Introduction

The frequency of animal life in the Universe must be some function of how often it arises, and then how long it survives after evolving. Both of these factors may be significantly influenced by the frequency and intensity of mass extinctions, brief intervals of time when significant proportions of a planet's biota are killed off. They are killed by one or some combination of too much heat or cold, not enough food or nutrients, too little (or too much) water, oxygen, or carbon dioxide, excess radiation, incorrect acidity in the environment, or environmental toxins. Based on the history of Earth's life, mass extinctions seemingly have the potential to end life on any planet where it has arisen.

Mass extinctions do more than threaten biota. They may also play a large part in evolutionary novelty. On Earth there have been about 15 such episodes during the last 500 Myr, five of which eliminated more than half of all species then inhabiting our planet. These events significantly affected the evolutionary history of Earth's biota: for example, if the dinosaurs had not been suddenly killed off following a comet collision with the Earth 65 Ma, there probably would not have been an Age of Mammals, since mammals were held in evolutionary check so long as dinosaurs existed. The wholesale evolution of mammalian diversity took place only after the dinosaurs were swept from the scene. Mass extinctions are thus both instigators as well as foils to evolution and innovation.

By the end of the Galileo mission, many characteristics of Io's volcanism had been identified for the first time. Processes were observed that continue to yield insight into the way volcanic activity helped shape the Earth and other terrestrial planets. This chapter summarizes the post-Galileo view of Io.

Volcanism and crustal structure

Galileo revealed a world dominated by silicate volcanism, putting to rest a major point of contention that dated back to Voyager. Galileo discovered many active volcanic centers on Io (Appendix 1), considerably more than were detected by Voyager. The global distribution of activity and volcanic features favors the heating of Io by tidal dissipation taking place, in large part, in the aesthenosphere – the partially molten upper mantle.

The upper few kilometers of Io's crust appear to be mostly mafic silicates interbedded with deposits of sulphur, sulphur dioxide, and silicate pyroclasts. The thickness of this volatile-rich layer probably varies from location to location but provides plenty of material that is easily mobilized by thermal interaction with mafic magma. Ultramafic magmas may be present, but the wide error bars ascribed to the data do not show this conclusively. If recent data analyses (Keszthelyi et al., 2004a, 2005a) are confirmed, ultramafic volcanism may not be present or widespread, and the presence of a global magma ocean (e.g., Keszthelyi et al., 1999) may not be required.

After the Voyager encounters, analysis of the data continued. More hot spots were identified in IRIS data (see McEwen et al., 1996; also see Appendix 1). Early post-Voyager models of Io incorporated a sulphur-rich crust with an ocean of sulphur beneath (e.g., Sagan, 1979; Smith et al., 1979b, 1979c). The sulphur was kept liquid and mobile through interaction with hot or molten silicates within the crust (e.g., Lunine and Stevenson, 1985). This model was supported by the abundance of sulphur compounds on Io's surface and sulphur in the Io plasma torus and neutral clouds, and by the fact that silicates had not been detected on the surface. The occasional eruption of silicate lava was still possible, but Io's volcanism was dominated by sulphur lavas.

Silicate volcanism on Io?

The case for ubiquitous silicate activity was made by Carr (1986), who argued on the basis of crustal strength that the “sulphur ocean” postulated to lie close to the surface was unlikely, based on the strength of the crust needed to support the steep walls of many paterae, where temperatures in excess of 650 K would literally melt the caldera walls. Heat transport from Io's interior to the surface was via silicate volcanism. To support the observed mountains, Io required a predominantly silicate lithosphere at least as thick as the mountains were high (>15 km in some cases).

I have always been fascinated by volcanoes, and especially by Io, a tiny moon that beyond any expectation turned out to be the most volcanically active body in the Solar System. Now that the NASA Galileo mission is over and initial data analyses have been completed, this is an appropriate time to assess the “state of the satellite” and review what has been learned about Io over the past few decades.

A fascination with volcanoes is understandable, but I am also inspired to understand, through modeling of volcanic processes, how volcanoes work. Such motivation was instilled in me as a post-graduate student by Lionel Wilson and Harry Pinkerton at Lancaster University in the UK.

In this book, therefore, I have endeavored not only to describe what Galileo saw, but also to provide the necessary background for understanding the physical, volcanological processes taking place on Io, and to demonstrate how remote-sensing data of volcanic activity can be used to peel back the layers of a planet to reveal interior processes and structure. To put the majestic scale of volcanism on Io into proper context, comparison is made wherever possible with volcanic activity on Earth.

It has taken nearly two years to write this book. Along the way, I have had a great deal of help from friends, family, and colleagues.

This chapter reviews the history of Io observations up to and including the Voyager encounters. The material in this chapter is drawn primarily from Satellites of Jupiter (Morrison, 1982) and Satellites (Burns and Matthews, 1986), both published by the University of Arizona Press, and Time-Variable Phenomena in the Jovian System, NASA Special Publication 494 (Belton et al., 1989).

Io before Voyager

The study of Io dates from the very beginning of telescope-based astronomy. In his observation notes for January 7, 1610, Galileo Galilei wrote: “when I was viewing the heavenly bodies with a spyglass, Jupiter presented itself to me; and because I had prepared a very excellent instrument for myself I perceived that beside the planet there were three little stars, small indeed, but very bright.” Subsequent observations revealed a fourth “little star.”

Thus were discovered the Galilean satellites, named by Simon Marius (a contemporary of Galileo) Io, Europa, Ganymede, and Callisto. Subsequently, little attention was paid to the satellite system except as a means of measuring the speed of light (Roemer's method). Not until the nineteenth century did physical observations become important. In 1805, Laplace used the orbital resonant properties to estimate satellite masses. New refracting telescopes at Lick and Yerkes measured satellite sizes, and bulk densities were obtained.

This table of Io's active hot spots has been compiled from identifications of hot spots in data from Voyager IRIS (Pearl and Sinton, 1982; McEwen et al., 1992, 1996); SSI (McEwen et al., 1997, 1998a; Keszthelyi et al., 2001a; Geissler, 2003; Turtle et al., 2004); NIMS (Lopes-Gautier et al., 1997, 1999, 2000; Lopes et al., 2001, 2004); PPR (Spencer et al., 2000b; Rathbun et al., 2004); and terrestrial telescopes (Goguen et al., 1998; Marchis et al., 2001; de Pater et al., 2004; Marchis et al., 2005). Hot spots in bold type have temperatures in excess of 1000 K as determined from SSI data (summarized in Geissler [2003]), NIMS data (Davies et al., 1997, 2000b; Lopes-Gautier et al., 2000; Davies et al., 2001, 2003b), and ground-based instruments (Marchis et al., 2002). This list will no doubt expand as more data are obtained and analyzed. Provisional names for features are in italics. Detections are indexed as follows: I = Voyager IRIS; S = Galileo SSI; N = Galileo NIMS; P = Galileo PPR; E = Earth-based or Earth-orbiting telescopes; C = Cassini ISS. All of the hot spots are plotted in Plates 13b and 16a.

The positions of hot spots are subject to uncertainties caused by different observation resolutions and movement of thermal sources, as well as by reference to nearest feature rather than the hot spot itself.

Chapters 6 and 7 describe models that are used to quantify thermal emission from volcanic processes and the subsequent use of those results to quantify the movement of magma from the interior of a planet to the surface. This chapter demonstrates how different types of eruptions generate different thermal emission spectra, characteristic “thermal signatures” of eruption style, and how those spectra change with time. Determination of eruption style is important because it allows application of the appropriate model to determine effusion rate. Observed temporal behavior allows eruption mechanism to be constrained.

Because most ionian volcanic thermal-emission data are at resolutions rendering the entire emitting area sub-pixel, it is necessary to understand how eruption style affects the shape of the integrated thermal emission spectrum. Such analyses are dependent on, first, adequate spectral coverage and, second, temporal coverage. Spectral resolution need not be particularly high so long as data at appropriate wavelengths are available. Two wavelengths (2 μm and 5 μm are particularly effective) can be used to constrain the silicate volcanism mode because the resulting 2-μm to 5-μm ratio (2:5-μm ratio) is very sensitive to the changing surface temperatures of lava surfaces of ages ranging from seconds to months, sometimes to years (Davies and Keszthelyi, 2005). Table 8.1 shows the timescales of different volcanic processes and the instruments best suited to observe them.

Volcanism has shaped the surface of the terrestrial planets. Earth, the Moon, Venus, and Mars have all been heavily modified by volcanism during at least some part of their history. Active high-temperature volcanism, with magma at temperatures in excess of 1000 K issuing onto the surface, however, has been observed on only two planetary bodies – Earth and Io (compared in Table 4.1). This high-temperature volcanism is clear evidence of the triumph of interior heating processes over planetary heat loss mechanisms. Interior heating has melted at least part of the planetary mantle to form silicate magmas. The forces primarily responsible for heating within these two bodies, however, are very different.

Why, then, are Earth and Io currently volcanically active? What are the origins of the heat that is generating and being lost through volcanism, and why is Io unique in the jovian system in having high-temperature volcanism? How have these bodies evolved over time?

Global heat flow

First and foremost, it is necessary to quantify surface heat flow in order to model heating and heat transport mechanisms. Determination of heat flow from remote-sensing data is complex, being dependent on the thermal properties of surface and subsurface materials. In Io's case, these properties and materials are not known with any certainty. In comparison, measurement of heat flow on Earth is easily accomplished with in situ measurements.

Using Galileo data, volcanic activity at individual volcanic centers was monitored and quantified on timescales of weeks to months, for a period of seven years. This monitoring was a huge step beyond the two snapshots of activity, four months apart, obtained by the Voyager spacecraft. Galileo's most exciting discovery was the large number of active volcanoes on Io, of which at least 50 had temperatures >1000 K (identified in Appendix 1). Having identified the locations of these high-temperature lavas, similar locations (e.g., dark-floored paterae; McEwen et al., 1985) were inferred also to be the result of silicate volcanism (see Chapter 9). By the end of the Galileo mission, more than 160 active hot spots had been identified from Voyager, Galileo, and Earth-based telescope data (see lists in Lopes et al. [2001, 2004]). In NIMS regional observations around Prometheus, where spatial resolution improved by a factor of 10 (Plate 7), the number of detected hot spots increased by a factor of ≈ 3 (Lopes et al., 2004), confirming earlier predictions that these small hot spots were present (Blaney et al., 1998, 2000). PPR detected many more low-temperature thermal sources not detectable by SSI or NIMS (Rathbun et al., 2004). Extrapolating to a global scale suggested as many as 300 large and small active hot spots on Io (Lopes et al., 2001), although the region covered by the aforementioned NIMS observation has a particularly high density of volcanic features (Schenk et al., 2001).

Because of its persistent red deposits, Pele is the most distinctive of Io's volcanoes (Plates 9a, b). Appropriately named after the Hawaiian goddess of volcanoes, Pele is the source of a giant plume more than 300 km high, as seen by Voyager (Strom et al., 1981). The reddish deposits laid down by the plume are ≈ 1200 km across and are rich in sulphur and sulphur dioxide (Geissler et al., 1999; Spencer et al., 2000a). Closer to the vent, dark pyroclastic material, most likely of silicate composition, streaks the surface (Strom et al., 1981; Geissler et al., 1999). SSI and the Hubble Space Telescope (Plate 9f) have shown that the Pele plume can exceed 400 km in height (Spencer et al., 1997b; McEwen et al., 1998a).

Voyager observed dramatic changes in the shape of the plume deposits between encounters when the deposits changed from a “cloven hoof” shape (almost certainly caused by a vent obstruction of some kind) to a circular appearance (Figures 1.3d, e). The change in the shape of the plume deposits took place in the four months between Voyager encounters, so the resurfacing for some period of time was relatively rapid. Although the Pele plume was not seen by Voyager 2, it may have changed to a more tenuous, gas-rich form – a “stealth” plume (Johnson et al., 1995).