flow. At some point, erosion of the bed begins and the channel becomes fixed in position (Fig 19c). Once the channel depth becomes larger than the floor depth, the width of the channel also becomes fixed (Fig 19d). Ground erosion occurs along the entire length of the rille, but less so further from the vent. Once the flow has cooled enough to prevent the flow from proceeding, the rille has reached its final form (Fig 19e). Eventually, the sides of the rille will slump under the influence of gravity, filling the floor with the slumped material and give the rille a V- shaped profile, as opposed to the U-shape of fresher rilles (Fig 19f). As a result of this, the channel rims widen (just as slumping of crater rims during modification will act to widen the transient crater diameter). This is the form we see on the Moon today. Channels of another form found only on Venus are called canali. They often start in an irregular depression and also travel in a sinuous way. Their widths (~3km) remain constant for large portions of the length and are often longer than 1000km, close to ten times the longest lava flow on Earth. In one case, the channel Hildr, the flow reaches 6800km, the length of the Nile. Where the canali are well developed, they may bifurcate, sometimes rejoining, sometimes staying separate. In the freshest form, they may give rise to a large spread similar to the appearance of a flow of lava, the diameters of which can be hundreds of kilometres. Close to the vent, canali seem to have a box-shaped profile with no raised rim, but further away from the vent, they develop a raised rim relative to the surface. The length of these channels is puzzling, since even with a flow rate comparable with the Amazon river, the most fluid lava should have cooled and slowed to a halt only 2000-3000 km from the source. That Hildr reached 6800km testifies to some extraordinary flow conditions, not only in terms of magma output, temperature, and fluidity, but also in terms of heat insulation - implying lava tubes or other forms of thermal protection. An alternative explanation could be that they are formed by carbonate lava, which would have a melting temperature close to that at the surface of Venus, preventing such lavas from cooling effectively, allowing them to travel long distances. FIGURE 20: Morphology of canali3C11 Planetary Geology Volcanism 5.1.2 Wrinkle ridges Prominent ridges, stretching several hundreds of kilometres in length and up to several kilometres wide, are common on all regions of flood volcanism. These ridges have a two-component morphology on the plains, comprising of a broad arch up to 10km across and a superimposed narrower summit ridge, usually offset into roughly parallel segments. When these ridges pass onto highland material, they continue as simple scarps, showing ridges to be related in some way to the underlying topography. Various explanations for these features have been given, both volcanic and tectonic, and it seems that they derive from the interplay of volcanic intrusion and extrusion from fissures, with both extensional and compressive tectonic movements while they grew. FIGURE 21: Cross-section of a wrinkle ridge complex 5.2 Plains volcanism Plains volcanism occurs when extensive basalt eruptions take place at a more modest rate than that seen in flood volcanism. This results in lower discharge rates through narrower fissures, imposing severe cooling and flow constraints, so that only a few individual feeding pipes remain active, as opposed to entire fissures. The discharged lavas, less voluminous and quick cooling, build shield-shaped aprons around central vents - shields - which often overlap, and funnel; their lava flows in the saddles between them. Tube and channel fed lava flows are also widespread, producing compound flows, whose course is determined by the existing topography. The exposed surface of plains volcanism thus represents a complex succession of overlapping extrusions, forming from a number of different centres. FIGURE 22: The multi-layered features of plains volcanism The Moon has a few examples of small-scale plains volcanism, where a number of subdued shields are found on the edge of the Lunar Mare. Averaging 10km in diameter and only a couple of hundred meters in height, these are most likely the result of a drop in magma discharge rates in the final stages of mare flooding, resulting in the closing off of fissures and shortened flows building up around the vents. On Venus, fields of small volcanoes are common, termed as “dome fields” or “shield fields”. These vary in scale significantly, but may contain tens or hundreds of volcanoes, with a density of 4 to 10 per 10 3 km 2 in an area >10 4 km 2 . The fields can have diameters from 50 to 350km and can be broken down into 4 basic classes: I. Simple field. Randomly scattered on the plains unit with no apparent association between the plains and edifices. II. Apron shield fields. Clusters of volcanoes spatially associated with bright or dark volcanic flows. III. Companion shield fields. Spatially associated with a large volcanic centre. IV. Plains units with abundant small shields that have consistent stratigraphic relations with other mappable units.3C11 Planetary Geology Volcanism FIGURE 23: Seamounts in a mid-ocean ridge Plains volcanism is believed to account for the vast majority of the lava flows on Mars, and are also seen on Io and Mercury. While relatively rare on the surface of the Earth, whose prime example is Snake River Plain in Idaho, the formation of plains style volcanism along the mid-ocean ridges are what drives sea floor spreading. The ubiquitous nature of plains volcanism shows that it is possibly the most common form of volcanic style. 5.3 Shield volcanoes FIGURE 24: Profile of a shield volcano Shield volcanoes are characterised as low-profile curved structures, with gently sloping flanks, known as shields. These are built from at least 90% lava, of generally low-viscosity mafic (low silicate) lava composition, so that lava flows from the summit spread out to significant distances from the vent. In this way, a gentle slope is built to the central conduit, but the volcano maintains a low profile. The size of shield volcanoes varies greatly. The smallest shields are just a few kilometres across, have very low slopes, and can stand alone or be associated with plains volcanism. While many are circular, others are elliptical due to lateral flows caused by sub-volcanic rifts. The shield formation process can continue to build volcanoes up to a very large size, through the successive eruption of large volumes of fluid mafic magma, which in larger shields can erupt from the summit, from lower on the flanks or along rift zones that transect the volcanic edifice. Shield volcanoes occur on most terrestrial planets, but their scale varies greatly from planet to planet. The Moon has only small shields produced through plains volcanism, while Earth’ s largest shields are built over hot spot volcanoes, reaching, from a base of several hundred kilometres on the ocean floor, to heights of up to ten kilometres. On Venus, over 160 volcanoes larger than 100km have been identified, with a few having diameters larger than 700km. Their slopes are gentle, rarely greater than a few degrees, but still attain heights of several kilometres. FIGURE 25: The differences in scale of Martian shield volcanoes Shield volcanoes on Mars can grow to a far greater size than on other planets. This is due to the combined presence of a thick lithosphere to support the weight of such volcanoes, and a continuous supply of magma to the surface in the same region for periods of billions of years. This results in volcanoes which include the largest known shield volcano in the solar system, Olympus Mons, with a diameter of ~500km and a summit which lies 20,000m above the surrounding plateau, and 27,000m above the reference level on Mars. 5.3.1 Caldera formation A caldera is a large, usually circular depression at the summit of a volcano formed when magma is withdrawn or erupted from the underlying magma chamber. The removal of large volumes of magma leaves the ’ honeycomb’ like chamber with no mechanical resistance, resulting in a loss of structural support for the overlying rock, thereby leading to the collapse of the ground and formation of a large depression.3C11 Planetary Geology Volcanism Caldera on Earth are generally limited to less than 25 kilometres in diameter, and several kilometres deep, but those on Mars and Venus can exceed 100 kilometres in diameter. Caldera can also form as several overlapping depressions, produced by multiple cycles of magma withdrawal and crater collapse. 5.3.2 Paterae The term patera has been applied to certain very large, low relief volcanic land forms found on Mars, Venus and on Jupiter’ s moon, Io. These volcanoes resemble flattened shields, having a shallow saucer shape (from which they derive their Greek name). One of the best examples of this is Alba Patera on Mars, which has volcanic flows that extend for 1000km from the summit, making it the largest diameter volcanic feature in the solar system. Those paterae on Mars associated with the northern volcanic regions, along with those of Io, follow this form of extended lava flows. There are also paterae within the cratered highland regions of Mars, which share the very flat profile, but were probably formed as collapsed shields built by massive ash eruptions early in the Martian history. 5.4 Volcanic cones Volcanic cones are formed under similar conditions to shield volcanoes, but the lava generally tends to have a higher viscosity, leading to a greater proportion of the eruptive products being expelled explosively. This results in shorter, thicker flows and the ejection of pyroclastic material falling close to the vent, building up the slopes of the volcano into a cone. FIGURE 26: Profile of a cone volcano 5.4.1 Spatter and cinder cones Spatter and cinder cones are steep-sided cones, generally relatively small in size, built from particles and blobs of congealed lava ejected from a single vent. Long-lived basaltic lava fountains erupt either very fluid fragments of molten lava ejected from a vent that flatten and congeal on the ground (spatter), or gas-charged lava that is blown violently into the air, breaking into small fragments that solidify before reaching the ground (cinder). These cones are usually formed over relatively short periods of time, typically weeks to perhaps centuries, and generally vary little during this time in eruptive behaviour. Studies of suspected lunar and Martian spatter and cinder cones indicate that generally they are only one quarter as large as those of Earth. 5.4.2 Stratovolcanoes Stratovolcanoes (also known as composite volcanoes) are large steep-sided, symmetrical cones built of alternating layers of lava flows and pyroclastic material. Usually constructed over a period of tens to hundreds of thousands of years, stratovolcanoes may erupt a variety of magma types, with basalt producing lava flows, and more developed silicic lavas generating pyroclastic material in explosive eruptions. A stratovolcano typically consists of many separate vents, some of which may have erupted cinder cones and domes on the volcano’ s flanks. These are most common on Earth, where magma often progresses to more silicic compositions, though some examples have been seen on Mars. FIGURE 27: Profile of a stratovolcano 5.4.3 Maars Maars are a special case of small cone, produced by high-explosive phreatic activity where a rising column of hot magma intersects with subsurface water. They are generated by single-event, explosive eruptions, converting the water almost instantly to steam. As a result the lava is scattered into tiny particles which are spread out around the vent by base surges around the vent, giving rise to a circular and somewhat broader rim than is usual. While maars are likely to have been formed on Mars, due to the explosive interaction of magma with subsurface ice, maar-type structures have not as yet been recognised on other planets. FIGURE 28: Profile of a maar volcano3C11 Planetary Geology Volcanism 5.4.4 Dark halo craters Dark halo craters, as their name suggests are crater-like forms with a dark halo around them. They are usually found along basin margins or along rilles and lineaments, with rim deposits between 2-10km in diameter. Dark halo craters are formed by explosive ejection of sprays of molten rock, which are deposited around the explosively opened vent or fissure similar in form to a Vulcanian eruption, where gaseous material breaks through a plug and explosively decompresses onto the airless Lunar surface. FIGURE 29: Formation of dark halo craters These dark ring deposits have similarities with pyroclastic rings on Io, both involving silicate pyroclasts accelerated from the vent by the gas until they decouple and continue on ballistic trajectories, affected only by gravity. 5.5 Volcanic domes Volcanic domes are mounds that form when viscous lava is erupted slowly and piles up over the vent, rather than moving away as a lava flow. The sides of most domes are very steep and typically are mantled with unstable rock debris formed during or shortly after dome emplacement. Most domes are composed of highly developed silica-rich lava, and this may result in pressurised gas within the melt exploding during dome extrusion, forming pyroclastic flows. Volcanic domes commonly occur within the crater, near the summit or, in a modified form, on the flanks of large composite volcanoes. Domes are rare on most planets, because of the silicic magma needed to form them, but they are more common on both the Earth and Venus. FIGURE 30: Profile of a dome volcano 5.5.1 Pancake domes While domes are usually relatively small features, surrounding major volcanic land forms, on Venus there is a form of dome called a pancake dome, which is much larger. These have relatively steep sides, are circular with relatively flat or upwardly convex profiles and variable fractures and pits. With diameters in the range of 20-50km and heights of 100m to 1000m, they have volumes of ~100km 3 compared to only ~1km 3 on Earth. If pancake domes are as silicic as domes found on Earth, then they are the first extra- terrestrial example of such developed magmas, but they could just as easily owe their viscosity to low-temperature emplacement of a crystal-rich ’ mush’ , with no need for an exotic magma. 5.5.2 Ticks FIGURE 31: Profile of a tick feature Often referred to as “ticks”, “scalloped-margin”, “fluted” or “modified” domes, these are steep-sided domes, with flanks that are extensively gullied and widely depressed summits. They are probably formed due to the collapse of pancake domes, and some are surrounded by curved fractures, indicating subsidence of the dome following eruptive activity.3C11 Planetary Geology Volcanism 6 Volcanic provinces While volcanoes can occur in relative isolation, they are more usually created as the individual surface expressions of a much larger region of general cooling. As such, across all the terrestrial planets, volcanoes tend to appear in clusters, often resulting in a planetary surface split into distinct regions of differing volcanic output. 6.1 Hotspot volcanism Hot spot volcanism occurs when convection within a planet results in a continually rising plume of heat within the mantle. This plume causes enough melting to produce a replenishing source of magma, which can then allow the repeated eruption of material to the surface. 6.1.1 Island chains FIGURE 32: The Hawaiian/Emperor island chain On Earth, where the ocean bearing crust is almost everywhere in motion relative to the hot mantle beneath, volcanic features formed on the surface are carried away, with volcanoes drifting out of the zone of influence of the rising plume, and eventually shutting down. On drifting plates there is thus a chain of medium-sized shield volcanoes, lined up from the hot spot in the direction of plate drift. The active volcano of the chain is the one growing over the hot spot, and other volcanoes display progressively older ages the farther they are from their place of origin. 6.1.2 Volcanic rises Volcanic rises are the sites of large-scale active mantle upwelling, and their associated volcanism occurs in response to pressure-release partial melting in the underlying mantle. Most volcanic rises show a wide diversity of volcanic features, are significantly uplifted, and tend to cluster along major rift zones, lines of extensional tectonics. FIGURE 33: The extended region of a Volcanic Rise Volcanic rises form a major part of the highlands on both Mars and Venus. They are often associated with broad tectonic junctions and so demonstrate well the interaction between volcanism and tectonism at large scales. Often the distinction between the volcanic and tectonic relief is uncertain, but they do contain obvious centres of volcanism. They are very large scale, often reaching sizes of 1000km to 3000km across, and have large gravitational anomalies, which are perhaps due to dynamic mantle phenomena. 6.1.3 Coronae, arachnoids and novae Coronae are roughly circular features a few hundred km in diameter. They can come in a variety of forms, but in general have raised rims above the surrounding terrain with a pattern of fracturing and ridges associated with it. They are interpreted as surface expressions of thermal plumes rising from the mantle to the near surface, probably causing a doming of the surface. Closely associated with coronae are the smaller arachnoids, which are raised regions which get their name from the extensive ’ web-like’ faults which cross them, and nova volcanoes, which are large volcanoes whose slopes are crossed by extensive radial faults.3C11 Planetary Geology Volcanism FIGURE 34: Morphology of a corona Coronae are smaller and more numerous than volcanic rises, and could represent plumes originating at lesser depths, probably at an upper/lower mantle boundary. Simulations show that a plume reaching a rigid lithosphere will stretch it into a domical uplift, cracked by radial fault, and that with time the hot spot will spread laterally under the lithosphere, cooling and causing the overlying topography to subside. FIGURE 35: Scales of mantle upwelling on Venus Novae, arachnoids and coronae are thus the surface manifestation of developing plumes within the mantle. Matching these features to the ’ spreading drop’ model, begins with nova volcanoes, volcanoes whose radial, star-like pattern of extensional faults, shows the early stages of lithospheric uplift into a dome. Arachnoids are uplifted regions with both radial and concentric faults, illustrating the next stage of plume evolution, when a plume begins to spread outwards, and superimposes concentric strain onto the initially radial fault pattern. The larger the arachnoid, the more concentric features appear to dominate over radial ones, consistent with this size-dependant switch from radial to compressional regimes. The third stage of plume spreading leads to the formation of coronae, with less topographic relief as the plume thins out and cools, leaving a near absence of radial features, and a distinctive ring of compressional ridges. Once convection has switched off, there is a surface collapse above the plume causing a sagging of the domed structure, and resulting in the corona-like form observed. Coronae are often associated with large flood lavas and other features. FIGURE 36: Corona formation process 6.2 Plate volcanism Plate tectonics is the major cooling mechanism for the Earth, and dominates the volcanic output of the planet as a result. New plate material is made at mid-ocean ridges, and old plate material is destroyed as it is subducted beneath adjacent plates. 6.2.1 Rift volcanism Mid-ocean ridges are arguably the most prominent geological feature on Earth, snaking across the sea floor over a total distance of 60,000km. On a regional scale, each few hundred kilometres in length is laterally offset by transform faults, reflecting the rising diapirs that underpin each section. The emplacement of fresh magma at mid-ocean ridges occurs episodically, and during an eruption, magma rises through dykes to the surface, forming regions of plains volcanism. When the eruption comes to an end, the last dregs of magma solidify in the feeder dykes, and further crustal extension is needed for the next eruption, resulting in the spreading outwards of the sea floor. Volcanic rifting does not occur only on the sea floor, with continental rifts showing the same features of linear fissures and flattened cones. They are underlain by3C11 Planetary Geology Volcanism hot spots, and are suggestive of how hot spot activity can lead to continental break-up and the formation of new oceans. Extensive rifting occurs in volcanic rises on both Mars and Venus, but there is no conclusive evidence to suggest that these rifts have progressed to the point where the surrounding plains are being extensively spread outwards, as on the Earth. 6.2.2 Subduction volcanism Subduction volcanism is the most varied, and also the most explosive form of volcanism found on the Earth. The subduction environment occurs where old and dense lithosphere is curved down into the viscous mantle to be melted and ’ recycled’ , carrying sediments, water and crustal rocks with it. This results in strong concentrations of dissolved volatiles, and often highly exotic magmas. It is these viscous, volatile entrained, silicic magmas which form such explosive volcanoes upon reaching the surface. Subduction volcanism is believed to be unique to the Earth. FIGURE 37: Subduction induced volcanism 6.3 Impact associated volcanism There is little evidence to suggest that the formation of impact craters directly produces volcanic activity on any of the planets, but the formation of impact basins do enhance the movement of magma to the surface, thus preferentially increasing volcanic activity in their locality. Large impact basins cause significant upwelling of the underlying mantle material, bringing it closer to the surface. They also often result in massive ring fractures that penetrate deep into the crust, providing an easy escape route to the underlying pressurised magma. This has resulted in large-scale flood lavas on the Moon, producing the Mare, but has also occurred on Mars, resulting in a ring of explosive volcanoes around the edges of the Hellas basin. FIGURE 38: Mare formation Reference List Hulme, G. (1973) Turbulent Lava Flow and the Formation of Lunar Sinuous Rilles. Mod. Geol. 4 pp. 107-117. Cattermole, P., 1996, “Planetary Volcanism”, Wiley Praxis Publishing Frankel C., 1996, “Volcanoes of the Solar System”, Cambridge University Press