Physical Geology - 2nd Edition

Physical Geology - 2nd Edition

Steven Earle

Karla Panchuk


Victoria, B.C.



About BCcampus Open Education

BCcampus Open Education began in 2012 as the B.C. Open Textbook Project with the goal of making post-secondary education in British Columbia more accessible by reducing student costs through the use of openly licensed textbooks and other OER. BCcampus supports the post-secondary institutions of British Columbia as they adapt and evolve their teaching and learning practices to enable powerful learning opportunities for the students of B.C. BCcampus Open Education is funded by the British Columbia Ministry of Advanced Education, Skills & Training, and the Hewlett Foundation.

Open textbooks are open educational resources (OER) created and shared in ways so that more people have access to them. This is a different model than traditionally copyrighted materials. OER are defined as teaching, learning, and research resources that reside in the public domain or have been released under an intellectual property license that permits their free use and re-purposing by others."Open Educational Resources," Hewlett Foundation, (accessed September 27, 2018). Our open textbooks are openly licensed using a Creative Commons licence, and are offered in various e-book formats free of charge, or as printed books that are available at cost. For more information about open education in British Columbia, please visit the BCcampus Open Education website. If you are an instructor who is using this book for a course, please fill out our Adoption of an Open Textbook form.


Exceptions to the CC BY Licence

Physical Geology – 2nd Edition by Steven Earle is under a Creative Commons Attribution 4.0 International Licence, except where otherwise noted. The following materials have been included in this text but are under licences with additional restrictions.

All rights reserved

Permission to include All Rights Reserved pieces has been granted for non-commercial purposes in this open textbook by the copyright holders. If you plan to adapt any content including these pieces, please reconfirm permission with the copyright holders or remove them from your version.

Permission from Google

The following images are used with permission from Google, which allows for non-commercial use.






I am grateful to the members of the BC Earth Science Articulation Committee for their encouragement and support during this project, and to the following colleagues from institutions around BC and elsewhere for acting as subject matter experts and chapter reviewers: Sandra Johnstone, Kathleen Jagger, Tim Stokes, Cathie Hickson, Michelle Lamberson, Casey Brant, Alan Gilchrist, Deirdre Hopkins, Todd Redding, Duncan Johansen, Craig Nicol, John Martin, Mark Smith, Jeff Lewis, and Russel Hartlaub. I am also grateful to Karla Panchuk of the University of Saskatchewan for conceiving and writing Chapter 22, The Origin of the Earth and the Solar System.

I thank the staff of BCcampus, especially Amanda Coolidge for her excellent guidance and devotion to this project, and also Clint Lalonde and Lauri Aesoph.

And finally, I thank my family for inspiration and help, especially Justine and Kate, and also Isaac, Rosie, Heather, and Tim.



This book was born out of a 2014 meeting of earth science educators representing most of the universities and colleges in British Columbia, and nurtured by a widely shared frustration that many students are not thriving in our courses because textbooks have become too expensive for them to buy. But the real inspiration comes from a fascination for the spectacular geology of western Canada and the many decades that I have spent exploring this region along with colleagues, students, family, and friends. My goal has been to provide an accessible and comprehensive guide to the important topics of geology, richly illustrated with examples from western Canada. Although this text is intended to complement a typical first-year course in physical geology, its contents could be applied to numerous other related courses.

As a teacher for many years, and as someone who is constantly striving to discover new things, I am well aware of that people learn in myriad ways, and that for most, simply reading the contents of a book is not one of the most effective ones. For that reason, this book includes numerous embedded exercises and activities that are designed to encourage readers to engage with the concepts presented, and to make meaning of the material under consideration. It is strongly recommended that you try the exercises as you progress through each chapter. You should also find it useful, whether or not assigned by your instructor, to complete the questions at the end of each chapter.

Over many years of teaching earth science I have received a lot of feedback from students. What gives me the most pleasure is to hear that someone, having completed my course, now sees Earth with new eyes, and has discovered both the thrill and the value of an enhanced understanding of how our planet works. I sincerely hope that this textbook will help you see Earth in a new way.

Steven Earle, Gabriola Island, 2015


Chapter 1 Introduction to Geology

Learning Objectives

After carefully reading this chapter, completing the exercises within it, and answering the questions at the end, you should be able to:

  • Explain what geology is, how it incorporates the other sciences, and how it is different from the other sciences.
  • Discuss why we study Earth and what type of work geologists do.
  • Define some of the properties of a mineral and explain the differences between minerals and rocks.
  • Describe the nature of Earth’s interior and some of the processes that take place deep beneath our feet.
  • Explain how those processes are related to plate tectonics and describe a few of the features that are characteristic of plate boundaries.
  • Use the notation for geological time, gain an appreciation for the vastness of geological time, and describe how very slow geological processes can have enormous impacts over time.


1.1 What is Geology?

In its broadest sense, geology is the study of Earth—its interior and its exterior surface, the minerals, rocks and other materials that are around us, the processes that have resulted in the formation of those materials, the water that flows over the surface and through the ground, the changes that have taken place over the vastness of geological time, and the changes that we can anticipate will take place in the near future. Geology is a science, meaning that we use deductive reasoning and scientific methods to understand geological problems. It is, arguably, the most integrated of all of the sciences because it involves the understanding and application of all of the other sciences: physics, chemistry, biology, mathematics, astronomy, and others. But unlike most of the other sciences, geology has an extra dimension, that of time—deep time—billions of years of it. Geologists study the evidence that they see around them, but in most cases, they are observing the results of processes that happened thousands, millions, and even billions of years in the past. Those were processes that took place at incredibly slow rates—millimetres per year to centimetres per year—but because of the amount of time available, they produced massive results.

Geology is displayed on a grand scale in mountainous regions, perhaps nowhere better than the Rocky Mountains in Canada (Figure 1.1.1). The peak on the right is Rearguard Mountain, which is a few kilometres northeast of Mount Robson, the tallest peak in the Canadian Rockies (3,954 metres). The large glacier in the middle of the photo is the Robson Glacier. The river flowing from Robson Glacier drains into Berg Lake in the bottom right. There are many geological features portrayed here. The sedimentary rock that these mountains are made of formed in ocean water over 500 million years ago. A few hundred million years later, these beds were pushed east for tens to hundreds of kilometres by tectonic plate convergence and also pushed up to thousands of metres above sea level. Over the past two million years this area—like most of the rest of Canada—has been repeatedly glaciated, and the erosional effects of those glaciations are obvious.

The Robson Glacier is now only a small remnant of its size during the Little Ice Age of the 15th to 18th centuries, and even a lot smaller that it was just over a century ago in 1908. The distinctive line on the slope on the left side of both photos shows the elevation of the edge of the glacier a few hundred years ago. Like almost all other glaciers in the world, it receded after the 18th century because of natural climate change, is now receding even more rapidly because of human-caused climate change.

Two photos of a glacier. One taken in 1908 and the other in 2012. In that time, the glacier has melted substantially
Figure 1.1.1 Rearguard Mountain and Robson Glacier in Mount Robson Provincial Park, BC. Left: Robson Glacier in 2012. Right: Robson Glacier circa 1908.

Geology is also about understanding the evolution of life on Earth; about discovering resources such as water, metals and energy; about recognizing and minimizing the environmental implications of our use of those resources; and about learning how to mitigate the hazards related to earthquakes, volcanic eruptions, and slope failures. All of these aspects of geology, and many more, are covered in this textbook.

What are scientific methods?

A group of people peer down at rocks in a stream
Figure 1.1.2

There is no single method of inquiry that is specifically the “scientific method”; furthermore, scientific inquiry is not necessarily different from serious research in other disciplines. The most important thing that those involved in any type of inquiry must do is to be skeptical. As the physicist Richard Feynman once said: the first principle of science is that “you must not fool yourself—and you are the easiest person to fool.” A key feature of serious inquiry is the creation of a hypothesis (a tentative explanation) that could explain the observations that have been made, and then the formulation and testing (by experimentation) of one or more predictions that follow from that hypothesis.

For example, we might observe that most of the cobbles in a stream bed are well rounded (see photo above), and then derive the hypothesis that the rocks are rounded by transportation along the stream bed. A prediction that follows from this hypothesis is that cobbles present in a stream will become increasingly rounded as they are transported downstream. An experiment to test this prediction would be to place some angular cobbles in a stream, label them so that we can be sure to find them again later, and then return at various time intervals (over a period of  years) to carefully measure their locations and roundness.

A critical feature of a good hypothesis and any resulting predictions is that they must be testable.  For example, an alternative hypothesis to the one above is that an extraterrestrial organization creates rounded cobbles and places them in streams when nobody is looking. This may indeed be the case, but there is no practical way to test this hypothesis. Most importantly, there is no way to prove that it is false, because if we aren’t able to catch the aliens at work, we still won’t know if they did it!

Media Attributions


1.2 Why Study Earth?

The simple answer to this question is that Earth is our home—our only home for the foreseeable future—and in order to ensure that it continues to be a great place to live, we need to understand how it works. Another answer is that some of us can’t help but study it because it’s fascinating. But there is more to it than that:

An example of the importance of geological studies for minimizing risks to the public is illustrated in Figure 1.2.1. This is a slope failure that took place in January 2005 in the Riverside Drive area of North Vancouver. The steep bank beneath the house shown gave way, and a slurry of mud and sand flowed down, destroying another house below and killing one person. This event took place following a heavy rainfall, which is a common occurrence in southwestern B.C. in the winter.

A steep, muddy cliff at the end of a residential yard where the embankment gave way
Figure 1.2.1 The aftermath of a deadly debris flow in the Riverside Drive area of North Vancouver in January 2005.

The irony of the 2005 slope failure is that the District of North Vancouver had been warned in a geological report written in 1980 that this area was prone to slope failure and that steps should be taken to minimize the risk to residents. Very little was done in the intervening 25 years, and the consequences of that were deadly.

Media Attributions


1.3 What Do Geologists Do?

Geologists are involved in a range of widely varying occupations with one thing in common: the privilege and responsibility of studying this fascinating planet. In Canada, many geologists work in the resource industries, including mineral exploration and mining and energy exploration and extraction. Other major areas where geologists work include hazard assessment and mitigation (e.g., assessment of risks from slope failures, earthquakes, and volcanic eruptions); water supply planning, development, and management; waste management; and assessment of geological issues in the forest industry, and on construction projects such as highways, tunnels, and bridges. Most geologists are employed in the private sector, but many work for government-funded geological organizations, such as the Geological Survey of Canada or one of the provincial geological surveys. And of course, many geologists are involved in education at the secondary and the post-secondary levels.

Some people are attracted to geology because they like to be outdoors, and it is true that many geological opportunities involve fieldwork in places that are as amazing to see as they are interesting to study. But a lot of geological work is also done in offices or laboratories. Geological work tends to be varied and challenging, and for these reasons and many others, geologists are among those who are the most satisfied with their employment.

Figure 1.3.1 Geologists examining ash-layer deposits at Kilauea Volcano, Hawaii.

In Canada, most working geologists are required to be registered with an association of professional geoscientists. This typically involves meeting specific post-secondary educational standards and gaining several years of relevant professional experience under the supervision of a registered geoscientist. More information can be found at Engineers and Geoscientists British Columbia.

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1.4 Minerals and Rocks

The rest of this chapter is devoted to a brief overview of a few of the important aspects of physical geology, starting with minerals and rocks. This is followed by a review of Earth’s internal structure and the processes of plate tectonics, and an explanation of geological time.

The Earth is made up of varying proportions of the 90 naturally occurring elements—hydrogen, carbon, oxygen, magnesium, silicon, iron, and so on. In most geological materials, these combine in various ways to make minerals. Minerals will be covered in some detail in Chapter 2, but here we will briefly touch on what minerals are, and how they are related to rocks.

A mineral is a naturally occurring combination of specific elements that are arranged in a particular repeating three-dimensional structure or lattice.Terms in bold are defined in the glossary at the end of the book. The mineral halite is shown as an example in Figure 1.4.1.

Figure 1.4.1 The lattice structure and composition of the mineral halite (common table salt).

In this case, atoms of sodium (Na: purple) alternate with atoms of chlorine (Cl: green) in all three dimensions, and the angles between the bonds are all 90°. Even in a tiny crystal, like the ones in your salt shaker, the lattices extend in all three directions for thousands of repetitions. Halite always has this composition and this structure.

Note: Element symbols (e.g., Na and Cl) are used extensively in this book. In Appendix 1, you will find a list of the symbols and names of the elements common in minerals and a copy of the periodic table. Please use those resources if you are not familiar with the element symbols.

There are thousands of minerals, although only a few dozen are mentioned in this book. In nature, minerals are found in rocks, and the vast majority of rocks are composed of at least a few different minerals. A close-up view of granite, a common rock, is shown in Figure 1.4.2. Although a hand-sized piece of granite may have thousands of individual mineral crystals in it, there are typically only a few different minerals, as shown here.

Hornblende is black, quartz is white and shiny, and feldspar is white-ish or grey
Figure 1.4.2 A close-up view of the rock granite and some of the minerals that it typically contains (H = hornblende (amphibole), Q = quartz and F = feldspar). The crystals range from about 0.1 to 3 millimetres (mm) in diameter. Most are irregular in outline, but some are rectangular.

Rocks can form in a variety of ways. Igneous rocks form from magma (molten rock) that has either cooled slowly underground (e.g., to produce granite) or cooled quickly at the surface after a volcanic eruption (e.g., basalt). Sedimentary rocks, such as sandstone, form when the weathered products of other rocks accumulate at the surface and are then buried by other sediments. Metamorphic rocks form when either igneous or sedimentary rocks are heated and squeezed to the point where some of their minerals are unstable and new minerals form to create a different type of rock. An example is schist.

A critical point to remember is the difference between a mineral and a rock. A mineral is a pure substance with a specific composition and structure, while a rock is typically a mixture of several different minerals (although a few types of rock may include only one type of mineral). Examples of minerals are feldspar, quartz, mica, halite, calcite, and amphibole. Examples of rocks are granite, basalt, sandstone, limestone, and schist.

Key Takeaway: Know the difference between minerals and rocks!

If you are currently taking a geology course, you’ll likely be asked more than once to name a mineral or a rock that has specific characteristics or composition, or was formed in a specific environment. Please make sure that if you’re asked for a rock name that you don’t respond with a mineral name, and vice versa. Confusing minerals and rocks is one of the most common mistakes that geology students make.

Exercise 1.1 Find a piece of granite

The rock granite is very common in most parts of North America, and unless everything is currently covered with snow where you live, you should have no trouble finding a sample of it near you. The best places to look are pebbly ocean or lake beaches, a gravel bar of a creek or river, a gravel driveway, or somewhere where gravel has been used in landscaping. In Figure 1.4.3, taken on a beach, the granitic pebbles are the ones that are predominantly light-coloured with dark specks. The one in the very centre is a good example.

A pile of different coloured rocks. Granite rocks are white with dark specks
Figure 1.4.3

Select a sample of granite and, referring to Figure 1.4.2, see if you can identify some of the minerals in it. It may help to break it in half with a hammer to see a fresh surface, but be careful to protect your eyes if you do so. You should be able to see glassy-looking quartz, dull white plagioclase feldspar (and maybe pink potassium feldspar), and black hornblende or, in some cases, flaky black biotite mica (or both).

In addition to identifying the minerals in your granite, you might also try to describe the texture in terms of the range sizes of the mineral crystals (in millimetres) and the shapes of the crystals (some may be rectangular in outline, most will be irregular). Think about where your granite might have come from and how it got to where you found it.

See Appendix 3 for Exercise 1.1 answers.

Media Attributions


1.5 Fundamentals of Plate Tectonics

Plate tectonics is the model or theory that has been used for the past 60 years to understand and explain how the Earth works—more specifically the origins of continents and oceans, of folded rocks and mountain ranges, of earthquakes and volcanoes, and of continental drift. Plate tectonics is explained in some detail in Chapter 10, but is introduced here because it includes concepts that are important to many of the topics covered in the next few chapters.

Figure 1.5.1 The components of the interior of the Earth (click on the image to see a full-size version).

Key to understanding plate tectonics is an understanding of Earth’s internal structure, which is illustrated in Figure 1.5.1. Earth’s core consists mostly of iron. The outer core is hot enough for the iron to be liquid. The inner core—although even hotter—is under so much pressure that it is solid. The mantle is made up of iron and magnesium silicate minerals. The bulk of the mantle surrounding the outer core is solid rock, but is plastic enough to be able to flow slowly. The outermost part of the mantle is rigid. The crust—composed mostly of granite on the continents and mostly of basalt beneath the oceans—is also rigid. The crust and outermost rigid mantle together make up the lithosphere. The lithosphere is divided into about 20 tectonic plates that move in different directions on Earth’s surface.

An important property of Earth (and other planets) is that the temperature increases with depth, from close to 0°C at the surface to about 7000°C at the centre of the core. In the crust, the rate of temperature increase is about 30°C every kilometre. This is known as the geothermal gradient.

Heat is continuously flowing outward from Earth’s interior, and the transfer of heat from the core to the mantle causes convection in the mantle (Figure 1.5.2). This convection is the primary driving force for the movement of tectonic plates. At places where convection currents in the mantle are moving upward, new lithosphere forms (at ocean ridges), and the plates move apart (diverge). Where two plates are converging (and the convective flow is downward), one plate will be subducted (pushed down) into the mantle beneath the other. Many of Earth’s major earthquakes and volcanoes are associated with convergent boundaries.

The movement of currents in the Earth's mantle puts pressure on the Lithosphere are causes plates to move
Figure 1.5.2 Depiction of the convection in the mantle and it’s relationship to plate motion

Earth’s major tectonic plates and the directions and rates at which they are diverging at sea-floor ridges, are shown in Figure 1.5.3.

Exercise 1.2 Plate

Using either a map of the tectonic plates from the Internet or Figure 1.5.3 determine which tectonic plate you are on right now, approximately how fast it is moving, and in what direction. How far has that plate moved relative to Earth’s core since you were born?

Figure 1.5.3 A map showing 15 of the Earth’s tectonic plates and the approximate rates and directions of plate motions.


See Appendix 3 for Exercise 1.2 answers.

Media Attributions


1.6 Geological Time

In 1788, after many years of geological study, James Hutton, one of the great pioneers of geology, wrote the following about the age of Earth: The result, therefore, of our present enquiry is, that we find no vestige of a beginning — no prospect of an end.Hutton, J, 1788. Theory of the Earth; or an investigation of the laws observable in the composition, dissolution, and restoration of land upon the Globe. Transactions of the Royal Society of Edinburgh. Of course he wasn’t exactly correct, there was a beginning and there will be an end to Earth, but what he was trying to express is that geological time is so vast that we humans, who typically live for less than a century, have no means of appreciating how much geological time there is. Hutton didn’t even try to assign an age to Earth, but we now know that it is approximately 4,570 million years old. Using the scientific notation for geological time, that is 4,570 Ma (for mega annum or “millions of years”) or 4.57 Ga (for giga annum or billions of years). More recent dates can be expressed in ka (kilo annum); for example, the last cycle of glaciation ended at approximately 11.7 ka or 11,700 years ago. This notation will be used for geological dates throughout this book.

Exercise 1.3 Using geological time notation

To help you understand the scientific notation for geological time—which is used extensively in this book—write the following out in numbers (for example, 3.23 Ma = 3,230,000 years).

  1. 2.75 ka
  2. 0.93 Ga
  3. 14.2 Ma

We use this notation to describe geological events in the same way that we might say “they arrived at 2 pm.”  For example, we can say “this rock formed at 45 Ma.” But this notation is not used to express elapsed time. We don’t say: “I studied for 4 pm for that test.” And we don’t say: “The dinosaurs lived for 160 Ma.”  Instead, we could say: “The dinosaurs lived from 225 Ma to 65 Ma, which is 160 million years.”

See Appendix 3 for Exercise 1.3 answers.

Unfortunately, knowing how to express geological time doesn’t really help us understand or appreciate its extent. A version of the geological time scale is included as Figure 1.6.1. Unlike time scales you’ll see in other places, or even later in this book, this time scale is linear throughout its length, meaning that 50 Ma during the Cenozoic is the same thickness as 50 Ma during the Hadean—in each case about the height of the “M” in Ma. The Pleistocene glacial epoch began at about 2.6 Ma, which is equivalent to half the thickness of the thin grey line at the top of the yellow bar marked “Cenozoic.” Most other time scales have earlier parts of Earth’s history compressed so that more detail can be shown for the more recent parts. That makes it difficult to appreciate the extent of geological time.

Geological time scale. Image description available.
Figure 1.6.1 The geological time scale. [Image Description]

To create some context, the Phanerozoic Eon (the last 542 million years) is named for the time during which visible (phaneros) life (zoi) is present in the geological record. In fact, large organisms—those that leave fossils visible to the naked eye—have existed for a little longer than that, first appearing around 600 Ma, or a span of just over 13% of geological time. Animals have been on land for 360 million years, or 8% of geological time. Mammals have dominated since the demise of the dinosaurs around 65 Ma, or 1.5% of geological time, and the genus Homo has existed since approximately 2.8 Ma, or 0.06% (1/1,600th) of geological time.

Geologists (and geology students) need to understand geological time. That doesn’t mean memorizing the geological time scale; instead, it means getting your mind around the concept that although most geological processes are extremely slow, very large and important things can happen if such processes continue for enough time.

For example, the Atlantic Ocean between Nova Scotia and northwestern Africa has been getting wider at a rate of about 2.5 centimetres (cm) per year. Imagine yourself taking a journey at that rate—it would be impossibly and ridiculously slow. And yet, since it started to form at around 200 Ma (just 4% of geological time), the Atlantic Ocean has grown to a width of over 5,000 kilometres (km)!

A useful mechanism for understanding geological time is to scale it all down into one year. The origin of the solar system and Earth at 4.57 Ga would be represented by January 1, and the present year would be represented by the last tiny fraction of a second on New Year’s Eve. At this scale, each day of the year represents 12.5 million years; each hour represents about 500,000 years; each minute represents 8,694 years; and each second represents 145 years. Some significant events in Earth’s history, as expressed on this time scale, are summarized on Table 1.1.

Table 1.1 A summary of some important geological dates expressed as if all of geological time was condensed into one year.
[Skip Table]
Event Approximate Date Calendar Equivalent
Formation of oceans and continents 4.5 to 4.4 Ga January
Evolution of the first primitive life forms 3.8 Ga early March
Formation of British Columbia’s oldest rocks 2.0 Ga July
Evolution of the first multi-celled animals 0.6 Ga or 600 Ma November 15
Animals first crawled onto land 360 Ma December 1
Vancouver Island reached North America and the Rocky Mountains were formed 90 Ma December 25
Extinction of the non-avian dinosaurs 65 Ma December 26
Beginning of the Pleistocene ice age 2 Ma or 2000 ka 8 p.m., December 31
Retreat of the most recent glacial ice from southern Canada 14 ka 11:58 p.m., December 31
Arrival of the first people in British Columbia 10 ka 11:59 p.m., December 31
Arrival of the first Europeans on the west coast of what is now Canada 250 years ago 2 seconds before midnight, December 31

Exercise 1.4 Take a trip through geological time

We’re going on a road trip! Pack some snacks and grab some of your favourite music. We’ll start in Tofino on Vancouver Island and head for the Royal Tyrrell Museum just outside of Drumheller, Alberta, 1,500 km away. Along the way, we’ll talk about some important geological sites that we pass by, and we’ll use the distance as a way of visualizing the extent of geological time. Of course it’s just a “virtual” road trip, but it will be fun anyway. To join in, go to: Virtual Road Trip.

Once you’ve had a chance to do the road trip, answer these questions:

  1. We need oxygen to survive, and yet the first presence of free oxygen (O2 gas) in the atmosphere and the oceans was a “catastrophe” for some organisms. When did this happen and why was it a catastrophe?
  2. Approximately how much time elapsed between the colonization of land by plants and animals?
  3. Explain why the evolution of land plants was such a critical step in the evolution of life on Earth.

See Appendix 3 for Exercise 1.4 answers.

Image descriptions

Figure 1.6.1 image description: The Hadean eon (3800 Ma to 4570 Ma), Archean eon (2500 Ma to 3800 Ma), and Proterozoic eon (542 Ma to 2500 Ma) make up 88% of geological time. The Phanerozoic eon makes up the last 12% of geological time. The Phanerozoic eon (0 Ma to 542 Ma) contains the Paleozoic, Mesozoic, and Cenozoic eras. [Return to Figure 1.6.1]

Media Attributions



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

Section Summary
1.1 What is Geology? Geology is the study of Earth. It is an integrated science that involves the application of many of the other sciences, but geologists also have to consider geological time because most of the geological features that we see today formed thousands, millions, or even billions of years ago.
1.2 Why Study Earth? Geologists study Earth out of curiosity and for other more practical reasons, including understanding the evolution of life on Earth, searching for resources, understanding risks from geological events such as earthquakes, volcanoes, and slope failures, and documenting past environmental and climate changes so that we can understand how human activities are affecting Earth.
1.3 What Do Geologists Do? Geologists work in the resource industries and in efforts to protect our natural resources and the environment in general. They are involved in ensuring that risks from geological events (e.g., earthquakes) are minimized and that the public understands what the risks are. Geologists are also engaged in fundamental research about Earth and in teaching.
1.4 Minerals and Rocks Minerals are naturally occurring, specific combinations of elements that have particular three-dimensional structures. Rocks are made up of mixtures of minerals and can form though igneous, sedimentary, or metamorphic processes.
1.5 Fundamentals of Plate Tectonics The Earth’s mantle is convecting because it is being heated from below by the hot core. Those convection currents contribute to the movement of tectonic plates (which are composed of the crust and the uppermost rigid mantle). Plates are formed at divergent boundaries and consumed (subducted) at convergent boundaries. Many important geological processes take place at plate boundaries.
1.6 Geological Time Earth is approximately 4,570,000,000 years old; that is, 4.57 billion years or 4.57 Ga or 4,570 Ma. It’s such a huge amount of time that even extremely slow geological processes can have an enormous impact.

Questions for Review

Answers to Review Questions at the end of each chapter are provided in Appendix 2.

  1. In what way is geology different from the other sciences, such as chemistry and physics?
  2. How would some familiarity with biology be helpful to a geologist?
  3. List three ways in which geologists can contribute to society.
  4. Describe the lattice structure and elemental composition of the mineral halite.
  5. In what way is a mineral different from a rock?
  6. What is the main component of Earth’s core?
  7. What process leads to convection in the mantle?
  8. How does mantle convection contribute to plate tectonics?
  9. What are some of the processes that take place at a divergent plate boundary?
  10. Dinosaurs first appear in the geological record in rocks at 225 Ma and then disappear at 65 Ma. For what proportion (%) of geological time did dinosaurs exist?
  11. If a typical rate for the accumulation of sediments is 1 mm/year, what thickness (metres) of sedimentary rock could accumulate over a period of 30 million years?


Chapter 2 Minerals

Learning Objectives

After reading this chapter, completing the exercises within it, and answering the questions at the end, you should be able to:

  • Describe the nature of atoms and their constituents, particularly the behaviour of electrons and the formation of ions.
  • Apply your understanding of atoms to explain bonding within minerals.
  • Describe mineral lattices and explain how they influence mineral properties.
  • Categorize minerals into groups based on their compositions.
  • Describe a silica tetrahedron and the ways in which tetrahedra combine to make silicate minerals.
  • Differentiate between ferromagnesian and other silicate minerals.
  • Explain some of the mechanisms of mineral formation.
  • Describe some of the important techniques for identifying minerals.

Minerals are all around us: the graphite in your pencil, the salt on your table, the plaster on your walls, and the trace amounts of gold in your computer. Minerals can be found in a wide variety of consumer products including paper, medicine, processed foods, cosmetics, electronic devices, and many more. And of course, everything made of metal is also derived from minerals.

As defined in Chapter 1, a mineral is a naturally occurring combination of specific elements arranged in a particular repeating three-dimensional structure (Figure 1.4.1).

Naturally occurring” implies that minerals are not artificially made.  Many minerals (e.g., diamond) can be made in laboratories, but if they can also occur naturally, they still qualify as minerals.

Specific elements” means that most minerals have a specific chemical formula or composition. The mineral pyrite, for example, is FeS2 (two atoms of sulfur for each atom of iron), and any significant departure from that would make it a different mineral. But many minerals can have variable compositions within a specific range. The mineral olivine, for example, can range all the way from Fe2SiO4 to FeMgSiOto Mg2SiO4. Intervening compositions are written as (Fe,Mg)2SiO4 meaning that Fe and Mg can be present in any proportion, and that there are two of them for each Si present. This type of substitution is known as solid solution.

Most important of all, a mineral has a specific “repeating three-dimensional structure” or “lattice,” which is the way in which the atoms are arranged. We’ve already seen in Chapter 1 how sodium and chlorine atoms in halite alternate in a regular pattern. That happens to be about the simplest mineral lattice of all; most mineral lattices are much more complicated, as we’ll see.

Some substances that we think must be minerals are not because they lack that repeating 3-dimensional structure of atoms.  Volcanic glass is an example, as is pearl or opal. As shown in Figure 2.0.1, opal appears to have a regular structure, but it’s not an atomic structure.

Figure 2.0.1  Opal (left) is like a mineral, but does not have a crystalline structure. Instead, it is made up of layers of closely packed silica spheres (right).

Media Attributions


2.1 Electrons, Protons, Neutrons, and Atoms

All matter that we are familiar with, including mineral crystals, is made up of atoms, and all atoms are made up of three main particles: protons, neutrons, and electrons. As summarized in Table 2.1, protons are positively charged, neutrons are uncharged and electrons are negatively charged. The −1 charge of one electron balances the +1 charge of one proton. Both protons and neutrons have a mass of 1, while electrons have almost no mass.

Table 2.1 Charges and masses of the particles within atoms
Elementary Particle Charge Mass
Proton +1 1
Neutron 0 1
Electron −1 ~0

The element hydrogen has the simplest atoms, each with just one proton and one electron. The proton forms the nucleus, while the electron orbits around it. All other elements have neutrons as well as protons in their nucleus, such as helium, which is depicted in Figure 2.1.1. The positively charged protons tend to repel each other, but the neutrons help to hold the nucleus together. The number of protons is the atomic number, and the number of protons plus neutrons is the atomic mass. For hydrogen, the atomic mass is 1 because there is one proton and no neutrons. For helium, it is 4: two protons and two neutrons.

For most of the 16 lightest elements (up to oxygen) the number of neutrons is equal to the number of protons. For most of the remaining elements there are more neutrons than protons because extra neutrons are needed to keep the nucleus together by overcoming the mutual repulsion of the increasing numbers of protons concentrated in a very small space. For example, silicon has 14 protons and 14 neutrons. Its atomic number is 14 and its atomic mass is 28. The most common isotope of uranium has 92 protons and 146 neutrons. Its atomic number is 92 and its atomic mass is 238 (92 + 146).

A dot inside a circle that is darker in the centre and grows lighter near the edge.
Figure 2.1.1 A depiction of a helium atom.

A helium atom is depicted on Figure 2.1.1. The dot in the middle is the nucleus, and the surrounding cloud represents where the two electrons might be at any time. The darker the shade, the more likely that an electron will be there. The helium atom is about 1 angstrom across.  An angstrom (Å) is 10−10 metres (m). The helium nucleus is about 1 femtometre across.  A femtometre (fm) is 10−15 m. In other words, a helium atom’s electron cloud is about 100,000 times bigger than its nucleus.  Stanley Park in Vancouver is about 2 km across.  If Stanley Park was a helium atom, the nucleus would be the size of a walnut.

Electrons orbiting around the nucleus of an atom are arranged in shells—also known as “energy levels.” The first shell can hold only two electrons, while the next shell holds up to eight electrons. Subsequent shells can hold more electrons, but the outermost shell of any atom holds no more than eight electrons. As we’ll see, the electrons in the outermost shell play an important role in bonding between atoms. The electron shell configurations for 29 of the first 36 elements are listed in Table 2.2.

Table 2.2 Electron shell configurations up to element 36. (The inert elements, with filled outer shells, have a *.)
[Skip Table]
Element Symbol Atomic No. Number of Electrons in the 1st Shell Number of Electrons in the 2nd Shell Number of Electrons in the 3rd Shell Number of Electrons in the 4th Shell
Hydrogen H 1 1  0  0 0
Helium * He 2 2  0  0  0
Lithium Li 3 2 1  0  0
Beryllium Be 4 2 2  0  0
Boron B 5 2 3  0  0
Carbon C 6 2 4  0  0
Nitrogen N 7 2 5  0  0
Oxygen O 8 2 6  0  0
Fluorine F 9 2 7  0  0
Neon * Ne 10 2 8  0  0
Sodium Na 11 2 8 1  0
Magnesium Mg 12 2 8 2  0
Aluminum Al 13 2 8 3  0
Silicon Si 14 2 8 4  0
Phosphorus P 15 2 8 5  0
Sulphur S 16 2 8 6  0
Chlorine Cl 17 2 8 7  0
Argon * Ar 18 2 8 8  0
Potassium K 19 2 8 8 1
Calcium Ca 20 2 8 8 2
Scandium Sc 21 2 8 9 2
Titanium Ti 22 2 8 10 2
Vanadium V 23 2 8 11 2
Chromium Cr 24 2 8 13 1
Manganese Mn 25 2 8 13 2
Iron Fe 26 2 8 14 2
. . . . . . .
Selenium Se 34 2 8 18 6
Bromine Br 35 2 8 18 7
Krypton * Kr 36 2 8 18 8


Media Attributions


2.2 Bonding and Lattices

As we’ve just seen, an atom seeks to have a full outer shell (i.e., eight electrons for most elements, or two electrons for hydrogen and helium) to be atomically stable. This is accomplished by transferring or sharing electrons with other atoms.

An ionic bond. Image description available
Figure 2.2.1 A very simplified 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). [Image Description]

Sodium has 11 electrons: two in the first shell, eight in the second, and one in the third (Figure 2.2.1). Sodium readily gives up that single third-shell electron, and when it loses this one negative charge, it becomes positively charged (because it now has 11 protons and only 10 electrons). By giving up its lone third-shell electron, sodium ends up with a full outer shell. Chlorine, on the other hand, has 17 electrons: two in the first shell, eight in the second, and seven in the third. Chlorine readily accepts an eighth electron to fill its third shell, and therefore becomes negatively charged because it has 17 protons and 18 electrons. In changing their number of electrons, these atoms become ions—the sodium loses an electron to become a positive ion or cation, and the chlorine gains an electron to become a negative ion or anion (Figure 2.2.1).

Since negative and positive charges attract, sodium and chlorine ions can stick together, creating an ionic bond. Electrons can be thought of as being transferred from one atom to another in an ionic bond. Common table salt (NaCl) is a mineral composed of chlorine and sodium linked together by ionic bonds (Figure 1.4.1). The mineral name for NaCl is halite.

An element like chlorine can also form bonds without forming ions. For example, two chlorine atoms, which each seek an eighth electron in their outer shell, can share an electron in what is known as a covalent bond to form chlorine gas (Cl2) (Figure 2.2.2). Electrons are shared in a covalent bond.

Each chlorine atom has 7 electrons in its outer shell. They each share one electron to form a covalent bond
Figure 2.2.2 Depiction of a covalent bond between two chlorine atoms. The electrons are black in the left atom and blue in the right atom. Two electrons are shared (one black and one blue) so that each atom “appears” to have a full outer shell.

Exercise 2.1 Cations, anions, and ionic bonding

A number of elements are listed below along with their atomic numbers. Assuming that the first electron shell can hold two electrons and subsequent electron shells can hold eight electrons, sketch in the electron configurations for these elements. Predict whether the element is likely to form a cation (+) or an anion (−) when electron transfer takes place, and what charge it would have (e.g., +1, +2, −1).

The first one is done for you.  Fluorine needed an extra electron to have 8 in its outermost shell, and in gaining that electron it became negatively charged.

See Appendix 3 for Exercise 2.1 answers.

Fluorine (9)
The Fluorine atom has 2 electrons in inner shell, 7 electrons in outer shell. It is an anion, charge negative 1

anion (−1)

Lithium (3)
A Lithium atom with no electrons filled in.


Magnesium (12)
A Magnesium atom with no electrons filled in.


Argon (18)
An Argon atom with no electrons filled in.


 Chlorine (17)
A Chlorine atom with no electrons filled in.


Beryllium (4)
An Beryllium atom with no electrons filled in.


Oxygen (8)
An Oxygen atom with no electrons filled in.


 Sodium (11)
A Sodium atom with no electrons filled in.


A carbon atom has six protons and six electrons; two of the electrons are in the inner shell and four in the outer shell (Figure 2.2.3). Carbon would need to gain or lose four electrons to have a filled outer shell, and this would create too great a charge imbalance for the ion to be stable. On the other hand, carbon can share electrons to create covalent bonds. In the mineral diamond, the carbon atoms are linked together in a three-dimensional framework, where each carbon atom is bonded to four other carbon atoms and every bond is a very strong covalent bond. In the mineral graphite, the carbon atoms are linked together in sheets or layers (Figure 2.2.3), and each carbon atom is covalently bonded to three others. Graphite-based compounds, which are strong because of the strong intra-layer covalent bonding, are used in high-end sports equipment such as ultralight racing bicycles. Graphite itself is soft because the bonding between these layers is relatively weak, and it is used in a variety of applications, including lubricants and pencils.

Carbon covalent bond. Image description available.
Figure 2.2.3 The electron configuration of carbon (left) and the sharing of electrons in covalent bonding of diamond (right). The electrons shown in blue are shared between adjacent Carbon atoms. Although shown here in only two dimensions, diamond has a three-dimensional structure as shown on Figure 2.2.5. [Image description]

Silicon and oxygen bond together to create a silica tetrahedron, which is a four-sided pyramid shape with O at each corner and Si in the middle (Figure 2.2.4). This structure is the building block of the silicate minerals (which are described in Section 2.4). 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 (SiO4) is: 4 + 4(−2) = 4 − 8 = −4. As we will see later, silica tetrahedra (plural of tetrahedron) link together in a variety of ways to form most of the common minerals of the crust.

A silicon ion bonded to four oxygen ions to form a pyramid shape
Figure 2.2.4 The silica tetrahedron, the building block of all silicate minerals. (Because the silicon ion has a charge of +4 and the four oxygen ions each have a charge of −2, the silica tetrahedron has a net charge of −4.)

Most minerals are characterized by ionic bonds, covalent bonds, or a combination of the two, but there are other types of bonds that are important in minerals, including metallic bonds and weaker electrostatic forces (hydrogen or Van der Waals bonds). Metallic elements have outer electrons that are relatively loosely held. (The metals are highlighted on the periodic table in Appendix 1.) When bonds between such atoms are formed, these electrons can move freely from one atom to another. A metal can thus be thought of as an array of positively charged atomic nuclei immersed in a sea of mobile electrons. This feature accounts for two very important properties of metals: their electrical conductivity and their malleability (they can be deformed and shaped).

Molecules that are bonded ionically or covalently can also have other weaker electrostatic forces holding them together. Examples of this are the force holding graphite sheets together and the attraction between water molecules.

What’s with all of these “sili” names?

The element silicon is one of the most important geological elements and is the second-most abundant element in Earth’s crust (after oxygen). Silicon bonds readily with oxygen to form a silica tetrahedron (Figure 2.2.4). Pure silicon crystals (created in a lab) are used to make semi-conductive media for electronic devices. A silicate mineral is one in which silicon and oxygen are present as silica tetrahedra. Silica also refers to a chemical component of a rock and is expressed as % SiO2. The mineral quartz is made up entirely of silica tetrahedra, and some forms of quartz are also known as “silica”. Silicone is a synthetic product (e.g., silicone rubber, resin, or caulking) made from silicon-oxygen chains and various organic molecules. To help you keep the “sili” names straight, here is a summary table:

“Sili” name Definition
Table 2.3 Summary of “Sili” names
Silicon The 14th element
Silicon wafer A crystal of pure silicon sliced very thinly and used for electronics
Silica tetrahedron A combination of one silicon atom and four oxygen atoms that form a tetrahedron
% silica The proportion of a rock that is composed of the component SiO2
Silica A solid made out of SiO(but not necessarily a mineral – e.g., opal)
Silicate A mineral that contains silica tetrahedra (e.g., quartz, feldspar, mica, olivine)
Silicone A flexible synthetic material made up of Si–O chains with attached organic molecules

Elements that have a full outer shell are described as inert because they do not tend to react with other elements to form compounds.  That’s because they don’t need to lose or gain any electrons to become stable, and so they don’t become ions. They all appear in the far-right column of the periodic table. Examples are: helium, neon, argon, etc.

As described in Chapter 1, all minerals are characterized by a specific three-dimensional pattern known as a lattice or crystal structure. These structures range from the simple cubic pattern of halite (NaCl) (Figure 1.4.1), to the very complex patterns of some silicate minerals. Two minerals may have the same composition, but very different crystal structures and properties. Graphite and diamond, for example, are both composed only of carbon, but while diamond is the hardest substance known, graphite is softer than paper. Their lattice structures are compared in Figure 2.2.5.

Graphite has a mix of strong covalent bonds and weak inter-layer bonds. Diamonds only have strong covalent bonds
Figure 2.2.5 A depiction of the lattices of graphite and diamond.

Mineral lattices have important implications for mineral properties, as exemplified by the hardness of diamond and the softness of graphite. Lattices also determine the shape that mineral crystals grow in and how they break. For example, the right angles in the lattice of the mineral halite (Figure 1.4.1) influence both the shape of its crystals (cubic), and the way those crystals break (Figure 2.2.6).

Figure 2.2.6 Cubic crystals (left) and right-angle cleavage planes (right) of the mineral halite. If you look closely at the cleavage fragment on the right, you can see where it would break again (cleave) along a plane parallel to an existing surface. In most minerals, cleavage planes do not align with crystal surfaces.

Image Descriptions

Figure 2.2.1 image description: Sodium has one electron in its outer shell and chlorine has 7 electrons in it its outer shell. Sodium’s one outer electron goes to chlorine which makes Chlorine slightly negative and Sodium slightly positive. They attract each other and together they form Sodium Chloride. [Return to Figure 2.2.1]

Figure 2.2.3 image description: (Left) A carbon atom has two electrons in its inner shell and four electrons in its outer shell. (Right) One Carbon atom shares electrons with four other carbon atoms to form a complete outer shell. [Return to Figure 2.2.3]

Media Attributions


2.3 Mineral Groups

Most minerals are made up of a cation (a positively charged ion) or several cations, plus an anion (a negatively charged ion (e.g., S2−)) or an anion complex (e.g., SO42−). For example, in the mineral hematite (Fe2O3), the cation is Fe3+ (iron) and the anion is O2− (oxygen). The two Fe3+ ions have an overall charge of +6 and that balances the overall charge of −6 from the three O2− ions.

We group minerals into classes on the basis of their predominant anion or anion complex. These include oxides, sulphides, carbonates, silicates, and others. Silicates are by far the predominant group in terms of their abundance within the crust and mantle. (They will be discussed in Section 2.4). Some examples of minerals from the different mineral groups are given in Table 2.4.

Table 2.4 The main mineral groups and some examples of minerals in each group.
[Skip Table]
Group Examples
Oxides Hematite (iron oxide Fe2O3), corundum (aluminum oxide Al2O3), water ice (H2O)
Sulphides Galena (lead sulphide PbS), pyrite (iron sulphide FeS2), chalcopyrite (copper-iron sulphide CuFeS2)
Sulphates Gypsum (calcium sulphate CaSO4·H2O), barite (barium sulphate BaSO4) (Note that sulphates are different from sulphides. Sulphates have the SO4−2 ion while sulphides have the S−2 ion)
Halides Fluorite (calcium flouride CaF2), halite (sodium chloride NaCl) (Halide minerals have halogen elements as their anion — the minerals in the second last column on the right side of the periodic table, including F, Cl, Br, etc. — see the periodic table in Appendix 1: List of Geologically Important Elements and the Periodic Table.)
Carbonates Calcite (calcium carbonate CaCO3), dolomite (calcium-magnesium carbonate (Ca,Mg)CO3)
Phosphates Apatite (Ca5(PO4)3(OH)), Turquoise (CuAl6(PO4)4(OH)8·5H2O)
Silicates Quartz (SiO2), feldspar (sodium-aluminum silicate NaAlSi3O8), olivine (iron or magnesium silicate (Mg,Fe)2SiO4)   (Note that in quartz the anion is oxygen, and while it could be argued, therefore, that quartz is an oxide, it is always classed with the silicates.)
Native minerals Gold (Au), diamond (C), graphite (C), sulphur (S), copper (Cu)

Oxide minerals have oxygen (O2−) as their anion, but they exclude those with oxygen complexes such as carbonate (CO32−), sulphate (SO42−), and silicate (SiO44−). The most important oxides are the iron oxides hematite and magnetite (Fe2O3 and Fe3O4, respectively). Both of these are common ores of iron. Corundum (Al2O3) is used as an abrasive, but can also be a gemstone in its ruby and sapphire varieties. If the oxygen is also combined with hydrogen to form the hydroxyl anion (OH) the mineral is known as a hydroxide. Some important hydroxides are limonite and bauxite, which are ores of iron and aluminium respectively. Frozen water (H2O) is a mineral (an oxide), but liquid water is not because it doesn’t have a regular lattice.

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. Some other sulphide minerals are pyrite (FeS2), bornite (Cu5FeS4), stibnite (Sb2S3), and arsenopyrite (FeAsS).

Sulphates are minerals with the SO4−2 anion, and these include anhydrite (CaSO4) and its cousin gypsum (CaSO4.2H2O) and the sulphates of barium and strontium: barite (BaSO4) and celestite (SrSO4). In all of these minerals, 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), cryolite (Na3AlF6), 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)The notations of two (or more) elements enclosed in parentheses with a comma between them: (Ca,Mg), indicates that both can be present, in varying proportions, but that there is still only one of them for each anion present., and siderite (FeCO3). The copper minerals malachite and azurite are also carbonates.

In phosphate minerals, the anion is the PO4−3 complex. An important phosphate mineral is apatite (Ca5(PO4)3(OH)), which is what your teeth are made of. Note that it is called a phosphate, not a hydroxide, even though it has a hydroxyl ion.

The silicate minerals include the elements silicon and oxygen in varying proportions ranging from Si : O2 to Si : O4. These are discussed at length in Section 2.4.

Native minerals are single-element minerals, such as gold, copper, sulphur, and graphite.

Exercise 2.2 Mineral groups

We classify minerals according to the anion part of the mineral formula, and mineral formulas are always written with the anion part on the right. For example, for pyrite (FeS2), Fe2+ is the cation and S is the anion. This helps us to know that it’s a sulphide, but it is not always that obvious. Hematite (Fe2O3) is an oxide; that’s easy, but anhydrite (CaSO4) is a sulphate because SO42− is the anion, not O. Along the same lines, calcite (CaCO3) is a carbonate, and olivine (Mg2SiO4) is a silicate. Minerals with only one element (such as S) are native minerals, while those with an anion from the halogen column of the periodic table (Cl, F, Br, etc.) are halides. Provide group names for the following minerals:

Table 2.5 Provide group names for the following minerals
[Skip Table]
Name Formula Group
sphalerite ZnS
magnetite Fe3O4
pyroxene MgSiO3
anglesite PbSO4
sylvite KCl
silver Ag
fluorite CaF2
ilmenite FeTiO3
siderite FeCO3
feldspar KAlSi3O8
sulphur S
xenotime YPO4

See Appendix 3 for Exercise 2.2 answers.


2.4 Silicate Minerals

The vast majority of the minerals that make up the rocks of Earth’s crust are silicate minerals. These include minerals such as quartz, feldspar, mica, amphibole, pyroxene, olivine, and a variety of clay minerals. The building block of all of these minerals is the silica tetrahedron, a combination of four oxygen atoms and one silicon atom. As we’ve seen, it’s called a tetrahedron because planes drawn through the oxygen atoms form a shape with 4 surfaces (Figure 2.2.4). Since the silicon ion has a charge of 4 and each of the four oxygen ions has a charge of −2, the silica tetrahedron has a net charge of −4.

In silicate minerals, these tetrahedra are arranged and linked together in a variety of ways, from single units to complex frameworks (Table 2.6). The simplest silicate structure, that of the mineral olivine, is composed of isolated tetrahedra bonded to iron and/or magnesium ions. In olivine, the −4 charge of each silica tetrahedron is balanced by two divalent (i.e., +2) iron or magnesium cations. Olivine can be either Mg2SiO4 or Fe2SiO4, or some combination of the two (Mg,Fe)2SiO4. The divalent cations of magnesium and iron are quite close in radius (0.73 versus 0.62 angstromsAn angstrom is the unit commonly used for the expression of atomic-scale dimensions. One angstrom is 10−10 metres or 0.0000000001 metres. The symbol for an angstrom is Å.). Because of this size similarity, and because they are both divalent cations (both can have a charge of +2), iron and magnesium can readily substitute for each other in olivine and in many other minerals.

Table 2.6 Silicate mineral configurations. The triangles represent silica tetrahedra.
[Skip Table]
Tetrahedron Configuration Picture Tetrahedron Configuration Name Example Minerals
 One triangle Isolated (nesosilicates) Olivine, garnet, zircon, kyanite
Two triangles joined at their tips. Pairs (sorosilicates) Epidote, zoisite
 Six triangles joined together in a circle to form a star Rings (cyclosilicates) Tourmaline
Five triangles joined together in a line. Single chains (inosilicates) Pyroxenes, wollastonite
 Two rows of triangles joined together Double chains (inosilicates) Amphiboles
 Multiple rows of triangles joined together Sheets (phyllosilicates) Micas, clay minerals, serpentine, chlorite
3-dimensional structure Framework (tectosilicates) Feldspars, quartz, zeolite

Exercise 2.3 Make a Tetrahedron

Cut around the outside of the shape (solid lines and dotted lines), and then fold along the solid lines to form a tetrahedron. If you have glue or tape, secure the tabs to the tetrahedron to hold it together. If you don’t have glue or tape, make a slice along the thin grey line and insert the pointed tab into the slit.

If you are doing this in a classroom, try joining your tetrahedron with others into pairs, rings, single and double chains, sheets, and even three-dimensional frameworks.

See Appendix 3 for Exercise 2.3 answers.

In olivine, unlike most other silicate minerals, the silica tetrahedra are not bonded to each other. Instead they are bonded to the iron and/or magnesium ions, in the configuration shown on Figure 2.4.1.

Figure 2.4.1 A depiction of the structure of olivine as seen from above. The formula for this particular olivine, which has three Fe ions for each Mg ion, could be written: Mg0.5Fe1.5SiO4.

As already noted, the 2 ions of iron and magnesium are similar in size (although not quite the same). This allows them to substitute for each other in some silicate minerals. In fact, the ions that are common in silicate minerals have a wide range of sizes, as depicted in Figure 2.4.2. All of the ions shown are cations, except for oxygen. Note that iron can exist as both a +2 ion (if it loses two electrons during ionization) or a +3 ion (if it loses three). Fe2+  is known as ferrous iron. Fe3+  is known as ferric iron. Ionic radii are critical to the composition of silicate minerals, so we’ll be referring to this diagram again.

Figure 2.4.2 The ionic radii (effective sizes) in angstroms, of some of the common ions in silicate minerals. [Image Description]

The structure of the single-chain silicate pyroxene is shown on Figures 2.4.3 and 2.4.4. In pyroxene, silica tetrahedra are linked together in a single chain, where one oxygen ion from each tetrahedron is shared with the adjacent tetrahedron, 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), and the net charge per silicon atom is less (−2 instead of −4).  Therefore, fewer cations are necessary to balance that charge. Pyroxene compositions are of the type MgSiO3, FeSiO3, and CaSiO3, or some combination of these. Pyroxene can also be written as (Mg,Fe,Ca)SiO3, where the elements in the brackets can be present in any proportion. In other words, pyroxene has one cation for each silica tetrahedron (e.g., MgSiO3) while olivine has two (e.g., Mg2SiO4). Because each silicon ion is +4 and each oxygen ion is −2, the three oxygens (−6) and the one silicon (+4) give a net charge of −2 for the single chain of silica tetrahedra. In pyroxene, the one divalent cation (2) per tetrahedron balances that −2 charge. In olivine, it takes two divalent cations to balance the −4 charge of an isolated tetrahedron.The structure of pyroxene is more “permissive” than that of olivine—meaning that cations with a wider range of ionic radii can fit into it. That’s why pyroxenes can have iron (radius 0.63 Å) or magnesium (radius 0.72 Å) or calcium (radius 1.00 Å) cations (see Figure 2.4.2 above).

Three parallel chains with a rows of positive 2 cations in between them
Figure 2.4.3 A depiction of the structure of pyroxene. The tetrahedral chains continue to left and right and each is interspersed with a series of divalent cations. If these are Mg ions, then the formula is MgSiO3.
Figure 2.4.4 A single silica tetrahedron (left) with four oxygen ions per silicon ion (SiO4). Part of a single chain of tetrahedra (right), where the oxygen atoms at the adjoining corners are shared between two tetrahedra (arrows). For a very long chain the resulting ratio of silicon to oxygen is 1 to 3 (SiO3).

Exercise 2.4 Oxygen deprivation

The diagram below represents a single chain in a silicate mineral. Count the number of tetrahedra versus the number of oxygen ions (yellow spheres). Each tetrahedron has one silicon ion so this should give you the ratio of Si to O in single-chain silicates (e.g., pyroxene).

A chain of six tetrahedra and 21 oxygen ions

The diagram below represents a double chain in a silicate mineral. Again, count the number of tetrahedra versus the number of oxygen ions. This should give you the ratio of Si to O in double-chain silicates (e.g., amphibole).

A double chain of 14 tetrahedra and 48 oxygen ions

See Appendix 3 for Exercise 2.4 answers.

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).

In mica structures, the silica tetrahedra are arranged in continuous sheets, where each tetrahedron shares three oxygen anions with adjacent tetrahedra. There is even more sharing of oxygens between adjacent tetrahedra and hence fewer cations are needed to balance the charge of the silica-tetrahedra structure in sheet silicate minerals. Bonding between sheets is relatively weak, and this accounts for the well-developed one-directional cleavage in micas (Figure 2.4.5). 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 that commonly includes magnesium. In muscovite mica, the only cations present are aluminum and potassium; hence it is a non-ferromagnesian silicate mineral.

Figure 2.4.5 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 (a.k.a. phyllosilicates), many of which exist as clay-sized fragments (i.e., less than 0.004 millimetres). These include the clay minerals kaolinite, illite, and smectite, and although they are difficult to study because of their very small size, they are extremely important components of rocks and especially of soils.

All of the sheet silicate minerals also have water molecules within their structure.

Silica tetrahedra are bonded in three-dimensional frameworks in both the feldspars and quartz. These are non-ferromagnesian minerals—they don’t contain any iron or magnesium. In addition to silica tetrahedra, feldspars include the cations aluminum, potassium, sodium, and calcium in various combinations. Quartz contains only silica tetrahedra.

The three main feldspar minerals are potassium feldspar, (a.k.a. K-feldspar or K-spar) and two types of plagioclase feldspar: albite (sodium only) and anorthite (calcium only). As is the case for iron and magnesium in olivine, there is a continuous range of compositions (solid solution series) between albite and anorthite in plagioclase. Because the calcium and sodium ions are almost identical in size (1.00 Å versus 0.99 Å) any intermediate compositions between CaAl2Si3O8 and NaAlSi3O8 can exist (Figure 2.4.6). This is a little bit surprising because, although they are very similar in size, calcium and sodium ions don’t have the same charge (Ca2+ versus Na+ ). This problem is accounted for by the corresponding substitution of Al+3  for Si+4 . Therefore, albite is NaAlSi3O8 (1 Al and 3 Si) while anorthite is CaAl2Si2O8 (2 Al and 2 Si), and plagioclase feldspars of intermediate composition have intermediate proportions of Al and Si. This is called a “coupled-substitution.”

The intermediate-composition plagioclase feldspars are oligoclase (10% to 30% Ca), andesine (30% to 50% Ca), labradorite (50% to 70% Ca), and bytownite (70% to 90% Ca). K-feldspar (KAlSi3O8) has a slightly different structure than that of plagioclase, owing to the larger size of the potassium ion (1.37 Å) and because of this large size, potassium and sodium do not readily substitute for each other, except at high temperatures. These high-temperature feldspars are likely to be found only in volcanic rocks because intrusive igneous rocks cool slowly enough to low temperatures for the feldspars to change into one of the lower-temperature forms.

Figure 2.4.6 Compositions of the feldspar minerals.

In quartz (SiO2), the silica tetrahedra are bonded in a “perfect” three-dimensional framework. Each tetrahedron is bonded to four other tetrahedra (with an oxygen shared at every corner of each tetrahedron), and as a result, the ratio of silicon to oxygen is 1:2. 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 result from the strong covalent/ionic bonds characteristic of the silica tetrahedron.

Exercise 2.5 Ferromagnesian silicates?

Silicate minerals are classified as being either ferromagnesian or non-ferromagnesian depending on whether or not they have iron (Fe) and/or magnesium (Mg) in their formula. A number of minerals and their formulas are listed below. For each one, indicate whether or not it is a ferromagnesian silicate.

Mineral Formula Ferromagnesian silicate?
olivine (Mg,Fe)2SiO4 .
pyrite FeS2 .
plagioclase feldspar CaAl2Si2O8 .
pyroxene MgSiO3 .
hematite Fe2O3 .
orthoclase feldspar KAlSi3O8 .
quartz SiO2 .
amphibole Fe7Si8O22(OH)2 .
muscovite K2Al4Si6Al2O20(OH)4 .
magnetite Fe3O4 .
biotite K2Fe4Al2Si6Al4O20(OH)4 .
dolomite (Ca,Mg)CO3 .
garnet Fe2Al2Si3O12 .
serpentine Mg3Si2O5(OH)4 .

See Appendix 3 for Exercise 2.5 answers.*Some of the formulas, especially the more complicated ones, have been simplified.

Image Descriptions

Figure 2.4.2 image description: The ionic radii of elements in angstroms and their charges.
Element Ionic Radii (in angstroms) Charge
Oxygen 1.4 −2 (Anion)
Potassium 1.37 1 (Cation)
Calcium 1.00 2 (Cation)
Sodium 0.99 1 (Cation)
Magnesium 0.72 2 (Cation)
Iron 0.63 2 (Cation)
0.49 3 (Cation)
Aluminum 0.39 3 (Cation)
Silicon 0.26 4 (Cation)
Carbon 0.15 4 (Cation)

[Return to Figure 2.4.2]


2.5 Formation of Minerals

In order for a mineral crystal to grow, the elements needed to make it must be present in the appropriate proportions, the physical and chemical conditions must be favourable, and there must be sufficient time for the atoms to become arranged.

Physical and chemical conditions include factors such as temperature, pressure, presence of water, pH, and amount of oxygen available. Time is one of the most important factors because it takes time for atoms to become ordered. If time is limited, the mineral grains will remain very small. The presence of water enhances the mobility of ions and can lead to the formation of larger crystals over shorter time periods.

Most of the minerals that make up tehe rocks around us formed through the cooling of molten rock, known as magma. At the high temperatures that exist deep within Earth, some geological materials are liquid. As magma rises up through the crust, either by volcanic eruption or by more gradual processes, it cools and minerals crystallize. If the cooling process is rapid (minutes, hours, days, or years), the components of the minerals will not have time to become ordered and only small crystals can form before the rock becomes solid. The resulting rock will be fine-grained (i.e., with crystals less than 1 mm). If the cooling is slow (from decades to millions of years), the degree of ordering will be higher and relatively large crystals will form. In some cases, the cooling will be so fast (seconds) that the texture will be glassy, which means that no crystals at all form. Volcanic glass is not composed of minerals because the magma has cooled too rapidly for crystals to grow, although over time (millions of years) the volcanic glass may crystallize into various silicate minerals.

Minerals can also form in several other ways:

Opal is a mineraloid (i.e., not an actual mineral) because although it has all of the other properties of a mineral, it does not have a specific structure. Pearl is not a mineral because it can only be produced by organic processes.

Exercise 2.7 Making crystals from solution

Light-coloured crystals
Figure 2.5.1

Place about ½ teaspoon (~2.5 cm3) of any kind of table salt into a small bowl. Add about 2 teaspoons (~10 mL) of very hot water and swirl it around for a few minutes until all or almost all of the salt has dissolved. (Be careful not to splash yourself with the hot water.)

Place the bowl in a safe place (windowsill, bookshelf), and check back every 24 hours to see what has happened.  Depending on the level of humidity in the room, you should see crystals forming within 24 hours, and all of the water should be gone, with reasonably large crystals formed, within about 3 days.  They should look a little like those shown here. In other words, they should be cubes.

Now try the same experiment again, but this time put the salt and water into a small sauce pan on the stove top at the lowest heat possible.  Within 10 to 20 minutes all of the water should be gone and you should be left with some very small salt crystals—too small to even see their shapes.  It takes time for mineral crystals to form.

Where does lithium come from?

The global demand for lithium has increased dramatically in the past decade, and will increase even more in the future as long as there is increasing demand for lithium-ion batteries in electronic devices, electric vehicles and for storage of solar- and wind-generated energy. Most of the world’s lithium supply comes from salt lakes (salars in Spanish) like the one shown below in southwestern Bolivia.

A lake with a large amount of salt deposits
Figure 2.5.2
The salty water of this and other lakes in the region has enough lithium in to make it a viable source of the metal, especially because, in the dry climate, that concentration can be increased by more evaporation.  When this water is evaporated lithium crystallizes out as the mineral lithium carbonate (Li2CO3).  For use in batteries the lithium is converted to other mineral forms—such as lithium cobalt oxide or lithium iron phosphate.

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2.6 Mineral Properties

Minerals are universal. A crystal of hematite on Mars will have the same properties as one on Earth, and the same as one on a planet orbiting another star. That’s good news for geology students who are planning interplanetary travel since we can use those properties to help us identify minerals anywhere. That doesn’t mean that it’s easy, however; identification of minerals takes a lot of practice. Some of the mineral properties that are useful for identification are as follows: colour, streak, lustre, hardness, crystal habit, cleavage/fracture, density and a few others.


For most of us, colour is one of our key ways of identifying objects. While some minerals have particularly distinctive colours that make good diagnostic properties, many do not, and for many, colour is simply unreliable. The mineral sulphur (2.6.1 left) is always a distinctive and unique yellow. Hematite, on the other hand, is an example of a mineral for which colour is not diagnostic. In some forms hematite is deep dull red, but in others it is black and shiny metallic (Figure 2.6.2). Many other minerals can have a wide range of colours (e.g., quartz, feldspar, amphibole, fluorite, and calcite). In most cases, the variations in colours are a result of varying proportions of trace elements within the mineral. In the case of quartz, for example, yellow quartz (citrine) has trace amounts of ferric iron (Fe3+), rose quartz has trace amounts of manganese, purple quartz (amethyst) has trace amounts of iron, and milky quartz, which is very common, has millions of fluid inclusions (tiny cavities, each filled with water).

The piece of sulphur is bright yellow. One piece of Hematite is a redish brown and the other is a silvery metalic colour.
Figure 2.6.1 Examples of the colours of the minerals sulphur and hematite.


In the context of minerals, “colour” is what you see when light reflects off the surface of the sample. One reason that colour can be so variable is that the type of surface is variable. It may be a crystal face or a fracture surface or a cleavage plane, and the crystals may be large or small depending on the nature of the rock. If we grind a small amount of the sample to a powder we get a much better indication of its actual colour. This can easily be done by scraping a corner of the sample across a streak plate (a piece of unglazed porcelain) to make a streak. The result is that some of the mineral gets ground to a powder and we can get a better impression of its “true” colour (Figure 2.6.2).

Figure 2.6.2 The streak colours of specular (metallic) hematite (left) and earthy hematite (right). Hematite leaves a distinctive reddish-brown streak whether the sample is metallic or earthy.


Lustre is the way light reflects off the surface of a mineral, and the degree to which it penetrates into the interior. The key distinction is between metallic and non-metallic lustres. Light does not pass through metals, and that is the main reason they look “metallic.” Even a thin sheet of metal—such as aluminum foil—will not allow light to pass through it. Many non-metallic minerals may look as if light will not pass through them, but if you take a closer look at a thin edge of the mineral you can see that it does. If a non-metallic mineral has a shiny, reflective surface, then it is called “glassy.” If it is dull and non-reflective, it is “earthy.” Other types of non-metallic lustres are “silky,” “pearly,” and “resinous.” Lustre is a good diagnostic property since most minerals will always appear either metallic or non-metallic. There are a few exceptions to this (e.g., hematite in Figure 2.6.1).


One of the most important diagnostic properties of a mineral is its hardness. In 1812 German mineralogist Friedrich Mohs came up with a list of 10 reasonably common minerals that had a wide range of hardnesses. These minerals are shown in Figure 2.6.3, with the Mohs scale of hardness along the bottom axis. In fact, while each mineral on the list is harder than the one before it, the relative measured hardnesses (vertical axis) are not linear. For example apatite is about three times harder than fluorite and diamond is three times harder than corundum. Some commonly available reference materials are also shown on this diagram, including a typical fingernail (2.5), a piece of copper wire (3.5), a knife blade or a piece of window glass (5.5), a hardened steel file (6.5), and a porcelain streak plate (7). These are tools that a geologist can use to measure the hardness of unknown minerals. For example, if you have a mineral that you can’t scratch with your fingernail, but you can scratch with a copper wire, then its hardness is between 2.5 and 3.5. And of course the minerals themselves can be used to test other minerals.

Mohs hardness versus measured hardness. Image description available.
Figure 2.6.3 Minerals and reference materials in the Mohs scale of hardness. The “measured hardness” values are Vickers Hardness numbers. [Image Description]

Crystal Habit

When minerals form within rocks, there is a possibility that they will form in distinctive crystal shapes if they formed slowly and if they are not crowded out by other pre-existing minerals. Every mineral has one or more distinctive crystal habits, but it is not that common, in ordinary rocks, for the shapes to be obvious. Quartz, for example, will form six-sided prisms with pointed ends (Figure 2.6.4a), but this typically happens only when it crystallizes from a hot water solution within a cavity in an existing rock. Pyrite can form cubic crystals (Figure 2.6.4b), but can also form crystals with 12 faces, known as dodecahedra (“dodeca” means 12). The mineral garnet also forms dodecahedral crystals (Figure 2.6.4c).

The quartz crystals poke out in multiple directions. They look like glass
Figure 2.6.4a Hexagonal prisms of quartz.
Cubes of pyrite are opaque and silvery in colour
Figure 2.6.4b Cubic crystals of pyrite.
A dark red garnet embedded in a rock
Figure 2.6.4c A dodecahedral crystal of garnet.

Because well-formed crystals are rare in ordinary rocks, habit isn’t as useful a diagnostic feature as one might think. However, there are several minerals for which it is important. One is garnet, which is common in some metamorphic rocks and typically displays the dodecahedral shape. Another is amphibole, which forms long thin crystals, and is common in igneous rocks like granite (Figure 1.4.2).

Mineral habit is often related to the regular arrangement of the molecules that make up the mineral. Some of the terms that are used to describe habit include bladed, botryoidal (grape-like), dendritic (branched), drusy (an encrustation of minerals), equant (similar in all dimensions), fibrous, platy, prismatic (long and thin), and stubby.

Cleavage and Fracture

Crystal habit is a reflection of how a mineral grows, while cleavage and fracture describe how it breaks. Cleavage and fracture  are the most important diagnostic features of many minerals, and often the most difficult to understand and identify. Cleavage is what we see when a mineral breaks along a specific plane or planes, while fracture is an irregular break. Some minerals tend to cleave along planes at various fixed orientations, some do not cleave at all (they only fracture). Minerals that have cleavage can also fracture along surfaces that are not parallel to their cleavage planes.

As we’ve already discussed, the way that minerals break is determined by their atomic arrangement and specifically by the orientation of weaknesses within the lattice. Graphite and the micas, for example, have cleavage planes parallel to their sheets (Figures 2.2.5 and 2.4.5), and halite has three cleavage planes parallel to the lattice directions (Figure 2.2.6).

Quartz has no cleavage because it has equally strong Si–O bonds in all directions, and feldspar has two cleavages at 90° to each other (Figure 2.6.5).

A piece of potassium feldspar that shows a fracture surface, which is rough, and two cleavage planes, which are smooth
Figure 2.6.5 Cleavage and fracture in potassium feldspar

One of the main difficulties with recognizing and describing cleavage is that it is visible only in individual crystals. Most rocks have small crystals and it’s very difficult to see the cleavage within those crystals. Geology students have to work hard to understand and recognize cleavage, but it’s worth the effort since it is a reliable diagnostic property for most minerals.

One last thing: it is important to recognize the difference between cleavage planes and crystal surfaces.  As already noted, crystal surfaces are related to how a mineral grows while cleavage planes are related to how it breaks. In most minerals cleavage planes and crystal surfaces do not align with one-another.  An exception is halite, which grows in cubic crystals and has cleavage along those same planes (Figure 1.4.1 and 2.2.6).  But this doesn’t hold for most minerals. Quartz has crystal surfaces but no cleavage at all.  Fluorite forms cubic crystals like those of halite, but it cleaves along planes that differ in orientation from the crystal surfaces.  This is illustrated in Figure 2.6.6.

Figure 2.6.6 Crystal faces and cleavage planes in the mineral fluorite. The top-left photo shows a natural crystal of fluorite.  It has crystal surfaces  but you can see some future cleavage planes inside the crystal. The top-right photo shows what you can create if you take a crystal like the one on the left and carefully break it along its cleavage planes.


Density is a measure of the mass of a mineral per unit volume, and it is a useful diagnostic tool in some cases. Most common minerals, such as quartz, feldspar, calcite, amphibole, and mica, have what we call “average density” (2.6 to 3.0 grams per cubic centimetre (g/cm3)), and it would be difficult to tell them apart on the basis of their density. On the other hand, many of the metallic minerals, such as pyrite, hematite, and magnetite, have densities over 5 g/cm3. They can easily be distinguished from the lighter minerals on the basis of density, but not necessarily from each other. A limitation of using density as a diagnostic tool is that one cannot assess it in minerals that are a small part of a rock that is mostly made up of other minerals.

Other Properties

Several other properties are also useful for identification of some minerals. For example, calcite is soluble in dilute acid and will give off bubbles of carbon dioxide. Magnetite is magnetic, so will affect a magnet. A few other minerals are weakly magnetic.

Image Descriptions

Figure 2.6.3 image description
Talc Gypsum Calcite Fluorine Apatite Feldspar Quartz Topaz Corundum Diamond
Measured Hardness 50 60 105 200 659 700 1100 1648 2085 7000
Mohs Hardness 1 2 3 4 5 6 7 8 9 10

[Return to Figure 2.6.3]

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

Section Summary
2.1 Electrons, Protons, Neutrons, and Atoms An atom is made up of protons and neutrons in the nucleus and electrons arranged in energy shells around the nucleus. The first shell holds two electrons, and outer shells hold more, but atoms strive to have eight electrons in their outermost shell (or two for H and He). They either gain or lose electrons (or share) to achieve this, and in so doing become either cations (if they lose electrons) or anions (if they gain them).
2.2 Bonding and Lattices The main types of bonding in minerals are ionic bonding (electrons transferred) and covalent bonding (electrons shared). Some minerals have metallic bonding or other forms of weak bonding. Minerals form in specific three-dimensional lattices, and the nature of the lattices and the type of bonding within them have important implications for mineral properties.
2.3 Mineral Groups Minerals are grouped according to the anion part of their formula, with some common types being oxides, sulphides, sulphates, halides, carbonates, phosphates, silicates, and native minerals.
2.4 Silicate Minerals Silicate minerals are, by far, the most important minerals in Earth’s crust. They all include silica tetrahedra (four oxygens surrounding a single silicon atom) arranged in different structures (chains, sheets, etc.). Some silicate minerals include iron or magnesium and are called ferromagnesian silicates.
2.5 Formation of Minerals Most minerals in the crust form from the cooling and crystallization of magma. Some form from hot water solutions, during metamorphism or weathering, or through organic processes.
2.6 Mineral Properties Some of the important properties for mineral identification include hardness, cleavage/fracture, density, lustre, colour, and streak colour.  It’s critical to be able to recognize these properties in order to be able to identify minerals.

Questions for Review

Answers to Review Questions can be found in Appendix 2.

  1. What is the electrical charge on a proton? A neutron? An electron? What are their relative masses?
  2. Explain how the need for an atom’s outer shell to be filled with electrons contributes to bonding.
  3. Why are helium and neon non-reactive?
  4. What is the difference in the role of electrons in an ionic bond compared to a covalent bond?
  5. What is the electrical charge on an anion? A cation?
  6. What chemical feature is used in the classification of minerals into groups?
  7. Name the mineral group for the following minerals:
    • calcite
    • gypsum
    • hematite
    • quartz
    • biotite
    • galena
    • graphite
    • fluorite
    • pyrite
    • orthoclase
    • magnetite
    • olivine
  8. What is the net charge on an unbonded silica tetrahedron?
  9. What allows magnesium to substitute freely for iron in olivine?
  10. How are the silica tetrahedra structured differently in pyroxene and amphibole?
  11. Why is biotite called a ferromagnesian mineral, while muscovite is not?
  12. What are the names and compositions of the two end-members of the plagioclase series?
  13. Why does quartz have no additional cations (other than Si+4)?
  14. Why is colour not necessarily a useful guide to mineral identification?
  15. You have an unknown mineral that can scratch glass but cannot scratch a porcelain streak plate. What is its approximate hardness?


Chapter 3 Intrusive Igneous Rocks

Learning Objectives

After carefully reading this chapter, completing the exercises within it, and answering the questions at the end, you should be able to:

  • Describe the rock cycle and the types of processes that lead to the formation of igneous, sedimentary, and metamorphic rocks, and explain why there is an active rock cycle on Earth.
  • Explain the concept of partial melting and describe the geological processes that lead to melting.
  • Describe, in general terms, the range of chemical compositions of magmas.
  • Discuss the processes that take place during the cooling and crystallization of magma, and the typical order of crystallization according to the Bowen reaction series.
  • Explain how magma composition can be changed by fractional crystallization and partial melting of the surrounding rocks.
  • Apply the criteria for igneous rock classification based on mineral proportions.
  • Describe the origins of phaneritic, porphyritic, and pegmatitic rock textures.
  • Identify plutons on the basis of their morphology and their relationships to the surrounding rocks.
  • Explain the origin of a chilled margin.
Figure 3.0.1 A fine-grained mafic dyke (dark green) intruded into a felsic dyke (pink) and into coarse diorite (grey), Quadra Island, B.C. All of these rocks are composed of more than one type of mineral. The mineral components are clearly visible in the diorite, but not in the other two rock types.

A rock is a consolidated mixture of minerals. By consolidated, we mean hard and strong; real rocks don’t fall apart in your hands! A mixture of minerals implies the presence of more than one mineral grain, but not necessarily more than one type of mineral (Figure 3.0.1). A rock can be composed of only one type of mineral (e.g., limestone is commonly made up of only calcite), but most rocks are composed of several different minerals. A rock can also include non-minerals, such as fossils or the organic matter within a coal bed or in some types of mudstone.

Rocks are grouped into three main categories based on how they form:

  1. Igneous: formed from the cooling and crystallization of magma (molten rock)
  2. Sedimentary: formed when weathered fragments of other rocks are buried, compressed, and cemented together, or when minerals precipitate directly from solution
  3. Metamorphic: formed by alteration (due to heat, pressure, and/or chemical action) of a pre-existing igneous or sedimentary rock

Media Attributions


3.1 The Rock Cycle

The rock components of the crust are slowly but constantly being changed from one form to another and the processes involved are summarized in the rock cycle (Figure 3.1.1). The rock cycle is driven by two forces: (1) Earth’s internal heat engine, which moves material around in the core and the mantle and leads to slow but significant changes within the crust, and (2) the hydrological cycle, which is the movement of water, ice, and air at the surface, and is powered by the sun.

The rock cycle is still active on Earth because our core is hot enough to keep the mantle moving, our atmosphere is relatively thick, and we have liquid water. On some other planets or their satellites, such as the Moon, the rock cycle is virtually dead because the core is no longer hot enough to drive mantle convection and there is no atmosphere or liquid water.

The rock cycle. Image description available
Figure 3.1.1 A schematic view of the rock cycle. [Image description]

In describing the rock cycle, we can start anywhere we like, although it’s convenient to start with magma. As we’ll see in more detail below, magma is rock that is hot to the point of being entirely molten, with a temperature of between about 800° and 1300°C, depending on the composition and the pressure.

Red hto magma runs down rocks
Figure 3.1.2 Magma forming pahoehoe basalt at Kilauea Volcano, Hawaii.

Magma can either cool slowly within the crust (over centuries to millions of years)—forming intrusive igneous rock, or erupt onto the surface and cool quickly (within seconds to years)—forming extrusive igneous rock (volcanic rock) (Figure 3.1.2). Intrusive igneous rock typically crystallizes at depths of hundreds of metres to tens of kilometres below the surface. To change its position in the rock cycle, intrusive igneous rock has to be uplifted and then exposed by the erosion of the overlying rocks.

Through the various plate-tectonics-related processes of mountain building, all types of rocks are uplifted and exposed at the surface. Once exposed, they are weathered, both physically (by mechanical breaking of the rock) and chemically (by weathering of the minerals), and the weathering products—mostly small rock and mineral fragments—are eroded, transported, and then deposited as sediments. Transportation and deposition occur through the action of glaciers, streams, waves, wind, and other agents, and sediments are deposited in rivers, lakes, deserts, and the ocean.

Exercise 3.1 Rock around the rock-cycle clock

Referring to the rock cycle (Figure 3.1.1), list the steps that are necessary to cycle some geological material starting with a sedimentary rock, which then gets converted into a metamorphic rock, and eventually a new sedimentary rock.

A conservative estimate is that each of these steps would take approximately 20 million years (some may be less, others would be more, and some could be much more). How long might it take for this entire process to be completed?

See Appendix 3 for Exercise 3.1 Answers.

Marine sandstone sticks out over a wall of marine mudstone
Figure 3.1.3 Cretaceous-aged marine sandstone overlying marine mudstone, Gabriola Island, B.C.

Unless they are re-eroded and moved along, sediments will eventually be buried by more sediments. At depths of hundreds of metres or more, they become compressed and cemented into sedimentary rock (See Figure 3.1.3 for example). Again through various means, largely resulting from plate-tectonic forces, different kinds of rocks are either uplifted, to be re-eroded, or buried deeper within the crust where they are heated up, squeezed, and changed into metamorphic rock (Figure 3.1.4)

Dark stone with white streaks that looks like it has been bent and folded
Figure 3.1.4 Metamorphosed and folded Triassic-aged limestone, Quadra Island, B.C.

Image Descriptions

Figure 3.1.1 image description: The rock cycle takes place both above and below the Earth’s surface. The rock deepest beneath the earth’s surface, and under extreme heat and pressure, is metamorphic rock. This metamorphic rock can melt and become magma. When magma cools below the earth’s surface, it becomes “intrusive igneous rock.” If magma cools above the earth’s surface, it is “extrusive igneous rock” and becomes part of the outcrop. The outcrop is subject to weathering and erosion, and can be moved and redeposited around the earth by forces such as water and wind. As the outcrop is eroded, it becomes sediment which can be buried, compacted, and cemented beneath the Earth’s surface to become sedimentary rock. As sedimentary rock gets buried deeper and comes under increased heat and pressure, it returns to its original state as metamorphic rock. Rocks in the rock cycle do not always make a complete loop. It is possible for sedimentary rock to be uplifted back above the Earth’s surface and for intrusive and extrusive igneous rock to be reburied and become metamorphic rock. [Return to Figure 3.1.1]

Images Attributions


3.2 Magma and Magma Formation

Magmas can vary widely in composition, but in general they are made up of only eight elements; in order of importance: oxygen, silicon, aluminum, iron, calcium, sodium, magnesium, and potassium (Figure 3.2.1). Oxygen, the most abundant element in magma, comprises a little less than half the total, followed by silicon at just over one-quarter. The remaining elements make up the other one-quarter. Magmas derived from crustal material are dominated by oxygen, silicon, aluminum, sodium, and potassium.

The composition of magma depends on the rock it was formed from (by melting), and the conditions of that melting. Magmas derived from the mantle have higher levels of iron, magnesium, and calcium, but they are still likely to be dominated by oxygen and silicon. All magmas have varying proportions of elements such as hydrogen, carbon, and sulphur, which are converted into gases like water vapour, carbon dioxide, and hydrogen sulphide as the magma cools.

Elements in the Earth's crust. Long decriptions available
Figure 3.2.1 Average elemental proportions in Earth’s crust, which is close to the average composition of magmas within the crust. [Image Description]

Virtually all of the igneous rocks that we see on Earth are derived from magmas that formed from partial melting of existing rock, either in the upper mantle or the crust. Partial melting is what happens when only some parts of a rock melt; it takes place because rocks are not pure materials. Most rocks are made up of several minerals, each of which has a different melting temperature. The wax in a candle is a pure material. If you put some wax into a warm oven (50°C will do as the melting temperature of most wax is about 40°C) and leave it there for a while, it will soon start to melt. That’s complete melting, not partial melting. If instead you took a mixture of wax, plastic, aluminum, and glass and put it into the same warm oven, the wax would soon start to melt, but the plastic, aluminum, and glass would not melt (Figure 3.2.2a). That’s partial melting and the result would be solid plastic, aluminum, and glass surrounded by liquid wax (Figure 3.2.2b). If we heat the oven up to around 120°C, the plastic would melt too and mix with the liquid wax, but the aluminum and glass would remain solid (Figure 3.2.2c). Again this is partial melting. If we separated the wax/plastic “magma” from the other components and let it cool, it would eventually harden. As you can see from Figure 3.2.2d, the liquid wax and plastic have mixed, and on cooling, have formed what looks like a single solid substance. It is most likely that this is a very fine-grained mixture of solid wax and solid plastic, but it could also be some other substance that has formed from the combination of the two.

Figure 3.2.2 Partial melting of “pretend rock”: (a) the original components of white candle wax, black plastic pipe, green beach glass, and aluminum wire, (b) after heating to 50˚C for 30 minutes only the wax has melted, (c) after heating to 120˚C for 60 minutes much of the plastic has melted and the two liquids have mixed, (d) the liquid has been separated from the solids and allowed to cool to make a “pretend rock” with a different overall composition.

In this example, we partially melted some pretend rock to create some pretend magma. We then separated the magma from the source and allowed it to cool to make a new pretend rock with a composition quite different from the original material (it lacks glass and aluminum).

Of course partial melting in the real world isn’t exactly the same as in our pretend-rock example. The main differences are that rocks are much more complex than the four-component system we used, and the mineral components of most rocks have more similar melting temperatures, so two or more minerals are likely to melt at the same time to varying degrees. Another important difference is that when rocks melt, the process takes thousands to millions of years, not the 90 minutes it took in the pretend-rock example.

Contrary to what one might expect, and contrary to what we did to make our pretend rock, most partial melting of real rock does not involve heating the rock up. The two main mechanisms through which rocks melt are decompression melting and flux melting. Decompression melting takes place within Earth when a body of rock is held at approximately the same temperature but the pressure is reduced. This happens because the rock is being moved toward the surface, either at a mantle plume (a.k.a., hot spot), or in the upwelling part of a mantle convection cell.Mantle plumes are described in Chapter 4 and mantle convection in Chapter 9. The mechanism of decompression melting is shown in Figure 3.2.3a. If a rock that is hot enough to be close to its melting point is moved toward the surface, the pressure is reduced, and the rock can pass to the liquid side of its melting curve. At this point, partial melting starts to take place. The process of flux melting is shown in Figure 3.2.3b. If a rock is close to its melting point and some water (a flux that promotes melting) is added to the rock, the melting temperature is reduced (solid line versus dotted line), and partial melting starts.

Graphs describing the melting curve for dry and wet mantle rock. Long description available
Figure 3.2.3 Mechanisms for (a) decompression melting (the rock is moved toward the surface) and (b) flux melting (water is added to the rock) and the melting curve is displaced. [Image Description]

The partial melting of rock happens in a wide range of situations, most of which are related to plate tectonics. The more important of these are shown in Figure 3.2.3. At both mantle plumes and in the upward parts of convection systems, rock is being moved toward the surface, the pressure is dropping, and at some point, the rock crosses to the liquid side of its melting curve. At subduction zones, water from the wet, subducting oceanic crust is transferred into the overlying hot mantle. This provides the flux needed to lower the melting temperature. In both of these cases, only partial melting takes place—typically only about 10% of the rock melts—and it is always the most silica-rich components of the rock that melt, creating a magma that is more silica-rich than the rock from which it is derived. (By analogy, the melt from our pretend rock is richer in wax and plastic than the “rock” from which it was derived.) The magma produced, being less dense than the surrounding rock, moves up through the mantle, and eventually into the crust.

Figure 3.2.4 Common sites of magma formation in the upper mantle. The black circles are regions of partial melting. The blue arrows represent water being transferred from the subducting plates into the overlying mantle.

As it moves toward the surface, and especially when it moves from the mantle into the lower crust, the hot magma interacts with the surrounding rock. This typically leads to partial melting of the surrounding rock because most such magmas are hotter than the melting temperature of crustal rock. (In this case, melting is caused by an increase in temperature.) Again, the more silica-rich parts of the surrounding rock are preferentially melted, and this contributes to an increase in the silica content of the magma.

At very high temperatures (over 1300°C), most magma is entirely liquid because there is too much energy for the atoms to bond together. As the temperature drops, usually because the magma is slowly moving upward, things start to change. Silicon and oxygen combine to form silica tetrahedra, and then, as cooling continues, the tetrahedra start to link together to make chains (polymerize). These silica chains have the important effect of making the magma more viscous (less runny), and as we’ll see in Chapter 4, magma viscosity has significant implications for volcanic eruptions. As the magma continues to cool, crystals start to form.

Exercise 3.2 Making magma viscous

This is an experiment that you can do at home to help you understand the properties of magma. It will only take about 15 minutes, and all you need is half a cup of water and a few tablespoons of flour.

If you’ve ever made gravy, white sauce, or roux, you’ll know how this works.

Place about 1/2 cup (125 mL) of water in a saucepan over medium heat. Add 2 teaspoons (10 mL) of white flour (this represents silica) and stir while the mixture comes close to boiling. It should thicken like gravy because the gluten in the flour becomes polymerized into chains during this process.

Now you’re going to add more “silica” to see how this changes the viscosity of your magma. Take another 4 teaspoons (20 mL) of flour and mix it thoroughly with about 4 teaspoons (20 mL) of water in a cup and then add all of that mixture to the rest of the water and flour in the saucepan. Stir while bringing it back up to nearly boiling temperature, and then allow it to cool. This mixture should slowly become much thicker — something like porridge — because there is more gluten and more chains have been formed (see the photo).

Thick, white goop in a pot.
Figure 3.2.5 Flour-and-water magma experiment.

This is analogous to magma, of course. As we’ll see below, magmas have quite variable contents of silica and therefore have widely varying viscosities (“thicknesses”) during cooling.

See Appendix 3 for Exercise 3.2 answers.

Image Descriptions

Figure 3.2.1 image description: The average elemental proportions in the Earth’s crust from the largest amount to the smallest amount. Oxygen (46.6%), Silicon (27.7%), Aluminum (8.1%), Iron (5.0%), Calcium (3.6%), Sodium (2.8%), Potassium (2.6%), Magnesium (2.1%), Others (1.5%). [Return to Figure 3.2.1]

Figure 3.2.3a image description: Dry mantle rock is predominately solid. However, its melting point is dependent on the temperature and pressure the rock is under. The higher the pressure (meaning the farther the rock is from the Earth’s surface), the more likely dry mantle rock is going to be solid. Dry mantle rock under extreme pressure requires a much higher temperature to melt than dry mantle rock under less pressure. As pressure drops (meaning as the rock rises towards the Earth’s surface), the required temperature to melt the mantle rock drops as well.

Figure 3.2.3b image description: In comparison to dry mantle rock, wet mantle rock under the same amount of pressure (at the same distance from the earth’s surface) requires a lower temperature to melt. When liquid is added to dry mantel rock at a pressure and temperature point in which wet mantle rock would be melted, flux melting occurs. [Return to Figure 3.2.3]

Media Attributions


3.3 Crystallization of Magma

The minerals that make up igneous rocks crystallize at a range of different temperatures. This explains why a cooling magma can have some crystals within it and yet remain predominantly liquid. The sequence in which minerals crystallize from a magma is known as the Bowen reaction series (Figure 3.3.1 and Figure 3.3.3).

Of the common silicate minerals, olivine normally crystallizes first, at between 1200° and 1300°C. As the temperature drops, and assuming that some silica remains in the magma, the olivine crystals will react (combine) with some of the silica in the magma to form pyroxene. As long as there is silica remaining and the rate of cooling is slow, this process continues down the discontinuous branch: olivine to pyroxene, pyroxene to amphibole, and (under the right conditions) amphibole to biotite.

At about the point where pyroxene begins to crystallize, plagioclase feldspar also begins to crystallize. At that temperature, the plagioclase is calcium-rich (anorthite) (see Figure 2.6.1). As the temperature drops, and providing that there is sodium left in the magma, the plagioclase that forms is a more sodium-rich variety.

Figure 3.3.1 The Bowen reaction series describes the process of magma crystallization.
A rectangular crystal of feldspar viewed in a microscope. It is concentrically zoned from black around the outside to light grey in the centre.
Figure 3.3.2 A zoned plagioclase crystal. The central part is calcium-rich and the darker outside part is sodium-rich.

In some cases, individual plagioclase crystals can be zoned from calcium-rich in the centre to more sodium-rich around the outside. This occurs when calcium-rich early-forming plagioclase crystals become coated with progressively more sodium-rich plagioclase as the magma cools. Figure 3.3.2 shows a zoned plagioclase under a microscope.

Finally, if the magma is quite silica-rich to begin with, there will still be some left at around 750° to 800°C, and from this last magma, potassium feldspar, quartz, and maybe muscovite mica will form.

Who was Bowen, and what is a reaction series?

Figure 3.3.3

Norman Levi Bowen, born in Kingston Ontario, studied geology at Queen’s University and then at MIT in Boston. In 1912, Norman Levi Bowenhe joined the Carnegie Institution in Washington, D.C., where he carried out groundbreaking experimental research into the processes of cooling magmas. Working mostly with basaltic magmas, he determined the order of crystallization of minerals as the temperature drops. The method, in brief, was to melt the rock to a magma in a specially-made kiln, allow it to cool slowly to a specific temperature (allowing some minerals to form), and then quench it (cool it quickly) so that no new minerals form (only glass). The results were studied under the microscope and by chemical analysis. This was done over and over, each time allowing the magma to cool to a lower temperature before quenching.

The Bowen reaction series is one of the results of his work, and even a century later, it is an important basis for our understanding of igneous rocks. The word reaction is critical. In the discontinuous branch, olivine is typically the first mineral to form (at just below 1300°C). As the temperature continues to drop, olivine becomes unstable while pyroxene becomes stable. The early-forming olivine crystals react with silica in the remaining liquid magma and are converted into pyroxene, something like this:

Mg2SiO4 + SiO2 (olivine) becomes 2MgSiO3 (proxene)

This continues down the chain, as long as there is still silica left in the liquid.

The composition of the original magma is critical to magma crystallization because it determines how far the reaction process can continue before all of the silica is used up. The compositions of typical mafic, intermediate, and felsic magmas are shown in Figure 3.3.4. Note that, unlike Figure 3.2.1, these compositions are expressed in terms of “oxides” (e.g., Al2O3 rather than just Al). There are two reasons for this: one is that in the early analytical procedures, the results were always expressed that way, and the other is that all of these elements combine readily with oxygen to form oxides.

Figure 3.3.4 The chemical compositions of typical mafic, intermediate, and felsic magmas and the types of rocks that form from them.

Mafic magmas have 45% to 55% SiO2, about 25% total of FeO and MgO plus CaO, and about 5% Na2O + K2O. Felsic magmas, on the other hand, have much more SiO2 (65% to 75%) and Na2O + K2O (around 10%) and much less FeO and MgO plus CaO (about 5%).

Exercise 3.3 Determining rock types based on magma composition

The proportions of the main chemical components of felsic, intermediate, and mafic magmas are listed in the table below. (The values are similar to those shown in Figure 3.3.4.)

Table 3.1 Proportions of the main chemical components in felsic, intermediate, and mafic magma
[Skip Table]
Oxide Felsic Magma Intermediate Magma Mafic Magma
SiO2 65% to 75% 55% to 65% 45% to 55%
Al2O3 12% to 16% 14% to 18% 14% to 18%
FeO 2% to 4% 4% to 8% 8% to 12%
CaO 1% to 4% 4% to 7% 7% to 11%
MgO 0% to 3% 2% to 6% 5% to 9%
Na2O 2% to 6% 3% to 7% 1% to 3%
K2O 3% to 5% 2% to 4% 0.5% to 3%

Chemical data for four rock samples are shown in the following table. Compare these with those in the table above to determine whether each of these samples is felsic, intermediate, or mafic.

Table 3.2 Chemical Data for Four Unidentified Rock Samples
 [Skip Table]
Rock Sample SiO2 Al2O3 FeO CaO MgO Na2O K2O What type of magma is it?
Rock 1 55% 17% 5% 6% 3% 4% 3%
Rock 2 74% 14% 3% 3% 0.5% 5% 4%
Rock 3 47% 14% 8% 10% 8% 1% 2%
Rock 4 65% 14% 4% 5% 4% 3% 3%

See Appendix 3 for Exercise 3.3 answers.

As a mafic magma starts to cool, some of the silica combines with iron and magnesium to make olivine. As it cools further, much of the remaining silica goes into calcium-rich plagioclase, and any silica left may be used to convert some of the olivine to pyroxene. Soon after that, all of the magma is used up and no further changes takes place. The minerals present will be olivine, pyroxene, and calcium-rich plagioclase. If the magma cools slowly underground, the product will be gabbro; if it cools quickly at the surface, the product will be basalt (Figure 3.3.5).

Felsic magmas tend to be cooler than mafic magmas when crystallization begins (because they don’t have to be as hot to remain liquid), and so they may start out crystallizing pyroxene (not olivine) and plagioclase. As cooling continues, the various reactions on the discontinuous branch will proceed because silica is abundant, the plagioclase will become increasingly sodium-rich, and eventually potassium feldspar and quartz will form. Commonly even very felsic rocks will not have biotite or muscovite because they may not have enough aluminum or enough hydrogen to make the OH complexes that are necessary for mica minerals. Typical felsic rocks are granite and rhyolite (Figure 3.3.5).

The cooling behaviour of intermediate magmas lie somewhere between those of mafic and felsic magmas. Typical mafic rocks are gabbro (intrusive) and basalt (extrusive). Typical intermediate rocks are diorite and andesite. Typical felsic rocks are granite and rhyolite (Figure 3.3.5).

Figure 3.3.5 Examples of the igneous rocks that form from mafic, intermediate, and felsic magmas.

A number of processes that take place within a magma chamber can affect the types of rocks produced in the end. If the magma has a low viscosity (i.e., it’s runny)—which is likely if it is mafic—the crystals that form early, such as olivine (Figure 3.3.6a), may slowly settle toward the bottom of the magma chamber (Figure 3.3.6b). The means that the overall composition of the magma near the top of the magma chamber will become more felsic, as it is losing some iron- and magnesium-rich components. This process is known as fractional crystallization. The crystals that settle might either form an olivine-rich layer near the bottom of the magma chamber, or they might remelt because the lower part is likely to be hotter than the upper part (remember, from Chapter 1, that temperatures increase steadily with depth in Earth because of the geothermal gradient). If any melting takes place, crystal settling will make the magma at the bottom of the chamber more mafic than it was to begin with (Figure 3.3.6c).

Figure 3.3.6 An example of crystal settling and the formation of a zoned magma chamber.

If crystal settling does not take place, because the magma is too viscous, then the process of cooling will continue as predicted by the Bowen reaction series. In some cases, however, partially cooled but still liquid magma, with crystals in it, will either move farther up into a cooler part of the crust, or all the way to the surface during a volcanic eruption. In either of these situations, the magma that has moved toward the surface is likely to cool much faster than it did within the magma chamber, and the rest of the rock will have a finer crystalline texture. An igneous rock with large crystals embedded in a matrix of much finer crystals is indicative of a two-stage cooling process, and the texture is porphyritic (Figure 3.3.7).  For the rock to be called “porphyritic” there has to be a significant difference in crystal size, where the larger crystals are at least 10 times larger than the average size of the smaller crystals.

Figure 3.3.7 Porphyritic textures, left: 1.3 cm long amphibole crystals in an intrusive igneous rock in which most of the crystals are less than 1 mm, right: 1 to 2 mm long feldspar crystals and 1 mm long amphibole crystals in a volcanic rock where most of the crystals are less than 0.1 mm.

Exercise 3.4 Porphyritic minerals

As a magma cools below 1300°C, minerals start to crystallize within it. If that magma is then involved in a volcanic eruption, the rest of the liquid will cool quickly to form a porphyritic texture. The rock will have some relatively large crystals (phenocrysts) of the minerals that crystallized early, and the rest will be very fine grained or even glassy. Using Figure 3.3.8, predict what phenocrysts might be present where the magma cooled as far as line a in one case, and line b in another.

Figure 3.3.8 Bowen reaction series. Line a – at high temperature – intersects olivine, Line b – at a lower temperature – intersects pryroxene and amphibole on the left, and plagioclase feldspar on the right

See Appendix 3 for Exercise 3.4 answers.

Media Attributions


3.4 Classification of Igneous Rock

As has already been described, igneous rocks are classified into four categories: felsic, intermediate, mafic, and ultramafic, based on either their chemistry or their mineral composition. The diagram in Figure 3.4.1 can be used to help classify igneous rocks by their mineral composition. An important feature to note on this diagram is the red line separating the non-ferromagnesian silicates in the lower left (K-feldspar, quartz, and plagioclase feldspar) from the ferromagnesian silicates in the upper right (biotite, amphibole, pyroxene, and olivine). In classifying intrusive igneous rocks, the first thing to consider is the percentage of ferromagnesian silicates. In most igneous rocks the ferromagnesian silicate minerals are clearly darker than the others, but it is still quite difficult to estimate the proportions of minerals in a rock.

Based on the position of the red line in Figure 3.4.1, it is evident that felsic rocks can have between 1% and 20% ferromagnesian silicates (the red line intersects the left side of the felsic zone 1% of the distance from the top of the diagram, and it intersects the right side of the felsic zone 20% of the distance from the top). Intermediate rocks have between 20% and 50% ferromagnesian silicates, and mafic rocks have 50% to 100% ferromagnesian silicates. To be more specific, felsic rocks typically have biotite and/or amphibole; intermediate rocks have amphibole and, in some cases, pyroxene; and mafic rocks have pyroxene and, in some cases, olivine.

Figure 3.4.1 A simplified classification diagram for igneous rocks based on their mineral compositions. [Image Description]

If we focus on the non-ferromagnesian silicates, it is evident that felsic rocks can have from 0% to 35% K-feldspar, from 25% to 35% quartz (the vertical thickness of the quartz field varies from 25% to 35%), and from 25% to 50% plagioclase (and that plagioclase will be sodium-rich, or albitic). Intermediate rocks can have up to 25% quartz and 50% to 75% plagioclase. Mafic rocks only have plagioclase (up to 50%), and that plagioclase will be calcium-rich, or anorthitic.

Exercise 3.5 Mineral proportions in igneous rocks

Figure 3.4.2

The dashed blue lines (labelled a, b, c, d)  in Figure 3.4.2 represent four igneous rocks. Complete the table by estimating the mineral proportions (percent) of the four rocks (to the nearest 10%).

Hint: Rocks b and d are the easiest; start with those.

 Rock Biotite/amphibole Pyroxene Olivine Plagioclase Quartz K-feldspar

See Appendix 3 for Exercise 3.5 answers.

Figure 3.4.3 provides a diagrammatic representation of the proportions of dark minerals in light-coloured rocks. You can use that when trying to estimate the ferromagnesian mineral content of actual rocks, and you can get some practice doing that by completing Exercise 3.6.  Be warned!  Geology students almost universally over-estimate the proportion of dark minerals.

Figure 3.4.3 A guide to estimating the proportions of dark minerals in light-coloured rocks.

Exercise 3.6 Proportions of ferromagnesian silicates

The four igneous rocks shown below have differing proportions of ferromagnesian silicates. Estimate those proportions using the diagrams in Figure 3.4.3, and then use Figure 3.4.1 to determine the likely rock name for each one.

"" "" "" ""
___% ___% ___% ___%
__________ __________ __________  __________

See Appendix 3 for Exercise 3.6 answers.

Igneous rocks are also classified according to their textures. The textures of volcanic rocks will be discussed in Chapter 4, so here we’ll only look at the different textures of intrusive igneous rocks. Almost all intrusive igneous rocks have crystals that are large enough to see with the naked eye, and we use the term phaneritic (from the Greek word phaneros meaning visible) to describe that. Typically that means they are larger than about 0.5 millitmeres (mm) — the thickness of a strong line made with a ballpoint pen. (If the crystals are too small to distinguish, which is typical of most volcanic rocks, we use the term aphanitic (from the Greek word aphanos – unseen) The intrusive rocks shown in Figure 3.3.5 are all phaneritic, as are those shown in Exercise 3.6.

In general, the size of crystals is proportional to the rate of cooling. The longer it takes for a body of magma to cool, the larger the crystals can grow. It is not uncommon to see an intrusive igneous rock with crystals up to 1 centimetre (cm) long. In some situations, especially toward the end of the cooling stage, the magma can become water rich. The presence of liquid water (still liquid at high temperatures because it is under pressure) promotes the relatively easy movement of ions, and this allows crystals to grow large, sometimes to several centimetres (Figure 3.4.4). Finally, as already described, if an igneous rock goes through a two-stage cooling process, its texture will be porphyritic (Figure 3.3.7).

Figure 3.4.4 A pegmatitic rock with large crystals

Image Descriptions

Figure 3.4.1 image description: Mineral composition of igneous rocks
 Igneous Rocks Felsic Intermediate Mafic Ultramafic
K-feldspar 0 to 35% 0% 0% 0%
Quartz 25 to 35% 0 to 25% 0% 0%
Plagioclase feldspar 25 to 50% 50 to 70% 0 to 50% 0%
Biotite and/or Amphibole 0 to 20% 20 to 40% 0 to 30% 0%
Pyroxene 0% 0 to 20% 20 to 75% 0% to 75%
Olivine 0% 0% 0 to 25 % 25% to 100%
Intrusive Granite Diorite Gabbro Peridotite
Extrusive Rhyolite Andesite Basalt Komatiite

[Return to Figure 3.4.1]



3.5 Intrusive Igneous Bodies

In most cases, a body of hot magma is less dense than the rock surrounding it, so it has a tendency to move very slowly up toward the surface. It does so in a few different ways, including filling and widening existing cracks, melting the surrounding rock (called country rock“Country rock” is not necessarily music to a geologist’s ears. The term refers to the original “rock of the country” or region, and hence the rock into which the magma intruded to form a pluton.), pushing the rock aside (where it is somewhat plastic), and breaking the rock. Where some of the country rock is broken off, it may fall into the magma, a process called stoping. The resulting fragments, illustrated in Figure 3.5.1, are known as xenoliths (Greek for “strange rocks”).

Figure 3.5.1 Xenoliths of mafic rock in granite, Victoria, B.C. The fragments of dark rock have been broken off and incorporated into the light-coloured granite.

Some upward-moving magma reaches the surface, resulting in volcanic eruptions, but most cools within the crust. The resulting body of rock is known as a pluton. Plutons can have various different shapes and relationships to the surrounding country rock as shown in Figure 3.5.2.

Figure 3.5.2 Depiction of some of the types of plutons. a: stocks (if they coalesce at depth then they might become large enough to be called a batholith), b: sill (a tabular body, in this case parallel to bedding), c: dyke (cross-cuts bedding), d: laccolith (a sill that has pushed up the overlying rock layers), e: pipe (a cylindrical conduit feeding a volcano). The two features labelled f could be pipes or dykes, but from this perspective it’s not possible to determine if they are cylindrical or tabular.

Large irregular-shaped plutons are called either stocks or batholiths. The distinction between the two is made on the basis of the area that is exposed at the surface: if the body has an exposed surface area greater than 100 square kilometres (km2), then it’s a batholith; smaller than 100 km2 and it’s a stock. Batholiths are typically formed only when a number of stocks coalesce beneath the surface to create one large body. One of the largest batholiths in the world is the Coast Range Plutonic Complex, which extends all the way from the Vancouver region to southeastern Alaska (Figure 3.5.3). More accurately, it’s many batholiths.

Tabular (sheet-like) plutons are distinguished on the basis of whether or not they are concordant with (i.e., parallel to) existing layering (e.g., sedimentary bedding or metamorphic foliation) in the country rock. A sill is concordant with existing layering, and a dyke is discordant. If the country rock has no bedding or foliation, then any tabular body within it is a dyke. Note that the sill-versus-dyke designation is not determined simply by the orientation of the feature. A dyke can be horizontal and a sill can be vertical (if the bedding is vertical). A large dyke can be seen in Figure 3.5.3.

A laccolith is a sill-like body that has expanded upward by deforming the overlying rock.

Finally, a pipe is a cylindrical body (with a circular, ellipitical, or even irregular cross-section) that served as a conduit for the movement of magma from one location to another. Most known pipes fed volcanoes, although pipes can also connect plutons. It is also possible for a dyke to feed a volcano.

Figure 3.5.3 The Stawamus Chief, part of the Coast Range Plutonic Complex, near to Squamish, B.C. The cliff is about 600 metres (m) high. Most of the dark stripes are a result of algae and lichen growth where the surface is frequently wet, but there is a large (about 10 m across) vertical dyke that extends from bottom to top.

As discussed already, plutons can interact with the rocks into which they are intruded, sometimes leading to partial melting of the country rock or to stoping and formation of xenoliths. And, as we’ll see in Chapter 7, the heat of a body of magma can lead to metamorphism of the country rock. The country rock can also have an effect on the magma within a pluton. The most obvious such effect is the formation of a chilled margin along the edges of the pluton, where it came in contact with country rock that was significantly colder than the magma. Within the chilled margin, the magma cooled more quickly than in the centre of the dyke, so the texture is finer and the colour may be different. An example is shown in Figure 3.5.4.

Figure 3.5.4 A mafic dyke with chilled margins within basalt at Nanoose, B.C. The coin is 24 mm in diameter. The dyke is about 25 centimetres (cm) across and the chilled margins are 2 cm wide.

Exercise 3.7 Pluton Problems

Figure 3.5.5 shows a cross-section through part of the crust showing a variety of intrusive igneous rocks. Except for the granite (a), all of these rocks are mafic in composition. Indicate whether each of the plutons labelled a to e on the diagram below is a dyke, a sill, a stock, or a batholith.


Figure 3.5.5

See Appendix 3 for Exercise 3.7 answers.

Media Attributions



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

Section Summary
3.1 The Rock Cycle The three types of rocks are igneous: formed from magma; sedimentary: formed from fragments of other rocks or precipitation from solution; and metamorphic: formed when existing rocks are altered by heat, pressure, and/or chemical action. The rock cycle summarizes the processes that contribute to cycling of rock material among these three types. The rock cycle is driven by Earth’s internal heat, and by processes happening at the surface, which are driven by solar energy.
3.2 Magma and Magma Formation Magma is molten rock, and in most cases, it forms from partial melting of existing rock. The two main processes of magma formation are decompression melting and flux melting. Magmas range in composition from ultramafic to felsic. Mafic rocks are rich in iron, magnesium, and calcium, and have around 50% silica. Felsic rocks are rich in silica (~75%) and have lower levels of iron, magnesium, and calcium and higher levels of sodium and potassium than mafic rocks.  Intermediate rocks have compositions between felsic and mafic.
3.3 Crystallization of Magma As a body of magma starts to cool, the first process to take place is the polymerization of silica tetrahedra into chains. This increases the magma’s viscosity (makes it thicker) and because felsic magmas have more silica than mafic magmas, they tend to be more viscous. The Bowen reaction series allows us to predict the order of crystallization of magma as it cools. Magma can be modified by fractional crystallization (separation of early-forming crystals) and by incorporation of material from the surrounding rocks by partial melting.
3.4 Classification of Igneous Rock Igneous rocks are classified based on their mineral composition and texture. Felsic igneous rocks have less than 20% ferromagnesian silicates (amphibole and/or biotite) plus varying amounts of quartz and both potassium and plagioclase feldspars. Mafic igneous rocks have more than 50% ferromagnesian silicates (primarily pyroxene) plus plagioclase feldspar. Most intrusive igneous rocks are phaneritic (crystals are visible to the naked eye). If there were two stages of cooling (slow then fast), the texture may be porphyritic (large crystals in a matrix of smaller crystals). If the cooling was extremely slow, or if water was present during cooling, the texture may be pegmatitic (very large crystals).
3.5 Intrusive Igneous Bodies Magma intrudes into country rock by pushing it aside or melting through it. Intrusive igneous bodies tend to be either irregular (stocks and batholiths), tabular (dykes and sills), or pipe-like. Batholiths have exposed areas of greater than 100 km2, while stocks are smaller. Sills are parallel to existing layering in the country rock, while dykes cut across layering. A pluton that intruded into cold rock it is likely to have a chilled margin.

Questions for Review

Answers to Review Questions at the end of each chapter are provided in Appendix 2.

  1. What processes must take place to transform rocks into sediment?
  2. What processes normally take place in the transformation of sediments to sedimentary rock?
  3. What are the processes that lead to the formation of a metamorphic rock?
  4. What is the significance of the term reaction in the name of the Bowen reaction series?
  5. Why is it common for plagioclase crystals to be zoned from relatively calcium-rich in the middle to more sodium-rich on the outside?
  6. What must happen within a magma chamber for fractional crystallization to take place?
  7. Explain the difference between aphanitic and phaneritic textures.
  8. Explain the difference between porphyritic and pegmatitic textures.
  9. Name the following rocks:
    1. An extrusive rock with 40% Ca-rich plagioclase and 60% pyroxene
    2. An intrusive rock with 65% plagioclase, 25% amphibole, and 10% pyroxene
    3. An intrusive rock with 25% quartz, 20% orthoclase, 50% feldspar, and minor amounts of biotit
  10. With respect to tabular intrusive bodies, what is the difference between a concordant body and a discordant body?
  11. Why does a dyke commonly have a fine-grained margin?
  12. What is the difference between a batholith and a stock?
  13. Describe two ways in which batholiths intrude into existing rock.
  14. Why is compositional layering a common feature of mafic plutons but not of felsic plutons?


Chapter 4 Volcanism

Learning Objectives

After carefully reading this chapter, completing the exercises within it, and answering the questions at the end, you should be able to:

  • Explain the relationships between plate tectonics, the formation of magma, and volcanism.
  • Describe the range of magma compositions formed in differing tectonic environments, and discuss the relationship between magma composition and eruption style.
  • Explain the geological and eruption-style differences between different types of volcanoes, especially shield volcanoes, composite volcanoes, and cinder cones.
  • Understand the types of hazards posed to people and to infrastructure by the different types of volcanic eruptions.
  • Describe the symptoms that we can expect to observe when a volcano is ready to erupt, and the techniques that we can use to monitor those volcanic symptoms and predict eruptions.
  • Summarize the types of volcanoes that have erupted in British Columbia since 2.6 Ma, and the characteristics of some of those eruptions.

A volcano is any location where magma comes to the surface, or has done so within the past several million years. This can include eruptions on the ocean floor (or even under the water of a lake), where they are called subaqueous eruptions, or on land, where they are called subaerial eruptions. Not all volcanic eruptions produce the volcanic mountains with which we are familiar; in fact most of Earth’s volcanism takes place along the spreading ridges on the sea floor and does not produce volcanic mountains at all—not even sea-floor mountains.

Canada has a great deal of volcanic rock, but most of it is old, some of it billions of years old. Only in B.C. and the Yukon are there volcanoes that have been active since 2.6 Ma (Pleistocene or younger), and the vast majority of these are in B.C. We’ll look at those in some detail toward the end of this chapter, but a few of them are shown on Figures 4.0.1 and 4.0.2.

The study of volcanoes is critical to our understanding of the geological evolution of Earth, and to our understanding of significant changes in climate. But, most important of all, understanding volcanic eruptions allows us to save lives and property. Over the past few decades, volcanologists have made great strides in their ability to forecast volcanic eruptions and predict the consequences—this has already saved thousands of lives.

Figure 4.0.1 Mount Garibaldi, near Squamish B.C., is one of Canada’s tallest (2,678 metres (m)) and most recently active volcanoes. It last erupted approximately 10,000 years ago.
Figure 4.0.2 Mount Garibaldi (background left, looking from the north) with Garibaldi Lake in the foreground. The lower volcanic peak in the centre is Mount Price and the dark flat–topped peak to the left of it is The Table. All three of these volcanoes were active during the last glaciation.

Media Attributions


4.1 Plate Tectonics and Volcanism

The relationships between plate tectonics and volcanism are shown on Figure 4.1.1. As summarized in Chapter 3, magma is formed at three main plate-tectonic settings: divergent boundaries (decompression melting), convergent boundaries (flux melting), and mantle plumes (decompression melting).

Figure 4.1.1 The plate-tectonic settings of common types of volcanism. Composite volcanoes form at subduction zones, either on ocean-ocean convergent boundaries (left) or ocean-continent convergent boundaries (right). Both shield volcanoes and cinder cones form in areas of continental rifting. Shield volcanoes form above mantle plumes, but can also form at other tectonic settings. 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 Figure 4.1.2. At a spreading ridge, hot mantle rock moves slowly upward by convection (centimetre/year), and within about 60 kilometres (km) of the surface, partial melting starts because of decompression. Over the triangular area shown in Figure 4.1.2a, 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 spills out onto 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.

Exercise 4.1 How thick is the oceanic crust?

Figure 4.1.2a shows a triangular zone about 60 km thick; within this zone, approximately 10% of the mantle rock melts to form oceanic crust. Based on this information, approximately how thick do you think the resulting oceanic crust should be?

Figure 4.1.2 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 boundary (Similar processes take place at an ocean-ocean convergent boundary), and (c) volcanism related to a mantle plume.


See Appendix 3 for Exercise 4.1 answers.

At an ocean-continent convergent boundary, part of a plate that is made up of oceanic crust is subducting beneath part of another plate made up of continental crust. At an ocean-ocean convergent boundary, oceanic crust is being subducted beneath another oceanic-crust plate.[/footnote] (Figure 4.1.2b). In both situations the oceanic crust is heated up, and while there isn’t enough heat to melt the subducting crust, there is enough heat to force the water out of some of its minerals. This released water rises into the overlying mantle where it contributes to flux melting of the 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 thus it assimilates much more felsic material. That magma, now likely intermediate in composition, continues to rise and assimilate crustal material.  In the upper part of the crust, it accumulates into plutons. From time to time, the magma from the plutons rises toward surface, leading to volcanic eruptions. Mount Garibaldi (Figures 4.0.1 and 4.0.2) is an example of subduction-related volcanism. 

A mantle plume is an ascending column of hot rock (not magma) that originates deep in the mantle, possibly just above the core-mantle boundary. Mantle plumes are thought to rise approximately 10 times faster than the rate of mantle convection. The ascending column may be on the order of kilometres to tens of kilometres across, but near the surface it spreads out to create a mushroom-style head that is several tens to over 100 km across. Near 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 beneath the oceans, the early stages of volcanism typically take place on the sea floor. Over time, islands may form like those in Hawaii.

Volcanism in northwestern B.C. (Figures 4.1.3 and 4.1.4) is related to continental rifting. This area is not at a divergent or convergent boundary, and there is no evidence of an underlying mantle plume. A likely explanation is that the crust of northwestern B.C. is being stressed by the northward movement of the Pacific Plate against the North America Plate, and the resulting crustal fracturing provides a conduit for the flow of magma from the mantle. This may, or may not, be an early stage of continental rifting, such as that found in eastern Africa.

Tuya Butte, Edziza, Stikine River, Iskut-Unuk, and Tseax River are volcanic fields in north western BC
Figure 4.1.3 Volcanoes and volcanic fields in the Northern Cordillera Volcanic Province, B.C.
A field of grey rocks molded together to form large, uneven bumps
Figure 4.1.4 Volcanic rock at the Tseax River area, northwestern B.C.

Media Attributions


4.2 Magma Composition and Eruption Style

As noted in the previous section, the types of magma produced in the various volcanic settings can differ 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—leads to increases in the felsic character of the magma.

Processes occuring in magma chambers. Image decription available
Figure 4.2.1 The important processes that lead to changes in the composition of magmas stored within magma chambers within relatively felsic rocks of the crust. [Image Description]

As illustrated in Figure 4.2.1, several processes can make magma that is stored in a chamber within the crust more felsic than it was to begin with, 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 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 felsic components of the country rock melt preferentially. Settling of ferromagnesian crystals from the upper part of the magma, and possible remelting 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 we’ve 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).

Graph comparing volatile compositions of magmas to silica content. Image description available
Figure 4.2.2 Variations in the volatile compositions of magmas as a function of silica content. [Image Description]

The general relationship between the SiO2 content of magma and the amount of volatiles is shown in Figure 4.2.2. 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.

Differences in viscosity and volatile levels 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 form. If the magma is runny enough for gases to rise up through it and escape to surface, the pressure will not become excessive. Assuming that it can break through to the surface, the magma will flow out relatively gently. An eruption that involves a steady non-violent flow of magma is called effusive.

Exercise 4.2 Under pressure!

A cork flying off the top of a champange bottle
Figure 4.2.3

A good analogy for a magma chamber in the upper crust is a plastic bottle of pop on the supermarket shelf. 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 (you don’t have to drink it!) and open it. 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 in 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, forcing the pop with them.

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 conduit for it to move towards surface, it may not flow out. Under these circumstances pressure will continue to build as more magma moves up from beneath and gases continue to exsolve. Eventually some part of the volcano will break and then all of that pent-up pressure will lead to an explosive eruption.

Mantle plume and spreading-ridge magmas tend to be consistently mafic, 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—even likely. Eruption styles can be correspondingly variable.

Image Descriptions

Figure 4.2.1 image description: Four processes that can affect the composition of magma stored in chambers.

  1. Loss of olivine or pyroxene (by crystal settling) makes the upper magma more felsic.
  2. Partial melting of country rock makes the magma more felsic.
  3. Partial or complete melting of xenoliths makes the magma more felsic.
  4. Possible re-melting of olivine or pyroxene can make the lower magma more mafic.

[Return to Figure 4.2.1]

Figure 4.2.2 image description: The following table describes the range of data points adapted from the original scatter plot graph.
Type of Magma Weight % of gasses (mostly H2O) Weight % of SiO2
Mafic 1 to 3% 47 to 50%
Intermediate 3 to 4% 57 to 64%
Felsic 4 to 7% 69 to 72%

[Return to Figure 4.2.2]

Media Attributions


4.3 Types of Volcanoes

There are numerous types of volcanoes or volcanic sources; some of the more common ones are summarized in Table 4.1.

Table 4.1 A summary of the important types of volcanism
[Skip Table]
Type Tectonic Setting Size and Shape Magma and Eruption Characteristics Example
Cinder cone Various; some form on the flanks of larger volcanoes Small (10s to 100s of metres) and steep (Greater than 20°) Most are mafic and form from the gas-rich early stages of a shield- or rift-associated eruption Eve Cone, northern B.C.
Composite volcano Almost all are at subduction zones Medium size (1000s of metres high and up to 20 km across) and moderate steepness (10° to 30°) Magma composition varies from felsic to mafic, and from explosive to effusive Mount St. Helens
Shield volcano Most are at mantle plumes; some are on spreading ridges Large (up to several 1,000 metres high and up to 200 kilometres across), not steep (typically 2° to 10°) Magma is almost always mafic, and eruptions are typically effusive, although cinder cones are common on the flanks of shield volcanoes Kilauea, Hawaii
Large igneous provinces Associated with “super” mantle plumes Enormous (up to millions of square kiometres) and 100s of metres thick Magma is always mafic and individual flows can be 10s of metres thick Columbia River basalts
Sea-floor volcanism Generally associated with spreading ridges but also with mantle plumes Large areas of the sea floor associated with spreading ridges Pillows form at typical eruption rates; lava flows develop if the rare of flow is faster Juan de Fuca ridge
Kimberlite Upper-mantle sourced The remnants are typically 10s to 100s of metres across Most appear to have had explosive eruptions forming cinder cones; the youngest one is dated at about 10 ka, and all others are at least 30 Ma Lac de Gras Kimberlite Field, N.W.T.

The sizes and shapes of typical shield, composite, and cinder-cone volcanoes are compared in Figure 4.3.1, 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 diameter is in the order of 200 km. Its elevation is 4,169 m above sea level. Mount St. Helens, a composite volcano of average size, rises above the surrounding hills of the Cascade Range. Its diameter is about 6 km, and its height is 2,550 m above sea level. Cinder cones are much smaller. On this drawing, even a large cinder cone is just a dot.

Figure 4.3.1 Profiles of Mauna Loa shield volcano, Mount St. Helens composite volcano, and a large cinder cone.

Cinder Cones

Cinder cones, like Eve Cone in northern B.C. (Figure 4.3.2), are typically only a few hundred metres in diameter, and few are more than 200 m high. Most are made up of fragments of vesicular mafic rock (scoria) that were expelled as the magma boiled when it approached the surface, creating fire fountains. In many cases, these later became the sites of effusive lava flows when the gases were depleted. Most cinder cones are monogenetic, meaning that they formed 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. They can be easily, and relatively quickly, eroded away.

Figure 4.3.2 Eve Cone, situated near to Mount Edziza in northern B.C., formed approximately 700 years ago.

Composite Volcanoes

Composite volcanoes, like Mount St. Helens in Washington State (Figure 4.3.3), are almost all associated with subduction at convergent plate boundaries—either ocean-continent or ocean-ocean boundaries (Figure 4.1.2b). They can extend up to several thousand metres from the surrounding terrain, and, with slopes ranging up to 30˚ They can be up to about 20 km across.   At many such volcanoes, magma is stored in a magma chamber in the upper part of the crust. For example, at Mount St. Helens, there is evidence of a magma chamber that is approximately 1 kilometre wide and extends from about 6 km to 14 km below the surface (Figure 4.3.4). Systematic variations in the composition of volcanism over the past several thousand years at Mount St. Helens imply that the magma chamber is zoned, from more felsic at the top to more mafic at the bottom.

Figure 4.3.3 The north side of Mount St. Helens in southwestern 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 mafic and less viscous lava led to construction of a dome inside the crater.
Figure 4.3.4 A cross-section through the upper part of the crust at Mount St. Helens showing the zoned magma chamber. [Image Description]

Mafic eruptions (and some intermediate eruptions), on the other hand, produce lava flows; the one shown in Figure 4.3.5b is thick enough (about 10 m in total) to have cooled in a columnar jointing pattern (Figure 4.3.7). Lava flows both flatten the profile of the volcano (because the lava typically flows farther than pyroclastic debris falls) and protect the fragmental deposits from erosion. Even so, composite volcanoes tend to erode quickly. Patrick Pringle, a volcanologist with the Washington State Department of Natural Resources, describes Mount St. Helens as a “pile of junk.” The rock that makes up Mount St. Helens ranges in composition from rhyolite (Figure 4.3.5a) to basalt (Figure 4.3.5b); this implies that the types of past eruptions have varied widely in 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 a conduit opens, and then an explosive eruption results from the gas-rich upper part of the magma chamber, producing pyroclastic debris, as shown on Figure 4.3.5a. This type of eruption can also lead to rapid melting of ice and snow on a volcano, which typically triggers large mudflows known as lahars (Figure 4.3.5a). Hot, fast-moving pyroclastic flows and lahars are the two main causes of casualties in volcanic eruptions. Pyroclastic flows killed approximately 30,000 people during the 1902 eruption of Mount. 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 people in the Colombian town of Armero, about 50 km from the volcano.

In a geological context, composite volcanoes tend to form relatively quickly and do not last very long. Mount 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, 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 not strong.

Figure 4.3.5 Mount 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. [Image Description]

Exercise 4.3 Volcanoes and Subduction

Figure 4.3.6

The map shown here illustrates the interactions between the North America, Juan de Fuca, and Pacific Plates off the west coast of Canada and the United States. The Juan de Fuca Plate is forming along the Juan de Fuca ridge, and is then subducted beneath the North America Plate along the red line with teeth on it (“Subduction boundary”).

  1. Using the scale bar in the lower left of the map, estimate the average distance between the subduction boundary and the Cascadia composite volcanoes.
  2. If the subducting Juan de Fuca Plate descends 40 km for every 100 km that it moves inland, what is its likely depth in the area where volcanoes are forming?

See Appendix 3 for Exercise 4.3 answers.

Figure 4.3.7 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.

Shield Volcanoes

Most shield volcanoes are associated with mantle plumes, although some form at divergent boundaries, either on land or on the sea floor. Because of their non-viscous mafic magma they tend to have relatively gentle slopes (2 to 10˚) and the larger ones can be over 100 km in diameter. 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, has been erupting, virtually without interruption, since 1983. 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 4.3.8). In this area, the Pacific Plate is moving northwest at a rate of about 7 centimetres (cm) per year. This means that the earlier formed — and now extinct — volcanoes have now moved well away from the mantle plume. As shown on Figure 4.3.8, 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 km and 11 km below surface.Lin, G, Amelung, F, Lavallee, Y, and Okubo, P, 2014, Seismic evidence for a crustal magma reservoir beneath the upper east rift zone of Kilauea volcano, Hawaii. Geology. V.

Figure 4.3.8 The mantle plume beneath the volcanoes of the island of Hawaii

Although it is not a prominent mountain (Figure 4.3.2), Kilauea volcano has a large caldera in its summit area (Figure 4.3.9). 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, which 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 4.3.9 Aerial view of the Kilauea caldera. The caldera is about 4 km across, and up to 120 m deep. It encloses a smaller and deeper crater known as Halema’uma’u.

One of the conspicuous features of Kilauea caldera is rising water vapour (the white cloud in Figure 4.3.9) and a strong smell of sulphur (Figure 4.3.10). 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 35 years, has been effusive, not explosive.

Figure 4.3.10 A gas-composition monitoring station (left) within the Kilauea caldera close to the edge of Halema’uma’u crater. The rising clouds are mostly composed of water vapour, but also include carbon dioxide and sulphur dioxide. Sulphur crystals (right) have formed around a gas vent in the caldera.

The Kilauea eruption that began in 1983 started with the formation of a cinder cone at Pu’u ’O’o, approximately 15 km east of the caldera (Figure 4.3.11). The magma feeding this eruption flowed along a major conduit system known as the East Rift, which extends for about 20 km from the caldera, first southeast and then east. Lava fountaining and construction of the Pu’u ’O’o cinder cone (Figure 4.3.12a) continued until 1986 at which time the flow became effusive. From 1986 to 2014, lava flowed from a gap in the southern flank of Pu’u ’O’o down the slope of Kilauea through a lava tube (Figure 4.3.12d), emerging at or near the ocean. During 2014 and 2015, the lava flowed northeast toward the community of Pahoa (see Exercise 4.4).  In May of 2018 a new eruption started another 15 km east of the 2014/15 flow in the area known as Leilani Estates. The Lower East Rift Flow was active for 6 months. During that time, 35 km2 of existing land was covered in lava and 3.5 km2 of new land was created (Figure 4.3.11), about 48 km of road were covered in lava and 716 dwellings were destroyed (see USGS Overview of Kilauea Volcanoe’s 2018 eruption[PDF]).  Volcanic activity on the East Rift ceased in August 2018, and there has been no activity on Kilauea since then.  This appears to mark the end of the eruption cycle that lasted—with only a few short interruptions—for 35 years. Kilauea will almost certainly erupt again within years or decades.

Figure 4.3.11 Satellite image of Kilauea volcano showing the East rift and Pu’u ’O’o, the site of the eruption that started in 1983 and continues to the present day.  The 2014-2015 lava flow described in Exercise 4.4 is shown with a yellow arrow, and the area of the 2018 lava flow is shown in red. The puffy white blobs are clouds.

The two main types of textures created during effusive subaerial eruptions are pahoehoe and aa. Pahoehoe, ropy lava that forms as non-viscous lava, flows gently, forming a skin that gels and then wrinkles because of ongoing flow of the lava below the surface (Figure 4.3.12b, and “lava flow video”). Aa, or blocky lava, forms when magma is forced to flow faster than it is able to (down a slope for example) (Figure 4.3.12c). Tephra (lava fragments) is produced during explosive eruptions, and accumulates in the vicinity of cinder cones.

Figure 4.3.12d is a view into an active lava tube on the southern edge of Kilauea. The red glow is from a stream of very hot lava (~1200°C) that has flowed underground for most of the 8 km from the Pu’u ’O’o vent. Lava tubes form naturally and readily on both shield and composite volcanoes because flowing mafic lava preferentially cools near its margins, forming solid lava levées that eventually close over the top of the flow. The magma within a lava tube is not exposed to the air, so it remains hot and fluid and can flow for tens of km, thus contributing to the large size and low slopes of shield volcanoes. The Hawaiian volcanoes are riddled with thousands of old lava tubes, some as long as 50 km.

Figure 4.3.12 Images of Kilauea volcano. (a) Pu’u’O’o cinder cone in the background with tephra in the foreground and aa lava in the middle, (b) Formation of pahoehoe on the southern edge of Kilauea, (c) Formation of aa on a steep slope on Kilauea, (d) Skylight in an active lava tube, Kilauea. Photos B & C taken in 2002 photos A & D taken in 2007.

Kilauea started forming at approximately 300 ka, while neighbouring Mauna Loa dates back to 700 ka and nearby Mauna Kea ito around 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 have emerged from the sea floor, and its other neighbours, Mauna Loa and Mauna Kea, will have become significantly eroded, like their cousins, the islands to the northwest (Figure 4.3.8).

Exercise 4.4 Kilauea’s 2014 lava flow

The U.S. Geological Survey Hawaii Volcano Observatory (HVO) map shown here, dated January 29, 2015, shows the outline of lava that started flowing northeast from Pu’u ’O’o on June 27, 2014 (the “June 27th Lava flow,” a.k.a. the “East Rift Lava Flow”). The flow reached the nearest settlement, Pahoa, on October 29, 124 days later. After damaging some infrastructure west of Pahoa, the flow stopped advancing. A new outbreak occurred November 1, branching out to the north from the main flow.

What is the average rate of advance of the flow front from June 27 to October 29, 2014, in metres per day and metres per hour?

The U.S. Geological Survey Hawaii Volcano Observatory (HVO) map. Image description available.
Figure 4.3.13 [Image Description]

See Appendix 3 for Exercise 4.4 Answers.

Large Igneous Provinces

While the Hawaii mantle plume has produced a relatively low volume of magma for a very long time (~85 Ma), 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 4.3.14). This volcanism, which covered an area of about 160,000 square kilometres (km2) with basaltic rock up to several hundred metres thick, took place between 17 and 14 Ma.

Figure 4.3.14 A part of the Columbia River Basalt Group at Frenchman Coulee, eastern Washington. All of the flows visible here have formed large (up to two m in diameter) columnar basalts, a result of relatively slow cooling of flows that are tens of metres thick. The inset map shows the approximate extent of the 17 to 14 Ma Columbia River Basalts, with the location of the photo shown as a star. [Image Description]

Most other LIP eruptions are much bigger. The Siberian Traps (also basalt), which erupted at the end of the Permian period at 250 Ma, are estimated to have produced approximately 40 times as much lava as the CRBG.

The mantle plume that is assumed to be responsible for the CRBG is now situated beneath the Yellowstone area, where it leads to felsic volcanism. Over the past 2 Ma three very large explosive eruptions at Yellowstone have yielded approximately 900 cubic kilometres (km3) of felsic magma, about 900 times the volume of the 1980 eruption of Mount 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 known being the one that created the Ontong Java plateau in the western Pacific Ocean at around 122 Ma. But most sea-floor volcanism originates at divergent boundaries and involves relatively low-volume eruptions. Under these conditions, hot lava that oozes out into 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 4.3.15). In terms of area, there is very likely more pillow basalt on the sea floor than any other type of rock on Earth.

Figure 4.3.15 (Left) Modern sea-floor pillows in the south Pacific. (Right) Ancient sea-floor pillow basalts. Eroded 40 to 50 Ma pillows on the shore of Vancouver Island, near to Sooke. The pillows are 30 to 40 cm in diameter.


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 called kimberlites that have their origins much deeper in the mantle, at depths of 150 km to 450 km. During a kimberlite eruption, material from this depth may make its way to surface quickly (hours to days) with little interaction with the surrounding rocks. As a result, kimberlite eruptive material is representative of mantle compositions: it is ultramafic.

Kimberlite eruptions that originate at depths greater than 200 km, within areas beneath old thick crust (shields), traverse the region of stability of diamond in the mantle, and in some cases, bring diamond-bearing material 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 4.3.16).

Figure 4.3.16 Ekati diamond mine, Northwest Territories, part of the Lac de Gras kimberlite field

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

How frequently do volcanoes erupt?

The Smithsonian Institution maintains a comprehensive catalogue of the world’s volcanoes, with information and eruptive history for nearly 2700 volcanic sites.  If you spend some time looking around that site you’ll discover the frequency of eruptions at different volcanoes is enormously variable, although we can make some generalizations. Focusing just on shield volcanoes and composite volcanoes some of the data are as follows:

Table 4.2 Eruptions of composite and shield volcanoes
Composite volcanoes Shield volcanoes
Avachinsky (Russia): 5 eruptions over the past 7000 years Fernandina (Galapagos): 31 eruptions over the past 1000 years
Pinatubo (Philippines): 4 eruptions over the past 9000 years Kilauea (Hawaii): 62 eruptions over the past 250 years
Adams (Oregon, USA): 6 eruptions over the past 7000 years Nyamuragira (Congo): 48 eruptions over the past 154 years

Based only on these numbers it is evident that, in general, shield volcanoes are much more active than composite volcanoes, but there are many exceptions to this trend.  Some composite volcanoes are nearly as active as the shield volcanoes listed here, and some shield volcanoes that are still considered to be “active” are almost as inactive as the composite volcanoes listed here.

Image Descriptions

Figure 4.3.4 image description: Mount St. Helens rises over 2.5 kilometres above sea lever and consists mostly of rock less than 3,000 years old. Underneath the mountain is older volcanic rock. Just below sea level is a small magma chamber, which is a probable reservoir for 1981 and later eruptions. Down 5 to 14 kilometres below sea level is the main magma chamber. Variations in the composition of the erupted magma imply this chamber is stratified, with more magma at the bottom. [Return to Figure 4.3.4]

Figure 4.3.5 image description: Image (A) shows a cliff wall with grey/brown and orange horizontal layers. The sides look soft like they would be easily worn away.  The grey/brown layers are lahar deposits and the orange layers are felsic pyroclastic deposits. Image (B) shows a columnar basalt lava flow which looks like a rocky, stone cliff with vertical layers. [Return to Figure 4.3.5]

Figure 4.3.13 image description: The U.S. Geological Survey Hawaii Volcano Observatory (HVO) map, dated January 29, 2015, shows the outline of lava that started flowing northeast from Pu’u ’O’o on June 27, 2004 (the “June 27th Lava flow,” a.k.a. the “East Rift Lava Flow”). The flow reached the nearest settlement, Pahoa, on October 29, after covering a distance of 20 km in 124 days. After damaging some infrastructure west of Pahoa, the flow stopped advancing. A new outbreak occurred November 1, branching out to the north from the main flow about 6 km southwest of Pahoa. [Return to Figure 4.3.13]

Figure 4.3.14 image description: The Columbia River Basalt Group covers most of south eastern Washington state and stretches along the borders between Washington, Idaho, and Oregon. The columnar basalts shown in the photo are in in eastern Washington. They rise up out of a flat valley as tall cliffs. [Return to Figure 4.3.14]

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4.4 Volcanic Hazards

There are two classes of volcanic hazards, direct and indirect. Direct hazards are forces that 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 degradation. It is estimated that indirect effects of volcanism have accounted for approximately 8 million deaths during historical times, while direct effects have accounted for fewer than 200,000, or 2.5% of the total. Some of the more important types of volcanic hazards are summarized in Table 4.3.

Table 4.3 A summary of the important volcanic hazards
[Skip Table]
Type Description Risk
Tephra emissions Small particles of volcanic rock emitted into the atmosphere
  • Respiration problems for some individuals
  • Significant climate cooling leading to crop failure and famine
  • Damage to aircraft
Gas emissions The emission of gases before, during, and after an eruption
  • Climate cooling leading to crop failure and famine
  • In some cases, widespread poisoning
Pyroclastic density current A very hot (several 100°C) mixture of gases and volcanic fragments (tephra) that flows rapidly (up to 100s of kilometres per hour (km/h)) down the side of a volcano Extreme hazard — destroys anything in the way
Pyroclastic fall Vertical fall of tephra in the area surrounding an eruption
  • Thick tephra coverage of areas close to the eruption (1 km to 10s of kilometres)
  • Collapsed roofs
Lahar A flow of mud and debris down a channel leading away from a volcano, triggered either by an eruption or a severe rain event Severe risk of destruction for anything within the channel—lahar mud flows can move at 10s of km/h
Sector collapse/ debris avalanche The failure of part of a volcano, either due to an eruption or for some other reason, leading to the failure of a large portion of the volcano Severe risk of destruction for anything in the path of the debris avalanche
Lava flow The flow of lava away from a volcanic vent People and infrastructure at risk, but lava flows tend to be slow (less than km/h) and are relatively easy to avoid

Volcanic Gas and Tephra Emissions

Large volumes of tephra (rock fragments, mostly pumice) and gases are emitted during major plinian eruptions (large explosive eruptions with hot gas and tephra columns extending into the stratosphere) at composite volcanoes, and a large volume of gas is released during some very high-volume effusive eruptions. One of the major effects is cooling of the climate by 1° to 2°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 significant event of this type was in 1991 and 1992 following the large eruption of Mount Pinatubo in the Philippines. A temperature decrease of 1° to 2°C may not seem like very much, but that is the global average amount of cooling, and cooling was much more severe in some regions and at some times.

Over an eight-month period in 1783 and 1784, a massive 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, along with a significant volume of hydrofluoric acid (HF). The sulphate aerosols that formed in the atmosphere led to dramatic cooling in the northern hemisphere. There were serious crop failures in Europe and North America, and a total of 6 million people are estimated to have died from famine and respiratory complications. 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 more than 10,000 human deaths—about 25% of the population.

Volcanic ash can also have serious implications for aircraft because it can destroy jet engines. For example, over 5 million airline passengers had their travel disrupted by the 2010 Eyjafjallajökull volcanic eruption in Iceland.

Pyroclastic Density Currents

In a typical explosive eruption at a composite volcano, the tephra and gases are ejected with explosive force and are hot enough to be forced high up into the atmosphere. As the eruption proceeds, and the amount of gas in the rising magma starts to decrease, parts will become heavier than air, and they can then flow downward along the flanks of the volcano (Figure 4.4.1). As they descend, they cool more and flow faster, reaching speeds up to several hundred kilometres per hour. A pyroclastic density current (PDC) consists of tephra ranging in size from boulders to microscopic shards of glass (made up of the edges and junctions of the bubbles of shattered pumice), plus gases (dominated by water vapour, but also including other gases). The temperature of this material can be as high as 1000°C. Among the most famous PDCs are the one that destroyed Pompeii in the year 79 CE, killing an estimated 18,000 people, and the one that destroyed the town of St. Pierre, Martinique, in 1902, killing an estimated 30,000.

The buoyant upper parts of pyroclastic density currents can flow over water, in some cases for several kilometres. The 1902 St. Pierre PDC flowed out into the city’s harbour and destroyed several wooden ships anchored there.

Figure 4.4.1 The plinian eruption of Mount 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 Fall

Most of the tephra from an explosive eruption ascends high into the atmosphere, and some of it is distributed around 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 Mount 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, leading to the collapse of thousands of roofs and to at least 300 of the 700 deaths attributed to the eruption.


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 Colombian town of Armero in 1985 (described earlier). 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.

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! Mudflows and debris flows occurred in many areas, especially in Honduras and Nicaragua. An example is at the 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 4.4.2). 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 of lahars.

Figure 4.4.2 Part of the path of the lahar from Casita Volcano, October 30, 1998.

Sector Collapse and Debris Avalanche

In the context of volcanoes, sector collapse or flank collapse is the catastrophic failure of a significant part of an existing volcano, creating a large debris avalanche. This hazard was first recognized with the failure of the north side of Mount St. Helens immediately prior to the large eruption on May 18, 1980. In the weeks before the eruption a large bulge had formed on the side of the volcano, the result of magma transfer from depth into a satellite magma body within the mountain itself. Early on the morning of May 18, a moderate earthquake struck nearby; this is thought to have destabilized the bulge, leading to Earth’s largest ever observed slope failure. The failure of this part of the volcano exposed the underlying satellite magma chamber, causing it to explode sideways, which then exposed the conduit leading to the magma chamber below. The resulting plinian eruption—with a 24 kilometre high eruption column—lasted for nine hours.

In August 2010, a massive part of the flank of B.C.’s Mount Meager gave way and about 48 million cubic metres (m3) of rock rushed down the valley, one of the largest slope failures in Canada in historical times (Figure 4.4.3). More than 25 slope failures have taken place at Mount Meager in the past 8,000 years, some of them more than 10 times larger than the 2010 failure.

Figure 4.4.3 The August 2010 Mount Meager rock avalanche, showing where the slide originated (arrow, 4 km upstream), its path down a steep narrow valley, and the debris field (and the stream that eventually cut through it) in the foreground.

Lava Flows

As we saw 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.

Exercise 4.5 Volcanic hazards in Squamish

Figure 4.4.4

The town of Squamish is situated approximately 10 km from Mount Garibaldi, as shown in the photo. If Mount Garibaldi were to erupt, which of the following hazards could be an issue for people in and around Squamish? Explain why or why not.

  1. Tephra emission.
  2. Gas emission.
  3. Pyroclastic denisty current.
  4. Pyroclastic fall.
  5. Lahar.
  6. Sector collapse.
  7. Lava flow.

See Appendix 3 for Exercise 4.5 answers.

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4.5 Monitoring Volcanoes and Predicting Eruptions

In 2005 USGS geologist Chris Newhall 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 into the atmosphere through cracks in the overlying rock
  2. Bit of a bulge — the deformation of part of the volcano, indicating that a magma chamber at depth is swelling or becoming 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, which may indicate that magma has stalled, which could mean that something is about to give way
  5. Big bump — a pronounced bulge on the side of the volcano (like the one at Mount St. Helens in 1980), which may indicate that magma has moved close to surface
  6. Blowing off steam — steam eruptions (a.k.a. phreatic eruptions) that happen when magma near the surface heats groundwater to the boiling point. The water eventually explodes, sending 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 can take to monitor a volcano and predict when it might erupt.

Assessing Seismicity

The simplest and cheapest way to monitor a volcano is with seismometers. In an area with several volcanoes that have the potential to erupt (e.g., the Squamish-Pemberton area), a few well-placed seismometers can provide us with an early warning that something is changing beneath one of the volcanoes, and that we need to take a closer look. There are currently enough seismometers in the Lower Mainland and on Vancouver Island to provide this information.See:

If there is seismic evidence that a volcano is coming to life, more seismometers should be placed in locations within a few tens of kilometres of the source of the activity (Figure 4.5.1). 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 4.5.1 A seismometer and related power and communication equipment installed in 2007 in the vicinity of the Nazco Cone, British Columbia

Detecting Gases

Water vapour quickly turns into clouds of liquid water droplets and is relatively easy to detect just by looking, but CO2 and SO2 are not as 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 monitored from a distance (from the ground or even from the air) using infrared devices, but to obtain more accurate data, we need to 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 4.3.12), or by collecting samples of the air and analyzing them in a lab.

Measuring Deformation

Measurement of deformation allows us to determine if a volcano is expanding or contracting. That can typically be related to the movement of magma into or out of a magma chamber near to surface and so is an indicator of the potential for an eruption.  There are two main ways to measure ground deformation. One is known as a tiltmeter, which is a sensitive three-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 system) technology (Figure 4.5.2). Both are effective, but GPS provides more information than a tiltmeter because it can tell us how far the ground has actually moved in the east-west, north-south, and up-down directions.

Figure 4.5.2 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, along with careful observations made on the ground and from the air, and a thorough knowledge of how volcanoes work, 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, there is no longer a large risk of surprise eruptions, 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 like 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 Earth.

Exercise 4.6 Volcano alert!

You’re the chief volcanologist for the Geological Survey of Canada (GSC), based in Vancouver. At 10:30 a.m. on a Tuesday, you receive a report from a seismologist at the GSC in Sidney saying that there has been a sudden increase in the number of small earthquakes in the vicinity of Mount Garibaldi. You have two technicians available, access to some monitoring equipment, and a four-wheel-drive vehicle. At noon, you meet with your technicians and a couple of other geologists. By the end of the day, you need to have a plan to implement, starting tomorrow morning, and a statement to release to the press. What should your first day’s fieldwork include? What should you say later today in your press release?

See Appendix 3 for Exercise 4.6 Answers.

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4.6 Volcanoes in British Columbia

As shown on the Figure 4.6.1, three types of volcanic environments are represented in British Columbia:

Subduction Volcanism

Southwestern British Columbia is at the northern end of the Juan de Fuca (Cascadia) subduction zone, and the volcanism there is related to magma generation by flux melting in the upper mantle above the subducting plate. In general, there has been a much lower rate and volume of volcanism in the B.C. part of this belt than in the U.S. part. One possible reason for this is that the northern part of the Juan de Fuca Plate (i.e., the Explorer Plate) is either not subducting, or is subducting at a slower rate than the rest of the plate. There are several volcanic centres in the Garibaldi Volcanic Belt: the Garibaldi centre (including Mount Garibaldi and the Black Tusk-Mount Price area adjacent to Garibaldi Lake (Figures 4.0.1 and 4.0.2), Mount Cayley, and Mount. Meager (Figure 4.4.3). The most recent volcanic activity in this area was at Mount Meager. Approximately 2,400 years ago, an explosive eruption of about the same magnitude as the 1980 Mount St. Helens eruption took place at Mount Meager. Ash spread as far east as Alberta. There was also significant eruptive activity at Mounts Price and Garibaldi approximately 12,000 and 10,000 years ago during the last glaciation; in both cases, lava and tephra built up against glacial ice in the adjacent valley (Figure 4.6.2). The Table in Figure 4.0.2 at the beginning of this chapter is a tuya, a volcano that formed beneath glacial ice and had its top eroded by the lake that formed around it in the ice.

Figure 4.6.2 Perspective view of the Garibaldi region (looking east) showing the outlines of two lava flows from Mount Price. Volcanism in this area last took place when the valley in the foreground was filled with glacial ice. The cliff known as the Barrier formed when part of the Mount Price lava flow failed after deglaciation. The steep western face of Mount Garibaldi formed by sector collapse, also because the rocks were no longer supported by glacial ice.

Mantle Plume Volcanism

The chain of volcanic complexes and cones extending from Milbanke Sound to Nazko Cone is interpreted as being related to a mantle plume currently situated close to the Nazko Cone, just west of Quesnel. The North America Plate is moving in a westerly direction at about 2 cm per year with respect to this plume, and the series of now partly eroded shield volcanoes between Nazco and the coast is interpreted to have been formed by the plume as the continent moved over it.

The Rainbow Range, which formed at approximately 8 Ma, is the largest of these older volcanoes. It has a diameter of about 30 km and an elevation of 2,495 m (Figure 4.6.3). The name “Rainbow” refers to the bright colours displayed by some of the volcanic rocks as they weather.

Figure 4.6.3 Rainbow Range, Chilcotin Plateau, B.C.

Rift-Related Volcanism

While B.C. is not about to split into pieces, two areas of volcanism are related to rifting—or at least to stretching-related fractures that might extend through the crust. These are the Wells Gray-Clearwater volcanic field southeast of Quesnel, and the Northern Cordillera Volcanic Field, which ranges across the northwestern corner of the province (as already discussed in section 4.1). This area includes Canada’s most recent volcanic eruption, a cinder cone and mafic lava flow that formed around 250 years ago at the Tseax River Cone in the Nass River area north of Terrace. According to Nisga’a oral history, as many as 2,000 people died during that eruption, in which lava overran their village on the Nass River. Most of the deaths are attributed to asphyxiation from volcanic gases, probably carbon dioxide.

The Mount Edziza Volcanic Field near the Stikine River is a large area of lava flows, sulphurous ridges, and cinder cones. The most recent eruption in this area was about 1,000 years ago. While most of the other volcanism in the Edziza region is mafic and involves lava flows and cinder cones, Mount Edziza itself (Figure 4.6.4) is a composite volcano with rock compositions ranging from rhyolite to basalt. A possible explanation for the presence of composite volcanism in an area dominated by mafic flows and cinder cones is that there is a magma chamber beneath this area, within which magma differentiation is taking place.

Figure 4.6.4 Mount Edziza, in the Stikine area, B.C., with Eve Cone in the foreground.

Exercise 4.7 Volcanoes down under

This map shows the plate tectonic situation in the area around New Zealand.

Figure 4.6.5
  1. Based on what you know about volcanoes in B.C., predict where you might expect to see volcanoes in and around New Zealand.
  2. What type of volcanoes would you expect to find in and around New Zealand?

See Appendix 3 for Exercise 4.7 Answers.

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

Section Summary
4.1 Plate Tectonics and Volcanism Volcanism is closely related to plate tectonics. Most volcanoes are associated with convergent plate boundaries (at subduction zones), and there is also a great deal of volcanic activity at divergent boundaries and areas of continental rifting. At convergent boundaries magma is formed where water from a subducting plate acts as a flux to lower the melting temperature of the adjacent mantle rock. At divergent boundaries magma forms because of decompression melting. Decompression melting also takes place within a mantle plume.
4.2 Magma Composition and Eruption Style The initial magmas in most volcanic regions are mafic in composition, but they can evolve into more felsic types through interaction with crustal rock, and as a result of crystal settling within a magma chamber. Felsic magmas tend to have higher gas contents than mafic magmas, and they are also more viscous. The higher viscosity prevents gases from escaping from the magma, and so felsic magmas are more pressurized and more likely to erupt explosively.
4.3 Types of Volcanoes Cinder cones, which can form in various volcanic settings, are relatively small volcanoes that were formed during a single eruptive event and are composed mostly of mafic rock fragments. Composite volcanoes are normally associated with subduction, and while their magma tends to be intermediate on average, it can range all the way from felsic to mafic. The corresponding differences in magma viscosity lead to significant differences in eruptions style. Most shield volcanoes are associated with mantle plumes, and have consistently mafic magma which typically erupts as lava flows.
4.4 Volcanic Hazards Most direct volcanic hazards are related to volcanoes that erupt explosively, especially composite volcanoes. Pyroclastic density currents, some as hot as 1000˚C can move at hundreds of kilometres  per hour and will kill anything in the way. Lahars, volcano-related mudflows, can be large enough to destroy entire towns.  Lava flows will destroy anything in their paths, but tend to move slowly enough so that people can get to safety.  But the indirect effects of volcanism have been more deadly in the past, mostly because volcanic ash and gases can lead to short-term significant climate cooling.
4.5 Monitoring Volcanoes and Predicting Eruptions We have the understanding and technology to predict volcanic eruptions with some success, and to ensure that people are not harmed. The prediction techniques include monitoring seismicity in volcanic regions, detecting volcanic gases, and measuring deformation of the flanks of a volcano.
4.6 Volcanoes in British Columbia There are examples of all of the important types of volcanoes in British Columbia, including subduction volcanism north of Vancouver, mantle-plume volcanism along the Nazco trend, and rift-related volcanism in the Wells Gray and Stikine regions.

Questions for Review

Answers to Review Questions at the end of each chapter are provided in Appendix 2.

  1. What are the three main tectonic settings for volcanism on Earth?
  2. What is the primary mechanism for partial melting at a convergent plate boundary?
  3. Why are the viscosity and gas content of a magma important in determining the type of volcanic rocks that will be formed when that magma is extruded?
  4. Why do the gases in magma not form gas bubbles when the magma is deep within the crust?
  5. Where do pillow lavas form? Why do they form and from what type of magma?
  6. What two kinds of rock textures are typically found in a composite volcano?
  7. What is a lahar, and why are lahars commonly associated with eruptions of composite volcanoes?
  8. Under what other circumstances might a lahar form?
  9. Explain why shield volcanoes have such gentle slopes.
  10. In very general terms, what is the lifespan difference between a composite volcano and a shield volcano?
  11. Why is weak seismic activity (small earthquakes) typically associated with the early stages of a volcanic eruption?
  12. How can GPS technology be used to help monitor a volcano in the lead-up to an eruption?
  13. What type of eruption at Mount St. Helens might have produced columnar basalts?
  14. What is the likely geological origin of the Nazko Cone?
  15. What might be the explanation for southwestern B.C. having much less subduction-related volcanism than adjacent Washington and Oregon?
  16. What was the likely cause of most of the deaths from the most recent eruption at the Tseax River Cone?


Chapter 5 Weathering and Soil

Learning Objectives

After carefully reading this chapter, completing the exercises within it, and answering the questions at the end, you should be able to:

  • Explain why rocks formed at depth in the crust are susceptible to weathering at the surface.
  • Describe the main processes of mechanical weathering, and the types of materials that are produced when mechanical weathering predominates.
  • Describe the main processes of chemical weathering, and the products of chemical weathering of minerals such as feldspar, ferromagnesian silicates, and calcite.
  • Explain the type of weathering processes that are likely to have taken place to produce a particular sediment deposit.
  • Discuss the relationships between weathering and soil formation, and the origins of soil horizons and some of the different types of soil.
  • Describe and explain the distribution of some of the important soil types in Canada.
  • Explain the geological carbon cycle, and how variations in rates of weathering can lead to climate change.
Figure 5.01 The Hoodoos, near Drumheller, Alberta, have formed from the differential weathering of sedimentary rock that was buried beneath other rock since about 100 Ma, but has now been exposed for several thousand years.

Weathering is what takes place when a body of rock is exposed to the “weather”—in other words, to the forces and conditions that exist at Earth’s surface. With the exception of volcanic rocks and some sedimentary rocks, most rocks are formed at some depth within the crust. There they experience relatively constant temperature, high pressure, no contact with the atmosphere, and little or no moving water. Once a rock is exposed at the surface, which is what happens when the overlying rock is eroded away, conditions change dramatically. Temperatures vary widely, there is much less pressure, oxygen and other gases are plentiful, and in most climates, water is abundant (Figure 5.01).

Weathering includes two main processes that are quite different. One is the mechanical breakdown of rock into smaller fragments, and the other is the chemical change of the minerals within the rock to forms that are stable in the surface environment. Mechanical weathering provides fresh surfaces for attack by chemical processes, and chemical weathering weakens the rock so that it is more susceptible to mechanical weathering. Together, these processes create two very important products, one being the sedimentary clasts and ions in solution that can eventually become sedimentary rock, and the other being the soil that is necessary for our existence on Earth.

The various processes related to uplift and weathering are summarized in the rock cycle in Figure 5.02.

Figure 5.02 Weathering can take place once a rock is exposed at surface by uplift and the removal of the overlying rock. [Image Description]

Image Descriptions

Figure 5.02 image description: “The Rock Cycle.” The rock cycle takes place both above and below the earth’s surface. The rock deepest beneath the earth’s surface and under extreme heat and pressure is metamorphic rock. This metamorphic rock can melt and become magma. When magma cools, if below the earth’s surface it becomes “intrusive igneous rock.” If magma cools above the earth’s surface it is “extrusive igneous rock” and becomes part of the outcrop.  The outcrop is subject to weathering and erosion, and can be moved and redeposited around the earth by forces such as water and wind. As the outcrop is eroded, it becomes sediment which can be buried, compacted, and cemented beneath the earth’s surface to become sedimentary rock. As sedimentary rock gets buried deeper and comes under increased heat and pressure, it returns to its original state as metamorphic rock. Rocks in the rock cycle do not always make a complete loop. It is possible for sedimentary rock to be uplifted back above the Earth’s surface and for intrusive and extrusive igneous rock to be reburied and became metamorphic rock. [Return to Figure 5.02]

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5.1 Mechanical Weathering

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. 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 an outcrop.To a geologist, an outcrop is an exposure of bedrock, the solid rock of the crust.

The most important agents of mechanical weathering are:

When a mass of rock is exposed by weathering and 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 the granitic rock in Figure 5.1.1., which, in places, is peeling off like the layers of an onion.

Figure 5.1.1 Exfoliation fractures in granitic rock exposed on the side of the Coquihalla Highway north of Hope, B.C.
Figure 5.1.2 Exfoliation of slate at a road cut in the Columbia Mountains west of Golden, B.C.

Granitic rock tends to exfoliate parallel to the exposed surface because the rock is typically homogenous, and it doesn’t have predetermined planes along which it must fracture. Sedimentary and metamorphic rocks, on the other hand, tend to exfoliate along predetermined planes (Figure 5.1.2).

Figure 5.1.3 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 farther 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 water seeps into cracks in a rock, expands on freezing, and thus enlarges the cracks (Figure 5.1.3). The effectiveness of frost wedging is related to the frequency of freezing and thawing. Frost wedging is most effective in a climate like Canada’s. 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.  If you are ever hiking in the mountains you might hear the effects of frost wedging when the Sun warms a steep rocky slope and the fragments of rock that were pried away from the surface by freezing the night before are released as that ice melts.

In many parts of Canada, the transition between freezing nighttime temperatures and thawing daytime temperatures is frequent — tens to hundreds of times a year. Even in warm coastal areas of southern B.C., freezing and thawing transitions are common at higher elevations. A common feature in areas of effective frost wedging is a talus slope—a fan-shaped deposit of fragments removed by frost wedging from the steep rocky slopes above (Figure 5.1.4).

Figure 5.1.4 An area with very effective frost-wedging near Keremeos, B.C. 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 damage to roads all over North America.

When salt water seeps into rocks and then evaporates on a hot 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 the rocky shorelines of Vancouver Island and the Gulf Islands, where sandstone outcrops are common and salty seawater is readily available (Figure 5.1.5). Salt weathering can also occur away from the coast, because most environments have some salt in them.

Figure 5.1.5 Honeycomb weathering of sandstone on Gabriola Island, B.C. 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 grow, widening the cracks and breaking the rock (Figure 5.1.6). 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.

Figure 5.1.6 Conifers growing on granitic rocks at The Lions, near Vancouver, B.C.

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 shown in Figure 5.1.4. On the steep rock faces at the top of the cliff, rock fragments have been 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 15. Other important agents of erosion that also have the effect of removing the products of weathering include water in streams (Chapter 13), glacial ice (Chapter 16), and waves on the coasts (Chapter 17).

Exercise 5.1 Mechanical weathering

This photo shows granitic rock at the top of Stawamus Chief near Squamish, B.C. Identify the mechanical weathering processes that you can see taking place, or you think probably take place at this location.

Figure 5.1.7

See Appendix 3 for Exercise 5.1 answers.

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5.2 Chemical Weathering

Chemical weathering results from chemical changes to 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 the environmental conditions. Some minerals, like quartz, are virtually unaffected by chemical weathering, while others, like feldspar, are 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 produces weak carbonic acid when combined with water. That process, which is fundamental to most chemical weathering, can be shown as follows:

H2O + CO2 ↔ H2CO3   then    H2CO3 ↔ H+ + HCO3

water + carbon dioxide ↔  carbonic acid    then     carbonic acid  ↔  dissolved hydrogen ions + dissolved bicarbonate ions

Yikes! Chemical formulas

Lots of people seize up when they are asked to read chemical or mathematical formulas.  It’s OK, you don’t necessarily have to!  If you don’t like the formulas just read the text underneath them.  In time you may get used to reading the formulas.

The double-ended arrow “” indicates that the reaction can go either way, but for our purposes these reactions are going towards the right.

Here we have water (e.g., as rain) plus carbon dioxide in the atmosphere, combining to create carbonic acid. Then carbonic acid dissociates (comes apart) to form hydrogen and bicarbonate ions. The amount of CO2 in the air is enough to make weak carbonic acid.  There is typically much more CO2 in the soil, so water that percolates through the soil can become more acidic.  In either case, this acidic water is a critical to chemical weathering.

In some types of chemical weathering the original mineral becomes altered to a different mineral. For example, feldspar is altered—by hydrolysis—to form clay minerals plus some ions in solution. In other cases the minerals dissolve completely, and their components go into solution. For example, calcite (CaCO3) is soluble in acidic solutions.

The hydrolysis of feldspar can be written like this:

CaAl2Si2O+ H2CO3  + ½O2 ↔ Al2Si2O5(OH)4 +   Ca2+ +  CO32−

plagioclase feldspar + carbonic acid kaolinite + dissolved calcium ions + dissolved carbonate ions

This reaction shows calcium-bearing 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 (probably in the ocean) to form the mineral calcite. The hydrolysis of feldspar to clay is illustrated in Figure 5.2.1, 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, and olivine can be converted to the clay mineral serpentine.

Figure 5.2.1 Unweathered (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 much of 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) dissolved iron ions + dissolved carbonate ions + dissolved silicic acid

But in the presence of oxygen and carbonic acid, the dissolved iron is then quickly converted to the mineral hematite:

2Fe2+  + 4HCO3 + ½ O2  +  2H2O ↔ Fe2O3   + 4H2CO3

dissolved iron ions + dissolved bicarbonate ions + oxygen + water hematite + carbonic acid

The equation shown here is for olivine, but it could apply to almost any other ferromagnesian silicate, including pyroxene, amphibole, or biotite. Iron in the sulphide minerals (e.g., pyrite) can also be oxidized in this way. And the mineral hematite is not the only possible end result, as there is a wide range of iron oxide minerals that can form in this way. The results of this process are illustrated in Figure 5.2.2, which shows a granitic rock in which some of the biotite and amphibole have been altered to form the iron oxide mineral limonite.

Figure 5.2.2 A granitic rock containing biotite and amphibole which have been altered near to the rock’s surface to limonite, which is a mixture of iron oxide minerals.

A special type of oxidation takes place in areas where the rocks have elevated levels of sulphide minerals, especially pyrite (FeS2). Pyrite reacts with water and oxygen to form sulphuric acid, as follows:

2FeS2  + 7O2 + 2H2O ↔ 2Fe2+   H2SO+ 2H+

pyrite + oxygen + water dissloved iron ions + sulphuric acid + dissolved hydrogen ions

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 have been mined from deep underground and then piled up and left exposed to water and oxygen. One example of that is the Mt. Washington Mine near Courtenay on Vancouver Island (Figure 5.2.3), but there are many similar sites across Canada and around the world.

Figure 5.2.3 Exposed oxidizing and acid generating rocks and mine waste at the abandoned Mt. Washington Mine, B.C. (left), and an example of acid drainage downstream from the mine site (right).

At many ARD sites, the pH of the runoff water is less than 4 (very acidic). Under these conditions, metals such as copper, zinc, and lead are quite soluble, and this can lead to toxicity for aquatic and other organisms. For many years, the river downstream from the Mt. Washington Mine had so much dissolved copper in it that it was toxic to salmon. Remediation work has since been carried out at the mine and the situation has improved.

The hydrolysis of feldspar and other silicate minerals 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+ or Fe2+). 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  + H+   + HCO3  ↔  Ca2+  + 2HCO3

calcite + dissolved hydrogen ions + dissolved bicarbonate ions ↔  dissolved  calcium ions + dissolved bicarbonate ions

Calcite is the major component of limestone (typically more than 95%), and under surface conditions, limestone can dissolve completely, as shown in Figure 5.2.4. 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 5.2.4 A limestone outcrop on Quadra Island, B.C. 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 chert, which stands out because it is not soluble.

Exercise 5.2 Chemical Weathering

The main processes of chemical weathering are hydrolysis, oxidation, and dissolution. Indicate which process is primarily involved during each of the following chemical weathering changes:

  1. Pyrite to hematite
  2. Calcite to calcium and bicarbonate ions
  3. Feldspar to clay
  4. Olivine to serpentine

See Appendix 3 for Exercise 5.2 answers.

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5.3 The 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 several 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 5.1. 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.

Table 5.1 A list of the typical weathering products of some of the minerals in common rocks
Common Mineral Typical Weathering Products
Quartz Quartz as sand grains
Feldspar Clay minerals plus potassium, sodium, and calcium in solution
Biotite and amphibole Chlorite plus iron and magnesium in solution
Pyroxene and olivine Serpentine plus iron and magnesium in solution
Calcite Calcium and carbonate in solution
Pyrite Iron oxide minerals plus iron in solution and sulphuric acid

Some examples of the products of weathering are shown in Figure 5.3.1. 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 5.3.1 Products of weathering and erosion formed under different conditions. [Image Description]

We’ll talk more about the nature and interpretation of sediments and sedimentary rocks in Chapter 6, but it’s worth considering here why the sand-sized sediments shown in Figure 5.3.1 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 types of weathering that occur at Earth’s surface. It is not affected by weak acids 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.

So when a rock like granite is subject to chemical weathering the feldspar and the ferromagnesian silicates get converted to clays and 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 5.3 Describing the weathering origins of sand

In the left side of the following table, a number of different sands are pictured and described. Describe some of the important weathering processes that might have led to the development of these sands.

See Appendix 3 for Exercise 5.3 answers.

Image Description and Location
Fragments of coral, algae, and urchin from a shallow water area (roughly 2 metres deep) near a reef in Belize. The grain diameters are between 0.1 and 1 millimetres.
Angular quartz and rock fragments from a glacial stream deposit near Osoyoos, B.C. The grain diameters are between 0.25 and 0.5 milimetres.
Rounded grains of olivine (green) and volcanic glass (black) from a beach on the big island of Hawaii. The grains are approximately 1 millimetre across.

Image Descriptions

Figure 5.3.1 image description: Examples of weathering and erosion.

  1. Boulders in a talus deposit at Keremeos. All are angular fragments from the same rock source.
  2. Pebbles on a beach in Victoria. All are rounded fragments of rock from different sources.
  3. Sand from a beach at Gabriola Island. most are angular quartz grains, some are sand-sized fragments of rock.
  4. Sand from a dune in Utah. All are rounded quartz grains.

[Return to Figure 5.3.1]

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5.4 Weathering and the Formation of Soil

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.

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 5.4.1. 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 5.4.1 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, and the consequent types and intensity of vegetation), the type of parent material, the slope of the surface, and the amount of time available.


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 5.4.2).

Figure 5.4.2 Poorly developed soil on wind-blown silt (loess) in an arid part of northeastern Washington State. The thickness shown is about 1 m, and the “soil” is just the upper 2 or 3 cm.

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 5.4.2. 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.

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.

Some unconsolidated materials, such as river-flood deposits, make for especially good soils because they 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.


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.


Even under ideal conditions, soil takes thousands of years to develop. Virtually all of southern Canada was still glaciated up until 14 ka, and most of the central and northern parts of B.C., the prairies, Ontario, and Quebec were still glaciated at 12 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.

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 5.4.3, are:

Although rare in Canada, another type of layer that develops in hot arid regions is known as caliche (pronounced ca-lee-chee). It forms from the downward (or in some cases upward) movement of calcium ions, and the precipitation of calcite within the soil. When well developed, caliche cements the surrounding material together to form a layer that has the consistency of concrete.

Figure 5.4.3 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. 

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 5.4.4). Raindrops can disaggregate exposed soil particles, putting the finer material (e.g., clays) into suspension in the water. Sheetwashunchannelled flow across a surface—carries suspended material away, and channels erode right through the soil layer, removing both fine and coarse material.

Figure 5.4.4 Soil erosion by rain and channelled 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 5.4.5).

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 5.4.5 Soil erosion by wind in Alberta.

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5.5 The Soils of Canada

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 5.2. The distribution of these types of soils (and a few others) in Canada is shown in Figure 5.5.1.

Table 5.2 The nature, origins and distributions of the more important soil orders in Canada
Order Type Brief Description Environment
Forest soils Podsol Well-developed A and B horizons Coniferous forests throughout Canada
Luvisol Clay rich B horizon Northern prairies and central B.C., mostly on sedimentary rocks
Brunisol Poorly developed or immature soil, that does not have the well-defined horizons of podsol or luvisol Boreal-forest soils in the discontinuous permafrost areas of central and western Canada, and also in southern B.C.
Grassland soils Chernozem High levels of organic matter and an A horizon at least 10 centimetres thick Southern prairies (and parts of B.C.’s southern interior), in areas that experience water deficits during the summer
Solonetzic A clay-rich B horizon, commonly with a salt-bearing C horizon Southern prairies, in areas that experience water deficits during the summer
Other important soils Organic Dominated by organic matter; mineral horizons are typically absent Wetland areas, especially along the western edge of Hudson Bay, and in the area between the prairies and the boreal forest
Cryosol Poorly developed soil, mostly C horizon Permafrost areas of northern 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.

Figure 5.5.1 The soil order map of Canada. [Image Description]

In the grasslands of the dry southern parts of the prairie provinces and in some of the drier parts of southern B.C., 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. Development of caliche layers is rare in Canada.

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.

Exercise 5.4 The soils of Canada

Examine Figure 5.5.1, which shows the distribution of soils in Canada. Briefly describe the distributions of the five soils types listed. For each one, explain its distribution based on what you know about the conditions under which the soil forms and the variations in climate and vegetation related to it.

  1. Chernozem
  2. Luvisol
  3. Podsol
  4. Brunisol
  5. Organic

See Appendix 3 for Exercise 5.4 answers.

Image Descriptions

Figure 5.5.1 image description: Soil order map of Canada. Chernozem is found in the southern parts of the prairies provinces (Alberta, Saskachewan, and Manitoba). Luvisol is found in the BC interior, most of central and northern Alberta, and a strip through central Saskatchewan and Manitoba. Podsol is common on the eastern and western coastal areas, as well as northern BC, and much of Quebec and the maritime provinces. Brunisol is common in the southern parts of the territories, northern Sakachewan and Manitoba, as well as parts of Ontario and Quebec. Organic soil is scattered all over the country, but most concentrated in northern Ontario and the lowlands around Hudson Bay.  Cryosol is predominant in the far north. [Return to Figure 5.5.1]

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5.6 Weathering and Climate Change

Earth has two important carbon cycles. One is the biological one, wherein living organisms—mostly plants—consume carbon dioxide from the atmosphere to make their tissues.  After they die almost all of that carbon is released back into the atmosphere when they decay over a period of years or decades. A small proportion of this biological-cycle carbon becomes buried in sedimentary rocks: during the slow formation of coal, as tiny fragments and molecules in organic-rich shale, and as the shells and other parts of marine organisms in limestone. This then becomes part of the geological carbon cycle, a cycle that actually involves a majority of Earth’s carbon, but one that operates very slowly.

The geological carbon cycle is shown diagrammatically in Figure 5.6.1. The various steps in the process (not necessarily in this order) are as follows (the letters are shown on Figure 5.6.1):

a) Organic matter from plants is stored in lake sediments, peat, and permafrost for up to tens of thousands of years, and some may be buried deeper to form coal that can be stored for tens of millions of years.

b) Weathering of silicate minerals converts atmospheric carbon dioxide to dissolved bicarbonate, which is stored in the oceans for thousands to tens of thousands of years.

c) Dissolved carbon is converted by marine organisms to calcite, which is stored in carbonate rocks for hundreds of millions of years.

d) Organic and inorganic carbon compounds are stored in sediments for tens to hundreds of millions of years; some end up in petroleum deposits.

e) Carbon-bearing sediments are transferred to the mantle, where the carbon may be stored for tens of millions to billions of years.

f) During volcanic eruptions, carbon dioxide is released back to the atmosphere, where it is stored for years to decades.

Figure 5.6.1 A representation of the geological carbon cycle. The processes represented by the letters are described in the text above.

During much of Earth’s history, the geological carbon cycle has been balanced, with carbon being released by volcanism at approximately the same rate that it is stored by the other processes. Under these conditions, the climate remains relatively stable.

During some periods of Earth’s history, that balance has been upset. This can happen during prolonged periods of greater than average volcanism. One example is the eruption of the Siberian Traps at around 250 Ma, which appears to have led to strong climate warming over a few million years because of the slow but steady input of extra volcanic CO2 into the atmosphere.

A carbon imbalance is also associated with significant mountain-building events. For example, the Himalayan Range was formed between about 40 and 10 Ma and over that time period—and still today—the rate of weathering on Earth has been enhanced because those mountains are so high and steep and the range is so extensive. The weathering of these rocks—most importantly the hydrolysis of feldspar—has resulted in consumption of atmospheric carbon dioxide and transfer of the carbon to the oceans and to ocean-floor carbonate minerals. The steady drop in carbon dioxide levels over the past 40 million years, which led to the Pleistocene glaciations, is partly attributable to the formation of the Himalayan Range.

A non-geological form of carbon-cycle imbalance is happening today on a very rapid time scale. We are in the process of extracting vast volumes of fossil fuels (coal, oil, and gas) that was stored in rocks over the past several hundred million years, and then converting these fuels to energy and carbon dioxide. By doing so, we are changing the climate faster than has ever happened in the past, and putting both ecosystems and our descendants at significant risk.

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

Section Summary
5.1 Mechanical Weathering Rocks weather when they are exposed to surface conditions, which in most case are quite different from those at which they formed. The main processes of mechanical weathering include exfoliation, freeze-thaw, salt crystallization, and the effects of plant growth.
5.2 Chemical Weathering Chemical weathering takes place when minerals within rocks are not stable in their existing environment. Some of the important chemical weathering processes are hydrolysis of silicate minerals to form clay minerals, oxidation of iron in silicate and other minerals to form iron oxide minerals, and dissolution of calcite.
5.3 The Products of Weathering and Erosion The main products of weathering and erosion are grains of quartz (because quartz is resistant to chemical weathering), clay minerals, iron oxide minerals, rock fragments, and a wide range of ions in solution.
5.4 Weathering and the Formation of Soil Soil is a mixture of fine mineral fragments (including quartz and clay minerals), organic matter, and empty spaces that may be partially filled with water.   Soil formation is controlled by climate (especially temperature and humidity), the nature of the parent material, the slope (because soil can’t accumulate on steep slopes), and the amount of time available. Typical soils have layers called horizons which form because of differences in the conditions with depth.
5.5 The Soils of Canada Canada has a range of soil types related to our unique conditions. The main types of soil form in forested and grassland regions, but there are extensive wetlands in Canada that produce organic soils, and large areas where soil development is poor because of cold conditions.
5.6 Weathering and Climate Change The geological carbon cycle plays a critical role in balancing Earth’s climate. Carbon is released to the atmosphere during volcanic eruptions. Carbon is extracted from the atmosphere during weathering of silicate minerals and this is eventually stored in the ocean and in sediments. Atmospheric carbon is also transferred to organic matter and some of that is later stored in soil, permafrost, and rocks. Our use of geologically stored carbon (fossil fuels) has upset this balance and that has created a climate crisis.

Questions for Review

Answers to Review Questions at the end of each chapter can be found in Appendix 2.

  1. What has to happen to a body of rock before exfoliation can take place?
  2. The climate of central B.C. is consistently cold in the winter and consistently warm in the summer. At what times of year would you expect frost wedging to be most effective?
  3. What are the likely products of the hydrolysis of the feldspar albite (NaAlSi3O8)?
  4. Oxidation weathering of the sulphide mineral pyrite (FeS2) can lead to development of acid rock drainage (ARD). What are the environmental implications of ARD?
  5. Most sand deposits are dominated by quartz, with very little feldspar. Under what weathering and erosion conditions would you expect to find feldspar-rich sand?
  6. What ultimately happens to most of the clay that forms during the hydrolysis of silicate minerals?
  7. Why are the slope and the parent materials important factors in soil formation?
  8. Which soil constituents move downward to produce the B horizon of a soil?
  9. What are the main processes that lead to the erosion of soils in Canada?
  10. Where in Canada would you expect to find a chernozemic soil? What characteristics of this region produce this type of soil?
  11. Where are luvisolic soils found in B.C.?
  12. Why does weathering of silicate minerals, especially feldspar, lead to consumption of atmospheric carbon dioxide? What eventually happens to the carbon that is involved in that process?


Chapter 6 Sediments and Sedimentary Rocks

Learning Objectives

After carefully reading this chapter, completing the exercises within it, and answering the questions at the end, you should be able to:

  • Describe the differences between cobbles, pebbles, sand, silt, and clay and explain the relationship between clast size and the extent to which clasts can be transported by moving water or by wind.
  • Describe the characteristics of the various types of clastic sedimentary rock, including the significance of differences in the composition of sandstones.
  • Explain the differences in the characteristics and depositional environments of various types of chemical sedimentary rocks.
  • Differentiate between various sedimentary depositional environments in both terrestrial and marine environments, and explain how the formation of sedimentary basins can be related to plate tectonic processes.
  • Apply your understanding of the features of sedimentary rocks, including grain characteristics, sedimentary structures, and fossils, to the interpretation of past depositional environments and climates.
  • Explain the importance of and differences between groups, formations, and members.
Figure 6.0.1 The Cretaceous Dinosaur Park Formation at Dinosaur Provincial Park, Alberta, one the world’s most important sites for dinosaur fossils. The rocks in the foreground show cross-bedding, indicative of deposition in a fluvial (river) environment

In Chapter 5, we talked about weathering and erosion, which are the first two steps in the transformation of existing rocks into sedimentary rocks. The remaining steps in the formation of sedimentary rocks are transportation, deposition, burial, and lithification (Figure 6.0.2). Transportation is the movement of sediments or dissolved ions from the site of erosion to a site of deposition; this can be by wind, flowing water, glacial ice, or mass movement down a slope. Deposition takes place where the conditions change enough so that sediments being transported can no longer be transported (e.g., a current slows). Burial occurs when more sediments are piled onto existing sediments, and layers formed earlier are covered and compacted. Lithification is what happens—at depths of hundreds to thousands of metres—when those compacted sediments become cemented together to form solid sedimentary rock.

Figure 6.0.2 The rock cycle, showing the processes related to sedimentary rocks on the right-hand side.

In this textbook, we divide sedimentary rocks into two main types: clastic and chemical. Clastic sedimentary rocks are mainly composed of material that has been transported as solid fragments (clasts). Chemical sedimentary rocks are mainly composed of material that has been transported as ions in solution. It’s important not to assume that mechanical weathering leads only to clastic sedimentary rocks, while chemical weathering leads only to chemical sedimentary rocks. In most cases, millions of years separate the weathering and depositional processes, and both types of sedimentary rocks tend to include at least some material derived from both types of weathering.

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6.1 Clastic Sedimentary Rocks

A clast is a fragment of rock or mineral, ranging in size from less than a micronA micron is a millionth of a metre. There are 1,000 microns in a millimetre. (too small to see) to as big as an apartment block. Various types of clasts are shown in Figure 5.3.1 and in Exercise 5.3. The smaller ones tend to be composed of a single mineral crystal, and the larger ones are typically composed of pieces of rock. As we’ve seen in Chapter 5, most sand-sized clasts are made of quartz because quartz is more resistant to weathering than any other common mineral. Many of the clasts that are smaller than sand size (less than 1/16th millimetre) are made of clay minerals. Most clasts larger than sand size (greater than 2 millimetres) are actual fragments of rock, and commonly these might be fine-grained rock like basalt or andesite, or if they are bigger, coarse-grained rock like granite or gneiss. Sedimentary rocks that are made up of “clasts” are called clastic sedimentary rocks.  A comparable term is “detrital sedimentary rocks”.

Grain-Size Classification

Geologists that study sediments and sedimentary rocks use the Udden-Wentworth grain-size scale for describing the sizes of the grains in these materials (Table 6.1).

Table 6.1 The Udden-Wentworth grain-size scale for classifying sediments and the grains that make up sedimentary rocks
[Skip Table]
Type Description Size range (millimetres) Size range (microns)
Boulder large 1024 and up
medium 512 to 1024
small 256 to 512
Cobble large 128 to 256
small 64 to 128
Pebble (Granule) very coarse 32 to 64
coarse 16 to 32
medium 8 to 16
fine 4 to 8
very fine 2 to 4
Sand very coarse 1 to 2 1000 to 2000
coarse 0.5 to 1 500 to 1000
medium 0.25 to 0.5 (1/4 to 1/2 mm) 250 to 500
fine 0.125 to 0.25 (1/8th to 1/4 mm) 125 to 250
very fine 0.063 to 0.125 (or 1/16th to 1/8th mm) 63 to 125
Silt very course 32 to 63
course 16 to 32
medium 8 to 16
fine 4 to 8
very fine 2 to 4
Clay clay 0 to 2

There are six main grain-size categories; five are broken down into subcategories, with clay being the exception. The diameter limits for each successive subcategory are twice as large as the one beneath it. In general, a boulder is bigger than a toaster and difficult to lift. There is no upper limit to the size of boulder.The largest known free-standing rock (i.e., not part of bedrock) is Giant Rock in the Mojave Desert, California. It’s about as big as an apartment building—seven stories high! A small cobble will fit in one hand, a large one in two hands. A pebble is something that you could throw quite easily. The smaller ones—known as granules—are gravel size, but still you could throw one. You can’t really throw a single grain of sand. Sand ranges from 2 millimetres down to 0.063 millimetres, and its key characteristic is that it feels “sandy” or gritty between your fingers—even the finest sand grains feel that way. Silt is essentially too small for individual grains to be visible, and while sand feels sandy to your fingers, silt feels smooth to your fingers but gritty in your mouth. Clay is so fine that it feels smooth even in your mouth.

Exercise 6.1 Describe the sediment on a beach

Providing that your landscape isn’t covered in deep snow at present, visit a beach somewhere nearby—an ocean shore, a lake shore, or a river bank. Look carefully at the size and shape of the beach sediments. Are they sand, pebbles, or cobbles? If they are not too fine, you should be able to tell if they are well rounded or more angular.

The beach in Figure 6.1.1 is at Sechelt, B.C. Although there is a range of clast sizes, it’s mostly made up of well-rounded cobbles interspersed with pebbles. This beach is subject to strong wave activity, especially when winds blow across the Strait of Georgia from the south. That explains why the clasts are relatively large and are well rounded.


Figure 6.1.1 Pebbles on an ocean beach at Sechelt, B.C.

See Appendix 3 for Exercise 6.1 answers.

If you drop a granule into a glass of water, it will sink quickly to the bottom (less than half a second). If you drop a grain of sand into the same glass, it will sink more slowly (a second or two depending on the size). A grain of silt will take several seconds to get to the bottom, and a particle of fine clay may never get there. The rate of settling is determined by the balance between gravity and friction, as shown in Figure 6.1.2.  Large particles settle quickly because the gravitational force (which is proportional to the mass, and therefore to the volume of the particle) is much greater than the frictional resistance (which is proportional to the surface area of the particle). For smaller particles the difference between gravitational push and frictional resistance is less, so they settle slowly.

Figure 6.1.2 The two forces operating on a grain of sand in water. Gravity is pushing it down, and the friction between the grain and the water is resisting that downward force.

Small particles that settle slowly spend longer suspended in the water, and therefore tend to get moved farther than large particles if the water is moving.


One of the key principles of sedimentary geology is that the ability of a moving medium (air or water) to move sedimentary particles—and keep them moving—is dependent on the velocity of flow. The faster the medium flows, the larger the particles it can move. This is illustrated in Figure 6.1.3. Parts of the river are moving faster than other parts, especially where the slope is greatest and the channel is narrow. Not only does the velocity of a river change from place to place, but it changes from season to season.  During peak dischargeDischarge of a stream is the volume of flow passing a point per unit time. It’s normally measured in cubic metres per second (m3/s). at the location of Figure 6.1.3, the water is high enough to flow over the embankment on the right, and it flows fast enough to move the boulders that cannot be moved during low flows.

Figure 6.1.3 Variations in flow velocity on the Englishman River near Parksville, B.C. When the photo was taken the river was not flowing fast enough anywhere to move the boulders and cobbles visible here.  During flood events the water flows right over the snow-covered bank on the right, and is fast enough to move boulders.

Clasts within streams are moved in several different ways, as illustrated in Figure 6.1.4. Large bed load clasts are pushed (by traction) or bounced along the bottom (by saltation), while smaller clasts are suspended in the water and kept there by the turbulence of the flow. As the flow velocity changes, different-sized clasts may be either incorporated into the flow or deposited on the bottom. At various places along a river, there are always some clasts being deposited, some staying where they are, and some being eroded and transported. This changes over time as the discharge of the river changes in response to changing weather conditions.

Other sediment transportation media, such as waves, ocean currents, and wind, operate under similar principles, with flow velocity as the key underlying factor that controls transportation and deposition.

Figure 6.1.4 Transportation of sediment clasts by stream flow. The larger clasts, resting on the bottom (bedload), are moved by traction (sliding) or by saltation (bouncing). Smaller clasts are kept in suspension by turbulence in the flow. Ions (depicted as + and – in the image, but invisible in real life) are dissolved in the water.

Clastic sediments are deposited in a wide range of environments, including glaciers, slope failures, rivers—both fast and slow—lakes, deltas, and ocean environments—both shallow and deep. If the sedimentary deposits last long enough to get covered with other sediments they may eventually form into rocks ranging from fine mudstone to coarse breccia and conglomerate.

Lithification is the term used to describe a number of different processes that take place within a deposit of sediment to turn it into solid rock (Figure 6.1.5). One of these processes is burial by other sediments, which leads to compaction of the material and removal of some of the intervening water and air. After this stage, the individual clasts are touching one another. Cementation is the process of crystallization of minerals within the pores between the small clasts, and especially at the points of contact between clasts. Depending on the pressure, temperature, and chemical conditions, these crystals might include a range of minerals, the common ones being calcite, hematite, quartz and clay minerals.

Figure 6.1.5  Lithification turns sediments into solid rock. Lithification involves the compaction of sediments and then the cementation of grains by minerals that precipitate from groundwater in the spaces between these grains. Source: Karla Panchuk (2016) CC BY 4.0

The characteristics and distinguishing features of clastic sedimentary rocks are summarized in Table 6.2. Mudrock is composed of at least 75% silt- and clay-sized fragments. If it is dominated by clay, it is called claystone. If it shows evidence of bedding or fine laminations, it is shale; otherwise, it is mudstone. Mudrocks form in very low energy environments, such as lakes, river backwaters, and the deep ocean.

Table 6.2 The main types of clastic sedimentary rocks and their characteristics.
[Skip Table]
Group Examples Characteristics
Mudrock mudstone Greater than 75% silt and clay, not bedded
shale Greater than 75% silt and clay, thinly bedded
Coal Dominated by fragments of partially decayed plant matter often enclosed between beds of sandstone or mudrock.
Sandstone quartz sandstone Dominated by sand, greater than 90% quartz
arkose Dominated by sand, greater than 10% feldspar
lithic wacke dominated by sand, greater than 10% rock fragments, greater than 15% silt and clay
Conglomerate Dominated by rounded clasts, granule size and larger
Breccia Dominated by angular clasts, granule size and larger

Most coal forms in fluvial or delta environments where vegetation growth is vigorous and where decaying plant matter accumulates in long-lasting swamps with low oxygen levels. To avoid oxidation and breakdown, the organic matter must remain submerged for centuries or millennia, until it is covered with another layer of either muddy or sandy sediments. It is important to note that in some textbooks coal is described as an “organic sedimentary rock.” In this book, coal is included with the clastic rocks for two reasons: first, because it is made up of fragments of organic matter; and second, because coal seams (sedimentary layers) are almost always interbedded with layers of clastic rocks, such as mudrock or sandstone. In other words, coal accumulates in environments where other clastic rocks accumulate.

Figure 6.1.6 A compositional triangle for arenite sandstones, with the three most common components of sand-sized grains: quartz, feldspar, and rock fragments. Arenites have less than 15% silt or clay. Sandstones with more than 15% silt and clay are called wackes (e.g., quartz wacke, lithic wacke).

It’s worth taking a closer look at the different types of sandstone because sandstone is a common and important sedimentary rock. Typical sandstone compositions are shown in Figure 6.1.6. Sandstones are mostly made up of sand grains of course, but they also include finer material—both silt and clay. The term arenite applies to a so-called clean sandstone, meaning one with less than 15% silt and clay. Considering the sand-sized grains only (the grains larger than 1/16th mm), arenites with 90% or more quartz are called quartz arenites. If they have more than 10% feldspar and more feldspar than rock fragments, they are called feldspathic arenites or arkosic arenites (or just arkose). If they have more than 10% rock fragments, and more rock fragments than feldspar, they are lithic arenites.“Lithic” means “rock.” Lithic clasts are rock fragments, as opposed to mineral fragments. A sandstone with more than 15% silt or clay is called a wacke (pronounced wackie). The terms quartz wacke, lithic wacke, and feldspathic wacke are used with limits similar to those on the arenite diagram. Another name for a lithic wacke is greywacke.

Some examples of sandstones, magnified in thin section are shown in Figure 6.1.7. (A thin section is rock sliced thin enough so that light can shine through.)


Figure 6.1.7 Microscope photos of three types of sandstone in thin-section. Some of the minerals are labelled: Q=quartz, F=feldspar and L= lithic (rock fragments). The quartz arenite and arkose have relatively little silt-clay matrix, while the lithic wacke has abundant matrix.

Clastic sedimentary rocks in which a significant proportion of the clasts are larger than 2 millimetres are known as conglomerate if the clasts are well rounded, and breccia if they are angular. Conglomerates form in high-energy environments such as fast-flowing rivers, where the particles can become rounded. Breccias typically form where the particles are not transported a significant distance in water, such as alluvial fans and talus slopes. Some examples of clastic sedimentary rocks are shown on Figure 6.1.8.

Figure 6.1.8 Examples of various clastic sedimentary rocks. [Image Description]

Exercise 6.2 Classifying sandstones

Table 6.3 below shows magnified thin sections of three sandstones, along with descriptions of their compositions. Using Table 6.1 and Figure 6.1.6, find an appropriate name for each of these rocks.

Table 6.3 Classifying sandstones
Magnified Thin Section Description
Thin white, grey, and black pieces with jagged edges. Angular sand-sized grains are approximately 85% quartz and 15% feldspar. Silt and clay make up less than 5% of the rock.
Small, flat pieces of light, earthy colours. Rounded sand-sized grains are approximately 99% quartz and 1% feldspar. Silt and clay make up less than 2% of the rock.
Small pieces of various sizes and colours. Angular sand-sized grains are approximately 70% quartz, 20% lithic, and 10% feldspar. Silt and clay make up about 20% of the rock.

See Appendix 3 for Exercise 6.2 answers.


Image Descriptions

Figure 6.1.8 image description: (A) Mudrock with bivalve impressions, Cretaceous Nanaimo group, Browns River, Vancouver Island. A very fine-grained rock with shell impressions. (B) Coarse sandstone with cross-bedding, Cambrian Tapeats Formation Chino Valley, Arizona. (C) Conglomerate with imbricate (aligned, tilted down to the left) cobbles, Cretaceous Geoffrey Formation, Hornby Island, BC. (D) Sedimentary breccia, the Pre-Cambrian Toby Formation, east of Castlegar, BC. [Return to Figure 6.1.8]

Media Attributions

  • Figures 6.1.1, 6.1.2, 6.1.3, 6.1.4, 6.1.5, 6.1.6, 6.1.7, 6.1.8: © Steven Earle. CC BY.
  • Exercise 6.2, first image: Aplite Red © Rudolf Pohl. CC BY-SA.


6.2 Chemical Sedimentary Rocks

Whereas clastic sedimentary rocks are dominated by components that have been transported as solid clasts (clay, silt, sand, etc.), chemical sedimentary rocks are dominated by components that have been transported as ions in solution (Na+, Ca2+, HCO3, etc.). There is some overlap between the two because almost all clastic sedimentary rocks contain cement formed from dissolved ions, and many chemical sedimentary rocks include some clasts. Since ions can stay in solution for tens of thousands of years (some much longer), and can travel for tens of thousands of kilometres, it is virtually impossible to relate chemical sediments back to their source rocks.

Chemical weathering and chemical sedimentary rocks

Many students confuse chemical weathering with chemical sedimentary rocks, or mistakenly assume that when and where chemical weathering is taking place, chemical sedimentary rocks will accumulate. Most ions in solution in rivers, lakes and the ocean are produced during chemical weathering, but those ions can remain in solution for millions of years, and during that time they can travel hundreds of thousands of km (yes, literally around the world, several times). They might eventually come out of solution as a result of a biological process or a change in the chemical conditions and will then become a mineral crystal that can settle to form a chemical sediment.

So the calcium ions that are part of a calcite mud on the sea floor near Australia’s Great Barrier Reef could literally have come from anywhere on Earth (and almost certainly came from many different places), and might have been in solution for as little as a few days or for as long as tens of millions of years.

The most common chemical sedimentary rock, by far, is limestone. Others include chert, banded iron formation, and evaporites. Biological processes are important in the formation of some chemical sedimentary rocks, especially limestone and chert. For example, limestone is made up almost entirely of fragments of marineWe use the word marine when referring to salt water (i.e., oceanic) environments, and the word aquatic when referring to freshwater environments. organisms that manufacture calcite for their shells and other hard parts, and most chert includes at least some of the silica tests (shells) of tiny marine organisms (such as diatoms and radiolarians).


Almost all limestone forms in the oceans, and most of that forms on the shallow continental shelves, especially in tropical regions with coral reefs. Reefs are highly productive ecosystems populated by a wide range of organisms, many of which use calcium and bicarbonate ions in seawater to make carbonate minerals (especially calcite) for their shells and other structures. These include corals, of course, but also green and red algae, urchins, sponges, molluscs, and crustaceans. The hard parts of these organisms are eroded by waves and currents to produce carbonate fragments that accumulate in the surrounding region, as illustrated in Figure 6.2.1.

Figure 6.2.1 Various corals and green algae on a reef at Ambergris, Belize. The light-coloured sand consists of carbonate fragments eroded from the reef organisms, and this is the type of material that will eventually become limestone.

Figure 6.2.2 shows a cross-section through a typical reef in a tropical environment (normally between 40° N and 40° S). Reefs tend to form in areas with clear water (e.g., not close to the mouths of large rivers), and near the edges of steep drop-offs because the reef organisms thrive on nutrient-rich upwelling currents. As the reef builds up, it is eroded by waves and currents to produce carbonate sediments that are transported into the steep offshore fore-reef area and the shallower inshore back-reef area. These sediments are dominated by reef-type carbonate fragments of all sizes, including mud. In many such areas, carbonate-rich sediments also accumulate in quiet lagoons, where mud and mollusc-shell fragments predominate (Figure 6.2.3a) or in offshore areas with strong currents, where either foraminifera tests accumulate (Figure 6.2.3b) or calcite crystallizes inorganically to form ooids—spheres of calcite that form in shallow tropical ocean water with strong currents (Figure 6.2.3c).

Figure 6.2.2 Schematic cross-section through a typical tropical reef.
Figure 6.2.3 Carbonate rocks and sediments: (a) mollusc-rich limestone formed in a lagoon area at Ambergris, Belize, (b) foraminifera-rich sediment from a submerged carbonate sandbar in Belize (c) ooids from a beach at Joulters Cay, Bahamas.

Limestone also accumulates in deeper water, from the steady rain of the carbonate shells of tiny organisms that lived near the ocean surface. The lower limit for limestone accumulation is around 4,000 metres. Beneath that depth, calcite is soluble so limestone does not accumulate.

Calcite can also form on land in a number of environments. Tufa forms at springs (Figure 6.2.4) and travertine forms at hot springs. Similar material precipitates within limestone caves to form stalactites, stalagmites, and a wide range of other speleothems. Tufa, travertine and speleothems make up only a tiny proportion of all limestone.

Figure 6.2.4 Tufa formed at a spring at Johnston Creek, Alberta. The bedded grey rock to the left is limestone.

Dolomite (CaMg(CO3)2) is another carbonate mineral, but dolomite is also the name for a rock composed of the mineral dolomite (although some geologists use the term dolostone to avoid confusion). Dolomite rock is quite common (there’s a whole Italian mountain range named after it), which is surprising since marine organisms don’t make dolomite. All of the dolomite found in ancient rocks has been formed through magnesium replacing some of the calcium in the calcite in carbonate muds and sands. This process is known as dolomitization, and it is thought to take place where magnesium-rich water percolates through the sediments in carbonate tidal flat environments.


As we’ve seen, not all marine organisms make their hard parts out of calcite; some, like radiolarians and diatoms, use silica, and when they die their tiny shells (or tests) settle slowly to the bottom where they accumulate as chert. In some cases, chert is deposited along with limestone in the moderately deep ocean, but the two tend to remain separate, so chert beds within limestone are quite common (Figure 6.2.5), as are nodules, like the flint nodules of the Cretaceous chalk of southeastern England. In other situations, and especially in very deep water, chert accumulates on its own, commonly in thin beds.

Figure 6.2.5 Chert (brown layers) interbedded with Triassic Quatsino Fm. limestone on Quadra Island, B.C. All of the layers have been folded, and the chert, being insoluble and harder than limestone, stands out.

Banded iron formation

Banded iron formation (BIF) is a deep sea-floor deposit of iron oxide that is a common ore of iron (Figure 6.2.6). BIF forms when iron dissolved in seawater is oxidized, becomes insoluble, and sinks to the bottom in the same way that silica tests do to form chert. BIF is prevalent in rocks dating from 2400 to 1800 Ma, a result off changes in the atmosphere and oceans that took place over that time period. Photosynthetic bacteria (i.e., cyanobacteria, a.k.a. blue-green algae) consume carbon dioxide from the atmosphere and use solar energy to convert it to oxygen. These bacteria first evolved around 3500 Ma, and for the next billion years, almost all of that free oxygen was used up by chemical and biological processes, but by 2400 Ma free oxygen levels started to increase in the atmosphere and the oceans. Over a period of 600 million years, that oxygen gradually converted soluble ferrous iron (Fe2+) to insoluble ferric iron (Fe3+), which combined with oxygen to form the mineral hematite (Fe2O3), leading to the accumulation of BIFs on the sea floor. After 1800 Ma, little dissolved iron was left in the oceans and the formation of BIF essentially stopped.

Figure 6.2.6 An example of a banded iron formation with dark iron oxide layers interspersed with chert stained red by hematite. This rock is 2.1 billion years old.


In arid regions many lakes and inland seas have no stream outlet and the water that flows into them is removed only by evaporation. Under these conditions, the water becomes increasingly concentrated with dissolved salts, and eventually some of these salts reach saturation levels and start to crystallize (Figure 6.2.7). Although all evaporite deposits are unique because of differences in the chemistry of the water, in most cases minor amounts of carbonates start to precipitate when the solution is reduced to about 50% of its original volume. Gypsum (CaSO4·H2O) precipitates at about 20% of the original volume and halite (NaCl) precipitates at 10%. Other important evaporite minerals include sylvite (KCl) and borax (Na2B4O7·10H2O). Sylvite is mined at numerous locations across Saskatchewan (Figure 6.2.8) from evaporites that were deposited during the Devonian (~385 Ma) when an inland sea occupied much of the region.

Figure 6.2.7 Spotted Lake, near Osoyoos, B.C. The patterns on the surface are salt. This photo was taken in May when the water was relatively fresh because of winter rains. By the end of the summer the surface of this lake is typically fully encrusted with salt deposits.
Figure 6.2.8 A mining machine at the face of potash ore (sylvite) in the Lanigan Mine near Saskatoon, Saskatchewan. The mineable potash layer (on the right) is about 3 metres thick.

Exercise 6.3 Making evaporite

Figure 6.2.9

This is an easy experiment that you can do at home. Pour about 50 mL (just less than 1/4 cup) of very hot water into a cup and add 2 teaspoons (10 mL) of salt. Stir until all or almost all of the salt has dissolved, then pour the salty water (leaving any undissolved salt behind) into a shallow wide dish or a small plate. Leave it to evaporate for a few days and observe the result.  What is the size range and shape of the crystals you grew?

It may look a little like Figure 6.2.9. These crystals are up to about 3 millimetres across.

See Appendix 3 for Exercise 6.3 answers.

Media Attributions

  • Figures 6.2.1, 6.2.2, 6.2.3ab 6.2.4, 6.2.5, 6.2.7, 6.2.9: © Steven Earle. CC BY.
  • Figure 6.2.3c: JoultersCayOoids by Wilson44691. Public domain.
  • Figure 6.2.6: © Andre Karwath. CC BY-SA.
  • Figure 6.2.8: Photo courtesy of PotashCorp. All rights reserved. Used with permission.



6.3 Depositional Environments and Sedimentary Basins

Sediments accumulate in a wide variety of environments, both on the continents and in the oceans. Some of the more important of these environments are illustrated in Figure 6.3.1.

Figure 6.3.1 Some of the important depositional environments for sediments and sedimentary rocks.

Table 6.4 provides a summary of the processes and sediment types that pertain to the various depositional environments illustrated in Figure 6.3.1. We’ll look more closely at the types of sediments that accumulate in these environments in the last section of this chapter. The characteristics of these various environments, and the processes that take place within them, are also discussed in later chapters on glaciation, mass wasting, streams, coasts, and the sea floor.

Table 6.4 The important terrestrial depositional environments and their characteristics
Environment Important transport processes Depositional environments Typical sediment types
Glacial gravity, moving ice, moving water valleys, plains, streams, lakes glacial till, gravel, sand, silt, and clay
Alluvial gravity steep-sided valleys coarse angular fragments
Fluvial moving water streams gravel, sand, silt, and organic matter (in swampy parts only)
Aeolian wind deserts and coastal regions sand, silt
Lacustrine moving water (flowing into a lake) lakes sand (near the edges only), silt, clay, and organic matter
Evaporite moving water (flowing into a lake) lakes in arid regions salts, clay
Table 6.5 The important marine depositional environments and their characteristics
Environment Important Transport Processes Depositional Environments Typical Sediment Types
Deltaic moving water deltas sand, silt, clay, and organic matter (in swampy parts only)
Beach waves, longshore currents beaches, spits, sand bars gravel, sand
Tidal tidal currents tidal flats silt, clay
Reefs waves and tidal currents reefs and adjacent basins carbonates
Shallow water marine waves and tidal currents shelves and slopes, lagoons carbonates in tropical climates,  sand/silt/clay elsewhere
Lagoonal little transportation lagoon bottom carbonates in tropical climates
Submarine fan underwater gravity flows continental slopes and abyssal plains gravel, sand, mud
Deep water marine ocean currents deep-ocean abyssal plains clay, carbonate mud, silica mud

Most of the sediments that you might see around you, including talus on steep slopes, sand bars in streams, or gravel in road cuts, will never become sedimentary rocks because they have only been deposited relatively recently—perhaps a few centuries or millennia ago—and are likely to be re-eroded before they are buried deep enough beneath other sediments to be lithified. In order for sediments to be preserved long enough to be turned into rock—a process that takes millions or tens of millions of years—they need to have been deposited in a basin that will last that long. Most such basins are formed by plate tectonic processes, and some of the more important examples are shown in Figure 6.3.2.

Figure 6.3.2 Some of the more important types of tectonically produced basins: (a) trench basin, (b) forearc basin, (c) foreland basin, and (d) rift basin.

Trench basins form where a subducting oceanic plate dips beneath the overriding continental or oceanic crust. They can be several kilometres deep, and in many cases, host thick sequences of sediments from eroding coastal mountains. There is a well-developed trench basin off the west coast of Vancouver Island. A forearc basin lies between the subduction zone and the volcanic arc, and may be formed in part by friction between the subducting plate and the overriding plate, which pulls part of the overriding plate down. The Strait of Georgia is a forearc basin. A foreland basin is caused by the mass of the volcanic range depressing the crust on either side. Foreland basins are not only related to volcanic ranges, but can form adjacent to fold belt mountains like the Canadian Rockies. A rift basin forms where continental crust is being pulled apart, and the crust on both sides of the rift subsides. As rifting continues this eventually becomes a narrow sea, and then an ocean basin. The East African rift basin represents an early stage in this process.

Media Attributions


6.4 Sedimentary Structures and Fossils

Through careful observation over the past few centuries, geologists have discovered that the accumulation of sediments and sedimentary rocks takes place according to some important geological principles, as follows:

In addition to these principles, that apply to all sedimentary rocks (as well as volcanic rocks), a number of other important characteristics of sedimentary processes result in the development of distinctive sedimentary features in specific sedimentary environments. By understanding the origins of these features, we can make some very useful inferences about the processes that led to deposition the rocks that we are studying.

Bedding, for example, is the separation of sediments into layers that either differ from one another in textures, composition, colour, or weathering characteristics, or are separated by partings—narrow gaps between adjacent beds (Figure 6.4.1). Bedding is an indication of changes in depositional processes that may be related to seasonal differences, changes in climate, changes in locations of rivers or deltas, or tectonic changes. Partings may represent periods of non-deposition that could range from a few decades to a few millennia. Bedding can form in almost any sedimentary depositional environment.

Figure 6.4.1 The Triassic Sulphur Mt. Formation near Exshaw, Alberta. Bedding is defined by differences in colour and texture, and also by partings (gaps) between beds that may otherwise appear to be similar.

Cross-bedding is bedding that contains angled layers within otherwise horizontal beds, and it forms when sediments are deposited by flowing water or wind. Some examples are shown in Figures 6.0.11, 6.1.7b, and 6.4.2. Cross-beds formed in streams tend to be on the scale of centimetres to tens of centimetres, while those in aeolian (wind deposited) sediments can be on the scale of metres to several metres.

Figure 6.4.2 Cross-bedded Jurassic Navajo Formation aeolian sandstone at Zion National Park, Utah. In most of the layers the cross-beds dip down toward the right, implying a consistent wind direction from right to left during deposition.

Cross-beds form as sediments are deposited on the leading edge of an advancing ripple or dune under steady state conditions (similar flow rate and same flow direction). Each layer is related to a different ripple that advances in the direction of flow, and is partially eroded by the following ripple (Figure 6.4.3). Cross-bedding is a very important sedimentary structure to be able to recognize because it can provide information on the process of deposition, the direction of current flows and, when analyzed in detail, on other features like the rate of flow and the amount of sediment available.

Figure 6.4.3 Formation of cross-beds as a series of ripples or dunes migrates with the flow. Each ripple advances forward (right to left in this view) as more sediment is deposited on its leading face (small arrows). (On each ripple the last deposited layer is represented by small dots.)

Graded bedding is characterized by a gradation in grain size from bottom to top within a single bed. “Normal” graded beds are coarse at the bottom and become finer toward the top.  They are a product of deposition from a slowing current (Figure 6.4.4).  Most graded beds form in a submarine-fan environment (see Figure 6.4.1), where sediment-rich flows descend periodically from a shallow marine shelf down a slope and onto the deeper sea floor. Some graded beds are reversed (coarser at the top), and this normally results from deposition by a fast-moving debris flow (see Chapter 15).

Figure 6.4.4 A graded turbidite bed in Cretaceous Spray Formation rocks on Gabriola Island, B.C. The lower several centimetres of sand and silt probably formed over the duration of less than an hour. The upper few centimetres of fine clay may have accumulated over several hundred years.

Ripples, which are associated with the formation of cross-bedding, may be preserved on the surfaces of sedimentary beds. Ripples can also help to determine flow direction as they tend to have their steepest surface facing in the direction of the flow (see Figure 6.4.3).

In a stream environment, boulders, cobbles, and pebbles can become imbricated, meaning that they are generally tilted in the same direction. Clasts in streams tend to tilt with their upper ends pointing downstream because this is the most stable position with respect to the stream flow (Figure 6.4.5 and Figure 6.1.7c).

Figure 6.4.5 An illustration of imbrication of clasts in a fluvial environment.

Mud cracks form when a shallow body of water (e.g., a tidal flat or pond or even a puddle), into which muddy sediments have been deposited, dries up and cracks (Figure 6.4.6). This happens because the clay in the upper mud layer tends to shrink on drying, and so it cracks because it occupies less space when it is dry.

Figure 6.4.6 Mudcracks in volcanic mud at a hot-spring area near Myvatn, Iceland.

The various structures described above are critical to understanding and interpreting the conditions that existed during the formation of sedimentary rocks. In addition to these, geologists also look very closely at sedimentary grains to determine their mineralogy or lithology (in order to make inferences about the type of source rock and the weathering processes), their degree of rounding, their sizes, and the extent to which they have been sorted by transportation and depositional processes.  Some of the types of differences that we might want to look for are illustrated in Figure 6.4.7.

Figure 6.4.7 Thin section photos of two sandstones with very different grain characteristics. The one on the left has angular grains with a wide range of different types (quartz, feldspar, biotite, rock fragments), and is poorly sorted (grains range from less than 0.05 mm to ~1 mm).  The one on the right has relatively well-rounded grains of quartz only, and the size range is much less (approx. 0.25 to 1 mm).  (Scale bars are 1 mm.)

We won’t be covering fossils in any detail in this book, but they are extremely important for understanding sedimentary rocks. Of course, fossils can be used to date sedimentary rocks, but equally importantly, they tell us a great deal about the depositional environment of the sediments and the climate at the time. For example, they can help to differentiate marine versus terrestrial environments; estimate the depth of the water; detect the existence of currents; and estimate average temperature and precipitation. For example, the tests of tiny marine organisms (mostly foraminifera) have been recovered from deep-ocean sediment cores from all over the world, and their isotopic signatures have been measured. As we’ll see in Chapter 19, this has provided us with information about the changes in average global temperatures over the past 65 million years.

Exercise 6.4 Interpretation of past environments

Sedimentary rocks can tell us a great deal about the environmental conditions that existed during the time of their formation. Make some inferences about the source rock, weathering environment, type and distance of sediment transportation, and deposition conditions that existed during the formation of the following rocks:

  1. Quartz sandstone: no feldspar, well-sorted and well-rounded quartz grains, cross-bedding
  2. Feldspathic sandstone and mudstone: feldspar, volcanic fragments, angular grains, repetitive graded bedding from sandstone upwards to mudstone
  3. Conglomerate: well-rounded pebbles and cobbles of granite and basalt; imbrication
  4. Breccia: poorly sorted, angular limestone fragments; orange-red matrix

See Appendix 3 for Exercise 6.4 answers.

Media Attributions


6.5 Groups, Formations, and Members

Geologists who study sedimentary rocks need ways to divide them into manageable units, and they also need to give those units names so that they can easily be referred to and compared with other rocks deposited in other places. The International Commission on Stratigraphy (ICS) has established a set of conventions for grouping, describing, and naming sedimentary rock units.

The main stratigraphic unit is a formation, which according to the ICS, should be established with the following principles in mind:

The contrast in lithology between formations required to justify their establishment varies with the complexity of the geology of a region and the detail needed for geologic mapping and to work out its geologic history. No formation is considered justifiable and useful that cannot be delineated at the scale of geologic mapping practiced in the region. The thickness of formations may range from less than a meter to several thousand meters.

In other words, a formation is a series of beds that is distinct from other beds above and below, and is thick enough to be shown on the geological maps that are widely used within the area in question. In most parts of the world, geological mapping is done at a relatively coarse scale, and so most formations are in the order of a few hundred metres thick. At that thickness, a typical formation would appear on a typical geological map as an area that is at least a few millimetres thick.

A series of formations can be classified together to define a group, which could be as much as a few thousand metres thick, and represents a series of rocks that were deposited within a single basin (or a series of related and adjacent basins) over a few million to a few tens of millions of years.

In areas where detailed geological information is needed (for example, within a mining or petroleum district) a formation might be divided into members, where each member has a specific and distinctive lithology. For example, a formation that includes both shale and sandstone might be divided into members, each of which is either shale or sandstone. In some areas, where particular detail is needed, members may be divided into beds, but this is only applicable to beds that have a special geological significance. Groups, formations, and members are typically named for the area where they are found.

Figure 6.5.1 The distribution of the Upper Cretaceous Nanaimo Group rocks on Vancouver Island, the Gulf Islands, and in the Vancouver area. [Image Description]

The sedimentary rocks of the Nanaimo Group provide a useful example for understanding groups, formations, and members. During the latter part of the Cretaceous Period, from about 90 Ma to 65 Ma, a thick sequence of clastic rocks was deposited in a foreland basin between what is now Vancouver Island and the B.C. mainland (Figure 6.5.1). The Nanaimo Group strata comprise a 5000-metre-thick sequence of conglomerate, sandstone, and mudstone layers. Coal was mined from Nanaimo Group rocks from around 1850 to 1950 in the Nanaimo region, and even more recently in the Campbell River area.

The Nanaimo Group is divided into 11 formations as described in Table 6.6.  In general, the boundaries between formations are based on major lithological differences. As can be seen in the far-right column of Table 6.6, a wide range of depositional environments existed during the accumulation of the Nanaimo Group rocks, from nearshore marine for the Comox and Haslam Formation, to fluvial and deltaic with backwater swampy environments for the coal-bearing Extension, Pender, and Protection Formations, to a deep-water submarine fan environment for the upper six formations.
Table 6.6 The formations of the Nanaimo Group[Based on data in Mustard, P., 1994, The Upper Cretaceous Nanaimo Group, Georgia Basin, in J. Monger (ed) Geology and Geological Hazards of the Vancouver Region, Geol. Survey of Canada, Bull. 481, p. 27-95.]
[Skip Table 6.6]
Approximate Age (Ma) Formation name Lithologies Depositional Environment
65 to 66 Gabriola Sandstone with minor mudstone Submarine fan, high energy
66 to 67 Spray (Fine grained) Mudstone/sandstone turbidites Submarine fan, low energy
67 to 68 Geoffrey Sandstone and conglomerate Submarine fan, high energy
68 to 70 Northumberland (Fine grained) Mudstone turbidites Submarine fan, low energy
70 De Courcy Sandstone Submarine fan, high energy
70 to 72 Cedar District (Fine grained) Mudstone turbidites Submarine fan, low energy
72 to 75 Protection Sandstone and minor coal Nearshore marine and onshore deltaic and fluvial
75 to 80 Pender Sandstone and minor coal Nearshore marine and onshore deltaic and fluvial
80 Extension Conglomerate, with minor sandstone and some coal Nearshore marine and onshore deltaic and fluvial
80 to 85 Haslam (Fine grained) Mudstone and siltstone Shallow marine
85 to 90 Comox Conglomerate, sandstone, mudstone (coal in the Campbell River area) Nearshore fluvial and marine

In tables like this one, the layers are always listed in order, with the oldest at the bottom and the youngest at the top.

The five lower formations of the Nanaimo Group are all exposed in the Nanaimo area, and were well studied during the coal mining era between 1850 and 1950. All of these formations (except Haslam) have been divided into members, as that was useful for understanding the rocks in the areas where coal mining was taking place.  Within some of those members even some individual beds have been named if they were of specific importance to the mining industry.

Although there is a great deal of variety in the Nanaimo Group rocks, and it would take hundreds of photographs to illustrate all of the different types of rocks, a few representative examples are provided in Figure 6.5.2.

Alternating layers of light and dark rock of various textures
Figure 6.5.2a Representative photos of Nanaimo Group rocks. Turbidite layers in the Spray Formation on Gabriola Island. Each turbidite set consists of a lower sandstone layer (light colour) that grades upward into siltstone, and then into mudstone. (See Figure 6.4.2 for detail.)
A tall, light coloured rock wall with a black streak running through the middle.
Figure 6.5.2b Two separate layers of fluvial sandstone with a thin (approx. 75 centimetres) coal seam in between. Pender Formation in Nanaimo.

Looks like smooth black rocks stuck into cement and worn flat.
Figure 6.5.2c Comox Formation conglomerate at the very base of the Nanaimo Group in Nanaimo. The metal object is the end of a rock hammer that is 3 centimetres wide. Almost all of the clasts in this view are well-rounded basalt pebbles cobbles eroded from the Triassic Karmutsen Formation which makes up a major part of Vancouver Island.

Image Descriptions

Figure 6.5.1 image description: A map showing that Nanaimo Group rocks are present along the east coast of Vancouver Island from Nanaimo to Campbell River, farther inland in areas around Port Alberni and Duncan, on much of the Gulf Islands and a bit in the Vancouver area. [Return to Figure 6.5.1]

Media Attributions

  • Figure 6.5.1: Redrawn based on Mustard, P., 1994, The Upper Cretaceous Nanaimo Group, Georgia Basin, in J. Monger (ed) Geology and Geological Hazards of the Vancouver Region, Geol. Survey of Canada, Bull. 481, pp. 27-95. © Steven Earle. CC BY.
  • Figure 6.5.2abc: © Steven Earle. CC BY.



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

Section Summary
6.1 Clastic Sedimentary Rocks Sedimentary clasts are classified based on their size, and variations in clast size have important implications for transportation and deposition. Clastic sedimentary rocks range from conglomerate to mudstone. Clast size, sorting, composition, and shape are important features that allow us to differentiate clastic rocks and understand the processes that took place during their deposition.
6.2 Chemical Sedimentary Rocks Chemical sedimentary rocks form from ions that were transported in solution, and then converted into minerals by biological and/or chemical processes. The most common chemical rock, limestone, typically forms in shallow tropical environments, where biological activity is a very important factor. Chert and banded iron formation are deep-ocean sedimentary rocks. Evaporites form where the water of lakes and inland seas becomes supersaturated due to evaporation.
6.3 Depositional Environments and Sedimentary Basins There is a wide range of depositional environments, both on land (glaciers, lakes, rivers, etc.) and in the ocean (deltas, reefs, shelves, and the deep-ocean floor). In order to be preserved, sediments must accumulate in long-lasting sedimentary basins, most of which form through plate-tectonic processes.
6.4 Sedimentary Structures and Fossils The deposition of sedimentary rocks takes place according to a series of important principles, including original horizontality, superposition, and faunal succession. Sedimentary rocks can also have distinctive structures, such as cross bedding, graded bedding and mud cracks, that are important in determining their depositional environments. Fossils are useful for determining the age of a rock, the depositional environment, and the climate at the time of deposition.
6.5 Groups, Formations, and Members Sedimentary sequences are classified into groups, formations, and members so that they can be referred to easily and without confusion.

Questions for Review

Answers to Review Questions at the end of each chapter can be found in Appendix 2.

  1. What are the minimum and maximum sizes of sand grains?
  2. How can you easily distinguish between a silty deposit and one that has only clay-sized material?
  3. What factors control the rate at which a clast settles in water?
  4. The material that makes up a rock such as conglomerate cannot be deposited by a slow-flowing river. Why not?
  5. Describe the two main processes of lithification.
  6. What is the difference between a lithic arenite and a lithic wacke?
  7. How does a feldspathic arenite differ from a quartz arenite?
  8. What can we say about the source area lithology and the weathering and transportation history of a sandstone that is primarily composed of rounded quartz grains?
  9. What is the original source of the carbon that is present within carbonate deposits such as limestone?
  10. What long-term environmental change on Earth led to the deposition of banded iron formations?
  11. Name two important terrestrial depositional environments and two important marine ones.
  12. What is the origin of a foreland basin, and how does it differ from a forearc basin?
  13. Explain the origin of  (a) bedding, (b) cross-bedding, (c) graded bedding, and (d) mud cracks.
  14. Under what conditions is reverse graded bedding likely to form?
  15. What are the criteria for the application of a formation name to a series of sedimentary rocks?
  16. Explain why some of the Nanaimo Group formations have been divided into members, while others have not.



Chapter 7 Metamorphism and Metamorphic Rocks

Learning Objectives

After carefully reading this chapter, completing the exercises within it, and answering the questions at the end, you should be able to:

  • Summarize the factors that influence the nature of metamorphic rocks and explain why each one is important.
  • Describe foliation and explain the mechanisms for its formation in metamorphic rocks.
  • Classify metamorphic rocks on the basis of their texture and mineral content, and explain the origins of these differences.
  • Describe the various settings in which metamorphic rocks are formed and the links between plate tectonics and metamorphism.
  • Summarize the important processes of regional metamorphism, and explain how rocks that were metamorphosed at depths of 10 kilometres or 20 kilometres can now be found on Earth’s surface.
  • Describe the important processes of contact metamorphism and metasomatism, and the key role hydrothermal fluids.

Metamorphism is the change that takes place within a body of rock as a result of it being subjected to conditions that are different from those in which it formed. In most cases—but not all—this involves the rock being deeply buried beneath other rocks, where it is subjected to higher temperatures and pressures than those under which it formed. Metamorphic rocks typically have different mineral assemblages and different textures from their parent rocks (Figure 7.0.1) but they may have the same overall chemical composition.

Figure 7.0.1 Metamorphic rock (gneiss) of the Okanagan Metamorphic and Igneous Complex at Skaha Lake, B.C. The dark bands are amphibole-rich, the light bands are feldspar-rich.

Most metamorphism results from the burial of igneous, sedimentary, or pre-existing metamorphic rocks to the point where they experience different pressures and temperatures than those at which they formed (Figure 7.0.2). Metamorphism can also take place if cold rock near the surface is intruded and heated by a hot igneous body. Although most metamorphism involves temperatures above 150°C, some metamorphism takes place at temperatures lower than those at which the parent rock formed.

Figure 7.0.2 The rock cycle, showing the processes related to metamorphic rocks at the bottom. [Image description]

Image Descriptions

Figure 7.0.2 image description: As sedimentary rock (or igneous rock) gets buried deeper and comes under increased heat and pressure, it can turn into metamorphic rock. That rock may be returned to surface for us to see, but if it gets buried deeper still it may partially melt to become magma. [Return to Figure 7.0.2]

Media Attributions


7.1 Controls Over Metamorphic Processes

The main factors that control metamorphic processes are:

Parent Rock

The parent rock is the rock that exists before metamorphism starts. Sedimentary or igneous rocks can be considered the parent rocks for metamorphic rocks.  Although an existing metamorphic rock can be further metamorphosed or re-metamorphosed, metamorphic rock doesn’t normally qualify as a “parent rock”.  For example, if a mudstone is metamorphosed to slate and then buried deeper where it is metamorphosed to schist, the parent rock of the schist is mudstone, not slate. The critical feature of the parent rock is its mineral composition because it is the stability of minerals that counts when metamorphism takes place. In other words, when a rock is subjected to increased temperatures, certain minerals may become unstable and start to recrystallize into new minerals.


The temperature that the rock is subjected to is a key variable in controlling the type of metamorphism that takes place. As we learned in the context of igneous rocks, mineral stability is a function of temperature, pressure, and the presence of fluids (especially water). All minerals are stable over a specific range of temperatures. For example, quartz is stable from environmental temperatures (whatever the weather can throw at it) all the way up to about 1800°C. If the pressure is higher, that upper limit will be even higher. If there is water present, it will be lower. On the other hand, most clay minerals are only stable up to about 150° or 200°C; above that, they transform into micas. Most feldspars are stable up to between 1000°C and 1200°C.  Most other common minerals have upper limits between 150°C and 1000°C.

Some minerals will crystallize into different polymorphs (same composition, but different crystalline structure) depending on the temperature and pressure. The minerals kyanite, andalusite, and sillimanite are polymorphs with the composition Al2SiO5. They are stable at different pressures and temperatures, and, as we will see later, they are important indicators of the pressures and temperatures that existed during the formation of metamorphic rocks (Figure 7.1.1).

Figure 7.1.1 The temperature and pressure stability fields of the three polymorphs of Al2SiO5 (Pressure is equivalent to depth. Kyanite is stable at low to moderate temperatures and low to high pressures, andalusite at moderate temperatures and low pressures, and sillimanite at higher temperatures.) [Image Description]


Pressure is important in metamorphic processes for two main reasons. First, it has implications for mineral stability (Figure 7.1.1). Second, it has implications for the texture of metamorphic rocks. Rocks that are subjected to very high confining pressures are typically denser than others because the mineral grains are squeezed together (Figure 7.1.2a), and also because they may contain minerals that have greater density because the atoms are more closely packed.

Because of plate tectonics, pressures within the crust are typically not applied equally in all directions. In areas of plate convergence, for example, the pressure in one direction (perpendicular to the direction of convergence) is typically greater than in the other directions (Figure 7.1.2b). In situations where different blocks of the crust are being pushed in different directions, the rocks will likely be subjected to sheer stress (Figure 7.1.2c).

One of the results of directed pressure and shear stress is that rocks become foliated—meaning that they’ll have a directional fabric. Foliation a very important aspect of metamorphic rocks, and is described in more detail later in this chapter.

Figure 7.1.2 An illustration of different types of pressure on rocks. (a) confining pressure, where the pressure is essentially equal in all directions, (b) directed pressure, where the pressure form the sides is greater than that from the top and bottom, and (c) shear stress caused by different blocks of rock being pushed in different directions. (In a and b there is also pressure in and out of the page.)


Water is the main fluid present within rocks of the crust, and the only one that we’ll consider here. The presence of water is important for two main reasons. First, water facilitates the transfer of ions between minerals and within minerals, and therefore increases the rates at which metamorphic reactions take place. So, while the water doesn’t necessarily change the outcome of a metamorphic process, it speeds the process up so metamorphism might take place over a shorter time period, or metamorphic processes that might not otherwise have had time to be completed are completed.

Secondly, water, especially hot water, can have elevated concentrations of dissolved elements (ions), and therefore it is an important medium for moving certain elements around within the crust. So not only does water facilitate metamorphic reactions on a grain-to-grain basis, it also allows for the transportation of elements from one place to another. This is very important in hydrothermal processes, which are discussed toward the end of this chapter, and in the formation of mineral deposits.


Most metamorphic reactions take place at very slow rates. For example, the growth of new minerals within a rock during metamorphism has been estimated to be about 1 millimetre per million years. For this reason, it is very difficult to study metamorphic processes in a lab.

While the rate of metamorphism is slow, the tectonic processes that lead to metamorphism are also very slow, so in most cases, the chance for metamorphic reactions to be completed is high. For example, one important metamorphic setting is many kilometres deep within the roots of mountain ranges. A mountain range takes tens of millions of years to form, and tens of millions of years more to be eroded to the extent that we can see the rocks that were metamorphosed deep beneath it.

Exercise 7.1

Figure 7.1.3 Garnets in a rock. Euro coin (23 mm) is for scale.

This photo shows a sample of garnet-mica schist from the Greek island of Syros. The large reddish crystals are garnet, and the surrounding light coloured rock is dominated by muscovite mica. The Euro coin is 23 millimetres in diameter. Assume that the diameters of the garnets increased at a rate of 1 millimetre per million years.

Based on the approximate average diameter of the garnets visible, estimate how long this metamorphic process might have taken.

See Appendix 3 for Exercise 7.1 answers.

Image Descriptions

Figure 7.1.1 image description: The temperature ranges that polymorphs of Al2SiO5 are stable at at various depths. 
Depth (kilometres) Kyanite Andalusite Sillimanite
5 Less than 300°C 300 to 650°C Greater than 670°C
10 Less than 400°C 410 to 580°C Greater than 590°C
15 Less than 500°C Not stable Greater than 500°C
20 Less than 570°C Not stable Greater than 590°C
25 Less than 640°C Not stable Greater than 620°C
30 Less than 700°C Not stable Greater than 700°C

[Return to Figure 7.1.1]

Media Attributions


7.2 Classification of Metamorphic Rocks

There are two main types of metamorphic rocks: those that are foliated because they have formed in an environment with either directed pressure or shear stress, and those that are not foliated because they have formed in an environment without directed pressure or relatively near the surface with very little pressure at all. Some types of metamorphic rocks, such as quartzite and marble, which can form whether there is directed-pressure or not, do not typically exhibit foliation because their minerals (quartz and calcite respectively) do not tend to show alignment (see Figure 7.2.8).

When a rock is squeezed under directed pressure during metamorphism it is likely to be deformed, and this can result in a textural change such that the minerals appear elongated in the direction perpendicular to the main stress (Figure 7.2.1). This contributes to the formation of foliation.

Figure 7.2.1 The textural effects of squeezing during metamorphism.  In the original rock (left) there is no alignment of minerals.  In the squeezed rock (right) the minerals have been elongated in the direction perpendicular to the squeezing.

When a rock is both heated and squeezed during metamorphism, and the temperature change is enough for new minerals to form from existing ones, there is a strong tendency for new minerals to grow with their long axes perpendicular to the direction of squeezing. This is illustrated in Figure 7.2.2, where the parent rock is shale, with bedding as shown. After both heating and squeezing, new minerals have formed within the rock, generally parallel to each other, and the original bedding has been largely obliterated.

Figure 7.2.2 The textural effects of squeezing and mineral growth during regional metamorphism. The left diagram is shale with bedding slanting down to the right. The right diagram represents schist (derived from that shale), with mica crystals orientated perpendicular to the main stress direction and the original bedding no longer easily visible.

Figure 7.2.3 shows an example of this effect. This large boulder has bedding visible as dark and light bands sloping steeply down to the right. The rock also has a strong slaty foliation, which is horizontal in this view (parallel to the surface that the person is sitting on), and has developed because the rock was being squeezed during metamorphism. The rock has split from bedrock along this foliation plane, and you can see that other weaknesses are present in the same orientation.

Squeezing and heating alone (as shown in Figure 7.2.1) can contribute to foliation, but most foliation develops when new minerals are formed and are forced to grow perpendicular to the direction of greatest stress (Figure 7.2.2). This effect is especially strong if the new minerals are platy like mica or elongated like amphibole. The mineral crystals don’t have to be large to produce foliation. Slate, for example, is characterized by aligned flakes of mica that are too small to see.

Figure 7.2.3 A slate boulder on the side of Mt. Wapta in the Rockies near Field, BC. Bedding is visible as light and dark bands sloping steeply to the right (yellow arrows). Slaty cleavage is evident from the way the rock has broken (along the flat surface that the person is sitting on) and also from lines of weakness that are parallel to that same trend (red arrows).

The various types of foliated metamorphic rocks, listed in order of the grade or intensity of metamorphism and the type of foliation are: slate, phyllite, schist, and gneiss (Figure 7.2.4). As already noted, slate is formed from the low-grade metamorphism of shale, and has microscopic clay and mica crystals that have grown perpendicular to the stress. Slate tends to break into flat sheets. Phyllite is similar to slate, but has typically been heated to a higher temperature; the micas have grown larger and are visible as a sheen on the surface. Where slate is typically planar, phyllite can form in wavy layers. In the formation of schist, the temperature has been hot enough so that individual mica crystals are big enough to be visible, and other mineral crystals, such as quartz, feldspar, or garnet may also be visible. In gneiss, the minerals may have separated into bands of different colours. In the example shown in Figure 7.2.4d, the dark bands are largely amphibole while the light-coloured bands are feldspar and quartz. Most gneiss has little or no mica because it forms at temperatures higher than those under which micas are stable. Unlike slate and phyllite, which typically only form from mudrock, schist, and especially gneiss, can form from a variety of parent rocks, including mudrock, sandstone, conglomerate, and a range of both volcanic and intrusive igneous rocks.

Schist and gneiss can be named on the basis of important minerals that are present. For example a schist derived from basalt is typically rich in the mineral chlorite, so we call it chlorite schist. One derived from shale may be a muscovite-biotite schist, or just a mica schist, or if there are garnets present it might be mica-garnet schist. Similarly, a gneiss that originated as basalt and is dominated by amphibole, is an amphibole gneiss or, more accurately, an amphibolite.

Figure 7.2.4 Examples of foliated metamorphic rocks: (A) Slate, (B) Phyllite, (C) Schist, (D) Gneiss.

If a rock is buried to a great depth and encounters temperatures that are close to its melting point, it may partially melt. The resulting rock, which includes both metamorphosed and igneous material, is known as migmatite (Figure 7.2.5).

Figure 7.2.5 Migmatite from Prague, Czech Republic

As already noted, the nature of the parent rock controls the types of metamorphic rocks that can form from it under differing metamorphic conditions. The kinds of rocks that can be expected to form at different metamorphic grades from various parent rocks are listed in Table 7.1. Some rocks, such as granite, do not change much at the lower metamorphic grades because their minerals are still stable up to several hundred degrees.

Table 7.1 A rough guide to the types of metamorphic rocks that form from different parent rocks at different grades of regional metamorphism.
Parent Rock Very Low Grade (150-300°C) Low Grade (300-450°C) Medium Grade (450-550°C) High Grade (Above 550°C)
Mudrock slate phyllite schist gneiss
Granite no change no change almost no change granite gneiss
Basalt chlorite schist chlorite schist amphibolite amphibolite
Sandstone no change little change quartzite quartzite
Limestone little change marble marble marble

Metamorphic rocks that form under either low-pressure conditions or just confining pressure do not become foliated. In most cases, this is because they are not buried deeply, and the heat for the metamorphism comes from a body of magma that has moved into the upper part of the crust. This is contact metamorphism. Some examples of non-foliated metamorphic rocks are marble, quartzite, and hornfels.

Marble is metamorphosed limestone. When it forms, the calcite crystals tend to grow larger, and any sedimentary textures and fossils that might have been present are destroyed. If the original limestone was pure calcite, then the marble will likely be white (as in Figure 7.2.6), but if it had various impurities, such as clay, silica, or magnesium, the marble could be “marbled” in appearance.  Marble that forms during regional metamorphism—and in fact that includes most marble—may or may not develop a foliated texture, but foliation is typically not easy to see in marble.

Figure 7.2.6 Marble with visible calcite crystals (left) and an outcrop of banded marble (right).

Quartzite is metamorphosed sandstone (Figure 7.2.7). It is dominated by quartz, and in many cases, the original quartz grains of the sandstone are welded together with additional silica. Most sandstone contains some clay minerals and may also include other minerals such as feldspar or fragments of rock, so most quartzite has some impurities with the quartz.

Figure 7.2.7 Quartzite from the Rocky Mountains, found in the Bow River at Cochrane, Alberta.

Even if formed during regional metamorphism, quartzite (like marble) does not tend to look foliated because quartz crystals don’t align with the directional pressure. On the other hand, any clay present in the original sandstone is likely to be converted to mica during metamorphism, and any such mica is likely to align with the directional pressure. An example of this is shown in Figure 7.2.8. The quartz crystals show no alignment, but the micas are all aligned, indicating that there was directional pressure during regional metamorphism of this rock. Since these micas are very small, this rock would not appear to be foliated to the naked eye.

Figure 7.2.8 Magnified thin section of quartzite in polarized light. The irregular-shaped white, grey, and black crystals are all quartz. The small, thin, brightly coloured crystals are mica. This rock is foliated, even though it might not appear to be if examined without a microscope, and so it must have formed under directed-pressure conditions.

Hornfels is another non-foliated metamorphic rock that normally forms during contact metamorphism of fine-grained rocks like mudstone or volcanic rock (Figure 7.2.9). In some cases, hornfels has visible crystals of minerals like biotite or andalusite. If the hornfels formed in a situation without directed pressure, then these minerals would be randomly orientated, not aligned with one-another, as they would be if formed with directed pressure.

Figure 7.2.9 Hornfels from the Novosibirsk region of Russia. The dark and light bands are bedding. The rock has been recrystallized during contact metamorphism and does not display foliation. (scale in centimetres).

Exercise 7.2 Naming metamorphic rocks

Provide reasonable names for the following metamorphic rocks based on the description:

  1. A rock with visible crystals of mica and with small crystals of andalusite. The mica crystals are consistently parallel to one another.
  2. A very hard rock with a granular appearance and a glassy lustre. There is no evidence of foliation.
  3. A fine-grained rock that splits into wavy sheets. The surfaces of the sheets have a sheen to them.
  4. A rock that is dominated by aligned crystals of amphibole.

See Appendix 3 for Exercise 7.2 answers.

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7.3 Plate Tectonics and Metamorphism

All of the important processes of metamorphism that we are familiar with can be understood in the context of geological processes related to plate tectonics. The relationships between plate tectonics and metamorphism are summarized in Figure 7.3.1, and in more detail in Figures 7.3.2, 7.3.3, 7.3.4, and 7.3.6.

Figure 7.3.1 Environments of metamorphism in the context of plate tectonics: (a) regional metamorphism related to mountain building at a continent-continent convergent boundary, (b) regional metamorphism of oceanic crust in the area on either side of a spreading ridge, (c) regional metamorphism of oceanic crustal rocks within a subduction zone, (d) contact metamorphism adjacent to a magma body at a high level in the crust, and (e) regional metamorphism related to mountain building at a convergent boundary.

Most regional metamorphism takes place within the continental crust. While rocks can be metamorphosed at depth in most areas, the potential for metamorphism is greatest in the roots of mountain ranges where there is a strong likelihood for burial of relatively young sedimentary rock to great depths, as depicted in Figure 7.3.2. An example would be the Himalayan Range. At this continent-continent convergent boundary, sedimentary rocks have been both thrust up to great heights (nearly 9,000 metres above sea level) and also buried to great depths. Considering that the normal geothermal gradient (the rate of increase in temperature with depth) is around 30°C per kilometre, rock buried to 9 kilometres below sea level in this situation could be close to 18 kilometres below the surface of the ground, and it is reasonable to expect temperatures up to 500°C. Metamorphic rocks formed there are likely to be foliated because of the strong directional pressure (compression) of converging plates.

Figure 7.3.2 Regional metamorphism beneath a mountain range related to continent-continent collision (typical geothermal gradient). (Example: Himalayan Range) [Image Description]

At an oceanic spreading ridge, recently formed oceanic crust of gabbro and basalt is slowly moving away from the plate boundary (Figure 7.3.3). Water within the crust is forced to rise in the area close to the source of volcanic heat, and this draws more water in from farther out, which eventually creates a convective system where cold seawater is drawn into the crust and then out again onto the sea floor near the ridge. The passage of this water through the oceanic crust at 200° to 300°C promotes metamorphic reactions that change the original pyroxene in the rock to chlorite and serpentine. Because this metamorphism takes place at temperatures well below the temperature at which the rock originally formed (~1200°C), it is known as retrograde metamorphism. The rock that forms in this way is known as greenstone if it isn’t foliated, or greenschist if it is. Chlorite ((Mg5Al)(AlSi3)O10(OH)8) and serpentine ((Mg, Fe)3Si2O5(OH)4) are both “hydrated minerals” meaning that they have water (as OH) in their chemical formulas. When metamorphosed ocean crust is later subducted, the chlorite and serpentine are converted into new non-hydrous minerals (e.g., garnet and pyroxene) and the water that is released migrates into the overlying mantle, where it contributes to flux melting (Chapter 3, section 3.2).

Figure 7.3.3 Regional metamorphism of oceanic crustal rock on either side of a spreading ridge. The dotted rectangles are the areas where metamorphism is taking place.  (Example: Juan de Fuca spreading ridge)

At a subduction zone, oceanic crust is forced down into the hot mantle. But because the oceanic crust is now relatively cool, especially along its sea-floor upper surface, it does not heat up quickly, and the subducting rock remains several hundreds of degrees cooler than the surrounding mantle (Figure 7.3.4). A special type of metamorphism takes place under these very high-pressure but relatively low-temperature conditions, producing an amphibole mineral known as glaucophane (Na2(Mg3Al2)Si8O22(OH)2), which is blue in colour, and is an important component of a rock known as blueschist.

You’ve probably never seen or even heard of blueschist; that’s not surprising. What is a little surprising is that anyone has seen it! Most blueschist forms in subduction zones, continues to be subducted, turns into eclogite at about 35 kilometres depth, and then eventually sinks deep into the mantle—never to be seen again because that rock will eventually melt. In only a few places in the world, where the subduction process has been interrupted by some other tectonic process, has partially subducted blueschist rock returned to the surface. One such place is the area around San Francisco; the rock is known as the Franciscan Complex (Figure 7.3.5).

Figure 7.3.4  Regional metamorphism of oceanic crust at a subduction zone. (Example: Cascadia subduction zone. Rock of this type is exposed in the San Francisco area.)
Figure 7.3.5 Franciscan Complex blueschist rock exposed north of San Francisco. The blue colour of rock is due to the presence of the amphibole mineral glaucophane.

Magma is produced at convergent boundaries and rises toward the surface, where it can form magma bodies in the upper part of the crust. Such magma bodies, at temperatures of around 1000°C, heat up the surrounding rock, leading to contact metamorphism (Figure 7.3.6). Because this happens at relatively shallow depths, in the absence of directed pressure, the resulting rock does not normally develop foliation. The zone of contact metamorphism around an intrusion is very small (typically metres to tens of metres) compared with the extent of regional metamorphism in other settings (tens of thousands of square kilometres).

Figure 7.3.6 d: Contact metamorphism around a high-level crustal magma chamber (Example: the magma chamber beneath Mt. St. Helens.)  e: Regional metamorphism in a volcanic-arc related mountain range (volcanic-region temperature gradient) (Example: The southern part of the Coast Range, B.C.)

Regional metamorphism also takes place within volcanic-arc mountain ranges, and because of the extra heat associated with the volcanism, the geothermal gradient is typically a little steeper in these settings (somewhere between 40° and 50°C per kilometre). As a result higher grades of metamorphism can take place closer to surface than is the case in other areas (Figure 7.3.6).

Another way to understand metamorphism is by using a diagram that shows temperature on one axis and depth—which is equivalent to pressure—on the other (Figure 7.3.7). The three heavy dotted lines on this diagram represent Earth’s geothermal gradients under different conditions. In most areas, the rate of increase in temperature with depth is 30°C per kilometre. In other words, if you go 1,000 metres down into a mine, the temperature will be roughly 30°C warmer than the average temperature at the surface. In most parts of southern Canada, the average surface temperature is about 10°C, so at a 1,000 metre depth, it will be about 40°C. That’s uncomfortably hot, so deep mines must have effective ventilation systems. This typical geothermal gradient is shown by the green dotted line in Figure 7.3.7. At a 10 kilometre depth, the temperature is about 300°C and at 20 kilometres it’s about 600°C.

In volcanic areas, the geothermal gradient is more like 40° to 50°C per kilometre, so the temperature at a 10 kilometre depth is in the 400° to 500°C range. Along subduction zones, as described above, the cold oceanic crust keeps temperatures low, so the gradient is typically less than 10°C per kilometre. The various types of metamorphism described above are represented in Figure 7.3.7 with the same letters (a through e) used in Figures 7.3.1 to 7.3.4 and 7.3.6.

Figure 7.3.7 Types of metamorphism shown in the context of depth and temperature under different conditions. The metamorphic rocks formed from mudrock under regional metamorphosis with a typical geothermal gradient are listed. The letters a through e correspond with those shown in Figures 7.3.1 to 7.3.4 and 7.3.6.

By way of example, if we look at regional metamorphism in areas with typical geothermal gradients, we can see that burial in the 5 kilometre to 10 kilometre range puts us in the zeoliteZeolites are silicate minerals that typically form during low-grade metamorphism of volcanic rocks. and clay mineral zone (see Figure 7.3.7), which is equivalent to the formation of slate. At 10 to 15 kilometres, we are in the greenschist zone (where chlorite would form in mafic volcanic rock) and very fine micas form in mudrock, to produce phyllite. At 15 to 20 kilometres, larger micas form to produce schist, and at 20 to 25 kilometres amphibole, feldspar, and quartz form to produce gneiss. Beyond a depth of 25 kilometres in this setting, we cross the partial melting line for granite (or gneiss) with water present, and so we can expect migmatite to form.

Exercise 7.3 Metamorphic rocks in areas with higher geothermal gradients

Figure 7.3.7 shows the types of rock that might form from mudrock at various points along the curve of the “typical” geothermal gradient (dotted green line). Looking at the geothermal gradient for volcanic regions (dotted yellow line in Figure 7.3.7), estimate the depths at which you would expect to find the same types of rock forming from a mudrock parent.

  1. Slate
  2. Phyllite
  3. Schist
  4. Gneiss
  5. Migmatite

See Appendix 3 for Exercise 7.3 answers.

Image Descriptions

Figure 7.3.2 image description: Regional metamorphism occurring beneath a mountain range due to continent-continent collision. The typical geothermal gradient for slate is 100°C, for phyllite 200°C, for schist 300°C, for gneiss °C, for migmatite 500°C. [Return to Figure 7.3.2]

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7.4 Regional Metamorphism

As described above, regional metamorphism occurs when rocks are buried deep in the crust. This is commonly associated with convergent plate boundaries and the formation of mountain ranges. Because burial to 10 to 20 kilometres is required, the areas affected tend to be large—thousands of square kilometres.

Rather than focusing on metamorphic rock textures (slate, schist, gneiss, etc.), geologists tend to look at specific minerals within the rocks that are indicative of different grades of metamorphism. Some common minerals in metamorphic rocks are shown in Figure 7.4.1, arranged in order of the temperature ranges over which they tend to be stable. The upper and lower limits of the ranges are intentionally vague because these limits depend on a number of different factors, such as the pressure, the amount of water present, and the overall composition of the rock.

Figure 7.4.1 Metamorphic index minerals and their approximate temperature ranges [Image Description]

The southern and southwestern parts of Nova Scotia were regionally metamorphosed during the Devonian Acadian Orogeny (around 400 Ma), when a relatively small continental block (the Meguma TerraneNo, it’s not a spelling mistake! A terrane is a distinctive block of crust that is now part of a continent, but is thought to have come from elsewhere, and was added on by plate-tectonic processes.) was pushed up against the existing eastern margin of North America. As shown in Figure 7.4.2, clastic sedimentary rocks within this terrane were variably metamorphosed, with the strongest metamorphism in the southwest (the sillimanite zone), and progressively weaker metamorphism toward the east and north. The rocks of the sillimanite zone were likely heated to over 700°C, and therefore must have buried to depths between 20 and 25 kilometres. The surrounding lower-grade rocks were not buried as deep, and the rocks within the peripheral chlorite zone were likely not buried to more than about 5 kilometres.

Figure 7.4.2 Regional metamorphic zones in the Meguma Terrane of southwestern Nova Scotia.

A probable explanation for this pattern is that the area with the highest-grade rocks was buried beneath the central part of a mountain range formed by the collision of the Meguma Terrane with North America. As is the case with all mountain ranges, the crust became thickened as the mountains grew, and it was pushed farther down into the mantle than the surrounding crust. This happens because Earth’s crust is floating on the underlying mantle—and that is known as an isostatic relationship. As the formation of mountains adds weight, the crust in that area sinks farther down into the mantle to compensate for the added weight. The likely pattern of metamorphism in this situation is shown in cross-section in Figure 7.4.3a. The mountains were eventually eroded (over tens of millions of years), allowing the crust to rebound upward, thus exposing the metamorphic rock (Figure 7.4.3b).

Figure 7.4.3a Schematic cross-section through the Meguma Terrane during the Devonian period. The crust is thickened underneath the mountain range to compensate for the added weight of the mountains above and has sunk into the mantle.
Figure 7.4.3b Schematic present-day cross-section through the Meguma Terrane. The mountains have been eroded. As they lost mass the base of the crust gradually rebounded, pushing up the core of the metamorphosed region so that the once deeply buried metamorphic zones are now exposed at surface.

The metamorphism in Nova Scotia’s Meguma Terrane is just one example of the nature of regional metamorphism. Obviously many different patterns of regional metamorphism exist, depending on the parent rocks, the geothermal gradient, the depth of burial, the pressure regime, and the amount of time available. The important point is that regional metamorphism happens only at significant depths. The greatest likelihood of attaining those depths, and then having the once-buried rocks eventually exposed at the surface, is where mountain ranges existed and have since been largely eroded away. As this happens typically at convergent plate boundaries, directed pressures can be strong, and regionally metamorphosed rocks are almost always foliated.

Exercise 7.4 Scottish metamorphic zones

Figure 7.4.4

The map shown here represents the part of western Scotland between the Great Glen Fault and the Highland Boundary Fault. The shaded areas are metamorphic rock, and the three metamorphic zones represented are garnet, chlorite, and biotite.

Label the three coloured areas of the map with the appropriate zone names (garnet, chlorite, and biotite).  Hint: refer to Figure 7.4.1 above to work out which of these zones might represent the peripheral area of low-grade metamorphism, and which might represent the core area of higher-grade metamorphism.

Indicate which part of the region was likely to have been buried the deepest during metamorphism.

British geologist George Barrow studied this area in the 1890s and was the first person anywhere to map metamorphic zones based on their mineral assemblages. This pattern of metamorphism is sometimes referred to as “Barrovian.”

See Appendix 3 for Exercise 7.4 answers.

Image Descriptions

Figure 7.4.1 image description: Approximate temperature range of metamorphic minerals: Chlorite, 50 to 450°C. Muscovite, 175 to 625°C. Biotite, 350 to 725°C. Garnet, 375 to 900°C. Andalusite, 400 to 850°C. Sillimanite, 575 to 1000°C. [Return to Figure 7.4.1]

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7.5 Contact Metamorphism and Hydrothermal Processes

Contact metamorphism takes place where a body of magma intrudes into the upper part of the crust. Any type of magma body can lead to contact metamorphism, from a thin dyke to a large stock. The type and intensity of the metamorphism, and width of the metamorphic aureole will depend on a number of factors, including the type of country rock, the temperature of the intruding body and the size of the body (Figure 7.5.1). A large intrusion will contain more thermal energy and will cool much more slowly than a small one, and therefore will provide a longer time and more heat for metamorphism. That will allow the heat to extend farther into the country rock, creating a larger aureole.

Figure 7.5.1 Schematic cross-section of the middle and upper crust showing two magma bodies. The upper body has intruded into cool unmetamorphosed rock near to the surface and has created a zone of contact metamorphism. The lower body is surrounded by rock that is already hot (and probably already metamorphosed), and so it does not have a significant metamorphic aureole.

Contact metamorphic aureoles are typically quite small, from just a few centimetres around small dykes and sills, to several 10s of metres around a large stock. As was shown in Figure 7.3.7, contact metamorphism can take place over a wide range of temperatures—from around 300° to over 800°C—and of course the type of metamorphism, and new minerals formed, will vary accordingly. The nature of the country rock (or parent rock) is also important.

A hot body of magma in the upper crust can create a very dynamic situation that may have geologically interesting and economically important implications. In the simplest cases, water does not play a big role, and the main process is transfer of heat from the pluton to the surrounding rock, creating a zone of contact metamorphism (Figure 7.5.2a). In that situation mudrock or volcanic rock will likely be metamorphosed to hornfels (Figure 7.2.9), limestone will be metamorphosed to marble (Figure 7.2.6), and sandstone to quartzite (Figure 7.2.7).  (But don’t forget that marble and quartzite can also form during regional metamorphism!)

In many cases, however, water is released from the magma body as crystallization takes place, and this water is dispersed along fractures in the country rock (Figure 7.5.2b). The water released from a magma chamber is typically rich in dissolved minerals. As this water cools, is chemically changed by the surrounding rocks, or boils because of a drop in pressure, minerals are deposited, forming veins within the fractures in the country rock. Quartz veins are common in this situation, and they might also include pyrite, hematite, calcite, and even silver and gold.

Figure 7.5.2 Depiction of metamorphism and alteration around a pluton in the upper crust. (a) Thermal metamorphism only (within the purple zone). (b) Thermal metamorphism plus veining (white) related to dispersal of magmatic fluids into the overlying rock. (c) Thermal metamorphism plus veining from magmatic fluids plus alteration and possible formation of metallic minerals (hatched yellow areas) from convection of groundwater.

Heat from the magma body will heat the surrounding groundwater, causing it to expand and then rise toward the surface. In some cases, this may initiate a convection system where groundwater circulates past the pluton. Such a system could operate for many thousands of years, resulting in the circulation of billions of litres of groundwater from the surrounding region past the pluton. Hot water circulating through the rocks can lead to significant changes in the mineralogy of the rock, including alteration of feldspars to clays, and deposition of quartz, calcite, and other minerals in fractures and other open spaces (Figure 7.5.3). As with the magmatic fluids, the nature of this circulating groundwater can also change adjacent to, or above, the pluton, resulting in deposition of other minerals, including ore minerals. Metamorphism in which much of the change is derived from fluids passing through the rock is known as metasomatism. When hot water contributes to changes in rocks, including mineral alteration and formation of veins, it is known as hydrothermal alteration.

Figure 7.5.3 Calcite veins in limestone of the Comox Formation, Nanaimo, B.C

A special type of metasomatism can take place where a hot pluton intrudes into carbonate rock such as limestone. When magmatic fluids rich in silica, calcium, magnesium, iron, and other elements flow through the carbonate rock, their chemistry can change dramatically, resulting in the deposition of minerals that would not normally exist in either the igneous rock or limestone. These include garnet, epidote (another silicate), magnetite, pyroxene, and a variety of copper and other minerals (Figure 7.5.4). This type of metamorphism is known as skarn, and again, some important types of mineral deposits can form this way.

Figure 7.5.4 A skarn rock from Mount Monzoni, Northern Italy, with recrystallized calcite (blue), garnet (brown), and pyroxene (green). The rock is 6 centimetres across.

Exercise 7.5 Contact metamorphism and metasomatism

Figure 7.5.5

This diagram shows a pluton that has intruded into a series of sedimentary rocks. What type of metamorphic rock would you expect to see at locations:

a) Mudstone? _____________________

b) Limestone? _____________________

c) Sandstone? _____________________

See Appendix 3 for Exercise 7.5 answers.

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

Section Summary
7.1 Controls Over Metamorphic Processes Metamorphism is controlled by five main factors: the composition of the parent rock, the temperature to which the rock is heated, the amount and type of pressure, the volumes and compositions of aqueous fluids that are present, and the amount of time available for metamorphic reactions to take place.
7.2 Classification of Metamorphic Rocks Metamorphic rocks are classified on the basis of texture and mineral composition. Foliation is a key feature of metamorphic rocks formed under directed pressure.  Foliated metamorphic rocks include slate, phyllite, schist, and gneiss. Metamorphic rocks formed in environments without strong directed pressure include hornfels, marble, and quartzite, although the latter two may form in high-pressure situations but not develop obvious foliated textures.
7.3 Plate Tectonics and Metamorphism Almost all metamorphism can be explained by plate-tectonic processes. Oceanic crustal rock can be metamorphosed near the spreading ridge where it was formed, but most other regional metamorphism takes place in areas where mountain ranges have been created, which are most common at convergent boundaries. Contact metamorphism takes place around magma bodies in the upper part of the crust, which are also most common above convergent boundaries.
7.4 Regional Metamorphism Geologists classify metamorphic rocks based on some key minerals—such as chlorite, garnet, andalusite, and sillimanite—that form at specific temperatures and pressures. Most regional metamorphism takes place beneath mountain ranges because the crust becomes thickened and rocks are pushed down to great depths because of the isostatic relationship between the crust and mantle. When mountains erode, those metamorphic rocks are uplifted by crustal rebound.
7.5 Contact Metamorphism and Hydrothermal Processes Contact metamorphism takes place around magma bodies that have intruded into cool rocks at high levels in the crust. Heat from the magma is transferred to the surrounding country rock, resulting in mineralogical and textural changes. Water from a cooling body of magma, or from convection of groundwater produced by the heat of the pluton, can also lead to metasomatism, hydrothermal alteration, and accumulation of valuable minerals in the surrounding rocks.

Questions for Review

Answers to Review Questions can be found in Appendix 2.

  1. What are the two main agents of metamorphism, and what are their respective roles in producing metamorphic rocks?
  2. Into what metamorphic rocks will a mudrock be transformed at very low, low, medium, and high metamorphic grades?
  3. Why doesn’t granite change very much at lower metamorphic grades?
  4. Describe the main process of foliation development in a metamorphic rock such as schist.
  5. What process contributes to metamorphism of oceanic crust at a spreading ridge?
  6. How do variations in the geothermal gradient affect the depth at which different metamorphic rocks form?
  7. Blueschist metamorphism takes place within subduction zones. What are the unique temperature and pressure characteristics of this geological setting?
  8. Rearrange the following minerals in order of increasing metamorphic grade: biotite, garnet, sillimanite, chlorite.
  9. Why does contact metamorphism not normally take place at significant depth in the crust?
  10. What is the role of magmatic fluids in metamorphism that takes place adjacent to a pluton?
  11. How does metasomatism differ from regional metamorphism?
  12. How does the presence of a hot pluton contribute to the circulation of groundwater that facilitates metasomatism and hydrothermal processes?
  13. What must be present in the country rock to produce a skarn?
  14. Two things that a geologist first considers when looking at a metamorphic rock are what the parent rock might have been, and what type of metamorphism has taken place. This can be difficult to do, even if you have the actual rock in your hand, but give it a try for the following metamorphic rocks:
    1. Chlorite schist
    2. Slate
    3. Mica-garnet schist
    4. Amphibolite
    5. Marble


Review of Minerals and Rocks

Mineral and rock review

Steven Earle

Mineral and Rock Review

Crystals of the mineral native sulphur growing on the rock basalt at an outlet of volcanic gases, Kilauea volcano, Hawaii

Now that we’ve covered minerals and all three types of rocks it’s important for you to convince yourself that you’ve got them straight in your mind.  As already noted, one of the most common mistakes that geology students make on assignments, tests and exams is to confuse minerals with rocks and then give a wrong answer when asked to name one or the other based on information provided.

In this exercise you are given a list of names of minerals and rocks and asked to determine which ones are minerals and which are rocks.  For those that you think are minerals you should then indicate which mineral group it belongs to (e.g., oxide, sulphate, silicate, carbonate, halide etc.). For those that you think are rocks, you should describe what type of rock it is (e.g., intrusive igneous, extrusive igneous, clastic sedimentary, chemical sedimentary, foliated metamorphic and non-foliated metamorphic).  The answers can be found in Rock and mineral review exercise answers in Appendix 3.

Mineral or rock name Rock or mineral? If it’s a mineral, which group does it belong to?  If it’s a rock, what type is it?

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Chapter 8 Measuring Geological Time

Learning Objectives

After carefully reading this chapter, completing the exercises within it, and answering the questions at the end, you should be able to:

  • Apply basic geological principles to the determination of the relative ages of rocks.
  • Explain the difference between relative and absolute age-dating techniques.
  • Summarize the history of the geological time scale and the relationships between eons, eras, periods, and epochs.
  • Understand the importance and significance of unconformities.
  • Estimate the age of a rock based on the fossils that it contains.
  • Describe some applications and limitations of isotopic techniques for geological dating.
  • Use isotopic data to estimate the age of a rock.
  • Describe the techniques for dating geological materials using tree rings and magnetic data.
  • Explain why an understanding of geological time is critical to both geologists and the public in general.

Time is the dimension that sets geology apart from most other sciences. Geological time is vast, and Earth has changed enough over that time that some of the rock types that formed in the past could not form today. Furthermore, as we’ve discussed, even though most geological processes are very, very slow, the vast amount of time that has passed has allowed for the formation of extraordinary geological features, as shown in Figure 8.0.1.

Figure 8.0.1 Arizona’s Grand Canyon is an icon for geological time; 1,450 million years are represented by this photo. The light-coloured layered rocks at the top formed at around 250 Ma, and the dark rocks at the bottom (within the steep canyon) at around 1,700 Ma.

We have numerous ways of measuring geological time. We can tell the relative ages of rocks (for example, whether one rock is older than another) based on their spatial relationships; we can use fossils to date sedimentary rocks because we have a detailed record of the evolution of life on Earth; and we can use a range of isotopic techniques to determine the actual ages (in millions of years) of igneous and metamorphic rocks.

But just because we can measure geological time doesn’t mean that we understand it. One of the biggest hurdles faced by geology students—and geologists as well—in mastering geology, is to really come to grips with the slow rates at which geological processes happen and the vast amount of time involved.

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8.1 The Geological Time Scale

William Smith worked as a surveyor in the coal-mining and canal-building industries in southwestern England in the late 1700s and early 1800s. While doing his work, he had many opportunities to look at the Paleozoic and Mesozoic sedimentary rocks of the region, and he did so in a way that few had done before. Smith noticed the textural similarities and differences between rocks in different locations, and more importantly, he discovered that fossils could be used to correlate rocks of the same age. Smith is credited with formulating the principle of faunal succession (the concept that specific types of organisms lived during different time intervals), and he used it to great effect in his monumental project to create a geological map of England and Wales, published in 1815. For more on William Smith, including a large-scale digital copy of the famous map, see the William Smith Wikipedia page.

Inset into Smith’s great geological map is a small diagram showing a schematic geological cross-section extending from the Thames estuary of eastern England all the way to the west coast of Wales. Smith shows the sequence of rocks, from the Paleozoic rocks of Wales and western England, through the Mesozoic rocks of central England, to the Cenozoic rocks of the area around London (Figure 8.1.1). Although Smith did not put any dates on these—because he didn’t know them—he was aware of the principle of superposition (the idea, developed much earlier by the Danish theologian and scientist Nicholas Steno, that young sedimentary rocks form on top of older ones), and so he knew that this diagram represented a stratigraphic column. And because almost every period of the Phanerozoic is represented along that section through Wales and England, it is a primitive geological time scale.

Figure 8.1.1 William Smith’s “Sketch of the succession of strata and their relative altitudes,” an inset on his geological map of England and Wales (with era names added).

Smith’s work set the stage for the naming and ordering of the geological periods, which was initiated around 1820, first by British geologists, and later by other European geologists. Many of the periods are named for places where rocks of that age are found in Europe, such as Cambrian for Cambria (Wales), Devonian for Devon in England, Jurassic for the Jura Mountains in France and Switzerland, and Permian for the Perm region of Russia. Some are named for the type of rock that is common during that age, such as Carboniferous for the coal- and carbonate-bearing rocks of England, and Cretaceous for the chalks of England and France.

The early time scales were only relative because 19th century geologists did not know the ages of the rocks. That information was not available until the development of isotopic dating techniques early in the 20th century.

The geological time scale is currently maintained by the International Commission on Stratigraphy (ICS), which is part of the International Union of Geological Sciences. The time scale is continuously being updated as we learn more about the timing and nature of past geological events. You can view the ICS time scale online. It would be a good idea to print a copy (in colour) to put on your wall while you are studying geology.

Geological time has been divided into four eons: Hadean (4570 to 4850 Ma), Archean (3850 to 2500 Ma), Proterozoic (2500 to 540 Ma), and Phanerozoic (540 Ma to present). As shown in Figure 8.1.2, the first three of these represent almost 90% of Earth’s history. The last one, the Phanerozoic (meaning “visible life”), is the time that we are most familiar with because Phanerozoic rocks are the most common on Earth, and they contain evidence of the life forms that we are familiar with to varying degrees.

Figure 8.1.2 The four eons of Earth’s history.

The Phanerozoic eon—the past 540 Ma of Earth’s history—is divided into three eras: the Paleozoic (“early life”), the Mesozoic (“middle life”), and the Cenozoic (“new life”), and each of these is divided into a number of periods (Figure 8.1.3). Most of the organisms that we share Earth with evolved at various times during the Phanerozoic.

Figure 8.1.3 The eras (middle row) and periods (bottom row) of the Phanerozoic eon. [Image Description]

The Cenozoic era, which represents the past 65.5 Ma, is divided into three periods: Paleogene, Neogene, and Quaternary, and seven epochs (Figure 8.1.4). Dinosaurs became extinct at the start of the Cenozoic, after which birds and mammals radiated to fill the available habitats. Earth was very warm during the early Eocene and has steadily cooled ever since. Glaciers first appeared on Antarctica in the Oligocene and then on Greenland in the Miocene, and covered much of North America and Europe by the Pleistocene. The most recent of the Pleistocene glaciations ended around 11,700 years ago. The current epoch is known as the Holocene. Epochs are further divided into ages (a.k.a. stages), but we won’t be going into that level of detail here.

Figure 8.1.4 The periods (middle row) and epochs (bottom row) of the Cenozoic era. [Image Description]

Most of the boundaries between the periods and epochs of the geological time scale have been fixed on the basis of significant changes in the fossil record. For example, as already noted, the boundary between the Cretaceous and the Paleogene coincides exactly with a devastating mass extinction. That’s not a coincidence. The dinosaurs and many other types of organisms went extinct at this time, and the boundary between the two periods marks the division between sedimentary rocks with Cretaceous organisms (including dinosaurs) below, and Paleogene organisms above.

Image Descriptions

Figure 8.1.3 image description: The eras and periods that make up the Phanerozoic Eon.
Era Period Time span
Paleozoic Cambrian 488 to 540 Ma
Paleozoic Ordovician 488 to 444 Ma
Paleozoic Silurian 444 to 416 Ma
Paleozoic Devonian 416 to 359 Ma
Paleozoic Carboniferous 359 to 299 Ma
Paleozoic Permian 299 to 251 Ma
Mesozoic Triassic 251 to 202 Ma
Mesozoic Jurassic 202 to 146 Ma
Mesozoic Cretaceous 146 to 65.5 Ma
Cenozoic Paleogene 65.5 to 23 Ma
Cenozoic Neogene 23 to 2.6 Ma
Cenozoic Quaternary 2.6 Ma to present

[Return to Figure 8.1.3]

Figure 8.1.4 image description: The periods and epochs that make up the Cenozoic era.
Period Epoch Time span
Paleogene Paleocene 65.5 to 55.8 Ma
Paleogene Eocene 55.8 to 33.9 Ma
Paleogene Oligocene 33.9 to 23.0 Ma
Neogene Miocene 23.0 to 5.3 Ma
Neogene Pliocene 5.3 to 2.6 Ma
Quaternary Pleistocene 2.6 Ma to 11,700 years ago
Quaternary Holocene 11,700 years ago to the present

[Return to Figure 8.1.4]

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8.2 Relative Dating Methods

The simplest and most intuitive way of dating geological features is to look at the relationships between them. There are a few simple rules for doing this, some of which we’ve already looked at in Chapter 6. For example, the principle of superposition states that sedimentary layers are deposited in sequence, and, unless the entire sequence has been turned over by tectonic processes or disrupted by faulting, the layers at the bottom are older than those at the top. The principle of inclusions states that any rock fragments that are included in rock must be older than the rock in which they are included. For example, a xenolith in an igneous rock or a clast in sedimentary rock must be older than the rock that includes it (Figure 8.2.1).

Figure 8.2.1a A xenolith of diorite incorporated into a basalt lava flow, Mauna Kea volcano, Hawaii. The lava flow took place some time after the diorite cooled, was uplifted, and then eroded. (geological rock hammer head for scale).
Figure 8.2.1b Rip-up clasts of shale embedded in Gabriola Formation sandstone, Gabriola Island, B.C. The pieces of shale were eroded as the sandstone was deposited, so the shale is older than the sandstone.

The principle of cross-cutting relationships states that any geological feature that cuts across, or disrupts another feature must be younger than the feature that is disrupted. An example of this is given in Figure 8.2.2, which shows three different sedimentary layers. The lower sandstone layer is disrupted by two faults, so we can conclude that the faults are younger than that layer. But the faults do not appear to continue into the coal seam, and they certainly do not continue into the upper sandstone. So we can infer that coal seam is younger than the faults (because it cuts them off), and of course the upper sandstone is youngest of all, because it lies on top of the coal seam.

Figure 8.2.2 Superposition and cross-cutting relationships in Cretaceous Nanaimo Group rocks in Nanaimo, B.C. The coal seam is about 50 centimetres thick. The sequence of events is as follows:  a) deposition of lower sandstone, b) faulting of lower sandstone, c) deposition of coal seam and d) deposition of upper sandstone.

Exercise 8.1 Cross-Cutting Relationships

Figure 8.2.3

The outcrop shown here (at Horseshoe Bay, B.C.) has three main rock types:

  1. Buff/pink felsic intrusive igneous rock present as somewhat irregular masses trending from lower right to upper left
  2. Dark grey metamorphosed basalt
  3. A 50 centimetres wide light-grey felsic intrusive igneous dyke extending from the lower left to the middle right – offset in several places

Using the principle of cross-cutting relationships outlined above, determine the relative ages of these three rock types.

(The near-vertical stripes are blasting drill holes. The image is about 7 metres across.)

See Appendix 3 for Exercise 8.1 answers.

An unconformity represents an interruption in the process of deposition of sedimentary rocks. Recognizing unconformities is important for understanding time relationships in sedimentary sequences. An example of an unconformity is shown in Figure 8.2.4. The Proterozoic rocks of the Grand Canyon Group have been tilted and then eroded to a flat surface prior to deposition of the younger Paleozoic rocks. The difference in time between the youngest of the Proterozoic rocks and the oldest of the Paleozoic rocks is close to 300 million years. Tilting and erosion of the older rocks took place during this time, and if there was any deposition going on in this area, the evidence of it is now gone.

Figure 8.2.4 The great angular unconformity in the Grand Canyon, Arizona. The tilted rocks at the bottom are part of the Proterozoic Grand Canyon Group (aged 825 to 1,250 Ma). The flat-lying rocks at the top are Paleozoic (540 to 250 Ma). The boundary between the two represents a time gap of nearly 300 million years.

There are four types of unconformities, as summarized in Table 8.1, and illustrated in Figure 8.2.5.

Table 8.1 The characteristics of the four types of unconformities
Unconformity Type Description
Nonconformity A boundary between non-sedimentary rocks (below) and sedimentary rocks (above)
Angular unconformity A boundary between two sequences of sedimentary rocks where the underlying ones have been tilted (or folded) and eroded prior to the deposition of the younger ones (as in Figure 8.2.4)
Disconformity A boundary between two sequences of sedimentary rocks where the underlying ones have been eroded (but not tilted) prior to the deposition of the younger ones (as in Figure 8.2.2)
Paraconformity A time gap in a sequence of sedimentary rocks that does not show up as an angular unconformity or a disconformity
Figure 8.2.5 The four types of unconformities: (a) a nonconformity between older non-sedimentary rock and sedimentary rock, (b) an angular unconformity, (c) a disconformity between layers of sedimentary rock, where the older rock has been eroded but not tilted, and (d) a paraconformity where there is a long period (typically millions of years) of non-deposition between two parallel layers.

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8.3 Dating Rocks Using Fossils

Geologists get a wide range of information from fossils. They help us to understand evolution and life in general; they provide critical information for understanding depositional environments and changes in Earth’s climate; and, of course, they can be used to date rocks.

Although the recognition of fossils goes back hundreds of years, the systematic cataloguing and assignment of relative ages to different organisms from the distant past—paleontology—only dates back to the earliest part of the 19th century. The oldest undisputed fossils are from rocks dated around 3.5 Ga, and although fossils this old are tiny, typically poorly preserved and are not useful for dating rocks, they can still provide important information about conditions at the time. The oldest well-understood fossils are from rocks dating back to around 600 Ma, and the sedimentary record from that time forward is rich in fossil remains that provide a detailed record of the history and evolution of life on Earth. However, as anyone who has gone hunting for fossils knows, that does not mean that all sedimentary rocks have visible fossils, or that they are easy to find. Fossils alone cannot provide us with numerical ages of rocks, but over the past century geologists have acquired enough isotopic dates from rocks associated with fossil-bearing rocks (such as igneous dykes cutting through sedimentary layers, or volcanic layers between sedimentary layers) to be able to put specific time limits on most fossils.

A very selective history of life on Earth over the past 600 million years is provided in Figure 8.3.1. The major groups of organisms that we are familiar with evolved between the late Proterozoic and the Cambrian (approximately 600 to 520 Ma). Plants, which evolved in the oceans as green algae, came onto land during the Ordovician (approximately 450 Ma). Insects, which evolved from marine arthropods, came onto land during the Devonian (400 Ma), and amphibians (i.e., vertebrates) came onto land about 50 million years later. By the late Carboniferous, trees had evolved from earlier plants, and reptiles had evolved from amphibians. By the mid-Triassic, dinosaurs and mammals had evolved from very different branches of the reptiles; birds evolved from dinosaurs during the Jurassic. Flowering plants evolved in the late Jurassic or early Cretaceous. The earliest primates evolved from other mammals in the early Paleogene, and the genus Homo evolved during the late Neogene (roughly 2.8 Ma).

Figure 8.3.1 A summary of life on Earth during the late Proterozoic and the Phanerozoic. The top row shows geological eras, and the lower row shows the periods. [Image Description]

If we understand the sequence of evolution on Earth, we can apply knowledge to determining the relative ages of rocks. This is William Smith’s principle of faunal succession, although of course it doesn’t just apply to “fauna” (animals); it can also apply to fossils of plants and those of simple organisms.

The Phanerozoic has seen five major extinctions, as indicated in Figure 8.3.1. The most significant of these was at the end of the Permian, which saw the extinction of over 80% of all species and over 90% of all marine species. Most well-known types of organisms were decimated by this event, but only a few became completely extinct, including trilobites. The second most significant extinction was at the Cretaceous-Paleogene boundary (K-Pg, a.k.a. the K-T extinction). At that time, about 75% of marine species disappeared. Again, a few well-known types of organisms disappeared altogether, including the dinosaurs (but not birds) and the pterosaurs. Many other types were badly decimated by that event but survived, and then flourished in the Paleogene. The K-Pg extinction is thought to have been caused by the impact of a large extraterrestrial body (10 to 15 kilometres across), but it is generally agreed that the other four Phanerozoic extinctions had other causes, although their exact nature is not clearly understood.

As already stated, it is no coincidence that the major extinctions all coincide with boundaries of geological periods and even eras. Geologists have placed most of the divisions of the geological time scale at points in the fossil record where there are major changes in the type of organisms observed, and most of these correspond with minor or major extinctions.

Figure 8.3.2 The application of bracketing to constrain the age of a rock based on several fossils. In this diagram, the coloured bars represent the time range during which each of the four species (A, B, C, D) existed on Earth. Although each species lived for several million years, we can narrow down the likely age of the rock to just 700,000 years during which all four species coexisted.

If we can identify a fossil to the species level, or at least to the genus level, and we know the time period when the organism lived, we can assign a range of time to the rock. That range might be several million years because some organisms survived for a very long time. If the rock we are studying has several types of fossils in it, and we can assign time ranges to several of them, we might be able to narrow the time range for the age of the rock considerably. An example of this is given in Figure 8.3.2.

Some organisms survived for a very long time, and are not particularly useful for dating rocks. Sharks, for example, have been around for over 400 million years, and the great white shark has survived for 16 million years, so far. Organisms that lived for relatively short time periods are particularly useful for dating rocks, especially if they were distributed over a wide geographic area and so can be used to compare rocks from different regions. These are known as index fossils. There is no specific limit on how short the time span has to be to qualify as an index fossil. Some lived for millions of years, and others for much less than a million years.

Some well-studied groups of organisms qualify as biozone fossils because, although the genera and families lived over a long time, each species lived for a relatively short time and can be easily distinguished from others on the basis of specific features. For example, ammonites have a distinctive feature known as the suture line—where the internal shell layers that separate the individual chambers (septae) meet the outer shell wall, as shown in Figure 8.3.3. These suture lines are sufficiently variable to identify species that can be used to estimate the relative or absolute ages of the rocks in which they are found.

Figure 8.3.3 The septum of an ammonite (white part, left), and the suture lines where the septae meet the outer shell (right).

Foraminifera (small, carbonate-shelled marine organisms that originated during the Triassic and are still around today) are also useful biozone fossils. As shown in Figure 8.3.4, numerous different foraminifera lived during the Cretaceous. Some lasted for over 10 million years, but others for less than 1 million years. If the foraminifera in a rock can be identified to the species level, we can get a good idea of its age.

Figure 8.3.4 Time ranges for Cretaceous foraminifera (left).  Modern foraminifera from the Ambergris area of Belize (right).

Exercise 8.2 Dating rocks using index fossils

Figure 8.3.5 shows the age ranges for some late Cretaceous inoceramid clams in the genus Mytiloides:

Figure 8.3.5 shows the age ranges for some late Cretaceous inoceramid clams in the genus Mytiloides:

  • M. hattiru, 93.4 to 92.6 Ma
  • M. kossmati, 93.3 to 92.5 Ma
  • M. columbiarus, 93.2 to 92.5 Ma
  • M. subhercynius, 92.7 to 91.9 Ma
  • M. labiatus, 92.9 to 92.6 Ma

Using the bracketing method described above, determine the possible age range of the rock that these five organisms were found in.

How would that change if M. subhercynius was not present in these rocks?

See Appendix 3 for Exercise 8.2 answers.

Image Descriptions

Figure 8.3.1 image description: Life on earth during the late Proterozoic and the Phanerozoic.
Eon Ero Period Life on Earth
Proterozoic Ediacaran Worms, jellyfish, corals
Phanerozoic Paleozoic (540 to 251 Ma) Cambrian Brachiopods, molluscs, trilobites, primitive fish
Phanerozoic Paleozoic (540 to 251 Ma) Ordovician Land plants, the period ends with a major extinction
Phanerozoic Paleozoic (540 to 251 Ma) Silurian
Phanerozoic Paleozoic (540 to 251 Ma) Devonian Insects, amphibians, the period ends with a major extinction
Phanerozoic Paleozoic (540 to 251 Ma) Carboniferous Trees, reptiles
Phanerozoic Paleozoic (540 to 251 Ma) Permian The period ends with a major extinction
Phanerozoic Mesozoic (251 to 65 Ma) Triassic Dinosaurs, mammals, the period ends with a major extinction
Phanerozoic Mesozoic (251 to 65 Ma) Jurrassic Birds, flowering plants
Phanerozoic Mesozoic (251 to 65 Ma) Cretaceous The period ends with a major extinction
Phanerozoic Cenozoic, (65 Ma to present) Paleogene Primates
Phanerozoic Cenozoic, (65 Ma to present) Neogene Homo
Phanerozoic Cenozoic, (65 Ma to present) Quaternary  

[Return to Figure 8.3.1]

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8.4 Isotopic Dating Methods

Originally fossils only provided us with relative ages because, although early paleontologists understood biological succession, they did not know the absolute ages of the different organisms. It was only in the early part of the 20th century, when isotopic dating methods were first applied, that it became possible to discover the absolute ages of the rocks containing fossils. In most cases, we cannot use isotopic techniques to directly date fossils or the sedimentary rocks they are found in, but we can constrain their ages by dating igneous rocks that cut across sedimentary rocks, or volcanic layers that lie within sedimentary layers.

Figure 8.4.1 The decay of 40K over time. Each half-life is 1.3 billion years, so after 3.9 billion years (three half-lives) 12.5% of the original 40K will remain. The red-blue bars represent 40K and the green-yellow bars represent 40Ar. [Image Description]

Isotopic dating of rocks, or the minerals in them, is based on the fact that we know the decay rates of certain unstable isotopes of elements and that these rates have been constant over geological time. It is also based on the premise that when the atoms of an element decay within a mineral or a rock, they stay there and don’t escape to the surrounding rock, water, or air. One of the isotope pairs widely used in geology is the decay of 40K to 40Ar (potassium-40 to argon-40). 40K is a radioactive isotope of potassium that is present in very small amounts in all minerals that have potassium in them. It has a half-life of 1.3 billion years, meaning that over a period of 1.3 Ga one-half of the 40K atoms in a mineral or rock will decay to 40Ar, and over the next 1.3 Ga one-half of the remaining atoms will decay, and so on (Figure 8.4.1).

In order to use the K-Ar dating technique, we need to have an igneous or metamorphic rock that includes a potassium-bearing mineral. One good example is granite, which normally has some potassium feldspar (Figure 8.4.2). Feldspar does not have any argon in it when it forms. Over time, the 40K in the feldspar decays to 40Ar. Argon is a gas and the atoms of 40Ar remain embedded within the crystal, unless the rock is subjected to high temperatures after it forms. The sample must be analyzed using a very sensitive mass-spectrometer, which can detect the differences between the masses of atoms, and can therefore distinguish between 40K and the much more abundant 39K. Biotite and hornblende are also commonly used for K-Ar dating.

Figure 8.4.2 Crystals of potassium feldspar (pink) in a granitic rock are candidates for isotopic dating using the K-Ar method because they contained potassium and no argon when they formed.
Why can’t we use isotopic dating techniques to accurately date sedimentary rocks?
Figure 8.4.3

An important assumption that we have to be able to make when using isotopic dating is that when the rock formed none of the daughter isotope was present (e.g., 40Ar in the case of the K-Ar method). A clastic sedimentary rock is made up of older rock and mineral fragments, and when the rock forms it is almost certain that all of the fragments already have daughter isotopes in them. Furthermore, in almost all cases, the fragments have come from a range of source rocks that all formed at different times. If we dated a number of individual grains in the sedimentary rock, we would likely get a range of different dates, all older than the age of the rock.  That could be useful information, but it would not provide an accurate date for the rock in question.

It might be possible to directly date some chemical sedimentary rocks isotopically, but there are no useful isotopes that can be used on old chemical sedimentary rocks. Radiocarbon dating can be used on sediments or sedimentary rocks that contain carbon, but it cannot be used on materials older than about 60 ka.

K-Ar is just one of many isotope-pairs that are useful for dating geological materials. Some of the other important pairs are listed in Table 8.2, along with the age ranges that they apply to and some comments on their applications. When radiometric techniques are applied to metamorphic rocks, the results normally tell us the date of metamorphism, not the date when the parent rock formed.

Table 8.2 A few of the isotope systems that are widely used for dating geological materials
[Skip Table]
Isotope System Half-Life Useful Range Comments
Potassium-argon 1.3 Ga 10 Ka to 4.57 Ga Widely applicable because most rocks have some potassium
Uranium-lead 4.5 Ga 1 Ma to 4.57 Ga The rock must have uranium-bearing minerals, but most have enough.
Rubidium-strontium 47 Ga 10 Ma to 4.57 Ga Less precision than other methods at old dates
Carbon-nitrogen (a.k.a. radiocarbon dating) 5,730 years 100 to 60,000 years Sample must contain wood, bone, or carbonate minerals; can be applied to young sediments

Exercise 8.3 Isotopic dating

Assume that a feldspar crystal from the granite shown in Figure 8.4.2 was analyzed for 40K and 40Ar. The proportion of 40K remaining is 0.91. Using the decay curve shown on the graph below, estimate the age of the rock.

Figure 8.4.4 [Image Description]

An example is provided (in blue) for a 40K proportion of 0.95, which is equivalent to an age of approximately 96 Ma. This is determined by drawing a horizontal line from 0.95 to the decay curve line, and then a vertical line from there to the time axis.

See Appendix 3 for Exercise 8.3 answers.

Figure 8.4.5 Radiocarbon dates on wood fragments in glacial sediments in the Strait of Georgia.

Radiocarbon dating (using 14C) can be applied to many geological materials, including sediments and sedimentary rocks, but the materials in question must be younger than 60 ka. Fragments of wood incorporated into young sediments are good candidates for carbon dating, and this technique has been used widely in studies involving late Pleistocene glaciers and glacial sediments. An example is shown in Figure 8.4.5; radiocarbon dates from wood fragments in glacial sediments have been used to estimate the timing of the last glacial advance along the Strait of Georgia.  It is evident that the ice-front of the major glacier that occupied the Strait of Georgia was near to Campbell River at around 35 ka, near to Nanaimo and Vancouver at about 25 ka, and had reached the Victoria area by around 22 ka.

Over the past decade there has been increasing use of U-Pb dating to study sedimentary rocks, not necessarily to find out the age of the rock, but to discover something about its history and origins.  All clastic sedimentary rocks contain some tiny clasts of the silicate mineral zircon (ZrSiO4), derived from the weathering of the sediment parent rocks.  Zircon always has some uranium in it (but no lead) so it is a good candidate for U-Pb dating, and it isn’t too difficult to separate the grains of zircon from the other grains in a sandstone.  The procedure is to isolate a few hundred tiny zircons from a rock sample, and then carry out U-Pb dating on each one of them.  An example of the types of results obtained are shown on Figure 8.5.6.  All of the samples are from Nanaimo Gp. rocks on Vancouver Island and nearby Salt Spring Island.

The three samples from Vancouver Island have zircons aged around 90 Ma, 118 Ma and 150 Ma.  The Salt Spring Island sample has some zircons aged around 150 Ma, but most are much older, at 200 Ma and 340 to 360 Ma.  It is interpreted that the younger zircons (90 to 150 Ma) are mostly derived from granitic rocks in the Coast Range, while the older ones (>200 Ma) are from older rocks on Vancouver Island (Huang, 2018).

Figure 8.4.6 U-Pb dates for zircon samples from the Nanaimo Gp. (after Krause, 2018) a: a typical zircon clast (this one is about 1/4 mm long).  b: plots of zircon ages for 4 sandstone samples.

Image Descriptions

Figure 8.4.1 image description: Decay of 40K over time.
Number of half-lives Percent of 40K remaining Percent of 40Ar
0 100 0
1 50 50
2 25 75
3 12.5 87.5
4 6.25 93.75
5 3.125 96.875
6 1.5625 98.4375
7 0.78125 99.21875

[Return to Figure 8.4.1]

Figure 8.4.4 image description: isotopic dating graph
Proportion of Potassium-40 remaining Age (in millions of  Years)
1.00 0
0.99 19
0.98 37
0.97 55
0.96 75
0.95 96
0.94 114
0.93 134
0.92 156
0.91 175
0.90 194

[Return to Figure 8.4.4]

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8.5 Other Dating Methods

There are numerous other techniques for dating geological materials, but we will examine just two of them here: tree-ring dating (i.e., dendrochronology) and dating based on the record of reversals of Earth’s magnetic field.

Dendrochronology can be applied to dating very young geological materials based on reference records of tree-ring growth going back many millennia. The longest such records can take us back to 25 ka, to the height of the last glaciation. One of the advantages of dendrochronology is that, providing reliable reference records are available, the technique can be used to date events to the nearest year.

Dendrochronology has been used to date the last major subduction-zone earthquake on the coast of B.C., Washington, and Oregon. When large earthquakes strike in this setting, there is a tendency for some coastal areas to subside by one or two metres. Seawater then rushes in, flooding coastal flats and killing trees and other vegetation within a few months. There are at least four locations along the coast of Washington that have such dead trees (and probably many more in other areas). Wood samples from these trees have been studied and the ring patterns have been compared with patterns from old living trees in the region (Figure 8.5.1).

Figure 8.25.1 Example of tree-ring dating of dead trees.
Figure 8.5.2 Sites in Washington where dead trees are present in coastal flats. The outermost wood of eight trees was dated using dendrochronology, and of these, seven died during the year 1699, suggesting that the land near to the coast was inundated by water at that time.

At all of the locations studied, the trees were found to have died either in the year 1699, or very shortly thereafter (Figure 8.5.2). On the basis of these results, it was concluded that a major earthquake took place in this region sometime between the end of growing season in 1699 and the beginning of the growing season in 1700. Evidence from a major tsunami that struck Japan on January 27, 1700, narrowed the timing of the earthquake to sometime in the evening of January 26, 1700. For more information, see The 1700 Juan de Fuca Earthquake.

Changes in Earth’s magnetic field can also be used to date events in geologic history. The magnetic field makes compasses point toward the North Pole, but, as we’ll see in Chapter 10, this hasn’t always been the case. At various times in the past, Earth’s magnetic field has reversed itself completely, and during those times a compass would have pointed to the South Pole. By studying magnetism in volcanic rocks that have been dated isotopically, geologists have been able to delineate the chronology of magnetic field reversals going back to 250 Ma. About 5 million years of this record is shown in Figure 8.5.3, where the black bands represent periods of normal magnetism (“normal” meaning similar to the current magnetic field) and the white bands represent periods of reversed magnetism. These periods of consistent magnetic polarity are given names to make them easier to reference. The current normal magnetic field, known as Brunhes, has lasted for the past 780,000 years. Prior to that there was a short reversed period and then a short normal period known as Jaramillo.

Figure 8.5.3 The last 5 Ma of magnetic field reversals.

Oceanic crust becomes magnetized by the magnetic field that exists as the crust forms from magma. As it cools, tiny crystals of magnetite that form within the magma become aligned with the existing magnetic field and then remain that way after all of the rock has hardened, as shown in Figure 8.5.4. Crust that is forming today is being magnetized in a “normal” sense, but crust that formed 780,000 to 900,000 years ago, in the interval between the Brunhes and Jaramillo normal periods, was magnetized in the “reversed” sense.

Chapter 9 has a discussion of Earth’s magnetic field, including where and how it is generated and why its polarity changes periodically.

Figure 8.5.4 Depiction of the formation of magnetized oceanic crust at a spreading ridge. Coloured bars represent periods of normal magnetism, and the small capital letters denote the Brunhes, Jaramillio, Olduvai, and Gauss normal magnetic periods (see Figure 8.5.2).

Magnetic chronology can be used as a dating technique because we can measure the magnetic field of rocks using a magnetometer in a lab, or of entire regions by towing a magnetometer behind a ship or an airplane. For example, the Juan de Fuca Plate, which lies off of the west coast of B.C., Washington, and Oregon, is being and has been formed along the Juan de Fuca spreading ridge (Figure 8.5.5). The parts of the plate that are still close to the ridge have normal magnetism, while parts that are farther away (and formed much earlier) have either normal or reversed magnetism, depending on when the rock formed. By carefully matching the sea-floor magnetic stripes with the known magnetic chronology, we can determine the age at any point on the plate. We can see, for example, that the oldest part of the Juan de Fuca Plate that has not subducted (off the coast of Oregon) is just over 8 million years old, while the part that is subducting underneath Vancouver Island is between 0 and about 6 million years old.

Figure 8.5.5 The pattern of magnetism within the area of the Juan de Fuca Plate, off the west coast of North America. The coloured shapes represent parts of the sea floor that have normal magnetism, and the magnetic time scale is shown using the same colours. The blue bands represent Brunhes, Jaramillo, and Olduvai; the green represents Gauss; and so on. (Note that in this diagram, sea-floor magnetism is only shown for the Juan de Fuca Plate, although similar patterns exist on the Pacific Plate.)

Exercise 8.4 Magnetic dating

The fact that magnetic intervals can only be either normal or reversed places significant limits on the applicability of magnetic dating. If we find a rock with normal magnetism, we can’t know which normal magnetic interval it represents, unless we have some other information.

Using Figure 8.5.3 for reference, determine the age of a rock with normal magnetism that has been found to be between 1.5 and 2.0 Ma based on fossil evidence.

How about a rock that is limited to 2.6 to 3.2 Ma by fossils and has reversed magnetism?

See Appendix 3 for Exercise 8.4 answers.

Media Attributions


8.6 Understanding Geological Time

It’s one thing to know the facts about geological time—how long it is, how we measure it, how we divide it up, and what we call the various periods and epochs—but it is quite another to really understand geological time. The problem is that our lives are short and our memories are even shorter. Our experiences span only a few decades, so we really don’t have a way of knowing what 11,700 years means. What’s more, it’s hard for us to understand how 11,700 years differs from 65.5 million years, or even from 1.8 billion years. It’s not that we can’t comprehend what the numbers mean—we can all get that figured out with a bit of practice—but even if we do know the numerical meaning of 65.5 Ma, we can’t really appreciate how long ago it was.

You may be wondering why it’s so important to really “understand” geological time. There are some very good reasons. One is so that we can fully understand how geological processes that seem impossibly slow can produce anything of consequence. For example, we are familiar with the concept of driving from one major city to another: a journey of several hours at around 100 kilometres per hour. Continents move toward each other at rates of a fraction of a millimetre per day, or something in the order of 0.00000001 kilometres per hour, and yet, at this impossibly slow rate (try walking at that speed!), they can move thousands of kilometres. Sediments typically accumulate at even slower rates—less than a millimetre per year—but still they are thick enough to be thrust up into monumental mountains and carved into breathtaking canyons.

Another reason is that for our survival on this planet, we need to understand issues like extinction of endangered species and anthropogenic (human-caused) climate change. Some people, who don’t understand geological time, are quick to say that the climate has changed in the past, and that what is happening now is no different. And it certainly has changed in the past—many times. For example, from the Eocene (50 Ma) to the present day, Earth’s climate cooled by about 12°C. That’s a huge change that ranks up there with many of the important climate changes of the distant past, and yet the rate of change over that time was only 0.000024°C/century. Anthropogenic climate change has been 1.1°C over the past century,Climate change data from NASA Goddard Institute for Space Studies:; that is 45,800 times faster than the rate of natural climate change since the Eocene!

One way to wrap your mind around geological time is to put it into the perspective of single year, because we all know how long it is from one birthday to the next. At that rate, each hour of the year is equivalent to approximately 500,000 years, and each day is equivalent to 12.5 million years.

If all of geological time is compressed down to a single year, Earth formed on January 1, and the first life forms evolved in late March (roughly 3,500 Ma). The first large life forms appeared on November 13 (roughly 600 Ma), plants appeared on land around November 24, and amphibians on December 3. Reptiles evolved from amphibians during the first week of December and dinosaurs and early mammals evolved from reptiles by December 13, but the dinosaurs, which survived for 160 million years, were gone by Boxing Day (December 26). The Pleistocene Glaciation got started at around 6:30 p.m. on New Year’s Eve, and the last glacial ice left southern Canada by 11:59 p.m.

It’s worth repeating: on this time scale, the earliest ancestors of the animals and plants with which we are familiar did not appear on Earth until mid-November, the dinosaurs disappeared after Christmas, and most of Canada was periodically locked in ice from 6:30 to 11:59 p.m. on New Year’s Eve. As for people, the first to inhabit B.C. got here about one minute before midnight, and the first Europeans arrived about two seconds before midnight.

It is common for the popular press to refer to distant past events as being “prehistoric.” For example, dinosaurs are reported as being “prehistoric creatures,” even by the esteemed National Geographic Society. The written records of our history date back to about 6,000 years ago, so anything prior to that can be considered “prehistoric.” But to call the dinosaurs prehistoric is equivalent to—and about as useful as—saying that Singapore is beyond the city limits of Kamloops! If we are going to become literate about geological time, we have to do better than calling dinosaurs, or early horses (54 Ma), or even early humans (2.8 Ma), “prehistoric.”

Exercise 8.5 What happened on your birthday?

Using the “all of geological time compressed to one year” concept, determine the geological date that is equivalent to your birthday. First go to Day Number of the Year Calculator to find out which day of the year your birth date is. Then divide that number by 365, and multiply that number by 4,570 to determine the time (in millions since the beginning of geological time). Finally subtract that number from 4,570 to determine the date back from the present.

For example, April Fool’s Day (April 1) is day 91 of the year: 91/365 = 0.2493. 0.2493 x 4,570 = 1,139 million years from the start of time, and 4,570 – 1,193 = 3,377 Ma is the geological date.

Finally, go to the Foundation for Global Community’s “Walk through Time” website to find out what was happening on your day. The nearest date to 3,377 Ma is 3,400 Ma. Bacteria ruled the world at 3,400 Ma, and there’s a discussion about their lifestyles.

See Appendix 3 for Exercise 8.5 answers.



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

Section Summary
8.1 The Geological Time Scale The work of William Smith was critical to the establishment of the first geological time scale early in the 19th century, but it wasn’t until the 20th century that geologists were able to assign reliable dates to the various time periods. Geological time is divided into eons, eras, periods, and epochs and the geological time scale is maintained and updated by the International Commission on Stratigraphy.
8.2 Relative Dating Methods We can determine the relative ages of different rocks by observing and interpreting relationships among them, such as superposition, cross-cutting, and inclusions. Gaps in the geological record are represented by various types of unconformities.
8.3 Dating Rocks Using Fossils Fossils are useful for dating rocks date back to about 600 Ma. If we know the age range of a fossil, we can date the rock, but some organisms lived for many millions of years. Index fossils represent shorter geological times, and if a rock has several different fossils with known age ranges, we can normally narrow the time during which the rock formed.
8.4 Isotopic Dating Methods Radioactive isotopes decay at predictable and known rates, and can be used to date igneous and metamorphic rocks. Some of the more useful isotope systems are potassium-argon, rubidium-strontium, uranium-lead, and carbon-nitrogen.   Radiocarbon dating can be applied to sediments and sedimentary rocks, but only if they are younger than 60 ka.
8.5 Other Dating Methods There are many other methods for dating geological materials. Two that are widely used are dendrochronology and magnetic chronology. Dendrochronology, based on studies of tree rings, is widely applied to dating glacial events. Magnetic chronology is based on the known record of Earth’s magnetic field reversals.
8.6 Understanding Geological Time While knowing about geological time is relatively easy, actually comprehending the significance of the vast amounts of geological time is a great challenge. To be able to solve important geological problems and critical societal challenges, like climate change, we need to really understand geological time.

Questions for Review

Answers to Review Questions can be found in Appendix 2.

  1. A granitic rock contains inclusions (xenoliths) of basalt. What can you say about the relative ages of the granite and the basalt?
  2. Explain the differences between:
    1. a disconformity and a paraconformity
    2. a nonconformity and an angular unconformity
  3. What are the features of a useful index fossil?
  4. The bellow shows a geological cross-section. The granitic rock “f” at the bottom is the one that you estimated the age of in Exercise 8.3. A piece of wood from layer “d” has been sent for radiocarbon dating and the result was 0.55 14C remaining. How old is layer “d”? (You can use the carbon-14 decay curve of Figure 8.27 to answer this question.)
    Left: Cross section through the crust for questions 4 to 7.  Right: 14C decay curve for question 4.
  5. Based on your answer to question 4, what can you say about the age of layer “c” in Figure 8.27.?
  6. What type of unconformity exists between layer “c” and rock “f”?
  7. What about between layer “c” and layer “b”?
  8. We can’t use magnetic chronology to date anything younger than 780,000 years. Why not?
  9. How did William Smith apply the principle of faunal succession to determine the relative ages of the sedimentary rocks of England and Wales?
  10. Access a copy of the geological time scale (International Commission on Stratigraphy). What are the names of the last age (or stage) of the Cretaceous and the first age of the Paleogene? Print out the time scale and stick it on the wall above your desk!

Media Attributions


Chapter 9 Earth’s Interior

Learning Objectives

After carefully reading this chapter, completing the exercises within it, and answering the questions at the end, you should be able to:

  • Explain the variations in the composition and characteristics of Earth’s different layers.
  • Compare the characteristics and behaviour of the two types of seismic body waves.
  • Summarize the variations in seismic-wave velocity as a function of rock type and temperature and pressure conditions.
  • Explain some of the ways that seismic data can be used to understand planetary interiors.
  • Describe the temperature variations within Earth and their implications for internal processes such as mantle convection.
  • Explain the origins of Earth’s magnetic field and the timing of magnetic field reversals.
  • Describe the isostatic relationship between the crust and the mantle, and the implications of that relationship for geological processes on Earth.

In order to understand how Earth works, and especially the mechanisms of plate tectonics (covered in Chapter 10), we need to know something about the inside of our planet — what it’s made of, and what goes on in there. We have a variety of ways of knowing, and these will be discussed in this chapter, but the one thing we can’t do is go down and look! Fortunately there are a few places where mantle rock is exposed on Earth’s surface, and we have some samples of material from the insides of other planetary bodies, in the form of meteorites that have landed on Earth (Figure 9.0.1). We also have a great deal of seismic information that can help us understand the nature of Earth’s interior.

Figure 9.0.1 Left: a fragment of the Tagish Lake meteorite, discovered in 2000 on the ice of Tagish Lake, B.C. It is a “stony” meteorite that is dominated by ferromagnesian silicate minerals, and is similar in composition to Earth’s mantle. Right: part of the Elbogen meteorite discovered in Germany around 1400. It is an iron meteorite, similar in composition to Earth’s core. Both samples are a few centimetres across.

Earth’s interior is broadly divided by composition and depth into crust, mantle, and core (Figure 9.0.2). The crust is primarily (roughly 95%) made up of igneous rock and metamorphic rock with an overall composition between intermediate and felsic. The remaining 5% is made up of sedimentary rock, which is dominated by mudstone.

The mantle includes several layers, all with the same overall ultramafic composition. The upper mantle is typically composed of peridotite, a rock dominated by olivine and pyroxene. The lower mantle has a similar chemical composition, but because of the extreme pressures, different minerals are present, including spinels and garnets. The properties of the mantle also vary with depth, as follows:

The core is primarily composed of iron, with lesser amounts of nickel (about 5%) and several percent oxygen. It is extremely hot (roughly 3500° to 5000°C). The outer core is liquid while the inner core is solid—even though it is hotter—because the pressure is so much greater at that depth.

Although the CMB is just about half of the way to Earth’s centre, the mantle, being on the outside, is by far the major component of Earth. The mantle makes up 82.5% of the volume, the core 16.1%, and the crust only 1.4%.

In the remainder of this chapter, we’ll look first at how we know about Earth’s interior structure, and then at the properties of the different layers and the processes that take place within them.

Figure 9.0.2 Earth’s layers: crust is pink, mantle is green, core is blue. [Image Description]

Image Descriptions

Figure 9.0.2 image description: Layers of the earth
Layer Kilometres below the Earth’s surface Thickness (kilometres)
Crust and Lithospheric part of the mantle 0 to 100 100
Asthenosphere and Upper mantle 100 to 660 560
Lower mantle 660 to 2,700 2,040
D” layer 2,700 to 2,890 190
Outer liquid core 2,890 to 5,100 2,210
Inner solid core 5,100 to 6,370 1,270

[Return to Figure 9.0.2]

Media Attributions


9.1 Understanding Earth Through Seismology

Seismology is the study of vibrations within Earth. These vibrations are caused by various events: earthquakes, extraterrestrial impacts, explosions, storm waves hitting the shore, and tidal effects. Of course, seismic techniques have been most widely applied to the detection and study of earthquakes, but there are many other applications, and arguably seismic waves provide the most important information that we have concerning Earth’s interior. Before going any deeper into Earth, however, we need to take a look at the properties of seismic waves. The types of waves that are useful for understanding Earth’s interior are called body waves, meaning that, unlike the surface waves on the ocean, they are transmitted through Earth materials.

Figure 9.1.1 Hitting a large block of rock with a heavy hammer will create seismic waves within the rock. Please don’t try this at home!

Imagine hitting a large block of strong rock (e.g., granite) with a heavy sledgehammer (Figure 9.1.1). At the point where the hammer strikes it, a small part of the rock will be compressed by a fraction of a millimetre. That compression will transfer to the neighbouring part of the rock, and so on through to the far side of the rock—all in a fraction of a second. This is known as a compression wave, and it can be illustrated by holding a loose spring (like a Slinky) that is attached to something (or someone) at the other end. If you give it a sharp push so the coils are compressed, the compression propagates (travels) along the length of the spring and back (Figure 9.1.2). You can think of a compression wave as a “push” wave—it’s called a P wave (although the “P” stands for “primary” because P waves arrive first at seismic stations).

When we hit a rock with a hammer, we also create a different type of body wave, one that is characterized by back-and-forth vibrations (as opposed to compressions). This is known as a shear wave (S wave, where the “S” stands for “secondary”), and an analogy would be what happens when you flick a length of rope with an up-and-down motion. As shown in Figure 9.1.2, a wave will form in the rope, which will travel to the end of the rope and back.

Figure 9.1.2 A compression wave can be illustrated by a spring (like a Slinky) that is given a sharp push at one end. A shear wave can be illustrated by a rope that is given a quick flick.

Compression waves and shear waves travel very quickly through geological materials. As shown in Figure 9.1.3, typical P wave velocities are between 0.5 kilometres per second (km/s) and 2.5 km/s in unconsolidated sediments, and between 3.0 km/s and 6.5 km/s in solid crustal rocks. Of the common rocks of the crust, velocities are greatest in basalt and granite. S waves are slower than P waves, with velocities between 0.1 km/s and 0.8 km/s in soft sediments, and between 1.5 km/s and 3.8 km/s in solid rocks.

Figure 9.1.3 Typical velocities of P-waves (red) and S-waves (blue) in sediments and in solid crustal rocks. [Image Description]

Exercise 9.1 How soon will seismic waves get here?

Imagine that a strong earthquake takes place on Vancouver Island within Strathcona Park (west of Courtenay). Assuming that the crustal average P wave velocity is 5 km per second, how long will it take (in seconds) for the first seismic waves (P waves) to reach you in the following places (distances from the epicentre are shown)?

  1. Nanaimo (120 km away)
  2. Surrey (200 km away)
  3. Kamloops (390 km away)

See Appendix 3 for Exercise 9.1 answers.

Mantle rock is generally denser and stronger than crustal rock and both P- and S-waves travel faster through the mantle than they do through the crust. Moreover, seismic-wave velocities are related to how tightly compressed a rock is, and the level of compression increases dramatically with depth. Finally, seismic waves are affected by the phase state of rock. They are slowed if there is any degree of melting in the rock. If the material is completely liquid, P waves are slowed dramatically and S waves are stopped altogether.

Seismic Wave Velocity chart. Image description available
Figure 9.1.4 P wave (red) and S wave (blue) velocity variations with depth in Earth. The diagram on the right shows an expanded view of the upper 660 kilometres of the curves in the diagram on the left. [Image Description]

As shown on the right-hand part of Figure 9.1.4, the upper approximately 100 km of the Earth is known as the lithosphere.  This includes the rigid upper part of the mantle (or lithospheric mantle) and the crust.  The next 150 km is the asthenosphere or low velocity zone (because seismic waves are slowed as they pass through that material). As we’ll see below, that part of the mantle is close to it’s melting point and in some regions may be partially molten.

Accurate seismometers have been used for earthquake studies since the late 1800s, and systematic use of seismic data to understand Earth’s interior started in the early 1900s. The rate of change of seismic waves with depth in Earth (as shown in Figure 9.1.4) has been determined over the past several decades by analyzing seismic signals from large earthquakes at seismic stations around the world. Small differences in arrival time of signals at different locations have been interpreted to show that:

One of the first discoveries about Earth’s interior made through seismology was in 1909 when Croatian seismologist Andrija Mohorovičić (pronounced Moho-ro-vi-chich) realized that at certain distances from an earthquake, two separate sets of seismic waves arrived at a seismic station within a few seconds of each other. He reasoned that the waves that went down into the mantle, traveled through the mantle, and then were bent upward back into the crust, reached the seismic station first because although they had farther to go, they traveled faster through mantle rock (as shown in Figure 9.1.5). The boundary between the crust and the mantle is known as the Mohorovičić discontinuity (or Moho). Its depth is between  30 and 40 kilometres beneath most of the continental crust, and between 5 and 10 kilometres beneath the oceanic crust.

Figure 9.1.5 Depiction of seismic waves emanating from an earthquake (red star). Some waves travel through the crust to the seismic station (at about 6 km/s), while others go down into the mantle (where they travel at around 8 km/s) and are bent upward toward the surface, reaching the station before the ones that traveled only through the crust.

Our current understanding of the patterns of seismic wave transmission through Earth is summarized in Figure 9.1.6. Because of the gradual increase in density (and therefore rock strength) with depth, all waves are refracted (toward the lower density material) as they travel through homogenous parts of Earth and thus tend to curve outward toward the surface. Waves are also refracted at boundaries within Earth, such as at the Moho, at the core-mantle boundary (CMB), and at the outer-core/inner-core boundary.

S waves do not travel through liquids—they are stopped at the CMB—and there is an S wave shadow on the side of Earth opposite a seismic source. The angular distance from the seismic source to the shadow zone is 103° on either side, so the total angular distance of the shadow zone is 154°. We can use this information to infer the depth to the CMB.

P waves do travel through liquids, so they can make it through the liquid part of the core. Because of the refraction that takes place at the CMB, waves that travel through the core are bent away from the surface, and this creates a P wave shadow zone on either side, from 103° to 150°. This information can be used to discover the differences between the inner and outer parts of the core.

Figure 9.1.6 Patterns of seismic wave propagation through Earth’s mantle and core. S waves do not travel through the liquid outer core, so they leave a shadow on Earth’s far side where they cannot get to. P waves do travel through the core, but because the waves that enter the core are refracted, there are also P wave shadow zones.

Exercise 9.2 Liquid Cores in Other Planets

Figure 9.1.7

We know that other planets must have (or at least did have) liquid cores like ours, and we could use seismic data to find out how big they are. The S wave shadow zones on planets A and B are shown. Using the same method used for Earth (on the left), sketch in the outlines of the cores for these two other planets.

See Appendix 3 for Exercise 9.2 answers.

Figure 9.1.8 P-wave tomographic profile of area in the southern Pacific Ocean from southeast of Tonga to Fiji. Blue represents rock that has relatively high seismic velocities, while yellow and red represent rock with low velocities. Open circles are earthquakes used in the study.

Using data from many seismometers and hundreds of earthquakes, it is possible to create a two- or three-dimensional image of the seismic properties of part of the mantle. This technique is known as seismic tomography, and an example of the result is shown in Figure 9.1.8.

The Pacific Plate subducts beneath Tonga and appears in Figure 9.1.8 as a 100 kilometre thick slab of cold (blue-coloured) oceanic crust that has pushed down into the surrounding hot mantle. The cold rock is more rigid than the surrounding hot mantle rock, so it is characterized by slightly faster seismic velocities. There is volcanism in the Lau spreading centre and also in the Fiji area, and the warm rock in these areas has slower seismic velocities (yellow and red colours).

Image descriptions

Figure 9.1.3 image description: Wave velocity in different materials in kilometres per second.
Material S Wave (kilometres per second) P Wave (kilometres per second)
Dry sand 0.1 to 0.4 0.4 to 1.3
Clay 0.2 to 0.6 0.6 to 1.6
Wet sand 0.7 to 0.8 1.5 to 2.2
Till 0.8 to 1.0 1.9 to 2.6
Mudstone 2.1 to 2.3 3.0 to 4.3
Sandstone 1.4 to 2.5 3.0 to 5.0
Limestone 2.4 to 3.1 4.2 to 5.8
Granite 3.0 to 3.7 4.9 to 5.9
Basalt 3.3 to 4.0 5.2 to 6.2

[Return to Figure 9.1.3]

Figure 9.1.4 image description: P-wave and S-wave velocity variations with depth in Earth.
Layer Depth from surface (km) S-Wave velocity (kilometres per second) P-Wave velocity (kilometres per second)
Crust 0 to 30 3.0 to 4.6 5.3 to 7.0
Lithosphere 30 to 100 4.6 to 5.8 7.0 to 8.7
Asthenosphere 100 to 250 5.0 to 5.9 7.8 to 8.5
Mantle 250 to 2890 5. to 7.0 8.2 to 12.6
Outer core 2890 to 5100 0 8.0 to 10.1
Inner core 5100 to 6370 0 11.8 to 12.0

[Return to Figure 9.1.4]

Media Attributions


9.2 The Temperature of Earth’s Interior

As we’ve discussed in the context of metamorphism, Earth’s internal temperature increases with depth. However, as shown in Figure 9.2.1 (right), that rate of increase is not linear. The temperature gradient is around 15° to 30°C per kilometre within the upper 100 kilometres; it then drops off dramatically through the mantle, increases more quickly at the base of the mantle, and then increases slowly through the core. The temperature is around 1000°C at the base of the crust, around 3500°C at the base of the mantle, and around 5,000°C at Earth’s centre. The temperature gradient within the lithosphere (upper 100 kilometres) is quite variable depending on the tectonic setting. Gradients are lowest in the central parts of continents, higher in the vicinity of subduction zones, and higher still at divergent boundaries.

Figure 9.2.1 Right: generalized rate of temperature increase with depth within Earth. Temperature increases to the right, so the flatter the line, the steeper the temperature gradient. Our understanding of the temperature gradient comes from seismic wave information and knowledge of the melting points of Earth’s materials. Left: Rate of temperature increase with depth in Earth’s upper 500 kilometres, compared with the dry mantle rock melting curve (red dashed line). LVZ= low-velocity zone. [Image Description]

Figure 9.2.1 (left) shows a typical temperature curve for the upper 500 kilometres of the mantle in more detail, along with the melting curve for dry mantle rock. (Mantle rock will melt under conditions to the right of the dashed red line.) In general the mantle is not molten because the temperature lies to the left of the melting curve, but within the depth interval between 100 and 250 kilometres the temperature curve comes very close to the melting boundary for dry mantle rock. At these depths, therefore, mantle rock is either very nearly melted or partially melted. In some situations, where extra heat is present and the temperature line crosses over the melting line, or where water is present, it may be completely molten. This region of the mantle—the asthenosphere—is also known as the low-velocity zone because seismic waves are slowed within rock that is near its melting point. Below 250 kilometres the temperature stays on the left side of the melting line; in other words, the mantle is solid from here all the way down to the D” layer near the core-mantle-boundary.

The fact that the temperature gradient is much less in the main part of the mantle than in the lithosphere has been interpreted to indicate that the mantle is convecting, and therefore that heat from depth is being brought toward the surface faster than it would be with only heat conduction. As we’ll see in Chapter 10, a convecting mantle is an key feature of plate tectonics.

The convection of the mantle is a product of the upward transfer of heat from the core to the lower mantle. As in a pot of soup on a hot stove (Figure 9.2.2), the material near the heat source becomes hot and expands, making it lighter than the material above. The force of buoyancy causes it to rise, and cooler material flows in from the sides. The mantle convects in this way because the heat transfer from below is not perfectly even, and also because, even though mantle material is solid rock, it is sufficiently plastic to slowly flow (at rates of centimetres per year) as long as a steady force is applied to it.

As in the soup pot example, Earth’s mantle will no longer convect once the core has cooled to the point where there is not enough heat transfer to overcome the strength of the rock. This has already happened on smaller planets like Mercury and Mars, as well as on Earth’s Moon.

Figure 9.2.2 Convection in a pot of soup on a hot stove (left). As long as heat is being transferred from below, the liquid will convect. If the heat is turned off (right), the liquid remains hot for a while, but convection will cease.

Why is the inside of the Earth hot?

Figure 9.2.3 Heat flow on Earth from radioactive decay

The heat of Earth’s interior comes from two main sources, each contributing about 50% of the heat. One of those is the frictional heat left over from the collisions of large and small particles that created Earth in the first place, plus the subsequent frictional heat of redistribution of material within Earth by gravitational forces (e.g., sinking of iron to form the core).

The other source is radioactivity, specifically the spontaneous radioactive decay of the isotopes 235U, 238U, 40K, and 232Th, which are primarily present in the mantle. As shown on Figure 9.2.3, the total heat produced that way has been decreasing over time (because these isotopes are getting used up), and is now roughly 25% of what it was when Earth formed. This means that Earth’s interior is slowly becoming cooler.

Image Descriptions

Figure 9.2.1 image description: Temperature increase within the Earth.
Layer Depth (kilometres) Temperature increase (Celsius) Temperature increase rate (Degrees per kilometre)
Lithosphere 0 to 100 0 to 1400 14
Asthenosphere 100 to 250 1400 to 1700 2
Mantle 250 to 1000 1700 to 2100 0.53
1000 to 2000 2100 to 2600 0.5
2000 to 2890 2600 to 3800 1.35
Outer Core 2890 to 4000 3800 to 4600 0.72
4000 to 5100 4600 to 5000 0.36
Inner core 5100 to 6370 5000 to 5100 0.079

[Return to Figure 9.2.1]

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9.3 Earth’s Magnetic Field

Heat is also being transferred from the solid inner core to the liquid outer core, and this leads to convection of the liquid iron of the outer core. Because iron is a metal and conducts electricity (even when molten), its motion generates a magnetic field.

Figure 9.3.1 Depiction of Earth’s magnetic field as a bar magnet coinciding with the core. The south pole of such a magnet points to Earth’s North Pole. The red arrows represent the orientation of the magnetic field at various locations on Earth’s surface.

Earth’s magnetic field is defined by the North and South Poles that align generally with the axis of rotation (Figure 9.3.1). The lines of magnetic force flow into Earth in the northern hemisphere and out of Earth in the southern hemisphere. Because of the shape of the field lines, the magnetic force trends at different angles to the surface in different locations (red arrows of Figure 9.3.1). At the North and South magnetic poles, the force is vertical. Anywhere near to the equator the force is horizontal, and everywhere in between, the magnetic force is at some intermediate angle to the surface. As we’ll see in Chapter 10, the variations in these orientations provide a critical piece of evidence to the understanding of continental drift as an aspect of plate tectonics.

Earth’s magnetic field is generated within the outer core by the convective movement of liquid iron, but as we discovered in Chapter 8, the magnetic field is not stable over geological time. For reasons that are not completely understood, the magnetic field decays periodically and then becomes re-established. When it does re-establish, it may be oriented the way it was before the decay, or it may be oriented with the reversed polarity. Over the past 250 Ma, there have been a few hundred magnetic field reversals, and their timing has been anything but regular. The shortest ones that geologists have been able to define lasted only a few thousand years, and the longest one was more than 30 million years, during the Cretaceous (Figure 9.3.2).

Exercise 9.3 What would a magnetic dip meter tell you?

Regular compasses point only to the north magnetic pole, but if you have a magnetic dip meter you could also measure the angle of the magnetic field at your location in the up-and-down sense.

Using Figure 9.3.1 as a guide, describe where you’d be on Earth if the vertical angles are as follows:

  1. Up at a shallow angle
  2. Parallel to the ground
  3. Down at a steep angle
  4. Straight down

See Appendix 3 for Exercise 9.3 answers.

Figure 9.3.2 Magnetic field reversal chronology for the past 170 Ma. The first 5 Ma of the magnetic chronology are shown in more detail in Figure 8.5.3, although the time scale is in the opposite direction in that figure.

Changes in Earth’s magnetic field have been studied using a mathematical model, and reversals have been shown to take place when the model was run to simulate a period of several hundred thousand years. The fact that field reversals took place shows that the model is a reasonably accurate representation of the Earth. According to the lead author of the study, Gary Glatzmaier, of University of California at Santa Cruz: “Our solution shows how convection in the fluid outer core is continually trying to reverse the field but that the solid inner core inhibits magnetic reversals because the field in the inner core can only change on the much longer time scale of diffusion. Only once in many attempts is a reversal successful, which is probably the reason why the times between reversals of the Earth’s field are long and randomly distributed.” A depiction of Earth’s magnetic field lines during a stable period and during a reversal is shown in Figure 9.3.3. To read more about these phenomena see Glatzmaier’s Geodynamo website.

Figure 9.3.3 Depiction of Earth’s magnetic field between reversals (left) and during a reversal (right). The lines represent magnetic field lines: blue where the field points toward Earth’s centre and yellow where it points away. The rotation axis of Earth is vertical, and the outline of the core is shown as a dashed white circle. [Image Description]

Image Descriptions

Figure 9.3.3 image description: The earth’s magnetic fields is normally very uniform with the magnetic field pointing towards the earth in the north and away from the earth in the south. During a reversal, the Earth’s magnetic field becomes very convoluted. [Return to Figure 9.3.3]

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9.4 Isostasy

Theory holds that the mantle is able to convect because of its plasticity, and this plasticity also allows for another very important Earth process known as isostasy. The literal meaning of the word isostasy is “equal standstill,” but the importance behind it is the principle that Earth’s crust is floating on the mantle, like a raft floating in the water, rather than resting on the mantle like a raft sitting on the ground.

The relationship between the crust and the mantle is illustrated in Figure 9.4.1 On the right is an example of a non-isostatic relationship between a raft and solid concrete. It’s possible to load the raft up with lots of people, and it still won’t sink into the concrete. On the left, the relationship is an isostatic one between two different rafts and a swimming pool full of peanut butter. With only one person on board, the raft floats high in the peanut butter, but with three people, it sinks dangerously low. We’re using peanut butter here, rather than water, because its viscosity more closely represents the relationship between the crust and the mantle. Although it has about the same density as water, peanut butter is much more viscous (stiff), and so although the three-person raft will sink into the peanut butter, it will do so quite slowly.

Figure 9.4.1 Illustration of a non-isostatic relationship between a raft and solid ground (right) and of isostatic relationships between rafts and peanut butter (left).

The relationship of Earth’s crust to the mantle is similar to the relationship of the rafts to the peanut butter. The raft with one person on it floats comfortably high. Even with three people on it the raft is less dense than the peanut butter, so it floats, but it floats uncomfortably low for those three people. The crust, with an average density of around 2.6 grams per cubic centimetre (g/cm3), is less dense than the mantle (average density of approximately 3.4 g/cm3 near the surface, but more than that at depth), and so it is floating on the “plastic” mantle. When more weight is added to the crust, through the process of mountain building, it slowly sinks deeper into the mantle and the mantle material that was there is pushed aside (Figure 9.4.2, left). When that weight is removed by erosion over tens of millions of years, the crust rebounds and the mantle rock flows back (Figure 9.4.2, right).

Figure 9.4.2 Illustration of the isostatic relationship between the crust and the mantle. Following a period of mountain building, mass has been added to a part of the crust, and the thickened crust has pushed down into the mantle (left). Over the following tens of millions of years, the mountain chain is eroded and the crust rebounds (right). The green arrows represent slow mantle flow.

The crust and mantle respond in a similar way to glaciation and deglaciation as they do to the growth and erosion of mountain ranges. Thick accumulations of glacial ice add weight to the crust, and as the mantle beneath is squeezed to the sides, the crust subsides. This process is illustrated for the current ice sheet on Greenland in Figure 9.4.3 (a and b). The Greenland Ice Sheet at this location is over 2,500 metres thick, and the crust beneath the thickest part has been depressed to the point where it is below sea level over a wide area. When the ice eventually melts, the crust and mantle will slowly rebound, but full rebound will likely take more than 10,000 years (Figure 9.4.3c).

Figure 9.4.3 (a) A cross-section through the crust in the northern part of Greenland (The ice thickness is based on data from NASA and the Center for Remote Sensing of Ice Sheets, but the crust thickness is less than it should be for the sake of illustration.) The maximum ice thickness is over 2,500 m. The red arrows represent downward pressure on the mantle because of the mass of the ice. (b) Depiction of the situation after complete melting of the ice sheet, a process that could happen within 2,000 years if people and their governments continue to ignore climate change. The isostatic rebound of the mantle would not be able to keep up with this rate of melting, so for several thousand years the central part of Greenland would remain close to sea level, in some areas even below sea level.
Figure 9.4.3c Depiction of the crust beneath a post-glacial Greenland once isostatic equilibrium is achieved.  It is likely that complete rebound of the mantle would take more than 10,000 years.

How can the mantle be both solid and plastic?

Figure 9.4.4

You might be wondering how it is possible that Earth’s mantle is rigid enough to break during an earthquake, and yet it convects and flows like a very viscous liquid. The explanation is that the mantle behaves as a non-Newtonian fluid, meaning that it responds differently to stresses depending on how quickly the stress is applied. A good example of this is the behaviour of the material known as Silly Putty, which can bounce and will break if you pull on it sharply, but will deform like a liquid if stress is applied slowly. In this photo, Silly Putty was placed over a hole in a glass tabletop, and in response to gravity, it slowly flowed into the hole. The mantle will flow when placed under the slow but steady stress of a growing (or melting) ice sheet.


Figure 9.4.5  The current rates of post-glacial isostatic uplift (green, blue, and purple shades) and subsidence (yellow and orange). Subsidence is taking place where the mantle is slowly flowing back toward areas that are experiencing post-glacial uplift.

Large parts of Canada are still rebounding as a result of the loss of glacial ice over the past 12 ka, and as shown in Figure 9.4.5, other parts of the world are also experiencing isostatic rebound. The highest rate of uplift is in within a large area to the west of Hudson Bay, which is where the Laurentide Ice Sheet was the thickest (over 3,000 m). Ice finally left this region around 8,000 years ago, and the crust is currently rebounding at a rate of nearly 2 centimetres per year. Strong isostatic rebound is also occurring in northern Europe where the Fenno-Scandian Ice Sheet was thickest, and in the eastern part of Antarctica, which also experienced significant ice loss during the Holocene.

There are also extensive areas of subsidence surrounding the former Laurentide and Fenno-Scandian Ice Sheets. During glaciation, mantle rock flowed away from the areas beneath the main ice sheets, and this material is now slowly flowing back, as illustrated in Figure 9.4.3b.

Exercise 9.4 Rock density and isostasy

The densities (also known as “specific gravity”) of a number of common minerals are given in Table 9.1.

Table 9.1 Densities of common minerals.
Mineral Density (grams per cubic centimetre, g/cm3)
Quartz 2.65
Feldspar 2.63
Amphibole 3.25
Pyroxene 3.4
Olivine 3.3

The following table provides the approximate proportions of these minerals in the continental crust (typified by granite), oceanic crust (mostly basalt), and mantle (mainly the rock known as peridotite). Assuming that you have 1,000 cm3 of each rock type, estimate the respective rock-type densities. For each rock type, you will need to multiply the volume of the different minerals in the rock by their density, and then add those numbers to get the total weight for 1,000 cm3 of that rock. The density is that number divided by 1,000. The continental crust is done for you.

Table 9.2 Determine the density of different kinds of crusts
Rock Type Volumes of individual minerals in 1000 cm3. Grams of individual minerals in 1000 cm3 Total Weight (grams) Density (grams per cubic centimetre, g/cm3)
Continental Crust (Granite) Quartz – 180 cm3

Feldspar – 760 cm3

Amphibole – 70 cm3

Quartz – 477 g

Feldspar – 1999 g

Amphibole – 277 g

2703 g 2.70
Oceanic Crust (Basalt) Feldspar – 450 cm3

Amphibole – 50 cm3

Pyroxene – 500 cm3

Feldspar –

Amphibole –

Pyroxene –

Mantle (Peridotite) Pyroxene – 450 cm3

Olivine – 550 cm3

Pyroxene –

Olivine –

If continental crust (represented by granite) and oceanic crust (represented by basalt) are like rafts floating on the mantle, what does this tell you about how high or low they should float?

This concept is illustrated in Figure 9.4.6. The dashed line is for reference, showing points at equal distance from Earth’s centre.

Figure 9.4.6

See Appendix 3 for Exercise 9.4 answers.

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

Section Summary
9.1 Understanding Earth Through Seismology Seismic waves that travel through Earth are either P-waves (compression, or “push” waves) or S-waves (shear waves). P-waves are faster than S-waves, and can pass through fluids. By studying seismic waves, we can discover the nature and temperature characteristics of the various parts of Earth’s interior.
9.2 The Temperature of Earth’s Interior Earth’s temperature increases with depth (to around 5000°C at the centre), but there are significant variations in the rate of temperature increase. These variations are related to differences in composition and the existence of convection in the mantle and liquid part of the core.
9.3 Earth’s Magnetic Field Because of outer-core convection, Earth has a magnetic field. The magnetic force directions are different at different latitudes. The polarity of the field is not constant, and has flipped from “normal” (as it is now) to reversed and back to normal hundreds of times in the past.
9.4 Isostasy The “plastic” nature of the mantle, which allows for mantle convection, also determines the nature of the relationship between the crust and the mantle. The crust floats on the mantle in an isostatic relationship. Where the crust becomes thicker because of mountain building, it pushes farther down into the mantle. Oceanic crust, being heavier than continental crust, floats lower on the mantle.

Questions for Review

Answers to Review Questions can be found in Appendix 2.

  1. What parts of Earth are most closely represented by typical stony meteorites and typical iron meteorites?
  2. On the below diagram draw (from memory) and label the approximate locations of the following boundaries: crust/mantle, mantle/core, outer core/inner core.""
  3. Describe the important differences between P-waves and S-waves.
  4. Why does P-wave velocity decrease dramatically at the core-mantle boundary?
  5. Why do both P-waves and S-waves gradually bend as they move through the mantle?
  6. What is the evidence for mantle convection, and what is the mechanism that causes it?
  7. Where and how is Earth’s magnetic field generated?
  8. When were the last two reversals of Earth’s magnetic field?
  9. What property of the mantle is essential for the isostatic relationship between the crust and the mantle?
  10. How would you expect the depth to the crust-mantle boundary in the area of the Rocky Mountains to differ from that in central Saskatchewan?
  11. As you can see in Figure 9.22, British Columbia is still experiencing weak post-glacial isostatic uplift, especially in the interior, but also along the coast. Meanwhile offshore areas are experiencing weak isostatic subsidence. Why?


Chapter 10 Plate Tectonics

Learning Objectives

After carefully reading this chapter, completing the exercises within it, and answering the questions at the end, you should be able to:

  • Discuss some of the early evidence for continental drift and Alfred Wegener’s role in promoting this theory.
  • Explain some of the other models that were used early in the 20th century to understand global geological features.
  • Describe the numerous geological advances made in the middle part of the 20th century that provided the basis for understanding the mechanisms of plate tectonics and the evidence that plates have moved and lithosphere is created and destroyed.
  • List the seven major plates, their extents, and their general directions of motion, and identify the types of boundaries between them.
  • Describe the geological processes that take place at divergent and convergent plate boundaries, and explain the existence of transform faults.
  • Explain how super-continents form and how they break apart.
  • Describe the mechanisms for plate movement.

As we discovered in Chapter 1, plate tectonics is the model or theory that we use to understand how our planet works. More specifically it is a model that explains the origins of continents and oceans, folded rocks and mountain ranges, igneous and metamorphic rocks, earthquakes (Figure 10.0.1) and volcanoes, and continental drift. Plate tectonics was first proposed just over 100 years ago, but did not become an accepted part of geology until about 50 years ago. It took 50 years for this theory to be accepted for a few reasons. First, it was a true revolution in thinking about Earth, and that was difficult for many established geologists to accept. Second, there was a political gulf between the main proponent of the theory Alfred Wegener (from Germany) and the geological establishment of the day, which was mostly centred in Britain and the United States. Third, the evidence and understanding of Earth that would have supported plate tectonic theory simply didn’t exist until the middle of the 20th century.

Figure 10.0.1 The San Andreas Fault at Pt. Reyes Station, California. The two parts of the fence show the offset on the fault caused by the M7.9 San Francisco earthquake in 1906. The near side of the fence is on the Pacific Plate and the far side is on the North America Plate. The relationship between tectonic plates and earthquakes was not known in Alfred Wegener’s time.

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10.1 Alfred Wegener: The Father of Plate Tectonics

Figure 10.1.1 Alfred Wegener a few years before his death in 1930.

Alfred Wegener (1880-1930) (Figure 10.1.1) earned a PhD in astronomy at the University of Berlin in 1904, but he had always been interested in geophysics and meteorology and spent most of his academic career working in meteorology. In 1911 he happened on a scientific publication that included a description of the existence of matching Permian-aged terrestrial fossils in various parts of South America, Africa, India, Antarctica, and Australia (Figure 10.1.2).

Wegener concluded that this distribution of terrestrial organisms could only exist if these continents were joined together during the Permian, and he coined the term Pangea (“all land”) for the supercontinent that he thought included all of the present-day continents.

Figure 10.1.2 The distribution of several Permian terrestrial fossils that are present in various parts of continents that are now separated by oceans. [Image Description]

Wegener pursued his theory with determination—combing the libraries, consulting with colleagues, and making observations—looking for evidence to support it. He relied heavily on matching geological patterns across oceans, such as sedimentary strata in South America matching those in Africa (Figure 10.1.3), North American coalfields matching those in Europe, and the mountains of Atlantic Canada matching those of northern Britain—both in morphology and rock type.

Figure 10.1.3 A cross-section showing the geological similarities between parts of Brazil (South America) on the left and Angola (Africa) on the right. The pink layer is a salt deposit, which is now known to be common in areas of continental rifting.

Wegener referred to the evidence for the Carboniferous and Permian (~300 Ma) Karoo Glaciation in South America, Africa, India, Antarctica, and Australia (Figure 10.1.4). He argued that this could only have happened if these continents were once all connected as a single supercontinent. He also cited evidence (based on his own astronomical observations) that showed that the continents were moving with respect to each other, and determined a separation rate between Greenland and Scandinavia of 11 metres per year, although he admitted that the measurements were not accurate. In fact they weren’t even close—the separation rate is actually about 2.5 centimetres per year!

Figure 10.1.4 The distribution of the Carboniferous and Permian Karoo Glaciation (outlined in blue).

Wegener first published his ideas in 1912 in a short book called Die Entstehung der Kontinente (The Origin of Continents), and then in 1915 in Die Entstehung der Kontinente und Ozeane (The Origin of Continents and Oceans). He revised this book several times up to 1929. It was translated into French, English, Spanish, and Russian in 1924.

In fact the continental fits were not perfect and the geological matchups were not always consistent, but the most serious problem of all was that Wegener could not conceive of a credible mechanism for moving the continents around. It was understood by this time that the continents were primarily composed of sialic material (SIAL: silicon and aluminum dominated, similar to “felsic”), and that the ocean floors were primarily simatic (SIMA: silicon and magnesium dominated, similar to “mafic”). Wegener proposed that the continents were like icebergs floating on the heavier SIMA crust, but the only forces that he could invoke to propel continents around were poleflucht, the effect of Earth’s rotation pushing objects toward the equator, and the lunar and solar tidal forces, which tend to push objects toward the west. It was quickly shown that these forces were far too weak to move continents, and without any reasonable mechanism to make it work, Wegener’s theory was quickly dismissed by most geologists of the day.

Alfred Wegener died in Greenland in 1930 while carrying out studies related to glaciation and climate. At the time of his death, his ideas were tentatively accepted by only a small minority of geologists, and soundly rejected by most. However, within a few decades that was all to change. For more about his extremely important contributions to Earth science, visit the NASA website to see a collection of articles on Alfred Wegener.

Image Descriptions

Figure 10.1.2 image description: Fossils found across different continents suggest that these continents were once joined as a super-continent. Fossil remains of Cynognathus (a terrestrial reptile) and Mesosaurus (a freshwater reptile) have been found in South America and Africa. Fossil evidence of the Lystrosaurus, a land reptile from the Triassic period, has been found in India, Africa, and Antarctica. Fossils of the fern Glossopteris have been found in Australia, Antarctica, India, Africa, and South America. When you position these continents so they fit together, the areas where these fossils were found line up. [Return to Figure 10.1.2]

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10.2 Global Geological Models of the Early 20th Century

The untimely death of Alfred Wegener didn’t solve any problems for those who opposed his ideas because they still had some inconvenient geological truths to deal with. One of those was explaining the distribution of terrestrial species across five continents that are currently separated by hundreds or thousands of kilometres of ocean water (Figure 10.1.2), and another was explaining the origin of extensive fold-belt mountains, such as the Appalachians, the Alps, the Himalayas, and the Canadian Rockies.

Before we go any further, it is important to know what was generally believed about global geology before plate tectonics. At the beginning of the 20th century, geologists had a good understanding of how most rocks were formed and understood their relative ages through interpretation of fossils, but there was considerable controversy regarding the origin of mountain chains, especially fold-belt mountains. At the end of the 19th century, one of the prevailing views on the origin of mountains was the theory of contractionism—the idea that since Earth is slowly cooling, it must also be shrinking. In this scenario, mountain ranges had formed like the wrinkles on a dried-up apple, and the oceans had submerged parts of former continents. While this theory helped to address the dilemma of the terrestrial fossils, it came with its own set of problems, one being that the amount of cooling couldn’t produce the necessary amount of shrinking, and the other being the principle of isostasy (which had already been around for several decades), which wouldn’t allow continents to sink. (See Section 9.4 for a review of the important principle of isostasy.)

Another widely held view was permanentism, in which it was believed that the continents and oceans have always been generally as they are today. This view incorporated a mechanism for creation of mountain chains known as the geosyncline theory. A geosyncline is a thick deposit of sediments and sedimentary rocks, typically situated along the edge of a continent (Figure 10.2.1).

Figure 10.2.1 The development of a geosyncline along a continental margin. (Note that a geosyncline is not related to a syncline, which is a downward fold in sedimentary rocks.)

The idea of geosynclines developing into fold-belt mountains originated in the middle of the 19th century, proposed first by James Hall and later elaborated by Dwight Dana, both of whom worked extensively in the Appalachian Mountains of the eastern United States. The process of converting a geosyncline into a mountain belt was never really adequately explained, although it was widely believed that mountain belts formed when geosynclines were compressed by forces pushing from either side. The problem is that, without the lateral forces related to plate tectonics, no one was able to adequately describe what would do the pushing. The sediments that accumulate within a geosyncline are derived from erosion of the adjacent continent. Geosynclinal sediments—which eventually turn into sedimentary rocks—may be many thousands of metres thick. As they accumulate, they push down the pre-existing crustal rocks. Extensive geosynclinal deposits exist around much of the coastline of most of the continents; there is a large geosyncline along the eastern edge of North America.

Proponents of the geosyncline theory of mountain formation—and there were many well into the 1960s—also had the problem of explaining the intercontinental terrestrial fossil matchups. The simple explanation was that there were “land bridges” across the Atlantic along which animals and plants could migrate back and forth. One proponent of this idea was the American naturalist Ernest Ingersoll. Referring to evidence of past climate changes, Ingersoll contributed the following to the Encyclopedia Americana in 1920: “The most interesting feature of these changes, however, is that by which, now and again, the Old World was connected with the New by necks or spaces of land, known as “land-bridges”; especially as these permitted an interchange of plants and animals, giving to us many new ones from the other side of the ocean, including, finally, man himself.”

There are serious problems with the land-bridge theory. One is that it is completely inconsistent with isostasy, and another is that there is no evidence of the remnants of the land bridges. The Atlantic Ocean is several thousand metres deep over wide areas, and so the underwater slopes leading up to a land bridge would have to have been at least tens of kilometres wide in most places, and many times that in others. A land bridge of that size would certainly have left some trace.

Exercise 10.1 Fitting the continents together

Figure 10.2.2

The main continents around the Atlantic Ocean are depicted here in the shapes that they might have had during the Mesozoic, including the extents of their continental shelves. Cut these shapes out and see how well you can fit them together in the positions that these areas occupied within Pangea. You can refer to a map of Pangea to help you make the fit.

See Appendix 3 for Exercise 10.1 answers.

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10.3 Geological Renaissance of the Mid-20th Century

As the mineral magnetite (Fe3O4) crystallizes from magma, it becomes magnetized with an orientation parallel to that of Earth’s magnetic field at that time. This is called remnant magnetism. Rocks like basalt, which cool from a high temperature and commonly have relatively high levels of magnetite (up to 1 or 2%), are particularly susceptible to being magnetized in this way, but even sediments and sedimentary rocks, as long as they have small amounts of magnetite, will take on remnant magnetism because the magnetite grains gradually become reoriented following deposition. By studying both the horizontal and vertical components of the remnant magnetism, one can tell not only the direction to magnetic north at the time of the rock’s formation, but also the latitude where the rock formed relative to magnetic north.

In the early 1950s, a group of geologists from Cambridge University, including Keith Runcorn, Ted Irving,Ted Irving later set up a paleomagnetic lab at the Geological Survey of Canada in Sidney, B.C., and did a great deal of important work on understanding the geology of western North America. and several others, started looking at the remnant magnetism of Phanerozoic British and European volcanic rocks, and collecting paleomagnetic data. They found that rocks of different ages sampled from generally the same area showed quite different apparent magnetic pole positions (Figure 10.3.1). They initially assumed that this meant that Earth’s magnetic field had, over time, departed significantly from its present position—which is close to the rotational pole.

Figure 10.3.1 Apparent polar-wandering paths (APWP) for Eurasia and North America. The view is from the North Pole (black dot) looking down. The outer circle is the equator. In the diagram to the right the curve locations have been corrected taking continental drift into account.

The curve defined by the paleomagnetic data was called a polar wandering path because Runcorn and his students initially thought that their data represented actual movement of the magnetic poles (since geophysical models of the time suggested that the magnetic poles did not need to be aligned with the rotational poles). We now know that the magnetic data define movement of continents, and not of the magnetic poles, so we call it an apparent polar wandering path (APWP).

What is a polar wandering path?

The magnetic orientation of rocks in Europe from 500 Ma to the present. Image description available
Figure 10.3.2 [Image Description]

At around 500 Ma, what we now call Europe was south of the equator, and so European rocks formed then would have acquired an upward-pointing magnetic field orientation (see Figure 9.3.2 and Figure 10.3.2). Between then and now, Europe gradually moved north, and the rocks forming at various times acquired steeper and steeper downward-pointing magnetic orientations.When researchers evaluated magnetic data in this way in the 1950s, they plotted where the North Pole would have appeared to be based on the magnetic data and assumed that the continent was always where it is now. That means that the 500 Ma “apparent” north pole would have been somewhere in the South Pacific, and that over the following 500 million years it would have gradually moved north.Of course we now know that the magnetic poles don’t move around much (although polarity reversals do take place) and that the reason Europe had a magnetic orientation characteristic of the southern hemisphere is that it was in the southern hemisphere at 500 Ma.Runcorn and colleagues soon extended their work to North America, and this also showed apparent polar wandering, but the results were not consistent with those from Europe. For example, the 200 Ma pole for North America plotted somewhere in China, while the 200 Ma pole for Europe plotted in the Pacific Ocean. Since there could only have been one pole position at 200 Ma, this evidence strongly supported the idea that North America and Europe had moved relative to each other since 200 Ma. Subsequent paleomagnetic work showed that South America, Africa, India, and Australia also have unique polar wandering curves. In 1956, Runcorn changed his mind and became a proponent of continental drift.This paleomagnetic work of the 1950s was the first new evidence in favour of continental drift, and it led a number of geologists to start thinking that the idea might have some merit. Nevertheless, for a majority of geologists working on global geology at the time, this type of evidence was not sufficiently convincing to get them to change their views.

During the 20th century, our knowledge and understanding of the ocean basins and their geology increased dramatically. Before 1900, we knew virtually nothing about the bathymetry and geology of the oceans. By the end of the 1960s, we had detailed maps of the topography of the ocean floors, a clear picture of the geology of ocean floor sediments and the solid rocks underneath them, and almost as much information about the geophysical nature of ocean rocks as of continental rocks.

Up until about the 1920s, ocean depths were measured using weighted lines dropped overboard. In deep water this is a painfully slow process and the number of soundings in the deep oceans was probably fewer than 1,000. That is roughly one depth sounding for every 350,000 square kilometres of the ocean. To put that in perspective, it would be like trying to describe the topography of British Columbia with elevation data for only a half a dozen points! The voyage of the Challenger in 1872 and the laying of trans-Atlantic cables had shown that there were mountains beneath the seas, but most geologists and oceanographers still believed that the oceans were essentially vast basins with flat bottoms, filled with thousands of metres of sediments.

Following development of acoustic depth sounders in the 1920s (Figure 10.3.3), the number of depth readings increased by many orders of magnitude, and by the 1930s, it had become apparent that there were major mountain chains in all of the world’s oceans. During and after World War II, there was a well-organized campaign to study the oceans, and by 1959, sufficient bathymetric data had been collected to produce detailed maps of all the oceans (Figure 10.3.4).

Figure 10.3.3 Depiction of a ship-borne acoustic depth sounder. The instrument emits a sound (black arcs) that bounces off the sea floor and returns to the surface (white arcs). The travel time is proportional to the water depth.
Figure 10.3.4 Ocean floor bathymetry (and continental topography). Inset (a): the mid-Atlantic ridge, (b): the Newfoundland continental shelf, (c): the Nazca trench adjacent to South America, and (d): the Hawaiian Island chain.

The important physical features of the ocean floor are:

Seismic reflection sounding involves transmitting high-energy sound bursts and then measuring the echos with a series of geophones towed behind a ship. The technique is related to acoustic sounding as described above; however, much more energy is transmitted and the sophistication of the data processing is much greater. As the technique evolved, and the amount of energy was increased, it became possible to see through the sea-floor sediments and map the bedrock topography and crustal thickness. Hence sediment thicknesses could be mapped, and it was soon discovered that although the sediments were up to several thousands of metres thick near the continents, they were relatively thin — or even non-existent — in the ocean ridge areas (Figure 10.3.5). The seismic studies also showed that the crust is relatively thin under the oceans (5 km to 6 km) compared to the continents (30 km to 60 km) and geologically very consistent, composed almost entirely of basalt.

Figure 10.3.5 Topographic section at an ocean ridge based on reflection seismic data. Sediments are not thick enough to be detectable near the ridge, but get thicker on either side. The diagram represents approximately 50 km width, and has a 10x vertical exaggeration.

In the early 1950s, Edward Bullard, who spent time at the University of Toronto but is mostly associated with Cambridge University, developed a probe for measuring the flow of heat from the ocean floor. Bullard and colleagues found the rate to be higher than average along the ridges, and lower than average in the trench areas. Although Bullard was a plate-tectonics sceptic, these features were interpreted to indicate that there is convection within the mantle — the areas of high heat flow being correlated with upward convection of hot mantle material, and the areas of low heat flow being correlated with downward convection.

With developments of networks of seismographic stations in the 1950s, it became possible to plot the locations and depths of both major and minor earthquakes with great accuracy. It was found that there is a remarkable correspondence between earthquakes and both the mid-ocean ridges and the deep ocean trenches. In 1954 Gutenberg and Richter showed that the ocean-ridge earthquakes were all relatively shallow, and confirmed what had first been shown by Benioff in the 1930s — that earthquakes in the vicinity of ocean trenches were both shallow and deep, but that the deeper ones were situated progressively farther inland from the trenches (Figure 10.3.6).

Figure 10.3.6 Cross-section through the Aleutian subduction zone with a depiction of the increasing depth of earthquakes “inshore” from the trench. [Image Description]

In the 1950s, scientists from the Scripps Oceanographic Institute in California persuaded the U.S. Coast Guard to include magnetometer readings on one of their expeditions to study ocean floor topography. The first comprehensive magnetic data set was compiled in 1958 for an area off the coast of B.C. and Washington State. This survey revealed a bewildering pattern of low and high magnetic intensity in sea-floor rocks (Figure 10.3.7). When the data were first plotted on a map in 1961, nobody understood them — not even the scientists who collected them. Although the patterns made even less sense than the stripes on a zebra, many thousands of kilometres of magnetic surveys were conducted over the next several years.

Figure 10.3.7 Pattern of sea-floor magnetism off of the west coast of British Columbia and Washington.

The wealth of new data from the oceans began to significantly influence geological thinking in the 1960s. In 1960, Harold Hess, a widely respected geologist from Princeton University, advanced a theory with many of the elements that we now accept as plate tectonics. He maintained some uncertainty about his proposal however, and in order to deflect criticism from mainstream geologists, he labelled it geopoetry. In fact, until 1962, Hess didn’t even put his ideas in writing—except internally to the U.S. Navy (which funded his research)—but presented them mostly in lectures and seminars. Hess proposed that new sea floor was generated from mantle material at the ocean ridges, and that old sea floor was dragged down at the ocean trenches and re-incorporated into the mantle. He suggested that the process was driven by mantle convection currents, rising at the ridges and descending at the trenches (Figure 10.3.8). He also suggested that the less-dense continental crust did not descend with oceanic crust into trenches, but that colliding land masses were thrust up to form mountains. Hess’s theory formed the basis for our ideas on sea-floor spreading and continental drift, but it did not deal with the concept that the crust is made up of specific plates. Although the Hess model was not roundly criticized, it was not widely accepted (especially in the U.S.), partly because it was not well supported by hard evidence.

Figure 10.3.8 A representation of Harold Hess’s model for sea-floor spreading and subduction.

Collection of magnetic data from the oceans continued in the early 1960s, but still nobody could explain the origin of the zebra-like patterns. Most assumed that they were related to variations in the composition of the rocks—such as variations in the amount of magnetite—as this is a common explanation for magnetic variations in rocks of the continental crust. The first real understanding of the significance of the striped anomalies was the interpretation by Fred Vine, a Cambridge graduate student. Vine was examining magnetic data from the Indian Ocean and, like others before, he noted the symmetry of the magnetic patterns with respect to the oceanic ridge.

At the same time, other researchers, led by groups in California and New Zealand, were studying the phenomenon of reversals in Earth’s magnetic field. They were trying to determine when such reversals had taken place over the past several million years by analyzing the magnetic characteristics of hundreds of samples from basaltic flows. As discussed in Chapter 9, it is evident that Earth’s magnetic field becomes weakened periodically and then virtually non-existent, before becoming re-established with the reverse polarity. During periods of reversed polarity, a compass would point south instead of north.

The time scale of magnetic reversals is irregular. For example, the present “normal” event, known as the Bruhnes magnetic chron, has persisted for about 780,000 years. This was preceded by a 190,000-year reversed event; a 50,000-year normal event known as Jaramillo; and then a 700,000-year reversed event (see Figure 9.3.3).

In a paper published in September 1963, Vine and his PhD supervisor Drummond Matthews proposed that the patterns associated with ridges were related to the magnetic reversals, and that oceanic crust created from cooling basalt during a normal event would have polarity aligned with the present magnetic field, and thus would produce a positive anomaly (a black stripe on the sea-floor magnetic map), whereas oceanic crust created during a reversed event would have polarity opposite to the present field and thus would produce a negative magnetic anomaly (a white stripe). The same idea had been put forward a few months earlier by Lawrence Morley, of the Geological Survey of Canada; however, his papers submitted earlier in 1963 to Nature and The Journal of Geophysical Research were rejected. Many people refer to the idea as the Vine-Matthews-Morley (VMM) hypothesis.

Vine, Matthews, and Morley were the first to show this type of correspondence between the relative widths of the stripes and the periods of the magnetic reversals. The VMM hypothesis was confirmed within a few years when magnetic data were compiled from spreading ridges around the world. It was shown that the same general magnetic patterns were present straddling each ridge, although the widths of the anomalies varied according to the spreading rates characteristic of the different ridges. It was also shown that the patterns corresponded with the chronology of Earth’s magnetic field reversals. This global consistency provided strong support for the VMM hypothesis and led to rejection of the other explanations for the magnetic anomalies.

In 1963, J. Tuzo Wilson of the University of Toronto proposed the idea of a mantle plume or hot spot—a place where hot mantle material rises in a stationary and semi-permanent plume, and affects the overlying crust. He based this hypothesis partly on the distribution of the Hawaiian and Emperor Seamount island chains in the Pacific Ocean (Figure 10.3.9). The volcanic rock making up these islands gets progressively younger toward the southeast, culminating with the island of Hawaii itself, which consists of rock that is almost all younger than 1 Ma. Wilson suggested that a stationary plume of hot upwelling mantle material is the source of the Hawaiian volcanism, and that the ocean crust of the Pacific Plate is moving toward the northwest over this hot spot. Near the Midway Islands, the chain takes a pronounced change in direction, from northwest-southeast for the Hawaiian Islands and to nearly north-south for the Emperor Seamounts. This change is widely ascribed to a change in direction of the Pacific Plate moving over the stationary mantle plume, but a more plausible explanation is that the Hawaiian mantle plume has not actually been stationary throughout its history, and in fact moved at least 2,000 km south over the period between 81 and 45 Ma.J. A. Tarduno et al., 2003, The Emperor Seamounts: Southward Motion of the Hawaiian Hotspot Plume in Earth’s Mantle, Science 301 (5636): 1064–1069.

Figure 10.3.9 The ages of the Hawaiian Islands and the Emperor Seamounts in relation to the location of the Hawaiian mantle plume.

Exercise 10.2 Volcanoes and the Rate of Plate Motion

The Hawaiian and Emperor volcanoes shown in Figure 10.3.9 are listed in the table below along with their ages and their distances from the centre of the mantle plume under Hawaii (the Big Island).

Ages of Hawaiian and Emperor volcanoes and their distances from the centre of the mantle plume. Calculate their rate of movement in centimetres per year. 
Island Age Distance Rate
Hawaii 0 Ma 0 km
Necker 10.3 Ma 1,058 km 10.2 cm/y
Midway 27.7 Ma 2,432 km
Koko 48.1 Ma 3,758 km
Suiko 64.7 Ma 4,860 km

Plot the data on the graph provided here, and use the numbers in the table to estimate the rates of plate motion for the Pacific Plate in cm/year. (The first two are plotted for you.)

A blank graph. Distance (in kilometres) on y-axis. Age (Ma) on x-axis

See Appendix 3 for Exercise 10.2 answers.

There is evidence of many such mantle plumes around the world (Figure 10.3.10). Most are within the ocean basins—including places like Hawaii, Iceland, and the Galapagos Islands—but some are under continents. One example is the Yellowstone hot spot in the west-central United States, and another is the one responsible for the Anahim Volcanic Belt in central British Columbia. It is evident that mantle plumes are very long-lived phenomena, lasting for at least tens of millions of years, possibly for hundreds of millions of years in some cases.

Figure 10.3.10 Mantle plume locations. Selected Mantle plumes: 1: Azores, 3: Bowie, 5: Cobb, 8: Eifel, 10: Galapagos, 12: Hawaii, 14: Iceland, 17: Cameroon, 18: Canary, 19: Cape Verde, 35: Samoa, 38: Tahiti, 42: Tristan, 44: Yellowstone, 45: Anahim

Although oceanic spreading ridges appear to be curved features on Earth’s surface, in fact the ridges are composed of a series of straight-line segments, offset at intervals by faults perpendicular to the ridge (Figure 10.3.11). In a paper published in 1965, Tuzo Wilson termed these features transform faults. He described the nature of the motion along them, and showed why there are earthquakes only on the section of a transform fault between two adjacent ridge segments. The San Andreas Fault in California is a very long transform fault that links the southern end of the Juan de Fuca spreading ridge to the East Pacific Rise spreading ridges situated in the Gulf of California (see Figure 10.4.9). The Queen Charlotte Fault, which extends north from the northern end of the Juan de Fuca spreading ridge (near the northern end of Vancouver Island) toward Alaska, is also a transform fault.

Figure 10.3.11 A part of the mid-Atlantic ridge near the equator. The double white lines are spreading ridges. The solid white lines are fracture zones. As shown by the yellow arrows, the relative motion of the plates on either side of the fracture zones can be similar (arrows pointing the same direction) or opposite (arrows pointing opposite directions). Transform faults (red lines) are in between the ridge segments, where the yellow arrows point in opposite directions.

In the same 1965 paper, Wilson introduced the idea that the crust can be divided into a series of rigid plates, and thus he is responsible for the term plate tectonics.

Exercise 10.3 Paper transform fault model

Figure 10.3.12

Tuzo Wilson used a paper model, a little bit like the one shown here, to explain transform faults to his colleagues. To use this model either print this page or download the image above and print that, then cut around the outside, and then slice along the line A-B (the fracture zone) with a sharp knife. Fold down the top half where shown, and then pinch together in the middle. Do the same with the bottom half.

Figure 10.3.13

When you’re done, you should have something like the example shown on Figure 10.3.13, with two folds of paper extending underneath. Find someone else to pinch those folds with two fingers just below each ridge crest, and then gently pull apart where shown. As you do, the oceanic crust will emerge from the middle, and you’ll see that the parts of the fracture zone between the ridge crests will be moving in opposite directions (this is the transform fault) while the parts of the fracture zone outside of the ridge crests will be moving in the same direction. You’ll also see that the oceanic crust is being magnetized as it forms at the ridge. The magnetic patterns shown are accurate, and represent the last 2.5 Ma of geological time.

There are other versions of this model available here: Paper Models of Transform Faults.For more information see: Earle, S., 2004, A simple paper model of a transform fault at a spreading ridge, J. Geosc. Educ. V. 52, p. 391-2.

See Appendix 3 for Exercise 10.3 answers.

Image Descriptions

Figure 10.3.2 image description: At 500 Ma, rocks in Europe had upward-pointing magnetic orientations. At 400 Ma, the magnetic orientation leveled. From 300 Ma to the present, rocks in Europe shown an increasingly downward-pointing magnetic orientation. [Return to Figure 10.3.2]

Figure 10.3.6 image description: A cross section of the trench formed at the Aleutian subduction zone as the Pacific plate subducts under the North American plate in the middle of the Pacific Ocean. The farther away an earthquake is from this trench (on the North America plate side), the deeper it is. [Return to Figure 10.3.6]

Media Attributions

  • Figures 10.3.1, 10.3.2, 10.3.3, 10.3.5, 10.3.6, 10.3.8, 10.3.11, 10.3.12, 10.3.13: © Steven Earle. CC BY.
  • Figure 10.3.4: “Elevation” by NOAA. Adapted by Steven Earle. Public domain.
  • Figure 10.3.7: “Juan de Fuca Ridge” by USGS. Adapted by Steven Earle. Public domain. Based on Raff, A. and Mason, R., 1961, Magnetic survey off the west coast of North America, 40˚ N to 52˚ N latitude, Geol. Soc. America Bulletin, V. 72, p. 267-270.
  • Figure 10.3.9: “Hawaii Hotspot” by National Geophysical Data Center. Adapted by Steven Earle. Public domain.
  • Figure 10.3.10: “Hotspots” by Ingo Wölbern. Public domain.


10.4 Plate, Plate Motions, and Plate Boundary Processes

Continental drift and sea-floor spreading became widely accepted around 1965 as more and more geologists started thinking in these terms. By the end of 1967 the Earth’s surface had been mapped into a series of plates (Figure 10.4.1). The major plates are Eurasia, Pacific, India, Australia, North America, South America, Africa, and Antarctic. There are also numerous small plates (e.g., Juan de Fuca, Cocos, Nazca, Scotia, Philippine, Caribbean), and many very small plates or sub-plates. For example the Juan de Fuca Plate is actually three separate plates (Gorda, Juan de Fuca, and Explorer) that all move in the same general direction but at slightly different rates.

Figure 10.4.1 A map showing 15 of the Earth’s tectonic plates and the approximate rates and directions of plate motions. [Image Description]

Rates of motions of the major plates range from less than 1 cm/y to over 10 cm/y. The Pacific Plate is the fastest, followed by the Australian and Nazca Plates. The North American Plate is one of the slowest, averaging around 1 cm/y in the south up to almost 4 cm/y in the north.

Plates move as rigid bodies, so it may seem surprising that the North American Plate can be moving at different rates in different places. The explanation is that plates move in a rotational manner. The North American Plate, for example, rotates counter-clockwise; the Eurasian Plate rotates clockwise.

Boundaries between the plates are of three types: divergent (i.e., moving apart), convergent (i.e., moving together), and transform (moving side by side). Before we talk about processes at plate boundaries, it’s important to point out that there are never gaps between plates. The plates are made up of crust and the lithospheric part of the mantle (Figure 10.4.2), and even though they are moving all the time, and in different directions, there is never a significant amount of space between them. Plates are thought to move along the lithosphere-asthenosphere boundary, as the asthenosphere is the zone of partial melting. It is assumed that the relative lack of strength of the partial melting zone facilitates the sliding of the lithospheric plates.

Figure 10.4.2 The crust and upper mantle. Tectonic plates consist of lithosphere, which includes the crust and the lithospheric (rigid) part of the mantle.

At spreading centres, the lithospheric mantle may be very thin because the upward convective motion of hot mantle material generates temperatures that are too high for the existence of a significant thickness of rigid lithosphere (Figure 10.3.8). The fact that the plates include both crustal material and lithospheric mantle material makes it possible for a single plate to be made up of both oceanic and continental crust. For example, the North American Plate includes most of North America, plus half of the northern Atlantic Ocean. Similarly the South American Plate extends across the western part of the southern Atlantic Ocean, while the European and African plates each include part of the eastern Atlantic Ocean. The Pacific Plate is almost entirely oceanic, but it does include the part of California west of the San Andreas Fault.

Divergent Boundaries

Divergent boundaries are spreading boundaries, where new oceanic crust is created from magma derived from partial melting of the mantle caused by decompression as hot mantle rock from depth is moved toward the surface (Figure 10.4.3). The triangular zone of partial melting near the ridge crest is approximately 60 km thick and the proportion of magma is about 10% of the rock volume, thus producing crust that is about 6 km thick. Most divergent boundaries are located at the oceanic ridges (although some are on land), and the crustal material created at a spreading boundary is always oceanic in character; in other words, it is mafic igneous rock (e.g., basalt or gabbro, rich in ferromagnesian minerals). Spreading rates vary considerably, from 2 cm/y to 6 cm/y in the Atlantic, to between 12 cm/y and 20 cm/y in the Pacific. (Note that spreading rates are typically double the velocities of the two plates moving away from a ridge.)

Some of the processes taking place in this setting include:

Figure 10.4.3 The general processes that take place at a divergent boundary. The area within the dashed white rectangle is shown in Figure 10.4.4.
Figure 10.4.4 Depiction of the processes and materials formed at a divergent boundary.

Spreading is hypothesized to start within a continental area with up-warping or doming related to an underlying mantle plume or series of mantle plumes. The buoyancy of the mantle plume material creates a dome within the crust, causing it to fracture in a radial pattern, with three arms spaced at approximately 120° (Figure 10.4.5). When a series of mantle plumes exists beneath a large continent, the resulting rifts may align and lead to the formation of a rift valley (such as the present-day Great Rift Valley in eastern Africa). It is suggested that this type of valley eventually develops into a linear sea (such as the present-day Red Sea), and finally into an ocean (such as the Atlantic). It is likely that as many as 20 mantle plumes, many of which still exist, were responsible for the initiation of the rifting of Pangea along what is now the mid-Atlantic ridge (see Figure 10.3.10).

Figure 10.4.5 Depiction of the process of dome and three-part rift formation (left) and of continental rifting between the African and South American parts of Pangea at around 200 Ma (right).

Convergent Boundaries

Convergent boundaries, where two plates are moving toward each other, are of three types, depending on whether oceanic or continental crust is present on either side of the boundary. The types are ocean-ocean, ocean-continent, and continent-continent.

At an ocean-ocean convergent boundary, one of the plates (oceanic crust and lithospheric mantle) is pushed, or subducted, under the other. Often it is the older and colder plate that is denser and subducts beneath the younger and hotter plate. There is commonly an ocean trench along the boundary. The subducted lithosphere descends into the hot mantle at a relatively shallow angle close to the subduction zone, but at steeper angles farther down (up to about 45°). As discussed in the context of subduction-related volcanism in Chapter 4, the significant volume of water within the subducting material is released as the subducting crust is heated. Most of this water is present within the sheet silicate mineral serpentine which is derived from alteration of pyroxene and olivine near the spreading ridge shortly after the rock’s formation. It is released when the oceanic crust is heats and then rises and mixes with the overlying mantle. The addition of water to the hot mantle lowers the rocks’s melting point and leads to the formation of magma (flux melting) (Figure 10.4.6). The magma, which is lighter than the surrounding mantle material, rises through the mantle and the overlying oceanic crust to the ocean floor where it creates a chain of volcanic islands known as an island arc. A mature island arc develops into a chain of relatively large islands (such as Japan or Indonesia) as more and more volcanic material is extruded and sedimentary rocks accumulate around the islands.

As described above in the context of Benioff zones (Figure 10.3.6), earthquakes take place close to the boundary between the subducting crust and the overriding crust. The largest earthquakes occur near the surface where the subducting plate is still cold and strong.

Figure 10.4.6 Configuration and processes of an ocean-ocean convergent boundary.

Examples of ocean-ocean convergent zones are subduction of the Pacific Plate beneath the North America Plate south of Alaska (Aleutian Islands) and beneath the Philippine Plate west of the Philippines, subduction of the India Plate beneath the Eurasian Plate south of Indonesia, and subduction of the Atlantic Plate beneath the Caribbean Plate (see Figure 10.4.1).

At an ocean-continent convergent boundary, the oceanic plate is pushed under the continental plate in the same manner as at an ocean-ocean boundary. Sediment that has accumulated on the continental slope is thrust up into an accretionary wedge, and compression leads to thrusting within the continental plate (Figure 10.4.7). The mafic magma produced adjacent to the subduction zone rises to the base of the continental crust and leads to partial melting of the crustal rock. The resulting magma ascends through the crust, producing a mountain chain with many volcanoes.

Figure 10.4.7 Configuration and processes of an ocean-continent convergent boundary.

Examples of ocean-continent convergent boundaries are subduction of the Nazca Plate under South America (which has created the Andes Range) and subduction of the Juan de Fuca Plate under North America (creating the mountains Garibaldi, Baker, St. Helens, Rainier, Hood, and Shasta, collectively known as the Cascade Range).

A continent-continent collision occurs when a continent or large island that has been moved along with subducting oceanic crust collides with another continent (Figure 10.4.8). The colliding continental material will not be subducted because it is too light (i.e., because it is composed largely of light continental rocks [SIAL]), but the root of the oceanic plate will eventually break off and sink into the mantle. There is tremendous deformation of the pre-existing continental rocks, and creation of mountains from that rock, from any sediments that had accumulated along the shores (i.e., within geosynclines) of both continental masses, and commonly also from some ocean crust and upper mantle material.

Figure 10.4.8 Configuration and processes of a continent-continent convergent boundary.

Examples of continent-continent convergent boundaries are the collision of the India Plate with the Eurasian Plate, creating the Himalaya Mountains, and the collision of the African Plate with the Eurasian Plate, creating the series of ranges extending from the Alps in Europe to the Zagros Mountains in Iran. The Rocky Mountains in B.C. and Alberta are also a result of continent-continent collisions.

Figure 10.4.9 The San Andreas Fault extends from the north end of the East Pacific Rise in the Gulf of California to the southern end of the Juan de Fuca Ridge. All of the red lines on this map are transform faults.

Transform boundaries exist where one plate slides past another without production or destruction of crustal material. As explained above, most transform faults connect segments of mid-ocean ridges and are thus ocean-ocean plate boundaries (Figure 10.3.11). Some transform faults connect continental parts of plates. An example is the San Andreas Fault, which extends from the southern end of the Juan de Fuca Ridge to the northern end of the East Pacific Rise (ridge) in the Gulf of California (Figures 10.28 and 10.29). The part of California west of the San Andreas Fault and all of Baja California are on the Pacific Plate. Transform faults do not just connect divergent boundaries. For example, the Queen Charlotte Fault connects the north end of the Juan de Fuca Ridge, starting at the north end of Vancouver Island, to the Aleutian subduction zone.

A bridge stretches across the San Andreas Fault between the Pacific and the North American Plates.
Figure 10.4.10 The San Andreas Fault at Parkfield in central California. The person with the orange shirt is standing on the Pacific Plate and the person at the far side of the bridge is on the North American Plate. The bridge is designed to accommodate motion on the fault by sliding on its foundation.

Exercise 10.4 A different type of transform fault

Figure 10.4.11

This map shows the Juan de Fuca (JDF) and Explorer Plates off the coast of Vancouver Island. We know that the JDF Plate is moving toward the North American Plate at around 4 centimetres per year to 5 centimetres per year. We think that the Explorer Plate is also moving east, but we don’t know the rate, and there is evidence that it is slower than the JDF Plate.

The boundary between the two plates is the Nootka Fault, which is the location of frequent small-to-medium earthquakes (roughly up to magnitude 5), as depicted by the red stars. Explain why the Nootka Fault is a transform fault, and show the relative sense of motion along the fault with two small arrows.

See Appendix 3 for Exercise 10.4 answers.

As originally described by Wegener in 1915, the present continents were once all part of a supercontinent, which he termed Pangea (meaning all land). More recent studies of continental matchups and the magnetic ages of ocean-floor rocks have enabled us to reconstruct the history of the break-up of Pangea.

Pangea began to rift apart along a line between Africa and Asia and between North America and South America at around 200 Ma. During the same period, the Atlantic Ocean began to open up between northern Africa and North America, and India broke away from Antarctica. Between 200 and 150 Ma, rifting started between South America and Africa and between North America and Europe, and India moved north toward Asia. By 80 Ma, Africa had separated from South America, most of Europe had separated from North America, and India had separated from Antarctica. By 50 Ma, Australia had separated from Antarctic, and shortly after that, India collided with Asia. To see the timing of these processes for yourself, go to time lapse of Continental Movements.

Within the past few million years, rifting has taken place in the Gulf of Aden and the Red Sea, and also within the Gulf of California. Incipient rifting has begun along the Great Rift Valley of eastern Africa, extending from Ethiopia and Djibouti on the Gulf of Aden (Red Sea) all the way south to Malawi.

Over the next 50 million years, it is likely that there will be full development of the east African rift and creation of new ocean floor. Eventually Africa will split apart. There will also be continued northerly movement of Australia and Indonesia. The western part of California (including Los Angeles and part of San Francisco) will split away from the rest of North America, and eventually sail right by the west coast of Vancouver Island, en route to Alaska. Because the oceanic crust formed by spreading on the mid-Atlantic ridge is not currently being subducted (except in the Caribbean), the Atlantic Ocean is slowly getting bigger, and the Pacific Ocean is getting smaller. If this continues without changing for another couple hundred million years, we will be back to where we started, with one supercontinent.

Pangea, which existed from about 350 to 200 Ma, was not the first supercontinent. It was preceded by Pannotia (600 to 540 Ma), by Rodinia (1,100 to 750 Ma), and by others before that.

In 1966, Tuzo Wilson proposed that there has been a continuous series of cycles of continental rifting and collision; that is, break-up of supercontinents, drifting, collision, and formation of other supercontinents. At present, North and South America, Europe, and Africa are moving with their respective portions of the Atlantic Ocean. The eastern margins of North and South America and the western margins of Europe and Africa are called passive margins because there is no subduction taking place along them.

This situation may not continue for too much longer, however. As the Atlantic Ocean floor gets weighed down around its margins by great thickness of continental sediments (i.e., geosynclines), it will be pushed farther and farther into the mantle, and eventually the oceanic lithosphere may break away from the continental lithosphere (Figure 10.4.12). A subduction zone will develop, and the oceanic plate will begin to descend under the continent. Once this happens, the continents will no longer continue to move apart because the spreading at the mid-Atlantic ridge will be taken up by subduction. If spreading along the mid-Atlantic ridge continues to be slower than spreading within the Pacific Ocean, the Atlantic Ocean will start to close up, and eventually (in a 100 million years or more) North and South America will collide with Europe and Africa.

Figure 10.4.12 Development of a subduction zone at a passive margin. Times A, B, and C are separated by tens of millions of years. Once the oceanic crust breaks off and starts to subduct the continental crust (North America in this case) will no longer be pushed to the west and will likely start to move east because the rate of spreading in the Pacific basin is faster than that in the Atlantic basin.

There is strong evidence around the margins of the Atlantic Ocean that this process has taken place before. The roots of ancient mountain belts, which are present along the eastern margin of North America, the western margin of Europe, and the northwestern margin of Africa, show that these land masses once collided with each other to form a mountain chain, possibly as big as the Himalayas. The apparent line of collision runs between Norway and Sweden, between Scotland and England, through Ireland, through Newfoundland, and the Maritimes, through the northeastern and eastern states, and across the northern end of Florida. When rifting of Pangea started at approximately 200 Ma, the fissuring was along a different line from the line of the earlier collision. This is why some of the mountain chains formed during the earlier collision can be traced from Europe to North America and from Europe to Africa.

That the Atlantic Ocean rift may have occurred in approximately the same place during two separate events several hundred million years apart is probably no coincidence. The series of hot spots that has been identified in the Atlantic Ocean may also have existed for several hundred million years, and thus may have contributed to rifting in roughly the same place on at least two separate occasions (Figure 10.3.13).

Figure 10.4.13 A scenario for the Wilson cycle. (A) The cycle starts with continental rifting above a series of mantle plumes. (B) The continents separate, and then (C) re-converge some time later, forming a fold-belt mountain chain. (D) Eventually rifting is repeated, possibly because of the same set of mantle plumes, but this time the rift is in a different place.

Exercise 10.5 Getting to know the plates and their boundaries

This map shows the boundaries between the major plates. Without referring to the plate map in Figure 10.4.1, or any other resources, write in the names of as many of the plates as you can. Start with the major plates, and then work on the smaller ones. Don’t worry if you can’t name them all.

Figure 10.4.14

Once you’ve named most of the plates, draw arrows to show the general plate motions. Finally, using a highlighter or coloured pencil, label as many of the boundaries as you can as divergent, convergent, or transform.

See Appendix 3 for Exercise 10.5 answers.

Image descriptions

Figure 10.4.1 image description: Descriptions of 15 different plates and their movements.
Plate name Description of plate Bordering plates (ordered from longest border to shortest) Description of movement
Africa plate This plate includes all of Africa and the surrounding ocean, including the eastern Atlantic Ocean, the surrounding Antarctic Ocean, and the western Indian ocean. South America plate, Antarctic plate, Eurasia plate, North America plate, Arabia plate, India plate, Australia plate This plate is moving north east towards the Arabia and Eurasia plates.
Antarctic plate. This plate makes up all of Antarctica and much of the surrounding ocean. Pacific plate, Australia plate, Africa plate, Scotia plate, Nazca plate, South America plate. The part of the plate around the South America plate is moving northwards and a little east. The part of the plate around the Australia plate is moving southwards.
Arabia plate This plate includes all of Saudi Arabia, and much of the Levant (up to Iraq and Syria). Eurasia plate, Africa plate, India plate This plate is moving north east towards the Eurasia plate.
Australia plate This plate includes Australia and much of the surrounding ocean. New Guinea and the northern parts of New Zealand are part of the Australia plate. The ocean area along southern Asia  up to the India plate is also a part of the Australia plate. Antarctic plate, Pacfic plate, Eurasia plate, India Plate, Africa plate. This plate is moving north east towards the Eurasia plate and the Pacific plate.
Caribbean plate This plate is small. It includes the central Caribbean countries and runs along the northern edge of South America. North America plate, South America plate, Cocos plate. N/A
Cocos plate This plate is small. It runs along the west coast of Mexico and western Caribbean countries. Nazca plate, Pacific plate, North America plate, Caribbean plate. This plate is moving north east towards the Caribbean and North America plates.
Eurasia plate This plate includes the northeastern part of the Atlantic ocean, all of Europe, all of Russia (except its most eastern part), and down through southeast Asia, including China and Indonesia. North America plate, Africa plate, Australia plate, Arabia plate, India plate, Filipino plate. This plate is rotating in a clockwise direction towards the Pacific plate.
Filipino plate This plate includes the islands that make up the Philipines and north to include parts of southern Japan. Eurasia plate, Pacific plate. This plate is moving north west towards the Eurasia plate.
India plate This plate includes India and the surrounding India Ocean. Australia plate, Eurasia plate, Africa plate, Arabia plate. This plate is moving north north east towards the Eurasia plate.
Juan de Fuca plate This plate is small. It runs along the north western coast of the United States and the southern British Columbia coast. Pacific plate, North America plate. N/A
Nazca plate This plate is in the Pacific Ocean between the Pacific plate and the South America plate. South America plate, Pacific plate, Antarctic plate, Cocos plate This plate is moving directly east towards the South America plate.
North America Plate This plate includes all of North America, Greenland, the eastern most part of Russia, northern Japan, and the northwestern part of the Atlantic ocean. Eurasia plate, Pacific plate, Africa plate, Caribbean plate, South America plate, Cocos plate, Juan de Fuca plate This plate is rotating counter clockwise in towards the Pacific plate.
Pacific plate This plate makes up most of the Pacific Ocean. North America plate, Australia plate,  Antarctic plate, Nazca plate, Filipino plate, Cocos plate, Juan de Fuca plate This plate is moving northwest towards the Australia, Filipino, and Eurasia plates.
Scotia plate This plate is small. It runs from the tip of South America eastwards to form a barrier between the Antarctic plate and the South America plate. Antarctic plate, South America plate. N/A
South America plate This plate starts at the western edge of South America and stretches east into the southwestern parst of the Atlantic Ocean. Africa plate, Nazca plate, Scotia plate, Caribbean plate, Antarctic plate, North America plate This plate moves north and slightly west towards the Caribbean plate and the North America plate.

[Return to Figure 10.4.1]

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10.5 Mechanisms for Plate Motion

It has been often repeated in this text and elsewhere that convection of the mantle is critical to plate tectonics, and while this is almost certainly so, other forces likely play a significant role. One side in the argument holds that the plates are only moved by the traction caused by mantle convection. The other side holds that traction plays only a minor role and that two other forces, ridge-push and slab-pull, are more important (Figure 10.5.1). Some argue that the real answer lies somewhere in between.

Figure 10.5.1 Models for plate motion mechanisms. [Image Description]

Kearey and Vine (1996)Kearey and Vine , 1996, Global Tectonics (2ed), Blackwell Science Ltd., Oxford have listed some compelling arguments in favour of the ridge-push/slab-pull model, as follows: (a) plates that are attached to subducting slabs (e.g., Pacific, Australian, and Nazca Plates) move the fastest, and plates that are not (e.g., North American, South American, Eurasian, and African Plates) move significantly slower; (b) in order for the traction model to apply, the mantle would have to be moving about five times faster than the plates are moving (because the coupling between the partially liquid asthenosphere and the plates is not strong), and such high rates of convection are not supported by geophysical models; and (c) although large plates have potential for much higher convection traction, plate velocity is not related to plate area.

In the ridge-push/slab-pull model, which is the one that has been adopted by most geologists working on plate-tectonic problems, the lithosphere is the upper surface of the convection cells, as is illustrated in Figure 10.5.2.

Figure 10.5.2 The ridge-push/slab-pull model for plate motion, in which the lithosphere is the upper surface of the convective systems.

Although ridge-push/slab-pull is the widely favoured mechanism for plate motion, it’s important not to underestimate the role of mantle convection. Without convection, there would be no ridges to push from because upward convection brings hot buoyant rock to surface. Furthermore, many plates, including our own North American Plate, move along nicely—albeit slowly—without any slab-pull happening.

Image Descriptions

Figure 10.5.1 image description: In this model, there are three forces working to move the plates. Ridge-push forces cause two plates to pull apart on the surface. Slab-pull forces pull the plates down. This movement of out and down is also encouraged by convection traction, or clockwise and counterclockwise currents that are present beneath the plates. [Return to Figure 10.5.1]

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

Section Summary
10.1 Alfred Wegener: The Father of Plate Tectonics The evidence for continental drift in the early 20th century included the matching of continental shapes on either side of the Atlantic and the geological and fossil matchups between continents that are now thousands of kilometres apart.
10.2 Global Geological Models of the Early 20th Century The established theories of global geology were permanentism and contractionism, but neither of these theories was able to explain some of the evidence that supported the idea of continental drift.
10.3 Geological Renaissance of the Mid-20th Century Giant strides were made in understanding Earth during the middle decades of the 20th century, including discovering magnetic evidence of continental drift, mapping the topography of the ocean floor, describing the depth relationships of earthquakes along ocean trenches, measuring heat flow differences in various parts of the ocean floor, and mapping magnetic reversals on the sea floor. By the mid-1960s, the fundamentals of the theory of plate tectonics were in place.
10.4 Plate, Plate Motions, and Plate Boundary Processes Earth’s lithosphere is made up of over 20 plates that are moving in different directions at rates of between 1 cm/y and 10 cm/y. The three types of plate boundaries are divergent (plates moving apart and new crust forming), convergent (plates moving together and one being subducted), and transform (plates moving side by side). Divergent boundaries form where existing plates are rifted apart, and it is hypothesized that this is caused by a series of mantle plumes. Subduction zones are assumed to form where accumulation of sediment at a passive margin leads to separation of oceanic and continental lithosphere. Supercontinents form and break up through these processes.
10.5 Mechanisms for Plate Motion It is widely believed that ridge-push and slab-pull are the main mechanisms for plate motion, as opposed to traction by mantle convection. Mantle convection is a key factor for producing the conditions necessary for ridge-push and slab-pull.

Questions for Review

  1. List some of the evidence used by Wegener to support his idea of moving continents.
  2. What was the primary technical weakness with Wegener’s continental drift theory?
  3. How were mountains thought to be formed (a) by contractionists and (b) by permanentists?
  4. How were the trans-Atlantic paleontological matchups explained in the late 19th century?
  5. In the context of isostasy, what would prevent an area of continental crust from becoming part of an ocean?
  6. How did we learn about the topography of the sea floor in the early part of the 20th century?
  7. How does the temperature profile of the crust and the mantle indicate that part of the mantle must be convecting?
  8. What evidence from paleomagnetic studies provided support for continental drift?
  9. Which parts of the oceans are the deepest?
  10. Why is there less sediment in the ocean ridge areas than in other parts of the sea floor?
  11. How were the oceanic heat-flow data related to mantle convection?
  12. Describe the spatial and depth distribution of earthquakes at ocean ridges and ocean trenches.
  13. In the model for ocean basins developed by Harold Hess, what took place at oceanic ridges and what took place at oceanic trenches?
    Figure A
  14. What aspect of plate tectonics was not included in the Hess theory?
  15. Figure 10.36 shows the pattern of sea-floor magnetic anomalies in the area of a spreading ridge. Draw in the likely location of the ridge.
  16. What is a mantle plume and what is its expected lifespan?
  17. Describe the nature of movement at an ocean ridge transform fault (a) between the ridge segments, and (b) outside the ridge segments.
  18. How is it possible for a plate to include both oceanic and continental crust?
  19. What is the likely relationship between mantle plumes and the development of a continental rift?
  20. Why does subduction not take place at a continent-continent convergent zone?
  21. Divergent, convergent, and transform boundaries are shown in different colours on Fiugre 10.37. Which colours are the divergent boundaries, which are the convergent boundaries, and which are the transform boundaries?
  22. Name the plates on this map and show their approximate motion directions.
  23. Show the sense of motion on either side of the plate boundary to the west of Haida Gwaii (Queen Charlotte Islands).
  24. Where are Earth’s most recent sites of continental rifting and creation of new ocean floor?
  25. What is likely to happen to western California over the next 50 million years?
  26. What geological situation might eventually lead to the generation of a subduction zone at a passive ocean-continent boundary such as the eastern coast of North America?

Answers to Review Questions can be found in Appendix 2.

Image Descriptions

Figure B image description: A black line with triangles pointing towards the coast stretches from the Oregon and Washington state up just past Vancouver Island to the southern tip of Haida Gwaii. This line also appears along the Alaskan coast and stretches part way down the Alaskan Pan-Handle. A thin red line stretches from the Alaskan Pan-Handle down just past the southern tip of Haida Gwaii. From that point, it alternates from being a thin red to a thick blue line to form uneven angles zig zagging south past Oregon state. [Return to Figure B]

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Chapter 11 Earthquakes

Learning Objectives

After carefully reading this chapter, completing the exercises within it, and answering the questions at the end, you should be able to:

  • Explain how the principle of elastic deformation applies to earthquakes.
  • Describe how the main shock and the immediate aftershocks define the rupture surface of an earthquake, and explain how stress transfer is related to aftershocks.
  • Explain the process of episodic tremor and slip.
  • Describe the relationship between earthquakes and plate tectonics, including where we should expect earthquakes to happen at different types of plate boundaries and at what depths.
  • Distinguish between earthquake magnitude and intensity, and explain some of the ways of estimating magnitude.
  • Explain the importance of collecting intensity data following an earthquake.
  • Describe how earthquakes lead to the destruction of buildings and other infrastructure, fires, slope failures, liquefaction, and tsunami.
  • Discuss the value of earthquake forecasting, and describe some of the steps that governments and individuals can take to minimize the impacts of large earthquakes.
Figure 11.0.1 A schoolroom in Courtenay damaged by the 1946 Vancouver Island earthquake. If the earthquake had not happened on a Sunday, the casualties would have been much greater.

Earthquakes scare people … a lot! That’s not surprising because time and time again earthquakes have caused massive damage and many injuries and deaths. Anyone who has lived through a strong earthquake cannot forget the experience (Figure 11.0.1). But geoscientists and engineers are getting better at understanding earthquakes, minimizing the amount of damage they cause, and reducing the number of people affected. People living in western Canada don’t need to be frightened by earthquakes, but they do need to be prepared.

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Figure 11.0.1: “Courtenay B.C., Damage to interior of Elementary School” by Earthquakes Canada. Permission to reproduce for public, non-commercial purposes.


11.1 What is an Earthquake?

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 Earth’s surface.

A body of rock that is under stress becomes deformed. When the rock can no longer withstand the deformation, it breaks and the two sides slide past each other. Most earthquakes take place near plate boundaries—but not necessarily right on a boundary—and not necessarily even on a pre-existing fault.

The engineering principle of elastic deformation, which can be used to understand earthquakes, is illustrated in Figure 11.1.1. The stress applied to a rock—typically because of ongoing plate movement—results in strain or deformation of the rock (Figure 11.1.1b). 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 that limit is reached. At that point, in the case of rocks within the crust, the rock breaks and there is displacement along the rupture surface (Figure 11.1.1c). 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 11.1.1 Depiction of the concept of elastic deformation and rupture. The plate boundary is shown as a dashed red line. In b the two plates are moving as shown by the arrows, but they are locked against each along the plate boundary so they are both deforming and the rocks are stressed.  In c there has been a rupture along the boundary and the stress is released.

The concept of a rupture surface, which is critical to understanding earthquakes, is illustrated in Figure 11.1.2. 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 11.1.2), and, by definition, it decreases to zero at the edges of the rupture surface because the rock beyond that point isn’t displaced at all. The extent of a rupture surface and the amount of displacement will depend on a number of factors, including the type and strength of the rock, and the degree to which it was stressed beforehand.

Figure 11.1.2 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 part of the fault plane along which displacement occurred. In this example, the near side of the fault is moving to the left, and the lengths of the arrows within the rupture surface represent relative amounts of displacement.

Earthquake rupture doesn’t happen all at once; it starts at a single point and spreads rapidly from there. 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 (Figure 11.1.3). The initiation point isn’t necessarily in the centre of the rupture surface; it may be close to one end, near the top, or near the bottom.

Figure 11.1.3 Propagation of failure on a rupture surface. In this case, the failure starts at the dark blue heavy arrow in the centre and propagates outward, reaching the left side first (green arrows) and the right side last (yellow arrows).

Figure 11.1.4 shows the distribution of immediate aftershocks associated with the 1989 Loma Prieta earthquake. Panel (b) is a section along the San Andreas Fault; this view is equivalent to what is shown in Figures 11.1.2 and 11.1.3. The area of red dots is the rupture surface; each red dot is a specific aftershock that was recorded on a seismometer. The hexagon labelled “main earthquake” represents the first or main shock. When that 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 (aftershocks), in this case, over an area about 60 kilometres long and 15 kilometres wide.

Figure 11.1.4 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 exactly 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 in Figure 11.1.4 all happened within seconds or minutes of the main shock, but aftershocks can be delayed for hours, days, weeks, or even years. 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 30 kilometres along the fault in each direction and for 15 kilometres toward 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 modelled for numerous earthquakes and active faults around the world, is depicted in Figure 11.1.5. Stress was reduced in the area of the rupture (blue), but was increased at either end of the rupture surface (red and yellow).

Figure 11.1.5 Depiction of stress changes related to an earthquake. Stress decreases in the area of the rupture surface, but increases 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 may lead to increased stress on other faults in the region. 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 buildup of stress from plate motion and eventually leads to another earthquake.

Episodic Tremor and Slip

Episodic tremor and slip (ETS) 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). It was first discovered on the Vancouver Island part of the Cascadia subduction zone by Geological Survey of Canada geologists Herb Dragert and Garry Rogers.Rogers, G. and Dragert, H., 2003, Episodic tremor and slip on the Cascadia subduction zone: the chatter of silent slip, Science, V. 300, p. 1942-1943.

The boundary between the subducting Juan de Fuca Plate and the North America Plate can be divided into three segments (Figure 11.1.6). The cold upper part of the Juan de Fuca Plate boundary is locked. The plates are stuck and don’t move, except with very large earthquakes that happen approximately every 500 years (the last one was approximately M9 on January 26, 1700). The warm lower part of the boundary is sliding continuously because the warm rock is weaker. The central part of the boundary isn’t cold enough to be stuck, but isn’t warm enough to slide continuously. Instead it slips episodically, approximately every 14 months for about 2 weeks, moving a few centimetres each time.

Figure 11.1.6 The boundary between the subducting Juan de Fuca Plate and the North America Plate is locked in the upper part, slides continuously in the lower part and slides episodically in the middle part.

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 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, New Zealand and Japan.

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11.2 Earthquakes and Plate Tectonics

The distribution of earthquakes across the globe is shown in Figure 11.2.1. 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 at less than a 30 kilometre depth. Shallow earthquakes are also common along transform faults, such as the San Andreas Fault. Along subduction zones, as we saw in Chapter 10, earthquakes are very abundant, and they are increasingly deep on the landward side of the subduction zone.

Figure 11.2.1 General distribution of global earthquakes of magnitude 4 and greater from 2004 to 2011, colour coded by depth (red: 0 to 33 kilometres, orange 33 to 70 kilometres, green: 70 to 300 kilometres, blue: 300 to 700 kilometres).

Earthquakes are also relatively common at a few intraplate locations. Some are related to the buildup of stress due to continental rifting or the transfer of stress from other regions, and some are not well understood. Examples of intraplate earthquake regions include the Great Rift Valley area of Africa, the Tibet region of China, and the Lake Baikal area of Russia.

Earthquakes at Divergent and Transform Boundaries

Figure 11.2.2 provides a closer look at magnitude (M) 4 and larger earthquakes in an area of divergent boundaries in the mid-Atlantic region near the equator. Here, as we saw in Chapter 10, the segments of the mid-Atlantic ridge are offset by some long transform faults. Most of the earthquakes are located along the transform faults, rather than along the spreading segments, although there are clusters of earthquakes at some of the ridge-transform boundaries. Some earthquakes do occur on spreading ridges, but they tend to be small and infrequent because of the relatively high rock temperatures in the areas where spreading is taking place.

Figure 11.2.2 Distribution of earthquakes of M4 and greater in the area of the mid-Atlantic ridge near the equator from 1990 to 1996. All are at a depth of 0 to 33 kilometres.

Earthquakes at Convergent Boundaries

The distribution and depths of earthquakes in the Caribbean and Central America area are shown in Figure 11.2.3. In this region, the Cocos Plate is subducting beneath the North America and Caribbean Plates (ocean-continent convergence), and the South and North America Plates are subducting beneath the Caribbean Plate (ocean-ocean convergence). In both cases, the earthquakes get deeper with distance from the trench. In Figure 11.2.3, the South America Plate is shown as being subducted beneath the Caribbean Plate in the area north of Colombia, but since there is almost no earthquake activity along this zone, it is questionable whether subduction is actually taking place.

Figure 11.2.3 Distribution of earthquakes of M4 and greater in the Central America region from 1990 to 1996 (red: 0 to 33 kilometres, orange: 33 to 70 kilometres, green: 70 to 300 kilometres, blue: 300 to 700 kilometres) (Spreading ridges are heavy lines, subduction zones are toothed lines, and transform faults are light lines.)

There are also various divergent and transform boundaries in the area shown in Figure 11.2.3, and as we’ve seen in the mid-Atlantic area, most of these earthquakes occur along the transform faults.

Figure 11.2.4 Distribution of earthquakes in the area of the Kuril Islands, Russia (just north of Japan) (White dots represent the April 2009 M6.9 earthquake. Red and yellow dots are from background seismicity over several years prior to 2009.)

The distribution of earthquakes with depth in the Kuril Islands of Russia in the northwest Pacific is shown in Figure 11.2.4. 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 M6.9 earthquake in April 2009. The relatively large earthquake took place on the upper part of the plate boundary between 60 kilometres and 140 kilometres 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 a greater width of rupture zone on a gently dipping 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, and on other similar ones, is predominantly near the upper side of the subducting plate. The frequency of earthquakes is greatest near the surface and especially around the area where large subduction quakes happen, but it extends to at least a 400 kilometre depth. There is also significant seismic activity in the overriding North America Plate, again most commonly near the region of large quakes, but also extending for a few hundred kilometres away from the plate boundary.

Figure 11.2.5 Distribution of earthquakes in the area where the India Plate is converging with the Asia Plate (data from 1990 to 1996, red: 0 to 33 kilometres, orange: 33 to 70 kilometres, green: 70 to 300 kilometres). (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 millimetres per year).

The distribution of earthquakes in the area of the India-Eurasia plate boundary is shown in Figure 11.2.5. This is a continent-continent convergent boundary, and it is generally assumed that although the India Plate continues to move north toward the Asia Plate, there is no actual subduction taking place. There are transform faults on either side of the India Plate in this area.

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 11.2.5 are related to the thrust faults shown in Figure 11.2.6 (and to hundreds of other similar ones that cannot be shown at this scale). The southernmost thrust fault in Figure 11.2.6 is equivalent to the Main Boundary Fault in Figure 11.2.5.

Figure 11.2.6 Schematic diagram of the India-Asia convergent boundary, showing examples of the types of faults along which earthquakes are focused. The devastating Nepal earthquake of May 2015 took place along one of these thrust faults.

There is a very significant concentration of both shallow and deep (greater than 70 kilometres) earthquakes in the northwestern part of Figure 11.2.5. This is northern Afghanistan, and at depths of more than 70 kilometres, 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 part of the India Plate beneath the Asia Plate in this area.

Exercise 11.1 Earthquakes in British Columbia

Figure 11.2.7 [Image Description]

This map shows the incidence and magnitude of earthquakes in British Columbia over a one-month period in March and April 2015.

  1. What is the likely origin of the earthquakes between the Juan de Fuca (JDF) and Explorer Plates?
  2. The string of small earthquakes adjacent to Haida Gwaii (H.G.) coincides closely with the rupture surface of the 2012 M7.8 earthquake in that area. How might these earthquakes be related to that one?
  3. Most of the earthquakes around Vancouver Island (V.I.) are relatively shallow. What is their likely origin?
  4. Some of the earthquakes in B.C. are interpreted as being caused by natural gas extraction (including fracking). Which of the earthquakes here could fall into this category?

See Appendix 3 for Exercise 11.1 answers.

Image descriptions

Figure 11.2.7 image description: The incidence and magnitude of earthquakes in British Columbia over a one-month period in March and April 2015: There were a few dozen smaller earthquakes spread out around Vancouver Island and the sunshine coast with a magnitude of 2. Farther west along the Explorer Plate, which is between the North American plate, the Juan de Fuca Plate, and the Pacific Plate, there were quite a few earthquakes with a magnitude of 3 and at least one earthquake with a magnitude of 4. Between the North American Plate and the Pacific Plate off the south-west coast of Haida Gwaii, there was a large cluster of earthquakes with magnitudes of 2. Along the Alaskan panhandle, there was a collection of 2- and 3-magnitude earthquakes. In addition, there were two 3-magnitude earthquakes west of Fort St. John in northern British Columbia and one or two 2-magnitude earthquakes. In total, this map shows one hundred and forty-nine earthquakes with a magnitude less than 2, ninety-seven earthquakes with a magnitude of 2, thirty-nine earthquakes with a magnitude of 3, and two earthquakes with a magnitude of 4 [Return to Figure 11.2.7]

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11.3 Measuring Earthquakes

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 used 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, which is proportional to 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 therefore intensity can vary widely, depending on the proximity to the earthquake and the types of materials underneath and the local conditions.

Earthquake Magnitude

Before we look more closely at magnitude we need to review what we know about body waves, and look at surface waves. Body waves are of two types, P waves, or primary or compression waves (like the compression of the coils of a spring), and S waves, or secondary or shear waves (like the flick of a rope). An example of P and S seismic wave records is shown in Figure 11.3.1. The critical parameters for the measurement of 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 of the earthquake.

Figure 11.3.1 P-waves and S-waves from a small (M4) earthquake that took place near Vancouver Island in 1997. [Image Description]

When body waves (P or S) reach Earth’s surface, some of their energy is transformed into surface waves, of which there are two main types, as illustrated in Figure 11.3.2. 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 10% slower than 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 11.3.2 Depiction of seismic surface waves.

Other important terms for describing earthquakes are hypocentre (or focus) and epicentre. 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 vertically above the hypocentre (Figure 11.3.3).

Figure 11.3.3 Epicentre and hypocentre.

A number of methods for estimating magnitude are listed in Table 11.1. Local magnitude (ML) was widely used until late in the 20th century, but moment magnitude (MW) 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 measure distant large earthquakes.

Because of the increasing size of cities in earthquake-prone areas (e.g., China, Japan, California) 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 to determine magnitude 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 down systems so that damage and casualties can be limited.

Table 11.1 A summary of some of the different methods for estimating earthquake magnitude.Table 11.1 by Steven Earle.
[Skip Table]
Type M Range Dist. Range Comments
Local or Richter (ML) 2 to 6 0 to 400 kilometres The original magnitude relationship defined in 1935 by Richter and Gutenberg. It is based on the maximum amplitude of S-waves recorded on a Wood‑Anderson torsion seismograph. ML values can be calculated using data from modern instruments. L stands for local because it only applies to earthquakes relatively close to the seismic station.
Moment (MW) Greater than 3.5 All Based on the seismic moment of the earthquake, which is equal to the average amount of displacement on the fault times the fault area that slipped. It can also be estimated from seismic data if the seismometer is tuned to detect long-period body waves.
Surface wave (MS) 5 to 8 20 to 180° A magnitude for distant earthquakes based on the amplitude of surface waves measured at a period near 20 seconds.
P-wave 2 to 8 Local Based on the amplitude of P-waves. This technique is being increasingly used to provide very rapid magnitude estimates so that early warnings can be sent to utility and transportation operators to shut down equipment before the larger (but slower) S-waves and surface waves arrive.

Exercise 11.2 Moment magnitude estimates from earthquake parameters

Use this moment magnitude calculation tool to estimate the moment magnitude based on the approximate length, width, and displacement values provided in the following table:

Table 11.2 Calculate Moment Magnitude Based on Length, Width, and Displacement Values
[Skip Table]
Length (kilometres) Width (kilometres) Displacement (metres) Earthquake MW?
60 15 4 The 1946 Vancouver Island earthquake
0.4 0.2 .5 The small Vancouver Island earthquake shown in Figure 11.3.1
20 8 4 The 2001 Nisqually earthquake described in Exercise 11.3
1,100 120 10 The 2004 Indian Ocean earthquake
30 11 4 The 2010 Haiti earthquake

The largest recorded earthquake had a magnitude of 9.5. Could there be a 10? You can answer that question using this tool. See what numbers are needed to make MW = 10. Are they reasonable?

See Appendix 3 for Exercise 11.2 answers.

The magnitude scale is logarithmic; in fact, the amount of energy released by an earthquake of M4 is 32 times higher than that released by one of M3, and this ratio applies to all intervals in the scale. If we assign an arbitrary energy level of 1 unit to a M1 earthquake the energy for quakes up to M8 will be as shown on the following table:

Table 11.3 The energy of an earthquake increases by 32 times at each magnitude level.
Magnitude Energy
1 1
2 32
3 1,024
4 32,768
5 1,048,576
6 33,554,432
7 1,073,741,824
8 34,359,738,368

In any given year, when there is a large earthquake on Earth (M8 or M9), the amount of energy released by that one event will likely exceed the energy released by all smaller earthquake events combined.

Earthquake Intensity

The intensity of earthquake shaking at any location is determined not only 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 buildings (and their inherent natural vibrations) 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 (Table 11.4). 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 than those underlain by unconsolidated river or lake sediments.

Table 11.4 The modified Mercalli intensity scale.
[Skip Table]
Level of intensity Description
Not felt (1) Not felt except by a very few under especially favourable conditions
Weak (2) Felt only by a few persons at rest, especially on upper floors of buildings
Weak (3) Felt quite noticeably by persons indoors, especially on upper floors of buildings; many people do not recognize it as an earthquake; standing motor cars may rock slightly; vibrations similar to the passing of a truck; duration estimated
Light (4) Felt indoors by many, outdoors by few during the day; at night, some awakened; dishes, windows, doors disturbed; walls make cracking sound; sensation like heavy truck striking building; standing motor cars rocked noticeably
Moderate (5) Felt by nearly everyone; many awakened; some dishes, windows broken; unstable objects overturned; pendulum clocks may stop
Strong (6) Felt by all, many frightened; some heavy furniture moved; a few instances of fallen plaster; damage slight
Very Strong (7) Damage negligible in buildings of good design and construction; slight to moderate in well-built ordinary structures; considerable damage in poorly built or badly designed structures; some chimneys broken
Severe (8) Damage slight in specially designed structures; considerable damage in ordinary substantial buildings with partial collapse; damage great in poorly built structures; fall of chimneys, factory stacks, columns, monuments, walls; heavy furniture overturned
Violent (9) Damage considerable in specially designed structures; well-designed frame structures thrown out of plumb; damage great in substantial buildings, with partial collapse; buildings shifted off foundations
Extreme (10) Some well-built wooden structures destroyed; most masonry and frame structures destroyed with foundations; rails bent
Extreme (11) Few, if any (masonry), structures remain standing; bridges destroyed; broad fissures in ground; underground pipelines completely out of service; earth slumps and land slips in soft ground; rails bent greatly
Extreme (12) Damage total; waves seen on ground surfaces; lines of sight and level distorted; objects thrown upward into the air

An example of this effect is the 1985 M8 earthquake that struck the Michoacán region of western Mexico, southwest of Mexico City. There was relatively little damage in the area around the epicentre, but there was tremendous damage and about 5,000 deaths in heavily populated Mexico City some 350 kilometres from the epicentre. 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 two 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 amplitude of the seismic waves was amplified by the lake sediments. Survivors of the disaster recounted that the ground in some areas moved up and down by about 20 centimetres every two seconds for over two minutes. Damage was greatest to buildings between 5 and 15 storeys tall, because they also resonated at around two seconds, which amplified the shaking.

Exercise 11.3 Estimating intensity from personal observations

The following observations were made by residents of the Nanaimo area during the M6.8 Nisqually earthquake near Olympia, Washington in 2001. Estimate the Mercalli intensities using Table 11.4.

Table 11.5
[Skip Table]
Building Type Floor Shaking Felt How long it lasted (in seconds) Description of Motion Intensity?
House 1 no 10 Heard a large rumble lasting not even 10 seconds, mirror swayed
House 2 moderate 60 Candles, pictures and CDs on bookshelf moved, towels fell off racks
House 1 no Pots hanging over stove moved and crashed together
House 1 weak Rolling feeling with a sudden stop, picture fell off mantle, chair moved
Apartment 1 weak 10 Sounded like a big truck then everything shook for a short period
House 1 moderate 20-30 Teacups rattled but didn’t fall off
Institution 2 moderate 15 Creaking sounds, swaying movement of shelving
House 1 moderate 15-30 Bed banging against the wall with me in it, dog barking aggressively

See Appendix 3 for Exercise 11.3 answers.

The graduated intensity of a 7.3 earthquake. Image description available.
Figure 11.3.4 Intensity map for the 1946 M7.3 Vancouver Island earthquake. [Long Decsription]

An intensity map for the 1946 M7.3 Vancouver Island earthquake is shown in Figure 11.3.4. 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.

Image Descriptions

Figure 11.3.1 image description: P-waves and S-waves from a small (M4) earthquake near Vancouver Island in 1997. The P-wave arrived in 0.7 seconds with an amplitude ranging from negative 0.7 millimetres per second to 1.1 millimetres per second and lasting until the arrival of the S-wave. The S-wave arrived at 8.7 seconds, with a minimum amplitude of negative 2.8 millimetres per second and a maximum amplitude of 2.7 millimetres per second. The S-wave’s net amplitude gradually decreased over the next 5 seconds. [Return to Figure 11.3.1]

Figure 11.3.4 image description: The graduated intensity of the 1945 M7.3 Vancouver Island earthquake based on the modified Mercalli intensity scale. The area surrounding the epicentre of the earthquake which included central Vancouver Island ranged between a very strong (7) and severe (8) intensity. The next ring included the northern and southern parts of Vancouver Island, as well as a part of the main land coast including Vancouver and much of the Sunshine coast a strong (6) intensity. The next ring, which reached experienced a moderate (5) intensity, included Seattle and much of the BC interior. The outermost ring ranged between not felt (1) and light (4) intensity. It was felt as far north as Prince Rupert and the southern tip of Haida Gwaii, south eastern BC, and as far south as north western Oregon. [Return to Figure 11.3.4]

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11.4 The Impacts of Earthquakes

Some of the common impacts of earthquakes include structural damage to buildings, fires, damage to bridges, highways, pipelines and electrical transmission lines, initiation of slope failures, liquefaction, and tsunami. The types of impacts depend to a large degree on where the earthquake is located: whether it is predominantly urban or rural, densely or sparsely populated, highly developed or underdeveloped, 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 solid rock, while the lower frequency waves (with periods slower than one second) pass through the solid rock without being absorbed, but are eventually absorbed and amplified by soft sediments. It is therefore 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 San Francisco, where parts of a two-layer highway built on soft sediments collapsed during the 1989 Loma Prieta earthquake (Figure 11.4.1).

Figure 11.4.1 A part of the Cypress Freeway in Oakland California that collapsed during the 1989 Loma Prieta earthquake.

Building damage is also greatest in areas of soft sediments, and multi-storey 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 poor, 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 1999 M7.6 Izmit earthquake (Figure 11.4.2). After two devastating earthquakes that year, Turkish authorities strengthened the building code further, but the new code has been applied only 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 11.4.2 Buildings damaged by the 1999 earthquake in the Izmit area, Turkey.

Fires are commonly associated with earthquakes because fuel pipelines rupture and electrical lines are damaged when the ground shakes (Figure 11.4.3). Most of the damage in the great 1906 San Francisco earthquake was caused by massive fires in the downtown area of the city (Figure 11.4.4). Some 25,000 buildings were destroyed by those fires, which were fuelled by broken gas pipes. Fighting the fires was difficult because water mains had 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 11.4.3 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.
Figure 11.4.4 Fires in San Francisco following the 1906 earthquake.
A debris slide that wiped out a large section of a residential area.
Figure 11.4.5 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 weak. 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 11.4.5).

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 is 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 11.4.6). 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.

Figure 11.4.6 Collapsed apartment buildings in the Niigata area of Japan. The material beneath the buildings was liquefied to varying degrees by the 1964 earthquake.

Parts of the Fraser River delta are prone to liquefaction-related damage because the region is characterized by a 2 metre to 3 metre thick layer of fluvial silt and clay over top of at least 10 metres of water-saturated fluvial sand (Figure 11.4.7). Under these conditions, it can be expected that seismic shaking will be amplified and that the sandy sediments will liquefy. This could lead to subsidence and tilting of buildings, and to failure and sliding of the silt and clay layer. Current building-code regulations in the Fraser delta area require that measures be taken to strengthen the ground underneath multi-storey buildings prior to construction.

Figure 11.4.7 Recent unconsolidated sedimentary layers in the Fraser River delta area (top) and the potential consequences in the event of a damaging earthquake.

Exercise 11.4 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 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.

Figure 11.4.8

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.

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; 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.

See Appendix 3 for Exercise 11.4 answers.

Earthquakes that take place beneath the ocean have the potential to generate tsunami. (Tsunami is the Japanese word for harbour wave. It is the same in both singular and plural.) The most likely situation for a significant tsunami is a large (M7 or greater) subduction-related earthquake. As shown in Figure 11.4.9, during the time between earthquakes the overriding plate becomes distorted by elastic deformation; it is squeezed laterally (Figure 11.4.9B) and pushed up.

Figure 11.4.9 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.

When an earthquake happens (Figure 11.4.9C), the plate rebounds and there is both uplift and subsidence on the sea floor, in some cases by as much as several metres vertically over an area of thousands of square kilometres. This vertical motion is transmitted through the water column where it generates a series of waves that then spread across the ocean.

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 M8), don’t typically generate tsunami either, because the motion is mostly side to side, not vertical.

Tsunami waves travel at velocities of several hundred kilometres per hour and easily make it to the far side of an ocean in about the same time as a passenger jet. The simulated one shown in Figure 11.4.10 is similar to that created by the 1700 Cascadia earthquake off the coast of British Columbia, Washington, and Oregon, which was recorded in Japan nine hours later.

Figure 11.4.10 Model of the tsunami from the 1700 Cascadia earthquake (around M9) showing open-ocean wave heights (colours) and travel time contours. Tsunami wave amplitudes typically increase in shallow water.

Tsunami are discussed further in Chapter 17 under the topic of waves and coasts.

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