Introduction

Flood lavas, by definition, cover vast areas in great sheets of lava, without the construction of major edifices (e.g., Geikie, 1880; Washington, 1922; Tyrrell, 1937; Self et al., 1997). The flat terrain that flood lavas produce has led to the term “plateau volcanism” to be used as a synonym for flood volcanism. In addition, the classic erosion pattern of flood lavas leaves a series of topographic steps. Thus many flood basalt provinces are known as “traps” from the Scandinavian word for steps. Plateau volcanism transitions to “plains” volcanism when low shields become common (Greeley and King, 1977). It is not surprising that these large-volume eruptions are usually composed of the most common of volcanic rocks: basalt. Thus, the term “flood basalt” is often used interchangeably with “flood volcanism.” However, there can be interesting and significant compositional variability within flood “basalt” provinces. The most general term to describe all large-volume volcanism is “Large Igneous Province” (LIP) (e.g., Coffin and Eldholm, 1994).

LIPs represent a major geologic event with significant repercussions on the interior of a planetary body. The extraction of such large volumes of magma can alter the thermal state of the mantle, indicate major changes in the convection patterns within the mantle, and lead to geochemical evolution of the mantle on a regional scale (e.g., Coffin and Eldholm, 1994 and references therein). Flood lavas also alter the face of a planet for geologically significant time.

Introduction

Playas (dry lakes) are a type of lacustrine system that are dry most of the time, and can be flooded only occasionally. A playa environment, despite its dry conditions, is characterized by an active hydrological cycle. This is evidenced by a wide range of hydrogeological processes operating today or in the recent past. Therefore, playas are a fundamentally different environment from dry desiccated deserts, and identification of playas on Mars has significant implications for the planet's hydroclimatic history.

Mars currently is dominated by a hyperarid environment. Today, water appears to exist abundantly in the Martian polar caps, and also in the surrounding high-latitude regions, but as near-surface ice (Boynton et al., 2002), not liquid water. Whether or not there are localities with recent active hydrogeological processes is uncertain. However, there may have been sites of stable lakes (deep-water lakes) in the past. Such sites would have changed to playa environments, owing to the decline in the water budget, and eventually desiccated completely. Photogeologic surveys have identified possible paleoshorelines in the northern plains (Parker et al., 1989) and crater lakes (Cabrol and Grin, 1999; Ori et al., 2000a; Malin and Edgett, 2003). If these features are in fact paleoshorelines, it would necessarily imply that conditions suitable for stable oceans and lakes must have existed at some point in Mars' history. Ice-covered paleolakes could have also existed, and their shoreline geomorphology could differ from that of paleolakes without ice cover.

Introduction

A series of astrobiological high-altitude expeditions to the South American Andean Mountains were initiated in 2002 to explore the highest perennial lakes on Earth, including several volcanic crater lakes at or above 6000 m in elevation. During the next five years, they will provide the first integrated long-term astrobiological characterization and monitoring of lacustrine environments and their biology at such an altitude. These extreme lakes are natural laboratories that provide the field data, currently missing above 4000 m, to complete our understanding of terrestrial lakes and biota. Research is being performed on the effects of UV in low-altitude lakes and models of UV flux over time have been developed (Cockell, 2000). The lakes showing a high content of dissolved organic material (DOM) shield organisms from UV effects (McKenzie et al., 1999; Rae et al., 2000). DOM acts as a natural sunscreen by influencing water transparency, and therefore is a determinant of photic zone depth (Reche et al., 2000). In sparsely vegetated alpine areas, lakes tend to be clearer and offer less protection from UV to organisms living in the water. Transparent water, combined with high UV irradiance may maximize the penetration and effect of UV radiation as shown for organisms in alpine lakes (e.g., Vincent et al., 1984; Vinebrook and Leavitt, 1996). Shallow-water benthic communities in these lakes are particularly sensitive to UV radiation. Periphyton, which defines communities of microorganisms in bodies of water, can live on various susbtrates.

Introduction

The arid climate, extensional rift setting, range in type and age of volcanic eruptions, and generally widespread and geologically youthful volcanism in New Mexico contribute to an environment rich in geologic processes and landforms analogous to many of those on Mars. Young (<5 Ma) volcanoes and associated volcanic rocks are more widely distributed throughout the state than in many other volcanic localities on the North American continent. All of the principal volcanic landforms occur including long lava flows, viscous domes, calderas, composite volcanoes, monogenetic scoria cones, small shield volcanoes, and numerous hydromagmatic vents. The morphologies, volcanic emplacement processes, and dissected structures, and the arid environment, result in many volcanic landforms analogous to those on Mars. These features provide some clues to the details of geologic processes responsible for their Martian counterparts that are uncommon in areas where volcanism is less abundant and where the environments are less arid.

The largest young caldera (Valles Caldera), largest young lava flows (McCartys and Carrizozo), abundance of Quaternary volcanic fields, volatile-rich magmatism, including non-juvenile (maars) and juvenile types (Shiprock-Narbona Pass), spring deposits, and one of the great modern rift valleys on Earth (Rio Grande rift) occur in an arid setting where annual precipitation is between 8 and 15 inches (20–40 cm) per year. Combined with arid dissection and eolian in-fill, these contribute to a landscape that mimics the appearance of many volcanic terrains on Mars.

Introduction: Martian gullies and terrestrial debris flows

The observation of small gullies on Mars was one of the more unexpected discoveries of the Mars Observer Camera (MOC) aboard the Mars Global Surveyor (MGS) spacecraft (Malin and Edgett, 2000). Gullies are the flutes and narrow troughs formed by the debris flows process and not the process itself. They mostly occur in a latitudinal band higher than 30°. The upper parts of the slopes (mostly south facing slopes in the southern hemisphere) exhibit alcoves, with generally broad and deep channels. They are characterized by their distinct V-shaped channels with well-defined levees. Individual channels exhibit low sinuosity and deep erosion down to the fans that bury the lower parts of the crater walls (Figure 10.1). These debris fans correspond to one or several lobes.

The characteristics of these gullies suggest that they were formed by flowing water mixed with soil and rocks transported by these flows. They appeared to be surprisingly young, as if they had formed in the last few million years or even more recently. In their initial analysis, the MGS Camera investigators Mike Malin and Ken Edgett (2000) proposed a scenario involving ground water seepage from a sub-surface liquid water reservoir located a few hundred meters or less below the surface. However, the process capable of maintaining such a shallow aquifer at temperatures above the freezing point of water remains unclear.

Introduction

The origin of fluvial valleys has been one of the great problems of science ever since the debates in the eighteenth and early nineteenth centuries over the role of cataclysmic events in the shaping of Earth's valleys. The famous founders of geology, including Hutton, Lyell, and Cuvier all participated in these great debates. It can be argued that geology emerged as a science because of the scientific reasoning that was applied to this problem (Davies, 1969).

It was one of the great surprises of modern planetary science that fluvial valleys were discovered on the planet Mars by study of the vidicon images returned by the Mariner 9 spacecraft (Masursky, 1973). From the much more extensive coverage of the Viking mission we know that the heavily cratered Martian highlands are locally dissected by networks of tributary valleys with widths of 10 km or less and lengths of a few kilometers to nearly 1000 km. These valley networks are one major type of large elongate Martian troughs thought to have fluvial origins (Baker, 1982).

In distinction from fluvial valleys, large outflow channels show extensive evidence of large-scale fluid flow on their floors and walls (Baker et al., 1992a), so technically speaking they are channels, rather than valleys. This is a rather curious circumstance, since Earth experience leads us to suppose that fluvial channels are invariably much smaller than fluvial valleys. However, the Martian outflow channels may arise from the peculiar geological history of Mars and its water endowment (Baker et al. 1991).

Introduction

Solidified lava flow morphologies are a consequence of complex interactions between the moving, cooling lava and its environment. Because no active Martian lava flow has been observed, eruption and emplacement parameters must be determined from the resulting volcanic morphologies. Griffiths and Fink (1992a, b) demonstrated the effects that ambient conditions exert on the gross morphology of lava flows with Newtonian rheologies. Through the use of analog experiments, they concluded that typical lava flow morphologies are created by a balance between the rate at which heat is advected within the flow and the cooling rate – a ratio they quantified with the dimensionless parameter Ψ (Fink and Griffiths, 1990). Gregg and Fink (2000) examined the effect of underlying slope on lava flow morphologies, and concluded that increasing slope has a similar effect to increasing effusion rate. However, Gregg and Smith (2003) show that this relationship breaks down somewhat on slopes steeper than about 20°. Griffiths and Fink (1997) and Fink and Griffiths (1998) examined the effect of ambient conditions on laboratory flows with a Bingham rheology, and observed a similar dependence of morphology with Ψ.

Thus, the main parameters that appear to control lava flow morphologies for lavas with Newtonian or Bingham rheologies are effusion rate, eruption temperature, lava viscosity, underlying slope, and ambient conditions (e.g., Fink and Griffiths, 1990, 1998; Gregg and Fink, 2000).

Introduction

For more than a quarter of a century, the spectacular grabens of Canyonlands National Park, Utah, have provided planetologists with a fundamental analog for understanding what planetary grabens should look like and – more importantly – what may be implied about the depth variation of mechanical properties and horizontal extensional strain.

The seminal work on Canyonlands grabens was done by George McGill and coworkers in support of their investigations of the origin and kinematic significance of lunar and Martian straight rilles (McGill, 1971; McGill and Stromquist, 1975, 1979; Stromquist, 1976; Wise, 1976). McGill and Stromquist (1979) hoped to invert graben widths, assessed on an aerial or orbital image, for the depth of faulting (i.e., fault intersection depth). By equating this depth with stratigraphic layer thickness and assuming a symmetric graben geometry and plausible values of fault dip angles, grabens provided ready and seemingly reliable probes of the near-surface planetary stratigraphy and strain. Interestingly, the analog modeling of brittle-layer extension over a ductile (quasiplastic) substrate, appropriate to Canyonlands stratigraphy (McGill and Stromquist, 1975, 1979), anticipated the key role of faulting in triggering and mobilizing salt or shale diapirism at depth (Jackson and Vendeville, 1994; Jackson, 1995). Other observations and inferences made in the 1970s, including flexure of rock layers at ramps near graben terminations and incremental growth of fault slip (McGill and Stromquist, 1979), anticipated these fundamentally important ideas by at least a decade (Sibson, 1989; Peacock and Sanderson, 1991; Cowie and Scholz, 1992).

Introduction

The structure and morphology of Martian calderas have been well studied through analysis of the Viking Orbiter images (e.g., Mouginis-Mark, 1981; Wood, 1984; Mouginis-Mark and Robinson, 1992; Crumpler et al., 1996), and provide important information on the evolution and eruptive styles of the parent volcanoes. Using Viking data it has been possible, for numerous calderas, to define the sequence of collapse events, identify locations of intra-caldera activity, and recognize post-eruption deformation for several calderas. Inferences about the geometry and depth of the magma chamber and intrusions beneath the summit of the volcano can also be made from image data (Zuber and Mouginis-Mark, 1992; Scott and Wilson, 1999). In at least one case, Olympus Mons, analysis of compressional and extensional features indicates that, when active, the magma chamber was located within the edifice (i.e., at an elevation above the surrounding terrain). The summit areas of Olympus and Ascraeus Montes provide evidence of a dynamic history, with deep calderas showing signs of having been full at one time to the point that lava flows spilled over the caldera rim (Mouginis-Mark, 1981). Similarly, shallow calderas contain evidence that they were once deeper (e.g., the western caldera of Alba Patera; Crumpler et al., 1996). Some of the best evidence for circumferential vents on Mars can be found on Pavonis Mons, where several sinuous rilles can be identified that must have originated from vents close to the rim (Zimbelman and Edgett, 1992).

This part of the book is intended to act as a guide to the basic technological principles that are specific to landers, penetrators and atmospheric-entry probes, and to act as a pointer towards more detailed technical works. The chapters of this part aim to give the reader an overview of the problems and solutions associated with each sub-system/flight phase, without going into the minutiae.

Overview and fundamentals

The descent through the atmosphere is often the only part of a planetary probe mission, as for example the Pioneer Venus and Galileo probes; on other missions it is just the last stage of a long journey prior to surface operations. The key parameters are the altitude of deployment – usually the altitude at which the vehicle ends its entry phase, as defined by some Mach number threshold – and the required duration of descent.

The duration of descent for an atmospheric probe is often dictated by an external constraint on the mission duration, such as the visibility window of a flyby spacecraft that is to act as a communications relay. This imposes an upper limit on the descent duration – it may be that (as for the Huygens probe) some part of that mission window is desired to be spent on the surface.

The instantaneous rate of descent (and thus the total duration) is determined at steady state by the balance between weight and drag. The former is simply mass times gravity; the latter depends on ambient air density, the drag area of the vehicle and any drag-enhancement device such as a parachute or ballute. The drag area is usually expressed as a reference area and a drag coefficient. Often these parameters and the mass are lumped together into the so-called ballistic coefficient β.

Often the dynamic pressure of descent is used to force ambient air into sampling instruments such as gas chromatographs.

The transfer of material that is not native to a planet has been happening over the history of the Solar System, with meteorite delivery being a common example of this interchange. With the development of rocket launchers capable of injecting objects into interplanetary trajectories, mankind joined Nature in being able to alter another planet's composition. Generally spacecraft and their associated hardware are designed and assembled so as to minimize the amount of debris that they carry. This chapter examines the problems associated with the unintentional delivery of living or dead organic matter to celestial bodies; so-called ‘forward contamination’. The topic is often referred to by the phrase planetary protection, and its scope includes not only the possible contamination of planetary bodies, but also the potential introduction to the Earth of materal from a non-terrestrial biosphere. Furthermore, the threat that planetary protection seeks to minimize is not restricted to the introduction of non-native organisms to another planetary body. Non-living material, such as DNA fragments and other complex bio-relevant molecules might trigger false-positives from equipment designed to detect extant or extinct life.

A practical definition of a living entity might be that the agent processes matter and energy in such a way that it can reproduce, and in doing so prosper in the face of environmental stresses. If the environment of the organism changes too radically then the organism may be killed or rendered dormant.

While much can be achieved by purely passive observations and measurements of a planetary lander's immediate environment, some key science requires the landed system to interact with the surface mechanically. This may involve the acquisition of samples of material, either to be returned to Earth or delivered to instrumentation internal to the lander. Other instruments, while external, require intimate contact with target rocks – these include alpha-X-ray, X-ray fluorescence or Mössbauer spectrometers, and microscopes. Other interactions may include mechanical-properties investigations using a penetrometer, or current measurements of wheel-drive motors.

Thus a variety of mechanisms have been operated on planetary surfaces, including deployment devices and sampling arms of various types, together with drills, abrasion tools and instrumentation. Soviet/Russian landers have tended to feature simple but robust actuators, usually simple hinged arms, and often actuated by pyro or spring. These include the penetrometers on the Luna and Venera missions. Lunokhods 1 and 2 carried a cone-vane shear penetrometer that was lowered into the lunar regolith and rotated by a motor, to measure bearing strength and shear strength. The rovers made 500 and 740 such measurements, respectively, during their traverses across the lunar surface.

A more sophisticated arm was flown on the Surveyor 3, 4 and 7 lunar landers (Figure 12.1). The Surveyor soil mechanics surface sampler (SMSS) was a tubular aluminium pantograph, five segments long, with a total reach of 1.5 m.

Following the success of the Mars Pathfinder project in 1997, there was a resurgence of interest in the deployment of an untethered rover on the surface of Mars. The concept of a semi-autonomous and freely roving vehicle was mooted as a follow-on to the Viking missions of the late 1970s. Almost twenty years were to pass before a rover was to be operated on Mars. After the Mars Pathfinder mission, NASA had proposed to send a rover equipped with a geology/chemistry payload, dubbed the ‘Athena’ suite, to Mars in 2001. Various constraints led to the redesign of the mission for a 2003 launch, although experiments of the payload were carried on the ill-fated Mars Polar Lander. In 2000 the Mars Exploration Rover mission was selected, with a launch-date flight three years later. This time, the Athena payload was to be duplicated, carried on two identical 174 kg rovers. Designated MER-A and MER-B, the spacecraft carrying the rovers were launched to Mars on separate Delta 2 boosters, making use of the favourable 2003 window for low-energy trajectories. The rovers on each craft were targeted to different regions of Mars. The MER-A craft, carrying the ‘Spirit’ rover, arrived on 4 January 2004 and was directed toward Gusev crater (14.5°S, 175.5°E) in the Aeolis region of Mars. This crater is the terminus of the fluid-cut Ma'adim Vallis, and Gusev was thought to host geological clues to the presence of water on Mars.

Payload delivery penetrators are bullet-shaped vehicles designed to penetrate a surface and emplace experiments at some depth. The basic technology for these has existed for several decades based largely on military heritage (e.g. Simmons, 1977; Murphy et al., 1981a; Bogdanov et al., 1988), however only in the mid 1990s did proposals for their use in Solar System exploration begin to be adopted for actual flight. In the US, Mars penetrators were studied for several years (and, indeed, field tested) as part of a possible post-Viking mission, while in the Soviet Union planetary penetrator work seems to have started in the mid 1980s.

Impact speeds range from about 60 to 300 m s−1. The resulting impact load experienced by penetrators as they decelerate in geological materials routinely exceeds 500 g, and terrestrial systems in the military field can be rated at 10 000 g or even 100 000 g, although the choice of components at these levels is severely limited (being more suited to the relatively simple job of triggering a detonator than making planetary science measurements). Additional impact damping may be included in the form of crushable material (e.g. honeycomb or solid rocket motor casing), sacrificial ‘cavitator’ spikes protruding ahead of the penetrator's tip (e.g. Luna-Glob high-speed penetrator concept, with speeds exceeding 1.5 km s−1) and gas-filled cavities (e.g. the Mars 96 penetrators).

Masses have ranged from the tiny DS-2 Mars Microprobes at 2.5 kg each (excluding aeroshell) to 62.5 kg each for the Mars 96 penetrators.