more abundant elements, and their behavior tends often to be simpler and easier to treat than that of major elements (a property we will come to know as Henry's law). Geochemists have developed special tools for dealing with trace elements; the objective of Chapter 7 is to become familiar with them.
Chapters 8 and 9 are devoted to isotope geochemistry. In Chapter 8, we learn that radioactive decay adds the important element of time; radioactivity is nature's clock because the rate at which a radioactive nuclide decays is absolutely constant and independent of all external influences. We can read this clock by measuring the build-up of radiogenic daughter elements, for example 206Pb produced by decay of 238U. In this way we have established the age of the Solar System and the continents, and we have placed firm ages alongside the relative geologic time scale developed in the nineteenth century. Importantly, radiogenic isotope geochemistry has provided some perspective on the rate and manner of evolution of the Earth, and the evolution of our own species by answering questions such as how old are those bones and when were those cave paintings done? We can also use the products of radioactive decay, radiogenic elements, as tracers. By following these tracers much as we would dye in a fish tank, we can follow the evolution of a magma, the convection pattern of the mantle, and the circulation of the oceans, and determine from where sediments were derived. Radiogenic isotopes allow us to distinguish magmas produced by melting of the crust from those produced by melting of the mantle and to distinguish a number of distinct chemical reservoirs in the mantle; for example, magmas erupted by oceanic island volcanoes come from different reservoirs than those erupted at mid-ocean ridges.
The isotopes of another set of elements vary not because of radioactive decay, but because of subtle differences in their chemical behavior. These “stable isotopes” are the subject of Chapter 9. The subtle differences in isotopic abundances of elements such as H, C, N, O, and S has provided insights into the evolution of the Earth's atmosphere and reveal the diets of ancient peoples. One of the most significant contributions of stable isotope geochemistry has been to establish temperature changes of the oceans associated with the Pleistocene glacial cycles. Together with radiogenic isotope geochronology, the timing of these cycles and demonstrated that the ultimate driver of them has been small variations in the Earth's orbit and rotation, known as Milankovitch variations. Stable isotope geochemistry is the last of our geochemical tools.
With our toolbox full, we are ready to examine the Earth from the geochemical perspective in the second part of the book. Where else to start but at the beginning? The Earth today is the product of its long history, and of all the events in that history, none set the stage more for what Earth would become than its formation.
In Chapter 10 we'll begin by looking at “the big picture”: the cosmos and the Solar System. The cosmic beginning was some 13.8 billion years ago. The Big Bang, time's opening act, produced a universe of hydrogen and helium and very little else. Only once stars and galaxies had formed, perhaps half a billion years later, did the universe begin to be seeded with heavier elements. Stars the size of the Sun and larger synthesize the principal elements of life, carbon, nitrogen, and oxygen in their geriatric “red giant” phase and blow them back out into the cosmos in enormous stellar winds. That, however, is not enough to create a planet like Earth, or support life for that matter, both of which require heavier elements as well such as magnesium, silicon, phosphorous, and iron. These are synthesized during the death throes of giant stars and expelled into the cosmos in spectacular explosions called supernovae, which can radiate more energy than an entire galaxy.
Some 9.5 billion years later, part of a vast cloud of gas and dust, not unlike the Great Nebula in Orion visible in the northern hemisphere night sky in winter, began to collapse in on itself, spinning ever more rapidly as it did so like a skater pulling in her arms. The Sun formed in the center of this swirling mass and planets formed in the surrounding disk. The idea that the Solar System formed in this way is an old one: Immanuel Kant postulated it in 1755. But what are the details? We'll find that the details are revealed in leftovers from the process: chondritic meteorites. These meteorites consist of aggregations the dust from which the solar system is formed, although some were metamorphosed in their asteroidal parent bodies. Among other things, they reveal that this nebula, at least in the inner part, was so hot that almost all the dust had evaporated to gas. The first materials to condense, so-called calcium–aluminum inclusions, have been dated with exquisite precision by the decay of U to Pb at 4568.22 ± 0.17 million years. These meteorites also once contained short-lived radioactive nuclides that must have been synthesized within a million years or less of solar system formation, products of nucleosynthesis in our galactic neighborhood. The decay of these radionuclides resulted in the build-up of their daughter products, and we can put our tools of isotope geochemistry to good use to see how this can be used to produce a chronology of events in the young solar system. As samples of the solar system nebular dust, these meteorites provide an inventory of the elements available to build the Earth and in this way place important constraints on the Earth's composition.
The Earth differs in its composition from chondrites, mainly because the region in which it formed was too hot for the more volatile elements to condense. We'll put our thermodynamic tools to good use in understanding the sequence in which the elements condensed from the nebular gas. Other meteorites, the achondrites and irons, come from larger asteroids that broke apart when they collided with each other before they had become full-fledged planets. Remarkably, they had already differentiated into iron cores and silicate mantles within a few million years of the start of the solar system – we know this from radiogenic isotope geochemistry. These meteorites thus provide insights into the process of planetary formation. The chronology established by those short-lived radionuclides reveal that formation of the Earth was a much more drawn-out process that continued for tens of millions of years before a cataclysmic collision between Earth and a Mars-sized body produced the Moon. The abundance of certain trace elements in the Earth's mantle tell us, however, that a bit more material must have accreted to the Earth after that.
In Chapter 11, we turn our attention to solid Earth, its composition and its differentiation into layers: crust, mantle, and core. One question we would like to answer is what is the composition of the Earth? Since the mantle is the largest and most massive of these layers, we begin there. Because we only rarely find mantle material at the surface, geophysical observations such as seismic waves, gravity, and moment of inertia provide particularly critical constraints on the nature of the mantle. The composition of blocks of mantle occasionally thrust to the surface, small pieces of it carried to the surface in volcanic eruptions, and the composition of magma produced by melting of the mantle are also critical in constraining its nature. Finally, chondritic meteorites provide the inventory of materials available to form the Earth; while the Earth certainly differs in its composition from chondrites, we need to relate terrestrial composition to that of chondrites in a way consistent with thermodynamics and what we know of the behavior of the elements. We'll find that most modern estimates converge on an estimate of terrestrial composition within a few percent for the most abundant elements. For trace elements, estimates diverge by 25% or so, but that is nevertheless remarkable, considering how heterogeneous the Earth is.
The heterogeneous nature of the mantle comes into full focus when we examine the trace element and isotopic composition of basalts. Basalts are our most abundant mantle sample, but as partial melts they are not compositionally representative except for isotope ratios. The understanding of trace element behavior in partial melting and fractional crystallization we gain in Chapter 7 nevertheless allows us to constrain mantle trace element compositions. What we find from combining trace elements and isotope ratios is that the mantle consists of identifiable
