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The intent of this book is to introduce you to geochemistry, and through it, paraphrasing Schönbein, reveal the mysteries of our planet. To do this, we must first acquire the tools of the trade. Every trade has a set of tools. Carpenters have their saws and T-squares; plumbers have their torches and wrenches. Physicians have their stethoscopes, accountants their balance sheets, geologists have their hammers, compasses, and maps. Geochemists too have a set of tools. These include not only a variety of physical tools such as analytical instruments, but interpretative tools that allow them to make sense of the data these instruments produce. The first part of this book is intended to familiarize you with the tools of geochemistry. Once we have a firm grip on these tools, we can use them to dissect the Earth in the second part of the book. There, we begin at the beginning, with the formation of the Solar System and the Earth. We then work our way upward through the solid Earth, from core to mantle and crust, and on to the intersection between geochemistry and life: organic geochemistry, the carbon cycle and climate. We'll then examine the processes at the surface of the Earth, first on land, then in the oceans. Finally, we will briefly consider how geochemistry is applied to practical problems: finding resources and addressing pollution.

In filling our geochemical toolbox, we start with the tools of physical chemistry: thermodynamics and kinetics. Thermodynamics is perhaps the most fundamental tool of geochemistry; most other tools are built around this one. For this reason, Chapters 2, 3, and 4 are devoted to thermodynamics. In Chapter 2, we will introduce the laws of thermodynamics and from them develop a most useful tool: the Gibbs free energy. In Chapters 3 and 4, we'll expand our tool set to deal with solutions. These tools allows us to predict the outcome of chemical reactions under a given set of conditions. In geochemistry, we can, for example, predict the sequence of minerals that will crystallize from a magma under given conditions of temperature and pressure or which should replace them in weathering reactions at the Earth's surface. Thus, thermodynamics provides enormous predictive power for the petrologist. Since geologists and geochemists are more often concerned with understanding the past than with predicting the future, this might seem to be a pointless academic exercise. However, we can also use thermodynamics in the reverse sense: given a suite of minerals in a rock, we can use thermodynamics to determine the temperature and pressure conditions under which the rock formed. We can also use it to determine the temperature and composition of water or magma from which minerals crystallized. This sort of information has been invaluable in reconstructing the past and understanding how the Earth has come to its present condition.

Thermodynamics has an important limitation: it is useful only in equilibrium situations. The rate at which chemical systems achieve equilibrium increases exponentially with temperature. Thermodynamics will be most useful at temperatures relevant to the interior of the Earth, but at temperatures relevant to the surface of the Earth, many geochemical systems will not be in equilibrium and instead be governed by kinetics, the subject of Chapter 5. Kinetics deals with the rates and mechanisms of reactions. Reactions can occur only when reactants are brought together. Unlike gas phase reactions or ones within a solution, this often requires the reactants be transported across an interface. So in this chapter, we will also touch on such topics as diffusion and mineral surfaces.

In Chapter 6, we see how tools of physical chemistry are adapted for use in dealing with natural aqueous solutions. Much of the Earth's surface is covered by water, and water usually is present in pores and fractures to considerable depths even on the continents. This water is not pure but is instead a solution formed by interaction with minerals and atmospheric gases. In Chapter 6, we acquire tools that allow us to deal with the interactions among dissolved species and their interactions with the solids with which they come in contact. These interactions include phenomena such as dissolution and precipitation, complexation, adsorption and ion exchange. The tools of aquatic chemistry are essential to understanding processes such as weathering and precipitation of sedimentary minerals, as well as dealing with environmental problems.

In Chapter 7, we move on to trace element geochemistry. In this chapter we will see that trace elements, which comprise most of the periodic table, have provided remarkable insights into the origin and behavior of magmas. Without question, their value to geochemists far outweighs their abundance. There are several reasons for this. Their concentrations vary much more than do those of the 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 chemical reservoirs whose evolution we can partly reconstruct. Mid-ocean ridge basalts, easily the most voluminous on the planet, come from a shallow reservoir from which melt has previously been extracted to form the crust. Oceanic island and other basalts produced by melting of mantle plumes rising from the deep mantle clearly derive from different reservoirs. Although these too shows evidence of previous melt extraction, they have been reenriched in the elements lost. Furthermore, stable isotope ratios in these basalts demonstrate conclusively that they contain material once at the surface of the Earth. This is truly remarkable: the surface and deep Earth are connected by a grand geochemical cycle.

The core, as we noted earlier, consists of iron-nickel alloy. You might ask how can we be confident about the composition of something we have never sampled and have no prospect of ever sampling? The answer is again the geophysical constraints, which tell us that the core is very dense, and the composition of chondrites, which tell us that the only elements of sufficient abundance and density to form the core are iron and nickel. That conclusion is reinforced by iron meteorites, most of which are cores of asteroids. There is a problem, however; namely, that any combination of iron and nickel will be denser than the Earth's core at relevant temperatures and pressures. These elements must thus be diluted with perhaps some 5% or so of one or more lighter elements. The meteorite inventory of what was available and the isotopic composition of some of the candidate light elements, such as silicon, helps us narrow the possibilities, but we do not yet have a firm answer. Experiments showing how elements partition between silicate and iron liquids together with thermodynamics places important constraints on what is possible. Comparing the composition of the mantle with that of chondritic meteorites show that the mantle is highly depleted in elements, including the most valuable metals such as platinum and gold, that we expect to partition into iron liquid and since this partitioning is temperature and pressure dependent, we can begin to develop scenarios on how the core formed.

Then we turned to the crust, first the oceanic crust, then to the continental crust. The first question is its composition, an easier one to answer than the composition of mantle and core. It is not an easy task, however, given how heterogeneous the crust is and while the surface is easily sampled, the lower crust is not. Nevertheless, we continue to build on the work of Clark and Goldschmidt and refine estimates of crustal composition. Then we turn our attention to how the crust formed. We can certainly establish that the continental crust has formed through partial melting of the mantle, but in what tectonic environment under what circumstances, and when? Did it form early in Earth's history, steadily through time or in pulses, or perhaps only recently? And how permanent is it? We know a lot, but we're still struggling to completely answer these questions.

Life is, of course, ubiquitous at the surface of the Earth and has modified the planet in remarkable ways: life is a geologic force. Organisms produce a vast array of chemicals that find their way into the physical environment. As we noted, modern geochemistry differs from what Schönbein envisioned in that it encompasses organic as well as inorganic matter, and these organic substances are ubiquitous at the surface of the Earth. This is the subject to which we turn in Chapter 12. After briefly exploring the nature and structure of organic compounds and the role they play in life, we'll survey their presence in soils and natural waters. Once outside a cell, organic substances are subject to attack by microbes and begin to degrade almost immediately. Yet some can survive on millennial time scales and longer. An emerging paradigm emphasizes the importance of adsorption of mineral surfaces in resisting degradation. The ability of dissolved organic molecules to adsorb complex inorganic substances is important: it retains nutrients in soil and maintains otherwise insoluble metals in solution. Some of these long surviving molecules, or at least their hydrocarbon skeletons, can be associated with specific biomolecules. Some of these biomarkers, or chemical fossils, are restricted to specific groups of organisms and can thus help us reconstruct past environments and biological evolution. Others have proved useful in reconstructing past atmospheric CO₂ levels and paleotemperatures.

Organic substances are an important part of the carbon cycle. Photosynthesis and subsequent sequestration of organic matter in sedimentary rocks transformed the Earth's initial CO₂-rich atmosphere to one containing free oxygen, which first occurred 2.3 billion years ago in the Great Oxidation Event. For the next billion and a half years, some atmospheric oxygen was present, but not enough to support metazoans (animals). Then around 600 million years ago, atmospheric oxygen levels began to rise again and just at this time the first animals appear in the fossil record. But as oxygen was produced, atmospheric CO₂ was drawn down. As a greenhouse gas, CO₂ plays a critically important role governing climate and the times oxygen rose in the atmosphere were accompanied by glaciations in the Proterozoic and Paleozoic.

This was not the cause of the Pleistocene glaciations, however. Stable isotope studies demonstrated that glacial-interglacial cycles correlated with small changes in the Earth's orbit and rotation (the Milankovitch variations). These were the pacemaker of the Pleistocene glacial cycles, but it was shuffling of CO₂ between the atmosphere and oceans that actually caused the climate swings.

Burial of organic carbon in sediments has also produced the coal and petroleum that have provided the energy to power the global economy since the Industrial Revolution. We'll examine the processes that transform this buried organic matter into these energy resources. But in burning fossil fuels we are increasing atmospheric CO₂, which, not surprisingly, is warming the planet and initiating a host of other climate changes.

In Chapter 13, we will focus attention on the Critical Zone, which is the land surface from the top of vegetation to the bottom of circulating groundwater. It is so called because essentially all terrestrial life lives within it and ultimately all life, including marine life, depends on processes occurring within this zone. It is here that rock comes in contact with water and air, and primary minerals are replaced by new ones. These weathering reactions produce soil and release nutrients that make terrestrial life possible. Some fraction of these nutrients is carried to the oceans by streams and rivers and make marine life possible. Life is an integral part of the weathering and soil development process, as organic acids help to break down rock and movements of metals complexed by organic molecules contribute to the development of distinct soil horizons over time.

Weathering of silicate rocks is another important part of the carbon cycle and consequently influences climate on time scales of tens to hundreds of millions of years. This is because carbonic acid produced by dissolution of CO₂ in water provides most of the acidity necessary to drive weathering reactions. The result is a solution enriched in calcium and bicarbonate, which is then carried to the oceans by streams and rivers to be precipitated as carbonate sediment, thus removing CO₂ from the atmosphere until it is again released by metamorphism or volcanism to the atmosphere as CO₂. Over Earth's history, there has been a net transfer of CO₂ from the atmosphere to sedimentary carbonate, keeping Earth's surface temperature within the habitable range even as the Sun has grown steadily brighter. We'll examine weathering reactions and their rates from the perspective of field studies. We'll find that lithology, climate and hydrology, topography, and the biota all exert important controls on weathering rates. We'll then turn our attention to the composition of streams and rivers and see how these same factors control the composition of streams and rivers. Finally, we look at the composition of saline lakes and see how the process of fractional crystallization leads to a great diversity of their compositions.

In Chapter 14, we follow the rivers to where they lead: the oceans. The oceans are salty and alkaline because, as Anton Lavoisier put it, they are “the rinsings of the Earth,” that is, they contain the weathering products of the land surface. Just six components, Na⁺, Mg²⁺, Ca²⁺, K⁺, Cl^–, and , make up 99.3% of the dissolved solids, and these are always present in the same proportions and in the same proportion to the total, the salinity, which is about 35 parts per thousand by weight on average. A final component, , brings the total to 99.7%.

Ultimately, the concentrations of all components in seawater are controlled both by the rates at which they are added from sources and the rates at which they are removed by sinks. Rivers are the major source of most elements in seawater, but the atmosphere is the major source for dissolved gases as well as a few metals such as Al, Pb, and Th, which reach the ocean in wind-blown dust. Ridge crust hydrothermal activity is an important source of some elements, but it is also an important sink for others. Sediments are the major sink for most elements, and half the ocean floor is covered by the carbonate and siliceous shells of planktonic organisms. Evaporites are the major sink for Na⁺, K⁺, Cl^–, and , but these form discontinuously through time. The last major evaporite deposit formed in the Mediterranean when tectonics closed the Strait of Gibraltar between 6 and 5.3 million years ago. The Mediterranean dried up nearly entirely, and the resulting drop in base level allowed rivers running into it, such as the Rhone and Nile, to cut channels 1000 m below their present levels, which subsequently filled with sediment when the Gibraltar connection reopened. The vast, thick beds of salt deposited beneath the Mediterranean during this time were enough to decrease global ocean salinity by 5%.

Biological processes exert an extremely important influence on ocean chemistry. Unlike the other major components, the concentration of varies, mainly due to photosynthesis and respiration, although calcium carbonate precipitation and CO₂ exchange with the atmosphere also contribute to variations. Photosynthesis is restricted by light availability to the upper hundred meters or so, while respiration occurs throughout the ocean. Temperature and salinity establish a strong density gradient in the ocean that has the effect of limiting exchange between this photic zone and the deep ocean. Once it is cooled at high latitudes and sinks into the deep ocean, water remains there on time scales of ∼1000 years. Sinking organic remains can fall through the water column and this density barrier and can be remineralized through respiration in the deep water. This transports dissolved CO₂ from the surface to this deep water where it builds up, a phenomenon known as the biological pump. Consequently is present in higher concentration in deep water, which also results in a decrease in pH from ∼8.1 in the surface water to ∼7.6 in the deep water. Partly as a result of this variation in pH, the ocean becomes undersaturated with respect to calcium carbonate with depth so that carbonate shells formed in the surface water tend to dissolve of depth and do so completely below a depth at ∼4500 m. Falling carbonate shells also contribute to the biological pump, and as we noted above, this is also part of the long-term carbon cycle controlling climate. On much shorter time scales, changes in ocean circulation and biological productivity changed the efficiency of this biological pump between glacial and interglacial periods, resulting in a transfer of CO₂ from the atmosphere to the deep ocean, very much amplifying the Milankovitch climate signal.

Unlike the major elements, concentrations of most minor and trace elements are quite variable in the oceans and much of this variation is due to biologic activity that imposes vertical concentration gradients, as these elements are taken up by phytoplankton in the surface water and released by respiration in the deep water. This includes not only nutrients such as P, Si, and Fe, but also nonutilized elements such as Ge because organisms take them up incidentally. A few elements, such as Al and Pb, show the opposite pattern: enrichment in the surface water and depletion in deep water because wind-deposited dust is the primary source of these elements and they are quickly scavenged onto particle surfaces after deposition.

In the final chapter we see how geochemistry can be used to address the needs of society, specifically, its need for mineral resources and environmental protection. The story of civilization is in some respects the story of increasingly sophisticated tools. The Stone Age ended when people learned to produce copper metal from copper sulfide ores around 7000 years ago. Copper tools were subsequently replaced by bronze ones and then by iron ones beginning around 3000 years ago. In a sense, we still live in the Copper and Iron Ages, however, as 21 million tons of copper ore and 2.5 billion tons of iron ore were mined globally in 2018. In the United States, about half the demand for metals is met by recycling, but modern society still need enormous amounts. Furthermore, modern technology requires a great variety of metals, many of which were unknown as recently as two centuries ago. At least 80 different elements are incorporated in smartphones or used in their production, including exotic ones like neodymium, europium, and tantalum. Two other exotic elements, cadmium and tellurium, are used to produce CdTe solar panels, which have the highest efficiency and can be produced in thinner films than other solar cells.

We'll discuss the process of geochemical exploration and consider examples of the formation of a variety of ore deposit types. The first of these is the Bushveld complex of South Africa, which is an example of orthomagmatic ores, in which the ore had precipitated directly from magma. The Bushveld, which outcrops over an area the size of Ireland, is a layered mafic intrusion that formed 2 billion years ago and hosts the world's largest reserves of platinum group elements, Cr, and V. Decades of geochemical detective work have shown that these ores formed as fractional crystallization combined with repeated intrusions of magma and assimilation of surrounding crust periodically saturated the magma in ore-forming minerals, including chromite, magnetite, and sulfides that settled out of the magma chamber to formed distinct bands. In contrast, hydromagmatic ores such as porphyry copper deposits, which are the primary source of copper ore, form when a saline aqueous fluid exsolves from a magma and intrudes, often with violent force, into surrounding rock. Laboratory experiments together with analysis of fluid inclusions in these ores have revealed that many metals, including Cu, Zn, Pb, Co, Sn, and Au, form highly soluble chloride and sulfide complexes in these fluids at elevated temperatures and partition into the fluid phase from the magma, then precipitate when the solution cools. These form mainly from subduction-related magmas because they are rich in water and oxidizing; the latter prevents premature precipitation from the magma of the ore metals as sulfides. Many tin deposits form in a similar way but the magmas are produced by melting of Sn-rich sediments within the crust and reducing conditions allow Sn concentrations to build up through fractional crystallization and Sn is often complexed by F rather than Cl.

Hydrothermal ores also precipitate from aqueous solution and chloride complexes are also important in transporting metals in these deposits. The fluid, however, is derived from seawater or formation brines within the crust. These types of deposits include volcanogenic massive sulfides (VMS); mid-ocean ridge hydrothermal systems are actively forming examples of this type of deposit. The ore-forming fluids can be directly sampled and their chemistry determined; study of these systems has provided much insight into how VMS deposits form. Seawater is warmed as it penetrates the hot, young ocean crust and a series of reactions result in the solution becoming acidic and reducing. Under these conditions, metals, most notably Cu, Zn, and Pb, are leached from the rock. When temperatures reach 350–400°C, the fluid rises, eventually mixing with seawater whereupon the metals precipitate as sulfides.

We'll examine two examples of sedimentary ore deposits. The first is banded iron formations, which are the principal source of iron ore. Most of these formed around the time the atmosphere first became oxidizing about 2.3−2.4 billion years ago as ferrous iron-rich deep ocean water upwelled to the surface and the iron was oxidized to the insoluble ferric form. Directly or indirectly, the evolution of photosynthetic life appears responsible for them. Brines associated with saline lakes and salars, or salt flats, and their associated brines, particularly from the high plateaus of the Andes and Tibet, are becoming the most important source of lithium, which is needed for high performance batteries in everything from cell phones to electric cars. But not all such brines are Li-rich; we learn the conditions under which Li-rich brines form. Weathering-related ore deposits include bauxite, the ore of Al, and laterites, which are sources of Ni, Fe, and rare earths. These form through extreme weathering of soils such that little remains but these highly insoluble elements; what we have learned about weathering, soil-forming processes, and the geochemistry of these metals will serve us well understand how these deposits form. Because of their importance to everything from flat panel computer displays and televisions to high performance magnets in wind turbines, electric cars, and speakers, we briefly examine rare earth ore deposits, which fall into many of the above categories.

Finally, we put our geochemical toolbox to use to understand how human activities can degrade environmental quality and how this can be addressed. Like ore deposits, this is an enormous topic and we have space to consider only a few examples. We begin with the problem of eutrophication and associated anoxia in fresh water lakes, using Lake Erie, one of the Great Lakes of North America, as an example. Eutrophication refers to situations where nutrient levels in water allow excessive growth of algae, usually cyanobacteria, which produce microcystin toxins. Lakes typically become temperature-stratified in summer such that oxygen in the deep water is not replenished. Bacteria consuming the remains of algae falling into the deep water can consume all available oxygen leading to anoxic conditions in the deep water and consequent fish kills. Persistent eutrophication in Lake Erie was successfully addressed by regulations in the 1970s that severely limited nutrients from sewage, industrial effluents and particulate phosphorus in agricultural runoff and the lake was restored to health. In the late 1990s eutrophication is summer began to occasionally reoccur due to dissolved phosphorus from agricultural runoff. Solving the problem will require further modification of farming practices.

Toxic metals are another important environmental problem. One source is mining of sulfide deposits, such as the several types described above. Sulfides exposed to water and atmospheric oxygen quickly weather to produce sulfuric acid, resulting in a problem known as acid mine drainage. Not only is the acidity a problem with pH values as low as 2, but under these conditions many otherwise insoluble toxic metals become soluble. The solution is certainly not to simply shut down mines as when pumps are shut off, water penetrates in mine shafts, pits, and tailings ponds and the problem worsens. Indeed, the bigger problem is old, abandoned mines as a number of strategies are deployed in modern mining operations to prevent the problem. Lead and mercury are highly toxic metals and anthropogenic release of these elements to the atmosphere has polluted the entire surface of the planet. Lead, however, is an example of an environmental success story largely due to the efforts of one geochemist, Claire Patterson. Regulations that eliminated Pb from gasoline and emissions from smelters have dramatically reduced the amount of Pb in the environment. Regulations have also starkly reduced emissions of Hg, at least in developed countries, and local sources of extreme pollution, such as in Minamata, Japan, where mercury poisoning killed over 1700 people and disabled many more, have been eliminated in most cases. Nevertheless, levels in the atmosphere, soils, plants, the ocean, and many fish species remain high and will decrease only slowly in the future, even if all emissions are eliminated. An understanding of the unique geochemistry of Hg will enable us to understand why.

Finally, we examine the problem of acid rain. This results from burning of fossil fuels, particularly coal, which oxidizes sulfur and nitrogen ultimately to sulfuric and nitric acid, although use of nitrogen fertilizers also contributes. This can lower pH in rain to values as low as 4. Depending on the nature of the soil and bedrock this may or may not be a problem, and the understanding of weathering reactions we gained in earlier chapters will help understand why. In areas where soils have developed through weathering of rocks with low acid neutralizing capacity, the low pH alone can have deleterious effects on trees, fish, and aquatic invertebrates, but that is not the main problem. Instead, the principal problems are loss of cations such as Ca²⁺ and aluminum toxicity. Aluminum is one of the most abundant elements in the Earth's crust, yet natural Al toxicity is rare. Once we understand the geochemistry of Al, we'll be able to understand why this is usually not an issue but can be when rain is acidic. Acid rain is another environmental success story, although a still unfolding one. Regulations have greatly reduced emissions in the developed world, but it will take decades before soils and stream chemistry returns to natural levels and for damaged ecosystems to heal.