The Earth is a planetary oasis in the cold near-vacuum of space. The lifeless Moon is 30 Earth-diameters away (approximately 385,000 km). Mars—the nearest planet—is 9,000 Earth-diameters away (at its closest) and there is effectively no exchange of matter or energy between the Earth and either of these bodies—except as described in Box 1.1. In an atomic sense, virtually everything that’s here now has been here for all of the last 3 billion years of Earth’s existence (see Box 1.2), and relatively little matter enters or leaves. In other words, the Earth is a closed system with respect to matter: effectively nothing is added or taken away. However, the Earth is not a closed system in terms of energy. We receive a significant amount of radiant energy from the sun, and of course, without that, we could not live here. We also emit some radiant energy into space, both as reflected sunlight, and as radiated infrared light from the warm surfaces of the Earth.

The iconic 1972 Apollo 17 photograph of Earth, taken from about 1/12th of the way to the moon (Figure 1.1.1), has been reproduced more than almost any other photo, and has had a strong effect on billions of people. It is the first ever photo that provides a fully illuminated view of the Earth from space (as opposed to a crescent, for example), and it’s the last such photo taken by a human (the NASA moon-landing program was cancelled after Apollo 17). The image of a brightly sun-lit, and vigourously active and “alive” Earth, surrounded by the apparent empty blackness of space, highlights the extraordinary isolation, uniqueness and significant fragility of our planet. In the 1970s and 80s this image played a role in the growth of environmentalism, and also in the development of the concept of the Earth as a system.

Figure 1.1.1 The Earth, From the Apollo 17 Spacecraft en Route to the Moon on December 7, 1972.

There are some important features to observe in that 1972 image, such as the dark blue of the vast oceans, the dry sandy Sahara in northern Africa, the frozen expanse of Antarctica, and the swirling clouds obscuring the green equatorial region of Africa. Sunlight is reflecting brightly from high-albedo surfaces (ice and clouds), moderately from lower-albedo surfaces (desert sand and vegetation) and very little from the low-albedo surface of the oceans. In this image it is almost possible to visualize the interactions taking place between the different components of the Earth.

Box 1.2 Incoming!

The Earth formed from the accumulation of rocky debris that existed within the radius of our orbit from the sun. That started with the Early Bombardment, from around 4.6 Ga to about 4.1 Ga. During that time there was ongoing accumulation of material to make up the Earth’s mass. There was also one massive collision with a planet about the size of Mars (or about one-eighth the volume of Earth) (Figure 1.1.2) at about 4.5 Ga. Some of the material from that collision was blasted into Earth orbit and eventually coalesced to form the Moon. The Early Bombardment was followed by the Late Heavy Bombardment, which peaked at around 3.9 Ga, and gradually decreased in intensity until around 3.0 Ga.

Figure 1.1.2 Depiction of an Interplanetary Impact Similar to the Moon-Forming Impact With the Early Earth.

The present rate of accumulation of extra-terrestrial material on Earth is insignificant. Each year we receive around 14,000 tonnes of material from space, mostly as dust-sized particles. That is equivalent to about 1/400,000,000,000,000,000,000th of Earth’s mass per year, or about 1/400,000,000,000th of Earth’s mass over the past billion years.

Media Attributions

• Figure 1.1.1 NASA public domain image of Earth seen from Apollo, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:The_Earth_seen_from_Apollo_17.jpg
• Figure 1.1.2 NASA public domain image, Planetary Smash-up, https://www.nasa.gov/multimedia/imagegallery/image_feature_1454.html

Earth System Science is based on the premise that on our planet everything is connected to everything else via a complex web of energy and matter transfers. There is exchange of matter between rocky material of the geosphere and the hydrosphere, the biosphere and the atmosphere, and all of those systems are also connected with each other. This concept of interconnectivity dates back to 1926 with the publication (in Russian) of Vladimir Vernadsky’s “The Biosphere,” but it only really started to develop and become well known in the west in the 1970s through the work and writing of British scientist James Lovelock. The term Earth System Science was coined in 1983. In 2001 a number of organizations met to draft a declaration of Earth System Science, which is duplicated here in Box 1.3. The important parts of the 2001 Amsterdam Declaration are:

• The Earth behaves as an isolated system of parts connected by transfers of matter and energy and by feedbacks,
• That system is subject to change, and that change can be human caused,
• Such changes can cross critical thresholds leading the Earth System to start operating in a different way, and
• There is no precedent or analogue for the present high rate of change in the Earth System.

An overview of some of the interactions that take place in the Earth System is given on Figure 1.2.1. It is important to point out that this highly-simplified diagram shows only a small fraction of the important interactions, and that it excludes both the energy input from the sun and the role of humans. More on those later.

Figure 1.2.1 A Depiction of Some of the Natural Interactions That Take Place Within the Earth System (see text for an explanation of the labelled interactions).

The single and double arrows of Figure 1.2.1 represent transfer and interchange of matter between the different parts of the system, but they don’t come close to revealing the complexity of the system because they should be shown as a series of intricate paths, twisting in and out, connecting each part of the system to almost every other part. For example, some of the water that originates by evaporation from the ocean rains down onto the ground, is modified by interactions with rock, and then is taken up by plants. Some of the chemicals from that water eventually come down as leaves and seeds and other plant parts and are then washed into the ocean where the chemicals are transferred into fish or marine plants. A fish swims upriver and is caught and partially eaten by a bear, the remains of the fish—which includes components from the ocean and the stream—fertilize the nearby trees. We could keep going on from there, but you probably get the point.

The labelled interactions of Figure 1.2.1 can be summarized as follows:

1. During a volcanic eruption, gases (including H₂O, CO₂ and SO₂) from the geosphere mix with the atmosphere, while volcanic rock fragments (ash) remains suspended in the atmosphere for a short time (weeks).
2. Atmospheric gases (especially CO₂, H₂O and O₂) react with rocks of the geosphere during weathering (an example is provided below).
3. There is exchange of gases between the biosphere and the atmosphere as a result of photosynthesis in which CO₂ in the atmosphere combines with H₂O to produce glucose (C₆H₁₂O₆) and O₂, and respiration, in which glucose combines with O₂ to produce CO₂ and H₂O and energy.
4. There is exchange of water via evaporation and condensation, and of liquid water as rain, between the hydrosphere and the atmosphere.
5. Components of the biosphere (e.g., plants) extract chemicals (e.g., phosphorous, potassium, magnesium, sulphur and calcium) from the geosphere, and in turn deposit material that gets incorporated into the geosphere.
6. There are two-way chemical transfers between the rocks of the geosphere and the surface water of the hydrosphere (for example, some elements in rocks dissolve into surface water during weathering, and then may be deposited from surface water to form new rocks).
7. There are two-way chemical transfers (including elements such as calcium, potassium, sodium, sulphur and carbon) between the rocks of the geosphere and the groundwater of the hydrosphere.
8. There are two-way chemical transfers between groundwater and the biosphere.
9. Components of the biosphere (especially carbon-bearing organic matter) and the hydrosphere accumulate as sediments (geosphere) on the floor of the ocean (and lakes).

These are just a few of the important interactions—there are many others—and it is critical to remember that all components of the Earth System are inter-related. No part of the natural system will operate on its own without input of matter and/or energy from some other part, and each component of the system has the potential to significantly influence the conditions in other components.

A good example of that interdependence is the relationship between the biosphere and the atmosphere. Photosynthetic plants (on both land and in the water) exchange gases with the atmosphere via photosynthesis and respiration as noted above. They cannot survive without atmospheric carbon dioxide and water, and in turn they influence the levels of atmospheric gases. As we’ll see in Chapter 3, photosynthesizers have been changing the composition of the atmosphere for over three billion years and are largely responsible for keeping our planet habitable.

Most of the interactions shown on Figure 1.2.1 involve movement of water. It is important to have some idea of the volumes of water that move through the different reservoirs, as shown on Figure 1.2.2. The majority of the evaporation (87% of it) is from the oceans, and the rest is from land and that includes water produced by plant transpiration. The majority of the precipitation falls onto the oceans (79%), while the rest falls onto land. The 8% difference between 87% and 79% is made up by river runoff from the land to the oceans and groundwater flow into the oceans. Most of the precipitation that falls on the land (88%) falls in “external runoff” areas where it can flow back to the oceans (in rivers and groundwater). The remaining 12% falls in dry areas of the continents—areas of “internal runoff”—and it is returned to the cycle only by evaporation.

Figure 1.2.2 The Volumes of Water Transported per Year Within the Hydrological Cycle.

The water that is in the various reservoirs illustrated on Figure 1.2.2 stays in place for differing lengths of time, as summarized on Figure 1.2.3. For example, water in vegetation, the atmosphere, and rivers, typically stays there for days to weeks (although it can also be shorter or longer); for soil and lakes it may be years to decades; for glacial ice and the ocean it may be centuries to millennia; and for groundwater it might be as little as weeks, but it could also be decades to thousands of years, and even millions of years for some deep groundwater.

Figure 1.2.3 The Typical Residence Times of Water in the Various Reservoirs of the Earth. The solid bars represent likely residence times. The hatched parts indicate potential longer residence times.

Exercise 1.1 Identifying Earth-System Interactions

Figure 1.2.4 shows Rubble Creek and its surroundings (near to Mt. Garibaldi, British Columbia). Identify as many of the interactions shown in Figure 1.2.1 (and listed below that figure) as you can.

Figure 1.2.4 Rubble Creek, British Columbia

Exercise answers are provided in Appendix 2.

The sun, of course, is vitally important to the existence of life on Earth, but it also has a key role in many other Earth System processes. Some examples of the role of solar energy in Earth Systems are summarized on Figure 1.2.5. They include:

• Contributing to freeze-thaw and salt weathering by heating rock, and to weathering related to repeated expansion and contraction due to solar heating,
• Heating Earth surfaces (rock, soil, water, vegetation) which in turn emit infrared radiation, which adds heat to the atmosphere (via the greenhouse effect),
• Providing the radiative energy required for photosynthesis,
• Driving the hydrological cycle by evaporating water from lakes, oceans and vegetation, and
• Heating the near-surface ocean water, which generates wind (and sometimes massive storms as described in Chapter 14) and is a key driver for ocean currents.
  
  

  Figure 1.2.5 Some Aspects of the Role of the Sun Within the Earth System

Figure 1.2.6 A Cross-Section of the Upper Part of the Earth’s Interior Showing Geological Processes Associated with Mantle Convection, Important to Earth Systems

Earth’s systems are also influenced by what is happening deep within the Earth. Because of the significant size of the Earth, there is still a lot of heat left over from the collisions that took place during accumulation of fragments in its initial formation. Radioactive decay of elements like potassium, uranium, and thorium within the core also continue to contribute heat, and so the core is still very hot (up to about 7000⁰C—making it hotter than the surface of the sun) and partly liquid. This heat is slowly transferred from the inner core to the outer core and then from the core to the mantle, and that process leads to convection in the mantle (just like heat from your stove will cause convection in a pot of soup). In contrast, the Moon and Mars are much smaller than the Earth, and these bodies have lost enough heat to become too cold for continuing mantle convection. They don’t have plate tectonics, so mountain ranges are not being created, and they don’t have active volcanoes.

As shown in Figure 1.2.6, convection in the mantle is the key driver of plate motions in the crust and uppermost mantle (although not the only one), and plate motions are associated with a number of features that have implications for Earth System processes. One of these is volcanism, which takes place in several different settings related to plate boundaries. Volcanism creates new crustal rock of course, but it also contributes gases to the atmosphere and water to the hydrosphere. Over geological time, volcanism has led to significant changes in the Earth’s atmosphere, and it may be the source of most of the water in the oceans. A large proportion of those gases and water come from the mantle, and some part of that is recycled back into the mantle during subduction, along with sediments that accumulated on the sea floor.

Plate boundary processes lead to the formation of mountains, both where continents collide (as India is colliding with Asia) and where there is subduction-related volcanism (like that in southern British Columbia, Washington, Oregon and northern California). Mountains are important in controlling many Earth Systems (e.g., weather patterns and biological processes). They also influence the composition of the atmosphere because steep slopes lead to accelerated weathering of rocks, and weathering of rocks consumes carbon dioxide through the process described below.

water + carbon dioxidecarbonic acid

Here we have water (e.g., as rain) plus carbon dioxide in the atmosphere, combining to form carbonic acid. The carbonic acid then combines with feldspar (the most abundant mineral in the crust) to form clay minerals along with dissolved ions through a process called hydrolysis. This process (which is illustrated in Box 3.1 of Chapter 3) can be written like this:

feldspar +  carbonic acid + oxygen kaolinite + calcium ions + carbonate ions

The kaolinite is relatively easily eroded by water and most ends up on the ocean floor, while the dissolved calcium and carbonate ions are also transported to the ocean where they will eventually be combined, with help from marine organisms, to form calcite.

Continents are moved by plate tectonic processes, and as described in Chapter 2, that can have implications for how much of the sun’s energy is converted to heat here on Earth. Ocean basins are also changed by plate tectonics, and that can affect the flow of ocean currents, a process which is critical to the redistribution of heat on the Earth.
Some of the Earth’s internal heat does reach the surface. The obvious manifestation of that is the heat from volcanic eruptions (Figure 1.2.7), but there is also a very slow flow of heat towards surface in non-volcanic areas. Contrary to what you might think, the warming of Earth’s surface is almost exclusively the work of the sun. The contribution of heat from inside is tiny, approximately 0.03% of the amount of heat that we get from the sun.

Finally, Earth Systems are significantly and increasingly influenced by human activities. A few of the relevant changes that we have made and activities that we undertake are illustrated on Figure 1.2.8. Those shown can be summarized as follows:

• Destruction of forests and other natural vegetation,
• Use of cleared land for crops and livestock,
• Use of chemicals (both fertilizers and pesticides) in agriculture,
• Use of cleared land for buildings, roads, parking lots and airports,
• Exploitation of petroleum and other fossil fuels and use of fossil fuels for transportation, electricity generation, farming, heating and manufacturing,
• Unsustainable and wasteful exploitation of marine food resources,
• Indiscriminate dumping of human and industrial wastes onto land and into water and the air,
• Disruption of natural surface water bodies (including drainage in some cases and flooding in others), and
• Changes to natural slopes to permit construction of buildings and roads.

Figure 1.2.7 Examples of Changes Made to the Earth System by Humans.

Another way to think about Earth System Science would be to consider any natural or man-made object on the surface of the Earth. Every natural object, say a bush, a bear, a boulder, or a body of water is obviously interacting (exchanging matter and energy) with its surroundings (water, rock, air, organisms). Every man-made object is doing the same thing, but in different ways. For example, an old car will interact with water and the air by rusting, and will release chemicals (e.g., dissolved iron and iron oxide minerals) into the environment. A concrete wall can be worn away by water or broken by a tree root, and can be chemically changed by oxygen, carbonic acid, and by plants. It can also release chemicals (e.g., calcium ions, clay minerals) into the environment. A piece of glass or plastic will eventually break down and start to fall apart, although that will take decades to centuries.
If you’re interested in how human constructions eventually succumb to Earth’s systems you might want to have a look at Alan Weisman’s book The World Without Us.

Field Trip 1.1

Time to get outside! Go for a walk anywhere near to where you live and find a man-made object that has been around for at least a few decades. It could be a concrete sidewalk, a road, a building, a gravestone, a statue, an abandoned vehicle, an old wharf, or a sea wall. (See Figure 1.2.9 for example)

Figure 1.2.9 An old tractor (Author photo, CC BY 4.0)

Try to estimate how long it has been there (there could be a date on it) and think about how it has been interacting with the oxygen and carbon dioxide of the atmosphere, the hydrosphere (rain, for example), the biosphere, and how it has been affected by the sun. As it starts to break down, how will the materials released affect the soil, the water, the air and the biosphere?

The soil around the tractor in the photo above is bare of vegetation. Can you think of a reason why?

A useful review paper on Earth System Science was published in 2020 by Steffen and co-authors. They conclude with a statement on the importance of understanding Earth Systems now more than ever, because, in light of our changing climate, “humans are now the dominant force driving the trajectory of the Earth System.”

Media Attributions

• Figure 1.2.1 Steven Earle, CC BY 4.0
• Figure 1.2.2 Rekacewicz, P., and Digout, D. (2005). UNEP/GRID-Arendal. UN Environment Program (accessed May 2021).  https://www.grida.no/resources/5794 Public domain,
• Figure 1.2.3 Steven Earle, CC BY 4.0
• Figure 1.2.4 Steven Earle, CC BY 4.0
• Figure 1.2.5 Steven Earle, CC BY 4.0
• Figure 1.2.6 Steven Earle, CC BY 4.0
• Figure 1.2.7 Steven Earle, CC BY 4.0
• Figure 1.2.8 Steven Earle, CC BY 4.0

Earth System Science is an important part of Environmental Geology, and so interaction between systems is considered in almost every topic covered in this textbook. But Environmental Geology is about much more than Earth System interactions.

Environmental Geology is about the interface between geological processes and the environment, and we can choose to use a broad definition of “the environment”. In this textbook that covers all types of Earth Systems, including systems involving the biosphere (or ecosystems if you like); systems of the hydrosphere and atmosphere, such as the hydrological cycle and glaciation; and systems of the geosphere, such as volcanism, earthquakes and slope failure. But it also includes human activities that have implications for the geosphere (and hence the rest of the Earth System), such as mining, energy resource extraction, and waste disposal: into the ground, into water and into the air.
The major topics considered here (apart from a review of Physical Geology in Chapter 2) are as follows:

Chapter 3: Geological Control Over the Earth’s Climate. Climate change is the most significant issue of our time, and in order to understand how we are changing the climate now we need to know how it has changed naturally in the past. We will look back in time to consider the geological and other processes that have controlled our climate over the past 4.6 billion years. These include long-term changes in the sun, evolution of living organisms, continental positions, mountain building, changes to ocean currents, volcanic eruptions, variations in the Earth’s orbital shape and tilt and collisions with extra-terrestrial objects.

Chapter 4: Glaciation. Glaciation has significant implications for topography and surficial materials and the extent, motion and melting of glaciers are important aspects of the Earth System. In this chapter we will consider some of the Earth’s past glacial periods, the persistent cooling over the Cenozoic, and the cyclical controlling factors of the Quaternary glaciations. We will compare continental and alpine glaciation and will examine glacial erosion landforms and glacial deposits and their implications for other aspects of Environmental Geology.

Chapter 5: Slope Failure. Plate motions and volcanism create steep slopes and those slopes are subject to the pull of gravity. In looking at slope failure we will consider forces on slopes, the natural angle of repose of loose materials, the importance of water, types of failure motion, classification of slope failure, the implications of glaciation for slope failure, triggers for slope failure, and implications of climate change for slope failure.

Chapter 6: Earthquakes. Earthquakes cause massive destruction and death around the world and it is important to understand them so that we can all take steps to reduce our vulnerability. We will consider plate boundary processes, rock strength, elastic deformation, sticking and slipping, rupture surfaces, seismic waves, amplification, liquefaction, earthquake predictions and warnings, and public and personal earthquake preparation.

Chapter 7: Volcanoes. As noted above, volcanic eruptions are an important component of the Earth System, but they also represent significant geological hazards. We will look at volcanism in the context of plate boundary processes, the importance of magma characteristics, volcanic hazards (lava flows, pyroclastic flows, lahars, ash fall), benefits from volcanism, predicting eruptions, and preparing for potential eruptions.

Chapter 8: Resources. Our civilization is built around a supply of metals, so it is important to understand where they come from and the implications of their extraction and use. We will consider background metal contents in rocks and metal enrichment processes, alteration of surrounding rock, mining methods, mine wastes, ore processing wastes, acid rock drainage and metal contamination, mine-waste accidents, and the effects of use of metals on climate change. We will also take a look the sources and environmental issues related to some of the metals important to modern technology, including lithium for batteries.

Chapter 9: Energy Resources. Our current way of life is tied to an abundant supply of cheap energy, and for the past 200 years that has been provided mostly by fossil fuels. As we know, that cannot continue, so we must shift our focus to sustainable energy sources. In this chapter we will look at the formation, extraction, use and emissions of coal, oil and gas. We will also consider other energy sources such as: uranium, hydro, wind, solar, geothermal, and wave energy.

Chapter 10: Soils and Clay Minerals. Eight billion people cannot live on this planet unless we grow a lot of food, so an understanding of soil is critically important. We will discuss some of the variables in soil formation, such as climate, parent material, slope and time, and also the importance of soil conservation. Clay minerals are important components of soils, but they are also significant to many other geological processes. We will consider the origins, mineralogy, properties, importance of clay minerals in the context of agriculture, climate change, earthquakes, mineral exploration, groundwater, slope failure, waste disposal, and environmental geochemistry.

Chapter 11: Water. A supply of clean water is essential to our lives, and in many areas that comes from surface sources like rivers and lakes. We will discuss the basics of hydrology, hydrographs, flood recurrence intervals, dyking, dams, and flooding, along with natural and anthropogenic contamination, and implications of climate change for surface resources.
The other major water source is groundwater, and there are many ways in which surface water and groundwater are connected. In order to understand groundwater resources, we need to examine porosity and permeability, aquifers (unconfined and confined), the water table and the potentiometric surface, and hydraulic gradient. Some other issues of importance are wells and pumping, groundwater chemistry, contamination, and the implications of climate change.

Chapter 12: Karst and Caves. Caves typically develop in areas with soluble bedrock, such as limestone. They can have significant environmental implications. We’ll be looking at the surface and underground features of cave landscape, how water flows through caves, how caves form, some of the contents of cave systems, and how humans interact with caves and karst.

Chapter 13: Flooding. In terms of human and economic cost, flooding is the serious type of natural disaster, and as we have seen in recent years, it is only going to get more frequent and more serious with climate change.  In this chapter we will discuss the causes and consequences of flooding, and some of the steps that can be taken to reduce the impacts of flooding.

Chapter 14: Waste Disposal. Solid waste disposal is a geological problem because most of our waste is still placed in holes in the ground. We will discuss the sources and composition of waste, waste diversion, the components of a landfill, the generation and composition of leachate solutions, and landfill gases and their contribution to climate change.

Chapter 15: Consequences of Climate Change. Chapter 3 is about some of the natural processes of climate change that have taken place over Earth’s history. In Chapter 15 we’ll focus on anthropogenic climate change and some of the serious consequences we are currently seeing, and can expect to see more of in the future, including glacial ice melt, sea-ice melt, destabilization of permafrost, extreme drought and rainfall, wildfire, sea-level rise, tropical storms, changes to ocean currents, ocean acidification, and how those changes can affect geological processes.

It isn’t difficult to see that Environmental Geology is more important now than it has ever been. This is partly because environmental issues in general are more important than ever due to the crisis of climate change and the rapidly expanding human population, but also because climate change is, to a large degree, a geological problem. We can understand its past by studying the geological record going back thousands, millions and billions of years and we can understand its present by applying geological methods to data collection and analysis. Moreover, climate change is affecting many processes that fall into the realm of Environmental Geology, such as water supply, flooding, erosion, deglaciation, and slope failure.
Most important of all, we can affect the future of climate change by changing the way we do things, and every one of us has a role in making those changes.

The topics covered in this chapter can be summarized as follows:

Electrons, Protons, Neutrons and Atoms

Figure 2.1.1 A Part of the Periodic Table Showing Element Symbols (e.g., Cr), Names (e.g., chromium), Atomic Numbers and Atomic Weights

Bonding and Lattices

Sodium has 11 electrons, 2 in the first shell, 8 in the second, and 1 in the third (Figure 2.1.2). Sodium readily gives up this third shell electron, and when it does it loses one negative charge and becomes positively charged. Chlorine, on the other hand, has 17 electrons, 2 in the first shell, 8 in the second, and 7 in the third. Chlorine readily accepts an eighth electron for its third shell, and therefore becomes negatively charged. In changing their number of electrons these atoms become ions—the sodium a positive ion or cation, the chlorine a negative ion or anion. The resulting electronic attraction between these ions is known as an ionic bond. Electrons can be thought of as being transferred from one atom to another in an ionic bond.

Figure 2.1.2 The Electron Configuration of Sodium and Chlorine Atoms (top). Sodium gives up an electron to become a cation (bottom left) and chlorine accepts an electron to become an anion (bottom right).

Figure 2.1.3 The Alternating Arrangement of Sodium and Chlorine Atoms in the Mineral Halite (left); and An Example of a Cubic Crystal of Halite (right).

Exercise 2.1 Cations, Anions and Ionic Bonding

Several elements are listed below along with their atomic numbers.  Assuming that the first electron shell can hold 2 electrons, and that subsequent electron shells can hold 8 electrons, sketch in the electron configurations for these elements.  Predict whether the element is likely to form a cation (+) or an anion (-), and what valency it would have (e.g., +1, +2, -1 etc.)  The first one is done for you.

Figure 2.1.4 Predicting Cations, Anions and Ionic Bonding

Exercise answers are provided Appendix 2.

 

Figure 2.1.5 Electron Sharing in Carbon-Carbon Co-valent Bonds. The light blue circles represent shared electrons.

The elements silicon and oxygen are both abundant in the Earth’s crust and mantle and are present in the minerals that make up most rocks.  Silicon and oxygen bond together to create a four-sided pyramid shape with an oxygen at each corner and a silicon in the middle, known as a silica tetrahedron (Figure 2.1.5).  This is the building block of the many important silicate minerals. The bonds in a silica tetrahedron have some of the properties of covalent bonds and some of the properties of ionic bonds. As a result of the ionic character, silicon becomes a cation (with a charge of +4) and oxygen becomes an anion (with a charge of -2).  The net charge of a silica tetrahedron (Si0₄) is -4.  As we will see later, silica tetrahedra are linked together in a variety of ways to form most of the common minerals of the crust.

Figure 2.1.6 The Silica Tetrahedron – Building Block of All Silicate Minerals. Because the silicon has a charge of +4 and the four oxygens each have a charge of -2, the silica tetrahedron has a net charge of -4.

Non-Silicate Minerals and Mineral Groups

Oxide minerals have oxygen as their anion, but they exclude those with oxygen complexes such as carbonate (CO₃), sulphate (SO₄), or silicate (SiO₂). If the oxygen is also combined with hydrogen to form the hydroxyl anion (OH^–) the minerals is known as a hydroxide. Some important hydroxides are limonite (FeO(OH)) and gibbsite (Al(OH)₃), which are common within ores of iron and aluminum.

Sulphides are minerals with the S^-2 anion, and they include galena (PbS), sphalerite (ZnS), chalcopyrite (CuFeS₂) and molybdenite (MoS₂), which are the most important ores of lead, zinc, copper and molybdenum respectively. Another important sulphide minerals is pyrite (FeS₂).

Sulphates are minerals with the SO₄^-2 anion, and these include anhydrite (CaSO₄ and its cousin gypsum CaSO₄.2H₂0) and the sulphates of barium and strontium: barite (BaSO₄) and celestite (SrSO₄).  In all of these cases the cation has a +2 charge which balances the -2 charge on the sulphate ion.

The halides are so named because the anions include the halogen elements chlorine, fluorine, bromine etc.  Examples are halite (NaCl) and fluorite (CaF₂).

The carbonates include minerals in which the anion is the CO₃^-2 complex. The carbonate combines with +2 cations to form minerals such as calcite (CaCO₃), magnesite (MgCO₃), dolomite ((Ca,Mg)CO₃) and siderite (FeCO₃). The copper minerals malachite and azurite are also carbonates.

In phosphate minerals the anion is PO₄^-4. An important phosphate mineral is apatite (Ca₅(PO₄)₃(OH)), which is what your teeth are made of.

Native minerals include only one element (bonded to itself), such as gold, copper, sulphur or carbon (which could be graphite or diamond).

The silicate minerals include the elements silicon and oxygen in varying proportions ranging from SiO₂ to SiO₄. These are discussed at length below.

Silicate Minerals

Figure 2.1.7 Silicate Mineral Configurations. The triangles represent silica tetrahedra.

In pyroxene silica tetrahedra are linked together in a single chain, where one oxygen ion from each tetrahedron is shared with the adjacent tetrahedron, and hence there are fewer oxygens in the structure. The result is that the oxygen to silicon ratio is lower than in olivine (3:1 instead of 4:1), the net charge per silicon atom is atom is less (-2 instead of -4), and fewer cations are necessary to balance that charge.  Pyroxene compositions are of the type MgSiO₃, FeSiO₃ and CaSiO₃, or some combination of these.  In other words, pyroxene has one cation for each silica tetrahedron (e.g., MgSiO₃) while olivine has two (e.g., Mg₂SiO₄).

In amphibole structures the silica tetrahedra are linked in a double chain that has an oxygen to silicon ratio lower than that of pyroxene, and hence still fewer cations are necessary to balance the charge. Amphibole is even more permissive than pyroxene and its compositions can be very complex. Hornblende, for example, can include sodium, potassium, calcium, magnesium, iron, aluminum, silicon, oxygen, fluorine and the hydroxyl ion (OH^-1).

In mica structures the silica tetrahedra are arranged in continuous sheets.  There is even more sharing of oxygens between adjacent tetrahedra and hence fewer charge-balancing cations are needed.  Bonding between sheets is relatively weak, and this accounts for the well-developed one-directional cleavage (Figure 2.1.7). Biotite mica can have iron and/or magnesium in it and that makes it a ferromagnesian silicate mineral (like olivine, pyroxene and amphibole).  Chlorite is another similar mineral which commonly includes magnesium. In muscovite mica the only cations present are aluminum and potassium, and hence it is not a ferromagnesian silicate mineral.

Figure 2.1.8 Biotite Mica (left) and Muscovite Mica (right). Both are sheet silicates and split easily into thin layers along planes parallel to the sheets. Biotite is dark like the other iron- and/or magnesium-bearing silicates (e.g., olivine, pyroxene and amphibole), while muscovite is light coloured. Each sample is about 3 cm across.

Apart from muscovite, biotite and chlorite, there are many other sheet silicates (or phyllosilicates) which usually exist as clay-sized fragments (i.e., less than 0.004 mm). These include the clay minerals kaolinite, illite and smectite, and we will be looking at them more closely in Chapter 10.

Silica tetrahedra are bonded in three-dimensional frameworks in both the feldspars and in quartz. These minerals are non-ferromagnesian—they don’t contain any iron or magnesium. In addition to silica tetrahedra, the feldspars include aluminum, potassium, sodium and calcium in various combinations.  The three main feldspar minerals are orthoclase (a.k.a., potassium feldspar), and two types of plagioclase: albite (sodium only) and anorthite (calcium only).  Quartz contains only silica tetrahedra, so only has silicon and oxygen.

In quartz the silica tetrahedra are bonded in a “perfect” three-dimensional framework.  Each tetrahedron is bonded to four other tetrahedra (with oxygen shared at each corner of each tetrahedron), and as a result, the ratio of silicon to oxygen is 1:2 (quartz is SiO₂). Since the one silicon cation has a +4 charge and the two oxygen anions each have a -2 charge, the charge is balanced.  There is no need for aluminum or any of the other cations such as sodium or potassium.  The hardness, and lack of cleavage in quartz, are related to the fact that all of the bonds are the strong covalent/ionic bonds characteristic of the silica tetrahedron.

Rocks are aggregates of the crystals of one or more minerals, and there are three main types: igneous, sedimentary and metamorphic. They form in different ways and they have specific properties that help us to distinguish them. Through geological processes the rocks on Earth can be transformed from one type to another, and this concept is illustrated using the rock cycle (Figure 2.2.1). The three types of rock are depicted within rectangles, some of the intermediate forms of rocky materials (magma, outcrop and loose sediments) are shown as ellipses, and the processes that are important in the transformations are shown as arrows. Those processes (solid black arrows) include:

• Uplift (such as mountain formation) that results in rock that was present at depth in the crust being brought to surface and exposed to weathering,
• Erosion and transportation of weathered products (e.g., sand, clay, and ions in solution) and then deposition as sediments (e.g., within rivers or the ocean),
• Burial of those sediments beneath other sediments, followed by compaction (squeezing) and cementation to make sedimentary rock,
• Further burial which results in heating and more squeezing and then mineral transformations to form metamorphic rock,
• Further heating or other changes in conditions that result in melting to form magma, and
• Movement of magma towards surface where it can cool slowly to make intrusive igneous rock, or to surface where it can cool quickly to make extrusive igneous (volcanic) rock.

The rock cycle can be disrupted at any point by a change in conditions brought about by uplift, or burial (dashed blue and red arrows), or in rare cases by the nearby presence of magma.

Figure 2.2.1 The Rock Cycle

Igneous Rocks

Igneous rocks form from the cooling of magma (molten rock), either slowly at depth in the crust, or quickly at surface. In general, the longer that cooling process takes (up to millions of years), the larger the crystals will be. Volcanic rocks typically have mineral crystals that are less than 0.1 mm across (because they can cool in seconds or minutes), while intrusive igneous rocks have crystals that are typically larger than 1 mm across. Classification of igneous rocks is based mainly on the rock composition and also on the texture, and this is illustrated in Figure 2.2.2.

Figure 2.2.2 General Classification of Igneous Rocks

Three broad compositional classes of igneous rocks are shown, namely felsic, intermediate and mafic, and these are determined by the proportions of the dark silicate minerals (biotite, amphibole, pyroxene and olivine). Felsic rocks are light coloured, often close to white, with less than 20% dark minerals. Intermediate rocks are medium-dark (typically grey) with 20 to 50% dark minerals, and mafic rocks are close to black (with more than 50% dark minerals). The three main types of intrusive igneous rocks are granite, diorite and gabbro. The equivalent volcanic rocks—which are fine grained—are rhyolite, andesite and basalt. Igneous rocks with close to 100% dark minerals (not shown on Figure 2.2.2) are known as ultramafic; these are rare on the Earth’s surface, but common in the mantle.

Mafic igneous rocks (e.g., basalt or gabbro) are denser than felsic igneous rocks. Basalt has a specific gravity of about 3 g/cm³, while for granite the value is about 2.6 g/cm³. This small difference becomes very important in the context of plate tectonics. In comparison, the ultramafic rock of the mantle has a density of about 3.3 g/cm³.

Sedimentary Rocks

Sedimentary rocks form near to the Earth’s surface following the accumulation of fragments of rocks and minerals that have been weathered and eroded from outcrops, transported by gravity, rivers, waves, wind, or glacial ice, and then deposited as sediments (Figure 2.2.3). The sedimentary grains (e.g., grains of sand) are known as clasts, and the resulting sedimentary rocks are called clastic if they are composed mostly of such grains.

Figure 2.2.3 Cross-Bedded River-Deposited Sand on Vancouver Island. This is not sandstone because the deposit is soft and loose, and can be scraped away with fingers.

Clasts can range in size from tiny (invisible) clay fragments to boulders the size of buildings and we classify clastic sedimentary rocks on the basis of the clast sizes. The key piece of information to remember is that a grain of sand ranges in size from 1/16^th mm to 2 mm, and that means from about 1/4 the size of the period at the end of this sentence, to about the size of this capital O. (That depends, of course, on the type of device that you are reading this on.) Very fine sand will feel gritty (not slippery) between your fingertips. Clasts smaller than 1/16^th mm are classed as silt and clay, and those larger than 2 mm are granules, pebbles, cobbles and boulders (in order of increasing size). Sand grains can be transported by rivers with medium flow, by strong winds, and by waves, and so sand deposits tend to accumulate in rivers and deserts and on beaches. Silt and clay can be transported in similar environments, but they tend not to be deposited unless the medium slows, so they will be deposited in lakes, and the ocean. Granules and larger fragments can typically only be transported and deposited by fast-flowing water, and so they are commonly deposited in high-energy parts of streams. These sediments must then become buried beneath other layers of sediments and compressed and cemented before they can become sedimentary, as illustrated on Figure 2.2.1.

The three main types of clastic sedimentary rocks are illustrated on Figure 2.2.4

Figure 2.2.4 Examples of the Elastic Sedimentary Rocks Conglomerate, Sandstone and Mudstone. The pen is 15 cm long, and the coin in the mudstone photo is 2 cm across.

Sedimentary rocks can also form from the crystallization of ions that were transported in water as dissolved ions. These are known as chemical sedimentary rocks. For example, marine organisms extract bicarbonate and calcium ions (HCO₃^– and Ca²⁺) from ocean-water to make calcite shells (CaCO₃) which then accumulate on the sea floor (typically in tropical areas around reefs) to form calcite mud and sand that later gets buried and becomes limestone. Some organisms make their shells out of silica, and those can accumulate on the sea floor to make the rock chert. In some cases, minerals form from the evaporation of water in an inland sea or lake. Examples are rock salt (halite) and gypsum (which is used to make plaster board).

Metamorphic Rocks

Metamorphic rocks form when pre-existing sedimentary or igneous rocks are heated and squeezed in such a way that one or more of the minerals present becomes unstable. The result might be that crystals of those minerals are converted into different minerals or into larger crystals of the same type.
One way to understand this process is to consider the sedimentary rock mudstone, which is mostly made up to clay minerals. Clays are low-temperature minerals; they are not stable at temperatures higher than about 150⁰ C. As a clay-rich rock is heated the clay minerals tend to break down and are converted into micas. At even higher temperatures those might be converted to minerals such as quartz, feldspar and amphibole.

Most metamorphism takes place at depth in the crust in areas that have experienced mountain-building (and so crustal thickening) and where there is compression due to plate convergence. That means that the rocks get heated (because of burial) and squeezed (because of the converging plates) at the same time. New minerals that form under these conditions are typically forced to grow perpendicular to the direction of that pressure, and so the metamorphic rock becomes foliated–meaning that it takes on a fabric of aligned minerals or aligned bands of minerals.

Examples of foliated metamorphic rocks are shown on Figure 2.2.5. Slate forms from mudstone at relatively low metamorphic grade, as clay minerals are turned into tiny (invisible) mica crystals. The alignment of these gives slate a layered look and the tendency to split into sheets. Schist forms at higher temperatures that allow the micas to become large enough to see. The mica crystals are generally parallel to each other, but the rock doesn’t split into sheets so easily. Gneiss forms at temperatures thar are typically beyond the stability of micas, and so is characterized by minerals like quartz, amphibole and feldspar that have segregated into dark and light bands.

Figure 2.2.5 Examples of Foliated Metamorphic Rocks. (From left to right: slate, schist and gneiss.)

It is very important not to confuse foliation with bedding. The slate of Figure 2.2.5 is not splitting along pre-existing bedding planes in the parent mudstone, and the layers in the gneiss have no relationship to what bedding (if any) might have existed in the parent rock.

Some other metamorphic rocks include quartzite and marble. These form from the metamorphosis of sandstone and limestone respectively, and they tend not to be foliated even if they did form under directional squeezing. Quartzite, for example, is mostly made up of quartz crystals, and they tend not to take on a directional fabric.

Exercise 2.3 Rock Groups

For each of the rocks illustrated in Figure 2.2.6 indicate (a) the major rock type (e.g., igneous, sedimentary or metamorphic), (b) a sub-type (e.g., intrusive igneous, clastic sedimentary), (c) a rock name (e.g., sandstone, andesite), and (d) the likely geological environment and processes that led to the formation of this rock.

Figure 2.2.6 What Kinds of Rocks are These?

Exercise answers are provided Appendix 2.

Media Attributions

• Figure 2.2.1 Steven Earle, CC BY 4.0
• Figure 2.2.2 Steven Earle, CC BY 4.0
• Figure 2.2.3 Steven Earle, CC BY 4.0
• Figure 2.2.4 (all photos) Steven Earle, CC BY 4.0
• Figure 2.2.5 (all photos) Steven Earle, CC BY 4.0
• Figure 2.2.6 (all photos) Steven Earle, CC BY 4.0

As shown on Figure 2.3.1 the interior of the Earth is comprised of a number of concentric layers, starting with the crust at surface, down through the different parts of the mantle, and then into the outer and inner core.

Figure 2.3.1 The Parts of the Earth’s Interior

The crust is the Earth’s thin outer layer, and is shown in more detail in the inset of Figure 2.3.1. It makes up approximately 0.2% of the Earth’s radius, and therefore is proportionally no thicker than the skin on an apple. The continental crust is dominated by intrusive igneous rocks, with lesser amounts of various types of volcanic sedimentary, and metamorphic rocks, and with an overall composition similar to that of granite. It is 30 to 40 km thick in most areas (thicker in mountainous areas, and thinner near to the edges of the continents). The oceanic crust is made up mostly of mafic igneous rocks, including basalt in the upper several hundred metres, and gabbro at depth. The oceanic crust is 5 to 6 km thick in most areas. The mafic rocks of the oceanic crust are denser than the granitic rocks of the continental crust.

The crust rests on the mantle, which is made up of rocks with an ultramafic composition (dominated by ferromagnesian silicate minerals). The upper part of the mantle is rigid or brittle, just like the crust, and this is known as the lithospheric mantle. That rigid part of the mantle plus the crust is called the lithosphere. The temperature within the Earth increases with depth, from close to 0⁰ C at surface to almost 5000⁰ C in the centre of the core, but that rate of increase isn’t linear (Figure 2.3.2). The temperature rises quite rapidly within the lithosphere but the rate slows within the lower mantle. At between about 100 and 250 km depth the temperature is very close to, or even slightly above, the melting temperature for mantle rock and so the mantle rock within that depth range is partly molten. This layer is known as the asthenosphere because the rock there is weaker (softer, or more plastic-like) than in the rest of the mantle.

Figure 2.3.2 Temperature and Partial Melting in the Upper Part of the Mantle

Below the asthenosphere the mantle rock is solid, but it is not rigid, and in response to long-term stress created by heat transfer from the core, it is slowly convecting. As we’ll see in the next section, that convection is critical to the process of plate tectonics.

There is evidence that mantle plumes exist within the mantle, where a mantle plume is a column of hot mantle rock (not magma) that rises either from near to the core-mantle boundary, or from the mid-mantle region, and reaches the upper mantle just below the crust (Figure 2.3.3). Mantle plumes are thought to be responsible for chains of volcanoes like the Hawaiian Islands, and for other volcanic phenomena that will be discussed in Chapter 3 and Chapter 5.

Figure 2.3.3 Conceptual Models of Mantle Plumes

The Earth’s core is mostly made of iron, with up to 10% nickel and a small proportion of other elements, including silicon, oxygen and sulphur. The temperature at the core-mantle boundary is close to 4000⁰ C and that gradually increases to about 5000⁰ C at the Earth’s centre. In the outer part of the core that temperature is high enough for the iron-nickel alloy to be liquid, but the pressure is so extreme at greater depth that the inner core is solid. The transfer of heat from below leads to convection in the outer core, and, because the core is metallic, that convection generates the Earth’s magnetic field.

Media Attributions

• Figure 2.3.1 Steven Earle, CC BY 4.0
• Figure 2.3.2 Steven Earle, CC BY 4.0
• Figure 2.3.3 Steven Earle, CC BY 4.0

The lithosphere (the upper rigid mantle plus the crust) is divided into a number of pieces known as plates that can move around on the Earth’s surface, all in different directions and at different rates. The behaviour of these plates: their motions and their interactions with each other is known as plate tectonics, where the word “tectonics” refers to the deformation of rocks. The distributions of the plates are shown on Figure 2.4.1; the seven major ones are as follows:

• North America Plate: underlies most of North America, part of eastern Russia, part of the Arctic Ocean and the northwestern part of the Atlantic Ocean,
• Eurasia Plate: underlies Europe, much of Asia and the northeastern part of the Atlantic Ocean,
• Pacific Plate: underlies much of the Pacific Ocean,
• Africa Plate: underlies Africa, the southeastern Atlantic Ocean, and the western Indian Ocean,
• South America Plate: underlies South America and the southwestern Atlantic Ocean,
• Indo-Australian Plate (includes Australia and India Plates): underlies India, Australia, and parts of the surrounding Indian and Pacific Oceans and islands therein, and
• Antarctic Plate: underlies Antarctica and most of the Southern Ocean.

Figure 2.4.1 The Approximate Extents and Rates of Motion of the Earth’s Tectonic Plates

Some of the minor plates shown on Figure 2.4.1 include the Juan de Fuca, Cocos and Nazca Plates along the western edge of N. America, the Caribbean Plate, and the Scotia, Arabia and Filipino Plates. There are several other smaller plates not shown on this map.

The generalized plate motions are shown on Figure 2.4.1. Plate motions range from roughly 1 to 10 cm/y. As can be seen on the map, parts of individual plates are moving in different directions and at different rates. That’s not because the plates are squeezing and stretching (although that does happen near to plate boundaries) but because the plates are all moving in a rotational way, each around a different rotational axis. The North America Plate is moving counter-clockwise around an axis in the southern hemisphere, so the rate of motion is greater in the far north than in the south, and it changes from “towards the northwest” in the east to “towards the southeast” in the west.

At plate boundaries the interaction between plates can be: convergent (moving towards each other), divergent (moving away from each other), or transform (moving side by side). A convergent boundary is illustrated on Figure 2.4.2. In this case a plate comprised of oceanic crust is moving towards one comprised of continental crust. Because oceanic crustal is denser than continental crust, the oceanic plate gets pushed down—or subducted—beneath the continental plate. That has some important implications, as shown. First, is that there is friction between the two plates, which results in periodic earthquakes, some of which can be very large. Second, is that the subducted oceanic crust gets heated and some of the free water and water held in minerals is released and starts to rise towards surface (more on this below). This water mixes with the hot rocks of the asthenosphere and that reduces their melting temperature (so more melting takes place) creating magma that moves towards surface.

Figure 2.4.2 Depiction of Some of the Processes Taking Place at an Ocean-Continent Subduction Zone. The red stars represent potential earthquakes.

As an oceanic plate converges with a continental plate (in the manner shown on Figure 2.4.2) it is possible that a continent or island will be moving along with that oceanic lithosphere and that the two areas of continental lithosphere will eventually collide. This scenario is illustrated on Figure 2.4.3. In this situation the continental lithosphere cannot be subducted because it isn’t sufficiently dense to be pushed down into the mantle. As the continents collide the sediments that had accumulated along their edges get pushed up to form fold-belt mountains and the older crustal rocks also get deformed and pushed up. The leading edge of the subducted oceanic lithosphere eventually breaks off and descends into the mantle.

Figure 2.4.3 Depiction of Processes Taking Place at a Continent-Continent Convergence Zone

A divergent boundary exists where two plates are moving apart from each other, likely in response to convection in the mantle. This means that there is slow upward movement of mantle rock along the ridge axis. As the hot mantle rock moves upward it experiences reduced pressure which brings it closer to its melting point (decompression melting). Some of it (approximately 10%) melts, producing mafic magma that erupts at surface (in this case on the sea floor) to form basalt, and also cools beneath surface to form gabbro. New oceanic crust is made in this way.

Figure 2.4.4 Depiction of Processes Taking Place at a Divergent Boundary Between Two Oceanic Plates

Figure 2.4.5 Three Different Types of Plate Boundaries in the Western US and Canada

Some plate boundaries are shown in plan view on Figure 2.4.5. The Juan de Fuca Ridge, where the Juan de Fuca and Pacific Plates are moving away from each other and new oceanic crust is being made, is an example of a divergent boundary. The Cascadia Subduction Zone is an example of a convergent boundary. Here the Juan de Fuca Plate is pushing down underneath the North America Plate, resulting in earthquakes and volcanoes.

The third type of plate boundary—a transform boundary—where two plates are moving side-by-side relative to each other, also exists in this region. An example is the San Andreas Fault, which forms the boundary between the North America and Pacific Plates through California. The relative motion of these two plates is shown with small white arrows. There have been large earthquakes along this boundary, and there are frequent smaller ones.

Two smaller transform boundary segments are shown (as red lines) on the map, between the segments of the Juan de Fuca Ridge. In both cases the Juan de Fuca Plate is moving relative to the Pacific Plate, and there are frequent small earthquakes along these boundaries.

Exercise 2.4 Plate Boundary Processes

Using Figures 2.4.2 and 2.4.4 as examples, draw a cross section through the crust and upper mantle along the white dashed line labelled a-b on Figure 2.4.6 below. Label the plates and show their motion directions with arrows. Indicate where there might be earthquakes and volcanic activity.

Figure 2.4.6

Exercise answers are provided Appendix 2.

Media Attributions

• Figure 2.4.1 Steven Earle, CC BY 4.0, from a public domain base map, “Slabs,” created by US Geological Survey
• Figure 2.4.2 Steven Earle, CC BY 4.0
• Figure 2.4.3 Steven Earle, CC BY 4.0
• Figure 2.4.4 Steven Earle, CC BY 4.0
• Figure 2.4.5 Steven Earle, CC BY 4.0
• Figure 2.4.6 Steven Earle, CC BY 4.0

Most Earth system processes involve components of the geosphere and some critically important ones take place entirely within the geosphere. Most of these are not visible to us, and we only know about them because we can observe their effects when rocks and magma get brought to surface. We can then use chemistry and physics to understand how they happened.

One such process takes place within the crust on either side of a divergent boundary. Due to the heat of volcanism close to the boundary itself, water that is present in the cracks and pores of the oceanic crust gets heated and rises towards surface. This draws more ocean water into the crust from adjacent areas and a strong convection system develops that can continue for millions of years (Figure 2.5.1). One result is that the oceanic crust becomes metamorphosed through a process of hydrothermal alteration. There are several different metamorphic reactions that take place there. For example, at around 400⁰ C olivine (a magnesium or iron silicate) in the basalt reacts with water to create serpentine (a hydrated magnesium sheet-silicate mineral) plus brucite (a magnesium hydroxide mineral), as shown here:

olivine + water serpentine + brucite

As a result of reactions like this one, much of the ocean crust—everywhere—is rich in minerals like serpentine and brucite that include the water in their structures.

Figure 2.5.1 Hydrothermal Processes at a Divergent Boundary

Wherever that hot groundwater flows out onto the seafloor it creates some interesting conditions. The groundwater in this setting tends to be quite rich in sulphur, along with various metals and when the hot groundwater meets the cold ocean the sulphur (as hydrogen sulphide) quickly reacts with metals to form sulphide minerals, such as pyrite:

The plume of tiny crystals of pyrite appear black and the result is known as a black smoker (Figure 2.5.2). In some cases, the sulphur is present in an oxidized form (as sulphate ions) and that combines with calcium ions to create the calcium sulphate mineral anhydrite.

In this situation the plume is likely to be white, so these are called white smokers.

Black or white, sea-floor hot spring vents are attractive places for sea life. Various microorganisms get energy from the unique chemical conditions, forming microbial mats around the vents. These are eaten by shelled organisms and crustaceans, which in turn provide food for larger organisms. All of this happens in complete darkness, making this one of the few ecosystems that doesn’t depend on sunlight as an energy source. Some have speculated that sea-floor vents are possible candidates for the origin of life on Earth.

Figure 2.5.2 A black smoker on the East Pacific Rise off the coast of Mexico

Another important geosphere Earth-system process takes place following subduction of oceanic lithosphere. This results in rock, sediment, water and biological matter that was present on the sea floor and part of the oceanic lithosphere being forced down into the mantle. When this material gets heated some of the water—including that within minerals like serpentine and brucite—gets released and comes back towards the surface above the subduction zone, but most of it, and everything else, continues down into the middle and lower mantle, where it eventually mixes with the rest of the mantle rock. Some of that material eventually comes back to surface via convection and magmatism, and if we look carefully at the chemistry of rocks produced from that magma it can be possible to see evidence of their long past history on the Earth’s surface.

Media Attributions

• Figure 2.5.1 Steven Earle, CC BY 4.0
• Figure 2.5.2 Public Domain image by Normark, W. & Foster, D., East Pacific Rise, 21 degrees north. Base of “black smoker” chimney, Pacific Ocean. via Wikipedia, https://en.wikipedia.org/wiki/Hydrothermal_vent#/media/File:BlackSmoker.jpg

The topics covered in this chapter can be summarized as follows:

The solar system, which has been around just a little longer than the Earth (4.57 billion years), formed out of the gaseous and particulate remnants of one (or more) stars that had existed in this part of the galaxy and then exploded. The sun has changed significantly over that time and will continue to change in the future. The evolution of a typical star like ours is depicted on Figure 3.1.1.

Figure 3.1.1 The Life Cycle of a Main Sequence Star Like Our Sun. Note that the depiction of the diameter of the Sun is not to scale. During the Red Giant phase the Sun is expected to be about 200 times larger than it is now, almost big enough to engulf the Earth, and certainly too close for anything to live here.

Although it’s not obvious from Figure 3.1.1, the sun’s luminosity has increased substantially over 4.57 billion years, and it is now about one-third brighter that it was originally (Figure 3.1.2). It is now producing about 33% more heat than it did at the start. This increase in intensity is a result of the ongoing conversion (by nuclear fusion) of hydrogen to helium in the sun’s core. The growing proportion of helium results in an increase in the density of the solar core region, which causes the core to contract. The increased gravitational pressure forces the hydrogen atoms closer together, and that accelerates the rate of fusion and makes the sun hotter and brighter.

The sun will continue to get hotter in this way for another four billion years or so, until its core is made up entirely of helium, at which point it will evolve into a Red Giant (Figure 3.1.1) and will start to expand, first consuming Mercury, and then Venus and quite likely even the Earth. At that time the helium will start to fuse into carbon, almost half of the sun’s mass will be lost into space via massive explosions, and what remains will eventually collapse into a White Dwarf.

As noted, the solar luminosity has increased by about 33% over its entire 4.6 billion year history—so far (Figure 3.1.2). That’s a huge amount from the perspective of Earth’s climate, but it’s over a very long time. The sun is going to get hotter still. Within another 4.4 billion years it will be roughly twice as hot as it was originally.

Figure 3.1.2 The Change in Luminosity of Our Sun, and therefore, in the Amount of Solar Energy Received on Earth

Please don’t be fooled into thinking that a warming sun is behind the climate change we see happening right now, or a reason for us to be concerned for our distant future. The present rate of solar warming is about 8% every billion years. That’s 0.008% every million years or 0.0000008% every century. That is a very, very small amount of warming on a human time scale. Over the past century it has not been enough to noticeably warm the climate—not even close. For example, the increase in luminosity due to long-term solar evolution from 1920 to 2020 was only enough to increase the Earth’s surface temperature by about 0.0000016⁰ C. During that time the surface temperature has actually increased by about 1⁰ C, which is about 625,000 faster than the temperature increase that could be attributed to the change in solar luminosity over the same time period, so it’s clear that solar evolution is not the cause.

The earliest evidence of life on Earth is found in rocks about four billion years old. At that time the sun was about 80% as bright as it is today. An Earth with today’s atmosphere, and with an 80% sun, would have been completely frozen (no liquid water anywhere on the surface!). That might lead us to wonder how it could have been possible for any type of life to evolve, and this is known as ‘the faint young sun paradox’. The Earth wasn’t encased in ice four billion years ago because the Archean atmosphere was very different from today’s atmosphere. It was considerably thicker—approaching the density of the atmosphere of Venus—and, as shown on Figure 3.1.3, it was rich in carbon dioxide. The carbon dioxide proportion might have been in the order of 10%, as compared with today’s level of 0.04% (415 parts per million), and that provided a sufficiently strong greenhouse effect to keep the Earth’s surface temperature warm even with a relatively faint sun.

Figure 3.1.3 Evolution of the Earth’s Atmosphere Over Geological Time. The pie charts at the bottom show the relative proportions of CO₂, CH₄ and O₂ in our atmosphere at different times in the past (represented by the dashed lines) but do not include the gas N₂.

Life on Earth may have evolved in a sea-floor volcanic region close to scalding submarine hot springs (see Chapter 2), or in a shallow pool that was subject to repeated wetting and drying, or in some other environment with a source of heat and some useful molecules. In any case, the Earth’s atmosphere at the time was rich in carbon dioxide and had no free oxygen at all. Free oxygen (such as the O₂ in our present atmosphere, and unlike the oxygen in CO₂ or H₂O) would have been a deadly poison to the earliest microorganisms, just as it is today to some of the things that thrive in the dark and damp places in our bodies and our homes, and in boggy environments around us. Many early life forms were methanogens—meaning that they produced methane—and so methane levels in the atmosphere increased as life started to thrive. This contributed to the warming because methane is a potent greenhouse gas. Around 3.5 Ga (see Box 3.1) microorganisms developed the ability to use the sun as an energy source. The first of these were bacteria (possibly cyanobacteria, a.k.a. blue-green algae – see Figure 3.1.4) and they had a distinct advantage over other life forms of the day in that they didn’t have to rely on energy from volcanic hot springs or from chemical sources: they could live virtually anywhere that the sun’s light could reach. The photosynthetic process involves consumption of carbon dioxide and release of oxygen, but the proportion of free oxygen in the atmosphere remained very low for almost another billion years. That’s because any oxygen produced by photosynthetic organisms was first used up in chemical reactions with abundant elements like iron, or gases like methane, and was also consumed through the decay of dead organic matter.

Figure 3.1.4 Unicellular Cyanobacteria from a Microbial Mat, Found Near to Guerrero Negro, Mexico

Free oxygen first started to accumulate in our atmosphere around 2.4 Ga, initially in very small amounts. We call this transition the “oxygen crisis” because it led to extinction for many organisms. On the other hand, it really was the beginning of the beginning for organisms like us because the oxygen crisis appears to have pushed some existing organisms to evolve cells with a nucleus. Such organisms are eukaryotes, and they are our ancestors.

As more and more oxygen was produced it continued to react with methane (because methane and oxygen quickly react to form carbon dioxide and water) and the methane level dropped. Since methane is a much more powerful greenhouse gas than carbon dioxide, the climate cooled dramatically, triggering an extensive period of glaciation starting around 2.3 Ga—the Huronian glaciation. That was followed by a very warm period, and then the climate levelled out a bit, and there is no evidence of glaciation on Earth for another 1.6 billion years.

So, to summarize, it was relatively warm during most of the Earth’s first few billion years, despite a cooler sun. That’s because greenhouse gas levels were much higher than they are now. But there’s still a paradox here because life has evolved and flourished in liquid water (not frozen and not boiled off into space) for four billion years. How could the Earth have maintained a reasonable temperature throughout its history, a temperature comfortable enough (notwithstanding some ups and downs, such as the Huronian glaciation) for enough water to remain liquid? In other words, why has the Earth’s atmosphere changed as the sun has warmed, in such a way that a “Goldilocks” climate has existed here for four billion years?

Although we don’t fully understand why the Earth has been so habitable for so long, a key mechanism is the evolution of the atmosphere, and one of the drivers of that is photosynthesis. Various types of photosynthetic organisms—large and small, in the oceans and then later on land—have taken carbon dioxide from the atmosphere and released oxygen, converting the carbon to hydrocarbons and storing it in the rocks of the Earth’s crust (Figure 3.1.4). A parallel process, with a similar outcome, is the conversion of carbon from carbon dioxide into carbonate minerals. This is what most shelled organisms do to make their shells. Over time, as the sun has slowly warmed, life has used these two processes to reduce the greenhouse effect enough to keep the Earth from getting too hot.

The idea that life has controlled Earth’s climate to its own benefit forms the basis of the Gaia theory, first proposed by James Lovelock in 1972 and expanded upon by Lovelock and Lynn Margulis in 1974. According to Lovelock and Margulis, the Earth and the living organisms on it form a self-regulating system that ensures that conditions remain suitable for life to persist, and they have been able to do so even as the system’s main source of energy (the sun) has slowly changed in intensity. This self-regulation proceeds through various types of biological processes and climate feedbacks. (There’s more on climate feedbacks in Box 3.2, below.)

Don’t worry, this is not a sinister conspiracy on the part of the Earth’s organisms to control the climate, it is just a result of natural feedbacks that have led to conditions being reasonable for the life present at any one time, or for new types of organisms to evolve that can take advantage of changing conditions. The more important thing is that carbon has been stored in Earth’s crust as a way of getting it out of the atmosphere to keep the climate reasonable, and we—just another lifeform—are upsetting that balance by removing a lot of that carbon from the crust and using it to drive our cars.

Media Attributions

• Figure 3.1.1 Steven Earle, CC BY 4.0, after Solar Life Cycle image, public domain, by Oliver Beatson, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Solar_Life_Cycle.svg
• Figure 3.1.2 Steven Earle, CC BY 4.0; based on information in Ribas, I. (2009). The sun and stars as the primary energy input in planetary atmospheres. Proceedings of the International Astronomical Union, 5(S264), 3–18. https://doi.org/10.1017/S1743921309992298
• Figure 3.1.3 Steven Earle, CC BY 4.0, after Nisbet E., Fowler C., (2011). The evolution of the atmosphere in the Archaean and early Proterozoic. Chinese Science Bulletin, 56, 4−13. https://doi.org/10.1007/s11434-010-4199-8; and Large, R., Mukherjee, I., Gregory, D. et al. (2019). Atmosphere oxygen cycling through the Proterozoic and Phanerozoic. Mineral Deposita 54, 485–506. https://doi.org/10.1007/s00126-019-00873-9
• Figure 3.1.4 Cyanobacteria from Guerrero Negro, NASA public domain image, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Cyanobacteria_guerrero_negro.jpg

As described in Chapter 2, plate tectonics allows the plates of the Earth’s lithosphere to move around on the surface over time. This has resulted in many different configurations of continents over geological time, and also in plate-boundary interactions that have led to volcanism and the formation of mountain ranges.

Continental Positions

At present, the continents (which make up 29% of the Earth’s surface) are almost evenly distributed in a latitudinal sense, with approximately 33% in equatorial regions (between 30⁰ north and 30⁰ south), 38% in temperate regions (between 30⁰ and 60⁰ north and south), and 29% in polar regions (north of 60⁰ north and south of 60⁰ south).

This matters to climate because different types of Earth surfaces reflect different proportions of the energy that we get from the sun. That reflectivity is called albedo. Land surfaces are more reflective than open water, and some land surfaces are more reflective than others. Snow- and ice-covered surfaces have albedos in the range of 70 to 90%, unvegetated surfaces are generally in the range 15 to 40% (lower if wet), while vegetated surfaces are in the range 10 to 20%. Open water of oceans or lakes has an albedo of less than 10%. That means that 90% of the sun’s energy that shines on water is absorbed and converted into heat (in the form of warm water). Only about 10% of the sunlight that hits fresh snow is converted into heat; the rest is reflected back into space.

A critical factor in the context of albedo is latitude, because albedo makes a much bigger difference at low latitudes (equatorial regions), where solar intensity is high throughout the year, than it does at high latitudes, where solar intensity isn’t very high—even in the summer—because the sun never gets very far above the horizon.

At approximately 720 Ma the situation was much different than it is now, as shown on Figure 3.2.1. Most of the land area was part of the supercontinent Rodinia, with 50% of it in equatorial regions, 40% in temperate regions, and only 10% in polar regions. That much land in the sensitive equatorial zone had a cooling effect because of the higher albedo of the land (which was even higher than it is now since there was no land vegetation at that time ) compared to the ocean. Most of the sunlight striking that land reflected back into space, and was not converted into heat. The albedo implication of a supercontinent centred on the equator—more sunlight reflected back into space—is considered to be an important contributor to the first of the Cryogenian Period Snowball Earth glaciations.

Figure 3.2.1 The Approximate Location of the Supercontinent Rodinia in the Late Proterozoic

Starting at around 720 Ma the Earth entered its most intensive and the most extensive glacial period. The small temperature forcing caused by the albedo effect of an equatorial supercontinent (estimated at about 3⁰ C of cooling) led to snow and ice accumulation at high elevations and higher latitudes. Lower temperatures also favoured transfer of more of the atmosphere’s CO₂ into the oceans. Those feedbacks soon drove more intense cooling and before too long the land was mostly glaciated. This is known as the Sturtian Glaciation; it lasted about 60 million years. For much of that time the Earth’s mean annual temperature was about minus 40⁰ C, and the entire ocean—even at the equator—was covered in more than 200 meters of ice. Because there was little liquid water anywhere at surface, the hydrological cycle was essentially shut down.

Sixty million years is a very long winter but considering that the bright icy surface reflected most of the incoming solar energy, it might have been longer still (perhaps even until now). Our saving grace is that the Earth’s internal heat engine still motored on over that time, and volcanoes continued to erupt. Along with those eruptions came gases, including CO₂, and because there was no open ocean water almost all of that volcanic CO₂ stayed in the atmosphere, gradually building a greenhouse effect strong enough to start melting the ice. It is likely that the CO₂ level had to reach about 13% (about 325 times the current level of 0.04%) in order to overcome the cold. As some of the terrestrial glaciers receded, positive feedbacks started working to enhance the warming, including the decrease in albedo caused by melting ice, and the release of both carbon dioxide and methane from melting permafrost.

Eventually the sea ice started to melt, and it was probably mostly gone within several thousand years. This relatively rapid transformation from reflective ice and snow to dark open water for the entire ocean, under an atmosphere with at least several per cent CO₂, would have then contributed to an intense “hothouse” climate for at least another several thousands or tens of thousands of years.

Mountain Ranges

At about 100 Ma, the plate carrying the Indian continent started diverging from Antarctica and moved north towards Asia. While that was happening, the continents were eroding and sediments and sedimentary rocks were accumulating on the ocean floor adjacent to both continents. When India reached Asia, sometime between 55 and 45 Ma, the continental part of the Indian plate was unable to subduct (as illustrated on Figure 2.4.3). Instead, the rocks making up northern India and southern Asia, plus the sedimentary rocks in between, got crumpled, folded, faulted and uplifted to start construction of what is now—by a wide margin—the Earth’s highest and most extensive range of mountains (Figure 3.2.2). This uplift continued for tens of millions of years. In fact, the Indo-Australian Plate is still moving north, and still pushing the mountains up.

The Himalayans aren’t the only significant mountains to have been formed in relatively recent geological times. Others include the Zagros and adjacent mountain ranges of Iran, Iraq and Turkey, and the Alps of Europe, which were built mostly within the period of 65 to 40 Ma. All of these ranges can also be attributed to continental collisions.

Mountainous parts of the continents erode many times faster than plains (Figure 3.2.2). The Himalayan Range, which extends over 2400 km from Myanmar to Pakistan and well north into southern China, is eroding faster than any similar sized area on the planet and has been doing so for close to 50 million years. One of the processes associated with that erosion is chemical weathering of the rocks. This takes many forms, but the one of interest to us here is the hydrolysis (see Box 3.2) of silicate minerals, such as feldspar, to form clay minerals—and the resulting consumption of atmospheric carbon dioxide.

Figure 3.2.2 Rugged Terrain in the Himalayan Annapurna Region of Nepal. Rapid erosion is clearly evident from the thick accumulations of loose rocks on the lower slopes in the distance.

Hydrolysis is the process through which a molecule is split apart by water. In the context of mineral weathering, it can be represented like this:

water + carbon dioxide carbonic acid

This process can be written like this:

feldspar +  carbonic acid + oxygen kaolinite + calcium ions + carbonate ions

in which feldspar reacts with water, carbonic acid to form the clay mineral kaolinite along with calcium and carbonate ions in solution. The key thing happening here is that carbon dioxide is coming out of the atmosphere, first to form carbonic acid (H₂CO₃), and then reacting with feldspar to become carbonate ions that will eventually reach the ocean and get fixed into a mineral like calcite (CaCO₃), and will become part of a limestone deposit.

The effects of hydrolysis are illustrated on Figure 2.3.3, which shows unweathered (left) and weathered (right) parts of the same piece of granite. The surfaces of the feldspar crystals have been weathered to the clay mineral kaolinite.

Figure 3.2.3 A Fresh Surface of a Piece of Granite (left) and a Weathered Surface of the Same Piece (right).

The relationship between mountain building and global temperatures during the Cenozoic (since 66 Ma) is illustrated on Figure 3.2.4. Temperatures were consistently high through the Mesozoic (from 261 to 66 Ma) and that continued into the early part of the Cenozoic, but the climate started to cool around 50 Ma, and since then there has been a cumulative drop in global temperatures of about 14⁰ C. This long-term decline closely follows the atmospheric CO₂ curve, and most of that change can be attributed to the enhanced weathering associated with mountain ranges like the Himalayas, and therefore to plate tectonics. As described below, there are some other factors that contributed to climate change during the Cenozoic.

Figure 3.2.4 Temperature and CO₂ Change in the Context of Tectonic Events During the Cenozoic

Ocean Currents

Plate tectonic processes have also been responsible for climate change in other ways. For example, the movement of plates can change the characteristics of ocean basins, and that can change ocean currents, and therefore, the climate. Prior to about 40 Ma, the southern end of South America was either still connected to Antarctica, or at least the passage between them was too shallow to allow significant water flow. Sometime between 41 and 34 Ma that body of water—the Drake Passage—was widened and deepened by plate motion, and since then the strong Antarctic Circumpolar Current has flowed around the continent in a west to east direction (Figure 3.2.5).

Figure 3.2.5 The Approximate Path of the Antarctic Circumpolar Current

This current has the effect of isolating Antarctica from relatively warm ocean currents of the southern Pacific, Atlantic and Indian Oceans. That has kept warm water away from Antarctica and is responsible for the glaciation of the southern continent starting at about 35 Ma and continuing until today (with a possible interruption between 25 and 15 Ma) (Figure 3.2.4).

Between about 100 Ma and 10 Ma North and South America were separated from each other by a waterway hundreds of kilometers wide; under those conditions water was able to flow freely between the Pacific and Atlantic Oceans. But there was ongoing subduction of oceanic crust beneath what is now Central America. In a manner similar to what is shown on Figure 2.4.2, that process led to formation of magma above the subducting plate. That led to many millions of years of volcanic activity, and to the formation of a series of volcanic islands within what is now Central America (Figure 3.2.6). Finally, at around 10 Ma, those volcanic islands coalesced into an isthmus that opened the way for land animals to pass between North and South America but blocked the Central American Seaway.

Figure 3.2.6 The Final Stages in the Development of the Isthmus of Panama. At around 15 Ma the Nazca Plate was subducting (along the toothed lines) beneath North and South America, while a small part of the Caribbean Plate was subducting in the other direction beneath what is now Panama. The white triangles are possible locations of volcanoes. The dashed line shows the likely next step in the construction of the isthmus.

That change had the effect of making the Gulf Stream (and the entire Atlantic circulation system) more intense, and the warm water flowing north brought more warmth and more moisture to the northern Atlantic. Ironically, that additional warmth and moisture led to more intense snowfall in Iceland, Greenland, northern North America, and northern Europe, and thus to a lower albedo, and eventually to the beginning of the Pleistocene Glaciations. The northern hemisphere has been repeatedly glaciated since 2.5 Ma, in cycles that have remarkably regular periodicity. The origin of those cycles is discussed in Chapter 6.

Media Attributions

• Figure 3.2.1 Steven Earle, CC BY 4.0
• Figure 3.2.2 Photo by Isaac Earle, used with permission, CC BY 4.0
• Figure 3.2.3 Steven Earle, CC BY 4.0
• Figure 3.2.4 Steven Earle, CC BY 4.0, based on information from studies by James Hansen and others compiled by Root Routledge and posted at Alpine Analytics, http://alpineanalytics.com/Climate/DeepTime.html
• Figure 3.2.5 Steven Earle, CC BY 4.0
• Figure 3.2.6 Steven Earle, CC BY 4.0; based on León, S. et al. (2018). Transition from collisional to subduction‐related regimes: An example from Neogene Panama‐Nazca‐South America interactions. Tectonics, 37(1), 119-139. https://doi.org/10.1002/2017TC004785

Significant volumes of gases are emitted during volcanic eruptions, and some of these can have climate effects if the eruption is large. The most recent large volcanic event was the Pinatubo (Philippines) eruption in 1991. The amounts of three important gases emitted during that event are compared with the amounts of these same gases normally present in our atmosphere in Table 3.3.1. It’s easy to see that the amount of water emitted by an eruption such as Pinatubo is insignificant compared with the amount already in the atmosphere. It’s also important to recognize that water has a relatively short lifetime in the atmosphere (about 9 days), so most of the 400 million tonnes added by Pinatubo was likely rained out within a week and so would have had no climate impact.

Volcanic carbon dioxide emissions are also tiny compared with the current atmospheric reservoir, but CO₂ has a much longer residence time (hundreds to thousands of years), so there is the potential for volcanic CO₂ to lead to greenhouse-gas warming if a higher-than-average level of volcanism is sustained for centuries or more. The Pinatubo eruption lasted for less than a day, so there was no warming effect.

On the other hand, sulphur emissions from volcanic eruptions are typically large compared with the atmospheric reservoir of sulphur, and that’s why a major volcanic eruption can have a rapid and significant climate effect. In this case the effect is cooling, not warming, because sulphur gases get quickly converted to sulphate aerosols, tiny droplets or crystals that block incoming sunlight. Sulphate aerosols do not stay in the atmosphere for more than a few years in most cases, so the climate effect tends to be quite short.

The atmospheric effects of some recent volcanic eruptions are shown on Figure 3.3.1. At near-equatorial latitudes the Pinatubo eruption blocked almost 20% of incoming sunlight for several months and 10% of sunlight for almost 18 months. (Have a look at Exercise 3.2, below, and decide for yourself if and how the Pinatubo eruption affected the Earth’s climate.)

Figure 3.3.1 Evidence of Atmospheric Sulphate Aerosol Levels Associated With Major Volcanic Eruptions from 1960 to 2020 – Based on Solar Radiation Dimming . The inset shows the dimming associated with the 1991 Pinatubo eruption in a little more detail. The measurements shown on the graph were made at Mauna Loa, Hawaii, but the results are not influenced by volcanic activity in Hawaii.

Figure 3.3.2 Conceptual Model of Global Temperature Changes for One Million Years Following the Start of the Siberian Traps Eruption. The stars represent possible individual episodes of intense volcanism.

There have been numerous large volcanic eruptions in recent millennia, but none has affected the Earth’s climate by more than a few degrees C, nor for more than several years. Really significant climate change can be caused by really large eruptions and there have been some of those in the distant past. One example is the Deccan Traps in India at around 66 Ma. This event—which probably originated from a large mantle plume—involved the eruption of about 200,000 times as much magma as at Pinatubo—and correspondingly greater amounts of gases—over a period of about 30,000 years.

Although there would have been some cooling related to sulphur emissions, the main climate effect was significant warming from CO₂ emissions (up to 2⁰ C for as much as 500,000 years). However, the environmental effects are not fully understood because this coincided (and might have been partly caused by) the dinosaur-ending extra-terrestrial collision at the end of the Cretaceous.

The Siberian Traps eruption of 252 Ma was about four times larger than Deccan. It is estimated that about two-thirds of the magma erupted over approximately 300,000 years at the boundary between the Permian and Triassic Periods.There is evidence for strong warming, in the order of 10⁰ C, for at least the first ten million years of the Triassic. Figure 3.3.2 provides a conceptual model of how climate change might have progressed from the start of the Siberian Traps eruption (time 0) and then for the first million years of the Triassic. Assuming that the volcanic eruption proceeded episodically (which is typical), there would have been short periods of cooling caused by sulphate aerosol pulses, and increasingly intense warming caused by the progressive buildup of atmospheric carbon dioxide.

The Siberian Traps eruption coincided with the most catastrophic extinction of all time. Over 95% of all marine species and 70% of terrestrial species disappeared from the fossil record at the end of the Permian. Life on Earth was forever changed, or put differently, the future course of evolution—including the origin and evolution of mammals—was significantly affected by that event.

Exercise 3.2 Volcanism and the Climate

Figure 3.3.3 shows the global mean annual temperature for the period from 1970 to 2020, along with the timing of the large 1982 El Chichὀn and 1991 Pinatubo eruptions. Look closely at the temperature variations and decide whether, and how, those eruptions affected the global climate, and for how long.

Figure 3.3.3 Global Temperature Record for the Period, 1970 to 2020

Exercise answers are provided Appendix 2.

Media Attributions

• Figure 3.3.1 Steven Earle, CC BY 4.0, based on public domain data available at NOAA Earth System Research laboratory, Global Monitoring Division, https://www.esrl.noaa.gov/gmd/grad/mloapt.html
• Figure 3.3.2 Steven Earle, CC BY 4.0, based on a figure in Black, B. A., Neely, R. R., Lamarque, J. F. et al. (2018). Systemic swings in end-Permian climate from Siberian Traps carbon and sulfur outgassing. Nature Geoscience 11, 949–954.  https://doi.org/10.1038/s41561-018-0261-y
• Figure 3.3.3 Steven Earle, CC BY 4.0, based on public domain data from NASA Goddard Institute for Space Studies,  https://data.giss.nasa.gov/gistemp/tabledata_v4/GLB.Ts+dSST.txt

The Earth orbits around the sun in a nearly circular orbit, and it spins on an axis that is tilted at about 23.5⁰. The non-circular nature of the orbit varies over time, and both the tilt angle and the direction in which the spin axis points, also vary. It turns out that these small variations in the parameters of the Earth’s movement—known as Milanković cycles—have significant implications for our climate.

The shape of the Earth’s orbit around the sun is depicted on Figure 3.4.1. It is an elliptical shape, and the sun is not situated at the exact centre of that ellipse, but a little off to one side (this eccentricity is typical of all orbital relationships). On a consistent 100,000 year cycle, the shape changes from just a little bit elliptical, to a bit more elliptical, then back to just a little bit ellipitical.  When the orbit is more elliptical the eccentricity of the sun’s position within the orbit is also more extreme, and so the difference between the minimum and maximum Earth-sun distances is higher. The degree of eccentricity has an effect on our climate, and so, on a time scale of 100,000 years that effect increases and decreases.

Figure 3.4.1 A Representation of the Variations in the Elliptical Nature of Earth’s Orbit Around the Sun. The elliptical shapes of the orbits are greatly exaggerated in this diagram.

The second important feature of Earth’s motion is the tilt of the rotational axis (also known as obliquity) relative to the plane of the orbit around the sun (Figure 3.4.2). At present our axis is tilted at 23.5⁰ from “vertical”, but that varies from 22.1⁰ to 24.5⁰ and back to 22.1⁰ on a time scale of close to 41,000 years.

Figure 3.4.2 A Depiction of the Current Tilt of the Earth’s Axis

The tilt of the Earth’s rotational axis is what gives us seasons.  We have summer in the northern hemisphere when the Earth is in the part of its orbit where the northern hemisphere is pointed towards the sun, and summer in the southern hemisphere summer when the southern hemisphere is pointed towards the sun. At times of greater tilt, the seasons are slightly more exaggerated than they are now (colder winters, hotter summers).  At times of lesser tilt, they are less exaggerated (warmer winters, cooler summers). The tilt angle varies on a cycle of 40,000 years.

The third aspect of the Earth to consider is the variation in the direction of tilt (also known as precession). The Earth’s spin tends to keep it pointing in the same direction (just as a gyroscope tends to be stable), but in fact that direction is very slowly changing, such that in 23,000 years from now it will be pointing in the opposite direction. This cycle is important because it determines which hemisphere (north or south) will be farthest away from the sun during the summer, and that is really what drives glacial cycles.

The cyclical variations described above have implications for our climate, and especially for the advance and retreat of glaciers.  From year to year, and even over thousands of years, the amount of energy the whole Earth receives from the sun doesn’t change at all, but they do produce changes in the intensity of that solar energy at any particular latitude, and in the intensity at any particular time of year.  The Yugoslavian mathematician Milutin Milanković realized, for example, that glaciers grow best at temperate latitudes—in fact at around 65⁰ north or 65⁰ south—and that they can only start growing on land. 65⁰ north passes through Alaska, northern Canada, Greenland, Iceland, Scandinavia and Russia. It’s pretty much land the whole way. On the other hand, 65⁰ south is entirely in the Southern Ocean. There is almost no land at all, and so there is very little chance for glaciers to start forming in that area.

Based on this information Milanković decided that insolation variations at 65⁰ north were what mattered the most to the climate, and so he calculated the variations at that latitude. He also focused on insolation intensity in summer because he was aware that cool summers are more important than cold winters to the growth of glaciers. That may be counterintuitive, but it’s because less snow melts when the summers are cool, and also because cold winters tend to be drier than warm winters and so less snow falls.

The climate and glaciation effects of the three orbital parameters: eccentricity, tilt and tilt direction, are summarized in Table 3.4.1.

Milanković published his research on orbital cycles in 1941.  Alas, his theory was so far ahead of its time that there wasn’t enough evidence to demonstrate that it was reasonable. The first problem was that although it was widely accepted that glaciers had come and gone several times during the past million years, the timing of those events wasn’t well known. So, while Milanković was able to use his theory to estimate the timing of past glaciations, there was no way to verify his estimates. Another problem was that the calculated differences in insolation at 65⁰ N would not have been enough to drive glacial cycles. It turns out that the sceptics didn’t recognize the importance of climate feedbacks in amplifying the weak forcing of insolation differences.  For example, as it starts to cool at 65⁰ N more snow falls, and because the summers are especially cool, less of that snow melts. The reflectivity of surfaces goes up, and that leads to more cooling.  As the atmosphere cools so does the ocean, and that means that CO₂ becomes more soluble in the ocean water and more of it comes out of the atmosphere, leading to even more cooling.

During the middle part of the 20th century marine scientists and geologists started drilling into the soft sediments on the sea floor. The resulting cores provided a wealth of information about past marine life and conditions of sedimentation, and based on isotopic analyses, the temperature of the water. A key paper, published in 1976, clearly showed the relationship between those temperature variations and the astronomical cycles.  The authors wrote: “It is concluded that changes in the earth’s orbital geometry are the fundamental cause of the succession of Quaternary ice ages.”  This was the turning point for the Milanković concept.  Since then, thousands of studies of climate variations have corroborated the key role played by the Milanković cycles, both during the Quaternary glaciations and in other climate cycles (e.g., Monsoons) for millions of years before then.

It’s important to be clear that the Milanković cycles did not cause the Quaternary glaciations.  Instead, it was the long slow decline in global temperatures of the past 50 million years, which was mostly a result of enhanced weathering of the rocks in mountain ranges (as summarized on Figure 3.2.3). That long period of cooling brought the Earth to the point where glaciation was possible; since then the Milanković cycles have been the pacemaker of the glacial cycles.

In the 1970s a consortium of glaciologists from several countries started drilling through the thick glacial ice of Greenland and Antarctica. Over the next few decades, they recovered ice cores from holes extending down thousands of metres, retrieving samples of ice that had been laid down up to hundreds of thousands of years earlier. These cores not only allowed isotopic estimates of the temperatures at the time, but provided actual samples of the atmosphere (locked in bubbles in the ice) as it existed when the ice formed. And, unlike the deep-sea sediment cores, the ice cores have well-defined annual layers, so they can be time calibrated accurately (Figure 3.4.3).

The black line on Figure 3.4.3 shows the temperature of the ice that formed on Antarctica over the past 250,000 years (compared with today’s temperature).  The correlation between the temperature record and the July insolation levels is reasonably clear.  The 3rd-last interglacial—extending from 245,000 to 235,000 years—corresponds with a period of high insolation. The following very low insolation initiated the beginning of the 2nd-last glacial period. That was followed by a very high insolation period (at around 220,000 years)—which led to significant warming but wasn’t enough to break the glacial cycle. Glacial conditions then intensified over the next 90,000 years.

Figure 3.4.3 Insolation Levels and Antarctic Temperature Changes for the Past 250,000 Years. The temperature data are from the European Program for Ice-coring in Antarctica (EPICA) borehole at Dome C, Antarctica. For reference, the present-day average temperature at this site is about -50⁰ C. This core penetrated through ice as old as 800,000 years and is currently the longest ice-core record. A project to drill in an area with ice that could be as old as 1 million years, called “Beyond EPICA”, is in progress

Another period of very high insolation, culminating at around 120,000 years, was able to break the cycle, leading to the 2nd interglacial, which lasted from about 127,000 to 116,000 years ago. That was followed by a similar cycle of increasingly cold climates and strong glaciation until around 20,000 years ago, when the glacial cycle was again broken by a period of strong insolation.

Media Attributions

• Figure 3.4.1 Steven Earle, CC BY 4.0
• Figure 3.4.2 Steven Earle, CC BY 4.0
• Figure 3.4.3 Steven Earle, CC BY 4.0. The data shown here are from Jouzel, J. (2004).740,000-year Deuterium Record in an Ice Core from Dome C, Antarctica. EPICA Dome C Ice Cores Deuterium Data. IGBP PAGES, World Data Center for Paleoclimatology, Data Contribution Series # 2004 – 038. NOAA/NGDC Paleoclimatology Program, Boulder CO, USA. https://doi.org/10.3334/CDIAC/cli.007

Ocean currents are critically important in redistributing thermal energy and water on Earth. They have played important roles in past climate changes, and will be important in future climate changes.

The major surface currents of the world’s oceans are shown on Figure 3.5.1. Because of the Coriolis effect, currents tend to have a clockwise pattern in the northern hemisphere and a counter-clockwise pattern in the southern hemisphere. Currents that flow towards the equator generally bring cold water into warmer regions (blue arrows), while those that flow toward the poles bring warm water into colder regions (red arrows). In general, although not in all cases, currents that flow generally east-west have a neutral role when it comes to redistributing heat.

Figure 3.5.1 Major Surface Currents in the World’s Oceans. The “cold currents” bring cold water into relatively warm regions and the “warm currents” bring warm water into relatively cold regions.

The currents shown on Figure 3.5.1 are surface currents and, as such, they are confined to the upper 400 m of the oceans; most of the flow is within the upper 100 m. But there are also significant deep-flow currents, and those plays an equally important role in the redistribution of heat on the planet. Some of that deeper flow is shown on Figure 3.5.2. This is known as the thermohaline circulation system, where the term “thermohaline” implies that these flows are driven—at least in part—by both the temperature and the salinity of the water. These factors determine the density of the water in the currents, and that is critical to the thermohaline circulation.

Figure 3.5.2 Major Thermohaline Flow Patterns in the World’s Oceans. Cold salty surface water sinks in the north Atlantic, and deep water comes to surface in the Indian and the north Pacific oceans.

Pure water at 20⁰ C has a density of 998 grams/litre (g/L). Typical salty ocean water (3.5% salinity) at 20⁰ C has a density of 1025 g/L. Ocean water salinity ranges from about 3.3% in areas where there is a lot of rain or a lot of freshwater input from rivers, to about 3.8% where there is strong evaporation and relatively little freshwater input. The higher the salinity, the greater the density. Ocean water temperature ranges from around 30⁰ C in the tropics to a little less than 0⁰ C in polar regions. The lower the temperature, the greater the density, because a kilogram of cold water occupies less volume than a kilogram of warm water.

The water of the Gulf Stream flowing past Florida has a salinity of around 3.65%, a typical temperature of about 28⁰ C and density of about 1024 g/L. As this water flows north beyond Iceland its salinity drops only a little, to about 3.45% (due to rain and river input), while its temperature drops a lot, to about 2⁰ C. That cold salty water has a density of about 1028 g/L, making it the densest water anywhere in the open ocean, and quite a lot denser than the equally cold but much less salty water underneath. Because of this high density, the surface water in that region sinks and becomes part of deep flow system. It remains submerged as it moves south through the Atlantic and east past Africa, then it resurfaces either in the Indian Ocean—east of Madagascar—or in the northern part of the Pacific Ocean—north of Hawaii (Figure 3.5.2). This thermohaline circulation is important in controlling the Earth’s climate.

Glacial ice cores collected from Greenland and Antarctica have revealed some striking temperature variations over the past few hundred thousand years, especially during the more intense parts of the Quaternary glaciations, and it has been shown that much of that variability is related to changes in current flow patterns. Figure 3.5.3 shows the temperature record for the period from 44 to 26 ka, as determined in core samples from a drill hole in central Greenland.

Figure 3.5.3 Oxygen-Isotope Temperatures at the GISP2 Site for the Period From 44 to 26 Thousand Years Ago. The numbered peaks are Dansgaard-Oeschger events.

These temperature swings, in the order of 6 to 10⁰ C, on time scales of 1000 to 2000 years are known as Dansgaard-Oeschger cycles, in honour of the Danish (Willi Dansgaard) and Swiss (Hans Oeschger) scientists that first described them.

There is strong consensus, amongst marine scientists and climatologists, that the key factor in controlling Dansgaard-Oeschger cycles is variability in the salinity of the northern Atlantic Ocean. This process is known as the “salinity oscillator”. As already described, evaporation in the equatorial part of the Atlantic increases the salinity of the Gulf Stream. In the far north Atlantic, this still salty and now cold water sinks to become part of the subsurface flow that eventually comes back to surface in the Indian and Pacific Oceans (Figure 3.5.2). That represents a net movement of salt out of the Atlantic basin, and gradually (over hundreds of years) results in a decrease in the overall salinity of the Atlantic water. Another part of the process is related to the northward transportation of heat by the Gulf Stream, which makes the Arctic region warmer than it would be otherwise, and that leads to more melting of glacial ice in Greenland and northern Canada, further diluting the Atlantic.

As the Atlantic slowly becomes less salty and as the melting of northern glaciers make it fresher still, the tendency for the cooled Gulf Stream water to sink in the far north Atlantic is reduced, and therefore the strength of the overall thermohaline circulation system decreases. That means that less heat is transported north, the Arctic region cools, less salt is removed from the Atlantic basin, and glacier-melting slows so that less freshwater flows into the ocean.

The salinity oscillator process is illustrated on Figure 3.5.4, which is focused on Dansgaard-Oeschger events 6 and 7. The dashed curve is not based on data; it is simply a representation of slow changes in the salinity of the Atlantic Ocean and the strength of the thermohaline circulation (THC). When the THC is strong—because the salinity is high—western Europe and the Arctic region are warmed. A strong THC tends to slowly reduce Atlantic salinity, both because salty water is moved out of the Atlantic basin and because Arctic melting is enhanced. This slowly weakens the THC, and the Arctic cools. With a weaker THC and less Arctic melting, the salinity slowly increases. The entire cycle takes about 1500 years on average, although in the case illustrated the time between peaks is 1750 years.

Figure 3.5.4 The Driving Mechanisms and Processes of the North Atlantic Salinity Oscillator

Exercise 3.3 Ocean Water Densities

The density of ocean water is a function of its temperature and salinity, and the relationship is shown on Figure 3.5.5, where density (red lines) is expressed in g/L. For example, at a salinity of 2.5% and a temperature of 10⁰ C, the density is 1027 g/L. The table has a list of salinities and typical water temperatures for some offshore locations, some within the Gulf Stream in the Atlantic and two within the Pacific. Use the graph to estimate the water densities at these locations.

Figure 3.5.5 Water Density as a Function of Salinity and Temperature

Salinities and Typical Water Temperatures for Some Offshore Locations

Location	Salinity (%)	T (° C)	Density (g/L)
North Carolina	3.64	20
Newfoundland	3.58	15
Iceland	3.52	9
Svalbard	3.45	2
Baja Sur	3.4	20
Los Angeles	3.35	13

Exercise answers are provided Appendix 2.

Media Attributions

• Figure 3.5.1 Steven Earle, CC BY 4.0, based on a US Government public domain image, Corrientes-oceanicas, by Dr. Michael Pidwirny, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Corrientes-oceanicas.png
• Figure 3.5.2 Steven Earle, CC BY 4.0, based on maps of thermohaline circulation from various sources.
• Figure 3.5.3 Steven Earle, CC BY 4.0, data, public domain, from NOAA, https://www.ncdc.noaa.gov/data-access/paleoclimatology-data/datasets/ice-core, originally described by: Alley, R. B. (2000). The younger dryas cold interval as viewed from central Greenland. Quaternary Science Reviews, 19(1-5), 213-226. https://doi.org/10.1016/S0277-3791(99)00062-1. The Dansgaard-Oeschger cycles: Dansgaard, W. (1984) North Atlantic climate oscillations revealed by deep Greenland ice cores. In Hanson, J. E. & T. Takahashi (Eds.), Climate processes and climate sensitivity. American Geophysical Union, 288-298; and by: Dansgaard, W., Johnsen, S., Clausen, H. et al. (1993). Evidence for general instability of past climate from a 250-kyr ice-core record. Nature 364, 218–220. https://doi.org/10.1038/364218a0
• Figure 3.5.4 Steven Earle, CC BY 4.0, from the same data source as Figure 3.5.3.
• Figure 3.5.5 Steven Earle, CC BY 4.0

It is estimated that the Chicxulub object was approximately 12 km in diameter and that it hit the Earth at about 100,000 km/h. It landed in an area of relatively shallow sea water (a few hundred metres depth) that was underlain by about 5 km thickness of limestone and evaporite rocks. Limestone is primarily calcium carbonate (CaCO₃), and a major component of the evaporite layer is calcium sulphate (CaSO₄). The impactor exploded through those sedimentary layers in a fraction of a second, and then excavated a hole nearly 100 km wide and 20 km deep into the granitic rocks of the crust (Figure 3.6.1). All of that material was immediately melted or vaporized, and the solid particles were blasted out of the atmosphere.

Figure 3.6.1 A Depiction of the Crater that Formed Within Seconds of the Impact of an Approximately 12 km Diameter Object at Chicxulub, Mexico at 65 Ma.

The sudden change in the level of the sea floor within and around the crater created an enormous tsunami that spread across the Gulf of Mexico and out into the Atlantic and to the Pacific (as this predates the Isthmus of Panama). Within the gulf, the wave would have been as much as 1500 m high, and it was likely more than 15 m high in both the Pacific and Atlantic basins. This wreaked massive destruction in the extensive low-lying areas around the gulf, and then the water rebounded back towards Chicxulub, refilling much of the crater with sediments from around the region and with debris from the blast. The back-filled material included a large volume of charcoal, presumably from wildfires in areas around the gulf.

The rock fragments and glass of the impact plume were sent aloft through the atmosphere and dispersed across an area with a radius of at least 6000 km. On re-entry, friction turned this material into glowing bodies with enough intensity to outshine the sun by a factor of seven times, and enough heat to start wildfires within a radius of at least 6000 km from the impact site (i.e., all of North and South America), and possibly almost everywhere on Earth. Animals that could not hide underground or in water, would likely have perished in the resulting fires, if they hadn’t already been killed by the incandescent heat.

Wildfires of the magnitude described above would have created a massive amount of smoke, and the layer of soot from that smoke has been found in K-Pg boundary-layer deposits around the world. It amounts to approximately 15 giga tonnes (Gt), which is many hundreds of times the amount of soot produced by wildfires in a typical year, and is enough to have effectively blocked most of the incoming sunlight.  A digital model of the early-Paleogene climate following the emission of 15 Gt of wildfire soot shows that for several years the average insolation at the Earth’s surface was likely less than 1% of normal, and for low-latitude regions it was probably only a fraction of that. In other words, it was permanently dark—although not quite as dark as night-time—for years. Under those conditions, photosynthesis, and therefore growth of plants, on both land and in the oceans, was effectively stopped. Mean annual temperature in most continental areas, including equatorial regions, dropped to less than 0⁰ C, or 10 to 15⁰ C below normal (Figure 3.6.2), in some temperate and polar areas it was much less, although it probably stayed above freezing over tropical and temperate oceans. Precipitation dropped to about 20% of normal levels, typical of deserts in most regions. It is likely that the darkness started to lift after about two years, but the cold and dry conditions persisted for six to eight years.

Figure 3.6.2 Modeled Temperature and Precipitation Changes in the 15 Years Following the K-Pg Impact

Needless to say, survival would have been difficult under such conditions, even for the fittest. Animals that had survived the heat blast and the raging wildfires, would have emerged from burrows, crevices and swamps to find permanent near darkness, no new plant growth, wicked cold and, soon, almost no fresh water.

And that’s not all, because the modelling described above didn’t account for the 650 Gt of sulphur dioxide that would have been formed by the instantaneous vaporization of the thick beds of calcium sulphate—roughly 100 times the amount produced by the climate-cooling eruption of Pinatubo in 1991. This SO₂ would have quickly been converted to H₂SO₄ droplets (sulphuric acid) in the atmosphere, which as long as they stayed up there (likely a few years), would have made the cooling even stronger.  When significant rain finally returned, likely sometime in the sixth or seventh years, it would have been acidic.

In addition to soot, the wildfires would have produced a massive amount of CO₂, and CO₂ was also emitted by the vaporization of limestone at the impact site. Analysis of fish remains in a section spanning the K-Pg boundary in Tunisia provides evidence that—once the dust and sulphate aerosols had settled—there was an approximate 5⁰ C of CO₂-warming of the climate, and that warming lasted for about 100,000 years.

To summarize, the collision of a 12 km diameter extraterrestrial object with the Earth on the last day of the Cretaceous Period appears to have had both immediate and long-term climate effects.  First, there was intense heat from incoming solid impact ejecta. This lasted for several hours and was enough to kill any exposed organisms and start wildfires that covered entire continents. Next came several years of darkness, strong cooling and aridity, followed by acid rainfall. Finally, there was strong warming that lasted for about 100,000 years.

Ever since the first mention of the concept of the dinosaur-ending K-Pg impact in 1980, geologists have been looking for evidence that other extinction events might have been caused by similar events. So far, nothing convincing has been found, but another question that has been frequently asked since then is whether we could be in store for another event like that in the future.

An international program of tracking of potential “near Earth objects” (NEOs) has been in progress for the past 25 years.  Over that time thousands of objects have been identified and their orbits characterized.  A potentially hazardous NEO is one that is projected to come closer to Earth than 20 times the distance between the Earth and Moon. The current focus is on all NEOs that are over 140 m in diameter, as those are considered to represent significant hazards to people and infrastructure.

According to the NASA Jet Propulsion Laboratory there are estimated to be about 1000 NEOs greater than 1 km in diameter and about 15,000 greater than 140 m. As of May 2020, 731 NEOs greater than 1 km and 8,827 greater than 140 m have been discovered.. Of these, only two have a significant probability of coming close to the Earth. The probabilities of those actually hitting the Earth are in the order of 1 in 10,000, and the estimated dates are more than a century away. That represents a very small risk, but it’s important to remember that it is likely that there are still hundreds of other NEOs that have not yet been discovered.

So, although the risk of a large impact seems to be small, the implications for us on Earth of being hit by something several kilometres across are so extreme that it really is a good idea to keep an eye on the sky.

Media Attributions

• Figure 3.6.1 Steven Earle, CC BY 4.0, based on a diagram in: Gulick, S., Bralower, T. J., et al… Expedition 364 Scientists (2019). The first day of the Cenozoic. Proceedings of the National Academy of Sciences of the United States of America, 116(39), 19342–19351. https://doi.org/10.1073/pnas.1909479116
• Figure 3.6.2 Steven Earle, CC BY 4.0, based on drawings in Bardeen, C., Garcia, R., Toon, O., & Conley, A. (2017). On transient climate change at the Cretaceous-Paleogene boundary due to atmospheric soot injections. Proceedings of the National Academy of Sciences (PNAS) Sep 2017, 114(36), E7415-E7424; https://doi.org/10.1073/pnas.1708980114

The topics covered in this chapter can be summarized as follows:

We are currently in the middle of a glacial period that started about 34 million years ago, but became more intense about one million years ago. During that time glaciers have expanded and contracted on a time scale of around 100,000 years. As described in Chapter 3, this is not the only period of glaciation in Earth’s history; there have been many in the distant past, and the more significant ones are shown on Figure 4.1.1. In general, however, the Earth has been warm enough to be ice-free for much more of the time than it has been cold enough to be glaciated.

Figure 4.1.1 The Record of Major Glaciations During Earth’s History

The oldest known glacial period is the Huronian, and as we’ve seen, it was likely initiated because of the evolution of photosynthetic organisms. This resulted in an increase of free oxygen in the atmosphere and a consequent drop in the levels of the greenhouse gas methane—which triggered cooling. Based on evidence of glacial deposits from the area around Lake Huron in Ontario (and elsewhere), the Huronian glaciation lasted from approximately 2450 to 2400 Ma. Because rocks of that age are rare, we don’t know much about its intensity or global extent.

Late in the Proterozoic the climate cooled dramatically and the Earth was gripped by what appears to be the most intense glaciation ever. This may have been a result of a concentration of continents near to the equator, which had an albedo-related cooling effect (See figure 3.2.1). The glaciations of the Cryogenian Period (cryo is Latin for icy-cold) are also known as the Snowball Earth glaciations, because it is hypothesized that the entire planet was frozen—even in equatorial regions—with ice on the oceans up to several hundred metres thick. A visitor to our planet at that time might not have held out much hope for its habitability, although life still survived in the oceans. There were two main glacial periods within the Cryogenian, the Sturtian from 720 to 660 Ma, and the Marinoan from 645 Ma to 635 Ma. The end of the Cryogenian glaciations coincides with the evolution of relatively large and complex life forms on Earth. Some geologists think that the changing environmental conditions of the Cryogenian are what triggered the evolution of large and complex life first observed in the rocks of the late Proterozoic and then continuing with so-called “explosion” of life forms in the Cambrian.

As shown on Figure 4.1.1, there have been four major glaciations during the Phanerozoic (the past 540 million years), including the Andean/Saharan, the Late Devonian, the Karoo and the Cenozoic glaciations.

Evidence for the Andean-Saharan glaciation is found in rocks of the Andean region of South America and in the Sahara region of Africa during the Ordovician and Silurian (Figure 4.1.2). The glaciation may have lasted for up to 10 million years (from about 435 to 445 Ma) at a time when the supercontinent Gondwana was mostly close to the South Pole. Eyles (2008) has suggested that the cooling that triggered this glacial episode might have resulted from the formation of a major mountain range in the northern part of what is now Africa, which, as discussed in Chapter 3, would have led to enhanced erosion and weathering, and therefore to consumption of atmospheric carbon dioxide.

Figure 4.1.2 The Possible Extent of the Ordovician-Silurian Glaciation (white) on Gondwana at Around 440 Ma

Eyles (2008) describes the glaciation in the late Devonian (around 375 Ma) as being “short-lived”, although the volume of ice may have been similar to what is present on Earth now. This glaciation, which is recorded in rocks of Bolivia and Brazil, is also though to have been the result of uplift related to mountain building at continent-continent collisions associated with the formation of the Pangea supercontinent.

The Karoo was the longest of the Phanerozoic glaciations, and perhaps the longest of all known glaciations (although not the coldest), persisting from approximately 360 to 260 Ma – which includes the latest Devonian, all of the Carboniferous and much of the Permian. For most of that time that the southern (Gondwana) part of Pangea was still situated over the South Pole (Figure 4.1.3). The glaciation is named after a sequence of rocks in southern Africa (Karoo Supergroup), but there is abundant evidence for coincident glaciation in South America, Australia and Antarctica.

Figure 4.1.3 The Likely Extent (white) of the Karoo Glaciation Between 360 and 260 Ma

In the late 20th century Yale University geochemist Robert Berner devised a way to use geological data, including sediment volumes and isotopic compositions, to estimate the composition of the atmosphere during the Phanerozoic (the last 540 million years). The results of that work are summarized on Figure 4.1.4, along with the durations of the major Phanerozoic glaciations. The Karoo glaciation is coincident with a period of persistent significantly low CO2 levels, which, according to Berner are a result of enhanced weathering and the consequent consumption of CO2. The strong weathering is mostly ascribed to crustal uplift from collisions, but is also a product of vigorous plant growth at this time. The world’s first forests evolved at around 385 Ma, and the roots of those trees are thought to have played a significant role in loosening the soil and breaking up near-surface rock, making it more susceptible to weathering because tree roots loosen the soil and break up near-surface rocks, making both more susceptible to chemical weathering. Forests would also have had a direct role in removing CO2 from the atmosphere, and some of that would have been permanently sequestered in organic-rich sediments and coal (which first started to accumulate at around this time).

Figure 4.1.4 Calculated Atmospheric CO2 Levels and Timing of Glaciations During the Phanerozoic

The Earth was warm and essentially unglaciated throughout the Mesozoic. Although there may have been some alpine glaciation at this time, there is no longer any record of it; the dinosaurs, which dominated terrestrial habitats during the Mesozoic, did not have to endure icy conditions.

A warm climate persisted into the Cenozoic, in fact there is evidence that the late Paleocene and early Eocene (~55 to 45 Ma) were the warmest parts of the Phanerozoic since the Cambrian (Figure 3.2.3). As discussed in section 3.4, a number of tectonic events during the Cenozoic contributed to persistent and significant planetary cooling since 45 Ma. For example, the collision of India with Asia, and the formation of the Himalayan range and Tibetan Plateau, resulted in a dramatic increase in the rate of weathering and erosion. Higher than normal rates of weathering of rocks with silicate minerals, especially feldspar, consumes carbon dioxide from the atmosphere and therefore reduces the greenhouse effect, resulting in long-term cooling.

At 40 Ma ongoing plate motion widened the narrow gap between South America and Antarctica resulting in the opening of the Drake Passage. This allowed for the unrestricted west-to-east flow of water around Antarctica, the Antarctic Circumpolar Current (Figure 3.2.4), which effectively isolated the Southern Ocean from the warmer waters of the Pacific, Atlantic and Indian Oceans. The region cooled significantly, and by 35 Ma (Oligocene) glaciers had formed on Antarctica, the first in over 200 million years.

Global temperatures remained relatively steady during the Oligocene and early Miocene, and the Antarctic glaciation waned during that time. At around 15 Ma subduction-related volcanism between central and South America created the connection between North and South America, preventing water from flowing between the Pacific and Atlantic Oceans (Figure 3.2.5). This further restricted the transfer of heat from the tropics to the poles leading to a rejuvenation of the Antarctic glaciation. The expansion of that ice sheet increased the Earth’s reflectivity enough to promote a positive feedback loop of further cooling: more reflective glacial ice, more cooling, more ice, etc. Ice sheets started to grow in Greenland by around 3.5 Ma, and in North America and northern Europe by around 1 Ma (Figure 4.1.5). The most intense part of the current (“Pleistocene”) glaciation—and the coldest climate—was during the Pleistocene, but if we count Antarctic glaciation it really extends from the Oligocene to the Holocene, and will likely continue into the future.

The Pleistocene has been characterized by significant temperature variations (through a range of approximately 8˚ C) on time scales of 40,000 to 100,000 years, and to corresponding expansion and contraction of ice sheets in the northern hemisphere. These variations are attributed to subtle changes in the orbital parameters of the Earth (Milankovitch Cycles), which are described in section 3.4. Over the past million years the glaciation cycles have been approximately 100,000 years, and this variability is clearly visible on Figure 4.1.5.

Figure 4.1.5 Foram Oxygen Isotope Record for the Past 5 Million Years. Based on O isotope data from sea-floor sediments.

Exercise 4.1  Pleistocene Glacials and Interglacials

Figure 4.1.6 shows the past 500,000 y of the same data set shown in Figure 4.1.5 . The last five glacial periods are marked with snowflakes. The most recent one, which peaked at around 20 ka, is known as the Wisconsin Glaciation. Describe the nature of temperature changes that led up to each of these glacial periods, and how the temperature changed in the immediate post-glacial period.

Figure 4.1.6 Temperature Fluctuations Over the Past 500,000 Years

The current interglacial (Holocene) is marked with an H. Identify the previous five interglacial periods.

Exercise answers are provided Appendix 2.

At the height of the last glaciation (Wisconsin Glaciation) massive ice sheets covered virtually all of Canada and much of the northern United States (Figure 4.1.7). The massive Laurentide Ice Sheet covered most of eastern Canada, as far west as the Rockies, and the smaller Cordilleran Ice Sheet covered most of the western region. At various other glacial peaks during the Pleistocene and Pliocene the ice extent was similar to this, and in some cases even more extensive. The combined Laurentide and Cordilleran Ice Sheets were comparable in volume to the current Antarctic Ice Sheet.

Figure 4.1.7 The Extent of the Cordilleran and Laurentide Ice Sheets Near to the Peak of the Wisconsin Glaciation, Around 18.5 ka

Media Attributions

• Figure 4.1.1 Steven Earle, CC BY 4.0, after the International Geological Timescale, https://stratigraphy.org/chart
• Figure 4.1.2 Steven Earle, CC BY 4.0, after Eyles, N. (2008)
• Figure 4.1.3 Steven Earle, CC BY 4.0, after Eyles, N. (2008)
• Figure 4.1.4 Steven Earle, CC BY 4.0, based on data in: Berner, R. A., & Kothavala, Z. (2001). GEOCARB III: A revised model of atmospheric CO2 over Phanerozoic time. American Journal of Science, 301, 182–204. https://doi.org/10.2475/ajs.301.2.182
• Figure 4.1.5 Steven Earle, CC BY 4.0) from Lisiecki, L. E., & Raymo, M. E. (2005). A Pliocene-Pleistocene stack of 57 globally distributed benthic d18O records. Paleoceanography, 20, PA1003. https://doi.org/10.1029/2004PA001071
• Figure 4.1.6 Steven Earle, CC BY 4.0, using data from Lisiecki & Raymo (2005)
• Figure 4.1.7 Steven Earle, CC BY 4.0, based on a public domain NOAA map, Paleo Glaciation, https://www.ncdc.noaa.gov/paleo/glaciation.html

There are two main types of glaciers. Continental glaciers (a.k.a. ice sheets) cover vast areas of land and only exist now in extreme polar regions, including Antarctica and Greenland (Figure 4.2.1). Alpine glaciers (a.k.a. valley glaciers) originate on mountains, mostly in temperate and polar regions, but even in tropical regions if the mountains are high enough. They are typically confined to valleys.

Figure 4.2.1 Part of the Continental Ice Sheet in Greenland With Some Outflow Valley Glaciers in the Foreground

The Earth’s two great continental glaciers, on Antarctica and Greenland, comprise about 99% of all of the world’s glacial ice, and approximately 68% of all of the Earth’s fresh water. As is evident from Figure 4.2.2, the Antarctic ice sheet is vastly bigger than the Greenland ice sheet; it contains about 17 times as much ice. If the entire Antarctic ice sheet was to melt, sea level would rise by about 80 m and almost all of the Earth’s major cities would be submerged.

Figure 4.2.2 Simplified Cross-Sectional Profiles the Continental Ice Sheets in Greenland and Antarctica. Both drawn to the same scale.

Continental glaciers do not flow “downhill” because the large areas that they cover are generally flat. Instead, ice flows from the region where it is thickest towards the edges where it is thinner. This is shown schematically on Figure 4.2.3. It means that in the central thickest parts the ice flows almost vertically down towards the base, while in the peripheral parts it flows outwards towards the margins. In continental glaciers like Antarctica and Greenland, the thickest parts (4000 and 3000 m respectively) are the areas where the rate of snowfall (and therefore of ice accumulation) are highest.

Figure 4.2.3 Schematic Ice Flow Diagram for the Antarctic Ice Sheet

The flow of alpine glaciers is primarily controlled by the slope of the land beneath (Figure 4.2.4). In the zone of accumulation the rate of snowfall is greater than the rate of melting. In other words, not all of the snow that falls each winter melts during the following summer, and the ice surface is always covered with snow. In the zone of ablation more ice melts than accumulates as snow. The equilibrium line marks the boundary between the zones of accumulation and ablation.

Figure 4.2.4 Schematic Ice Flow Diagram for an Alpine Glacier (Steven Earle, CC BY 4.0)

The equilibrium line of the Overlord Glacier, near to Whistler BC, is shown in the photo on Figure 4.2.5. Below that line, in the zone of ablation, bare ice is exposed because last winter’s snow has all melted; above that line the ice is still mostly covered with snow from the last winter. The position of the equilibrium line changes from year to year as a function of the balance between snow accumulation in the winter and snow-melt during the summer. More winter snow and less summer melting obviously favours the advance of the equilibrium line (and of the glacier’s leading edge), but of these two variables it is the summer melt that matters most to a glacier’s budget. Cool summers promote glacial advance and warm summers promote glacial retreat.

Figure 4.2.5 The Approximate Location of the Equilibrium Line (red) in September 2013 on the Overlord Glacier, Near to Whistler, BC.

Above the equilibrium line of a glacier not all of the winter snow melts in the following summer, and thus snow gradually accumulates. The snow layer from each year is covered and compacted by subsequent snow and it gradually gets compressed and turned into firn within which the snowflakes lose their delicate shapes and become granules. With more compression the granules are pushed together and air is squeezed out. Eventually the granules are “welded” together to create glacial ice. Downward percolation of water, from melting taking place at surface, contributes to the process of ice formation (Figure 4.2.6).

Figure 4.2.6 A Depiction of the Evolution of Snowflakes into Glacial Ice

Glaciers move because the surface of the ice is sloped. This generates a stress on the ice, which is proportional to the slope and the depth below the surface. As shown on Figure 4.2.7, the stresses are quite small near to the ice surface but much larger at depth, and also greater in areas where the ice surface is relatively steep. Ice will deform, meaning that it will behave in a plastic manner, at stress levels of around 100 kilopascals, and so it’s evident that, in this case, the upper 50 to 100 m of the ice (above the dashed red line) is not plastic (it is rigid) while the lower ice is plastic and will flow. The rigid layer will be thinner where the ice surface is steeper and thicker where it is flatter.

When the lower ice of a glacier flows it moves the upper ice along with it, so although it might seem from the stress patterns (red numbers and red arrows) shown on Figure 4.2.7 that the lower part should move the most, in fact while the lower part deforms (and flows) and the upper part doesn’t deform at all, the upper part moves the fastest because it is pushed along by the lower ice.

Figure 4.2.7 Stress Within a Valley Glacier (red numbers) as Determined from the Slope of the Ice Surface and the Depth Within the Ice. The ice will deform and flow where the stress is greater than 100 kilopascals, and the relative extent of that deformation is depicted by the red arrows. Any deformation motion in the lower ice will be transmitted to the ice immediately above it, so although the red arrows get shorter towards the top, the ice velocity increases upwards (blue arrows). The upper ice (above the red dashed line) does not flow, but it is pushed along with the lower ice.

The plastic lower ice of a glacier can flow like a very viscous fluid, and it can therefore flow over irregularities in the base of the ice, and also around corners. But the upper rigid ice cannot flow in this way, and because it is being carried along by the lower ice it tends to crack where the lower ice has to flex. This leads to the development of crevasses in areas where the rate of flow of the plastic ice is changing. In the area shown on Figure 4.2.8, for example, the glacier is speeding up over the steep terrain, and the rigid surface ice has to crack to account for the change in velocity.

Figure 4.2.8 Crevasses on Overlord Glacier in the Whistler area, BC

The base of a glacier can be cold (below the freezing point of water) or warm (above the freezing point). If it is warm there will likely be a film of water between the ice and the material underneath, and the ice will be able to slide over that surface. This is known as basal sliding (Figure 4.2.9, left). If the base is cold the ice will be frozen to the material underneath and it will be stuck—unable to slide along its base. In this case all of the movement of the ice will be by internal flow (Figure 4.2.9, right).

Figure 4.2.9 Differences in Glacial Ice Motion with Basal Sliding (left) and Without Basal Sliding (right). The dashed red line indicates the upper limit of plastic internal flow.

One of the factors that affects the temperature of the base of a glacier is the thickness of the ice. Ice is a good insulator. The slow transfer of heat from the Earth’s interior will provide enough heat to warm up the base if the ice is thick, but not enough if it is thin and that heat can escape. It is typical for the leading edge of an alpine glacier to be relatively thin (see Figure 4.2.7), and so it is common for that part to be frozen to its base while the rest of the glacier is still sliding. This is illustrated on Figure 4.2.10 for the Athabasca Glacier. Because the leading edge of the glacier is stuck to its frozen base, while the rest continues to slide, the ice coming from behind has pushed (or thrust) itself over top of the part that is stuck fast.

Figure 4.2.10 Thrust Faults at the Leading Edge of the Athabasca Glacier, Alberta. The arrows show how the trailing ice has been thrust over the leading ice. The dark vertical stripes are mud from sediments that have been washed off of the lateral moraine lying on the surface of the ice.

Just as the base of a glacier moves slower than the surface, the edges, which are more affected by friction along the sides, move slower than the middle. If we were to place a series of markers across an alpine glacier and come back a year later we would see that the ones in the middle have moved further forward than the ones near to the edges (Figure 4.2.11).

Figure 4.2.11 Markers on an Alpine Glacier Will Move Forward Over a Period of Time

Glacial ice always moves downhill (or down from an area of thicker ice in the case of continental glaciers), in response to gravity, but the front edge of a glacier is almost always melting or else calving into water (shedding icebergs). Alpine glaciers can flow up over bumps in the terrain if the ice is thick enough. If the rate of forward motion of the glacier is faster than the rate of ablation (melting) the leading edge of the glacier will advance (move forward). If the rate of forward motion is about the same as the rate of ablation, the leading edge will remain stationary, and if the rate of forward motion is slower than the rate of ablation, the leading edge will retreat (move backward). Even if a glacier is retreating, the ice of the glacier will be moving forward.

Calving of icebergs is an important process for glaciers that terminate in lakes or the ocean. An example is shown on Figure 4.2.12.

Figure 4.2.12 Icebergs in Jökullsaárlón (A Pro-Glacial Lake) at the Front of Breiđmerkurjökull (Breiđmerkur Glacier) in Iceland

Exercise 4.2 Moving Ice

Figure 4.2.13 represents a glacier onto the surface of which some markers (yellow dots) have been placed to determine the rate of ice motion over a one-year period. The ice is flowing from left to right.

Figure 4.2.13 Glacial advance and retreat

In the middle diagram the leading edge of the glacier has advanced. Draw in where the markers might have moved to.

In the lower diagram the leading edge of the glacier has retreated. Draw in where the markers might have moved to.

Exercise answers are provided Appendix 2.

Media Attributions

• Figure 4.2.1 Photo by Ginny McLean, used with permission, CC BY 4.0
• Figure 4.2.2 Steven Earle, CC BY 4.0
• Figure 4.2.3 Steven Earle, CC BY 4.0
• Figure 4.2.4 Steven Earle, CC BY 4.0
• Figure 4.2.5 Modified by Steven Earle, CC BY 4.0, from a photo by Isaac Earle, 2013. Used with permission.
• Figure 4.2.6 Steven Earle, CC BY 4.0
• Figure 4.2.7 Steven Earle, CC BY 4.0
• Figure 4.2.8 Photo by Isaac Earle, used with permission, CC BY 4.0
• Figure 4.2.9 Steven Earle, CC BY 4.0
• Figure 4.2.10 Steven Earle, CC BY 4.0
• Figure 4.2.11 Steven Earle, CC BY 4.0
• Figure 4.2.12 Photo by Isaac Earle, 2015, used with permission, CC BY 4.0
• Figure 4.2.13 Steven Earle, CC BY 4.0

Glaciers are effective agents of erosion, especially in situations where the ice is not frozen to its base and therefore can slide over the bedrock or other sediment.  The ice itself is not particularly effective at erosion because it is relatively soft; instead, it is the rock fragments embedded in the ice that are pushed down onto the underlying surfaces and do most of the erosion.  A useful analogy would be to imagine the effect of rubbing a piece of paper against a wooden surface—not much happens—but, if it is a piece of sandpaper with embedded angular fragments of garnet, the wood will be significantly abraded.

Some of the results of glacial erosion are different in areas with continental glaciation versus alpine glaciation.  Continental glaciation tends to produce relatively flat bedrock surfaces, especially where the rock beneath is uniform in strength and there hasn’t been tectonic activity for hundreds of millions of years.  In areas with recent tectonic activity and where there are differences between the strength of rocks a glacier will obviously tend to erode the softer and weaker rock more effectively than the harder and stronger rock.  Much of central and eastern Canada, which was completely covered by the huge Laurentide Ice Sheet at various times during the Pleistocene, has been eroded to a relatively flat surface. In contrast, the Cordilleran Ice Sheet eroded and accentuated deep valleys and plateaus in the mountainous regions of western Canada and northwestern United States.  In many cases the existing relief is also due the presence of glacial deposits—such as drumlins, eskers and moraines (all discussed below)—rather than to differential erosion (Figure 4.3.1).

Figure 4.3.1 Drumlins—Streamlined Hills Formed Beneath a Glacier, Here Comprised of Sediment—In the Amundsen Gulf Region of Nunavut, Canada. The drumlins are tens of metres high, a few hundred metres across and a few kilometres long. One of the drumlins in this view has been highlighted with a yellow dotted line. Ice-flow was from the upper left towards the lower right.

Alpine glaciers produce very different topography than continental glaciers, and much of the topographic variability of mountains in temperate regions can be attributed to glacial erosion. In general, glaciers are much wider than rivers of similar length, and since they tend to erode more at their bases than their sides, they produce wide valleys with relatively flat bottoms and steep sides—known as U-shaped valleys.  Howe Sound, north of Vancouver, was occupied by a large glacier that originated in the Squamish, Whistler and Pemberton areas, and then joined the much larger glacier in the Strait of Georgia.  Howe Sound and most of its tributary valleys have a pronounced U-shaped profiles (Figure 4.3.2).

Figure 4.3.2 The View Down the U-shaped Valley of Mill Creek Valley Towards the U-Shaped Valley of Howe Sound, British Columbia, With the Village of Britannia on the Opposite Side

U-shaped valleys and their tributaries provide the basis for a wide range of alpine glacial topographic features, examples of which are visible on the Space Station view of the Swiss Alps shown on Figure 4.3.3.  This area was much more intensely glaciated during the past glacial maximum.  At that time the large U-shaped valley in the lower right was occupied by glacial ice, and all of the other glaciers shown here were longer and much thicker than they are now.  But even at the peak of Pleistocene glaciation some of the higher peaks and ridges would have been exposed and not directly affected by glacial erosion.  In these areas, and in the areas above the glaciers today, most of the erosion is related to freeze-thaw effects.  A peak that extends above the surrounding glacier is called a nunatuk.

Some of the important features visible on Figure 4.3.3 are arêtes: sharp ridges between U-shaped glacial valleys, cols: low points along arêtes that constitute passes between glacial valleys, horns: steep peaks that have been glacially and freeze-thaw eroded on three or more sides, cirques: bowl-shaped basins that form at the head of a glacial valley, hanging valleys: U-shaped valleys of tributary glaciers that hang above the main valley because the larger main-valley glacier eroded more deeply into the terrain, and truncated spurs: the ends of arêtes (a.k.a. “spurs”) that have been eroded into steep triangle-shaped cliffs by the glacier in the corresponding main valley.

Figure 4.3.3 A Space Station View of the Swiss Alps in the Area of the Aletsch Glacier. A variety of alpine glacial erosion features are labelled.

Some of these alpine-glaciation erosional features are also shown on Figure 4.3.4 in diagram form.

Figure 4.3.4 Some of the Important Alpine-Glaciation Erosion Features.

Exercise 4.3  Identify Glacial Erosion Features

Figure 4.3.5 shows Mt. Assiniboine in the BC Rockies . What are the features at locations a through e?  Look for one of each of the following:  a horn, an arête, a truncated spur, a cirque and a col.  Try to identify some of the numerous other arêtes in this view, as well as another horn.

Figure 4.3.5 Mt. Assiniboine, British Columbia

Exercise answers are provided Appendix 2.

A number of other glacial erosion features exist at smaller scales.  For example, a drumlin is an elongated feature that is streamlined at the down-ice end (Figure 4.3.6).  The one shown is larger than most others, and is almost entirely made up of rock.  Drumlins comprised of glacial sediments are very common in some areas of continental glaciation (Figure 4.3.1).

Figure 4.3.6 Bowyer Island, A Drumlin in Howe Sound, BC. Ice flow was from right to left.

A roche moutonée is another type of elongated erosional feature that has a steep and sometimes jagged down-ice end (Figure 4.3.7 left).  On a smaller scale still, glacial grooves (10s of cm to m wide) and glacial striae (mm to cm wide) are created by fragments of rock embedded in the ice at the base of a glacier (Figure 4.3.7 left and right).  Glacial striae are very common on rock surfaces eroded by both alpine and continental glaciers.

Figure 4.3.7 Left: A Roche Moutonée with Glacial Striae Near to Squamish, BC. Right: Glacial Striae at the Same Location Near to Squamish. Ice flow was from right to left in both cases.

Lakes are typical features of glacial environments.  A lake that is confined to a glacial cirque is known as a tarn (Figure 4.3.8).  Tarns are common in areas of alpine glaciation because the ice that forms a cirque typically carves out a depression that then fills with water.

Figure 4.3.8 Lower Thornton Lake, a Tarn, in the Northern Cascades National Park, Washington.

A lake that occupies a glacial valley, but is not confined to a cirque, is known as a finger lake.  In some cases a finger lake is confined by a dam formed by an end moraine, in which case it may be called a moraine lake (or moraine-dammed lake) (Figure 4.3.9).

Figure 4.3.9 Peyto Lake in the Canadian Rockies: Both a Finger Lake and a Moraine lake as it is Dammed by an End-Moraine (on the right).

In areas of continental glaciation the crust is depressed by the weight of glacial ice that is up to 4000 m thick.  Basins are formed along the edges of continental glaciers (except for those that cover entire continents like Antarctica and Greenland and terminate in the ocean), and these basins fill with glacial melt water.  Many such lakes, some of them huge, existed at various times along the southern edge of the Laurentide Ice Sheet.  One example is Glacial Lake Missoula, which formed within Idaho and Montana, just south of the BC border.  During the latter part of the last glaciation (30 to 15 ka) the ice holding back Lake Missoula retreated enough to allow some of the lake water to start flowing out and this then escalated into a massive and rapid outflow (over days to weeks) during which time much of the volume of the lake drained along the valley of the Columbia River to the Pacific Ocean.  It is estimated that this type of flooding happened at least 25 times over that period, and in many cases the rate of outflow was equivalent to the discharge of all of the Earth’s current rivers combined.  The record of these massive floods is preserved in the Channelled Scablands of Washington and Oregon (Figure 4.3.10).

Figure 4.3.10 Potholes Coulee Near to Wenatchee, Washington, One of Many Basins that Received Lake Missoula Floodwaters During the Late Pleistocene. Here the water flowed from right to left, over the cliff and into this basin.

Another type of glacial lake is a kettle lake. These are discussed in Section 4.4 in the context of glacial deposits. An example is shown on Figure 4.4.7.

Media Attributions

• Figure 4.3.1 Public domain NASA image,  Photo ID: 85506
• Figure 4.3.2 Photo by keefmon@hotmail.com, 2005, CC BY SA 2.5, via Wikimedia Commons
• Figure 4.3.3 Modified by Steven Earle, CC BY 4.0, from a public domain NASA image, Photo ID: 7195
• Figure 4.3.4 Modified by Steven Earle, CC BY 4.0, from a glacial landscape drawing by Luis Maria Benitez, 2005, Public Domain, via Wikimedia Commons.
• Figure 4.3.5 Modified by Steven Earle, CC BY 4.0, from a photo by Kurt Stegmüller, 2008, CC BY 3.0, via Wikimedia Commons
• Figure 4.3.6 Photo by Steven Earle, CC BY 4.0
• Figure 4.3.7 Photos by Steven Earle, CC BY 4.0
• Figure 4.3.8 Photo by Jeffrey Pang from Madison, NJ, USA, 2007, CC BY 2.0, via Wikimedia Commons
• Figure 4.3.9 Photo by chensiyuan, 2006, CC BY SA 4.0, via Wikimedia Commons
• Figure 4.3.10 Photo by Steven Earle, CC BY 4.0

 

Sediments transported and deposited during the Pleistocene glaciations are abundant throughout Canada and the northern US. They are important sources of construction materials and as reservoirs for groundwater, and, because they are almost all unconsolidated, they have significant implications for mass wasting.

Figure 4.4.1 illustrates some of the ways that sediments are transported and deposited. The Bering Glacier is the largest in North America, and although most of it is in Alaska, it flows from an ice field that extends into southwestern Yukon. The surface of the ice is partially, or in some cases completely covered with rocky debris that has fallen from surrounding steep rock faces. There are muddy rivers issuing from the glacier in several locations, depositing sediment on land, into Vitus Lake and directly into the ocean. There are dirty icebergs shedding their sediment into the lake. And, not visible in this view, there are sediments being moved along beneath the ice.

Figure 4.4.1 Part of the Bering Glacier in Southeast Alaska, the Largest Glacier in North America. It is about 14 km across in the centre of this view.

The formation and movement of sediments in glacial environments is shown diagrammatically on Figure 4.4.2. The main types of sediment in a glacial environment are as follows: Supraglacial (on top of the ice) and englacial (within the ice) sediments that slide off the melting front of a stationary glacier can form a ridge of unsorted sediments called an end moraine. The end moraine that represents the furthest advance of glacier is a terminal moraine.

Figure 4.4.2 A Depiction of the Various Types of Sediments Associated With Glaciation. The glacier is shown in cross section.

Sub-glacial till (the most abundant of which is lodgement till) is material that has been eroded from the underlying rock by the ice, and is moved by the ice, and emplaced on the bed by friction generated by the weight of overlying ice. It has a wide range of grain sizes, including a relatively high proportion of silt and clay. The larger clasts (pebbles to boulders in size) tend to get partly rounded by abrasion. When a glacier eventually melts the lodgement till is exposed as a sheet of well-compacted sediment ranging from several centimetres to many metres in thickness. Lodgement till is normally unbedded. An example is shown on Figure 4.4.3a.

Figure 4.4.3 Examples of Till. a. Lodgement till from the front of the Athabasca Glacier, Alberta; b. Ablation till at the Horstman Glacier, Blackcomb Mt., BC. Ablation till typically has less clay and is less well compacted than lodgement till.

Supra-glacial sediments are primarily derived from freeze-thaw eroded material that has fallen onto the ice from rocky slopes above. These sediments form lateral moraines (Figure 4.4.4) and, where two glaciers meet: medial moraines. (Medial moraines are visible on the Aletsch Glacier on Figure 4.3.3.) Most of this material is deposited on the ground when the ice melts, and is therefore called ablation till, a mixture of fine and coarse angular rock fragments, with much less sand, silt and clay than lodgement till. An example is shown on Figure 16.30b. When supraglacial sediments get incorporated into the body of the glacier they are known as englacial sediments (Figure 4.4.2).

Figure 4.4.4 A lateral moraine (dark grey material in the background) adjacent to the Athabasca Glacier (Author photo, CC BY 4.0)

Massive amounts of water flow on the surface, within and at the base of a glacier, even in cold areas and even when the glacier is advancing. Depending on its velocity, this water is able to move sediments of various sizes and most of that material is washed out of the lower end of the glacier and deposited as outwash sediments. These sediments accumulate in a wide range of environments in the proglacial region (the area in front of a glacier), most in fluvial (river) environments, but some in lakes and some in the ocean.

A large proglacial plain of sediment deposition is called a sandur, and within that area glacio-fluvial deposits can be tens of metres thick. The sandur shown on Figure 4.4.5 covers an area of over 1000 km² along the southern coast of Iceland near to Vatnajökull (the largest glacier in Iceland). Glacio-fluvial sediments are generally similar to sediments deposited in normal fluvial environments, and are dominated by silt, sand and gravel. The grains tend to be moderately well rounded, and the sediments have sedimentary structures (e.g., bedding, cross bedding, clast imbrication) that are similar to those formed by non-glacial streams (Figure 4.4.6).

Figure 4.4.5 Glaciofluvial sediments that make up the extensive sandur in front of Vatnajökull in Iceland (Author photo, CC BY 4.0)

Figure 4.4.6 Examples of Glacio-Fluvial Sediments: a. Glacio-fluvial sand of the Quadra Sand Fm. at Comox; BC, b. glacio-fluvial gravel and sand, Nanaimo, BC

In situations where a glacier is receding, a block of ice might become separated from the main ice sheet and then could get buried in glacio-fluvial sediments. When the ice block eventually melts a depression will form, and if this fills with water it is known as a kettle lake (Figure 4.4.7).

Figure 4.4.7 A Kettle Lake Amidst Vineyards in the Osoyoos Area of Southern BC

A subglacial stream will create its own channel within the ice, and sediments that are being transported and deposited by the stream will build up within that channel. When the ice recedes, that sediment will remain to form a long sinuous ridge known as an esker. Eskers are most common in areas of continental glaciation. They can be several metres high, tens of metres wide and tens of kilometres long (Figure 4.4.8).

Figure 4.4.8 Part of an Esker That Formed Beneath the Laurentide Ice Sheet in Northern Canada

Outwash streams can flow into proglacial lakes where glacio-lacustrine sediments are deposited. These are dominated by silt- and clay-sized particles and are typically laminated on the millimetre scale. In some cases varves develop: a series of beds that each has distinctive summer and winter layers. Relatively coarse sediments are deposited in the summer when melt discharge is high, and finer sediments are deposited in the winter, when discharge is very low. Ice bergs are common on pro-glacial lakes (see Figure 4.2.12), and most of them contain englacial sediments of various sizes. As these bergs melt, the released clasts sink to the bottom and get incorporated into the glacio-lacustrine layers as drop stones (Figure 4.4.9a).

The processes that occur in proglacial lakes can also take place where a glacier terminates in the ocean. These are called glacio-marine sediments (Figure 4.4.9b).

Figure 4.4.9 Examples of glacial sediments formed in quiet water, a: glacio-lacustrine sediment with a drop stone, Nanaimo, BC, and b: a laminated glacio-marine sediment, Englishman River, BC

Media Attributions

• Figure 4.4.1 NASA Earth Observatory‘s Image of the Day for Aug 3, 2004, public domain
• Figure 4.4.2 Steven Earle, CC BY 4.0
• Figure 4.4.3 Photos by Steven Earle, CC BY 4.0
• Figure 4.4.4 Photo by Steven Earle, CC BY 4.0
• Figure 4.4.5 Photo by Steven Earle, CC BY 4.0
• Figure 4.4.6 Photos by Steven Earle, CC BY 4.0
• Figure 4.4.7 Photo by Steven Earle, CC BY 4.0
• Figure 4.4.8 Photo from Agriculture and Agri-food Canada, https://sis.agr.gc.ca/cansis/images/nt/locsf/index.html, Open Government Licence – Canada, 
• Figure 4.4.9 Photos by Steven Earle, CC BY 4.0

Glaciers are receding as a result of the warming caused by anthropogenic climate change. A graphic example of this warming can be seen at the Athabasca Glacier, which is receding at a rate of around 5 m per year (Figure 4.5.1).

Figure 4.5.1 The Position of the Front of the Athabasca Glacier (Alberta) in 1992. The photo was taken in August 2006, at which time the ice front was about 300 m distant. The rock in the foreground has clearly evident glacial striations.

Most glaciers around the world are receding. Although some are holding steady and just a few are even advancing, the average ice mass balance of glaciers, large and small, is persistently negative. As shown on Figure 4.5.2, the total glacial ice loss (from all of the world’s major ice sheets and alpine glaciers) has been several hundred Gt/y (giga-tonnes per year) for most of the past three decades, and, as the climate has continued to warm, that rate has increased.

Figure 4.5.2 Annual Changes in the Mass of Glaciers Across the World, 1992 to 2016

Some of the significant environmental and social implications of ice loss on this scale are summarized in a 2019 article by Rasul and Molden. These, and others, can be summarized as follows:

• When glaciers recede, bare rock is exposed, and that has a lower albedo than the glacier did, so absorbs more solar energy and heats up more.
• Sea level rise is an obvious consequence of glacial melting and the amount of melting experienced in the past several years contributes approximately 2 mm to sea level rise each year.
• Enhanced melting of glaciers contributes more than normal amounts of fresh water to the oceans. As described in section 3.5, changes to ocean salinity—especially in the north Atlantic—can lead to changes in both surface and deep ocean currents which can have significant climate implications.
• Glacial recession is typically accompanied by melting of permafrost and that results in the release of carbon dioxide and methane, which contribute to more warming. Melting of permafrost also leads to the destabilization of slopes.
• Enhanced glacial melting can enlarge proglacial lakes and that will contribute to the risk of dangerous glacial outburst floods.
• After deglaciation slopes may no longer be supported by ice, and so the risk of slope failure is increased.
• In many regions the supply of drinking and irrigation water is dependent on the slow melting of glacial ice over the summer. As glaciers get smaller, this supply will be less reliable.
• Loss of glacial meltwater will have implications for fish and other aquatic life.
• Glacial regions are important draws for tourists and recreation. The associated economic opportunities will be threatened by glacier loss.

Glaciers and Earth Systems

Glaciers are a special part of the hydrological cycle, and they contribute significantly to Earth systems.

Glaciers, especially large continental ice sheets, contain a massive amount of frozen water, so they exert a strong control over sea level. During periods of extensive glaciation, average sea level can be over 100 m lower than it is now (or significantly more than that at some times in the past). On the other hand, when there is little or no glacial ice sea level would be about 70 m higher than it is now. These variations lead to changes in coastal processes and where those processes take place. For example, coastal regions tend to have low relief and are made of soft sediments following a drop in sea level and tend to have higher relief and are made up of hard rocks when sea level rises. Sea-level changes also change the Earth’s overall albedo, because of changes in the area of reflective ice and non-reflective sea water. Glaciers also have their own albedo effect because of their generally bright surfaces.

Glaciers represent storage of water that would otherwise flow quickly through the hydrological system, so they have significant implications for flow rates at different times of year.
Glaciers also play an important role in shaping mountainous areas into steep slopes, leading to slope failures and enhanced erosion—which reduces the carbon dioxide content of the atmosphere, and they produce a tremendous amount of sediment which has implications for biological processes on land and in the ocean.

Media Attributions

• Figure 4.5.1 Photo by Steven Earle, CC BY 4.0
• Figure 4.5.2 Steven Earle, CC BY 4.0, based on data in: Bamber, J., Westaway, R., Marzeion, B., & Wouters, B. (2018). The land ice contribution to sea level during the satellite era. Environmental Research Letters, 13(6), 063008. https://iopscience.iop.org/article/10.1088/1748-9326/aac2f0/meta

The topics covered in this chapter can be summarized as follows:

Mass wasting, which is synonymous with “slope failure”, is the failure and down-slope movement of rock or unconsolidated materials in response to gravity. The term “landslide” is almost synonymous with mass wasting, but not quite because some people reserve “landslide” for relatively rapid slope failures, while others do not. Because of that ambiguity, we will avoid the use of “landslide” in this textbook. Instead, wherever possible, mass wasting events will be referred to by specific names that describe the type of material that failed and the type of motion that took place.

Mass wasting happens because the Earth’s surface is made up of sloped surfaces, and that has happened because tectonic processes have produced uplift. Erosion, driven by gravity, is the inevitable response to that uplift, and mass wasting is a type of erosion. Slope stability is ultimately determined by two factors: the angle of the slope, and the strength of the materials on it.

Figure 5.1.1 shows a block of rock situated on a rock slope. It is being pulled towards the earth’s centre (straight down) by gravity. We can split the vertical gravitational force into two components relative to the slope, one pushing the block down the slope (the shear force), and the other pushing it into the slope (the normal force). The shear force—which wants to push the block down the slope—has to overcome the strength of the connection between the block and the slope, which may be quite weak if the block has split away from the main body of rock, or may be very strong if the block is still a part of the rock. This is the shear strength, and in Figure 5.1.1a it is greater than the shear force, so the block should not move. In Figure 5.1.1b the slope is steeper and the shear force is approximately equal to the shear strength. The block may or may not move under these circumstances. In Figure 5.1.1c the slope is steeper still, so the shear force is considerably greater than the shear strength, and the block will very likely move.

Figure 5.1.1 Differences in the Shear and Normal Components of the Gravitational Force on Slopes with Differing Steepness. The gravitational force is the same in all three cases. In “a” the shear force is substantially less than the shear strength, so the block should be stable. In “b” the shear force and shear strength are about equal, so the block may or may not move. In “c” the shear force is substantially greater than the shear strength, so the block is very likely to move.

As already noted, slopes are created by uplift. In areas with relatively recent uplift (most regions that have angular-looking mountains) slopes tend to be quite steep, and this is especially the case where glaciation has taken place because glaciers in mountainous terrain create steep-sided valleys. In areas without recent uplift (such as most of eastern North America), slopes are less steep because hundreds of millions of years of erosion (including mass wasting) has made them that way. However, as we’ll see, mass wasting can happen even on relatively gentle slopes.

The strength of the materials on slopes can vary widely. Solid rocks tend to be strong, but there is a very wide range of rock strength. If we consider just the strength of the rocks, and ignore issues like fracturing and layering, then most crystalline rocks—like granite, basalt or gneiss—are very strong, while some metamorphic rocks—like schist—are moderately strong. Sedimentary rocks have variable strength. Limestone is strong, some sandstone and conglomerate are moderately strong, while some types of sandstone and all mudstones are quite weak.

Fractures, metamorphic foliation or bedding can significantly reduce the strength of a body of rock, and, in the context of mass wasting, this is most critical if the planes of weakness are parallel to the slope (see Figure 5.0.1) and least critical if they are perpendicular to the slope. This is illustrated on Figure 5.1.2. At locations a and b the bedding is nearly perpendicular to the slope and the situation is relatively stable. At location d the bedding is nearly parallel to the slope and the situation is quite unstable. At location c the bedding is nearly horizontal, and the stability is intermediate between the other two extremes.

Figure 5.1.2 Cross-Section Through an Area with Topographic Relief That is Underlain by Folded Sedimentary Rocks. Relative stability of slopes as a function of the orientation of weaknesses (in this case bedding planes) relative to the slope orientations.

Internal variations in the compositions of rocks can significantly affect their strength. Schist, for example, may have layers that are rich in sheet silicates (mica or chlorite) and these will tend to be weaker than other layers. Some minerals tend to be more susceptible to weathering than others, and the weathered products are commonly quite weak. The side of Johnson Peak that failed in 1965 (Hope Slide) is made up of chlorite schist (metamorphosed sea-floor basalt) that has feldspar-bearing sills within it (they are evident within the inset area of Figure 5.0.1). The foliation and the sills are parallel to the steep slope. The schist was relatively weak to begin with, and the feldspar in the sills, which had been altered to clay, made it even weaker.

Unconsolidated sediments are generally weaker than sedimentary rocks because they are not cemented and, in most cases, have not been significantly compressed by overlying materials. Sand and silt tend to be particularly weak, clay is generally a little stronger (unless it is wet, see below), and sand mixed with clay can be stronger still. The deposits that make up the cliffs at Point Grey in Vancouver include sand, silt and clay overlain by sand. As shown on Figure 5.1.3 (left) the finer deposits are relatively strong (they maintain a steep slope), while the overlying sand is relatively weak, has maintained a shallower slope and has recently failed. Glacial till—typically a mixture of clay, silt, sand, gravel and larger clasts—that forms beneath tens to thousands of metres of glacial ice and is well compressed, can be as strong as some sedimentary rock (Figure 5.1.3 – right).

Figure 5.1.3 Left: Glacial Outwash Deposits at Pt. Grey, in Vancouver. The dark lower layer is comprised of sand, silt and clay. The light coloured upper layer is well-sorted sand. Right: Glacial Till at Quadra Island, BC. The till is strong enough to have formed a near-vertical slope.

Apart from the type of material on a slope, the amount of water it contains is the most important factor controlling its strength. This is especially true for unconsolidated materials, like those shown on Figure 5.1.3, but it also applies to bodies of rock. Granular sediments, like the sand at Pt. Grey, have lots of spaces between the grains. Those spaces may be completely dry (filled only with air), or moist, (often meaning that some spaces are water filled, some grains have a film of water around them and small amounts of water are present where grains are touching each other), or completely saturated (Figure 5.1.4). Unconsolidated sediments tend to be strongest when they are moist because the small amounts of water at the grain boundaries hold the grains together with surface tension. Dry sediments adhere together only by the friction between grains, and if they are well sorted or well rounded, or both, that adhesion is weak. Saturated sediments tend to be the weakest of all because the large amount of water actually pushes the grains apart reducing the mount friction between grains. This is especially true if the water is under pressure.

Figure 5.1.4 Depiction of Dry, Moist and Saturated Sand. Air-filled pores are white, water is blue.

Exercise 5.1 Sand and Water

If you’ve ever been to the beach, you’ll already know that sand behaves differently when it’s dry than it does when it’s wet, but it’s worth taking a systematic look at the differences in its behaviour. Find about half a cup of clean dry sand (or get some wet sand and dry it out), and then pour it from your hand onto a piece of paper. You should be able to make a cone-shaped pile that has a slope of around 30˚. If you pour more sand on the pile it will get bigger, but the slope should remain the same. (This is known as the angle of repose, which is the steepest angle that a pile of dry sediment can maintain without failing.)

(Photo by Steven Earle, CC BY 4.0)

Now return the sand to a container and add some water so that it is moist. An easy way to do this is to make it completely wet and then let the water drain away for a minute. You should be able to form this moist sand into a steep pile (with slopes of around 80˚). Finally put the same sand into a container and fill it up with water so the sand is just covered. Swirl it around so that the sand remains in suspension, and then quickly tip it out onto a flat surface (best to do this outside). It should spread out over a wide area, forming a pile with a slope of only a few degrees.

Water will also reduce the strength of solid rock, especially if it has fractures or bedding planes, or clay-bearing zones. This effect is most significant when the water is under pressure, and that’s why you’ll often see holes drilled into rocks on road cuts. One of the hypotheses advanced to explain the 1965 Hope Slide is that the very cold conditions that winter caused small springs in the lower part of the slope to freeze over, preventing water from flowing out. It is possible that water pressure gradually built up within the slope, weakening the rock mass to the extent that the shear strength was no longer greater than the shear force.

Water also has a particular effect on clay-bearing materials. All clay minerals will absorb a little bit of water, and this reduces their strength. The smectite clays (such as the bentonite used in cat litter) can absorb a lot of water, and that water pushes the sheets apart on a molecular level and makes the mineral swell. Smectite that has expanded in this way has almost no strength; it is extremely slippery.

And finally, water can significantly increase the mass of the material on a slope. This will increase the gravitational force pushing it down, but it will also increase the normal force pushing mass against the slope and that will increase the friction, so while adding water may make the material on the slope weaker and more prone to fail, the additional weight doesn’t necessarily contribute to failure.

Mass Wasting Triggers

In the foregoing discussion we talked about the shear force and the shear strength of materials on slopes, and about factors that can reduce the shear strength. Shear force is primarily related to slope angle, and this does not change quickly. But shear strength can change quickly for a variety of reasons, and events that lead to a rapid reduction in shear strength are considered to be triggers for mass wasting.

An increase in water content is the most common mass wasting trigger because of the reduction in strength. This can result from rapid melting of snow or ice, by heavy rain, or by some type of event that changes the pattern of water flow on the surface. Rapid melting can be caused by a dramatic increase in temperature (e.g., in spring or early summer) or by a volcanic eruption. Changes in water flow patterns can be caused by earthquakes, or previous slope failures that dam up streams or human structures that interfere with runoff (e.g., building, roads or parking lots). An example of this is the deadly 2005 debris flow in North Vancouver (Figure 5.1.5). The 2005 failure took place in an area that had failed previously, and in a report written in 1980 it was recommended that steps be taken by the municipal authorities and the residents to address drainage issues. Little was done to improve the situation.

Figure 5.1.5 Site of the Debris Flow in the Riverside Drive Area of North Vancouver in January, 2005. This debris flow happened during a rainy period, but was likely triggered by excess runoff related to the roads at the top of this slope and by landscape features in the area surrounding the house visible here.

In some cases, a decrease in water content can lead to failure. This is most common with clean sand deposits (e.g., the upper layer in Figure 5.1.3, left), which lose strength when there is no more water around the grains.

Freezing and thawing can also trigger some forms of mass wasting, more specifically, the freezing can expand the crack between two parts of rock, and then thawing can release a block of rock that was attached to a slope by a film of ice, as illustrated on Figure 5.1.6.

Figure 5.1.6 Illustration of the Process of Ice Wedging (left) and then Release of a Fragment Because of Melting

Shaking is another process that can weaken a body of rock or sediment. The most obvious source of shaking is an earthquake, but shaking from highway traffic, construction or mining will also do the job. Several deadly mass wasting events (including snow avalanches) were trigged by the M7.8 earthquake in Nepal in April 2015.

Saturation with water and then seismic shaking led to the occurrence of thousands of slope failures in the Sapporo area of Hokkaido, Japan in September 2018, as shown on Figure 5.1.7. The area was drenched with rain from tropical storm Jebi on September 4th, and then shaken by a M 6.6 earthquake on September 6th. That combination appears to have triggered thousands of debris flows of water-saturated volcanic materials on steep slopes. There were 41 deaths related to these slope failures.

Figure 5.1.7 Left: Slope Failures in the Sapporo Area of Japan Following a Typhoon (Sept. 4th, 2018); and right: Earthquake (Sept. 6th, 2018). (Before and after Landsat 8 images: left: July 2017, right: September 2018).

Media Attributions

• Figure 5.1.1 Steven Earle, CC BY 4.0
• Figure 5.1.2 Steven Earle, CC BY 4.0
• Figure 5.1.3 Photos by Steven Earle, CC BY 4.0
• Figure 5.1.4 Steven Earle, CC BY 4.0
• Figure 5.1.5 Photo from The Province newspaper, used with permission.
• Figure 5.1.6 Steven Earle, CC BY 4.0
• Figure 5.1.7 Public domain image from Dauphin, L. (2018). Landslides in Hikkaido. NASA Earth Observatory, https://earthobservatory.nasa.gov/images/92832/landslides-in-hokkaido

• The type of material that failed (typically either bedrock or unconsolidated sediment),
• The mechanism of the failure (how the material moved), and
• The rate at which it moved.

Rock Fall

Rock fragments can break off relatively easily from steep bedrock slopes, most commonly due to frost-wedging in areas where there are many freeze-thaw cycles per year.  If you’ve ever hiked along a steep mountain trail on a cool morning you might have heard the occasional fall of rock fragments onto a talus slope as the sun melts the ice, releasing rock fragments that had been wedged out the night before.  This process is illustrated in Figure 5.1.6 above.

A typical talus slope, near to Keremeos in southern BC, is shown on Figure 5.2.1.  In December 2014 a large block of rock split away from a cliff in this same area.  It broke into smaller pieces, which fell and tumbled down the slope and crashed into the road, smashing the concrete barriers and gouging out large parts of the pavement.

Figure 5.2.1  Left: A Talus Slope Near to Keremeos, BC, Formed by Rock Fall from the Cliffs Above.  Right: The Results of a Rock Fall onto a Highway West of Keremeos in December 2014.

Rock Slide

A rock slide is the sliding motion of rock along a sloping surface. In most cases the movement is parallel to a fracture, bedding plane or metamorphic foliation plane, and it can range from very slow to moderately fast.  The word sackung describes the very slow motion of a block of rock (mm/y to cm/y) on a steep slope.  A good example is the Downie Slide north of Revelstoke BC, which is illustrated on Figure 5.2.2.   In this case a massive body of rock is very slowly sliding down a steep slope along a plane of weakness that is parallel to the slope. The Downie Slide, which was recognized prior to the construction of the Revelstoke Dam, was moving very slowly at the time (a few cm/year). Geological engineers were concerned that the presence of water in the reservoir (visible on Figure 5.2.2) could further weaken the plane of failure, leading to an acceleration of the motion.  The result could have been a catastrophic failure into the reservoir that sent a wall of water over the dam and into the community of Revelstoke.  During the construction of the dam, they tunneled into the rock at the base of the slide and drilled hundreds of drainage holes upward into the plane of failure.  This allowed water to drain out so that the pressure was reduced, and that stabilized the slide block.  BC Hydro monitors this site continuously; the slide block is currently moving slower than it was prior to the construction of the dam.

Figure 5.2.2 The Downie Slide, A Sackung, on the Shore of the Lake Revelstoke Reservoir (Above the Revelstoke Dam). The head scarp is visible at the top and a side-scarp along the left-had side.

In the summer of 2008, a large block of rock slid rapidly from a steep slope above Highway 99 near to Porteau Cove (40 km north of Vancouver).  The block slammed into the highway and the adjacent railway and broke into many pieces.  The highway was closed for several days, and the slope was subsequently stabilized with rock bolts and drainage holes.  As shown on Figure 5.2.3, the bedrock at this location is fractured parallel to the slope, and this almost certainly contributed to the failure. It is not actually known what trigged this event as the weather was dry and warm during the preceding weeks and there was no significant earthquake in the region.

Figure 5.2.3  Site of the 2008 Rock Slide at Porteau Cove. Notice the prominent fracture set parallel to the surface of the slope. The slope has been stabilized with rock bolts (visible at the top) and holes have been drilled into the rock to improve drainage (one is visible in the lower right). Risk to passing vehicles from rock fall has been reduced by hanging mesh curtains (background).

Rock Avalanche

If a rock slide starts moving quickly (m/s) the rock is likely to break into many small pieces, and at that point it can become a rock avalanche, in which the large and small fragments of rock move in a fluid manner supported by cushion of air within and beneath the moving mass.  The 1965 Hope Slide (Figure 5.0.1) was a rock avalanche, as was the famous 1903 Frank Slide in southwestern Alberta.  The 2010 slide at Mt Meager (west of Lillooet), also a rock avalanche (Figure 5.2.4).

Figure 5.2.4 The 2010 Mt. Meager Rock Avalanche, Showing Where the Slide Originated (top centre of image). It then raced down a steep narrow valley, into the valley of Meager Creek, and then out into the wider valley of the Lillooet River at the bottom of the image.

Creep

The very slow—mm/y to cm/y—movement of soil or other unconsolidated material on a slope is known as creep. Creep, which normally only affects the upper several centimetres of loose material, is typically a type of very slow flow, but in some cases sliding may take place.  Creep can be facilitated by freezing and thawing, because, as shown on Figure 5.2.5, particles get lifted perpendicular to the surface by the growth of ice crystals within the soil, and are then let down vertically by gravity when the ice melts.  The same effect can be produced by frequent wetting and drying of the soil.

Creep is most noticeable on moderate to steep slopes where trees, fence posts or gravestones are consistently leaning in a downhill direction (Figure 5.2.6).  In the case of trees, they try to correct their lean by growing upright, and this leads to a curved lower trunk known as a “pistol butt” (or “j-shaped tree trunk”).  Creep can take place on nearly flat surfaces.

Figure 5.2.5 A Hillside with Pistol-Butt Trees as Evidence of Persistent Creep

Slump

Slump is a type of slide (movement as a mass), that takes place within thick unconsolidated deposits (typically greater than 10 m).  Slumps involve movement along a curved surface, with downward motion near to the top and outward motion towards the bottom (Figure 5.2.7).  They are typically caused by an excess of water within the materials on a steep slope.

Figure 5.2.6 A Depiction of the Motion of Unconsolidated Sediments in an Area of Slumping

An example of a slump in the Lethbridge area, Alberta, is shown on Figure 5.2.8.  This feature has likely been active for many decades, and moves a little more whenever there are heavy spring rains and significant snow-melt runoff.  The toe of the slump is failing because it has been eroded by the small stream at the bottom.

Figure 5.2.7  A Slump Along the Banks of a Small Coulee Near to Lethbridge, Alberta.  The main head-scarp is clearly visible at the top, and a second smaller one is visible about one-quarter of the way down. The toe of the slump is being eroded by the seasonal stream that created the coulee.

Mud Flows and Debris Flows

Figure 5.2.8  A Slump (left) and an Associated Mudflow (centre) (at the same location as Figure 5.2.8), Near to Lethbridge, Alberta.

Figure 5.2.9 Effects of a Debris Flow Within a Steep Stream Channel Near to Buttle Lake, BC., November 2006

Exercise 5.2  Classifying Slope Failures

The four photos below show some of the different types of slope failures described above.  Try to identify the different types, and in each case provide some criteria for why you made that choice.

Figure 5.2.10 Slope Failures

Exercise answers are provided Appendix 2.

As already noted, to understand a slope failure we need to be able to determine what type of material moved, what type (or types) of motion were involved, and how quickly it moved.  The type of motion is the most important of these, and so Figure 5.2.11 is provided here to help you clearly understand how things moved in different types of slope failure.

Figure 5.2.11 Types of Slope-Failure Motion. A fall is typically a rock fall. A flow can be creep, mud-flow, debris flow, or rock avalanche. A translational slide is typically a rock slide.  A rotational slide is typically a slump.

Exercise 5.3  Slope Failure Field Trip

It’s time to get outside and look for evidence of slope failure close to home.  No matter where you live, even on the flattest plain, there is likely to be some kind of natural slope failure nearby. Venture to some sloping terrain, such as a valley eroded by a stream, a road embankment, a gently sloping cemetery, and look for evidence that there has been some down-slope movement.

When you find something, try to answer the following question:

• What has failed (is it loose sediments or solid rock)?
• How has it failed (slide or flow)?
• How quickly did the material move?
• How long ago did it happen (or is it still happening)?
• What is the risk for future failure at this location?

Box 5.1 The November 2020 Mass Wasting Events at Elliot Creek, BC

On December 10th, 2020 helicopter pilot Bastian Fleury was flying over Elliot Creek in southwestern BC when he noticed that the creek valley had been devastated by a massive debris flow. Fleury was the first person to be aware of this event, which had actually taken place on November 27th (based on some small (M4) seismic events from that location).  There is no evidence that anyone was in the area at the time and there was little damage to human infrastructure. Later flights by Fleury and others confirmed that the event had started with a rock slide at the upper end of Elliot Lake, and continued for 14 km down Elliot Creek and then another 10 km along the Southgate River to the ocean at Bute Inlet. This mass wasting event is expected to have a significant impact on the salmon (and the salmon fishery) in the waters of Bute Inlet.

A summary of the various different phases of this mass wasting event is provided on the graphic below.

Figure 5.2.12 Elliot Creek Debris Flow Interpretation

Media Attributions

• Figure 5.2.1 Photos by Steven Earle, CC BY 4.0
• Figure 5.2.2 Image from Google Earth.
• Figure 5.2.3 Photo by Steven Earle, CC BY 4.0
• Figure 5.2.4 2010 Mount Meager Landslide by Tim Gage, 2014, CC BY SA 2.0, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:2010_Mount_Meager_landslide.jpg
• Figure 5.2.5 Photo by Steven Earle, CC BY 4.0
• Figure 5.2.6 Steven Earle, CC BY 4.0
• Figure 5.2.7 Photo by Steven Earle, CC BY 4.0
• Figure 5.2.8 Photo by Steven Earle, CC BY 4.0
• Figure 5.2.9 Photo by Steven Earle, CC BY 4.0
• Figure 5.2.10 Photo by Steven Earle, CC BY 4.0
• Figure 5.2.11 Steven Earle, CC BY 4.0
• Figure 5.2.12 Steven Earle, CC BY 4.0. All of the photographs are public domain and have been provided by 49North Helicopters, https://49northhelicopters.com; satellite imagery is from Google Earth.

As already noted above, we cannot prevent mass wasting, but in many situations there are actions that we can take to delay it from happening. Where we cannot delay it, there are things we can do to mitigate its damaging effects on people and infrastructure. Where we can neither delay nor mitigate mass wasting, we should have the sense to stay out of the way.

Preventing and Delaying Mass Wasting

It is comforting to think that we can prevent mass wasting by mechanical means, such as the bolts in the road cut at Porteau Cove (Figure 5.2.3), or by draining water out of a slope, as was done at the Downie Slide (Figure 5.2.2), or by building physical barriers. What we have to remember is that the works of man are weak and pathetic things compared to the works of nature. The rock bolts in the road cut at Porteau Cove will start to corrode within a few years, and within a few decades many of them will lose strength. Unless they are replaced—and nobody can guarantee that that will happen—they will no longer be supporting that slope. The drainage holes at the Downie Slide will eventually get plugged up with sediment and chemical precipitates, and, unless they are periodically unplugged (which would be difficult) their effectiveness will be reduced. Eventually, unless new holes are drilled, the drainage will be so compromised that the slide will start to move again.

The important point is that our efforts to “prevent” mass wasting are only as good as our resolve to maintain those preventative measures. We may have created systems to ensure that others will carry out that maintenance in future decades or even future centuries, but there is nothing to ensure that it will be done.

Delaying mass wasting is a worthy endeavour, of course, because during the time that the measures are still effective, they can save lives and damage to property. The other side of the coin is that we must be careful to avoid doing things that could make mass wasting more likely. The most common anthropogenic cause of mass wasting is road construction, and this applies both to remote gravel roads built for resource extraction and to large highways. Road construction is a potential problem for two reasons. First, to create a flat road surface on a slope inevitably involves creating a cut bank that is steeper than the original slope, and might also involve creating a filled bank that is both steeper and weaker than the original slope (Figure 5.3.1). Second, roadways typically cut across natural drainage features, and unless great care is taken to reroute the runoff water, and to prevent it from pooling, oversaturation of materials can result. A specific example of the contribution of construction-related changes to drainage that led to slope instability is the debris flow in North Vancouver that is illustrated on Figure 5.1.5 (above).

Figure 5.3.1 An Example of a Road Constructed by Cutting into a Steep Slope, and the Use of the Cut Material as Fill

Apart from water issues, builders of roads and other infrastructure on bedrock slopes have to be acutely aware of the geology, and especially of any weaknesses in the rock related to bedding, fracturing or foliation. If possible, situations like that at Porteau Cove (Figure 5.2.3) should be avoided—by building somewhere else—rather than trying to stitch the slope back together with rock bolts.

It is commonly believed that construction of buildings on the tops of steep slopes can contribute to the instability of the slope. This may be true, but probably not because of the weight of the building. As you’ll see by completing Exercise 15.3, a typical house isn’t likely to be heavier than the hole in the ground that was made to build it. A more likely contributor to instability of the slope around a building is the effect that it has on drainage.

Exercise 5.3 How Much Does a House Weigh?

It is commonly believed that building a house (or some other building) at the top of a slope will add lots of extra weight to the slope and could contribute to slope failure. But what does a house actually weigh? A standard 1600 ft² wood-frame house with a basement and a concrete foundation weighs about 145 T (metric tonnes). But most houses are excavated into the ground, and that involves digging a hole and taking some material away, so we need to subtract what that excavated material weighs. Assuming that our 1600 ft² house required an excavation that was 15 m by 11 m by 1 m deep. That’s 165 m³ of “dirt”, which typically has a density of about 1.6 T per m³.

(Steven Earle, CC BY 4.0)

Calculate the weight of the soil that was removed and compare that with the weight of the house and its foundation.

If you’re thinking that a building a bigger building is going to add more weight, consider that bigger buildings need bigger and deeper excavations, and in many cases the excavations will be into solid rock, which is heavier than regular soil.

Exercise answers are provided Appendix 2.

Monitoring Mass Wasting

In some areas it is necessary to establish warning systems so that we know if conditions have changed at a known slide area, or if a rapid failure, such as a debris flow, is actually on its way. The Downie Slide is monitored 24-7 with optical devices. A simple mechanical device for monitoring the nearby Checkerboard Slide (which is also above Lake Revelstoke) is shown on Figure 5.3.2. Both of these are very slow-moving rock slides, but it’s very important to be able to detect changes in their rates of motion because at both of these locations a rapid failure would result in large bodies of rock plunging into the reservoir, sending a wall of water over the Revelstoke Dam, potentially destroying the nearby town of Revelstoke.

Figure 5.3.2 Part of a motion-monitoring device at the Checkerboard Slide near to Revelstoke, BC. The other end of the cable is attached to a block of rock that is unstable. Any incremental motion of that block will move the cable and this will be detectable on this device. (Author photo, CC BY 4.0)

Figure 5.3.3 Mount Rainier, from Auburn, Washington

Mt. Rainier in Washington State has the potential to produce massive mud flows or debris flows (technically lahars) with or without a volcanic eruption, and over 100,000 people in the Tacoma, Puyallup and Sumner areas are in harm’s way because they currently reside on deposits from past lahars (Figure 5.3.3). In 1998 a network of acoustic monitors was established around Mt. Rainier. The monitors are embedded in the ground adjacent to expected lahar paths. They provide warnings to emergency officials, and when a lahar is detected the residents of the area will have anywhere from 40 minutes to 3 hours to get to safe ground. There’s more on the risk of lahars at Mt. Rainier in Chapter 7.

It is critical to monitor the weather conditions and stream flows in order to assess changes to the risk of slope failure. Changing weather conditions can be failure triggers, especially in the following ways:

• extreme rain events saturate and weaken surficial materials (and even bedrock),
• extreme rain events lead to stream flooding and so to eroded banks, contributing to failures,
• sudden warming events can trigger rapid snow melt, increasing the amount of water in surface materials and streams, and
• persistent drought will dry out sandy materials making them more prone to failure.

Mitigating the Impacts of Mass Wasting

In situations where we can’t predict, prevent, or delay mass wasting hazards, there are some effective measures that can be taken to minimize the associated risk. For example, many highways in temperate mountainous regions have avalanche shelters like that shown on Figure 5.3.4. In some parts of the world similar features have been built to shelter infrastructure from other types of mass wasting.

Figure 5.3.4 A Snow Avalanche Shelter on the Coquihalla Highway. The expected path of the avalanche is the steep un-treed slope above.

Debris flows are inevitable, unpreventable and unpredictable in many parts of British Columbia (and elsewhere of course), but nowhere more so than along the Sea-to-Sky Highway between Vancouver and Squamish. The results have been deadly and expensive many times in the past. It’s too late now to close this region to the public, so provincial authorities have taken steps to protect residents and traffic on the highway and railway. Debris-flow defensive structures have been constructed on several drainage basins, as shown on Figure 5.3.5. One strategy is to allow the debris to flow quickly through to the ocean along a smooth channel. Another is to capture the debris within a constructed basin that allows water to continue through.

Figure 5.3.5 Two Strategies for Mitigating Debris Flows on the Sea-to-Sky Highway. Left: a concrete–lined channel on Alberta Creek allows debris to flow quickly through to the ocean. Right: a debris-flow catchment basin on Charles Creek. In 2010 a debris flow filled the basin to the level of the dotted white line.

Finally, in situations where we can’t do anything to delay, predict, contain or mitigate slope failures, we simply have to have the common sense to stay away. There is a famous example of this in British Columbia at a site known as Garibaldi, 25 km south of Whistler. In the early 1980s the village of Garibaldi had a population of about 100, with construction underway on some new homes, and plans for many more. In the months that followed the deadly 1980 eruption of Mt. St. Helens the Ministry of Transportation commissioned a geological study that revealed that a steep cliff known as The Barrier (Figure 5.3.6) had collapsed in 1855, leading to a large rock avalanche, and that it was likely to collapse again unpredictably, putting Garibaldi at extreme risk. In an ensuing court case, it was ruled that Garibaldi was not a safe place for people to live. Those that already had homes were compensated, and everyone was ordered to leave.

Figure 5.3.6 The Barrier, South of Whistler, BC. In 1855 a massive block of rock broke away from this cliff and transformed into a rock avalanche which flowed 6 km down the valley and across an area that was under residential development in the 1980s.

Media Attributions

• Figure 5.3.1 Steven Earle, CC BY 4.0
• Figure 5.3.2 Photo by Steven Earle, CC BY 4.0
• Figure 5.3.3 Mount Ranier from Centennial Viewpoint by Ron Clausen, 2017, CC BY SA 4.0, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Mount_Rainier_from_Centennial_Viewpoint_Park_Auburn_Washington.jpg
• Figure 5.3.4 Photo by Steven Earle, CC BY 4.0
• Figure 5.3.5 Photos by Steven Earle, CC BY 4.0
• Figure 5.3.6 Photo by Steven Earle, CC BY 4.0

• Formation of mountains due to plate motions, especially continental collisions, but also continental rifting.
• Formation of mountains due to volcanism, such as at subduction boundaries and above mantle plumes.
• Steepening of slopes due to erosion by rivers and glaciers.
• Accumulation of thick deposits of unconsolidated sediments through erosion and transportation by rivers, glaciers and wind.
• Strengthening of geological materials by vegetation, and loss of that strengthening due to wildfires and other events.
• Weakening of geological materials through weathering, especially where there is alteration of existing minerals to weaker minerals, by processes such as oxidation and hydrolysis.
• Weakening of geological materials through circulation of groundwater, especially in situations where that water becomes heated because of volcanism (see Figure 5.4.1).
• Breaking of geological materials and enhanced chemical weathering due to biological processes.
  
  

  Figure 5.4.1 Hydrothermal (Hot-Water) Alteration of Rock in the Krafla Area, Iceland. The brown colours are a result of oxidation of iron-bearing silicate minerals in the rock to soft iron-oxides. The grey colours are a result of hydrolysis of silicate minerals to weak clay minerals. The area is eroding quickly because the rock is now softer and weaker than it was originally.

On the other hand, mass wasting contributes to Earth systems processes in many ways, including the following:

• Production of new rock surfaces for weathering and new slopes for erosion.
• Release of greenhouse gases stored in surface materials (e.g., soil) because of slope failure.
• Changes to drainage patterns, including the formation of lakes and diversion of streams.
• Changes to sediment loads in streams, lakes and the ocean (see Box 5.1 for an example) that can have implications (both negative and positive) for ecosystems.
• Contribution to the transfer of materials from the continental crust into the oceans.

Media Attribution

• Figure 5.4.1 Photo by Steven Earle, CC BY 4.0

The topics covered in this chapter can be summarized as follows:

An earthquake is the shaking caused by the rupture (breaking) and subsequent displacement of rocks (one body of rock moving with respect to another) beneath the Earth’s surface. Most earthquakes are a result of the stresses placed on rocks in areas where adjacent tectonic plates are moving in different directions.

A body of rock that is under stress becomes deformed, and that deformation is elastic, meaning that the stressed rock can spring back into its original position. When the rock can no longer withstand the deformation, it breaks and the two sides slide past each other. Most earthquakes take place near to plate boundaries, but not necessarily right on a boundary, and not necessarily even on an existing or known fault.

The engineering principle of elastic deformation is illustrated on Figure 6.1.1. The stress applied to a rock—typically because of ongoing plate movement—results in strain or deformation of the rock (Figure 6.1.1b). Because most rock is strong (unlike loose sand for example), it can withstand a significant amount of deformation without breaking. But every rock has a deformation limit and will rupture (break) once pushed beyond that limit. At that point the rock breaks, there is displacement along the rupture surface, and the two bodies of rock that had been elastically deformed, rebound back to their original shape (Figure 6.1.1c). The magnitude of the earthquake depends on the extent of the area that breaks (the area of the rupture surface) and the average amount of displacement (sliding).

Figure 6.1.1 Depiction of the Concept of Elastic Deformation, Rupture and Elastic Rebound

Figure 6.1.2 A 3-Dimensional View of a Rupture Surface (Dark Pink), on a Steeply-Dipping Fault Plane (Light Pink). The diagram represents a part of the crust that may be up to tens or hundreds of kilometres long. The rupture surface is the area of the fault plane over which displacement occurred. The lengths of the arrows within the rupture surface represent relative amounts of displacement.

Figure 6.1.3 Propagation of Failure on a Rupture Surface. In this case, the failure starts at the dark blue heavy arrow and propagates outward, reaching the left side first (green arrows) and the right side last (yellow arrows).

The concept of a rupture surface—which is critical to understanding earthquakes—is illustrated on Figure 6.1.2. An earthquake does not happen at a point, it happens over an area within a plane, although not necessarily a flat plane. Within the area of the rupture surface the amount of displacement is variable (Figure 6.1.2), and, by definition, it decreases to zero at the edges of the rupture surface because the rock beyond that point is not displaced at all. The extent of a rupture surface and the amount of displacement will depend on a number of factors, including the strength of the rock, and the degree to which it was stressed beforehand.

Earthquake rupture doesn’t happen all at once; it will start at a single point and spread rapidly from there (Figure 6.1.3). Depending on the extent of the rupture surface, the propagation of failures out from the point of initiation is typically completed within seconds to several tens of seconds. The initiation point isn’t necessarily in the centre of the rupture surface; it may be close to one end, or near to the top or the bottom.

Figure 6.1.4 shows the distribution of immediate aftershocks associated with the 1989 Loma Prieta earthquake. Panel a-b is a section along the San Andreas Fault, and this view is equivalent to what is shown in Figures 6.1.2 and 6.1.3. The area of red dots is the rupture surface; each red dot represents a specific aftershock that was recorded on a seismometer. The red star represents the initial or main shock. When that initial shock happened the rock at that location broke and was displaced. That released the stress on that particular part of the fault, but it resulted in an increase of the stress on other nearby parts of the fault and contributed to a cascade of smaller ruptures (immediate aftershocks), in this case over an area about 50 km long and 15 km wide (or “deep” as shown on Figure 6.1.4).

Figure 6.1.4 Distribution of the Aftershocks of the 1989 M 6.9 Loma Prieta Earthquake. (a: plan view, b: section along the fault, c: section across the fault.)

So, what is an aftershock then? An aftershock is an earthquake just like any other, but it is one that can be shown to have been triggered by stress transfer from a preceding earthquake. Within a few tens of seconds of the main Loma Prieta earthquake there were hundreds of smaller aftershocks; their distribution defines the area of the rupture surface.

Aftershocks can be of any magnitude. Most are smaller than the earthquake that triggered them, but they can be bigger. The aftershocks shown on Figure 6.1.4 all happened within seconds or minutes of the main shock, but aftershocks can be delayed for hours, weeks, years or even decades. As already noted, aftershocks are related to stress transfer. For example, the main shock of the Loma Prieta earthquake triggered aftershocks in the immediate area, which triggered more in the surrounding area, eventually extending for about 25 km along the fault in both directions and for 15 km towards the surface. But the earthquake as a whole also changed the stress on adjacent parts of the San Andreas Fault. This effect, which has been modeled for numerous earthquakes and active faults around the world, is depicted on Figure 6.1.5. Stress was reduced in the area of the rupture (blue) but was likely to have increased at either end of the rupture surface (red).

Figure 6.1.5 Depiction of Stress Changes Related to an Earthquake. Stress is decreased (blue) in the area of the rupture surface, but it may be increased (red) on adjacent parts of the fault.

Stress transfer isn’t necessarily restricted to the fault along which an earthquake happened. It will affect the rocks in general around the site of the earthquake, and therefore may lead to increased stress on other faults in the region. And the effects of stress transfer don’t necessarily show up right away. Segments of faults are typically in some state of stress, and the transfer of stress from another area is only rarely enough to push a fault segment beyond its limits to the point of rupture. The stress that is added by stress transfer accumulates along with the ongoing build-up of stress from plate motion, to eventually lead to an earthquake.

Box 6.1 Episodic Tremor and Slip

Episodic tremor and slip is periodic slow sliding along part of a subduction boundary. It does not produce recognizable earthquakes, but does produce seismic tremor (rapid seismic vibrations on a seismometer that cannot be felt by humans). It was first discovered in 2003 on the Vancouver Island part of the Cascadia subduction zone.

The boundary between the subducting Juan de Fuca plate and the North America plate can be divided into three segments, as shown on Figure 6.1.6. The rocks are locked together in the cold upper part of the boundary, and only move when there is a very large earthquake, in this case approximately every 500 years (the last one was M8.5+ on January 26, 1700). The lower part of the boundary is sliding continuously because the rock is warm and weak. The central part of the boundary isn’t cold enough to be locked, but isn’t warm enough to slide continuously. Instead it slips episodically, approximately every 14 months, slowly moving a few centimetres each time over a duration of about 2 weeks.

Figure 6.1.6 Depiction of the Parts of the Cascadia Subduction Zone. Large infrequent earthquakes happen in the locked zone.

You might be inclined to think that it’s a good thing that there is periodic slip on this part of the plate because it releases some of the tension and reduces the risk of a large earthquake. In fact, the opposite is likely the case. The movement along the ETS part of the plate boundary acts like a medium-sized earthquake and leads to stress transfer to the adjacent locked part of the plate. Approximately every 14 months, during the two-week ETS period, there is a transfer of stress from the ETS zone up to the shallow locked part of the Cascadia subduction zone, and therefore an increased chance of a large earthquake.
Since 2003 ETS processes have also been observed on subduction zones in Mexico and Japan.

Media Attributions

• Figure 6.1.1 Steven Earle, CC BY 4.0
• Figure 6.1.2 Steven Earle, CC BY 4.0
• Figure 6.1.3 Steven Earle, CC BY 4.0
• Figure 6.1.4 Steven Earle, CC BY 4.0, based on figures in Ward, P, L. & Page, R.A., (1990). The Loma Prieta earthquake of October 17, 1989: A brief geologic view of what caused the Loma Prieta earthquake and implications for future California earthquakes: What happened … what is expected … what can be done, U.S. Geological Survey, https://doi.org/10.3133/70039527
• Figure 6.1.5 Steven Earle, CC BY 4.0
• Figure 6.1.6 Steven Earle, CC BY 4.0

The distribution of earthquakes across the globe is shown on Figure 6.2.1. It is relatively easy to see the relationships between earthquakes and the plate boundaries. Along divergent boundaries like the mid-Atlantic ridge and the East Pacific rise earthquakes are common, but restricted to a narrow zone close to the ridge, and consistently less than 30 km depth (red dots). Shallow earthquakes are also common along transform faults, such as the San Andreas Fault. Earthquakes are very common along subduction zones, and their depth below surface increases inshore from the subduction zone (green and blue dots).

Figure 6.2.1 General Distribution of Global Earthquakes of Magnitude 4 and Greater During the Period from 2004 to 2011. Colour coded by depth (red: 0-33 km, orange 33-70 km, green: 70-300 km, blue: 300-700 km)

Earthquakes also occur in a few intraplate locations, including the Rift Valley area of Africa, in the Tibet region of China and in the Lake Baikal area of Russia.

Earthquakes at Divergent and Transform Boundaries

Figure 6.2.2 provides a closer look at magnitude 4 and larger earthquakes in an area of divergent and transform boundaries in the mid-Atlantic region near to the equator. Here the segments of the mid-Atlantic ridge (divergent boundaries) are offset by some long transform faults, and there is side-by-side motion on those faults. Most of the earthquakes are located along the transform faults, rather than along the divergent boundaries, although there are clusters of earthquakes at some of the divergent-transform intersections. Some earthquakes do occur on divergent boundaries, but they tend to be small and infrequent because of the relatively high rock temperatures in the areas where spreading is taking place.

Figure 6.2.2 Distribution of Earthquakes of Magnitude 4 and Greater in the Area of the Mid-Atlantic Ridge Near to the Equator During the Period From 1990 to 1996. All are 0 to 33 km depth.

Earthquakes at Convergent Boundaries

The distribution and depths of earthquakes in the Caribbean and Central America area are shown on Figure 6.2.3. In this region the Cocos Plate is subducting beneath the North America and Caribbean Plates (an ocean-continent convergence), and the South and North America Plates are subducting beneath the Caribbean Plate (an ocean-ocean convergence). In both cases the earthquakes get deeper with distance from the trench. The South America Plate is shown to be subducting beneath the Caribbean Plate in the area north of Columbia, but since there are almost no earthquake along this zone, it is questionable whether subduction is actually taking place.

Figure 6.2.3 Distribution of Earthquakes of Magnitude 4 and Greater in the Central America Region From 1990 to 1996. (red: 0-33 km, orange 33-70 km, green: 70-300 km, blue: 300-700 km) (Spreading ridges are heavy blue lines, subduction zones are toothed lines, and transform faults are light lines.)

Figure 6.2.4 Distribution of Earthquakes in the Area of the Kuril Islands, Russia (Just North of Japan). White dots represent the April 2009 magnitude 6.9 earthquake. Red and yellow dots are from background seismicity over several years prior to 2009.

There are also various divergent and transform boundaries in the area shown on Figure 6.2.3, and, as in the mid-Atlantic area, most of the earthquakes are along the transform faults, and are relatively shallow.

The distribution of earthquakes with depth in the Kuril Islands area in the northwest Pacific is shown on Figure 6.2.4. This is an ocean-ocean convergent boundary. The small red and yellow dots show background seismicity over a number of years, while the larger white dots are individual shocks associated with a magnitude 6.9 earthquake in April 2009. The relatively large earthquake took place on the upper part of the plate boundary between 60 and 140 km inland from the trench. As we saw for the Cascadia subduction zone, this is where large subduction earthquakes are expected to occur.

In fact, all of the very large earthquakes—M9 or higher—take place at subduction boundaries because there is the potential for greater rupture zone width on a gently dipping subduction boundary than on a steep transform boundary. The largest earthquakes on transform boundaries are in the order of M8.

The background seismicity at this convergent boundary (Figure 6.2.4), and on other similar ones, is predominantly near to the upper side (crust side) of the subducting Pacific plate. The frequency of earthquakes is greatest near to surface and especially around the region of large subduction quakes (white dots), but it extends down to at least 400 km depth. There is also significant seismic activity in the over-riding North America plate, again most commonly near to region of large quakes, but also extending for a few hundred kilometres away from the plate boundary.

The distribution of earthquakes in the India-Asia area is shown on Figure 6.2.5. This is a continent-continent convergent boundary, and it is generally assumed that although the India Plate continues to move north towards the Asia plate, there is no actual subduction taking place. There are transform faults on either side of the India plate in this area.

Figure 6.2.5 Distribution of Earthquakes in the Area Where the India Plate is Converging With the Asia Plate. (Data from 1990 to 1996, red: 0-33 km, orange 33-70 km, green: 70-300 km, blue: 300-700 km) Spreading ridges are heavy lines, subduction zones are toothed lines, and transform faults are light lines. The double line along the northern edge of the India plate indicates convergence, but not subduction. Plate motions are shown in mm/y.

The entire northern India and southern Asia region is very seismically active. Earthquakes are common in northern India, Nepal, Bhutan, Bangladesh and adjacent parts of China, and throughout Pakistan and Afghanistan. Many of the earthquakes are related to the transform faults on either side of the India plate, and most of the others are related to the significant tectonic squeezing caused by the continued convergence of the India and Asia plates. That squeezing has caused the Asia plate to be thrust over top of the India plate, building the Himalayas and the Tibet Plateau to enormous heights. Most of the earthquakes of Figure 6.2.3 are related to the thrust faults shown on Figure 6.2.6 (and to hundreds of other similar ones that cannot be shown at this scale). The southernmost thrust fault on this diagram is equivalent to the Main Boundary Fault on Figure 6.2.4.

Figure 6.2.6 Schematic Diagram of the India-Asia Convergent Boundary, Showing Examples of the Types of Faults Along which earthquakes are focused.

There is a very significant concentration of both shallow and deep (greater than 70 km) earthquakes in the northwestern part of Figure 6.2.5. This is northern Afghanistan, and at depths of more than 70 km, many of these earthquakes are within the mantle as opposed to the crust. It is interpreted that these deep earthquakes are caused by northwestward subduction of the India plate beneath the Asia plate in this area.

Exercise 6.1 Earthquakes in Washington and British Columbia

Figure 6.2.7 shows the incidence and magnitude of earthquakes in southwestern British Columbia and northwestern Washington over a one-year period from January 2020 to January 2021.

Figure 6.2.7 Earthquakes Recorded in Southwestern B.C. and Neighbouring Parts of Washington, From January 2020 to January 2021

1. What is the likely origin of the earthquakes between the JDF and Explorer plates?
2. What is the sense of motion of the plate boundary that aligns with the string of earthquakes adjacent to Haida Gwaii.
3. Most of the small earthquakes around Vancouver Island and south into northwest Washington are relatively shallow. What is their likely origin?

Exercise answers are provided Appendix 2.

Media Attributions

• Figure 6.2.1 Dale Sawyer, Rice University, http://plateboundary.rice.edu Used with permission. All rights reserved.
• Figure 6.2.2 Dale Sawyer, Rice University. http://plateboundary.rice.edu. Used with permission. All rights reserved.
• Figure 6.2.3 Dale Sawyer, Rice University. http://plateboundary.rice.edu. Used with permission. All rights reserved.
• Figure 6.2.4 Steven Earle, CC BY 4.0, based on information in Hayes, G. P., Wald, D. J., and Johnson, R. L. (2012). Slab1.0: A three-dimensional model of global subduction zone geometries. Journal of Geophysical Research, 117, B01302. https://doi.org/10.1029/2011JB008524
• Figure 6.2.5 Dale Sawyer, Rice University. http://plateboundary.rice.edu. Used with permission. All rights reserved.
• Figure 6.2.6 Steven Earle, CC BY 4.0, based on a drawing by Vuichard, D. (n.d.). The Himalayan region: a geographical overview. In Ives, J. D. & Messerli, B. (Eds.), The Himalayan Dilemma. United Nations University. http://archive.unu.edu/unupress/unupbooks/80a02e/80A02E05.htm
• Figure 6.2.7 Steven Earle, CC BY 4.0, based on a Open Government Licence – Canada map by Natural Resources Canada at https://www.seismescanada.rncan.gc.ca/recent/maps-cartes/index-1y-en.php?tpl_region=west

There are two main ways to measure earthquakes. The first of these is an estimate of the energy released, and the value is referred to as magnitude. This is the number that is typically first released by the press when a big earthquake happens. It is often referred to as “Richter magnitude”, but that is a misnomer, and it should be just “magnitude”. There are many ways to measure magnitude—including Charles Richter’s method developed in 1935—but they are all ways to estimate the same number: the amount of energy released.

The other way of assessing the impact of an earthquake is to assess what people felt and how much damage was done. This is known as intensity. Intensity values are assigned to locations, rather than to the earthquake itself, and intensity can vary widely therefore, depending on the proximity to the earth quake and the type of ground underneath.

Earthquake Magnitude

Before we look more closely at magnitude we need to review what we know about body waves, and also look at surface waves. Body waves are of two types, P or primary or compression waves (like the compression of the coils of a spring), and S or secondary or shear waves (like the flick of a rope). An example of P and S seismic records is shown on Figure 6.3.1. The critical parameters for the measurement of Richter magnitude are labelled, including the time interval between the arrival of the P and S waves—which is used to determine the distance from the earthquake to the seismic station, and the amplitude of the S waves—which is used to estimate the magnitude.

Figure 6.3.1 P and S Waves From a Small (M 4) Earthquake that Took Place Near to Vancouver Island in 1997.

When a body wave (P or S) reaches the Earth’s surface some of its energy is transformed into surface waves, of which there are two main types, as illustrated in Figure 6.3.2. Rayleigh waves are characterized by vertical motion of the ground surface, like waves on water, while Love waves are characterized by horizontal motion. Both Rayleigh and Love waves are about 90% as fast as S waves (so they arrive later at a seismic station). Surface waves typically have greater amplitudes than body waves, and they do more damage.

Figure 6.3.2 Depiction of Seismic Surface Waves

Figure 6.3.3 Definition of Epicentre and Hypocentre

Two other important terms from the perspective of describing earthquakes are hypocentre and epicenter. The hypocentre is the actual location of an individual earthquake shock at depth in the ground, and the epicentre is the point on the land surface directly above the hypocentre (Figure 6.3.4).

A number of methods for estimating magnitude are listed in Table 6.1. Local magnitude (M_L) was widely used until late in the 20th century, but moment magnitude (M_W) is now more commonly used because it gives more accurate estimates (especially with larger earthquakes) and can be applied to earthquakes at any distance from a seismometer. Surface-wave magnitudes can also be applied to large distant earthquakes.

Because of the increasing size of cities in earthquake-prone areas (e.g., China, Japan, California and Turkey) and the increasing sophistication of infrastructure, it is becoming important to have very rapid warnings and magnitude estimates of earthquakes that have already happened. This can be achieved by using P-wave data because P waves arrive first at seismic stations, in many cases several seconds ahead of the more damaging S waves and surface waves. Operators of electrical grids, pipelines, trains and other infrastructure can use the information to automatically shut systems down so that damage and casualties can be limited.

The magnitude scale is logarithmic, in fact the amount of energy released by an earthquake of magnitude 4 is 32 times higher than that released by one of magnitude 3, and this ratio applies to all intervals in the scale. If we assign an arbitrary energy level of 1 unit to a magnitude 1 earthquake the energy for quakes up to magnitude 8 will be as shown on the following list:

In any given year when there is a large earthquake on Earth (M 8 or 9) the amount of energy released by that one event will likely exceed the energy released by all smaller events combined.

Earthquake Intensity

The intensity of earthquake shaking at any location is determined by the magnitude of the earthquake and its distance, but also by the type of underlying rock or unconsolidated materials. If buildings are present, the size and type of building are also important.

Intensity scales were first used in the late 19th century, and then adapted in the early 20th century by Giuseppe Mercalli and modified later by others to form what we know call the Modified Mercalli Intensity scale (Figure 6.3.5). Intensity estimates are important because they allow us to characterize parts of any region into areas that are especially prone to strong shaking versus those that are not. The key factor in this regard is the nature of the underlying geological materials, and the weaker those are the more likely it is that there will be strong shaking. Areas underlain by strong solid bedrock tend to experience much less shaking that those underlain by unconsolidated river or lake sediments.

Figure 6.3.4 The Modified Mercalli Intensity Scale

An example of this amplifying effect is provided by the 1985 M8 earthquake which struck the Pacific coast Michoacán region of western Mexico, about 350 km southwest of Mexico City. There was relatively little damage in the area around the epicentre, but there was tremendous damage and about 5000 deaths in heavily populated Mexico City. The key reason for this is that Mexico City was built largely on the unconsolidated and water-saturated sediment of former Lake Texcoco. These sediments resonate at a frequency of about 2 seconds, which was similar to the frequency of the body waves that reached the city. For the same reason that a powerful opera singer can break a wine glass by singing the right note, the seismic shaking was amplified by the lake sediments. Survivors of the disaster recounted that the ground in some areas moved up and down by about 20 cm every 2 seconds for over 2 minutes. Damage was greatest to buildings between 5 and 15 stories tall, because they also resonated at around 2 seconds, and amplified the shaking.

Figure 6.3.5 Intensity Map for the 1946 M7.3 Vancouver Island Earthquake.

An intensity map for the M7.3 June 1946 Vancouver Island Earthquake is shown on Figure 6.3.5. The intensity was greatest in the central island region where, in some communities, chimneys were damaged on more than 75% of buildings, some roads were made impassable and a major rock slide occurred. The earthquake was felt as far north as Prince Rupert, as far south as Portland Oregon and as far east as the Rockies.

Media Attributions

• Figure 6.3.1 Steven Earle, CC BY 4.0, after a Open Government Licence – Canada image provided by Natural Resources Canada
• Figure 6.3.2 Modified by Steven Earle, from images via Wikipedia: https://en.wikipedia.org/wiki/Rayleigh_wave#/media/File:Rayleigh_wave.jpg, Public domain, and https://en.wikipedia.org/wiki/Love_wave#/media/File:Love_wave.jpg, CC BY 4.0
• Figure 6.3.3 Steven Earle, CC BY 4.0
• Figure 6.3.4 Steven Earle, CC BY 4.0, based on the modified scale by the US Geological Survey (public domain), https://www.usgs.gov/natural-hazards/earthquake-hazards/science/modified-mercalli-intensity-scale?qt-science_center_objects=0#qt-science_center_objects
• Figure 6.3.5 Intensity Map for 1946 M7.3 Vancouver Island Earthquake by Earthquakes Canada, Open Government Licence – Canada  http://www.earthquakescanada.nrcan.gc.ca/historic-historique/events/19460623-eng.php

Figure 6.4.1 A Part of the Nimitz Freeway in Oakland California that Collapsed During the 1989 Loma Prieta Earthquake.

Some of the common impacts of earthquakes include structural damage to buildings, fires, damage to bridges and highways, slope failures, liquefaction and tsunami. The types of impacts will depend to a large degree on the type of area where the earthquake is located: whether it is predominantly urban or rural, densely or sparsely populated, highly developed or under-developed, and of course on the ability of the infrastructure to withstand shaking.

As we’ve seen from the example of the 1985 Mexico earthquake, the geological foundations on which structures are built can have a significant impact on earthquake shaking. When an earthquake happens the seismic waves produced have a wide range of frequencies. The energy of the higher frequency waves tends to be absorbed by the solid rock of the crust, while the lower frequency waves (with periods slower than 1 second) pass through the solid rock without being absorbed, but are eventually absorbed by soft sediments. In many cases the soft sediments amplify the seismic shaking, so it is very common to see much worse earthquake damage in areas underlain by soft sediments than in areas of solid rock. A good example of this is in the Oakland area near to San Francisco, where parts of a two layer highway built on soft sediments collapsed during the 1989 Loma Prieta earthquake (Figure 6.4.1).

Building damage is also worse in areas of soft sediments, and multi-story buildings tend to be more seriously damaged than smaller ones. Buildings can be designed to withstand most earthquakes, and this practice is increasingly applied in earthquake-prone regions. Turkey is one such region, and even though Turkey had a relatively strong building code in the 1990s, adherence to the code was weak, as builders did whatever they could to save costs, including using inappropriate materials in concrete and reducing the amount of steel reinforcing. The result was that there were over 17,000 deaths in the M7.6 1999 Izmit earthquake (Figure 6.4.2). After two devastating earthquakes in 1999 Turkish authorities strengthened the building code, but the new code has only been applied in a few regions, and enforcement of the code is still weak, as revealed by the amount of damage from a M7.1 earthquake in eastern Turkey in 2011.

Figure 6.4.2 Buildings Damaged by the 1999 Earthquake in the Izmit Area, Turkey

Some of the common impacts of earthquakes include structural damage to buildings, fires, damage to bridges and highways, slope failures, liquefaction and tsunami. The types of impacts will depend to a large degree on the type of area

Figure 6.4.3 Some of the Effects of the 2011 Tohoku Earthquake in the Sendai Area of Japan. An oil refinery is on fire, and a vast area has been flooded by a tsunami.

where the earthquake is located: whether it is predominantly urban or rural, densely or sparsely populated, highly developed or under-developed, underlain by weak sediments or strong hard rock, and of course on the ability of the infrastructure to withstand shaking.

Figure 6.4.4 Fires in San Francisco During the 1906 Earthquake

Fires are commonly associated with earthquakes because gas pipelines get ruptured and electrical lines get damaged when the ground shakes (Figure 6.4.3). Most of the damage in the great 1906 San Francisco earthquake was caused by massive fires in the downtown area of the city (Figure 6.4.4). Some 25,000 buildings were destroyed by those fires, which were fueled by broken gas pipes. Fighting the fires was difficult because water mains were also ruptured. The risk of fires can be reduced through P-wave early warning systems if utility operators can reduce pipeline pressure and close electrical circuits.

Figure 6.4.5 The Las Colinas Debris Flow at Santa Tecla (A Suburb of the Capital San Salvador) Triggered by the January 2001 El Salvador Earthquake. This is just one of many hundreds of slope failures that resulted from that earthquake. Over 500 people died in the area affected by this slide.

Earthquakes are important triggers for failures on slopes that are already prone to weakness. An example is the Las Colinas slide in the city of Santa Tecla, El Salvador, which was triggered by a M7.6 offshore earthquake in January 2001 (Figure 6.4.5).

Figure 6.4.6 Collapsed Apartment Buildings in the Niigata Area, Japan. The material beneath the buildings was liquefied to varying degrees by the 1964 Niigata earthquake.

Another example is from the 2018 Sapporo earthquake in Japan which caused thousands of debris flows in an area that had recently been soaked by summer rains and then a typhoon (see Figure 5.1.7 in Chapter 5).

Ground shaking during an earthquake can be enough to weaken rock and unconsolidated materials to the point of failure, but in many cases the shaking also contributes to a process known as liquefaction, in which an otherwise solid body of sediment gets transformed into a liquid mass that can flow. When water-saturated sediments are shaken the grains become rearranged to the point where they are no longer supporting one-another. Instead, the water between the grains is holding them apart and the material can flow. Liquefaction can lead to the collapse of buildings and other structures that might be otherwise undamaged. A good example is the collapse of apartment buildings during the 1964 Niigata earthquake (M7.6) in Japan (Figure 6.4.6). Liquefaction can also contribute to slope failures and to fountains of sandy mud (sand “volcanoes”) in areas where there is loose saturated sand beneath a layer of more cohesive clay.

Parts of the Fraser River delta near to Vancouver BC are prone to liquefaction-related damage because the region is characterized by a 2 to 3 m thick layer of fluvial silt and clay over top of at least 10 m of water-saturated fluvial sand (Figure 6.4.7). Under these conditions it can be expected that seismic shaking will be amplified and also that the sandy sediments will be liquefied. This could lead to subsidence of buildings to tilting in areas where liquefaction is inconsistent, and to failure and sliding of the silt and clay layer. Current building-code regulations in the Fraser Delta area require that measure be taken to strengthen the ground underneath before construction of multi-story buildings.

Figure 6.4.7 Recent Sedimentary Layers in the Fraser Delta Area and the Potential Consequences in the Event of a Damaging Earthquake.

Exercise 6.4  Creating Liquefaction and Discovering the Harmonic Frequency

There are a few ways that you can demonstrate the process of liquefaction for yourself. The simplest is to go to a sandy beach (lake, ocean or river) and find a place near to the water’s edge where the sand is wet. This is best done with your shoes off, so let’s hope it’s not too cold! While standing in one place on a wet part of the beach start moving your feet up and down at a frequency of about once per second. Within a few seconds the previously firm sand will start to lose strength and you’ll gradually sink in up to your ankles.

If you can’t get to a beach, or if the weather isn’t cooperating, put some sand (sandbox sand will do) into a small container, saturate it with water, and then pour the excess water off. You can shake it gently to get the water to separate and then pour the excess water away, and you may have to do that more than once. Place a small rock on the surface of the sand; it should sit there for hours without sinking in. Now, holding the container in one hand gently thump the side or the bottom with your other hand, about twice a second. The rock should gradually sink in as the sand around it becomes liquefied (Figure 6.4.8)

Figure 6.4.8 Demonstrating Liquefaction with Sand

As you were moving your feet up and down or thumping the pot, it’s likely that you soon discovered the most effective rate for getting the sand to liquefy, and this would have been close to the natural harmonic frequency for that body of material. Stepping up and down as fast as you can (several times per second) on the wet beach would not have been effective, nor would you have achieved much by stepping once every several seconds. The body of sand vibrates most readily in response to shaking that is close to its natural harmonic frequency, and liquefaction is also most likely to occur at that frequency.

One of the ways to reduce risks to people and property from earthquakes is to understand the types of materials that buildings are situated on in cities. This involves mapping the distribution, thickness and nature of surficial materials such as river sediments, lake sediments, and glacial sediments, and also the degree to which these materials are saturated with water. This type of work has been done in many areas around the world, and an example of such a map is provided on Figure 6.4.9 for the city of Victoria, British Columbia.

Figure 6.4.9 Amplification and Liquefaction Risk Map for Victoria, British Columbia

Large parts of Victoria are underlain by bedrock or a very thin layer of sediments on top of bedrock. In other areas there are significant thicknesses of glacial sediments, including till, clay deposited under marine conditions in post-glacial times and post-glacial peat deposits . In several near-shore areas land has been claimed using fill. The earthquake hazard map of Figure 6.4.8 shows lowest amplification risk in bedrock areas (grey), moderate risk in areas of marine clay (orange and pink) and in some areas of near-shore fill (green), and greatest risk in areas where there is peat over top of clay (red).

Liquefaction risk is greatest where the loose sediments are saturated with water, and this applies to some near-shore areas, especially around Victoria Harbour, and in some inland areas where there is peat over clay.

Tsunami

Earthquakes that take place beneath the ocean have the potential to generate tsunami under certain conditions. The most likely situation for a significant tsunami is a large (M7 or greater) subduction-related earthquake. As shown on Figure 6.4.10, during the time between earthquakes the overriding plate gets slowly distorted by elastic deformation; it is squeezed laterally (Figure 6.4.10B) and it is pushed up. When an earthquake happens that deformed crust springs back, resulting in either rapid subsidence of the sea floor, rapid uplift of the seafloor, or both.

Figure 6.4.10 Elastic Deformation and Rebound of Overriding Plate at a Subduction Setting (B). The release of the locked zone during an earthquake (C) results in both uplift and subsidence on the sea floor, and this is transmitted to the water overhead, resulting in a tsunami.

Subduction earthquakes with magnitude less than 7 do not typically generate significant tsunami because the amount of vertical displacement of the sea floor is minimal. Sea-floor transform earthquakes, even large ones (M7 to 8), don’t typically generate tsunami either, because the motion is mostly side-to-side, and not vertical. And, of course, earthquakes that take place entirely on land don’t generate tsunami.

Tsunami waves travel at velocities of several hundred km/h and easily make it to the far side of an ocean in about the same time as a passenger jet. The simulated one shown on Figure 6.4.11, is similar to that created by the January 1700 Cascadia earthquake off the coast of British Columbia, Washington, and Oregon, which was recorded in Japan 9 hours later.

Figure 6.4.11 Model of the Tsunami from the 1700 Cascadia Earthquake (~M9) Showing Open-Ocean Wave Heights (Colours) and Travel Time Contours. Tsunami wave amplitudes typically increase dramatically in shallow water.

In many earthquake events the damage and loss of life from a tsunami is much greater than those from the earthquake shaking. This certainly applies to the massive (M9.1) 2004 Sumatra earthquake and tsunami, for which the death toll was well over 200,000. An example of the damage from that event is shown of Figure 6.4.12. In fact there is no certainty about the proportion of deaths related to shaking and building collapse versus those from the tsunami because many of the buildings in coastal areas that might have collapsed in the shaking were soon destroyed by the waves.

Figure 6.4.12 Damage to a Community in Aceh, Indonesia from the Tsunami Associated with the 2004 Earthquake.  The only building still standing is a large mosque.

Approximately 16,000 of the deaths from Japan’s 2011 Tohoku earthquake were a result of drowning or in some other way related to the tsunami , while about 3000 are described as being related to the earthquake . Most of the damage to structures was also caused by the tsunami, including the devastating damage to the Fukushima Daiichi nuclear power station.

Media Attributions

• Figure 6.4.1 Nimitz Freeway by Joe Lewis, 1989, CC BY SA 2.0, via Flickr,  https://www.flickr.com/photos/sanbeiji/220646891
• Figure 6.4.2 Collapsed buildings, US Geological Survey public domain image, https://gallery.usgs.gov/media/galleries/1999-izmit-turkey
• Figure 6.4.3 Helicopter flying over Sendai, US Navy public Domain image, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:SH-60B_helicopter_flies_over_Sendai.jpg
• Figure 6.4.4 San Francisco Fire, 1906, public domain image via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:San_francisco_fire_1906.jpg
• Figure 6.4.5 El Salvador Slide, Public domain US Geological Survey image via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:ElSalvadorslide.jpg
• Figure 6.4.6 Liquefaction at Niigata, public domain image via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Liquefaction_at_Niigata.JPG
• Figure 6.4.7 Steven Earle, CC BY 4.0
• Figure 6.4.8 Photos by Steven Earle, CC BY 4.0
• Figure 6.4.9 Modified by Steven Earle, CC BY 4.0, after Monahan, P., Levson, V., McQuarrie, E., Bean, S., Henderson, P., & Sy, A, (2000). Geoscience Map 2000-1. BC Ministry of Energy and Mines, Geological Survey Branch. https://www2.gov.bc.ca/gov/content/industry/mineral-exploration-mining/british-columbia-geological-survey/publications/geosciencemaps
• Figure 6.4.10 Steven Earle, CC BY 4.0
• Figure 6.4.11 Tsunami from the 1700 Cascadia Earthquake simulated model, Public domain image, NOAA/PMEL/Center for Tsunami Research, http://nctr.pmel.noaa.gov/cascadia_simulated/
• Figure 6.4.12 Tsunami damage to a Community in Aceh, Indonesia, Public domain image, US Navy, https://nara.getarchive.net/media/a-lone-mosque-stands-among-the-damage-of-a-coastal-village-near-aceh-sumatra-86b33f

It has long been a dream of seismologists, geologists and public safety officials to be able to accurately predict earthquakes on time scales (weeks, days or hours) that would be useful for minimizing danger to the public and damage to infrastructure. Many different avenues have been explored, such as warning foreshocks, changes in magnetic fields, seismic tremor, changing groundwater levels, strange animal behaviour, observed earthquake periodicity, stress transfer considerations and a range of others. So far, none of the research into earthquake prediction has provided a reliable method. Although there are reports of successful earthquake predictions, they are rare, and many are surrounded by doubtful circumstances.

The problem with earthquake predictions, as with any other type of prediction, is that they have to be accurate most of the time, not just some of the time. We have come to rely on weather predictions because they are generally (and increasingly) accurate. But if we try to predict earthquakes and are only accurate 10% of the time (and even that isn’t likely with the current state of knowledge), the public will lose faith in the process very quickly, and then all of the predictions will be ignored.

There was a great hope for earthquake predictions late in the 1980s, when attention was focused on part of the San Andreas fault at Parkfield, about 200 km south of San Francisco. Between 1881 and 1966 there were 5 earthquakes at Parkfield, most spaced at approximately 20-year intervals, all confined to the same 20 km-long segment of the fault and all very close to M6 (Figure 6.5.1). Both the 1934 and 1966 earthquakes were preceded by small foreshocks exactly 17 minutes before the main quake.

Figure 6.5.1 Earthquakes on the Parkfield Segment of the San Andreas Fault, 1881 to 2004.

The U.S. Geological Survey recognized this as an excellent opportunity to understand earthquakes and earthquake prediction, so they armed the Parkfield area with a huge array of geophysical instruments and waited for the next quake, which was expected to happen around 1987. Nothing happened! The “1987 Parkfield earthquake” finally struck in September 2004, about 17 years late. Fortunately, all of the equipment was still there, but it was to no avail from the perspective of earthquake prediction. There were no significant precursors to the 2004 Parkfield earthquake in any of the parameters measured, including: seismicity, harmonic tremor, strain (rock deformation), magnetic field, the conductivity of the rock or creep, and there was no foreshock. In other words, even though every available technique was used to monitor it, the 2004 Parkfield earthquake came as a complete surprise, with no warning whatsoever.

The hope for earthquake prediction is not dead, but it was hit hard by the Parkfield experiment. The current focus in earthquake-prone regions is to provide forecasts of earthquake probabilities within a certain time period—typically a number of decades—and also to ensure that the population is educated about earthquake risks and that buildings and other infrastructure are as safe as can be. An example of this approach for the San Francisco Bay region of California is shown on Figure 6.5.2. Based on a wide range of information, including past earthquake history, accumulated stress from plate movement, and known stress transfer, seismologists and geologists have predicted the likelihood of a M6.7 or greater earthquake on each of 8 major faults that cut through the region. The greatest probabilities are on the San Andreas, Rogers Creek and Hayward faults. As shown on the diagram, there is a 72% chance that a major and damaging earthquake will take place somewhere in the region prior to 2043.

Figure 6.5.2 Probabilities of a M6.7 or Larger Earthquake on Various Faults in the San Francisco Bay Region of California, 2014 to 2043.

It is feasible to provide useful warnings of earthquakes that have already happened to urban areas that are some distance from the epicentre. As described in Section 6.4, the earthquake in Mexico in 1985 took place near to the coast but had devastating effects in Mexico City, several hundred kilometres away. In 1993 the Mexican Government established an early warning system for the capital. It consists of a network of seismometers along the coast that will relay an electronic warning to emergency officials in Mexico City who then activate alert systems in hospitals and schools, and public alarms on 12,000 pole-mounted speakers throughout the city. Because seismic waves can take well over a minute to reach Mexico City from the coast, and because the warning messages are almost instant, residents of the capital typically have up to two minute’s warning that an earthquake has happened and shaking can be expected. According to Cochran et al (2018) the system has wide public acceptance, even though many of the earthquakes that have been warned of were too small to cause any damage in Mexico City.

Japan established a nation-wide earthquake warning system in 2008 , and the US Geological Survey started work on the ShakeAlert system for Washington, Oregon and California in 2016, although it will be some years before that system is fully functional. Both of these systems are based on seismometers on land.

A unique earthquake early warning system, that uses seismometers on the sea floor, is being developed in British Columbia. The initiative is based on Ocean Networks Canada existing fibre-optics communication and instrumentation array on the sea floor off the west coast of Vancouver Island (Figure 6.5.3). Eight seismometers situated up to 200 km offshore, networked with roughly 100 seismometers on land, can detect a large earthquake up to 90 seconds before the shaking could reach major centres such as Vancouver and Victoria .

Figure 6.5.3 Ocean Networks Canada Earthquake Early Warning System for Southwestern British Columbia

Figure 6.5.4 Strong Diagonal Bracing in the Pearl River Tower in Guanzhou, China

A critical component of reducing the damage and casualties from earthquakes is to ensure that buildings and other infrastructure (bridges, dams, roads etc.) are built to withstand strong shaking. This starts from ensuring that the foundation is constructed on strong material, and continues right up to ensuring that furniture and movable items (bookcases, water heaters etc.) are secure. Buildings themselves can be constructed to resist and withstand shaking by including diagonal bracing (Figure 6.5.4) and flexible foundations. Wooden buildings tend to perform well during shaking because of the flexibility of the wood. Bridges can also be designed to resist shaking. An interesting example of a bridge designed to withstand earthquakes is illustrated on Figure 6.5.5.

Figure 6.5.5 A bridge that spans the San Andreas Fault south of Parkfield California. The bridge deck is resting on concrete piers and it can slide as necessary when the foundations at either end move in different directions during an earthquake

As we’ve discussed already, it’s not sufficient to have strong building codes, they have to be enforced. Building code compliance is quite robust in most developed countries, but is not adequate in many developing countries.

It’s also not enough just to focus on new buildings, we have to make sure that existing buildings— especially schools and hospitals—and also other structures such as bridges and dams, are as safe as they can be. Efforts to upgrade the safety of schools in British Columbia are described in Box 6.2.

Box 6.2 Making the Seismic Up-grade in BC’s schools

British Columbia is in the middle of a multi-billion dollar program to make schools safer for students. The program is focused on older schools, because, according to the government, those built since 1992 already comply with modern seismic codes. Some schools would require too much work to make upgrading economically feasible and they are replaced. Where upgrading is feasible, the school is assessed carefully before any upgrade work is initiated.

An example is Sangster Elementary in Colwood on southern Vancouver Island (Figure 6.5.6). The school was originally built in 1957, with a major addition in 1973. Ironically, the newer part of the school, built of concrete blocks, required strengthening with the addition of a steel framework, while the 1957 part, which is a wood-frame building, did not require seismic upgrading. The work was completed in 2014.

Figure 6.5.6 Sangster Elementary School, Victoria, BC

As of May 2021 upgrades had been completed at 186 B.C. schools, 30 were underway, and an additional 13 were ready to proceed with funding identified. Another 267 schools were listed as needing upgrades.

The final part of earthquake preparedness involves the formulation of public emergency plans, including escape routes, medical facilities, shelters, food and water supplies. It also includes personal planning, such as emergency supplies (food, water, shelter and warmth), escape routes from houses and offices, and communication strategies (with a focus on ones that don’t involve the cellular network, which may not be functioning after a large earthquake).

Media Attributions

• Figure 6.5.1 Steven Earle, CC BY 4.0
• Figure 6.5.2 Public domain image from Aagard, B. et al, (2014). Earthquake outlook for the San Francisco Bay region, 2014 to 2043. US Geological Survey, https://pubs.usgs.gov/fs/2016/3020/fs20163020.pdf
• Figure 6.5.3 Image from Innovation Center. (n.d.). Earthquake early warning. Ocean Networks Canada. https://www.oceannetworks.ca/. Used with permission.
• Figure 6.5.4 Pearl River Tower, public domain image by Brad Wilkins, via Wikimedia Commons,  https://commons.wikimedia.org/wiki/File:2009_03_03_Pearl_River_Tower.jpg
• Figure 6.5.5 Photo by Steven Earle, CC BY 4.0
• Figure 6.5.6 Google Maps – street view

The relationships between volcanism and plate tectonics are summarized on Figure 7.1.1. As outlined in Chapter 2, magma is formed at three main plate-tectonic settings: divergent boundaries (decompression melting), convergent boundaries (flux melting), and mantle plumes (decompression melting).

Figure 7.1.1 The Plate-Tectonic Settings of Common Types of Volcanism. Most composite volcanoes form at subduction zones, either on ocean-ocean convergent boundaries (left) or ocean-continent convergent boundaries (right). Most shield volcanoes form above mantle plumes, but can also form at divergent boundaries. Sea-floor volcanism can take place at divergent boundaries, mantle plumes and ocean-ocean-convergent boundaries.

The mantle and crustal processes that take place in areas of volcanism are illustrated in a little more detail in Figure 7.1.2. At a spreading ridge hot mantle rock moves slowly upward by convection (cm/year) and within about 60 km of surface partial melting starts because of decompression. Over the triangular area shown on Figure 7.1.2a about 10% of the ultramafic mantle rock melts, producing mafic magma that moves upward toward the axis of spreading (where the two plates are moving away from each other). The magma fills vertical fractures produced by the spreading, and it spills out on the sea floor to form basaltic pillows (more on that later) and lava flows. There is spreading-ridge volcanism taking place about 200 km offshore from the west coast of Vancouver Island.

Figure 7.1.2 The Processes that Lead to Volcanism in the Three Main Volcanic Settings on Earth. (a) Volcanism related to plate divergence, (b) volcanism at an ocean-continent or ocean-ocean convergent boundary, and (c) volcanism related to a mantle plume.

At an ocean-continent or ocean-ocean convergent boundary oceanic crust is pushed down into the mantle (Figure 7.1.2b). It gets heated up, and while there isn’t enough heat to melt the relatively cool subducting oceanic crust, there is enough to force the water out of some of its minerals (especially the sheet-silicate mineral serpentine). This water rises into the overlying mantle where it contributes to flux melting of the already hot ultramafic mantle rock. The mafic magma produced rises through the mantle to the base of the crust. There it contributes to partial melting of crustal rock, and that makes the magma more felsic than it was to begin with. That magma continues to rise and to assimilate crustal material, and in the upper part of the crust it accumulates within magma chambers. From time to time some of the magma from the plutons is forced up to surface, leading to volcanic eruptions. Washington State’s Mt. St. Helens, which last erupted in the 1980s, is an example of subduction-related volcanism.

A mantle plume is an ascending column of hot mantle rock (not magma) that originates deep in the mantle, possibly just above the core-mantle boundary. Mantle plumes are thought to rise at between 5 and 10 times the rate of mantle convection. The ascending column may be in the order of tens of kilometres to over a hundred kilometres across, but near to the surface it spreads out to create a mushroom-style head (Figure 7.1.2c). Near to the base of the lithosphere (the rigid part of the mantle) the mantle plume (and possibly some of the surrounding mantle material) partially melts to form mafic magma that rises to feed volcanoes. Since most mantle plumes are situated beneath the oceans, the early stages of volcanism typically take place on the sea floor. Over time islands may form like those in Hawaii.

Not all volcanic regions fit nicely into these three categories, and one of these exceptions is the volcanism in northwestern British Columbia (Figure 7.1.3). This area is not at a divergent or convergent boundary, and there is no evidence of an underlying mantle plume. The prevailing theory is that the crust of northwestern BC is being stressed by the northward movement of the Pacific Plate against the North America Plate, and that the resulting crustal fracturing provides a conduit for the flow of magma from the asthenospheric mantle.

Figure 7.1.3 (Left) Volcanoes and Volcanic Fields in the Northern Cordillera Volcanic Province, BC.;  and (Right) Volcanic Rock at the Tseax River Area (now Ksi Sii Aks), Northwestern BC

Media Attributions

• Figure 7.1.1 Steven Earle, CC BY 4.0, after a Public Domain drawing, Dynamic Planet, section by J. Vigil, from US Geological Survey, http://pubs.usgs.gov/gip/dynamic/Vigil.html
• Figure 7.1.2 Steven Earle, CC BY 4.0, after a Public Domain drawing, Dynamic Planet, section by J. Vigil, from US Geological Survey, http://pubs.usgs.gov/gip/dynamic/Vigil.html
• Figure 7.1.3 Map modified by Steven Earle, CC BY 4.0;  from base map, South-West Canada, public domain by Qyd/US Geological Survey, ESRI GIS data via Wikimedia Commons, http://commons.wikimedia.org/wiki/File:South-West_Canada.jpg; Volcanic locations (data) from Edwards & Russell (2000); Photo on right by Steven Earle, CC BY 4.0

As noted above, the types of magma produced in the differing volcanic settings can differ quite significantly. At divergent boundaries and oceanic mantle plumes, where there is little interaction with crustal materials, the magma tends to be consistently mafic. At subduction zones, where the magma ascends through significant thicknesses of crust, interaction between the magma and the crustal rock—some of which is quite felsic—results in magma that is relatively felsic.

Figure 7.2.1 The Important Processes that Lead to Changes in the Composition of Magmas Stored Within Magma Chambers Within Relatively Felsic Rocks of the Crust.

As shown on Figure 7.2.1, there are several processes that can make magma that is stored in a chamber within the crust more felsic, and can also contribute to development of vertical zonation from more mafic at the bottom to more felsic at the top. Partial melting of country rock and of country-rock xenoliths increases the overall felsic character of the magma, first because the country rocks tends to be more felsic than the magma, and second because the more silica-rich minerals (e.g., feldspar) within any rock tend to melt at a lower temperature than the silica-poor ones (e.g., amphibole). Settling of ferromagnesian crystals from the upper part of the magma, and re-melting of those crystals in the lower part can both contribute to the vertical zonation from relatively mafic at the bottom to more felsic at the top.

From the perspective of volcanism there are some important differences between felsic and mafic magmas. First, as already discussed, felsic magmas tend to be more viscous because they have more silica, and hence more polymerization. Second, felsic magmas tend to have higher levels of volatiles—that is components that behave as gases during volcanic eruptions. The most abundant volatile in magma is water (H₂O), followed, typically, by carbon dioxide (CO₂) and then by sulphur dioxide (SO₂). The general relationship between the SiO₂ content of magma and the amount of volatiles is shown on Figure 7.2.2. Although there are many exceptions to this trend, mafic magmas typically have 1 to 3% volatiles, intermediate magmas have 3 to 4% volatiles and felsic magmas have 4 to 7% volatiles.

Figure 7.2.2 Variations in the Volatile Compositions of Magmas as a Function of Silica Content

Differences in viscosity and volatile level have significant implications for the nature of volcanic eruptions. When magma is deep beneath the surface and under high pressure from the surrounding rocks, the gases remain dissolved.

As magma approaches the surface the pressure exerted on it decreases. Gas bubbles start to form, and the more gas there is in the magma the more bubbles will form. If the gas content is low or the magma is runny enough for gases to rise up through it and escape to surface, the pressure will not become excessive. Assuming there is a way for it to get to surface, the magma will flow out relatively gently. An eruption that involves a steady non-violent flow of magma is called effusive.

Exercise 7.2 Under Pressure!

A good analogy for a magma chamber in the upper crust is a plastic bottle of soda pop.

Figure 7.2.3 Champagne Uncorking.

Go to a supermarket and pick one up off the shelf (something not too dark). You’ll find that the bottle is hard because it was bottled under pressure, and you should be able to see that there are no gas bubbles inside.

Buy a small bottle of pop in a plastic bottle (you don’t have to drink it!) and open the lid. The bottle will become soft because the pressure is released, and small bubbles will start forming. If you put the lid back on and shake the bottle (best to do this outside!) you’ll enhance the processes of bubble formation, and when you open the lid the pop will come gushing out, just like an explosive volcanic eruption.

A pop bottle is a better analogue for a volcano than the old baking soda and vinegar experiment that you did at elementary school, because pop bottles—like volcanoes—come pre-charged with gas pressure. All we need to do is release the confining pressure and the gases come bubbling out, bringing some of the soda-pop along for the ride.

Champagne works equally well for this experiment, but doesn’t typically come in plastic bottles!

If the magma is felsic, and therefore too viscous for gases to escape easily, or if it has a particularly high gas content, it is likely to be under high pressure. Viscous magma doesn’t flow easily, so even if there is a way for it to get out, it may not be able to flow out readily. Under these circumstances pressure will continue to build as more magma moves up from beneath. Eventually some part of the volcano will break and then that pent up pressure will lead to an explosive eruption. More on that later.

Mantle plume and spreading-ridge magmas tend to be consistently mafic and so effusive eruptions are the norm. At subduction zones the average magma composition is likely to be close to intermediate, but, as we’ve seen, magma chambers can become zoned and so compositions ranging from felsic to mafic are possible, and different eruptions can have very different magma compositions. Eruption styles can be correspondingly variable.

Media Attributions

• Figure 7.2.1 Steven Earle, CC BY 4.0
• Figure 7.2.2 Steven Earle, CC BY 4.0, after Schmincke, H-U. (2004). Volcanism. Springer-Verlag, Heidelberg. https://link.springer.com/content/pdf/10.1007%2F978-3-642-18952-4.pdf
• Figure 7.2.3 Champagne Uncorking by Niells Noordhoek, 2012, CC BY SA 3.0,  via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Champagne_uncorking_photographed_with_a_high_speed_air-gap_flash.jpg

There are numerous types of volcanism; some of the more common ones are summarized in Table 7.3.1.

The sizes and shapes of typical shield, composite and cinder-cone volcanoes are compared on Figure 7.3.1, although, to be fair, Mauna Loa is the largest shield volcano on Earth, all others are smaller. Mauna Loa rises from the surrounding flat sea floor, and its full diameter is in the order of 200 km, with a diameter of about 100 km above sea level. Its elevation is 4169 m above sea level. Mt. St. Helens, a composite volcano, rises above the surrounding hills of the Cascade Range. It is about 6 km across at the base, and its height is 2550 m above sea level. Cinder cones are much smaller. On this drawing even a large cinder cone is just a dot.

Figure 7.3.1 Profiles of a Shield Volcano (Mauna Loa and Kilauea), a Composite Volcano (Mt. St. Helens), and a Large Cinder Cone

Cinder Cones

Cinder cones, like Eve Cone in northern BC (Figure 7.3.2), are typically only a few hundred metres in diameter and few are more than 200 m high. Most are comprised of fragments of vesicular mafic volcanic rock that were blasted out during a high-gas-pressure early phase of an eruption that may have subsequently become effusive (lava flows). Most cinder cones are monogenetic, meaning that they were created during a single eruptive phase that might have lasted weeks or months. Because cinder cones are made up almost exclusively of loose fragments, they have very little strength and can be easily, and relatively quickly, eroded away.

Figure 7.3.2 Eve Cone, Which Rises About 170 m Above the Surrounding Plateau, Formed Approximately 700 Years Ago.

Composite Volcanoes

Composite volcanoes, like Mt Merapi in Java (Figure 7.0.1) or Mt. St. Helens in Washington State (Figure 7.3.3), are almost all associated with subduction at convergent plate boundaries—either ocean-continent or ocean-ocean boundaries (Figure 7.1.2b). At many such volcanoes magma is stored in a magma chamber in the upper part of the crust. For example, at Mt. St. Helens, there is evidence of a magma chamber that is approximately 1 kilometre in width and extends from about 6 to 14 km depth below surface (Figure 7.3.4). Systematic variations in the composition of volcanism over the past several thousand years at Mt. St. Helens imply that the magma chamber is zoned, from more felsic at the top to more mafic at the bottom.

Figure 7.3.3 The North Side of Mt. St. Helens in South-Western Washington State, 2003. The large 1980 eruption reduced the height of the volcano by 400 m, and a sector collapse removed a large part of the northern flank. Between 1980 and 1986 the slow eruption of more lava led to construction of a dome inside the crater.

Figure 7.3.4 A Cross-Section Through the Upper Part of the Crust at Mt. St. Helens Showing the Zoned Magma Chamber.

The rock that makes up Mt. St. Helens ranges in composition from rhyolite (Figure 7.3.5a) to basalt (Figure 7.3.5b), and that implies that the types of past eruptions have varied widely in their character. As already noted, felsic magma doesn’t flow easily and doesn’t allow gases to escape easily. Under these circumstances pressure builds up until some part of the volcano gives way, and then an explosive eruption results, producing pyroclastic debris, as shown on Figure 7.3.5a. This type of eruption can also lead to rapid melting of ice and snow on a volcano, and that typically triggers large mudflows known as lahars (Figure 7.3.5a). Hot, fast moving pyroclastic flows and lahars are the two main causes of casualties in volcanic eruptions. Pyroclastic flows killed approximately 30,000 during the 1902 eruption of Mt. Pelée on the Caribbean island of Martinique. Most were incinerated in their homes. In 1985 a massive lahar, triggered by the eruption of Nevado del Ruiz, killed 23,000 in the Columbian town of Armero, about 50 km from the volcano.

In contrast, mafic eruptions (and some intermediate eruptions), produce lava flows and the one shown on Figure 7.3.5b is thick enough (about 10 m in total) to have cooled in a columnar jointing pattern (Figure 7.3.6). Lava flows serve to both flatten the profile of the volcano (because the lava typically flows farther than the pyroclastic debris falls) and also to protect it from erosion. Even so, composite volcanoes tend to erode quite quickly. Patrick Pringle, a volcanologist formerly with the Washington State Department of Natural Resources describes Mt. St. Helens as a “pile of junk”.

In a geological context, composite volcanoes tend to form relatively quickly and do not last very long. Mt. St. Helens, for example, is made up of rock that is all younger than 40,000 years; most of it is younger than 3,000 years. If its volcanic activity ceases, then it might erode away within a few tens of thousands of years. This is largely because of the presence of pyroclastic eruptive material, which is quite weak because it is made up of fragments that are not well stuck together.

Figure 7.3.5 Mt. St. Helens Volcanic Deposits. (a) lahar deposits (L) and felsic pyroclastic deposits (P); and (b) a columnar basalt lava flow. The two photos were taken at locations only about 500 m apart. 

Figure 7.3.6 The Development of Columnar Jointing in Basalt, Here Seen from the Top Looking Down. As the rock cools it shrinks, and because it is very homogenous it shrinks in a systematic way. When the rock breaks it does so with approximately 120˚ angles between the fracture planes. The resulting columns tend to be 6-sided but 5- and 7-sided columns also form.

Exercise 7.3 Volcanoes and Subduction

Figure 7.3.7 Interactions Between the North America, Juan de Fuca and Pacific plates Off the West Coast of Canada and the US

Figure 7.3.7 illustrates the interactions between the North America, Juan de Fuca and Pacific plates off the west coast of Canada and the US. The Juan de Fuca plate is being formed along the Juan de Fuca ridge, and is then subducted beneath the North America plate along the red line with teeth on it (“Cascadia subduction boundary”)

1. Using the scale bar in the lower left, estimate the average distance between the subduction boundary and the Cascadia composite volcanoes. (Compare that result with the distance between the subduction boundary and the volcanoes shown on Figure 7.0.2 above.)
2. If the subducting Juan de Fuca plate descends 40 km for every 100 km that it moves inland, what is its likely depth of the subducting plate in the area directly beneath the existing volcanoes?

Exercise answers are provided Appendix 2.

Shield Volcanoes

Most shield volcanoes are associated with mantle plumes, although some form at divergent boundaries, either on land or on the sea floor. The best-known shield volcanoes are those that make up the Hawaiian Islands, and of these the only active ones are on the big island of Hawaii. Mauna Loa, the world’s largest volcano and the world’s largest mountain (by volume) last erupted in 1984. Kilauea, arguably the world’s most active volcano, erupted almost continuously, from 1983 to 2018, and then started up again in late 2020. Loihi is an underwater volcano on the southeastern side of Hawaii. It is last known to have erupted in 1996, but may have erupted since then without being detected.

All of the Hawaiian volcanoes are related to the mantle plume that currently lies beneath Mauna Loa, Kilauea and Loihi (Figure 7.3.8). In this area the Pacific Plate is moving northwest at a rate of about 7 cm/year, and this means that the earlier formed—and now extinct—volcanoes have now moved well away from the mantle plume. As shown on Figure 7.3.8, there is evidence of crustal magma chambers beneath all three active Hawaiian volcanoes. At Kilauea the magma chamber appears to be several kilometres in diameter and is situated between 8 and 11 km below surface.

Figure 7.3.8 A Cross-Section Through the Crust and Upper Mantle in the Area of the Hawaii Mantle Plume

Although it is not a prominent mountain (seen in Figure 7.3.1), there is a large caldera in the summit area of Kilauea volcano (Figure 7.3.9). A caldera is a volcanic crater that is more than 2 km in diameter; this one is 4 km long and 3 km wide. It contains a smaller feature called Halema’uma’u crater that has a total depth of over 200 m below the surrounding area. Most volcanic craters and calderas are formed above magma chambers, and the level of the crater floor is influenced by the amount of pressure exerted by the magma body. During historical times the floors of both Kilauea caldera and Halema’uma’u crater have moved up—during expansion of the magma chamber—and down—during deflation of the chamber.

Figure 7.3.9 Aerial View of the Kilauea Caldera in 2012. The caldera is about 4 km across, and up to 120 m deep. It encloses the smaller and deeper crater Halema’uma’u Crater

One of the conspicuous features of Kilauea caldera is the sight of rising water vapour (the white cloud in Figure 7.3.9) and a strong smell of sulphur (Figure 7.3.10). As is typical in magmatic regions, water is the main volatile component, followed by carbon dioxide and sulphur dioxide. These, and some minor gases, originate from the magma chamber at depth and rise up through cracks in the overlying rock. This degassing of the magma is critical to the style of eruption at Kilauea, which, for most of the past 30 years, has been effusive, not explosive.

Figure 7.3.10 A Gas-Composition Monitoring Station (left) Within the Kilauea Caldera and at the Edge of Halema’uma’u Crater. The rising clouds are mostly comprised of water vapour, but also include carbon dioxide and sulphur dioxide. Sulphur crystals (right) that have formed around a gas vent in the caldera.

Kilauea started forming at approximately 300 ka, while neighbouring Mauna Loa began to form at about 700 ka and Mauna Kea at about 1 Ma. If volcanism continues above the Hawaii mantle plume in the same manner that it has since 85 Ma, it is likely that Kilauea will continue to erupt for at least another 500,000 years. By that time its neighbour, Loihi, will likely have emerged from the sea floor, and its other neighbours, Mauna Loa and Mauna Kea, will have become significantly eroded.

Exercise 7.4 Kilauea’s June 2015 Lava Flow

The U.S. Geological Survey Hawaii Volcano Observatory (HVO) map below (Figure 7.3.11), shows the outline of lava that started flowing northeast from Pu’u’o’o on June 27th 2015 (the “June 27th Lava flow”, a.k.a. the “East Rift Lava Flow”). The flow reached the nearest settlement, Pahoa, on October 29th, after covering a distance of 20 km in 124 days. After damaging some infrastructure west of Pahoa, the flow stopped advancing. A new outbreak formed November 1st, branching out to the north from the main flow about 6 km southwest of Pahoa.

What is the average rate of advance of the flow front from June 27th to October 29th, 2015 – in m/day and m/hour?

Figure 7.3.11 Small-Scale Map Showing Kīlauea’s Active East Rift Zone Lava Flow, Flowing Northeast from Pu’u’o’o on June 27th 2015.

Exercise answers are provided Appendix 2.

Large Igneous Provinces

While the Hawaii mantle plume has produced magma at a relatively slow rate for a very long time (at least 85 million years), other mantle plumes are less consistent, and some generate massive volumes of magma over relatively short time periods. Although their origin is still controversial, it is thought that the volcanism leading to large igneous provinces (LIP) is related to very high volume but relatively short duration bursts of magma from mantle plumes. An example of an LIP is the Columbia River Basalt Group (CRGB), which extends across Washington, Oregon and Idaho (Figure 7.3.12). This volcanism, which covered an area of about 160,000 km² with basaltic rock up to several hundred metres thick, took place between 17 and 14 Ma.

Figure 7.3.12 A Part of the Columbia River Basalt Group at Frenchman Coulee, Eastern Washington. All of the flows visible here have formed large (up to 2 metres in diameter) columnar basalts. The inset map shows the extent of the 17 to 14 Ma Columbia River Basalts, with the approximate location of the photo shown as a star.

Some other LIP eruptions have been much bigger. The eruption of the Siberian Traps (also basalt), which happened at the end of the Permian period, at 250 Ma, is estimated to have been 40 times the volume of the CRBG, and is thought to have been responsible for the greatest extinction of all time.

The mantle plume that is assumed to be responsible for the CRBG is now situated beneath the Yellowstone area in Wyoming, where it is associated with felsic volcanism. Over the past 2 million years three very large explosive eruptions at Yellowstone have yielded approximately 900 km³ of felsic magma, about 900 times the volume of the 1980 eruption of Mt. St. Helens, but only 5% of the volume of mafic magma in the CRBG.

Sea Floor Volcanism

Some LIP eruptions occur on the sea floor, the largest being the one that created the Ontong Java plateau in the western Pacific Ocean at around 122 Ma. But most seafloor volcanism originates at divergent boundaries and involves relatively low volume eruptions. Under these conditions, hot lava that oozes out in the cold seawater quickly cools on the outside and then behaves a little like toothpaste. The resulting blobs of lava are known as pillows, and they tend to form piles around a sea-floor lava vent (Figure 7.3.13). In terms of area, there is very likely more pillow basalt on the sea floor than any other type of rock on Earth.

Figure 7.3.13 Modern Sea-Floor Pillows in the South Pacific. (Left: public domain image by NOAA, 1988, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Pillow_basalt_crop_l.jpg; and Right: 40 to 50 Ma pillows on the shore of Vancouver Island, near to Sooke. The pillows are 30 to 40 cm in diameter.

Kimberlites

While all of the volcanism discussed so far is thought to originate from partial melting in the upper mantle or within the crust, there is a special class of volcanoes—kimberlites—that have their origins much deeper in the mantle, at depths of 150 to 450 km. During a kimberlite eruption material from this depth may make its way to surface quite quickly (hours to days) and with little interaction with the surrounding rocks. As a result, kimberlite eruptive material is representative of mantle compositions—it is ultramafic.

The pressure and temperature suitable for diamonds to form exist in the mantle at depths of 160 to 190 km within areas of old thick crust (shields). Kimberlite eruptions that originate at greater depth traverse this region of diamond stability, and, in some cases bring diamond-bearing rock to the surface. All of the diamond deposits on Earth are assumed to have formed in this way; an example is the rich Ekati Mine in the Northwest Territories (Figure 7.3.14).

Figure 7.3.14 Ekati Diamond Mine, Northwest Territories, Part of the Lac de Gras Kimberlite Field. There are at least 3 kimberlite bodies in the area of this view, represented by the three open-pit mines.

The kimberlites at Ekati erupted between 45 and 60 Ma. Many kimberlites are older— some much older—but there have been no kimberlite eruptions in historic times. The youngest known kimberlites are at the Igwisi Hills in Tanzania, aged about 10,000 years, and the next youngest known are dated to about 30 Ma.

Media Attributions

• Figure 7.3.1 Steven Earle, CC BY 4.0
• Figure 7.3.2 Eve Cone by FortGirl, 2007, via Flickr, CC BY SA NC 2.0, https://www.flickr.com/photos/fortgirl/2184491354/in/photostream/
• Figure 7.3.3 Photo by Steven Earle, CC BY 4.0
• Figure 7.3.4 Steven Earle, CC BY 4.0, after Pringle, P. T., & Washington (State). Division of Geology and Earth Resources. (1993). Roadside geology of Mount St. Helens National Volcanic Monument and vicinity. Information circular/Washington Department of Natural Resources, Division of Geology and Earth Resources, 88. https://www.dnr.wa.gov/Publications/ger_ic88_mount_st_helens_pt1.pdf
• Figure 7.3.5 Photos by Steven Earle, CC BY 4.0
• Figure 7.3.6 Steven Earle, CC BY 4.0
• Figure 7.3.7 Steven Earle, CC BY 4.0
• Figure 7.3.8 Hotspot Cross-Sectional Diagram by J. E. Robinson, (2006). US Geological Survey public domain image via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Hawaii_hotspot_cross-sectional_diagram.jpg)
• Figure 7.3.9 Kilauea, NASA, public domain, via Wikimedia Commons,  https://commons.wikimedia.org/wiki/File:Kilauea_ali_2012_01_28.jpg
• Figure 7.3.10 Photos by Steven Earle, CC BY 4.0
• Figure 7.3.11 Map from US Geological Survey & Hawaiian Volcano Observatory, Public domain, https://www.usgs.gov/observatories/hvo
• Figure 7.3.12 Photo and inset drawing by Steven Earle, CC BY 4.0
• Figure 7.3.13 Photo by Steven Earle, CC BY 4.0
• Figure 7.3.14 Ekati mine by Jason Pineau, 2010, CC BY SA 3.0, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Ekati_mine_640px.jpg

There are two classes of volcanic hazards, direct and indirect. Direct hazards are those that can directly kill or injure people, or destroy property or wildlife habitat. Indirect hazards are volcanism-induced environmental changes that lead to distress, famine or habitat destruction. Indirect hazards of volcanism have accounted for many times more deaths during historical times than direct hazards. Some of the more important types of volcanic hazards are summarized in Table 7.4.1.

Volcanic Gas and Tephra Emissions

Large volumes of tephra (rock fragments, mostly pumice, and volcanic ash) and gases are emitted during major explosive eruptions at composite volcanoes, and a large volume of gas is also released during some high volume effusive eruptions. One of the major implications of these emissions is cooling of the climate by up to 1˚ C for several months to a few years because the dust particles and tiny droplets and particles of sulphur compounds block the sun. The last time this happened was in 1991 and 1992 following the large eruption of Mt. Pinatubo in the Philippines. 1˚ C may not seem like a lot, but that was the global average amount of cooling, and cooling was more severe in some regions and at some times.
Over an 8-month period in 1783 and 1784 a large effusive eruption took place at the Laki volcano in Iceland. Although there was relatively little volcanic ash involved, a massive amount of sulphur dioxide was released into the atmosphere, and also a significant volume of hydrofluoric acid (HF). The sulphate aerosols that formed in the atmosphere led to dramatic cooling in the northern hemisphere. In Iceland poisoning from the HF resulted in the death of 80% of sheep, 50% of cattle, and the ensuing famine, along with HF poisoning, resulted in the over 10,000 human deaths, about 25% of the population. The Laki eruption also resulted in many deaths in Europe, although the total number isn’t known, it is estimated that there were approximately 20,000 deaths in the United Kingdom because of very cold weather, and it seems likely that other parts of northern Europe would have been similarly affected.

Volcanic ash can also have serious implications for aircraft because it can destroy jet engines. In 2010 the eruption of Iceland’s Eyjafjallajökull volcano led to the closure of the European airspace for several days, and the cancellation of numerous trans-Atlantic flights.

Not all of the environmental effects of volcanism are related to eruptions of magma. The craters of dormant volcanoes are commonly filled with water (such as Crater Lake in Oregon). Within Lake Nyos in west central Cameroun, gases emanating from the underlying magma chamber continually percolate upward into the muddy lake sediment and bottom waters. One August night in 1986 a landslide, an earthquake or a minor eruption disturbed the lake sediment and the water and released approximately 100 million cubic metres of carbon dioxide from the lake bottom. The CO₂ quickly bubbled up through the water and out into the air above the lake. The gas spilled over the lip of the crater and descended in a white cloud down into the valleys surrounding the crater. Over 1700 people and 3000 cattle were killed in their sleep.

There are other lakes in Africa with similar conditions, including nearby Lake Monoun, and the much larger Lake Kivu on the Congo-Rwanda border. An operation to reduce the risk of a future CO₂ eruption from Lake Nyos is currently underway. The procedure involves lowering a strong polyethylene pipe to the lake bottom. Some water is pumped out at the top, and as the deep water rises through the pipe the carbon dioxide starts to bubble out. The gas and water then become buoyant and suck more water in at the bottom in a self-sustaining process (Figure 7.4.1).

Figure 7.4.1 The Degassing Operation on Lake Nyos, Cameroun

Pyroclastic Density Currents

In a typical explosive eruption at a composite volcano the tephra and gases are hot enough to be buoyant in the air and they are forced high up into the atmosphere. As the eruption proceeds, and the ejected materials start to cool, parts become heavier than air and they can then flow downward along the flanks of the volcano (Figure 7.4.2). As they descend, they cool more and so flow faster, reaching speeds up to several hundred km/h. A pyroclastic density current (PDC) consists of tephra ranging in size from microscopic shards of glass to boulders, plus gases (dominated by water vapour, but also including other gases). The temperature of this material can be as high as 1000˚ C. The most famous PDCs are the one that destroyed Pompeii in the year 79 AD, killing an estimated 18,000 and the one that destroyed the town of St. Pierre, Martinique in 1902, killing an estimated 30,000.

Pyroclastic density currents can flow over water, in some cases for tens of kilometres. The 1902 the St. Pierre PDC flowed out into the harbour and destroyed (burned) several wooden ships anchored there.

Figure 7.4.2 The Plinian Eruption of Mt. Mayon, Philippines. in 1984. Although most of the eruption column is ascending into the atmosphere, there are pyroclastic density currents flowing down the sides of the volcano in several places.

Pyroclastic Falls

Most of the tephra from an explosive eruption ascends high into the atmosphere, and some of it is distributed around the Earth by high altitude winds. The larger components (larger than 0.1 mm) tend to fall relatively close to the volcano, and the amount produced by large eruptions can cause serious damage and casualties. The large 1991 eruption of Mt. Pinatubo in the Philippines resulted in the accumulation of tens of centimetres of ash in fields and on rooftops in the surrounding populated region. Heavy typhoon rains that hit the island at the same time added to the weight of the tephra and led to the collapse of thousands of roofs and to at least 300 of the 700 deaths attributed to the eruption .

Lahars

A lahar is any mudflow or debris flow that is related to a volcano. Most are caused by melting snow and ice during an eruption, as was the case with the lahar that destroyed the Columbian city of Armero in 1985 (Figure 7.4.3). Lahars can also happen when there is no volcanic eruption, and one of the reasons is that, as we’ve seen, composite volcanoes tend to be weak and easily eroded.

Figure 7.4.3 Part of the City of Armero Following the 1985 Lahar.

In October 1998 category 5 hurricane Mitch slammed into the coast of central America. Damage was extensive and 19,000 people died, not so much because of high winds but because of intense rainfall—some regions received almost 2 m of rain over a few days! Lahars (mudflows and debris flows) occurred in many areas, especially in Honduras and Nicaragua. An example is Casita Volcano in Nicaragua, where the heavy rains weakened rock and volcanic debris on the upper slopes, resulting in a debris flow that rapidly built in volume as it raced down the steep slope, and then ripped through the towns of El Porvenir and Rolando Rodriguez killing more than 2,000 people (Figure 7.4.4). El Porvenir and Rolando Rodriguez were new towns that had been built without planning approval in an area that was known to be at risk from lahars.

Figure 7.4.4 Part of the Path of the Lahar from Casita Volcano, October 30th, 1998

Central America was struck again by major tropical storms in November 2020. Hurricanes Eta (category 4) and Iota (category 5) also produced mudflows and lahars. The combined death toll in Central America was close to 200.

For several reasons, there are significant lahar risks from Washington State’s Mt. Rainier:

• It is the tallest and largest of the Cascade Range volcanoes,
• It has numerous large glaciers and accumulates a great deal of snow in winter,
• It has a significant volume of rock that has been weakened by alteration,
• It has been active within the past 125 years, and
• There are several large communities close to the mountain and within the channels of known past lahars (Figure 7.4.5).

According to a US Geological Survey (USGS) publication, Mt. Rainier is a lahar risk because there is a likelihood of melting snow and ice during an eruption, and also because there is a risk of slope failure at any time, but especially if there is a large earthquake in the region, or if there is movement of magma within the volcano itself. The USGS, Pierce County Department of Emergency Management, and Washington State Emergency Management Division have established an array of motion sensors in the drainage channels around Mt. Rainier. The system is designed to detect the vibrations associated with a lahar and then to alert at risk residents so that they have time to get to higher ground.

Figure 7.4.5 The Debris Flow and Lahar Risk Around Mt. Rainier in West-Central Washington.

Sector Collapse and Debris Avalanches

In the context of volcanoes, sector collapse or flank collapse is the catastrophic failure of part of an existing volcano, and the formation of a large debris avalanche. The best known example of sector collapse is the failure of the northern side of Mt. St. Helens immediately prior to the large eruption on May 18, 1980. In the weeks leading up to the eruption a large bulge had formed on the side of the volcano, likely the result of transfer of magma from depth into a satellite magma chamber within the mountain itself. Early on the morning of May 18 a moderate earthquake struck nearby and this is thought to have destabilized the bulge, leading to the Earth’s largest slope failure in historical times. The failure of this part of the volcano exposed the underlying satellite magma chamber, causing it to explode sideways, and that exposed the conduit leading to the magma chamber below. The resulting eruption—with a 20 km high eruption column—lasted for 9 hours. Approximately 1 cubic km of tephra (volcanic rock fragments and ash) erupted, making this a relatively small eruption.

In August 2010 a large part of the side of BC’s Mt. Meager gave way, sending about 48 million cubic metres of rock down the valley, the largest slope failure in Canada in historical times (although the volume was only about 1/60th of the Mt. St. Helens 1980 sector collapse). This represents only a tiny fraction of the volume of Mt. Meager, but it can be considered a partial sector collapse. More than 25 slope failures have taken place at Mt. Meager in the past 8,000 years, some of them more than 10 times the volume of the 2010 failure.

Lava Flows

As we’ve seen in Exercise 4.4, lava flows at volcanoes like Kilauea do not advance very quickly, and in most cases, people can get out of the way. Of course, it is more difficult to move infrastructure, and so buildings and roads are typically the main casualties of lava flows.

That was the case with the massive Kilauea eruption in 2018 where lava flowed through a mostly rural area (Figure 7.4.6), resulting in 24 injuries, no deaths, $800 million in damages, and the loss of over 800 homes as a well as a geothermal energy plant. But the situation at Mount Nyiragongo in the Republic of the Congo in 2002 was quite different. There a similar-sized lava stream flowed through Goma, a city of 200,000 people, destroying thousands of buildings, and leaving about 120,000 people homeless. A total of 245 people died, most from carbon-dioxide asphyxiation.

Figure 7.4.6 Kilauea’s Lower Puna Lava Flow on May 19th, 2018

The eruption of the Tseax Cone in Northwestern BC, sometime between 1668 and 1714 resulted in a lava flow that now covers about 100 km² (Figure 7.1.4). The flow destroyed two Nisga’a villages and killed about 2000 people, again, mostly as a result of carbon dioxide asphyxiation.

Exercise 7.5 Volcanic Hazards in Squamish

The town of Squamish, British Columbia, is situated approximately 10 km from Mt. Garibaldi, as shown on Figure 7.4.7. In the event of a major eruption of Garibaldi, which of the following hazards has the potential to be an issue for the residents of Squamish, or for those passing through on Highway 99?

Figure 7.4.7 Distance from Squamish, BC to Mt. Garibaldi

Table 7.4.2 Hazards and Associated Risks to Squamish, BC from Mt. Garibaldi

Hazard	Is it a risk in this area (yes/no) and brief explanation
Tephra emission
Gas emission
Pyroclastic density current
Pyroclastic fall Lahar Lava flow
Sector collapse

Exercise answers are provided Appendix 2.

Media Attributions

• Figure 7.4.1 Exploding Lakes in Cameroon, by Bill Evans, US Geological Survey photo, public domain. https://www.usgs.gov/media/images/exploding-lakes-cameroon
• Figure 7.4.2 Pyroclastic Flows at Mayon Volcano, by C. Newhall, US Geological Survey, public domain, via Wikipedia, https://en.wikipedia.org/wiki/File:Pyroclastic_flows_at_Mayon_Volcano.jpg)
• Figure 7.4.3 Armero Mudflow and Ruins by N. Banks, US Geological Survey, Public domain, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Armero_Mudflow_and_ruins.jpg
• Figure 7.4.4 Lahar from Casita Volcano, US Geological Survey, Public domain, http://volcanoes.usgs.gov/hazards/lahar/casita.php
• Figure 7.4.5 Debris Flow and Lahar risks map, Public domain, from Driedger, C. & Scott, W. (2008). Mount Rainier: Living safely with a volcano in your backyard. US Geological Survey, Fact Sheet 2008-3062, p.3. https://doi.org/10.3133/fs20083062
• Figure 7.4.6 Kīlauea Volcano’s Lower East Rift Zone from US Geological Survey, 2018, public domain, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:USGS_K%C4%ABlauea_multimediaFile-2062.jpg
• Figure 7.4.7 Steven Earle, CC BY 4.0, based on a an image from Google Earth

In 2005, geologist Chris Newhall, US Geological Survey, made a list of the six most important signs of an imminent volcanic eruption. They are as follows:

1. Gas leaks: the release of gases (mostly H₂O, CO₂ and SO₂) from the magma, through cracks in the overlying rock and into the atmosphere.
2. Bit of a bulge: the deformation of part of a volcano, indicating that a magma chamber at depth is swelling or getting more pressurized.
3. Getting shaky: many (hundreds to thousands) of small earthquakes, indicating that magma is on the move. The quakes may be the result of the magma forcing the surrounding rocks to crack, or a harmonic vibration that is evidence of magmatic fluids moving underground.
4. Dropping fast: a sudden decrease in the rate of seismicity may indicate that magma has stalled, and this could mean that something is about to give way.
5. Big bump: a pronounced bulge on the side of the volcano (like at Mt. St. Helens in 1980) may indicate that magma has moved close to surface.
6. Blowing off steam: steam eruptions (a.k.a. phreatic eruptions) happen when magma near to surface heats groundwater to the boiling point. The water eventually explodes, sending steam and fragments of the overlying rock far into the air.

With these signs in mind, we can make a list of the equipment we should have and the actions we should take to monitor a volcano and predict when it might erupt.

Assessing Seismicity

Perhaps the most effective, simplest and cheapest way to monitor a volcano is with seismometers. In an area with a volcano that has the potential to erupt seismometers can provide us with an early warning that something is changing beneath the volcano. If there is seismic evidence that a volcano is coming to life, then more seismometers should be placed in locations within a few tens of kilometres of the source of the activity so as to provide an ongoing picture of how things are changing. (Figure 7.5.1). This will allow geologists to determine the exact location and depth of the seismic activity so that they can see where the magma is moving.

Figure 7.5.1 A Seismometer Installed in a Vault Above Ground at Mount Baker, Washington.

Detecting Gases

Water vapour quickly turns into clouds of liquid water droplets and it is relatively easy to detect just by looking, but CO₂ and SO₂ are not so obvious. It’s important to be able to monitor changes in the composition of volcanic gases, and we need instruments to do that. Some can be done from a distance (from the ground or even from the air) using infrared devices, but to get more accurate data we need to actually sample the air and do chemical analysis. This can be achieved with instruments placed on the ground close to the source of the gases (see Figure 7.3.10 in Section 7.3), or by collecting samples of the air and analyzing them in a lab.

Measuring Deformation

There are two main ways to measure ground deformation at a volcano. One is known as a tiltmeter, which is a sensitive 3-directional level that can sense small changes in the tilt of the ground at a specific location. Another is through the use of GPS (global positioning satellite) technology. GPS is more effective than a tiltmeter because it provides information on how far the ground has actually moved – east-west, north-south and up-down (Figure 7.5.2), but either instrument can be used to assess deformation that might be related to the movement of magma beneath the surface, and so could be indicative of imminent eruptive activity.

Figure 7.5.2 A GPS Unit Installed at Hualalai Volcano, Hawaii. The dish-shaped antenna on the right is the GPS receiver. The antenna on the left is for communication with a base station.

By combining information from these types of sources and a thorough knowledge of how volcanoes work, along with careful observations made on the ground and from the air, geologists can get a good idea of the potential for a volcano to erupt in the near future (months to weeks, but not days). They can then make recommendations to authorities about the need for evacuations and restricting transportation corridors. Our ability to predict volcanic eruptions has increased dramatically in recent decades because of advances in our understanding of how volcanoes behave and in monitoring technology. Providing that careful work is done, the risk of a surprise eruption is now much lower than it used to be, and providing that public warnings are issued and heeded, it is less and less likely that thousands will die from sector collapse, pyroclastic flows, ash falls, or lahars. Indirect hazards are still very real, however, and we can expect the next eruption, similar to the one at Laki in 1783, to take an even greater toll than it did then—especially since there are now roughly eight times as many people on the Earth.

Media Attributions

• Figure 7.5.1 Mount Baker, US Geological Survey, public domain, https://www.usgs.gov/volcanoes/mount-baker/earthquake-monitoring-mount-baker
• Figure 7.5.2 Volcano Watch, US Geological Survey/ Hawaiian Volcano Observatory, public domain, https://www.usgs.gov/center-news/volcano-watch-volcano-monitoring-lower-puna-recent-vandalism-and-a-way-help

Humans have a love-hate relationship with volcanoes. For many reasons humans are attracted to areas with active volcanism, but for several others that we’ve already discussed, they would be wise to stay away.

The key reason that humans like living around potentially active volcanoes is that the soil tends to be fertile, and thus there is the potential to grow enough food to live. For example, some parts of the area around Mt. Merapi in Indonesia (Figure 7.0.1) can support subsistence populations of 8 to 10 people per hectare. In comparison, the typical farm in the United States can feed just under 1 person per hectare (US Farm Bureau).

Volcanic soil is good for a number of reasons. One is that volcanic ash and rock fragments are rich in volcanic glass and under weathering conditions glass breaks down quickly to clay minerals so that productive soil can form within 200 to 300 years in favorable climates. Another is that the clays that form from volcanic parent materials are effective at holding onto nutrients such as phosphorous. A third is that volcanic lava or tephra are typically quite rich in some important plant nutrients, such as magnesium and sulphur. Volcanic regions all over the world are know for their fertile soils. Some examples, apart from Indonesia, include the volcanic areas in Italy, much of northern New Zealand, Japan, Hawaii, parts of Africa, and much of the Caribbean.

Volcanoes are also valued for their scenic beauty and recreational opportunities. An example is the Mt. Garibaldi area of southwestern British Columbia (Figure 7.6.1), but there are hundreds of other scenic volcanoes around the world, some of which are immense tourist and hiker attractions (Figure 7.6.2). Many volcanoes are also venues for a wide range of winter sports, and for hot springs, spas and mudbaths. Volcanic regions are also an excellent source of geothermal heat for both electricity and district heating, and of hydroelectric energy from streams.

Figure 7.6.1 Looking From the North to Mt. Garibaldi (Background Left of Centre, Shrouded in Cloud) with Garibaldi Lake in the Foreground. The volcanic peak in the centre is Mt. Price and the dark flat–topped peak is The Table.

Figure 7.6.2 Hikers Near the 3776 m Summit of Japan’s Mt. Fuji

Many volcanoes are also venues for a wide range of winter sports, and for hot springs, spas and mudbaths. Volcanic regions are also an excellent source of geothermal heat for both electricity and district heating, and of hydroelectric energy from streams. Figure 7.6.3 provides an overview of some of the ways that humans interact with volcanoes, and some of the risks associated with living nearby.

Figure 7.6.3 Some of the Ways that Humans Interact with Volcanoes

Volcanism and Earth Systems

As already noted in Chapter 1 and Chapter 3, volcanic eruptions contribute to the Earth’s systems in important ways. For starters, it is widely believed that the water in the Earth’s oceans is at least partly derived from volcanism, and the Earth would not have much in the way of systems without water.

Some of the key roles of volcanic eruptions in Earth systems are as follows:

• Cycling solids (mostly silicates) from depth in the mantle and the crust to surface,
• Cycling volatiles (water and gases) from depth, and thereby influencing organisms and the climate,
• Ejecting both solids and volatiles high into the atmosphere,
• Cycling thermal energy from depth,
• Creating solid surfaces (e.g., islands) that will be colonized by organisms, and
• Creating sloped surfaces (mountains) that influence weather and climate patterns, and will be eroded and weathered.

All of these products subsequently contribute to other Earth system processes in myriad ways.

Media Attributions

• Figure 7.6.1 Photo by Isaac Earle, used with permission, CC BY 4.0
• Figure 7.6.2 Mt. Fuji Summit by Derek Mawhinney, public domain image via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Mt_Fuji_Summit.jpg
• Figure 7.6.3 Steven Earle, CC BY 4.0

The topics covered in this chapter can be summarized as follows:

The earliest known metal mines, in Bulgaria and Serbia and some neighbouring areas of southeastern Europe, date back about 7000 years, with the main metal of interest being copper. Copper was also mined and used in the Great Lakes region of North America at around the same time.

Today there are mines, for virtually every metal in the periodic table, on every continent except Antarctica (where mining is banned by treaty), and the value of the mined products are enormous—both to the companies that do the mining and the nations where the mining takes place, and to the people everywhere that use the metals in their lives every day. Iron represents, by far, the largest amount of metal mined. It is used mostly in construction (buildings, bridges, etc.), for railways, and for vehicles, although most manufactured products have some iron in them. Each year we mine and process 2.5 billion tonnes of iron, which is approximately the weight of all of the cars, trucks and buses currently in the entre world. Aluminum, which is used in electrical transmission lines, construction and aircraft, is second. Chromium is third because it is an important component of steel, while copper, fourth, is used widely for electrical components.

Figure 8.1.1 The Amounts of Some of the Important Metals Mined Each Year on Earth. Each sphere is proportional in volume to the weight of that metal mined. There is 39 times as much iron mined as aluminum, and 760,000 times as much iron as gold.

A metal deposit is a body of rock in which one or more metals has been concentrated to the point of being economically viable for recovery. Not all metal-bearing rock is useful as ore. The granitic rock of the crust has an aluminum (Al) content of about 8%, but most of that is within the mineral feldspar, and it is very difficult to separate that Al from the other elements. The main ore of Al is bauxite, in which the Al is present as Al-hydroxides. Bauxite is much more readily processed into aluminum metal than is granite.

Some background levels of important metals in average rocks are shown on Table 8.1, along with the typical grades necessary to make a viable deposit, and the corresponding concentration factors. Looking at copper, for example, we can see that while average rock has around 40 ppm (parts per million) of copper, a grade of around 10,000 ppm or 1% is necessary to make a viable copper deposit. In other words, copper ore has about 250 times as much copper as typical rock. For all of the other elements in the list the concentration factors are much higher. For gold it’s 2000 times and for silver it’s around 10,000 times.

The economic viability of any metal deposit depends on a wide range of factors including its grade, size, shape and depth below surface, proximity to infrastructure, the current price of the metal, ease of processing, the labour and environmental regulations in the area, and many others. One of the most important of these factors is the grade (the concentration of metal in the ore) and whether it is high enough to make it mineable at a profit. For example, a typical economic grade for copper is 1%, but that will vary widely depending on the status of the various factors listed above, including size, shape and depth of the deposit, ease of processing the ore, proximity to infrastructure, and the labour and environmental rules in the country where the orebody is located.

From Table 8.1 we can see that some significant concentration must take place to form a mineable deposit. This concentration may occur during the formation of the host rock, or after the rock is formed, through a number of different types of processes. There is a very wide variety of ore-forming processes, and there are hundreds of types of mineral deposits. The origins and characteristics of a few of them are described below.

Magmatic Deposits

A magmatic deposit is one in which the metal concentration takes place primarily at the same time as the formation and emplacement of the magma. Most of the nickel mined in the world, and much of the copper, comes from magmatic deposits such as those in Indonesia, Canada, southern Africa, Australia and Siberia (Figure 8.1.2). The magmas from which these deposits are formed are of mafic or ultra-mafic composition (they are derived from the mantle), and therefore they have relatively high nickel and copper contents to begin with (as much as 100 times that of normal rocks in the case of nickel). These elements can be further concentrated within the magma as a result of addition of sulphur from partial melting of the surrounding rocks. The heavy nickel and copper sulphide minerals are then concentrated further still by gravity segregation (i.e., crystal settling towards the bottom of the magma chamber). In some cases, there are significant concentrations of platinum-bearing minerals in deposits of this type.

Most of magmatic deposits around the world are Precambrian in age—probably because the mantle was significantly hotter at that time, and the necessary ultramafic magmas were more likely to exist close to the surface.

Figure 8.1.2 The Nickel Mine and Smelter at Thompson, Manitoba.

Volcanogenic Massive Sulphide Deposits

Much of the world’s copper, zinc, lead, silver and gold is mined from volcanogenic massive sulphide (VMS) deposits associated with submarine volcanism. There are similar deposits in many countries, with some of the larger ones in Spain and Portugal, in southern Africa, Canada, and Russia.

VMS deposits are formed from the water discharged at high-temperature (250 to 300˚ C) at ocean-floor vents, primarily in areas of subduction-zone volcanism. Although most VMS deposits are tens to hundreds of millions of years old, the environment of their formation is comparable to that of a modern-day black smoker (Figure 8.1.3) which form where hot metal- and sulphide-rich water issues from the sea floor. They are called massive-sulphide deposits because the sulphide minerals (including pyrite (FeS₂) , sphalerite (ZnS), chalcopyrite (CuFeS₂) and galena (PbS)) are generally present at very high concentrations (making up the majority of the rock in some cases). The metals and the sulphur are leached out of the sea-floor rocks by convecting groundwater driven by the volcanic heat, and then quickly precipitated where that hot water cools suddenly and changes chemically on entering the cold sea water, so the process of concentration of the metal is a sudden change in metal solubility where the hot water cools quickly at the sea-floor rock-water interface. The volcanic rock that hosts VMS deposits is formed in the same area and at the same general time as the accumulation of the ore minerals.

Figure 8.1.3 (Left) A Black Smoker on the Juan de Fuca Ridge Off the West Coast of Vancouver Island . (Right) A Model of the Formation of a Volcanogenic Massive Sulphide Deposit on the Sea Floor.

Porphyry Deposits

Porphyry deposits are the most important ores of copper and molybdenum in western North and South America, and in other areas of the Pacific Rim. Most porphyry deposits also include some gold, and in a few cases gold is the primary commodity.

A porphyry deposit forms around a cooling felsic stock (magma chamber) in the upper part of the crust. They are called “porphyry” because upper crustal stocks are typically porphyritic in texture, the result of a two-stage cooling process. Metal enrichment results from convection of groundwater related to the heat of the stock, and also from hot water expelled by the cooling magma (Figure 8.1.4). The host rocks, which commonly include the stock itself and the surrounding country rocks, are normally highly fractured and brecciated. During the ore-forming process some of the original minerals in these rocks get altered to potassium feldspar, biotite, epidote and various clay minerals. The important ore minerals include chalcopyrite (CuFeS₂), bornite (Cu₅FeS₄) and pyrite in copper porphyry deposits, or molybdenite (MoS₂) and pyrite in molybdenum porphyry deposits. Gold is present as minute flakes of native gold.

This type of environment (i.e., around and above an intrusive body) is also favourable for the formation of other types of deposits—particularly vein-type gold deposits (a.k.a. epithermal deposits). Many of the other gold deposits situated all along the western edge of both South and North America are of the vein type shown in Figure 8.1.4, and are related to nearby magma bodies.

Figure 8.1.4 A Model for the Formation of a Porphyry Deposit Around an Upper-Crustal Porphyritic Stock, and of Associated Vein Deposits.

Some porphyry deposits are huge, and if the ore is near to the surface they are most economically mined by open-pit methods (more on that below). The largest porphyry copper deposit in the world is at Chuquicamata in northern Chile (Figure 8.1.5). It has been in operation for over 100 years, has produced over 29 million tonnes of copper metal, and is expected to continue operation for at least another 50 years. The mine is owned by the government of Chile and supports a nearby city with a population of 160,000.

Figure 8.1.5 The Chuquicamata Open Pit Copper Mine in Northern Chile. The pit is 4.3 km long, 2.7 km wide and 1.1 km deep. Mining currently takes place underground, below the bottom of the pit.

Banded Iron Formation

Most of the world’s important iron deposits are of the banded iron formation type, and most of them formed during the initial oxygenation of the Earth’s atmosphere between 2400 and 1800 Ma. At that time abundant iron in dissolved form in the ocean (as Fe²⁺) became oxidized to its insoluble form (Fe³⁺) as the Earth’s atmosphere gradually became more oxygen-rich. The iron minerals accumulated on the sea floor, mostly as hematite and magnetite interbedded with chert (Figure 8.1.6). Unlike many other metals, which are economically viable at grades of around 1% or even much less, iron deposits are only viable if the grades are in the order of 50% iron.

Figure 8.1.6 Banded iron formation from an unknown location in North America on display at a museum in Germany . The rock is about 2 m across. The reddish bands are mostly hematite (Fe2O3) and the darker bands a mostly magnetite (Fe3O4). (by André Karwath , at Museum of Mineralogy and Geology, Dresden, Germany, Creative Commons Attribution-Share Alike 2.5 Generic, from https://commons.wikimedia.org/wiki/File:Black-band_ironstone_(aka).jpg)

Unconformity-Type Uranium Deposits

There are several different types of uranium deposits, but some of the largest and richest are those within the Athabasca Basin of northern Saskatchewan. They are called unconformity-type because they are all situated very close to the unconformity between the Proterozoic Athabasca Group sandstone and the much older Archean sedimentary, volcanic and intrusive igneous rock (Figure 8.1.7). The origin of unconformity-type U deposits is not perfectly understood, but it is thought that two particular features are important: the relative permeability of the Athabasca Group sandstone, and the presence of graphitic schist within the underlying Archean rocks. The permeability of the sandstone allowed groundwater to flow through it and leach out small amounts of U which stayed in solution in the oxidized form U⁶⁺. The graphite (C) created a reducing environment (non-oxidizing) that converted the U from U⁶⁺ to insoluble U⁴⁺, at which point it was precipitated as the mineral uraninite (UO₂).

Figure 8.1.7 Model of the Formation of Unconformity-Type Uranium Deposits of the Athabasca Basin, Saskatchewan

Cobalt Deposits

Cobalt is an important component of lithium-based batteries, and the recent surge in the demand for batteries for electric vehicles and other purposes has resulted in increased demand for cobalt. Most of the world’s cobalt currently comes from sedimentary-rock hosted deposits in Zambia and Congo, Africa, although there is a significant amount that is mined along with nickel in magmatic and other deposits, especially those in Australia and Canada.

Media Attributions

• Figure 8.1.1 Steven Earle, CC BY 4.0, based on data current in 2020 from Visual Capitalist, https://www.visualcapitalist.com/
• Figure 8.1.2 Vale Nickel Mine by Timkal, 2008, via Wikimedia Commons,  https://commons.wikimedia.org/wiki/File:Vale_Nickel_Mine.JPG CC BY-SA 3.0
• Figure 8.1.3 (Left) Black Smoker by Butterfield, D. and Holden, J., National Oceanographic and Atmospheric Administration, public domain,  http://oceanexplorer.noaa.gov/okeanos/explorations/10index/background/plumes/media/black_smoker.html; (Right) Diagram by Steven Earle, CC BY 4.0)
• Figure 8.1.4 Steven Earle, CC BY 4.0
• Figure 8.1.5 Panoramic view of Chuquicamata by Diego Delso, delso, CC-BY-SA 2.0, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Mina_de_Chuquicamata,_Calama,_Chile,_2016-02-01,_DD_110-112_PAN.JPG
• Figure 8.1.6 Black-band Ironstone by André Karwath, Museum of Mineralogy and Geology, Dresden, Germany, via Wikimedia Commons, CC BY-SA 2.5, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Black-band_ironstone_(aka).jpg
• Figure 8.1.7 Steven Earle, CC BY 4.0

Metal deposits are mined in a variety of different ways depending on their depth, shape, size and grade.  Relatively large deposits that are quite close to surface and somewhat regular in shape are mined using open-pit mine methods (Figure 8.1.5 in Section 8.1).  Creating a giant hole in the ground is generally cheaper than making an underground mine, but it is also less precise, so it is necessary to mine a lot of waste rock along with the ore, and that waste rock becomes an environmental liability. Relatively deep deposits, or those with elongated or irregular shapes are typically mined from underground with deep vertical shafts, declines (sloped tunnels) and levels (horizontal tunnels) (Figures 8.2.1 and 8.2.2).  In this way it is possible to focus the mining on the orebody itself. In some cases, the near-surface part of an orebody is mined with an open pit, while the deeper parts are mined underground.

Figure 8.2.1 Simplified Schematic of an Underground Mine. A shaft is a vertical mine opening, a level is horizontal and a decline is sloped so as to allow passage by wheeled vehicles.

Figure 8.2.2 Underground at the Myra Falls Mine, Vancouver Island

When ore is first processed (typically close to the mine) it is crushed to gravel-sized chunks and then ground to a fine powder and the ore minerals are physically separated from the rest of the rock to make a concentrate.  At a molybdenum mine, for example, this concentrate may be almost pure molybdenite (MoS₂).  The rest of the rock is known as tailings. It comes out of the concentrator as a wet slurry and is typically stored near to the mine, in most cases within a tailings pond.

The tailings pond at the Volcanogenic Massive Sulphide copper-zinc-silver-gold Myra Falls Mine on Vancouver Island is shown on Figure 8.2.3.  The tailings are contained by an embankment. Also visible in the foreground of Figure 8.2.3 is a pile of waste rock, which is non-ore rock that was mined in order to access the ore.  Although this waste rock has low levels of ore minerals, at many mines it contains up to a few percent pyrite. The tailings and the waste rock at most mines are an environmental liability because they both contain pyrite plus small amounts of ore minerals. When pyrite is exposed to oxygen and water it generates sulphuric acid—also known as acid rock drainage (ARD). Acidity itself is a problem to the environment, but because the ore elements, for example copper or lead, are more soluble in acidic water than in neutral water the ARD is also typically quite rich in metals, many of which are toxic. The causes and consequences of ARD are summarized below in Box 8.1.

Figure 8.2.3 The Tailings Pond at the Myra Falls Mine on Vancouver Island.  The dry rock in the foreground is waste rock. The structure in the background on the right is the head-frame for the mine shaft.  Myra Creek flows between the tailings pond and the head frame, and the tailings embankment has been constructed to keep the tailings out of the creek.

Figure 8.2.4 The Mt. Polley Mine Area Before (left) and After (right) to the Dam Breach of August 2014.  The tailings were stored in the area labelled “retention basin”.

At the time of the failure of the retention dam work was underway to raise the level of the dam so that more tailings could be stored. Because of changes to the mining regulations the structure had not been inspected by government officials for at least two years. The released water and metal-rich tailings flowed first into Polley Lake, and then down Hazeltine Creek into Quesnel Lake. Surface water samples collected in the affected water bodies showed elevated levels of suspended solids, chromium, copper, iron and phosphorous during 2014, but the levels had returned to “normal” by 2016. The remediation project has cost the mining company $70 million, and the mine was out of operation for 3 years (2014 to 2017). Three engineers that were involved in the design, construction and maintenance of the dam are facing disciplinary hearings by Engineers and Geoscientists B.C.

Most mines have concentrators on site because it is relatively simple to separate ore minerals from non-ore minerals, and thus significantly reduce the costs and other implications of transportation.  But separation of ore minerals is only the preliminary stage of metal refinement, for most metals the second stage involves separating the actual elements within the ore minerals.  For example, the most common ore of copper is chalcopyrite (CuFeS₂).  The copper must be separated from the iron and sulphur to make copper metal and that involves complicated and very energy-intensive processes that are done at smelters or other types of refineries.  Because of their cost and the economies of scale, there are far fewer refineries than there are mines.

Some of the potential environmental effects of metal mines are summarized on Table 8.2.  All mining activities result in a loss of natural environment because land has to be cleared and is no longer available for or useful to the plants and animals that once lived there. Mining involves blasting, breaking and crushing rock, and so dust is created. Some of this material is metal-rich, and is fine enough to be wind-borne, and so can be dispersed widely.  Some gets into water courses. Because sulphide minerals—especially pyrite—lead to acidification of water acid rock drainage is a common problem around mine sites (see Box 8.1). Materials, such as waste rock that is piled up around a mine, are subject to slope failure and such failures represent risks to people, infrastructure and to the environment. Tailings represent a special class of such materials because they are typically saturated with water, and are also rich in metals and so are potential sources of ARD. Tailings are a problem if their containment structure is breached (as at Mt. Polley), but can also be an ongoing issue of the water they contain is able to slowly leak from the storage facility. Finally, smelting processes involve application of enough heat to melt ore concentrates and that produces gases and ash that become airborne and hence are widely dispersed. In most cases the heat comes from fossil fuels.

The Sudbury region of Ontario is an example of the environmental risk of smelter stacks. Sudbury is home to the largest nickel smelting operation in the world. By the middle part of the 20th century a region of several hundred square km around mine district had either been completely de-vegetated or damaged by the contaminants released or the acidity they produced.  In 1970 a 385 m tall superstack was built in Sudbury to allow smelter contamination to be dispersed further. This allowed for eventual revegetation of the region around the city but resulted in the spread of the contamination over a much wider area, even into other provinces and states.

Box 8.1 Acid Rock Drainage at Mt Washington, BC

Any material that contains the mineral pyrite (FeS₂), or some other iron sulphide mineral, has the potential to produce acidity in the surrounding surface environment. That’s because pyrite oxidizes quite readily at surface, and its oxidation leads to the release of hydrogen ions (H⁺) into the surface water and groundwater. A simplified equation for this process is as follows:

2FeS₂ + 7O₂ + 2H₂O -> 2Fe²⁺ + 4SO₄^2- + 4 H⁺
pyrite + oxygen + water -> dissolved iron + sulphate + hydrogen ions

Materials around a mine site that might contain pyrite include rock outcrops exposed by mining, road-building or construction, waste rock or any rock that was mined but doesn’t have enough metal to be considered ore, and tailings from a processing plant. The important point is that these materials may have been underground and so not susceptible to weathering, but, because of construction and mining activities they are now exposed to water and oxygen at surface.

Figure 8.2.5 shows part of the Mt. Washington porphyry copper mine on Vancouver Island. Waste rock piles are present in the background, and rock that has been exposed by mining is visible in the foreground. Both are producing acid rock drainage, and the pH of the water in this area is as low as 4.

Figure 8.2.5 Waste Rock and Acid-Rock Drainage at the Mt. Washington Mine, BC

Acid water alone has a negative effect on aquatic and terrestrial organisms, but in most cases the acid drainage from a mine also includes elevated levels of metals. This is because the rocks that are weathering tend to be metal-rich, and also because most metals are more soluble in acidic water than in neutral water.

Water draining from the Mt. Washington Mine flows into the Tsolum River an important salmon stream in the Comox Valley. In the 1930s and 1940s this stream had runs of up to 200,000 pink salmon, 15,000 coho, 11,000 chum and 3,500 steelhead. The mine operated for only 3 years from 1964 to 1966, but by 1987 only 14 coho could be found in the river and other fish stocks were also significantly depleted. In 1982 a fish hatchery released 2.5 million pink salmon fry into the Tsolum River; none returned.

Very low copper concentrations, in the order of 10 µg/L, can be toxic to Pacific salmon. In the 1980s and 1990s copper concentrations in the Tsolum River ranged from 20 to 400 µg/L, while copper concentrations in one of the tributaries leading from the Mt. Washington Mine were in the range 5,000 to 15,000 µg/L.

A remediation plan for the Tsolum River was established in 2003. The plan included steps to reduce the level of acid drainage at source, and also to naturally treat some of the water flowing into the river. The main stream draining the mine site was diverted into an existing wetland, and the flow out of this wetland was controlled so that most flow coincided with peak discharge periods of the Tsolum River (to achieve dilution). Reduction in ARD at source was achieved by diverting some surface drainage around the old mine workings and covering a large part of the workings with a bituminous liner.
By 2010 the copper level in the Tsolum River was consistently below 11 µg/L, and it has continued to drop since then. In 2018 over 220,00 Pink salmon, 32,000 Chum, 4,100 Coho and 600 Steelhead were counted in the river.

Media Attributions

• Figure 8.2.1 Steven Earle, CC BY 4.0
• Figure 8.2.2 Photo by Steven Earle, CC BY 4.0
• Figure 8.2.3 Photo by Steven Earle, CC BY 4.0
• Figure 8.2.4 Steven Earle, CC BY 4.0, after Jesse Allen, 2014, based on public domain Landsat data from the U.S. Geological Survey, via Wikipedia, https://en.wikipedia.org/wiki/Mount_Polley_mine#/media/File:Mount_Polley_Mine_site.jpg; and https://en.wikipedia.org/wiki/Mount_Polley_mine#/media/File:Mount_Polley_Mine_dam_breach_2014.jpg
• Figure 8.2.5 Photo by Steven Earle, CC BY 4.0

Metals are critical for our technological age, but there are a lot of other not-so-shiny materials that are needed to facilitate our way of life.  For everything made out of concrete or asphalt we need sand and gravel.  For concrete we also need limestone to make lime (CaO).  For the glass in our computer screens, and for glass-sided buildings, we need silica sand plus sodium oxide (Na₂O), sodium carbonate (Na₂CO₃), and calcium oxide (CaO). Potassium is an essential nutrient for farming in many areas, and for a wide range of applications (e.g., ceramics and many industrial processes) we also need various types of clay.

Sand and gravel represent, the greatest volume of mined material on Earth, at around 50 billion tonnes/year, which is 20 times as much as iron . The best types of sand and gravel resources are those that have been sorted by streams, and in many regions of Canada and the northern USA (and elsewhere) the most abundant and accessible fluvial deposits are associated with glaciation.  That doesn’t include till of course, because it has too much silt and clay, but it does include glaciofluvial outwash, which is present in thick deposits in many glaciated regions, similar to the one shown on Figure 8.3.1.  In a typical gravel pit these materials are graded on-site according to size and then used in a wide range of applications from constructing huge concrete dams to filling children’s sandboxes.  Sand is also used to make glass, but for most types of glass it has to be at least 95% quartz (which the sandy layers shown in Figure 8.3.1 are definitely not), and for high-purity glass and for the silicon wafers used for electronics the source sand has to be over 98% quartz.

Figure 8.3.1 Sand and Gravel in an Aggregate Pit Near to Nanaimo, BC

Approximately 4 billion tonnes of concrete are used globally each year—a little over one-half tonne per person. The cement used for concrete is made from approximately 80% calcite (CaCO₃) and 20% clay.  This mixture is heated to 1450˚ C to produce the required calcium silicate compounds (e.g., Ca₂SiO₄) and during that process the carbonate is transformed into carbon dioxide which is released into the atmosphere. The calcite typically comes from limestone quarries like the one shown on Figure 8.3.2. Limestone is also used as the source material for many other products that require calcium compounds, including the manufacture of steel and glass, processing pulp and paper and in plaster products for construction.

Figure 8.3.2 Triassic Quatsino Fm. Limestone Being Quarried on Texada Island, BC

Sodium is required for a wide range of industrial processes, and the most convenient source is sodium chloride (rock salt), which is mined from evaporite beds in many parts of the world.  Rock salt is also used as a source of sodium and chlorine in the chemical industry, to melt ice on roads, as part of the process of softening water, and as a seasoning.  Under certain conditions the salt sylvite (KCl) accumulates in evaporite beds. The potassium is used as a fertilizer.

Another evaporite mineral, gypsum (CaSO₄.2H₂0) is the main component of plaster board (“drywall”) that is widely used in the construction industry.

Rocks are quarried or mined for many different uses, such as building facades, countertops, stone floors and head-stones.  In most of these cases the favoured rock types are granitic rocks and marble.  Quarried rock is also used in some applications where rounded gravel isn’t suitable, such as the ballast (road bed) for railways, where crushed angular rock is needed because it provides a more stable base.

Exercise 8.3  Sources of Important Lighter Metals

When we think of the manufacture of consumer products, plastics and the heavy metals (copper, iron, lead, zinc) easily come to mind, but we often forget about some of the lighter metals and non-metals that are important.  Consider the following elements and determine their sources.  Answers for all of these except magnesium can be found above.  See if you can figure out a likely mineral source of magnesium.

Exercise answers are provided Appendix 2.

Media Attributions

• Figure 8.3.1 Photo by Steven Earle, CC BY 4.0
• Figure 8.3.2 Photo by Steven Earle, CC BY 4.0

There are several types of fossil fuels, but all of them involve the storage of organic matter in sediments or sedimentary rocks. All fossil fuels are rich in carbon and almost all of that carbon ultimately originates from CO₂ taken out of the atmosphere millions of years ago during photosynthesis. That process, driven by solar energy, involves reduction (the opposite of oxidation) of the carbon, resulting in it being combined with hydrogen instead of oxygen. The resulting “organic matter” is made up of complex and varied carbohydrate molecules.

Most of the organic matter produced in this way is oxidized back to CO₂ relatively quickly once the organism dies (within weeks to decades in most cases), but any of it that gets isolated from the oxygen of the atmosphere, for example deep in the ocean or in a stagnant bog, may last long enough to be buried by sediments and, if so, may be preserved for tens to hundreds of millions of years. Under natural conditions, that means it will be stored until those rocks are eventually exposed at surface and weathered.

In this section we’ll discuss the origins and extraction of the important fossils fuels, including coal, oil and gas. Coal, the first fossil fuel to be widely used, forms mostly on land in swampy areas adjacent to rivers and deltas in areas with humid tropical to temperate climates. The vigorous growth of vegetation leads to an abundance of organic matter that accumulates within the stagnant water of swamps, ponds and lakes with little circulation, and thus does not decay and oxidize. This situation, where the dead organic matter is submerged in oxygen-poor water, must be maintained for centuries to millennia in order for enough material to accumulate to form a thick layer (Figure 8.4.1a). At some point the swamp deposit is covered with more sediment—typically because a river changes its course or overtops its bank (Figure 8.4.1b). As more sediments are added the organic matter starts to become compressed and heated. Low grade lignite coal will form at depths between a few 100 m and 1500 m and temperatures up to about 50˚ C (Figure 8.4.1c). At between 1000 to 5000 m and temperatures up to 150˚ C m bituminous coal will form (Figure 8.4.1d). At depths beyond 5000 m and temperatures over 150˚ C anthracite coal will form.

Figure 8.4.1 Formation of Coal. (a) Accumulation of organic matter within a swampy area, (b) the organic matter is covered and compressed by deposition of a new layer of clastic sediments, (c) with greater burial lignite coal is formed, and (d) at even greater depths bituminous (and eventually anthracite) coal are formed.

During the process of converting organic matter to coal some methane is produced and it is stored within the pores of the coal. When coal is mined methane is released into the mine where it can become a serious explosion hazard. Modern coal-mining machines have methane detectors on them and will actually stop operating if the methane levels are dangerous. It is possible to extract the methane from coal beds without mining the coal and the gas recovered in this way is known as coal bed methane.

While almost all coal is formed on land from terrestrial vegetation, most oil and gas is derived primarily from marine microorganisms that accumulate within sea-floor sediments. In areas where marine productivity is high dead organic matter is delivered to the seafloor fast enough that at least some of it escapes being oxidized. This material accumulates in the muddy sediments and then those get buried to significant depth beneath other sediments.

Figure 8.4.2 The Depth and Temperature Limits for Biogenic Gas, Oil, and Thermogenic Gas

As the depth of burial increases so does the temperature—due to the geothermal gradient—and gradually the organic matter within the sediments gets converted to hydrocarbons (Figure 8.4.2). The first stage is the biological production (involving anaerobic bacteria) of methane. Most of this escapes back to surface, but some is trapped in methane hydrates near to the sea floor. At depths beyond about 2 km, and at temperatures ranging from 60 to 120˚ C, the organic matter is converted by chemical processes to oil. This depth and temperature range is known as the oil window. Beyond 120˚ C most of the organic matter is chemically converted to methane (i.e., natural gas).

The rock within which the formation of gas and oil takes place is known to petroleum geologists as the source rock. This rock is typically rich in organic matter, and a good example would be a black shale. Both liquid oil and gaseous methane are lighter than water, so as liquids and gases are formed they tend to move slowly towards surface, out of the source rock and into reservoir rocks. Reservoir rocks are typically relatively porous and permeable rocks such as sandstone or fractured limestone, because that allows migration of the fluids from the source rocks, and also facilitates recovery of the oil or gas. In some cases, the liquids and gases make it all the way to surface, where they are oxidized and the carbon is returned naturally to the atmosphere, but in others they are contained by overlying impermeable rocks (a.k.a. “cap rock”, e.g., mudrock) in situations where anticlines, faults, stratigraphy changes and reefs or salt domes create traps (Figure 8.4.3).

Figure 8.4.3 Migration of Oil and Gas from Source Rocks into Traps in Reservoir Rocks

The liquids and gases that are trapped within reservoirs will become separated into layers based on their density, with gas rising to the top, then oil, and water underneath. The proportions of oil and gas will depend primarily on the temperature in the source rocks. Some petroleum fields, such as many of those in Alberta, are dominated by oil, while others, notably those in northeastern BC, are dominated by gas.

In general petroleum fields are not visible from surface, and their discovery involves the search for structures in the sub-surface that have the potential to form traps. Seismic surveys are the most commonly used tool for early-stage petroleum exploration, as they can reveal important information about the stratigraphy and structural geology of sub-surface sedimentary rocks. An example from the Gulf of Mexico south of Texas is shown on Figure 8.4.4. In this area a thick evaporite deposit (“salt”) has formed domes because salt is lighter than other sediments and tends to rise slowly towards surface, and this has created traps. The sequence of deformed rocks is capped with a layer of undeformed rock.

Figure 8.4.4 Seismic Section Through the East Breaks Field in the Gulf of Mexico. The dashed red line marks the approximate boundary between deformed rocks underneath and younger undeformed rocks lying horizontally on top. The wiggly arrows are interpreted migration paths.

The type of oil and gas reservoirs illustrated in Figures 8.4.3 and 8.4.4 are described as conventional reserves. Some unconventional types of oil and gas include oil sands, shale gas and coal-bed methane.

Oil sands are important because the reserves in Alberta are so large (the largest single reserve of oil in the world), but they are very controversial from an environmental and social perspective. They are “unconventional” because the oil is exposed at surface and is highly viscous because of microbial changes that have taken place at surface. The hydrocarbons that form this reserve originated in deeply buried Paleozoic rocks adjacent to the Rocky Mountains and migrated up and towards the east (Figure 8.4.5).

The oil sands are controversial primarily because of the environmental cost of their extraction. Since the oil is so viscous, it requires heat energy to make it sufficiently liquid to process. This energy comes from gas; approximately 25 m³ of gas is used to produce 0.16 m³ (~one barrel) of oil. (That’s bad, but not quite as bad as it sounds, as the energy equivalent of the required gas is about 20% of the energy embodied in the produced oil.) The other environmental cost of oil sands production is the devastation of vast areas of land where strip-mining is taking place, and the unavoidable release of contaminants into the groundwater and rivers of the region.

At present most oil sand recovery is achieved by mining the sand and processing it on site. Exploitation of oil sand that is not exposed at surface depends on in situ processes, an example being the injection of steam into the oil-sand layer to reduce the viscosity of the oil so that it can be pumped to surface.

Figure 8.4.5 Schematic Cross-Section of Northern Alberta Showing the Source Rocks and Location of the Athabasca Oil Sands

Shale gas is gas that is trapped within rock that is too impermeable for the gas to escape under normal conditions and can only be extracted by fracturing the reservoir rock using water and chemicals under extremely high pressure. This procedure is known as hydraulic fracturing or fracking. Fracking is controversial because of the volume of water used, and because the fracking companies are not required to disclose the nature of the chemicals used. Although fracking is typically done at significant depths there is always the risk that overlying water-supply aquifers could be contaminated (Figure 8.4.6). Fracking also induces low-level earthquakes, which do have the potential to cause damage.

Figure 8.4.6 Depiction of the Process of Directional Drilling and Fracking to Recover Gas From Impermeable Rocks. The light blue arrows represent the potential for release of fracking chemicals to aquifers.

It is important for us to understand the origins and exploitation of fossil fuels, but it is equally important to recognize that the Earth can no longer sustain the current rate of fossil fuel use, or in fact any use at all. If we want to avoid catastrophic climate change we need to reduce the use and production of fossil fuels quickly—eventually to zero, and that means there is no point in searching for new fossil fuel resources, or in further developing known resources.

At the United Nations Framework Convention on Climate Change meeting in Paris in 2016 the countries of the world agreed to a goal of limiting anthropogenic (human-caused) warming to 1.5° C above pre-industrial levels. In order to achieve that goal, the Intergovernmental Panel on Climate Change (IPCC) has stated that we must decrease anthropogenic CO₂ emissions (and therefore fossil fuel use) by 45% below 2010 levels by 2030, and that we must reach net-zero CO₂ emissions by 2050. If we wish to limit warming to 2.0° C above pre-industrial levels, we must decrease anthropogenic CO₂ emissions by 25% below 2010 levels by 2030, and that we must reach net-zero CO₂ emissions by 2070.

Media Attributions

• Figure 8.4.1 Steven Earle, CC BY 4.0
• Figure 8.4.2 Steven Earle, CC BY 4.0
• Figure 8.4.3 Steven Earle, CC BY 4.0
• Figure 8.4.4 Modified by Steven Earle, CC BY SA 4.0, based on AAPG data from Lovely and Ruggiero (1995, personal communication), via Wikimedia Commons, http://wiki.aapg.org/File:Sedimentary-basin-analysis_fig4-55.png
• Figure 8.4.5 Steven Earle, CC BY 4.0
• Figure 8.4.6 Modified by Steven Earle, based on Hydraulic Fracturing diagram by Mike Norton, 2013, CC BY-SA 3.0, via Wikimedia, https://en.wikipedia.org/wiki/Hydraulic_fracturing#/media/File:HydroFrac2.svg)

Mining and ore-processing require a huge amount of energy for operating machinery, transportation, and, especially, for smelting and refining. That includes about 15% of total global electricity and 11% of total global energy overall, and therefore acquiring the metals used to make all the things we buy has a huge climate impact. We need to think carefully about that every time we buy something that has metal in it. And, of course, we need to think about that every time we buy a manufactured product that has anything in it.

Recovering fossil fuels is also energy intensive and leads to emissions of greenhouse gases at every step in the process, but the really significant climate cost of fossil fuels is in their use, and that’s why we can no longer consider fossil fuels as an energy source for the future, and we have to stop looking for and developing more fossil fuel resources right now.

• Mining results in exposure of rock to weathering, both within the mines (especially surface mines) and because ore-processing involves crushing and grinding rock into small pieces, producing waste materials that are highly susceptible to weathering.
• Weathering includes oxidation of sulphide minerals, releasing acid into the environment,  leading to some of the outcomes summarized in Box 8.1, and might also include hydrolysis of silicates, in which case it could consume CO₂ from the atmosphere.
• Refining metals (especially smelting) introduces a wide range of toxic materials into the atmosphere, as well as acidity, that can have significant negative for plant life, and so for ecosystems in general.
• Production of cement by heating of calcium carbonate results in the release of carbon dioxide, and so contributes to climate change.
• Mining, and the related construction of roads and railways, contributes to slope failures and that can have a range of Earth systems implications, some of which are described in Chapter 1.

Solar Energy

We use solar energy to grow food of course, and for some industrial processes (like concentrating lithium brines), and to passively heat buildings, but early in the 21st century the most important and fastest growing use of solar energy is for generating electricity. Solar energy is abundant. The amount of sunlight energy received on Earth is approximately 10,000 times the amount of energy consumed by humans in 2005. As shown on Figure 9.1.1, solar potential is greatest in sub-tropical regions, but it is also significantly affected by weather conditions, and is highest in drier climates. That is why the dry central region of southern Canada, for example, has much greater solar potential than the wetter and cloudier west coast, even at the same latitude.

Figure 9.1.1 Global Solar Energy Potential

There are two main ways to generate electricity from the sun. One is through heat, the other through photovoltaic cells. The Crescent Dunes facility in Nevada is an example of solar heat electricity generation (Figure 9.1.2). This facility has 10,347 movable mirrors that reflect sunlight onto a tower in the centre of the complex. The highly concentrated sunlight is used to heat molten salt that powers a steam turbine to generate electricity. The hot molten salt retains enough heat to continue producing electricity for 9 hours after sundown, and therefore the plant can operate through the high electricity-demand period in the evenings. The Crescent Dunes station operated from 2016 to 2019 but is currently shut down because of technical issues with the molten salt container. Solar thermal energy has the disadvantage of relatively high capital and operating costs, such that the electricity production cost at most existing plants is currently higher than for some other sources, including photo-voltaic solar. But it has the advantage of short-term energy storage, which allows for energy production into the evening when it’s need most and that could outweigh higher production cost in areas where there are time-of-day premiums on the value of electricity.

Figure 9.1.2 Crescent Dunes Thermal Solar Plant Near to Tonopah, Nevada

Photo voltaic solar (or PV) is based on the use of cells that convert sunlight directly to electricity. Solar PV is easily scalable, from several modules on a single roof (Figure 9.1.3) to thousands of modules at a utility-scale facility (Figure 9.1.4). The efficiency of solar PV technology has increased significantly over the past several decades. Commercially available modules have evolved from about 10% efficiency in the 1970s to about 20% efficiency in the early 2020s. Experimental solar cells are now operating at about 40% efficiency—meaning that 40% of the sun’s energy that strikes them is converted into electricity. While efficiency has gone up over that period, costs have come down dramatically. According to the International Energy Agency the average cost of solar PV modules has dropped from over $100/watt in 1975, to $10/W in 1987, to $1/W in 2015, and to $0.2/W in 2020.

Figure 9.1.3 Installing Solar Modules on a Roof

Figure 9.1.4 A Commercial Solar PV Installation in Greece

Box 9.1 Watts and Watt Hours

(Steven Earle, CC BY 4.0)

In this chapter we are talking about power and energy using the terms Watts and Watt hours. A Watt is a measure of power, which is the rate that energy is produced or consumed by a device. A 10 W LED light bulb consumes electricity at a rate of 10 W. A Watt is already a rate, so we don’t need to use something like “W/hour”. A 300 W solar panel in full sunlight generates electricity at a rate of 300 W. Each hour at that rate it will produce 300 W hours of energy. A 300 W solar module could power thirty 10 W light bulbs for as long as the sunlight holds up.

During each hour of full sunlight, a 300 W solar module will generate 300 Wh (Watt hours) of energy. A solar installation like the one in Figure 9.1.3, has a capacity factor of about 20%, meaning it can only generate about 20% of its rated power over the course of a year (because of darkness and cloud cover). In one average day each module in that system will generate 24 x 300 x 0.2 = 1,440 Wh of energy (or 1.44 kWh). In one year, each module will generate 1.44 x 365 = 526 kWh.

In the context of large energy projects, we need bigger numbers to express the rate of power generation, such as MW (Mega-Watts or 1,000,000 W) or GW (Giga-Watts or 1,000,000,000 W), or energy production, such as MWh and GWh.

 

Solar PV modules (and thermal solar) do not generate electricity all of the time, certainly not when the sun is down, and not that well under cloudy conditions. In areas with a reasonably good solar resource the capacity of solar modules is about 20% of the rated wattage. In other words, 100 modules rated at 250 watts (25 kW) should produce electricity at an average rate of 5 kW, or should generate 43,800 kWh of energy over the course of a year, which is about enough to power 4 houses in North America, more than that in Europe.

According to the International Renewable Energy Agency solar represented 23% of the world’s renewable energy supply in 2019, and that proportion is growing. The increasing efficiency and decreasing cost of solar PV modules has made it a viable option for a large part of the energy supply in a world without fossil fuels. The US National Renewable Energy Lab has estimated that a national grid with 55% solar and wind sources—both of which are intermittent—could be viable if the remaining power sources can be ramped up and down to even out the supply. Energy storage (such as pumped hydro or batteries) could also help to make this a reality.

A typical solar PV module, like those shown on Figure 9.1.3, contains about 76% glass (by weight), 10% plastic, 8% aluminum (in the frame), 5% silicon and 1% other metals (as wires and connectors). It takes energy to make these components and assemble them, but the amount of embodied energy is small compared with the amount of energy that can be produced by the module. The energy output of a solar module depends significantly on the region and the setting in which it is installed, but a modern well situated solar module should recover its energy costs in approximately one year, and the module should last for at least 25 years. Furthermore, many parts of a solar module are easily recyclable (or re-usable) including the glass and the metal frame. The silicon in the cells can also be melted for re-use. Even though the modules can be recycled, a significant increase in the use of solar-PV energy (which is what we need) will require more mining to supply the raw materials.

Exercise 9.1 Solar Potential

Using Figure 9.1.5 (or Figure 9.1.1 if you live outside of this area), determine the solar energy potential of your region. If you live in a reddish area, the potential is very high (score 3), orange: high (score 2) and yellow: moderately high (score 1).

Figure 9.1.5 Global solar energy potential for North America  (by The World Bank, 2017, CC BY 4.0, https://commons.wikimedia.org/wiki/File:Global_Map_of_Photovoltaic_Power_Potential.png)

Now think about the potential of the place in which you live (the building, or the piece of land around it). If there are places (on the roof, on the ground, or even on the walls) that get at least 4 hours of direct sun in a day (averaged year-round, and not counting bad weather), then your potential is moderate (score 1), if it’s 6 hours/day then it’s good (score 2), and if it is at least 8 hours/day then it is very good (score 3).

If the sum of your two scores is 4 or higher, then your place is likely to be viable for solar power, and that means that the cost of installation of solar panels would probably be recovered in about 10 years (or less). Of course, there are many other factors, such as the nature of your building, the cost of electricity, and the willingness of your electric utility to let you tie in to the grid.

Wind Energy

Wind is a product of solar energy and gravity because air masses move in response to differences in air density created by solar heating. Wind energy has been used for centuries to turn windmills and to power sailing vessels, but it’s only in the past few decades that wind has been used to generate electricity. As shown on Figure 9.1.5, wind is geographically variable as an energy resource. Average wind speeds are higher on the oceans than on land, and consistently higher in flat areas compared with mountainous areas. In North America the best wind resources are in eastern coast offshore, especially north from Virginia, and in western coast offshore areas, from California north, within the Great Lakes and Hudson Bay, and within the plains in both the US and Canada. Mountain ridges also have consistently strong winds.

Figure 9.1.6 Typical Wind Speeds at 100 m Elevation Within and Around North America

Most utility-scale wind turbines are of the horizontal axis type, mounted on a tall tower. Over the past 30 years the power potential of wind turbines has increased by a factor of nearly 30 times (Figure 9.1.6) and the towers have become very tall. Of course, a 10 MW wind turbine doesn’t generate power at 10 MW all of the time. The typical capacity factors for land-based turbines are in the order of 30%, while offshore turbines are higher, at around 40%. The tall 10 MW and higher turbines being installed in offshore locations in the mid 2020s are likely to have even higher capacity factors because wind consistency increases with elevation. Most existing offshore turbines are embedded in the sea floor in areas with water depths of less than 50 m. Floating turbines have been developed for use in areas with deeper water. A single 10 MW offshore turbine operating at 50% capacity can generate enough electricity to power about 4,000 homes in North America.

Figure 9.1.7 Highest Rated Wind Turbines Available at Various Times Between 1990 and 2020

According to the International Renewable Energy Agency wind energy makes up about 25% of installed renewable energy. Based on data from 2019, China and the US have the highest installed wind generation capacity, but Denmark has the highest proportion of wind-generated electricity, at 48%, followed by Ireland (33%), Portugal (27%), Germany (26%), the UK (22% – see Figure 9.1.7) and Spain (20%).

Figure 9.1.8 An Offshore Wind Farm Under Construction in the UK in 2011. These turbines are rated at 3.5 MW.

Modern wind turbines are large and complex, and their manufacture and installation require a lot of materials and energy. For example, a typical 3 MW turbine, similar to the ones on Figure 9.1.7, includes about 1100 tonnes of concrete for the base on the sea floor (or on land), 276 tonnes of steel, for the base the tower and the nacelle (the housing that holds the generating mechanism, and to which the blades are attached), 2.6 tonnes of copper and 2.3 tonnes of aluminum, and 20 tonnes of fibreglass and epoxy resin for the blades. As for solar modules, the energy payback of a turbine will depend on the features of the location selected, but it is typically less than one year, compared with an expected lifetime of at least 20 years. Much of this material can be recycled when the turbine is decommissioned.

Every year thousands of birds (and bats) are killed by wind turbines, although the rate of bird-kill is decreasing because the newer larger turbines spin slower than older models, and because a range of strategies are being developed to decrease the risk. To put this in perspective, millions of birds are killed each year by power lines, moving vehicles and tall buildings, and the total number from all of these sources of mortality is dwarfed by the number killed by domestic cats.

Solar and wind power are currently the cheapest forms of electricity generation according to the US Energy Information Administration. The average levelized cost of solar is $0.30 per kWh, while that for onshore wind is $0.37 per kWh. All other forms of electricity generation are more expensive, including gas, coal, hydro and nuclear. While most other energy forms are likely to become more expensive in coming years, solar and wind are likely to get cheaper because of technological improvements and economies of scale.

Exercise 9.2 Intermittent Wind and Solar Resources

Figure 9.1.9 A Combined Wind Turbine and Solar PV Energy Facility

A combined wind turbine and solar PV energy facility (like the one shown on Figure 9.1.9). has been established to serve an area with about 50,000 homes in southern Alberta It comprises 150,000 solar modules at 600 W each (total capacity 90 MW) and 20 wind turbines at 5 MW each (total capacity 100 MW). The region is often sunny and has good wind resources, but of course there are cloudy days and dark winter days, and there are calm days.

The following table, which is based on the weather conditions for a week in late spring, shows how many hours of strong sunlight equivalent* there are on each day to power the modules at their 600 W capacity, and how many hours of strong wind equivalent are available to power the turbines at their rated 5 MW capacity.

Complete the other rows of the table to work out the average amount of energy produced by each system on each of the 7 days, and the total amount of energy produced. For example, to estimate the daily energy production for the solar array, in MWh, multiple the number of hours of strong sun equivalent by the total capacity of the system (90 MW). For wind, multiply the hours of strong wind equivalent by the total capacity (100 MW). The first column (Monday) is done for you.

An average house in North America uses about 10 MWh of electricity per year, which is equivalent to 0.027 MWh per day, so if you divide the “Daily total MWh” by 0.027 you can estimate the number of homes that can supplied with electricity by this system on each day.

As the facility operator, what do you plan to do with the extra electricity on days when you have a surplus, and to keep your customers supplied with electricity on days when you don’t have enough?

*For example, on a given day there might be 4 hours of strong sunlight, 3 hours of weak morning or evening sunlight, and 6 hours with conditions ranging from partial to complete cloud. These might all add up to 8 hours of “strong sunlight equivalent”.

Exercise answers are provided Appendix 2.

Media Attributions

• Figure 9.1.1 Global Map of Photovoltaic Power Potential by The World Bank, 2017, CC BY 4.0, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Global_Map_of_Photovoltaic_Power_Potential.png
• Figure 9.1.2 Crescent Dunes Solar by Amble, 2014, CC BY-SA 4.0, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Crescent_Dunes_Solar_December_2014.JPG
• Figure 9.1.3 Photo by Steven Earle, CC BY 4.0
• Figure 9.1.4 Horizontal Single Axis Tracker in Greece by Vinaykumar8687, 2013, CC BY-SA 4.0, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:8MW_horizontal_single_axis_tracker_in_Greece.JPG
• Figure 9.1.5 Global Map of Photovoltaic Power Potential by The World Bank, 2017, CC BY 4.0, via Wikimedia Commons. https://commons.wikimedia.org/wiki/File:Global_Map_of_Photovoltaic_Power_Potential.png
• Figure 9.1.6 From Global Wind Atlas 3.0, a free, web-based application developed, owned and operated by the Technical University of Denmark (DTU) and released in partnership with the World Bank Group, utilizing data provided by Vortex, using funding provided by the Energy Sector Management Assistance Program (ESMAP). CC BY 4.0 For additional information see: https://globalwindatlas.info
• Figure 9.1.7 Steven Earle, CC BY 4.0, based on data in Molina, M. and Mercado, P. (2011). Chapter 16 Modelling and control design of pitch-controlled variable speed wind turbines. In Al-Bahady, I. H. (Ed), Wind turbines. IntechOpen, https://www.intechopen.com/chapters/14810
• Figure 9.1.8 Walney Offshore Windfarm by David Dixon, 2011, Walney Offshore Windfarm CC BY-SA 2.0, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Walney_Offshore_Windfarm_-_geograph.org.uk_-_2391702.jpg
• Figure 9.1.9 Renewable Energy Park by hpgruesen, CC0 1.0, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Renewable_energy_park.jpg

Hydro power is a product of solar energy because water, evaporated by the sun, falls onto land at elevation, and then flows down to the ocean. If we interrupt that flow, we can use its potential energy to generate electricity. Hydro power is particularly effective in areas where there has been tectonic uplift to create mountains, and where there is abundant precipitation in the form or rain or snow.

There are two main types of hydro-electric facilities, those that involve construction of dams, and those that involve diversion of part of the water of a stream into a canal and/or a pipe. We’ll refer to the first type as “dam and reservoir hydro”, and to the second type as “run-of-river hydro”.

Hydro electricity is a very mature industry that goes back to the late 19th century. It currently represents 47% of the installed global renewable energy capacity (and 6.5% of all energy capacity), but the rate of growth in hydro is now very close to 0%. Although there have been many enhancements to hydro energy technology in the past 100 years, it’s not likely that we will see significant technological advances in the future, and while there are many rivers that have still hydro potential, the public appetite for changing how rivers flow, just to generate electricity, is waning.

Dam and Reservoir Hydro

Most large hydro projects involve construction of a dam across a river and the creation of a reservoir behind that dam. The difference in elevation between the surface of the reservoir and that of the river below the dam is the hydraulic head, and the amount of energy that can be generated is proportional to that difference and to the rate at which water can flow through the turbines. An example is the Revelstoke Dam on the Columbia River in British Columbia (Figure 9.2.1). The Revelstoke Dam is 175 m high. It has created a reservoir (Lake Revelstoke) that is 130 km long and has flooded just over 100 km² of land. The electricity production capacity at Revelstoke Dam is 2,480 MW.

The significant value of a dam and reservoir is that a lot of water is stored (except under long-term drought conditions), and that water can be used to generate electricity according to the demand. The rate of flow through the turbines can be slowed (or even stopped) during the night (or over the summer) when demand is low and increased in the daytime (or during the winter) when the demand is higher. This type of production flexibility is especially valuable in areas where there is a significant proportion of non-dispatchable electricity generation from sources like solar and wind.

Figure 9.2.1 The Revelstoke Dam on the Columbia River, British Columbia. Water flows through the large pipes (penstocks) to turbines that are within the structure in the centre, and then out into the river below. The inset shows part of the reservoir (Lake Revelstoke) on the upper side of the dam.

Two significant disadvantages of dam and reservoir hydro are that they block passage of migratory fish, and that they result in the flooding of large areas of land that was useful for other purposes or for nature. It’s true that fish ladders can be constructed, but they are not typically successful in the case of a high dam like the one at Revelstoke. Flooding of land is inevitable when dams are constructed, and the implications of that include:

• loss of human habitat (e.g., construction of the Three Gorges Dam in China, the world’s single largest hydro project, resulted in relocation of 1.24 million residents from 13 cities, 140 towns and 1350 villages),
• loss of farmland (which is a big issue for the Site C dam under construction in British Columbia),
• loss of natural terrestrial and aquatic (river) habitat,
• increased risk of slope failure and bank erosion (see Section 5.2 in relation to the Revelstoke Dam and the Downie Slide),
• potential for release of mercury naturally present in soils, and
• potential for release of carbon dioxide from breakdown of organic matter (this is most likely to be an issue in tropical areas and where the reservoir is shallow).

Run-of-River Hydro

A run-of-river hydro installation is illustrated on Figure 9.2.2. Sechelt Creek flows into a coastal inlet about 60 km northwest of Vancouver. The three parts of the project are the head pond, at an elevation of 360 m, where a weir has been built to ensure that the upper end of the penstock is always submerged. Water flows from there through a 4-km long 1.2 m diameter steel pipe (penstock) to the powerhouse, where it is used to turn either one or both of two 8 MW turbine generators. This project did not involve flooding any land (the head pond is entirely within the river’s normal channel) and it has not disrupted the passage or habitat of migratory fish. (A waterfall just upstream from the powerhouse has always prevented ocean fish from migrating to the middle and upper parts of Sechelt Creek.)

Figure 9.2.2 Components of a Run of River Hydro Project at Sechelt Creek, British Columbia. Left: the weir and head pond (at 360 m elevation); centre: part of the penstock; right: one of two 8 MW turbines at 10 m elevation.

Water is not stored in a reservoir in the case of run-of-river hydro, instead the generation of electricity is dependent on the discharge of water at any one time. Over the course of a year the discharge of a stream like Sechelt Creek can vary significantly from season to season and even from day to day.

Two examples of river discharge variations are provided on Figure 9.2.3. The Qualicum River is situated in a coastal area of Vancouver Island where although the surrounding mountains are frozen and snow-covered over the winter, the lowland areas tend to remain unfrozen and rain is common. The hydrograph is dominated by rainfall in the fall, winter and early spring months, and very little precipitation from June through later September. The Stikine River drains a region in northern BC in a mountainous area that remains frozen through the winter and there is almost no melting between December and April. The hydrograph is characterized by very low flows during the cold months, and very strong snow-melt and rain-driven flows in the spring, summer and early fall.

Figure 9.2.3 1986 Variations in Daily Discharge on the Qualicum River (Vancouver Island) and the Stikine River (Northern Interior, BC)

The significant differences in the flow patterns of the Qualicum and Stikine Rivers illustrates the importance of understanding supply and demand when it comes to run-of-river generation. Sechelt Creek has a hydrograph similar to that of the Qualicum River, and so it can typically generate electricity at peak power in the period from November through June, but produces little or no power over the summer months. This fits well with the electrical power demand in British Columbia, which is greatest in winter, but a river like the Stikine would not be very suitable for run-of-river generation in British Columbia because it has relatively little flow in the winter months.

Media Attributions

• Figure 9.2.1 Steven Earle, CC BY 4.0
• Figure 9.2.2 Photos by Steven Earle, CC BY 4.0
• Figure 9.2.3 Steven Earle, CC BY 4.0, based on public domain data at Water Survey of Canada, Environment Canada, https://wateroffice.ec.gc.ca/

Waves are generated by wind blowing over water, so this form of energy originates from the sun. On the other hand, tides are related to rotation in the sun-Moon-Earth system and from the resulting variations in the gravitational force on Earth. That energy comes from the original angular momentum (spin) of the solar-system and the galaxy, and can possibly be traced right back to the Big Bang.

Wave Energy

The size of waves at any location is a function of the strength of the wind, and also the extent of the body of water over which that wind blows. This factor is known as the “fetch”. The longer the fetch the bigger the waves. Waves tend to be short and steep close to a storm, but longer and shallower far away. Typical large ocean waves have wavelengths in the order of 100 m and amplitudes in the order of a few metres. While waves are driven by the wind, they tend to be somewhat more reliable than the wind as a source of energy. At most exposed coastal locations there will be some wave action, even if there is little or no wind at that time. In general, wave energy is greatest in the areas of greatest wind energy, as shown on Figure 9.1.5.

Several different wave energy devices have been designed and tested over the past several decades. They include shore-based systems, sea-floor tethered systems, and anchored floating systems that bend as waves pass. Some of these are illustrated on Figure 9.3.1. All of them have cylinders that are compressed and then relaxed every time a wave goes by, and the pressure created by that is converted into electricity.

Figure 9.3.1 Three Different Wave Energy Systems

Although all three of these types of systems have been demonstrated in the past, at present (spring 2021) the only operating and grid-connected wave energy installations are of the fixed-surge type, and those are small, with power ratings in the kilowatts, not megawatts.

Tidal Energy

The tides are generated by variations in the geometry of the sun-Earth-Moon system and the effect that those have on the level of the surface of the ocean. In most areas there are two high tides and two low tides every day (known as semidiurnal tides), although a few regions have only one high and one low per day (diurnal tides). The range between high and low tides is highly variable from place to place. The world’s highest tide range are in the order of 16 m difference between low and high tides in the Bay of Fundy between Nova Scotia and New Brunswick. On the other hand, many regions have tidal ranges of less than 1 m. Some examples of differing tidal ranges are shown on Figure 9.3.2.

Figure 9.3.2 Typical Tidal Ranges at Three Stations, Old Port Tampa (Florida), Vancouver, British Columbia, and Grindstone Island, New Brunswick, February 2021

There are several ways to convert tidal energy into electrical energy, but the two most commonly explored methods include the construction of a barrage (a dam) across a tidal estuary, and the installation of underwater turbines within channels that have significant tidal flow. The barrage method has been proven to be effective, but it can have serious environmental implications because a barrage results in changes to water-level fluctuations, salinity and turbidity in tidal flat areas, most of which are of great ecological importance. A tidal barrage is shown on Figure 9.3.3. The barrage acts like a dam, and the flow of both incoming and outgoing tides can be converted to electricity using turbines installed within the structure. There are three significant operating barrage tidal installations at present, including Sihwa Lake (South Korea, 254 MW, operating since 2011), La Rance River (France, 240 MW, operating since 1966) and Annapolis Royal (Nova Scotia, 20 MW, operated from to 1984 to 2019). Others have been proposed, but as of 2021, none is under construction.

Figure 9.3.3 Part of the Sihwa Lake Tidal Barrage in South Korea, Completed in 2011 and Rated at 254 MW

Figure 9.3.4 A Single-Rotor Seaflow Tidal Stream Generator, With the Turbine in the Raised Position

The other main type of tidal system is a tidal stream turbine that is typically installed in a location with strong tidal currents. An example is shown on Figure 9.3.4. Most such devices are either embedded in, or resting on, the sea floor, although some are buoyant and are anchored to the sea floor. Based on information available in 2021, only one tidal stream is currently in operation, and that is the MeyGen project on the north coast of Scotland, which has four 1.5 MW sea-floor turbines connected to the grid, with more turbines reportedly to be under construction.

The potential for tidal energy is huge, but the engineering and environmental challenges are significant, and the high capital cost has been a barrier to implementation of many planned projects.

Media Attributions

• Figure 9.3.1 Steven Earle, CC BY 4.0
• Figure 9.3.2 Steven Earle, CC BY 4.0, based on public domain and Open Government Canada data from National Oceanographic and Atmospheric Administration, https://tidesandcurrents.noaa.gov/, and Fisheries and Oceans Canada, https://www.waterlevels.gc.ca/
• Figure 9.3.3 Sihwa Lake Tidal Power Station by Arne Müseler (arne-museler.com), CC-BY-SA-3.0, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Sihwa_Lake_Tidal_Power_Station_aerial_view.jpg
• Figure 9.3.4 Seaflow Raised by Fundy, CC-BY-SA-3.0, via Wikimedia Commons,  https://commons.wikimedia.org/wiki/File:Seaflow_raised_16_jun_03.jpg

Geothermal energy is heat that originates within the Earth—which includes heat left over from the original formation of the Earth and heat produced by radioactive decay—and is either naturally present at surface or is accessed at depth by drilling. We can harness this energy where there is higher than average heat flow from depth, and this is most common in areas near to active volcanoes. About 50% of all geothermal energy use is for heating buildings and other infrastructure, about 33% for hot pools and spas (Figure 9.4.1), while only about 17% is used to generate electricity. In many cases the left-over heat from electrical generation facilities is used for district heating or for swimming and bathing facilities.

Figure 9.4.1 The Geothermal Pool at Fludir, Iceland

A geo-exchange system (a.k.a. geothermal heat pump, or ground-source heat pump) is not based on geothermal energy at all. Instead, geo-exchange relies on the relatively constant temperature of the ground at depths between about 1 and 5 m, and that temperature is maintained by energy from the sun. Geo-exchange is used for either heating or cooling, or both.

Geothermal

In the context of generating electricity, geothermal heat is used to boil water or some other fluid to power a turbine. In the case of very hot sources, water piped to surface from depth will spontaneously convert to steam because of the reduction in pressure. If the water from depth is less than about 180° C it won’t spontaneously boil but can be used in a binary cycle system to heat a working fluid, such as pentene or toluene, that will boil at a lower temperature than water.

Figure 9.4.2 The Components of a Typical Binary Cycle Geothermal System

The components of a typical geothermal system are illustrated on Figure 9.4.2. The heat at depth (hundreds to a few thousand metres) is accessed via deep production wells. Water from within that rock (5) is pumped to surface and its heat is used to boil the working fluid to power the turbines (4). The original water is returned to the ground via an injection well (6) along with any additional surface water needed to maintain volume (1). Most geothermal plants have numerous production wells (dozens in some cases). It is common for wells to be cycled on and off to allow for recovery of the heat reservoir around them.

The Krafla power station in northern Iceland (Figure 9.4.3) is a 60 MW flash steam system (the hot water from depth boils at surface). It is situated in a very active volcanic region on the mid-Atlantic spreading ridge. An eruption episode that lasted from 1975 to 1984 almost led to cancellation of the project while it was under construction.

Figure 9.4.3 The Krafla Geothermal Plant in Iceland

The contribution of geothermal to total electricity production is important in some countries, especially in Iceland, where 30% of electricity is generated by geothermal (with most of the rest coming from hydro), and also the Philippines, at 27%, and El Salvador at 25%. In Costa Rica, Kenya and Nicaragua and New Zealand geothermal makes up 14, 11, 10 and 10% of electrical energy. These are all countries with active volcanoes and hence significantly high heat flow.

In the past 10 years global geothermal energy capacity has grown by an average of 3.75% per year. That is significant growth, although lower than for wind and solar, both of which have grown by an average of 20% per year in this century. But the potential for sustained growth is limited by the limited geothermal resource, and by the high capital cost of geothermal facilities.

Geo-Exchange

The first several metres of the ground beneath us represents a reservoir of thermal energy that is maintained by heat from the sun and stays nearly constant year-round. We can take advantage of that thermal reservoir by living in caves or—like Hobbits—digging deep into the sides of hills to build homes. Alternatively, we can stay above ground and install plumbing systems that allow us to tap into that stored warmth and bring it inside. The concept of geo-exchange is illustrated on Figure 9.4.4. An exchange loop can be vertical, with pipes installed in drilled holes, or horizontal, with pipes installed in an excavation at a depth of around 2 m.

Figure 9.4.4 Conceptual Vertical and Horizontal Geo-Exchange Loops

As shown on Figure 9.4.5, a geo-exchange loop can be used for both heating and cooling. For heating relatively warm fluid is pumped from the ground and then passed through a heat exchanger and/or heat-pump in the building to provide heating. For cooling relatively cool fluid is pumped from the ground and then passed through a heat exchanger for cooling. In fact, the “relatively warm” and “relatively cool” fluid may be approximately the same temperature, but it will be warmer than the outside air in winter, so useful for heating, and cooler than the outside air in summer, so useful for cooling.

Figure 9.4.5 Use of a Geo-Exchange Loop for Heating and Cooling

Geo-exchange is not a source of energy, but it is a way to provide significant savings in heating and cooling energy costs, both financial and from the perspective of climate change and the environment. According to the Geothermal Exchange Organization, a geo-exchange system can save building owners between 30 and 70% in heating costs and between 25 and 50% in cooling costs. The organization states that there are currently 750,000 geo-exchange systems in place, although it isn’t clear about what date or which jurisdiction that number refers to.

Media Attributions

• Figure 9.4.1 Photo by Steven Earle, CC BY 4.0
• Figure 9.4.2 Modified by Steven Earle, CC BY SA 4.0, based on EGS Diagram by Fishx, 2009, CC-BY-SA-3.0, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:EGS_diagram.svg
• Figure 9.4.3 Steven Earle, CC BY 4.0
• Figure 9.4.4 Modified by Steven Earle, from Loop Diagrams by Wgisol, 2015, CC BY-SA 4.0, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:ETI-FS_Loop_Diagrams,_01.26.11.jpg
• Figure 9.4.5 Steven Earle, CC BY 4.0

Nuclear energy is the energy that is stored in the nuclei of atoms; it comes in two types. Fission energy is released when large atoms like uranium split apart into smaller atoms. This is constantly taking place within the Earth—and it is part of the source of geothermal energy—and all around us, but the rate is very slow. We can build devices that increase that rate enough to generate heat in a controlled way, but there are some downsides to the technology.

Fusion energy is released when small atoms like hydrogen are fused together to make larger atoms. This is what is happening inside the sun, so we use that energy all of the time. We can reproduce nuclear fusion here on Earth in in the form of a hydrogen bomb, but that technology doesn’t allow for a controlled rate of energy release. Over the past several decades governments of several countries have spent tens of billions on research into controllable nuclear fusion. Fifty years ago, commercial nuclear fusion was considered to be less than fifty years away. During that time significant progress has been made in understanding what is needed to make that happen, but in 2021, it is still likely that a proof of the concept of fusion energy production is 5 to 10 years away, and that a fusion energy industry is likely another 25 years beyond that.

Nuclear Fission

Nuclear fission has been used for generating electricity since the late 1950s and there are currently around 440 fission reactors operating in 30 countries, producing about 10% of the world’s electricity. In a few countries, including France, Slovakia and Ukraine nuclear fission makes up more than 50% of electricity generation.

While naturally occurring uranium in rock splits apart only very slowly, the rate of fission is accelerated many thousands of times in a fission reactor because the uranium is highly concentrated, and the fuel is packed close enough together so that the neutrons released by one fission event will almost immediately trigger another fission in a nuclear chain reaction. An important point is that the chain reaction process must be carefully controlled by substances called moderators and, in the event of some irregularity, the reaction can be slowed relatively quickly by inserting control rods that inhibit the chain reaction process. It is important to realize that although a fission reactor can be dramatically slowed by inserting the control rods, the massive reactor core will still remain very hot for several days. This residual heat has been the cause of several nuclear accidents. A diagram of a type of fission reactor is provided on Figure 9.5.1, and the important components are described.

Figure 9.5.1 The Components of a Gas-Cooled Nuclear Fission Reactor . The uranium fuel elements are shown in orange. The graphite moderator, which keeps the reaction from getting out of control, is grey. The control rods that can be used to dramatically slow the chain reaction are black. Heat generated by the fission in the fuel is heats a gas which is circulated through a heat exchanger to boil water that is then used to run a steam turbine.

Modern nuclear reactors have many levels of protection against irregular operating conditions and against more significant failure and out-of-control behaviour. While that might give us some comfort, it cannot be denied that there have been some major failures in the nuclear power industry, at least one of which has had significant human and environmental implications. Three such failures are described in Box 9.1. Each of these led to a loss of public confidence in the safety of nuclear power, and to reductions in the number of reactors ordered and constructed in subsequent years.

Box 9.2 Nuclear Fission Plant Failures

Three significant nuclear plant failures—Three Mile Island (Pennsylvania), Chernobyl (Ukraine) and Fukushima (Japan)—have shaped the public’s perception of the safety of the nuclear industry, and have also informed the industry’s efforts to make nuclear reactors as safe as possible.

Figure 9.5.2 The Chernobyl Nuclear Fission Reactor Following the Accident in 1986

The Three Mile Island accident in March 1979 involved an operator error, the failure of a cooling system valve and the failure of the system that was meant to warn the operators of the valve’s malfunction. This led to a loss of coolant within the reactor vessel and resulted in partial melting of the nuclear fuel. There were high levels of radioactivity within the plant, and some radioactive materials were released into the environment via steam. Health officials eventually concluded that there were no deaths or significant illnesses that could be attributed to the accident, however that conclusion was controversial at the time, and remains so in the minds of many people.

The Chernobyl accident of April 1986 also involved an operator error and a system failure that resulted in the reactor power level increasing when it should have been decreasing (Figure 9.5.2). This led to a steam explosion that ruptured the reactor core and triggered an open-air reactor fire that continued for 9 days, releasing airborne radioactive material into the environment. A total of 28 plant workers and firefighters died within several days of the accident. A 2006 United Nations report concluded that as many as 4000 cancer deaths are likely to occur amongst 5 million residents in the contaminated region around the plant over the 80 years following the accident, and that this represents a 0.3% increase over expected cancer deaths for that region.

The Fukushima accident of March 2011 was the direct result of a magnitude 9 earthquake and a subsequent tsunami. Although none of the reactors was damaged by the earthquake, they were automatically shut down when the earthquake struck and therefore could not provide power to the cooling systems. Six external power supply systems failed because of the earthquake and although an emergency diesel power plant continued to supply power to cooling pumps, this was destroyed 50 minutes later by the 15 metre-high tsunami that overtopped protective barriers and inundated the plant. The result was that there was no cooling available for the still-hot reactors, three of which suffered melting within the reactor cores. No deaths have been attributed to the Fukushima accident, and although it is estimated that approximately 130 deaths may eventually occur because of elevated cancer rates , that number is small compared with the number of expected cancer deaths from other causes.

In a typical reactor the fuel elements will last for 18 months to a few years. Some types of reactors can be refuelled while still operating, but most have to be shut down for periods of up to several weeks. The spent fuel is removed and replaced with fresh fuel. The highly radioactive spent fuel is stored in swimming-pool sized tanks at the reactor site for about 5 years, and then is transferred to dry storage facilities, which are also on-site in most cases. On-surface storage of nuclear waste is not a viable option because the materials remain radioactive for thousands of years and represent both a health risk and security risk (from the perspective of nuclear proliferation). Several countries, including Canada and the United States, have conducted research into the deep underground storage of nuclear waste, but as of 2021, only Finland has an operating long-term storage facility.

Although the waste from nuclear fission is a serious problem, the volumes are not huge. The amount of radioactive waste produced for a person who used nuclear energy for their lifetime supply of electricity is about 3 kg. Nuclear fuel is quite heavy, so that amount would likely fit within a coffee cup.

Nuclear Fusion

Nuclear fusion can be achieved when hydrogen nuclei are forced close enough together to allow them to fuse together. The advantage of nuclear fusion for electricity generation is that it’s fuel (hydrogen) is abundantly available and it doesn’t produce significant amounts of toxic or radioactive waste. Nuclear fusion has the potential to produce energy from fuel that is virtually inexhaustible, without producing a significant amount of waste, but the process of achieving it is extremely complex, requires large amounts of energy to initiate the process, and extremely high temperatures (in the order of 100,000,000° C). These conditions, and the existence of lots of fast-moving neutrons, have possible implications for the mechanism that are still poorly understood.

The most advanced nuclear fusion project is ITER (International Thermonuclear Experimental Reactor) which has been under construction in southern France since 2008 and is expected to be complete by 2025 (Figure 9.5.3). This project is funded by the European Union, China, India, Japan, Russia, South Korea and the United States.

Figure 9.5.3 A Model of the ITER Fusion Reactor. There is a person for scale at the bottom.

ITER will not be used to produce any electricity, but it is expected to be able to generate about 10 times more energy than the amount needed for its operation. It is estimated that the total construction and operating cost of ITER will be in the region of $65 billion.

The follow-up to ITER is called DEMO (DEMOnstration Power Station). It is currently being planned by the EUROfusion consortium, but is unlikely to be ready to operate until 2050. Although DEMO is intended to produce electricity that will be fed into the grid, it will not be a commercially operating power station.

In summary, while nuclear fusion could provide abundant energy without significant waste products or climate implications, its development has been and will continue to be extremely expensive and its realization is still at least 30 years away, probably more than that.

Media Attributions

• Figure 9.5.1 Magnox Reactor Chematic By Emoscopes, CC BY-SA 3.0 Unported, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Magnox_reactor_schematic.svg
• Figure 9.5.2 Chernobyl Nuclear Fission Reactor from IAEA image bank, CC BY-SA 2.0, via Wikimedia Commons,  https://commons.wikimedia.org/wiki/File:IAEA_02790015_(5613115146).jpg
• Figure 9.5.3 ITER Exhibit by IAEA Imagebank, 2013, CC BY-SA 2.0, via Wikimedia Commons,  https://commons.wikimedia.org/wiki/File:ITER_Exhibit_(01810402)_(12219071813)_(cropped).jpg

We have some important choices to make in the next few decades because we have no option but to stop using fossil fuels as an energy source very soon, and we have to decide what mix of other energy sources, or other strategies, can be used to fill the gap. Some of our alternatives—wind and solar—are attractive because they are already cost effective, and are becoming more cost effective every year, but they are only available to us when the wind is blowing or the sun is shining. Wave and tidal, are more reliably available, but their development has been slow, even though the potential is significant. Large-scale hydro is attractive because it is readily dispatchable, so can be used to fill the gaps in the wind and solar supply, but there are environmental, social and practical limits to the amount that can be developed. Geothermal is an obvious choice in some regions. Nuclear fission is a reliable source of energy, but it has a bad reputation due to some serious past accidents and because of significant issues with disposal of nuclear waste. Nuclear fusion is still decades away, so is not going to be a candidate to replace fossil-fuels.

It is evident that we need to keep working on all of these options. That includes rapidly expanding solar and wind (and tidal and wave), and then ensuring that we have enough dispatchable energy (e.g., hydro) or continuous energy (e.g., nuclear) to fill the gaps. There are some other strategies that will be important to make it work, as follows.

1. 1. We—especially those of us in North America—can start by using much less energy, and that means driving less and flying less, living in smaller homes that are more energy-efficient, taking advantage of passive solar heating (and shading in hot seasons), and buying much less of the manufactured stuff that we don’t need.
   2. As shown on Figure 9.6.1, energy demand varies significantly from season to season and also from hour to hour within any day. In warm climates like California the electricity demand is lower in the winter than in the summer (mostly because of air conditioning). It is lowest in the middle of the night (1 am to 6 am) and highest in the evening (6 pm to 10 pm). There is also an afternoon low, but only in winter. In colder climates, where heating is needed through the winter, and where air conditioning is not widely used, the demand is typically higher in the cold part of the year. Electricity providers have to be very nimble to make sure that they can meet the demand at any time, but not generate more than what is needed. In the example shown for northern California, the demand ranges from a low of 7,900 MW on a Saturday afternoon in February, to a high of 15,400 MW on a Monday evening in July. Based only on this limited data, we can see that Pacific Gas and Electric (PG&E) has to be able to produce at least 16,000 MW for this market, but typically needs to produce less than 13,000 MW, and often less than 10,000 MW. If the demand curve could be smoothed out PG&E (and other utilities) could get by with less generating capacity overall and especially less peaking capacity, much of which is currently met with fossil fuel sources. Utilities can lower the peaks by charging more in the evenings (as PG&E does), and electricity users can play an important role by reducing their demand at peak times. For example, they could cook evening meals before 6 pm, or turn the a/c off between 7 and 9 and spend time outdoors on summer evenings, or not charge electric cars until after 10 pm. For most people, these types of changes would not represent a reduction in quality of life, just a change in habits.
      
      

      Figure 9.6.1 Electricity Demand for Pacific Gas and Electric Customers on Two Days in July 2020 and Two Days in February 2021
   3. Another way to even out demand, and also maximize the benefit of sources such as solar and wind, is to store energy for short periods. An example of this is the Hornsdale wind farm in Southern Australia. When electricity production exceeds demand the energy is stored in a 100 MWh Li-ion battery (the world’s largest). That stored energy is used when demand exceeds production. Some of the presently available energy-storage options are listed in Table 9.6.1. It should be noted that these options can provide for storage of energy for a few hours to a few days, and so can help to smooth out daily changes in demand, but that they do not have the capacity for longer-term storage so as to smooth out seasonal changes in demand.While the options in Table 9.6.1 are described as “utility-scale”, lithium-ion batteries can be used for household-scale energy storage to help individuals avoid peak-time electricity rates, and even smooth the curve a little. The Tesla Powerwall has software that allows the user to store energy during low-tariff periods and discharge energy during high-tariff periods, and therefore provides an opportunity to save money and help reduce utility peaks.
      

      Table 9.6.1 Utility-scale energy storage options
      Data from the Environmental and Energy Study Institute (EESI)

      Type of Storage	Efficiency	Lifetime or Number of Cycles	Limitations
      Pumped hydro	70-85%	Many decades	Topography and land needed
      Compressed air	40-70%	A few decades	Low efficiency
      Molten salt	80-90%	A few decades	Heat source needed
      Li-ion battery	85-95%	1,000-10,000 cycles	Expensive at present, limited cycles
      Flow battery	60-85%	12,000-14,000 cycles
      Hydrogen	25-45%	A few decades	Low efficiency

   4. Another way to smooth peaks in electricity demand and in supply is to share electricity over wide areas. This way, for example, solar electricity produced in the afternoon in western and central North America could be used to help supply the peak evening demands of eastern North America. This type of sharing would involve the construction of super grids with efficient high voltage direct current transmission (which suffer energy losses of less than 2% per 1000 km). One example of this type of connection is the Québec–New England transmission line which has a capacity of 2250 MW and brings hydro power from northern Québec to the eastern USA. Another is the Pacific DC Intertie that brings 3100 MW of power from the Pacific Northwest to southern California.

Media Attributions

• Figure 9.6.1 Steven Earle, CC BY 4.0, based on public domain data from U.S. Energy Information Administration (Feb. 2021) https://www.eia.gov/todayinenergy
• Table 9.6.1 Steven Earle, from data provided by Environmental and Energy Study Institute (EESI), CC BY 4.0, https://www.eesi.org/papers/view/energy-storage-2019

The main topics of this chapter can be summarized as follows:

Intrusive igneous rocks form at depths of several hundreds of metres to several tens of kilometres. Sediments are turned into sedimentary rocks only when they are buried by other sediments to depths in excess of several hundreds of metres and up to several kilometres. Most metamorphic rocks are formed at depths of kilometres to tens of kilometres. Weathering cannot even begin until these rocks are uplifted through various processes of mountain building—most of which are related to plate tectonics—and the overlying material has been eroded away and the rock is exposed as outcrop.

The important agents of mechanical weathering are as follows:

1. the decrease in pressure that results from removal of overlying rock,
2. erosional forces related to gravity, water and wind,
3. freezing and thawing of water within cracks in the rock,
4. formation of salt crystals within pores in the rock, and
5. plant roots and burrowing animals.

When a mass of rock is exposed by weathering and by removal of the overlying rock there is a decrease in the confining pressure on the rock, and the rock expands. This unloading promotes cracking of the rock, known as exfoliation, as shown in granitic rock on Figure 10.1.1.

Figure 10.1.1 Exfoliation Fractures in Granitic Rock Exposed at Yak Peak, Coquihalla Summit Recreation area, BC

Granitic rock tends to exfoliate parallel to the exposed surface because the rock is typically homogenous and it may not have pre-determined planes along which to fracture. Sedimentary and metamorphic rocks, on the other hand tend to exfoliate along predetermined planes (Figure 10.1.2).

Figure 10.1.2 Exfoliation of Slate at a Road Cut in the Columbia Mountains West of Golden, BC

Figure 10.1.3 The Process of Frost-Wedging on a Steep Slope. Water gets into fractures and then freezes, expanding the fracture a little. When the water thaws it seeps a little further into the expanded crack. The process is repeated many times, and eventually a piece of rock will be wedged away.

Frost wedging is the process by which the water seeps into cracks in a rock, expands on freezing, and thus enlarges the cracks (Figure 10.1.3). The effectiveness of frost wedging is related to the frequency of freezing and thawing. Frost wedging is most effective in a climate where there are many days each year with temperatures close to freezing, so that it might freeze overnight and then thaw in the day. In warm areas where freezing is infrequent, in very cold areas where thawing is infrequent, or in very dry areas, where there is little water to seep into cracks, the role of frost wedging is limited.

A common feature in areas of effective frost wedging is a talus slope, a fan-shaped deposit of fragments removed by frost wedging (and other mechanical weathering) from the steep rocky slopes above (Figure 10.1.4).

Figure 10.1.4 An Area With Very Effective Frost-Wedging Near to Keremeos, BC. The fragments that have been wedged away from the cliffs above have accumulated in a talus deposit at the base of the slope. The rocks in this area have quite varied colours, and those are reflected in the colours of the talus.

A related process, frost heaving, takes place within unconsolidated materials on gentle slopes. In this case water in the soil freezes and expands pushing the overlying material up. Frost heaving is responsible for winter-time damage to roads in very cold regions.

When salty water seeps into rocks, and then the water evaporates on a sunny day, salt crystals grow within cracks and pores in the rock. The growth of these crystals exerts pressure on the rock and can push grains apart, causing the rock to weaken and break. There are many examples of this on rocky ocean shorelines worldwide, especially where sandstone outcrops are common (Figure 10.1.5). Salt weathering on this scale is known as honeycomb weathering. At larger scales it is called tafoni weathering. Salt weathering can also occur away from the coast, because most environments have some salt in them.

Figure 10.1.5 Honeycomb Weathering of Sandstone on Gabriola Island, BC. The holes are caused by crystallization of salt within rock pores, and the seemingly regular pattern is related to the original roughness of the surface. It’s a positive-feedback process because the holes collect salt water at high tide, and so the effect is accentuated around existing holes. This type of weathering is most pronounced on south-facing sunny exposures.

The effects of plants and animals are significant in mechanical weathering. Roots can force their way into even the tiniest cracks, and then they exert tremendous pressure on the rocks as they expand, widening the cracks and breaking the rock. Although animals do not normally burrow through solid rock, they can excavate and remove huge volumes of soil, and thus expose the rock to weathering by other mechanisms.

Mechanical weathering is greatly facilitated by erosion, which is the removal of weathering products, allowing for the exposure of more rock for weathering. A good example of this is Figure 10.1.4. On the steep rock faces above at the top of the cliff rock fragments are broken off by ice-wedging, and then removed by gravity. This is a form of “mass wasting”, which is discussed in more detail in Chapter 5. Other important agents of erosion, that also have the effect of removing the products of weathering, include water in streams (Chapter 11), ice in glaciers (Chapter 4), wind in deserts, and waves on the coasts.

Exercise 10.1 Mechanical Weathering

Figure 10.1.6 shows granitic rock at the top of Stawamus Chief near to Squamish, BC. Identify the mechanical weathering processes that you can see are taking place, or you think probably take place at this location.

Figure 10.1.6 The Summit of Stawamus Chief Near to Squamish, BC

Exercise answers are provided Appendix 2.

Media Attributions

• Figure 10.1.1 Photo by Steven Earle, CC BY 4.0
• Figure 10.1.2 Photo by Steven Earle, CC BY 4.0
• Figure 10.1.3 Steven Earle, CC BY 4.0
• Figure 10.1.4 Photo by Steven Earle, CC BY 4.0
• Figure 10.1.5 Photo by Steven Earle, CC BY 4.0
• Figure 10.1.6 Photo by Isaac Earle, CC BY 4.0

Chemical weathering results from the chemical changes to some minerals that become unstable when they are exposed to surface conditions. The kinds of changes that take place are highly specific to the mineral and to the environmental conditions. Some minerals, like quartz, are virtually unaffected by chemical weathering, while others, like feldspar, are quite easily altered. In general, the degree of chemical weathering is greatest in warm and wet climates, and least in cold and dry climates. The important characteristics of surface conditions that lead to chemical weathering are: the presence of water (in the air and on the ground surface), the abundance of oxygen, and the presence of carbon dioxide, which, when combined with water, produces weak carbonic acid. That process, which is fundamental to most chemical weathering, can be shown as follows:

H₂O + CO₂ -> H₂CO₃

Here we have water (e.g., as rain) plus carbon dioxide in the atmosphere, combining to create carbonic acid. The amount of CO₂ in the air is enough to make only very weak carbonic acid, but there is typically much more CO₂ in the soil, so water that percolates through the soil can become significantly more acidic.

There are two main types of chemical weathering. On the one hand, some minerals become altered to other minerals. For example, feldspar is altered—by hydrolysis—to clay minerals. On the other hand, some minerals dissolve completely, and their components go into solution. An example is calcite (CaCO₃) which is soluble in acidic solutions.

The hydrolysis of feldspar can be written like this:

CaAl₂Si₂O₈ + H₂CO₃ + ½O₂ -> Al₂Si₂O₅(OH)₄ + Ca²⁺ + CO₃^2-
(plagioclase + carbonic acid + oxygen -> kaolin + calcium & carbonate ions in solution)

This reaction shows calcium plagioclase feldspar, but similar reactions could also be written for sodium or potassium feldspars. In this case we end up with the mineral kaolinite along with calcium and carbonate ions in solution. Those ions can eventually combine to form the mineral calcite, and that will probably happen in the ocean. The hydrolysis of feldspar to clay is illustrated on Figure 10.2.1, which shows two images of the same granitic rock, a recently broken fresh surface on the left, and a clay-altered weathered surface on the right. Other silicate minerals can also go through hydrolysis, although the end results will be a little different. For example, pyroxene can be converted to the clay minerals chlorite or smectite, or olivine can be converted to the clay mineral serpentine.

Figure 10.2.1 Un-weathered (left) and Weathered (right) Surfaces of the Same Piece of Granitic Rock. On the unweathered surfaces the feldspars are still fresh and glassy-looking. On the weathered surface the feldspar has been altered to the chalky-looking clay mineral kaolinite.

Oxidation is another very important chemical weathering process. The oxidation of the iron in a ferromagnesian silicate starts with the dissolution of the iron. For olivine the process looks like this, where olivine in the presence of carbonic acid is converted to dissolved iron, carbonate and silicic acid:

Fe₂SiO₄ + 4H₂CO₃ -> 2Fe²⁺ + 4HCO₃^– + H₄SiO₄
(olivine + carbonic acid -> iron, carbonate & silica ions in solution)

The equation is shown here for olivine, but could apply to almost any other ferromagnesian silicate, including pyroxene, amphibole and biotite. Iron in sulphide minerals (e.g., pyrite) can also be oxidized in this way. And the mineral hematite is not the only possible end result either, as there is a wide range of iron-oxide minerals that can form in this way. The results of this process are illustrated on Figure 10.2.2, which shows a granitic rock in which some of the biotite and amphibole have been altered to form the iron oxide mineral limonite.

Figure 10.2.2 A Granitic Rock Containing Biotite and Amphibole, Altered Near to the Rock’s Surface to Limonite, A Mixture of Iron-Oxide Minerals.

As we’ve seen above in Chapter 8, a special type of oxidation takes place in areas where the rocks have elevated levels of sulphide minerals, especially pyrite (FeS₂). Pyrite will react with water and oxygen to form sulphuric acid, as follows:

2FeS₂ + 7O₂ + 2H₂O -> Fe²⁺ + 2H⁺ + H₂SO₄
(pyrite +m oxygen -> iron and hydrogen ions in solution + sulphuric acid)

The runoff from areas where this process is taking place is known as acid rock drainage (ARD), and even a rock with 1 or 2% pyrite can produce significant ARD. Some of the worst examples of ARD are at metal mine sites, especially where pyrite-bearing rock and waste material has been mined from deep underground and then piled up and left exposed to water and oxygen.

The hydrolysis of feldspar, and of other silicate minerals as well, and the oxidation of iron in ferromagnesian silicates all serve to create rocks that are softer and weaker than they were to begin with, and thus more susceptible to mechanical weathering.

The weathering reactions that we’ve discussed so far involved the transformation of one mineral to another mineral (e.g., feldspar to clay), and the release of some ions in solution (e.g., Ca²⁺). Some weathering processes involve the complete dissolution of a mineral. Calcite, for example, will dissolve in weak acid, to produce calcium and bicarbonate ions. The equation is as follows:

CaCO₃ +  H₂CO₃ -> Ca²⁺ + 2HCO₃^–
(calcite + carbonic acid -> calcium + bicarbonate ions in solution)

Calcite is the major component of limestone (typically more than 95%), and under surface conditions limestone will dissolve to varying degrees (depending on which minerals it has other than calcite), as shown on Figure 10.2.3. Limestone also dissolves at relatively shallow depths underground, forming limestone caves. This is discussed in more detail in Chapter 14, where we look at groundwater.

Figure 10.2.3 A Limestone Outcrop on Quadra Island, BC. The limestone, which is primarily made up of the mineral calcite, has been dissolved to different degrees in different areas because of compositional differences. The buff-coloured bands are either volcanic rock or chert, which are not soluble.

Products of Weathering and Erosion

The products of weathering and erosion are the unconsolidated materials that we find around us on slopes, beneath, beside and on top of glaciers, in stream valleys, on beaches, and in deserts. The nature of these materials—their composition, size, degree of sorting, and degree of rounding—is determined by the type of rock that is being weathered, the nature of the weathering, the erosion and transportation processes, and the climate.

In addition to these solid sediments, the other important products of weathering are many different types of ions in solution.

A summary of the weathering products of some of the common minerals present in rocks is provided in Table 10.1.1. In addition to the weathering products listed in the table, most of the larger fragments—larger than sand grains—that make up sediments will be pieces of rock as opposed to individual minerals.

Some examples of the products of weathering are shown in Figure 10.2.4. They range widely in size and shape depending on the processes involved in their transportation. If and when deposits like these are turned into sedimentary rocks, the textures of those rocks will vary significantly. Importantly, when we describe sedimentary rocks that formed millions of years in the past, we can use those properties to make inferences about the conditions that existed during their formation.

Figure 10.2.4 Products of Weathering and Erosion Formed Under Different Conditions.

It’s worth considering here why the sand-sized sediments shown in Figure 10.2.4 are so strongly dominated by the mineral quartz, even though quartz makes up less than 20% of Earth’s crust. The explanation is that quartz is highly resistant to the weathering that occur at Earth’s surface. It is not affected by weak acids, or water or the presence of oxygen. This makes it unique among the minerals that are common in igneous rocks. Quartz is also very hard, and doesn’t have cleavage, so it is resistant to mechanical erosion.

When a rock like granite is subject to chemical weathering the feldspar and the ferromagnesian silicates get converted to clays plus dissolved ions such as: Ca²⁺, Na⁺, K⁺, Fe²⁺, Mg²⁺, and H₄SiO₄, but the quartz is resistant to those processes and remains intact. The clay gradually gets eroded away, then the rock breaks apart leaving lots of grains of quartz. In other words, quartz, clay minerals, and dissolved ions are the most common products of weathering. Quartz and some of the clay minerals tend to form sedimentary deposits on and at the edges of continents, while the rest of the clay minerals and the dissolved ions tend to be washed out into the oceans to form sediments on the sea floor.

Exercise 10.3 The Weathering Origins of Sand

A number of different sands are pictured and described in the following table. Describe some of the important weathering processes that might have led to the development of these sands.

Image	Description and location	Important weathering process?
	Fragments of coral, algae, and urchin from a shallow water area (roughly 2 m deep) near a reef in Belize. The grain diameters are between 0.1 and 1 mm.
	Angular quartz and rock fragments from a glacial stream deposit near Osoyoos, BC. The grain diameters are between 0.25 and 0.5 mm.
	Rounded grains of olivine (green) and volcanic glass (black) from a beach on the big island of Hawaii. The grains are approximately 1 mm across.

(Photos by Steven Earle, CC BY 4.0)

Media Attributions

• Figure 10.2.1 Photo by Steven Earle, CC BY 4.0
• Figure 10.2.2 Photo by Steven Earle, CC BY 4.0
• Figure 10.2.3 Photo by Steven Earle, CC BY 4.0
• Figure 10.2.4 Photos by Steven Earle, CC BY 4.0

Weathering is a key part of the process of soil formation, and soil is critical to our existence on Earth. In other words, we owe our existence to weathering, and we need to take care of our soil!

Many people refer to any loose material on Earth’s surface as soil, but to geologists (and geology students) soil is the material that includes organic matter, lies within the top few tens of centimetres of the surface, and is important in sustaining plant growth. Other types of soft sediments can be described using terms that are indicative of how they formed, such as glacial till, river-deposited sand and gravel, lake-deposited clay, and wind-blown silt.

Soil is a complex mixture of minerals (approximately 45%), organic matter (approximately 5%), and empty space (approximately 50%, filled to varying degrees with air and water). The mineral content of soils is variable, but is dominated by clay minerals and quartz, along with minor amounts of feldspar and small fragments of rock. The types of weathering that take place within a region have a major influence on soil composition and texture. For example, in a warm climate, where chemical weathering dominates, soils tend to be richer in clay. Soil scientists describe soil texture in terms of the relative proportions of sand, silt, and clay, as shown in Figure 10.3.1. The sand and silt components in this diagram are dominated by quartz, with lesser amounts of feldspar and rock fragments, while the clay component is dominated by the clay minerals.

Figure 10.3.1 Variations in the Proportions of Clay-, Silt-, and Sand-Sized Fragments in Soils. This diagram applies only to the mineral component of soils, and the names are textural descriptions, not soil classes.

Soil forms through accumulation and decay of organic matter and through the mechanical and chemical weathering processes described above. The factors that affect the nature of soil and the rate of its formation include climate (especially average temperature and precipitation amounts), organisms (especially the types and intensity of vegetation), relief (the slope and aspect of the surface) the type of parent material, and the amount of time available. These are known by the acronym CLORPT, coined by Hans Jenny in 1941.

Climate

Soils develop because of the weathering of materials on Earth’s surface, including the mechanical breakup of rocks, and the chemical weathering of minerals. Soil development is facilitated by the downward percolation of water. Soil forms most readily under temperate to tropical conditions (not cold) and where precipitation amounts are moderate (not dry, but not too wet). Chemical weathering reactions (especially the formation of clay minerals) and biochemical reactions proceed fastest under warm conditions, and plant growth is enhanced in warm climates. Too much water (e.g., in rainforests) can lead to the leaching of important chemical nutrients and hence to acidic soils. In humid and poorly drained regions, swampy conditions may prevail, producing soil that is dominated by organic matter. Too little water (e.g., in deserts and semi-deserts), results in very limited downward chemical transportation and the accumulation of salts and carbonate minerals (e.g., calcite) from upward-moving water. Soils in dry regions also suffer from a lack of organic material (Figure 10.3.2).

Figure 10.3.2 Poorly Developed Soil on Wind-Blown Silt (Loess) in an Arid Part of Northeastern Washington State. The thickness shown is about 1 m, and only the upper 2 or 3 cm are actual “soil”.

Organic matter

Good soil is rich in organic matter, and that can only accumulate if there is sufficient plant growth in the area. There is only sparse vegetation growing in the area of Figure 10.3.2, and the soil is poor.

Relief

Soil can only develop where surface materials remain in place and are not frequently moved away by mass wasting. Soils cannot develop where the rate of soil formation is less than the rate of erosion, so steep slopes tend to have little or no soil.

Parent Material

Soil parent materials can include all different types of bedrock and any type of unconsolidated sediments, such as glacial deposits and stream deposits. Soils are described as residual soils if they develop on bedrock and transported soils if they develop on transported material such as glacial sediments. Other sources may use the term “transported soil” to imply that the soil itself has been transported, but in this text “transported soil” is soil that is developed on transported materials, like the very thin soil shown in Figure 10.3.2. When referring to such soil, it is better to be specific and say “soil developed on unconsolidated material,” because that distinguishes it from soil developed on bedrock.

Quartz-rich parent material, such as granite, sandstone, or loose sand, leads to the development of sandy soils. Quartz-poor material, such as shale or basalt, generates soils with little sand, and in some cases with elevated clay levels.

Parent materials provide important nutrients to residual soils. For example, a minor constituent of granitic rocks is the calcium-phosphate mineral apatite (Ca₅(PO₄)₃(F,Cl,OH)), which is a source of the important soil nutrient phosphorus. Basaltic parent material tends to generate very fertile soils because it also provides phosphorus, along with significant amounts of iron, magnesium, and calcium. As described in Chapter 7, soils in active volcanic regions tend to be generally quite fertile, partly because the fine-grained and glass-rich eruption products weather quickly and produce significant amounts of clay that hold trace elements.

Some unconsolidated materials, such as river-flood deposits, make for especially good soils because they also tend to be rich in clay minerals. Clay minerals have large surface areas with negative charges that are attractive to positively charged elements like calcium, magnesium, iron, and potassium—important nutrients for plant growth.

Time

Even under ideal conditions, soil takes thousands of years to develop. Virtually all of southern Canada and parts of the northern US were still glaciated up until 14 ka. Glaciers still dominated the central and northern parts of Canada until around 10 ka, and so, at that time, conditions were still not ideal for soil development even in the southern regions. Therefore, soils in Canada, and especially in central and northern Canada, are relatively young and not well developed.

The same applies to soils that are forming on newly created surfaces, such as recent deltas or sand bars, or in areas of mass wasting. Climate and parent material both exert a significant control on how long it takes for soil to develop.

Soil formation

The process of soil formation is summarized on Figure 10.3.3. The left-hand diagram shows the situation of the site within several years of the exposure of the rock surface (for example, following deglaciation). The rock is fractured and both mechanical and chemical weathering have started. Lichen and moss are present on the rock surface and small plants have started to grow in cracks and depressions in the outcrop where small amounts of sediment have accumulated.

The second diagram represents a time some hundreds to thousands of years later (longer in cold and dry climates). Both mechanical and chemical weathering of the rock are well advanced. The rock is softer and weaker than it was originally and the weathering products have accumulated in between the remaining rock fragments. The C horizon soil has developed from the products of weathering (small rock fragments, sand, and clay) and organic matter is accumulating near to surface.

The third diagram is a few thousand to several thousand years later. The soil profile has started to evolve, mostly through chemical changes involving both downward and upward motion of ions in water, and transfer of water and chemicals by the roots of plants and mycorrhizal networks.

The last diagram, on the right, represents a time several thousands to tens of thousands of years later.  The soil is now well developed. There has been significant weathering of minerals like feldspar and amphibole within the soil to produce clay minerals and there has also been upward and downward movement of chemicals (iron, manganese, potassium, sodium, calcium, magnesium, and aluminum). Along with all of the living organisms (roots, mycorrhizae, worms, insects), there is a great deal of carbon stored in the soil (as organic matter, charcoal, carbon dioxide and methane). Depending on the type of parent material and the rate of soil formation in the local climate, the soil in this last diagram could be from several centimetres to over a metre thick.

Figure 10.3.3 Steps in the Evolution of Soil

Soil Horizons

The process of soil formation generally involves the downward movement of clay, water, and dissolved ions, and a common result of that is the development of chemically and texturally different layers known as soil horizons. The typically developed soil horizons, as illustrated in Figure 10.3.4, are:

O – a layer of organic matter,
A – a layer of partially decayed organic matter mixed with mineral material
E – an eluviated (leached) layer from which some of the clay and iron have been removed to create a pale layer that may be sandier than the other layers
B – a layer of accumulation of clay, iron, and other elements from the overlying soil
C – a layer of incomplete weathering, which grades down into unaltered parent material

Figure 10.3.4 Soil Horizons in a Podsol From a Site in Northeastern Scotland.

Like all geological materials, soil is subject to erosion, although under natural conditions on gentle slopes, the rate of soil formation either balances or exceeds the rate of erosion. Human practices, especially those related to forestry and agriculture, have significantly upset this balance in many places.

Soils are held in place by vegetation. When vegetation is removed, either through cutting trees or routinely harvesting crops and tilling the soil, that protection is either temporarily or permanently lost.

The primary agents of the erosion of unprotected soil are water and wind. Water erosion is accentuated on sloped surfaces because fast-flowing water obviously has greater eroding power than slow-flowing or still water (Figure 10.3.5). Raindrops can disaggregate exposed soil particles, putting the finer material (e.g., clays) into suspension in the water. Sheet wash—unchanneled flow across a surface—carries suspended material away, and channels erode right through the soil layer, removing both fine and coarse material.

Soil Erosion

Figure 10.3.5 Soil Erosion by Rain, Sheet Wash and Channeled Runoff on a Field in Alberta

Wind erosion is exacerbated by the removal of trees that act as wind breaks and by agricultural practices that leave bare soil exposed (Figure 10.3.6).

Tillage is also a factor in soil erosion, especially on slopes, because each time the soil is lifted by a cultivator, it is moved a few centimetres down the slope.

Figure 10.3.6 Soil Erosion by Wind in Alberta

Media Attributions

• Figure 10.3.1 USDA Soil Texture by Christopher Aragón, CC BY SA 4.0, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:USDA_Soil_Texture.svg
• Figure 10.3.2 Photo by Steven Earle, CC BY 4.0
• Figure 10.3.3 Steven Earle, CC BY SA 4.0
• Figure 10.3.4 Modified photo by Steven Earle, CC BY SA 4.0, from Podzol by Ailith Stewart, via geograph.org, CC BY SA 2.0, https://www.geograph.org.uk/photo/218892
• Figure 10.3.5 Image from Alberta Agriculture and Rural Development, http://www1.agric.gov.ab.ca/$department/deptdocs.nsf/all/agdex9313. Used with permission, all rights reserved.
• Figure 10.3.6 Image from Alberta Agriculture and Rural Development, http://www1.agric.gov.ab.ca/$department/deptdocs.nsf/all/agdex9313. Used with permission; all rights reserved.

Up until the 1950s, the classification of soils in Canada was based on the system used in the United States. However, it was long recognized that the U.S .system did not apply well to many parts of Canada because of climate and environmental differences. The Canadian System of Soil Classification was first outlined in 1955 and has been refined and modified numerous times since then.

There are 10 orders of soil recognized in Canada. Each one is divided into groups, and then families, and then series, but we will only look at the orders, some of which are summarized in Table 10.4.1.

The distribution of these types of soils (and a few others) in Canada is shown in Figure 10.4.1.

Figure 10.4.1 The Soil Order Map of Canada

There is an excellent website on Canadian soils, with videos describing the origins and characteristics of the soils, at: Soil Classification: Soil Orders of Canada.

As we’ve discussed, the processes of soil formation are dominated by the downward transportation of clays and certain elements dissolved in water, and the nature of those processes depends in large part on the climate. In Canada’s predominantly cool and humid climate (which applies to most places other than the far north), podsolization is the norm. This involves downward transportation of hydrogen, iron, and aluminum (and other elements) from the upper part of the soil profile, and accumulation of clay, iron, and aluminum in the B horizon. Most of the podsols, luvisols, and brunisols of Canada form through various types of podsolization.

In the grasslands of the dry southern parts of the prairie provinces dark brown organic-rich chernozem soils are dominant. In some parts of these areas, weak calcification takes place with leaching of calcium from the upper layers and accumulation of calcium in the B layer.

Organic soils form in areas with poor drainage (i.e., swamps) and a rich supply of organic matter. These soils have very little mineral matter.

In the permafrost regions of the north, where glacial retreat was most recent, the time available for soil formation has been short and the rate of soil formation is very slow. The soils are called cryosols (cryo means “ice cold”). Permafrost areas are also characterized by the churning of the soil by freeze-thaw processes, and as a result, development of soil horizons is very limited.

Media Attribution

• Figure 10.4.1 Image from the Department of Soil Science, University of Saskatchewan. Used with permission, all rights reserved.

In Earth science the word “clay” has two meanings. Clay is broadly defined as any unconsolidated material with a grain diameter less than 0.004 mm.  That is about 1/100th as big as the period at the end of this sentence, so is not big enough to see with the naked eye. This could include finely ground up quartz, feldspar, calcite, hematite or any other mineral. “Clay” also refers to the clay minerals, which are sheet silicates or phyllosilicates (from the Greek word phyllo meaning leaf). Clay minerals typically only exist as very tiny crystals, so most true clays also conform to the fine-grained meaning of the word “clay”.

A clay mineral is a finely crystalline sheet silicate with hydroxyl ions, and in some cases with water as part of the structure. By sheet silicate we mean that the silica tetrahedra are arranged in flat sheets, with strong covalent bonding within the sheets, and that these sheets are arranged in stacks, where the bonding between the sheets is relatively weak. A hydroxyl ion is an oxygen-hydrogen pair (OH^–), and these form a part of all clay minerals. Some clay minerals also have H₂O as part of the structure, or in some cases water is simply attached onto the structure.

Most of the clay present in rocks at surface has formed as a result of weathering of other silicate minerals, primarily feldspars, micas, pyroxene and amphibole. The reactions involved are hydrolysis reactions, something like the following reaction of potassium feldspar plus water and carbon dioxide to kaolin.

2KAlSi₃O₈ + 2CO₂ + 11H₂O —> Al₂Si₂O₅(OH)₄ + 2K⁺ + 4H₄SiO₄ + 2HCO₃^–

K-feldspar + carbon dioxide + water ->  kaolin + potassium, silica and bicarbonate ions in solution

In simple terms, K-feldspar reacts with water and carbon dioxide to form kaolin plus bicarbonate ions. The potassium and some of the silicon that were originally present in the feldspar, are removed in solution.  The CO₂ comes from the atmosphere, and over geological time, this is type of reaction plays an important role in controlling the atmosphere’s composition and hence the greenhouse effect.

Clay minerals can also be formed when hot waters (known as hydrothermal solutions) circulate through a body of rock. As is the case for weathering, the hot solutions lead to alteration of pre-existing minerals. Hydrothermal solutions are often also associated with the formation of metal deposits (such as porphyry copper deposits) and the surrounding clay-mineral halos can be an important guide in the exploration for such deposits.

Unlike the primary silicate minerals that they form from, clay minerals are soft and easily eroded into tiny fragments and then transported.  They accumulate mostly as sediments in low-energy deposition environments (e.g., deep ocean or in lakes), sediments that are eventually turned into shale.

Clay Mineral Structures

Clay minerals are comprised of silica tetrahedra and alumina octahedra, which are illustrated on Figure 10.5.1. As we’ve seen in Chapter 2 (Figure 2.1.5), a silica tetrahedron is a silicon ion surrounded by four oxygen ions. Planes drawn through lines connecting the oxygens atoms define a tetrahedral (four-surfaced) shape. An alumina octahedron is an aluminum ion surrounded by six oxygen or hydroxyl ions. Planes drawn through lines connecting the oxygens and hydroxyl atoms define an octahedral (eight-surfaced) shape.

Figure 10.5.1 Representations of the Silica Tetrahedra and Aluminum Octahedra That Combine to Form Clay Minerals

An important feature of clay minerals results from the characteristics of their bonding. The tetrahedra and octahedra are strongly bonded to each other within the sheets, but the sheets are only weakly bonded one to another. The sheets that make up a clay mineral grain have a tendency to slide with respect to each other, and the result is that clay mineral masses tend to be soft and plastic, and not very strong.

The simplest clay mineral is kaolin. Each “sheet” within the kaolin structure is comprised of a silica tetrahedral layer and an aluminum octahedral layer (Figure 10.5.2). The combination of one tetrahedral layer and one octahedral layer makes this a 1:1 layer silicate.  For simplicity, it may be useful to describe this as a T-O structure (1 tetrahedral layer and 1 octahedral layer). This structure is also found in the mineral serpentine, in which magnesium substitutes for aluminum in the octahedral sites.

Figure 10.5.2 A Representation of the Tetrahedral-Octahedral Layer Structure of the 1:1 Clay Mineral Kaolin. The layers are held together with weak van der Waals bonds.

The structure of illite is more complicated than that of kaolin. In this case each “sheet” within the structure is comprised of an aluminum octahedral layer sandwiched between two tetrahedral layers (one “right side up” and the other “up-side down”) (Figure 10.5.3). Illite also has potassium ions situated at specific sites between the sheets.  The combination of two tetrahedral layers surrounding one octahedral layer is known as a 2:1 layer silicate. We can also describe this as a T-O-T structure.  Some other T-O-T clays include smectite, talc and chlorite.

Figure 10.5.3 A Representation of the Tetrahedral-Octahedral-Tetrahedral Structure of the 2:1 Clay Mineral Illite

The potassium cations (K⁺) of illite are held in place because the upper and lower surface of each layer is saturated with oxygen ions (O^-2) giving these surfaces a consistent negative charge.  These negatively charged surfaces are one of the fundamentally important features of clay minerals, because such surfaces are attractive to positively charged ions, such as heavy metals, or some organic pollutants. Not only do clays have these attractive surfaces, but they have very large surface areas. It is estimated that a cubic centimetre of clay has a reactive surface of around 2800 square metres, which is equivalent to the area of a football field!

Figure 10.5.4 Scanning Electron Micrograph Showing Plates of the Clay Mineral Kaolin

Some kaolin crystals are shown Figure 10.5.4. The kaolin plates are stacked up in loosely defined strands.  Each visible plate is between 1/10th and 1/5th of a micron thick, but each of these is made up of hundreds to thousands of the T-O layers illustrated in Figure 10.5.2. There is a significant amount of empty space between the plates.

Although there are many dozens of different clay minerals, there are just a handful that are important for us to be aware of here, and these are summarized in Table 10.3. The first thing to note is that there are only two 1:1 (T-O) clay minerals in this list, kaolin and serpentine, while all of the others are 2:1 clays (T-O-T). Most of these 2:1 clays have magnesium and/or iron in them, and so can be considered ferromagnesian silicates. The only non-ferromagnesian clays listed are kaolin, pyrophyllite and Illite, although Illite can also have small amounts of magnesium, and the glauconite variety has iron. The important point here is that while kaolin, pyrophyllite and Illite are typically derived from alteration of minerals like feldspar and muscovite, most of the other clays are derived from alteration of minerals like olivine, pyroxene, amphibole and biotite.

Table 10.5.1  Some Important Clay Minerals, Their Chemical Formulas and Variations

Clay Mineral	Type	Typical Chemical Formula	Variations and (other names)
Kaolin	1:1	Al2Si2O5(OH)4	kaolinite, dickite, halloysite, nacrite
Serpentine	1:1	Mg3Si2O5(OH)4	antigorite, chrysotile (asbestos), lizardite
Illite	2:1	K0.65Al2.0(Al0.65Si3.35O10)(OH)2	glauconite, (hydromuscovite, K-deficient muscovite)
Pyrophyllite	2:1	Al2Si4O10(OH)2
Smectite	2:1	(Na, Ca)0.33(Al, Mg)2(Si4O10)(OH)2 ∙ nH2O	montmorillonite (bentonite), saponite, nontronite
Vermiculite	2:1	(Mg, Fe2+, Fe3+)3 ((Al, Si)4O10)(OH)2 ∙ 4H2O
Talc	2:1	Mg3Si4O10(OH)2
Chlorite	2:1	(Mg, Fe)3(SI, Al)4O10(OH)2 ∙ (Mg, Fe)3(OH)6	clinochlore, pennantite, chamosite, sudoite

Formation of Clay Minerals

As already noted, clay minerals typically form from the alteration (hydrolysis) of pre-existing silicate minerals. The type of clay mineral that will form in any situation depends partly on what silicate mineral is being altered, but also on a range of other variables such as the temperature and pressure, and the chemistry of the solutions that are passing through or over the rock at the time.

Clay minerals form during weathering at surface, during hydrothermal alteration of rock within the crust, and during the diagenesis (mineral alterations that take place when sediments get buried beneath other sediments) and their transformation into sedimentary rock.

While the temperatures and pressures of hydrothermal alteration and diagenesis can vary dramatically, weathering conditions are broadly similar the world over, the main differences being the amount water available from precipitation and the average temperatures. Temperature differences of a few tens of degrees mainly control the rate of weathering, not the type, and so the major factor that determines which clay minerals will form during weathering is the type of primary silicate minerals present in the rock.

A summary of the clay products of weathering of primary silicate minerals is provided in Table 10.4. Quartz in not in this list because it isn’t subject to chemical weathering.

The clay mineral products listed Table 10.4 are specifically those that form under weathering conditions, which generally means temperatures under 40° C, atmospheric pressure, and water that has low levels of dissolved ions and close to neutral pH.

The higher temperatures associated with burial and diagenesis of sediments, or with hydrothermal alteration or even low grade metamorphism, result in the formation of some clays that are not typically produced during weathering.

Some of the clay-mineral transformations that can take place within sediments as they undergo progressively greater burial beneath other sediments are illustrated on Figure 10.5.5.  The assumption here is that the sediments already include some clay minerals, especially the low-temperature clays smectite and kaolin produced during weathering in the sediment source area. Starting at around 100° C, the smectite might first be altered to a mineral with mixed or alternating layers of smectite and illite. The mixed-layer clays may be altered to chlorite and illite at around 150° C, with the illite being altered to muscovite at over 200° C.  Any kaolin originally present in the sediment as kaolinite or halloysite might first get transformed to the higher temperature polymorphs (dickite or nacrite) and then to illite, and eventually to either chlorite or muscovite.

Figure 10.5.5  The Transformation of Clay Minerals During Burial Diagenesis of Sediments and Sedimentary Rocks. The temperature versus depth relationship will be variable, depending on the geothermal gradient in a particular area.

Hydrothermal alteration takes place where hot water circulates at depths of hundreds to thousands of metres within the crust. This is commonly associated with convection systems produced by magmatic heat, and as described in Chapter 8, it is a process that is commonly associated with ore-formation. Clay alteration is well known around porphyry, epithermal, volcanogenic massive sulphide and some uranium deposits. In these environments water chemistry can be extremely variable, and temperatures can range up to hundreds of degrees, so a very wide range of clay minerals and other alteration minerals can form. The details are beyond the scope of this book but can be found in works related to mineral deposits.

Hydrothermal alteration to clay minerals also takes place in volcanic areas because the heat source can drive convection and will also speed up the reaction rates. This is evident around Mount Meager area in BC’s Coast Range (Figure 10.5.6).  Some of the rock colours are a product of clay alteration. The scar from the 2010 rock slide and rock avalanche—which happened in part because the volcanic rock had been weakened by hydrothermal clay-alteration—is visible in the photo.

Figure 10.5.6 Pylon Peak (2481 m, left, behind cloud) and Mt. Meager (2650 m, centre-right). The bare patch to the right of Mt. Meager is the source area of Canada’s largest slope failure in historical times, the 2010 rock slide and rock avalanche.

One of the most important sites of clay alteration is at oceanic divergent boundaries where there is volcanic heat resulting in convective water circulation. Oceanic crustal rocks include basalt and gabbro and even ultramafic rock at greater depth, so they are rich in pyroxene and olivine, and these minerals are readily altered to minerals like chlorite and serpentine at temperatures of a few hundred degrees. An example of that process is as follows:

2Mg₂SiO₄ + 3H₂O → Mg₃Si₂O₅(OH)₄ + Mg(OH)₂

(olivine + water -> serpentine + brucite (not a clay mineral)

The resulting rock may look like that shown in Figure 10.5.7. The pyroxene and olivine of oceanic crust may also be altered to smectite (at low temperatures) and to talc and chlorite. A significant proportion of oceanic crust has been altered in this way, and since oceanic crust makes up about 70% of the Earth’s crust, this is likely the largest repository of clay minerals on the planet. That is significant from an Earth-systems perspective because when oceanic crust is eventually subducted the clay minerals are heated and converted (by metamorphism) back to non-hydrous silicate minerals and the water is released. This water contributes to partial melting above a subduction zone, and therefore to composite volcanoes along subduction boundaries.

Figure 10.5.7  Serpentine-Bearing Rock from the Vermont Verde Antique Quarry, Vermont

Exercise 10.5 Clay Mineral Origins

Some clay minerals are listed below. Indicate the environment (e.g., weathering, diagenesis, hydrothermal alteration) in which it likely formed, and a possible precursor mineral.

Table 10.5.3 Clay Minerals and the Environment of Formation with Possible Precursor (Exercise)

Mineral	Likely Environment of Formation and Possible Precursor
Halloysite
Mixed-layer clay
Serpentine
Illite
Dickite

Exercise answers are provided Appendix 2.

Properties of Clay Minerals

It is worthwhile to understand some of the properties of clay minerals, as they have important implications for many aspects of environmental geology. They play a role in a great variety of processes, from soil chemistry to the causes and effects of earthquakes, to the permeability of rocks. Some of their important properties are as follows:

They are soft and weak, primarily because of the weak bonds between sheets and the resulting tendency for the sheets to slide past each other under stress. Talc is number 1 on the Mohs scale, and most other clay minerals are similarly soft. The weakness of clay minerals has implications for slope failure (as noted above) because clay bearing rocks also tend to be weak, and also for earthquakes, because a plate boundary with clay-rich rocks is likely to slide smoothly, and so less likely to stick and cause large earthquakes.

Most clays are malleable when wet—also because of weak inter-layer bonds—and so can easily be formed into useful shapes for artistic, domestic, industrial and scientific purposes.

Clay minerals are crystals like other minerals, but they typically only form as very small crystals, so clay deposits are almost universally fine grained. Although a body of clay has significant porosity, the pores are extremely small and most of the water within them is close enough to a grain boundary to be held tightly by surface tension, making a clay deposit significantly impermeable.  This has implications for groundwater flow (Chapter 11) and for waste disposal (Chapter 13).

The tetrahedra and octahedra that make up clay minerals have negatively charged ions (anions) on their outsides (either O^2- or OH^–) making the surfaces of the individual layers negatively charged, and therefore attractive to positively charged ions (cations) in solution. Most metals exist as cations and many organic pollutants have positive charges, and so clay minerals are efficient scavengers of environmental pollutants, and can be used as barriers to prevent dispersal of contaminants and also in environmental rehabilitation projects. Different clays have different capacities to absorb cations (known as “cation exchange capacity”), and some of these are listed in Table 10.5. Smectite has a much higher cation exchange capacity than other clays because cations can get onto the sites in between the molecular layers within a crystal, as opposed to just the outside surfaces of the crystals.

Table 10.5.4 Surface Area and Cation Exchange Capacity of Some Clay Minerals
*Meq/g is milli-equivalents per gram, or milli-moles. E.g., 10 meq/g of Cu = 0.29 g of Cu per g of clay

–	Effective Surface Area in m2/g	–	Cation Exchange Capacity
Mineral	Interlayer	External	Meq/g*
Kaolin	0	15	1 to 10
Chlorite	0	15	<10
Illite	5	15	10 to 40
Smectite	750	50	80 to 150

The smectite clays have a 2:1 structure similar to that of illite (Figure 10.4.3) but the interlayer cations are typically either sodium, calcium or magnesium (instead of potassium).  This means that the layers are a little further apart than in illite, and for that reason smectites have a unique ability to absorb water molecules in the interlayer sites. This is especially true for sodium smectites, which can absorb up to 18 layers of water in between the sheets, and thereby will expand or swell dramatically when wet. Some swelling clay is shown on Figure 10.5.8.  The clay is present within a depression and was wet.  On drying it shrank in the typical mudcrack pattern. Swelling clays have a number of important industrial and domestic uses, but they also have some serious geological implications. A swollen wet smectite is even weaker than a dry one, so can weaken slopes significantly, and bodies of swollen clay can distort the materials around them, also potentially contributing to slope failure or problems for building or road foundations.

Figure 10.5.8 Smectite-Bearing Clay in the volcanic Krafla Area, Iceland.  On drying the clay shrank and the decrease in volume was accommodated by cracking.

Vermiculite will also swell slightly when wet, but, unlike other clays it will expand dramatically on heating (Figure 10.5.9).  When heated to 500 to 800° C the water trapped between the layers will boil and push the layers apart increasing the volume dramatically.  Expanded vermiculite has many uses, including as a growing medium, insulation, brake linings, and fireproof panels.

Figure 10.5.9 Expanded Vermiculite

Exercise 10.6 Find Some Clay in Your Neighbourhood

Clay minerals are everywhere.  Look around your neighbourhood or your region to see if you can find some. Likely places might include: a rock outcrop that is being weathered, a rock that has been hydrothermally altered, a fine-grained sedimentary rock (e.g., mudstone), a dried up puddle, the edge of swamp or pond, or a small bay. If you can’t find any place like those suggested, think about where there might be some clay that you can’t see such as in the middle of a lake or in the ocean.

You might also be able to find some clays at home. Look in the medicine cabinet for example, or amongst some arts and craft supplies.

And then there are likely to be things in your house made out of clay (or that were clay, and are no longer).

Clay Minerals and Earth Systems

Although most of the following points have already been made earlier in this chapter, or in other chapters, it is worth reviewing some of the key implications that clay minerals have for Earth systems:

• Clay minerals may have played a role in the initial evolution of life from organic chemicals because the regular structure of the clays could have acted as a template for the assembly of organic molecules,
• Conversion of silicate minerals to clay consumes atmospheric CO₂ and so has climate implications,
• Clay minerals accumulate trace elements that later become available to plants and microorganisms,
• Clay minerals accumulate trace elements that may eventually get concentrated into mineral deposits,
• Clay minerals can reduce rock strength and so contribute to erosion and slope failure,
• Clay minerals suspended in water or as clouds of dust can be vehicles for the transfer of trace and major elements from land into the ocean (See Figure 10.5.10), and
• Clay minerals are the vehicles for the transfer of water from subducted oceanic crust into the mantle, leading to magma formation (via flux melting) and volcanism.

Figure 10.5.10 A Cloud of Dust from the Sahara Desert Stretches into the Atlantic and Heads Towards Europe,  February 2021

Media Attributions

• Figure 10.5.1 Steven Earle, CC BY 4.0
• Figure 10.5.2 Steven Earle, CC BY 4.0
• Figure 10.5.3 Steven Earle, CC BY 4.0
• Figure 10.5.4 Plates of the Clay Mineral Kaolin from ACEMAC Nano Scale Electron Microscopy and Analysis Facility, University of Aberdeen by GSoil, CC BY 3.0, https://www.abdn.ac.uk/business-info/facilities-and-expertise/acemac-nano-scale-electron-microscopy-and-analysis-facility-846.php#panel861
• Figure 10.5.5 Steven Earle, CC BY 4.0, after Figure 7.13. In Prothero, D. and Schwab, F. (2004). Sedimentary geology: An introduction to sedimentary rocks and stratigraphy (2nd ed.). Freeman and Co.
• Figure 10.5.6 Photo by Isaac Earle, CC BY 4.0
• Figure 10.5.7 Serpentinite by James St. John, CC BY 2.0, via Flickr, https://www.flickr.com/photos/jsjgeology/16940796272
• Figure 10.5.8 Photo by Steven Earle, CC BY 4.0
• Figure 10.5.9 Vermiculite by KENPEI, 2008, CC BY-SA 2.1 JP, via Wikimedia Commons https://commons.wikimedia.org/wiki/File:Vermiculite1.jpg
• Figure 10.5.10 Africa Dust from NASA, public domain, https://eoimages.gsfc.nasa.gov/images/imagerecords/147000/147952/africadust_virs_202149_lrg.jpg

The topics covered in this chapter can be summarized as follows:

Water is constantly on the move. It is evaporated by solar energy from the oceans, lakes, streams, the surface of the land, and from plants (transpiration) (Figure 11.1.1). It is moved through the atmosphere by winds and condenses to form clouds of water droplets or ice crystals. In response to the pull of gravity it comes back down as rain or snow and then flows through streams, into lakes, and eventually back to the ocean. Water on the surface and in streams and lakes infiltrates the ground to become groundwater. Groundwater slowly moves through the surface materials and underlying bedrock. Some groundwater returns to streams and lakes, and some goes directly back to the oceans.

 

Figure 11.1.1 The Various Components of the Hydrologic Cycle. Black or white text indicates the movement or transfer of water from one reservoir to another. Yellow text indicates the storage of water.

Figure 11.1.2 Representation of the Volumes of Earth’s Water Reservoirs. The 1 litre jug is filled with salty sea water (97%). The ice-cube is glacial ice (2%). The 2 teaspoons represent groundwater (1%), and the three drops represent all of the fresh water in lakes, streams, and wetlands, plus all of the water in the atmosphere.

Even while it is moving around, water is stored in various reservoirs. The largest, by far, is the ocean, accounting for 97% of the total water volume at Earth’s surface. That water is salty. The remaining 3% is fresh water. Two-thirds of our fresh water is stored in ice and one-third is stored in the ground. The remaining fresh water—about 0.03% of the total—is stored in lakes, streams, vegetation, and the atmosphere. To put that in perspective, imagine putting all of Earth’s water into a 1 litre jug (Figure 11.1.2). We start with sea water, by almost filling the jug with 970 mL of water and 34 grams of salt (30 ml, 2 tablespoons, or 1/8th of a cup of salt). Then we add one regular-sized (roughly 20 mL) ice cube (representing glacial ice) and two teaspoons (roughly 10 mL) of groundwater. All of the fresh water that we see around us in lakes and streams and up in the sky can be represented by adding just three more drops from an eyedropper (about one-third of a millilitre).

The volumes of the reservoirs are listed in Table 11.1, along with the average residence time of water in each reservoir. A molecule of water that falls into or flows into the ocean will stay there for an average of about 3,100 years. During that time, it might get moved all around the Earth via surface currents, and even to the deep ocean by thermohaline circulation. It will eventually be returned to the atmosphere via evaporation. A molecule of water that gets frozen into snow and then falls on a glacier, stays there for as long as 16,000 years on average (although some deep Antarctic ice is over a million years old) before it flows out as meltwater or calves into the ocean.

Groundwater stays in aquifers for an average of 300 years, although some very deep groundwater is hundreds of thousands or even millions of years old. In freshwater lakes with flows in and out via rivers, the average residence time is 1 to 100 years, while in salt lakes, from which the only exit is likely to be by evaporation, the average residence time is up to 1000 years. Moisture in the soil typically stays there for less than a year. It is cycled out via plants, or seeps down to become groundwater. Water remains in a river for an average of 12 to 20 days before flowing into a lake or the ocean. Water is held in the atmosphere there for an average of only 8 days.

Although the proportion of Earth’s water that is in the atmosphere is tiny, the actual volume is significant. At any given time, there is the equivalent of approximately 13,000 cubic kilometres (km³) of water in the atmosphere in the form of water vapour and water droplets in clouds. Water is evapotranspired from vegetation, and evaporated from the oceans and lakes at a rate of 1,580 km³ per day, and just about exactly the same volume falls as rain and snow every day—over both the oceans and land. The precipitation that falls on land goes back to the ocean in the form of stream flow (117 km³/day) and groundwater flow (6 km³/day). Most of the rest of this chapter is about that 123 km³/day of stream and groundwater flow.

Surface water exists within lakes and ponds, as snow and ice, and within flowing streams. If we want to use surface water as a resource, we generally have to restrict ourselves to streams, and we have to be careful to leave enough of the stream water for the ecosystems (and other people) that depend on it. It is true that we can extract water from lakes, and many municipalities do that, but that can only be done to the extent that the lake is part of a stream system, and we cannot take more water from the lake than is being added by streams. If we do that on a long-term basis (more than 1 year) the lake level will start to drop and before too long the lake will not be viable as a water source.

Box 11.1 Water Resources

Which of the following countries has the greatest resource of fresh water? Brazil, Canada, China, Russia, or USA?

Figure 11.1.3 A Lake-Rich Region Within the Mackenzie River Delta, Canada

If you went to school in Canada, you almost certainly answered “Canada” because that’s what Canadian school children are taught. And it’s not surprising. Canada has more lakes and more lake surface area than any other country, so it looks like a country with lots of water (Figure 11.1.3). Indeed, for many decades Canadian school children have been unintentionally led to believe that Canada’s freshwater resources are so vast that they don’t need to be careful with water use.

But, as noted above, a water resource is only the fraction of the water in the water cycle that humans can use without causing harm somewhere else in the water cycle.

For example, rivers may have a great deal of water flow, but only a fraction can be removed from the stream for human purposes as the river must still have sufficient water for the ecosystem to be healthy.  Lakes are the same.  They have inflows and outflows, and only a fraction of the inflows could be diverted to a human water resource without affecting the lake.  In terms of water supply for human use, a lake is only as good as the amount of water flowing through it.  Once these restrictions are considered, Canada can still be understood to have lots of water, but in terms of water resources, it ranks 4th on this list.  And that’s only part of the story.  Water is only a really a resource for human use if it is near to humans.  Most of the freshwater in Canada flows north into the Arctic Ocean or into Hudson Bay, while most Canadians live close to the border with the US, and so only a fraction of that fresh water is actually available. Canadians, like everyone else in the world, need to be very careful with their water supplies.

Media Attributions

• Figure 11.1.1 Steven Earle, CC BY SA 4.0, after Water Cycle by Ingwik,  CC BY SA 3.0, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Water_cycle_blank.svg
• Figure 11.1.2 Steven Earle, CC BY 4.0
• Figure 11.1.3 Modified by Steven Earle, from Mackenzie River with Heavy Sediment Load by Pierre Markuse, 2018, CC BY 2.0 via Flickr, https://www.flickr.com/photos/pierre_markuse/48169370867/in/photolist-HgtP63-GZC16h-QFW5id-QFWJa5-2goyBLg

Human activities are common causes of water-supply problems, and, in this regard, the most important activities are agriculture, industry, landfills, sewage, and anything that can lead to elevated turbidity. Anthropogenic sources are divided into two types, nonpoint sources where the effect is distributed over a large area such as agriculture or logging, and point sources, like factories, landfills or mines, where the effect is localized to a specific site.

By a wide margin, agriculture represents the greatest threat to water quality. Agriculture is everywhere, and it occupies more land than all other uses combined. There is currently 30% more agricultural land than forested land. Modern agriculture, while very efficient in terms of production per input dollar, is also intensively dependent on the use of chemicals (fertilizers and pesticides) and a significant proportion of the chemicals applied either run off the surface into streams and lakes, or seep into aquifers.

A two-decade study by the US Geological Survey has shown that the water in over 50% of wells tested in agricultural regions has detectable amounts of pesticides, that the issue is also prevalent in urban areas (because of pesticide use on golf courses and lawns) and that it was worse during the period from 2002 to 2011 than it was from 1993 to 2001 (Figure 11.2.1). Although the figure doesn’t show levels of pesticides in these wells, another study has shown that over 1000 wells in Florida have been closed because of pesticide levels above acceptable limits.

Figure 11.2.1 Pesticide Incidence in Groundwater in the United States, 1993 to 2011

Large amounts of nitrogen, phosphorous and potassium fertilizers are applied to fields all over the world. These nutrients help the crops to grow, but if more fertilizer is applied that is really needed, then the excess will end up making its way into the environment. Nitrogen from fertilizers can volatilize into the air, to come down within rainfall somewhere else. Nitrogen, phosphorous and potassium can run off fields into surface water bodies. They can also infiltrate down into the soil, and move downwards out of reach of the plant roots where they are added to the groundwater. That groundwater may discharge to a surface water body.  In lakes and streams these excess nutrients help algae to thrive, and that can create significant problems for aquatic life and human water supplies and is also an issue for anyone who likes to be in or near the water.

Lake Erie has long been the poster child for algal blooms related to excess nutrients, as is illustrated in a satellite image from 2011 (Figure 11.2.2). Erie is the shallowest of the Great Lakes (average depth is only 19 m, compared with 147 m for Lake Superior) and so has the greatest potential to warm up in the summer. The Erie drainage basin is also the most heavily populated of the Great Lakes (home to 12.4 million people) and the lake is entirely surrounded by farmland and cities and industry, with very little remaining forest. The western part of Lake Erie has a springtime phosphorous level of over 12 µg/L, whereas Lakes Superior and Huron are in the range of 2 to 3 µg/L, and Lake Ontario is just over 6 µg/L.

Figure 11.2.2 Algal Growth in Lake Erie in 2011

Agriculture is a major source of the phosphorous that leads to excess algal growth in Lake Erie, but not the only one. Other important sources include municipal sewage effluent, industrial effluent, urban storm water runoff, and atmospheric deposition.
Runoff from fields is not only laced with excess nutrients and pesticides, it is also rich in suspended matter that contributes to the turbidity of surface water, degrading aquatic ecosystems, and potentially endangering human water supplies (Figure 11.2.3). The risk of turbid runoff is greatest when fields have been left unprotected by vegetation following harvesting or tillage.

Figure 11.2.3 Turbid Runoff From a Field in Iowa Following a Rainstorm

Nearly 80% of global farmland is dedicated to livestock, either for grazing or for growing animal feed (although livestock farming accounts for only 18% of the caloric output from farming) and farm animals produce manure—lots of it. Manure from all livestock is rich in nitrogen, potassium and phosphorous, and so contributes to the surface water algae problem described above, and to elevated levels of nitrogen in groundwater.

The Abbotsford-Sumas Aquifer of southwestern British Columbia and northwestern Washington provides a good example of how agriculture can affect groundwater quality. The area is intensively farmed with row crops (mostly raspberries and blueberries), large-scale poultry production and pastured livestock. In the past, manure from poultry operations was spread on raspberry fields. A 1992 study on the Canadian side of the border showed that 60% of 125 wells had nitrate levels above the 10 mg/L MAC for drinking water. The Abbotsford-Sumas Aquifer slopes towards the south and crosses into northwestern Washington. A 1999 report shows that the nitrogen contamination extended south across the border, although similar farming activities on the US side are responsible for much of that contamination. The extent of the nitrate contamination in the Abbotsford-Sumas Aquifer is shown on Figure 11.2.4.

Figure 11.2.4 Extent of Elevated Nitrate Levels in the Abbotsford-Sumas Aquifer, British Columbia and Washington

In addition to nitrates and other nutrients, animal manure also contains bacteria, and while most of those bacteria are harmless to humans, some present a serious risk.

Walkerton Ontario, a town of about 5000 residents situated 150 km northwest of Toronto, has a water supply system sourced by wells in a fractured bedrock aquifer. One of those wells (Well 5) was shallow, drawing water from a depth of only 6 m. A heavy rainstorm in early May 2000 soaked a pasture that had been fertilized with cattle manure a few weeks earlier. Some of that water was introduced into the Walkerton water system from Well 5, situated 80 metres away from the pasture. In the following week 2300 residents became ill with gastroenteritis from the bacterium Escherichia coli in the water supply, and seven people died. Although Walkerton’s Public Utility Commission did routine tests for bacterial contamination, the testing protocols were not always followed nor were they as frequent as prescribed, monitoring records were falsified, and the utility’s chlorination system was set to a level below that stipulated by the Ontario Clean Water Agency. A public inquiry concluded that much of the blame for the Walkerton disaster can be traced back to provincial government cuts to funding for programs that had been established to protect citizens and ensure safe water supplies. The Walkerton disaster spurred increased efforts to protect groundwater across Canada.

Logging, like agriculture, is a nonpoint source of pollution because it is carried out over wide areas, and it involves the application of fertilizers and pesticides and the removal of vegetation. But logging differs because it is often carried out on steep slopes—in the remaining forested areas that are too rugged for farming or urbanization—(Figure 11.2.5) and so the issue with erosion and turbid runoff is exacerbated. Most logging also operations also involve the construction of temporary roads on steep terrain. Road construction amplifies the risk of slope failure and that increases the likelihood of damage to surface water habitat and supplies.

Figure 11.2.5 A Clear Cut in the Coquille River Basin of Southwestern Oregon

Point sources of pollution include mines, petroleum extraction operations, plants for extracting and refining ores and petroleum, chemical works, manufacturing plants, sewage treatment systems, landfills and underground storage tanks (USTs) at filling stations. Some of these are discussed in other chapters (mining in Chapter 8, energy in Chapter 9, and landfills in Chapter 13) and therefore won’t be considered here.

Throughout the history of industrialization, owners and operators all of these types of facilities have failed catastrophically and repeatedly to protect the environment, water supplies, ecosystems and the lives and health of workers and innocent people. One of the most notorious examples was in the city of Niagara Falls, New York. In the 1920s the city set aside an abandoned canal—Love Canal—as a dump site for municipal waste. In 1942 the Hooker Chemical Company was given permission to dispose of chemical by-products from the manufacture of dyes, perfumes and rubber goods into the dump. From then until 1952 Hooker dumped nearly 20,000 tonnes of chemical wastes at the site, most of it in 200-litre metal drums, which they eventually covered up with clay. In 1952 they sold the site (which they had purchased from the city in 1947), along with the liability for the dumped chemicals, to the Niagara Falls Board of Education. A 400-student elementary school was built on top of the old dump and it was in use from 1955 to 1978. Soon after construction of the school, toxic chemicals started coming to surface, parts of the clay cover collapsed, and some of the metal drums were exposed at surface. During heavy storms contaminated surface water ran off the site and into the Niagara River (and then over the falls). Contaminated groundwater also migrated offsite and into the basements of neighbouring homes. The Love Canal site was eventually cleaned up, and the school and surrounding homes abandoned, but there are lingering health issues amongst those who lived in the area.

There are industrial sites like Love Canal all over the world and their contamination legacies will be with us for decades, even centuries. Increasingly however, corporations and individuals are being held to account for making the assumption that they can dump whatever waste they need to get rid of onto the ground and into water and the air.

But it isn’t just large operations that are problematic. In virtually every village, town and neighbourhood in North America (and elsewhere as well) there are abandoned sites surrounded by chain link fences and dotted with the plastic pipes of monitoring wells. Many of these are the remains of gas stations with underground storage tanks (USTs) that had leaked so badly that the operation was forced to close. If you think you haven’t seen one, you probably just haven’t realized what they are. Some UST sites have been cleaned up and rehabilitated. The one illustrated on Figure 11.2.6 has been vacant for over 20 years, possibly because the value of the land is less than the cost of removing the contamination and making the site usable for some other purpose.

Figure 11.2.6 The Site of a Filling Station With Leaking Underground Storage Tanks. The white pipes are groundwater monitoring wells.

A leaking fuel UST is at least three problems in one (Figure 11.2.7). Some components of the fuel are volatile, and so will form a vapour phase that rises to surface and may enter nearby buildings. Most fuels used for vehicles like cars, trucks and planes are lighter than water and so will float along the surface of the saturated zone (the water table). Some components of the fuel are soluble in water and so will mix with groundwater and be dispersed in that way. Some components of the fuel may be heavier than water so will sink through the saturated zone. The aquifer water will eventually come to surface somewhere and will bring some of those dissolved and heavy constituents with it.

Figure 11.2.7 Illustration of the Fate of Petroleum That Has Leaked from a UST

Media Attributions

• Figure 11.2.1 Steven Earle, CC BY 4.0, based on data in Toccalino, P., Gilliom, R., Lindsey, B., Rupert, M. (2014). Pesticides in groundwater in the United States: Decadal-scale changes 1993-2011. Groundwater, 52, 112-125, via US Geological Survey, https://ca.water.usgs.gov/pubs/2014/ToccalinoEtAl2014.pdf
• Figure 11.2.2 Excess Nutrients Flowing in Lake Erie, satellite image from N.A.S.A., public domain; inset photo is public domain from U.S. Geological Survey, https://www.usgs.gov/media/images/excess-nutrients-flowing-lake-erie-can-cause-serious-algal-blooms
• Figure 11.2.3 Runoff of Soil by Lynn Betts, U.S. Department of Agriculture, public domain, via Wikimedia Commons, https://commons.wikimedia.org/wiki/File:Runoff_of_soil_%26_fertilizer.jpg
• Figure 11.2.4 Steven Earle, CC BY 4.0, based on maps in Leibscher et al. (1992). and in Carey, B. and Cummings, R. (2012). Sumas Blaine Aquifer nitrate contamination summary. Department of Ecology, State of Washington, Publication 12-03-026, https://apps.ecology.wa.gov/publications/documents/1203026.pdf
• Figure 11.2.5 Plum Creek Logging by Francis Eatherington, CC BY-NC 2.0, via Flickr, https://www.flickr.com/photos/umpquawild/2514317367
• Figure 11.2.6 Photo by Steven Earle, CC BY 4.0
• Figure 11.2.7 Steven Earle, CC BY 4.0

In Section 11.1 we focused on the volume of water available, and argued that it is acceptable for water to be diverted from the water cycle for human use, as long as it does not adversely affect the rest of the hydrologic cycle. This makes for a potential water resource. For water to be useful, it must also be of suitable quality. There are a range of factors that affect the quality of potential water resources.

Fresh water on the surface (in streams or lakes) has limited opportunity to react with the surrounding materials, and so its natural chemical composition (excluding human-sourced contamination) is not typically a factor in terms of its quality. Groundwater, on the other hand, is in close contact with the rock or sediment through which it is moving and there is plenty of opportunity for the water chemistry to be changed by interaction with minerals. Groundwater can have specific natural chemical properties that make it suitable for some purposes. For example, many commercial spring water purveyors make (generally unsupported) claims that their water is healthier or tastes better than other water. On the other hand, groundwater can also have qualities that make it less than ideal, or even a risk to health. Concentrations of some of the common naturally occurring chemicals in water are summarized in this section.

Figure 11.3.1 shows the concentrations of the major constituents in groundwater from different parts of a sandstone aquifer on Vancouver Island and in surface waters from the same region. These values could be considered to be somewhat representative of aquifers in other locations, and certainly these seven constituents are the dominant components of most groundwaters. Bicarbonate, sodium and chloride are present at the highest concentrations in these waters—all between 70 and 120 mg/L—followed by sulphate and calcium at around 20 mg/L. If we were to add up all of these numbers the total would be close to 350 mg/L, which means that if we evaporated a litre of this water, we would be left with 0.35 g of salts, or a little less than 1/10th of a teaspoon. Water with over 1000 mg/L of dissolved solids is unlikely to be good to drink. Sea water has approximately 35,000 mg/L of dissolved solids (about 7 teaspoons), mostly as sodium and chloride.

Figure 11.3.1 The average concentrations of the major components of groundwater from a sandstone aquifer, and surface water in the Vancouver Island area, BC

Concentrations in surface waters (streams and lakes) from the same areas of Vancouver Island are much lower (20% as high on average), although K and Mg are a little higher in the surface waters than the groundwaters. One litre of this water contains about 70 mg of the salts of these elements.

Concentrations of some of the important minor constituents in water samples from the same area are shown on Figure 11.3.2. The most abundant of these elements is silicon (Si), followed by iron (Fe) and fluoride (F). Note that this graph has a logarithmic scale; the average concentration of Fe is less than 1/10th that of Si, and Al is about 1/100th that of Si. The total contribution of all of these elements is equivalent to 6 mg in a litre of water.

Figure 11.3.2 The Average Concentrations of Some of the Minor Components of Groundwater from a Sandstone Aquifer, and Surface Water in the Vancouver Island Area, BC. On average, the surface waters represented here have minor element concentrations that are about 1/5th as high as the groundwaters.

pH

The hydrogen ion concentration of water is expressed in pH units. Most waters have pH within the range 6 to 8, where anything less than 7 is acidic. It is not uncommon for surface waters to have pH levels of less than 6, especially in areas where the water is draining over rock with even low levels of pyrite (FeS₂). A low pH is not necessarily a concern in itself, but because most metals are more soluble at lower pH, acidic water tends to have higher concentrations of heavy metals such as copper and zinc (see Section 8.2). Acidic water can also be a problem in buildings because it will contribute to corrosion of metal pipes. It’s also quite common for surface water to have a pH well above 7, especially in areas underlain by limestone. Groundwater can also be acidic or basic (pH greater than 7). As discussed below in the section on fluoride, while a high pH is not typically a problem itself, it can be associated with other issues.

Hardness

One of the most common issues with groundwater is hardness, which is a measure of the combined concentrations of calcium and magnesium. Water is considered “hard” if it has hardness above 80 mg/L (Ca plus Mg expressed as CaCO₃ equivalent) and this is common with waters that are derived from sandy aquifers or from limestone aquifers. Hard water inhibits the activity of soaps and detergents, so makes washing of clothes or dishes less effective than it is with soft water. Water softeners are widely used to treat hard water. These are effective in removing most of the Ca and Mg, but do so by replacing them with Na. Excess Na consumption is a health risk to many people, so drinking or cooking with softened water is not recommended.

Box 11.2 Softening Water

Figure 11.3.3 Schematic of a Water Softener

The ions of different elements have different tendencies to be adsorbed or desorbed from substrates such as clay minerals*. The tendency for adsorption amongst the major cations in natural waters is as follows:

(strongly adsorbed) Ca > Mg > K > Na (weakly adsorbed).

This means that calcium ions are much more likely to be adsorbed onto surfaces than are sodium ions.

A water softener (Figure 11.3.3) takes advantage of this relationship. As the “hard” water is passed through the system calcium and magnesium ions in solution are preferentially adsorbed onto a substrate (an ion-exchange resin) and sodium ions are released into the water. After some time, most of the exchange sites are occupied by calcium and magnesium and the system ceases to function effectively. A sodium-chloride brine is periodically passed through the system, and because of the overwhelming amount of sodium in the solution, the calcium and magnesium ions on the exchange sites are replaced by sodium ions, thus “recharging” the ion exchange resin.

*Note that the word is adsorb not absorb. Adsorption means being attached onto a surface. Absorption means being incorporated into a solid. Water is absorbed into a sponge, but ions are adsorbed onto the surfaces of minerals.

Iron

Like most other elements, iron can exist in more than one oxidation state; the two most common are Fe²⁺ (ferrous iron) and Fe³⁺ (ferric iron). The ferrous form predominates in situations where the oxidation potential is low and this is a characteristic of groundwater in many situations, especially where there is something like organic matter or iron-sulphide minerals that can use up the oxygen present in the water. Ferrous iron is quite soluble, while ferric iron is virtually insoluble, except at very low pH (which is typical of acid rock drainage, as described in Chapter 8). While the ground and surface waters shown on Figure 11.3.2 have iron concentrations of around 0.2 mg/L, ferrous iron is present at up to 10 mg/L in some deep groundwater. When water of this type comes to surface the iron is quickly oxidized to the ferric state, and it precipitates as iron oxide and iron hydroxide minerals. An illustration of the effects this process is provided on Figure 11.3.3. Rusty stains like this are also common in kitchen and bathroom fixtures in places where the water source is anoxic and rich in iron.

Figure 11.3.4 Rusty Stains Associated With the Discharge of Iron-Bearing Groundwater from a Sandstone Aquifer on Vancouver Island

Fluoride

Fluoride is beneficial for dental health at low levels, but too much fluoride can lead to discolouration and malformation of teeth, and excessive fluoride over decades can lead to crippling skeletal problems. Almost all surface water and most groundwater has less than 0.5 mg/L (<500 µg/L) fluoride, but some groundwater has over 1.5 mg/L, which is the international maximum acceptable concentration (MAC) for fluoride. Fluoride solubility is low in water that has significant calcium concentrations and relatively low pH.

If groundwater has an elevated fluoride level its not likely because the aquifer materials (rock or sediments) have particularly high fluoride levels. Instead, there are water-rock interactions that make fluoride more soluble than it is under normal conditions, allowing fluoride levels to become elevated, even where there is relatively little fluorine in the surrounding material. That process is called base-exchange softening, and it is similar to what happens inside a water softener.

Most sandy aquifers also include clay minerals and, as described in Chapter 10, clay minerals are effective at ion adsorption. When water flows through a sandstone aquifer that has abundant adsorbed sodium ions, the calcium ions in the water will replace the sodium on the clay-mineral adsorption sites (see Box 11.2 for an explanation in the context of a water softener). This leads to softening of the water (more dissolved Na, less dissolved Ca) and also to a higher pH. Fluoride is relatively soluble under those conditions, so fluoride levels increase.

An example of this relationship is shown on Figure 11.3.4. Water samples from this aquifer have pH ranging from 5.5 to nearly 9.5. At low pH the fluoride levels remain well below 1 mg/L. Water that has been affected by base-exchange softening—resulting in an increase in Na at the expense of Ca, and an increase in pH—has higher fluoride levels, in many cases higher than the fluoride maximum acceptable concentration (MAC) of 1.5 mg/L fluoride.

Figure 11.3.5 pH Versus Fluoride in Groundwater Samples from Nanaimo Group Aquifers on Vancouver Island and Adjacent Islands

Arsenic

In most ground and surface waters arsenic (As) concentrations are typically below 5 µg/L, and so are below the international MAC of 10 µg/L, but in some situations arsenic levels can get much higher.

Figure 11.3.6 Areas of Significant Groundwater Arsenic Contamination in Bangladesh

In the floodplain of the Ganges and Brahmaputra Rivers in Bangladesh over 100 million residents extract groundwater from small “tube wells” (wells with diameters of less than about 10 cm). About 8 million such wells were installed in Bangladesh between 1960 and 1990, many with assistance from UNICEF. Prior to that time most rural Bangladeshi’s relied on surface water sources, many of which were contaminated by bacteria and viruses. Gastrointestinal illnesses from these water sources were major health threats. The sediments surrounding groundwater wells are often sufficient to filter out many bacterial and viral contaminants, and thus provision of clean groundwater was celebrated as a major public health success.

However, in the mid 1990s it was discovered that many of the tube wells in Bangladesh have As levels above 50 μg/L (Figure 11.3.5), and as many as 20 million Bangladeshi’s are at risk of As poisoning (outcomes include cancer, diabetes, thickening of the skin, liver disease and digestive problems).

According to McArthur et al. (2001), the high levels of As are directly related to the degree of oxidation of the water in the aquifer. Like iron, arsenic can have varying oxidation states, and as for iron, the less oxidized version, As³⁺, is much more soluble than the more oxidized version, As⁵⁺. Most of the tube wells in Bangladesh penetrate into ancient unconsolidated sand and gravel river sediments (dominated by quartz, feldspar and mica) that also include extensive peat layers. Oxygen is consumed as the organic matter in the peat reacts with groundwater. That generates the reducing conditions that result in the dissolution of the mineral limonite (FeOOH) to soluble ferrous iron (Fe²⁺), and release of the arsenic that was adsorbed onto the limonite. These same reducing conditions ensure that the arsenic remains in the more soluble arsenite (As³⁺) state, rather than the less soluble arsenate (As⁵⁺) state.

Saltwater Intrusion

Many important aquifers are in coastal areas, but aquifers close to a coast are at risk from intrusion by salt water. As illustrated on Figure 11.3.6, groundwater beneath the ocean is salty, while that beneath the land is fresh. Because fresh water is less dense than salt water the fresh groundwater in near-shore areas exists in a lens that is approximately 40 times as deep as the height of the water table above the ocean surface. A well drilled within this lens (A) will produce fresh water, but one that penetrates into the salty groundwater (B) will produce salt water. In the scenario shown, well A is at risk of producing salt water if it is overused, as that will draw the water table down, and bring the fresh-salt interface up.

Figure 11.3.7 Depiction of the Fresh Water–Salt Water Interface in Near-Shore Areas

Figure 11.3.8 The Areal Extent of Salt Water Intrusion in the Biscayne Aquifer, Florida

A significant salt water intrusion problem exists in the eastern part of Florida around Miami (Figure 11.3.7). The area is prone to intrusion for several reasons. First, it has very subdued topography; much of southern Florida is only a few metres above sea level. Second, the main near-surface aquifer in the region—the Biscayne Aquifer—includes limestone units with significant porosity and permeability related to dissolution that took place when the unit was above surface. In addition to those natural factors, there are also several anthropogenic factors that contribute to salt intrusion. The region is densely populated and most of the public water supply is extracted from wells in the Biscayne Aquifer. Starting in the 19th century large areas of the nearby Everglade wetlands were drained to create land for agriculture and urban expansion. The consequent drop in the water level allowed sea water to flow inland. Furthermore, draining the wetlands was largely accomplished by building canals, and those canals have allowed sea water to extend well inland and then seep into the underlying aquifer.

The extent of salt water intrusion along Biscayne Bay is illustrated on Figure 11.3.7. As of 2011 about 1200 km² were affected. Salt extended further inland in some areas in 1955—particularly along canals—but steps have been taken in recent decades to restrict the inland flow of salt water in the canals and so some of those areas have recovered.

Sea level rise resulting from climate change is going to exacerbate the salt water intrusion problem in southern Florida, and in many other low-relief coastal areas around the world. See Section 14.1 for more on that problem.

Turbidity

Turbidity is a measure of the amount of solid matter—clay and silt plus fine-grained organic matter—suspended in water. Turbid water may look cloudy, although at relatively low turbidity levels—which can still be dangerous—turbidity isn’t detectable without an instrument. Turbidity isn’t an issue because the suspended particles themselves are a health risk, but because they inhibit the effectiveness of disinfection processes. In turbid water biological pathogens attached onto clay mineral grains are effectively protected from chlorine or ozone disinfection, or they can literally be hidden from ultraviolet radiation.

Groundwater is rarely turbid because of its inherent self-filtering by the aquifer, but surface water is often turbid, especially in regions with steep terrain (Figure 11.3.8) or following a storm or a debris flow. Most turbidity results from natural processes, but it can be exacerbated by human activities, especially farming, logging and construction.

Figure 11.3.9 Turbid Waters of a Glacier-Sourced Stream in the Canadian Rockies

Although it may not seem like a serious issue, elevated turbidity is one of the most common reasons for the declaration of boil-water advisories. Water suppliers that rely on surface water need to monitor turbidity carefully and continuously and to ensure that their filtration measures are effective. In order to avoid this problem, some water suppliers maintain a surface water supply for most of the year, but have backup groundwater wells that they can draw upon when the turbidity of the surface water is too high.

Media Attributions

• Figure 11.3.1 Steven Earle, CC BY 4.0, based on data in Allen, D. and Suchy, M. (2001). Geochemical evolution of groundwater on Saturna Island, British Columbia. Canadian Journal of Earth Sciences, 38(7), 1059-1080. https://doi.org/10.1139/e01-007; and in Earle, S., Krogh, E. (2004) Groundwater geochemistry of Gabriola. Shale: Journal of the Gabriola, Historical & Museum Society, No. 7, 37-44. https://www.researchgate.net/publication/237654544_Geochemistry_of_Gabriola’s_groundwater
• Figure 11.3.2 Steven Earle, CC BY 4.0, based on data in Earle, S. and Krogh, E. (2004). Groundwater geochemistry of Gabriola. Shale, No. 7, 37-44. https://www.researchgate.net/publication/237654544_Geochemistry_of_Gabriola’s_groundwater
• Figure 11.3.3 How [a] Water Softener Works by Aqua mechanical, CC BY 2.0, via Flickr, https://www.flickr.com/photos/aquamech-utah/22391683206/
• Figure 11.3.4 Steven Earle, CC BY 4.0
• Figure 11.3.5 Steven Earle, CC BY 4.0, based on data in Earle, S. and Krogh, E. (2006). Elevated fluoride levels in a sandstone and mudstone aquifer system, eastern Vancouver Island, Canada. Proceedings of the Sea to Sky Geotechnique 2006 – the 59th Canadian Geotechnical Conference and the 7th Joint CGS/IAH-CNC Groundwater Specialty Conference, International Association of Hydrogeologists, 2006, 1584-1591. http://www.x-cd.com/SeatoSkyOnline/S4/0240-247.pdf
• Figure 11.3.6 Groundwater Arsenic Contamination in Bangladesh from Ahmad, S., Khan, M., Haque, M. (2018). Arsenic contamination in groundwater in Bangladesh: implications and challenges for healthcare policy. Risk Management and Healthcare Policy, 11, 251-261. https://doi.org/10.2147/RMHP.S153188. CC BY-NC 3.0)
• Figure 11.3.7 Steven Earle, CC BY 4.0
• Figure 11.3.8 Steven Earle, CC BY 4.0, based on information in Prinos et al. (2014)
• Figure 11.3.9 Steven Earle, CC BY 4.0