Geologists study Earth — its interior and its exterior surface, the rocks and other materials around us, and the processes that formed those materials. They study the changes that have occurred over the vast time-span of Earth’s history, and changes that might take place in the near future.

Geology is a science, meaning that geological questions are investigated with deductive reasoning and scientific methodology. Geology is arguably the most interdisciplinary of all of the sciences because geologists must understand and apply other sciences, including physics, chemistry, biology, mathematics, astronomy, and more.

An aspect of geology that is unlike most of the other sciences is the role played by time — deep time — billions of years of it. When geologists study the evidence around them, they are often observing the results of are observing the results of events that took place thousands, millions, and even billions of years in the past, and which may still be ongoing. Many geological processes happen at incredibly slow rates — millimetres per year to centimetres per year — but because of the amount of time available, tiny changes can result in expansive oceans forming, or entire mountain ranges being worn away.

Geology on a Grand Scale in the Canadian Rocky Mountains

The peak on the right of the photographs in Figure 11.2 is Rearguard Mountain, which is a few kilometres northeast of Mount Robson. Mount Robson is the tallest peak in the Canadian Rockies, at 3,954 m. 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.

Many geological features are shown here. The rocks that these mountains are made of formed in ocean water over 500 million years ago. A few hundred million years later, the rocks were pushed east for tens to hundreds of kilometres, and thousands of meters upward in a great collision between Earth’s tectonic plates.

Over the past two million years this area, like most of the rest of Canada, has been repeatedly covered by glaciers that scoured away rocks to form the valley to the left of Rearguard Mountain. The Robson Glacier itself is now only a fraction of its  size during the Little Ice Age of the 15th to 18th centuries. And, like almost all other glaciers on Earth, it is now receding even more rapidly because of climate change. Figure 11.2 (right) taken around 1908 by the Canadian geologist and artist Arthur Philemon Coleman, gives an indication of how much the glacier has receded in the last hundred years.

Figure 11.2 Rearguard Mountain and Robson Glacier in Mount Robson Provincial Park, BC. Left: Robson Glacier today, retreating up the valley. Right: Robson Glacier circa 1908. Sources: Left- Karla Panchuk (2017) CC BY-SA 4.0 with photo by Steven Earle (2015) CC BY 4.0 view source. Right: A.P. Coleman (c. 1908) Public Domain. Click the image for more attributions.

Geology is about understanding the evolution of Earth through time. It is about discovering resources such as metals and energy, and minimizing the environmental implications of our use of resources. It is about learning to mitigate the hazards of earthquakes, volcanic eruptions, and slope failures. All of these aspects of geology, and many more, are covered in this textbook.

References

Victoria University Library (2009) A. P. Coleman Exhibition. Retrieved 25 August 2017. Visit the website

Why?  Because 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.

• We can study rocks and the fossils they contain to understand the evolution of our environment and the life within it.
• We can learn to minimize our risks from earthquakes, volcanoes, slope failures, and damaging storms.
• We can learn how and why Earth’s climate changed in the past, and use that knowledge to understand both natural and human-caused climate change.
• We rely on Earth for resources such as soil, water, metals, industrial minerals, and energy, and we need to know how to find these resources and exploit them sustainably.
• We can recognize how human activities have altered the environment, and learn how to prevent and sometimes repair the damage.
• We can use knowledge of Earth to understand other planets in our solar system, and those around distant stars.

The Importance of Geological Studies for Minimizing Risks to the Public

Figure 1.3 shows 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. It destroyed another house below, and killed one person. The slope failure happened after a heavy rainfall, which is a common occurrence in southwestern B.C. in the winter.

Figure 1.3 Aftermath of a deadly debris flow in the Riverside Drive area of North Vancouver in January, 2005. Source: The Province (2005), used with permission.

A geological report written in 1980 warned the District of North Vancouver that the area was prone to slope failure, and that steps should be taken to minimize the risk to residents. Unfortunately, not enough was done in the intervening 25 years to prevent a tragedy.

Geologists do a lot of different things.  Many of the jobs are the things you would expect.  Geologists work in the resource industry, including mineral exploration and mining, and exploring for and extracting sources of energy. They do hazard assessment and mitigation (e.g., assessment of risks from slope failures, earthquakes, and volcanic eruptions).  They study the nature of the subsurface for construction projects such as highways, tunnels, and bridges. They use information about the subsurface for water supply planning, development, and management; and to decide how best to contain contaminants from waste.

Geologists also do the research that makes practical applications of geology possible.  Some geologists spend their summers trekking through the wilderness to make maps of the rocks in a particular location, and collect clues about the geological processes that occurred there.  Some geologists work in laboratories analyzing the chemical and physical properties of rocks to understand how the rocks will behave when forces act on them, or when water flows through them.  Some geologists specialize in inventing ways to use complex instruments to make these measurements.  Geologists study fossils to understand ancient animals and environments, and go to extreme environments to understand how life might have originated on Earth.  Some geologists help NASA understand the data they receive from objects in space.

Geological work can be done indoors in offices and labs, but some people are attracted to geology because they like to be outdoors.  Many geological opportunities involve fieldwork in places that are as amazing to see as they are interesting to study. Sometimes these are locations where few people have ever set foot, and where few ever will again.

Figure 1.4 Geologists at work on the island of Spitsbergen, part of the Svalbard archipelago. The islands are located in the Arctic Ocean north of Norway. Source: Gus MacLeod (2007) CC BY-NC-ND 2.0 view source

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 key features of serious inquiry are the following:

• Creation of a hypothesis. This is a tentative idea about how to explain a set of observations.
• Test the hypothesis.  A hypothesis can be used to make predictions, and experiments can be run to see if those predictions are correct.

An Example of the Scientific Method at Work

Consider a field trip to the stream shown in Figure 1.5. Notice that the rocks in and along the stream are rounded off rather than having sharp edges. We might hypothesize that the rocks were rounded because as the stream carried them, they crashed into each other and pieces broke off.

Figure 1.5 Hypothesizing about the origin of round rocks in a stream. Source: Steven Earle (2015) CC BY 4.0 view source

If the hypothesis is correct, then the further we go downstream, the rounder and smaller the rocks should be. Going upstream we should find that the rocks are more angular and larger. If we were patient we could also test the hypothesis by marking specific rocks and then checking back to see if those rocks have become smaller and more rounded as they moved downstream.

If the predictions turn out to be correct, we must still be careful about how much certainty to attach to our hypothesis.  Although our hypothesis might seem to us to be the only reasonable explanation, someone could argue that we have the mechanism wrong, and the rocks weren’t rounded by bumping into each other. If our experiment didn’t specifically check for the mechanism (e.g., by looking to see if chips fall off the rocks and the rocks are made smoother) then we would have to acknowledge the possibility.  We needn’t abandon the hypothesis as a useful tool for making predictions, but it is necessary to be open to the possibility that other things might be going on. If someone demonstrates conclusively that our hypothesis is wrong, then we have to discard the hypothesis and come up with a better one.

A good hypothesis is testable.  Someone might argue that an extraterrestrial organization creates rounded rocks and places them in streams when nobody is looking. There is no practical way to test this hypothesis to confirm it, and there is no way to prove it false. Even if we never see aliens at work, we still can’t say they haven’t been, because according to the hypothesis they only work when people aren’t looking. Compare this to our original hypothesis which allows us to make testable predictions such as rocks getting smaller and rounder downstream. Our original hypothesis gives us a way to see how realistic it is, whereas the alien hypothesis gives us no way to know if it makes sense or not.

Theories and Laws

Two other terms appear in discussions of the scientific method: theory and law. A theory starts out as a hypothesis, but over a long period of time and a great many tests, it has never come up short. That doesn’t mean it never will, but the odds of that are very unlikely given our present (and conceivable future) state of knowledge.  You may have heard someone dismiss an idea by saying it is “just a theory,” but they are using the term incorrectly if they mean to say it’s a wild and unproven guess.

A law is a description of a phenomenon rather than an explanation of it.  For example, you could do thousands of tests by dropping an object with known mass and measuring its acceleration and the force with which it hits the ground.  Again and again your results will yield the formula force = mass x acceleration.  However, that doesn’t mean you know what is responsible for the force accelerating it toward the ground.  Yes, we say that gravity is pulling it toward the Earth’s surface, but why? A law is true regardless of why a phenomenon happens as long as it describes the outcome of that phenomenon.

In geology there are three big ideas that are fundamental to the way we think about how Earth works.  The ideas are like the sound track to a movie- sometimes we might not even notice them, but at the same time they affect our perception of what is happening.  In the rest of this book these ideas may be mentioned explicitly in some cases, but in other cases it will be helpful for you to realize that they are relevant, even if they are not being discussed by name.

Geological Time (Deep Time)

Earth is approximately 4.57 billion years old (4,570,000,000 years), which is a long time for geological events to unfold and changes to happen. The changes themselves might be tiny. For example, over a year, a chemical reaction might eat away a few layers of atoms at the surface of a rock. But over time the changes accumulate and have a great impact. Over hundreds of millions of years the chemical reaction could cause a mountain range to crumble into grains of sand, and be swept away by rivers.

For geologists who study very, very slow processes, 10 million years might be a short time, and 1 million years might be trivial.  For these geologists, intervals of 1 million years aren’t even useful to consider, because the changes over that time are too small to see in the rocks that accumulated.

As you read through this book, keep in mind that the well of geologic time is indeed deep, and “ancient” is defined in a whole new way.

Expressing Geological Time in Numbers

Special notation is used for geological time because, as you might imagine, writing all those zeroes can become tiresome.  Table 1.1 shows common abbreviations you will see throughout this book.

Expressing Geological Time Using the Geological Time Scale

The geological time scale (Figure 1.6) is a way of breaking down geological time according to important events in Earth’s history.  Time is divided into eons, eras, periods, and epochs, and these intervals are referred to by names rather than by years.  Giving intervals of geologic time names rather than using numbers makes sense because we won’t always know the age in years (the absolute age) of a rock or fossil, but we can place it in context based on our knowledge of the geological record.  We can describe its relative age by saying that it is older than or younger than another rock or fossil.

Figure 1.6 Geologic Society of America Geologic Time Scale, 2012. Source: Walker, J.D., Geissman, J.W., Bowring, S.A., and Babcock, L.E., compilers (2012) Geologic Time Scale v. 4.0: Geological Society of America, doi: 10.1130/2012.CTS004R3C. Download PDF

The tricky thing about the geologic time scale is that the boundaries are always changing.  As our knowledge of the absolute age of an event improves with new discoveries, it might be necessary to nudge a boundary earlier or later.  Sometimes the original reason for defining a boundary no longer holds, but we agree to use it anyway.  For example, 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, and its start was meant to mark the first appearance of these organisms. In fact, we now have evidence that large organisms — those that leave fossils visible to the naked eye — have existed longer than that, first appearing by 600 Ma at the latest.

An Early Definition of the Proterozoic

Notice that in Figure 1.6 the Proterozoic Eon precedes the Phanerozoic Eon. This was not always the case. Figure 1.7 shows an excerpt from a periodical published in 1879, in which the Proterozoic is defined as covering the Cambrian through Silurian. The author refers to “the most extreme adherents of the Murchisonian party in geology,” a reference to the contentious assertion by Scottish geologist Roderick Murchison (1792-1871) that the Silurian Period should encompass the Cambrian and Ordovician periods as well.

Figure 1.7 An excerpt from the periodical The Annals and Magazine of Natural History (1879) in which the name “Proterozoic” is assigned to the Cambrian, Ordovician, and Silurian periods instead of to the time preceding the Cambrian. Source: Karla Panchuk (2017) CC BY 4.0 Read the book

A Way To Think About Geological Time

A useful mechanism for understanding geological time is to scale it 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 in Table 1.2.

Uniformitarianism

Uniformitarianism is the notion that the geological processes occurring on Earth today are the same ones that occurred in the past.  This is an important idea because it means that observations we make today about geological processes can be used to interpret and understand the rock record.  While this idea might not seem remarkable today, it was ground breaking and even controversial for its time.  Many people who heard about it for the first time thought about the age of the Earth in thousands of years, but uniformitarianism required them to think on timescales almost too vast to comprehend.  For some, this implied questioning their most deeply held religious beliefs.

The Scottish geologist James Hutton initially presented the idea in 1785Read James Hutton's abstract at http://bit.ly/1j6tIAN. Note that the typeface prints an "s" like an "f.".  Charles Lyell, also a Scottish geologist, paraphrased this idea as “the present is the key to the past” in his book Principles of Geology.The 7th edition of Charles Lyell's Principles of Geology (1847) can be found at http://bit.ly/1l3T6Zh  This is how it is often described today.

To be clear, “the present is the key to the past” can be viewed as an oversimplification. Not all geological processes occurring today occurred at all times in the geological past.  For example, some important chemical reactions that happened on Earth’s surface today require abundant oxygen in the atmosphere, and could not have occurred prior to Earth developing an oxygen-rich atmosphere.  Conversely, there was a time in Earth’s history when continents as we know them hadn’t yet developed. Some events, such as devastating impacts by objects from space, have never been witnessed on the same scale by humans. We must be cognizant of the fact that conditions were different at different times in Earth’s history, and take that into account when interpreting the rock record.

Despite the different past conditions on Earth as a whole, there still exist environments today where some of these conditions are present. These environments are like little samples of what Earth used to be like.  This means we can still use present conditions to inform us about the past, but we have to think carefully about ways that such environments today differ from the ancient environments that no longer exist.

Plate Tectonics

It is only within the last 50 years or so that we have been able to answer questions like, “How did that mountain range get there?” and “Why do earthquakes happen where they do?”  The theory of plate tectonics– the idea that Earth’s surface is broken into large moving fragments, called plates– profoundly changed our perspective on how the Earth works.  Figure 1.8 shows Earth’s 15 largest tectonic plates, along with arrows indicating the plates’ direction of motion, and how fast they go.  (Longer arrows mean faster motion.)  There are many more plates on Earth that are too small to show conveniently in Figure 1.8. A more detailed map of Earth’s tectonic plates can be found at here.

Figure 1.8 Earth’s fifteen largest tectonic plates. Black arrows show the direction of plate motions. The length of the arrow indicates velocity. Red arrows show how plates move relative to each other. Source: Steven Earle (2015) CC BY 4.0. view source Modified after U. S. Geological Survey (1996) Public Domain view original

Prior to plate tectonics, we made observations but could only guess at mechanisms.  It was like watching the hands on a clock and trying to guess what moves them.  After plate tectonics it was like being able to open the clock and not only watch the gears turn, but realize for the first time that there are such things as gears. Plate tectonics not only explains why things have happened, but also allows us to predict what might happen in the future.

Plate tectonics is covered in more detail later, however the key point is that Earth’s outer layer consists of rigid plates that are constantly interacting with each other as they move around the Earth.  The boundaries of plates move away from each other in some places, collide in others, and sometimes just slide past each other (illustrated by the red arrows in Figure 1.8). The plates can move because they are floating on a layer of weak rock that deforms as the plates travel, much the same way the filling in a peanut butter and jelly sandwich allows you to slide the top layer of bread across the bottom layer.

Whether the plates move away from each other, collide, or just slide past each other determines things like the locations of mountain belts and volcanoes, where earthquakes happen, and the shapes and sizes of oceans and continents.

References

Cottrell, M. (2006) History of Saskatchewan. Retrieved 26 August 2017. Visit the website

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

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. Geologists must take into account the fact that the geological features we see today may have formed thousands, millions, or even billions of years ago, and over very long time spans.

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 industry, and in efforts to protect the environment. Geologists work to minimize the risks from geological hazards (e.g., earthquakes), and to help the public understand those risks. Geologists investigate Earth materials in the field, in and in the lab.

1.4 We Study Earth Using the Scientific Method

Scientific inquiry requires a careful process of making a hypothesis and then testing it. If a hypothesis doesn’t pass the test, it’s time for a new one. A theory is a hypothesis that has been tested repeatedly and never failed a test. A law is a description of a natural process.

1.5 Three Big Ideas: Geological Time, Uniformitarianism, and Plate Tectonics

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.

Uniformitarianism: Processes that occur today also occurred in the geologic past.  We can use our observations of the present to understand the processes that shaped the Earth throughout its history.

Plate tectonics: Earth’s surface is broken into plates that move and interact with each other.  The interactions between these plates are key for understanding the mechanisms behind geologic processes.

Review Questions

1. How does the element of time make geology different from the other sciences, such as chemistry and physics?
2. List three ways in which geologists can contribute to society.
3. The following dates are written with the abbreviations Ga, Ma, and ka. Express the dates in years. (For example, 2.3 Ma = 2,300,000 years)
   1. 2.75 ka
   2. 0.93 Ga
   3. 4.2 Ma
   4. 0.2 ka.
4. Dinosaurs first appear in the geological record in rocks from about 215 Ma and then most became extinct at 65 Ma. What percentage of geological time does this represent?
5. If sediments typically accumulate at a rate of 1 mm/year, what thickness of sediment could accumulate over a period of 30 million years?
6. Does uniformitarianism mean that conditions on Earth are uniform, and never change?
7. Summarize the main idea behind plate tectonics.

1. Geology requires that we consider vast amounts of time, and think about the effects that accumulate over thousands, millions, or even billions of years.
2. There are many ways that geologists contribute. Geologists provide information to reduce the risk of harm from hazards such as earthquakes, volcanoes, and slope failures; they play a critical role in the discovery of important resources; they contribute to our understanding of life and its evolution through paleontological studies; and they play a leading role in the investigation of climate change, past and present and its implications.
3. Ages in years
   1. 2.75 ka = 2,750 years
   2. 0.93 Ga = 930,000,000 years
   3. 14.2 Ma = 14,200,000 years
   4. 0.2 ka = 200 years.
4. 215 – 65 = 150 Ma. Since the age of the Earth is 4570 Ma, this represents 150/4,570 = 0.033 or 3.3% of geological time.
5. At 1 mm/y 30,000,000 mm of sediment would accumulate over that 30 million years. This is equivalent to 30,000 m or 30 km. Few sequences of sedimentary rock are even close to that thickness because most sediments accumulate at much lower rates, more like 0.1 mm/y. Also, over time the sediments are compressed.
6. No. Uniformitarianism means that we can use the processes we observe today to help us understand what happened in the past.
7. Plate tectonics is the idea that Earth’s outer layer is broken into rigid plates. The plates move around and interact with each other along their margins.

According to the big bang theory, the universe blinked violently into existence 13.8 billion years ago. The big bang is often described as an explosion, but imagining it as an enormous fireball isn’t accurate. The big bang started with a sudden expansion of energy and space from a single point. The kind of Hollywood explosion that might come to mind involves expansion of matter and energy within space, but during the big bang, energy, space, and matter were created. In Figure 2.2 the pointed base of the universe “vessel” represents the big bang. Time advances moving up the diagram. The vessel gets wider as time progresses, reflecting the expansion of the universe.

Figure 2.2 The big bang. The universe began 13.8 billion years ago as a rapid expansion of space, energy, and matter. It continues to expand. Left: Timeline of the universe. The point at the base of the “vessel” represents the moment of the big bang. The vessel gets wider as time progresses, representing the expansion of the universe. Right: Mollwiede projection of the cosmic microwave background, a “fog” from when the universe was still very dense. Temperature variations correspond to clumping of matter in the early universe. Source: Karla Panchuk (2018) CC BY 4.0 modified after Ryan Kaldari (2006) Public Domain view source, derivative of NASA/WMAP Science Team (2006) Public Domain view source. CMB map by NASA/WMAP Science Team (2006) Public Domain view source. Click the image for data sources.

You might wonder how a universe can be created out of nothing. Creating a universe out of nothing is mostly beyond the scope of this chapter, but there is a way to think about it. The particles that make up the universe have opposites that cancel each other out, similar to the way that we can add the numbers 1 and -1 to get zero (also known as “nothing”). As far as the math goes, having zero is exactly the same as having a 1 and a -1. It is also exactly the same as having a 2 and a -2, a 3 and a -3, two -1s and a 2, and so on. In other words, nothing is really the potential for something if you divide it into its opposite parts.

Composition of the Universe

In Figure 2.2, the “contents” of the vessel change as time progresses. A few minutes after the big bang, the universe was still too hot and dense to be anything but a sizzle of particles smaller than atoms. But as it expanded, it also cooled. Eventually particles that collided were able to stick together to form atoms, rather than being smashed apart again when other particles crashed into them. Those collisions produced hydrogen and helium, the most common elements in the universe.

For a long time after the big bang, clouds of hydrogen and helium atoms drifted about a dark universe. The “dark ages” (bottom of Figure 2.2) were a time when the ingredients for stars existed, but stars themselves did not. It took approximately 500 million years for enough hydrogen atoms to clump together in clouds to allow the first stars to form and begin to shine.

Looking Back to the Early Stages of the Big Bang

The notion of seeing the past is often used metaphorically when we talk about ancient events, but in this case it is meant literally. In our everyday experience, when we watch an event take place, we perceive that we are watching it as it unfolds in real time. In fact, this isn’t true. To see the event, light from that event must travel to our eyes. Light travels very rapidly, but it does not travel instantly. If we were watching a digital clock 1 m away from us change from 11:59 a.m. to 12:00 p.m., we would actually see it turn to 12:00 p.m. three billionths of a second after it happened.

This isn’t enough of a delay to cause us to be late for an appointment, but the universe is a very big place, and the “digital clock” in question is often much, much farther away. In fact, the universe is so big that it is convenient to describe distances in terms of light years, or the distance light travels in one year. What this means is that light from distant objects takes so long to get to us that we see those objects as they were at some considerable time in the past. For example, the star Proxima Centauri is 4.24 light years from the sun. If you viewed Proxima Centauri from Earth on January 1, 2018, you would actually see it as it appeared in early October 2013.

We now have tools that are powerful enough to look deep into space and see the arrival of light from early in the universe’s history. Astronomers can detect light from approximately 380,000 years after the big bang is thought to have occurred. Physicists tell us that if the big bang happened, then particles within the universe would still be very close together at this time. They would be so close that light wouldn’t be able to travel far without bumping into another particle and getting scattered in another direction. The effect would be to fill the sky with glowing fog, the “afterglow” from the formation of the universe.

In fact, this is exactly what we see when we look at light from 380,000 years after the big bang. The fog is referred to as the cosmic microwave background (or CMB), and it has been carefully mapped throughout the sky. In Figure 2.2, the colourful patch at the base of the diagram represents the fog that is measured today as the CMB. Figure 2.2 (right) is a CMB map of the universe in Mollweide projection. This is a projection that is used to show Earth’s geography on a flat surface. In this case, the map of the CMB represents a sphere surrounding Earth rather than what’s beneath our feet.

Colour variations in the CMB map represent temperature variations. These variations translate to differences in the density at which matter was distributed in the early universe. The red patches are the highest density regions and the blue patches are the lowest density. Higher density regions represent the eventual beginnings of stars and planets. The CMB map in Figure 2.2 has been likened to a baby picture of the universe.

The Universe is Still Expanding

The expansion that started with the big bang never stopped. It continues today, and we can see it happen by observing that large clusters of billions of stars, called galaxies, are moving away from us. (An exception is the Andromeda galaxy with which we are on a collision course.) The astronomer Edwin Hubble came to this conclusion when he observed that the light from other galaxies was red-shifted. The red shift is a consequence of the Doppler effect. This refers to how we see waves when the object that is creating the waves is moving toward us or away from us.

Before looking at the Doppler effect as it pertains to light, it can be useful to see how it works on something more tangible. The duckling swimming in Figure 2.3 is generating waves as it moves through the water. It is generating waves that move forward as well as back, but notice that the ripples ahead of the duckling are closer to each other than the ripples behind the duckling. The distance from one ripple to the next is called the wavelength. The wavelength is shorter in the direction that the duckling is moving, and longer as the duckling moves away.

Figure 2.3 A duckling illustrates the Doppler effect in water. The ripples made in the direction the duckling is moving (blue lines) are closer together than the ripples behind the duckling (red lines). Source: Karla Panchuk (2015) CC BY 4.0. Photo by M. Harkin (2013) CC BY 2.0 view source

When waves are in air as sound waves rather than in water as ripples, the different wavelengths manifest as sounds with different pitches — the short wavelengths have a higher pitch, and the long wavelengths have a lower pitch. This is why an observer will hear a change in the pitch of a car’s engine as the car races past.

For light waves, wavelength translates to colour. In the spectrum of light that we can see, shorter wavelengths are on the blue end of the spectrum, and longer wavelengths are on the red end of the spectrum. In Figure 2.4, the longer or shorter wavelengths of the water ripples at the top of the diagram reflect the longer or shorter wavelengths of light in the visible spectra below. Does this mean that galaxies look red because they are moving away from us? No, but the colour we see is shifted toward the red end of the spectrum and longer wavelengths.

Figure 2.4 Red shift in light from the supercluster BAS11 compared to the sun’s light. Black lines represent wavelengths absorbed by atoms (mostly hydrogen and helium). For BAS11 the black lines are shifted toward the red end of the spectrum compared to the sun. Source: Karla Panchuk (2018) CC BY 4.0, spectra by Georg Wiora (2011) CC BY-SA 2.5 view source

Notice that the sun’s spectrum in the upper part of Figure 2.4 has black lines in it. The black lines are there because some colours are missing from the sun’s light that reaches Earth. Different elements absorb light of specific wavelengths, and many of the black lines in Figure 2.4 represent colours that are absorbed by hydrogen and helium within the sun. This means the black lines are like a bar code that can tell us what a star is made of.

The lower spectrum in Figure 2.4 is the light coming from BAS11, an enormous cluster of approximately 10,000 galaxies located 1 billion light years away. The black lines represent the same elements as in the sun’s spectrum, but they are shifted to the right toward the red end of the spectrum, because BAS11 is moving away from us as the universe continues to expand. To summarize, because almost all of the galaxies we can see have light that is red-shifted, it means they are all moving away from us. In fact, the farther away they are, the faster they are going. This is evidence that the universe is still expanding.

References

European Space Agency (2015). Planck reveals first stars were born late. Visit website

Knop, R. (2010). The History of the Universe. Visit blog

Lawrence, C. R. (2015, March). Planck 2015 Results. Paper presented to the Astrophysics Subcommittee, NASA HQ. View slides

Only four elements account for 95% of Earth’s mass: oxygen (O), magnesium (Mg), silicon (Si), and iron (Fe). Most of the remaining 5% comes from aluminum (Al), calcium (Ca), nickel (Ni), hydrogen (H), and sulphur (S). We know that the big bang made hydrogen, but where did the rest of the elements come from?

The answer is that the other elements were made by stars. Sometimes stars are said to “burn” their fuel, but burning is not what is going on within stars. The burning that happens when wood in a campfire is turned to ash and smoke is a chemical reaction — heat causes the atoms that were in the wood and in the surrounding atmosphere to exchange partners. Atoms group in different ways, but the atoms themselves do not change. What stars do is change the atoms. The heat and pressure within stars cause smaller atoms to smash together and fuse into new, larger atoms. For example, when hydrogen atoms smash together and fuse, helium is formed. Large amounts of energy are released when atoms fuse within stars, and this is what causes stars to shine. Stars can form large quantities of elements as heavy as iron during their normal burning process. Side reactions can form heavier elements in small amounts.

It takes larger stars to make elements as heavy as iron in large quantities. Our sun is an average star. After it uses up its hydrogen fuel to make helium, and some of that helium is fused to make small amounts of other elements, it will be at the end of its life. It will stop making new elements and will cool down and bloat until its middle reaches the orbit of Mars. In contrast, large stars end their lives in spectacular fashion. They explode as supernovae, casting off newly formed atoms into space, and triggering side reactions to make even more heavy atoms. It took many generations of stars creating heavier elements and casting them into space before heavier elements were abundant enough for planets like Earth to form.

Until recently, astronomers have only been able to see stars that already contain heavier elements in small amounts, but not the first-generation stars that started out before any of the heavier elements were produced. That changed in 2015 when it was announced that a distant galaxy called CR7 had been found that contained stars made only of hydrogen and helium. The galaxy is so far away that it shows us a view of the universe from approximately 800 million years after the big bang. Since then, more galaxies like CR7 have been discovered.

References

Pilipenko, S. V. (2013). Paper-and-pencil cosmological calculator. arXiv:1303.5961v1 [astro-ph.CO]

Royal Astronomical Society (2016, June 28). CR7 is not alone—A team of super bright galaxies in the early universe. Phys.org Visit website

Sobral, D., Matthee, J.,  Darvish, B., Schaerer, D., Mobasher, B., Röttgering, H., Santos, S., & Hemmati, S. (2015). Evidence for PopIII-like stellar populations in the most luminous Lyman-α emitters at the epoch of re-ionisation: spectroscopic confirmation. The Astrophysical Journal 808(2) doi: 10.1088/0004-637x/808/2/139.

A solar system consists of a collection of objects orbiting one or more central stars. All solar systems start out the same way. They begin in a cloud of gas and dust called a nebula. Nebulae are some of the most beautiful objects that have been photographed in space. They have vibrant colours from the gases and dust they contain, and brilliant twinkling from the many stars that have formed within them (Figure 2.5). The gas consists largely of hydrogen and helium, and the dust consists of tiny mineral grains, ice crystals, and organic particles.

Figure 2.5 The Pillars of Creation within the Eagle Nebula viewed in visible light (left) and near infrared light (right). Near infrared light captures heat from stars and allows us to view stars that would otherwise be hidden by dust. This is why the picture on the right appears to have more stars than the picture on the left. Source: NASA, ESA, and the Hubble Heritage Team (STScI/AURA) (2015) Public Domain. view source

Step 1: Collapse a Nebula

A solar system begins to form when a small patch within a nebula (small by the standards of the universe, that is) begins to collapse upon itself. Exactly how this starts isn’t clear, although it might be triggered by the violent behaviour of nearby stars as they progress through their life cycles. Energy and matter released by these stars might compress the gas and dust in nearby neighbourhoods within the nebula.

Once it is triggered, the collapse of gas and dust within that patch continues for two reasons. One of those reasons is that gravitational force pulls gas molecules and dust particles together. But early in the process, those particles are very small, so the gravitational force between them isn’t strong. So how do they come together? The answer is that dust first accumulates in loose clumps for the same reason dust bunnies form under the bed: static electricity. Given the role of dust bunnies in the early history of the solar system, one might speculate that an accumulation of dust bunnies poses a substantial risk to one’s home (Figure 2.6). In practice, however, this is rarely the case.

Figure 2.6 Public service announcement. If you don’t think housekeeping is important, then you don’t understand the gravity of the situation. Source: Karla Panchuk (2018) CC BY 4.0. Planets modified after NASA/JPL (2008) Public Domain. view source

Step 2: Make a Disk with a Star at Its Centre

As the small patch within a nebula condenses, a star begins to form from material drawn into the centre of the patch, and the remaining dust and gas settle into a protoplanetary disk that rotates around the star. The disk is where planets will eventually form. Figure 2.7 (upper left) is an artist’s impression of a protoplanetary disk, and Figure 2.7 (upper right) is an actual protoplanetary disk surrounding the star HL Tauri. Notice the dark rings within the HL Tauri protoplanetary disk. These are gaps formed by the collection of dust and debris by incipient planets, called protoplanets, as they orbit the star. There is an analogy for this in our own solar system, because the dark rings are akin to the gaps in the rings of Saturn (Figure 2.7, lower left), where moons can be found (Figure 2.7, lower right).

Figure 2.7 Protoplanetary disks and Saturn’s rings. Upper left: Artist’s impression of a protoplanetary disk containing gas and dust, surrounding a new star. Upper right: A photograph of the protoplanetary disk surrounding HL Tauri. The dark rings within the disk are thought to be gaps where newly forming planets are sweeping up dust and gas. Lower left: A photograph of Saturn showing similar gaps within its rings. The bright spot at the bottom is an aurora, similar to the northern lights on Earth. Lower right: a close-up view of a gap in Saturn’s rings showing a moon as a white dot. Source: Upper left- NASA/JPL-Caltech (2008) Public Domain view source; Upper right- ALMA (ESO/NAOJ/NRAO) (2014) CC BY 4.0 view source; Lower left- NASA, ESA, J. Clarke (Boston University), and Z. Levay (STScI) (2005) Public Domain view source; Lower right- NASA/JPL/Space Science Institute (2005) Public Domain view source

Step 3: Build Some Planets

In general, planets can be classified into three categories based on what they are made of (Figure 2.8). Terrestrial planets are those planets like Earth, Mercury, Venus, and Mars that have a core of metal surrounded by rock. Jovian planets (also called gas giants) are those planets like Jupiter and Saturn that consist predominantly of hydrogen and helium. Ice giants are planets such as Uranus and Neptune that consist largely of water ice, methane (CH₄) ice, and ammonia (NH₃) ice, and have rocky cores. Often, the ice giant planets Uranus and Neptune are grouped with Jupiter and Saturn as gas giants; however, Uranus and Neptune are very different from Jupiter and Saturn.

Figure 2.8 Three types of planets. Jovian (or gas giant) planets such as Jupiter consist mostly of hydrogen and helium. They are the largest of the three types. Ice giant planets such as Uranus are the next largest. They contain water, ammonia, and methane ice, and have rocky cores. Terrestrial planets such as Earth are the smallest, and they have metal cores covered by rocky mantles. Source: Karla Panchuk (2015) CC BY 4.0. Click the image for more attributions.

These three types of planets are not mixed together randomly within our solar system. Instead they occur in a systematic way, with terrestrial planets closest to the sun, followed by the Jovian planets and then the ice giants (Figure 2.9). Smaller solar system objects follow this arrangement as well. The asteroid belt contains bodies of rock and metal. Bodies ranging from metres to hundreds of metres in diameter are classified as asteroids, and smaller bodies are referred to as meteoroids. In contrast, the Kuiper belt (Kuiper rhymes with piper), and the Oort cloud (Oort rhymes with sort), which are at the outer edge of the solar system, contain bodies composed of large amounts of ice in addition to rocky fragments and dust.

Figure 2.9 Our solar system. Top: The solar system shown with distances to scale. Distances are in astronomical units (AU), where 1 AU is the average distance from Earth to the sun. The edge of the Kuiper belt extends to 50 AU (7.5 billion km), but this distance is minuscule compared to the size of the solar system as a whole, which extends to the edge of the Oort cloud, thought to be 15 trillion km away. Bottom: Solar system with the sun and planets to scale. The gas giants are the largest planets, followed by the ice giants, and then the terrestrial planets. Note that the planets in this diagram likely do not reflect the entire population of planets in our solar system because evidence suggests that large planets are present beyond the Kuiper belt. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0, Milky Way photo by R@pp (2017) CC BY-NC-SA 2.0 view source, planet photographs courtesy of NASA. Click the image for planet photo sources and attributions.

Part of the reason for this arrangement is the frost line (also referred to as the snow line). The frost line marks the division between the inner part of the protoplanetary disk closer to the sun, where it was too hot to permit anything but silicate minerals and metal to crystalize, and the outer part of the disk farther from the sun, where it was cool enough to allow ice to form. As a result, the objects that formed in the inner part of the protoplanetary disk consist largely of rock and metal, while the objects that formed in the outer part consist largely of gas and ice. The young sun blasted the solar system with raging solar winds (winds made up of energetic particles), which helped to drive lighter molecules toward the outer part of the protoplanetary disk.

Rules of the Accretion Game

The objects in our solar system formed by accretion. Early in this process, particles collected in fluffy clumps because of static electricity. As the clumps grew larger, gravity became more important and collected clumps into solid masses, and solid masses into larger and larger bodies. If you were one of these bodies in the early solar system, and participating in the “accretion game” with the goal of becoming a planet, you would have to follow some key rules:

• Keep your velocity just right. If you move too fast and collide with another body, you both smash up and have to start again. If you move slowly enough, gravity will keep you from bouncing off each other and you can grow larger.
• Your distance from the Sun will determine how big you can get. If you are closer, there is less material for you to collect than if you are farther away.
• To begin with, you can only collect mineral and rock particles. You have to grow above a certain mass before your gravity is strong enough to hang onto gas molecules, because gas molecules are very light.
• As your mass increases, your gravity becomes stronger and you can grab material from farther away. The bigger you are, the faster you grow.

You would also have to watch out for some dangers:

• In the early stages of the game, the protoplanetary disk is turbulent, and you and other objects can get thrown into different orbits or at each other. This might be a good thing, or it might not, depending on how the rules above apply to you.
• If the game progresses to the point where there is no more material within your reach and you are not yet a planet, then it’s game over.
• If you slow down too much (e.g., from bumping into other objects), you could spiral into the Sun (game over).
• If another planet gets big enough, it can:
  • Rip you apart and then swing the pieces around so fast that for the rest of the game you collide too hard with other pieces to grow any bigger (game over)
  • Fling you out of the solar system (game over)
  • Grab you for itself (game over)
  • Trap you in an orbit around it, turning you into a moon (game over, and incredibly humiliating)

Winners and Losers

The outcome of the game is evident in Figure 2.9. Today eight official winners are recognized, with Jupiter taking the grand prize, followed closely by Saturn. Both planets have trophy cases with more than 60 moons each, and each has a moon that is larger than Mercury. Prior to 2006, Pluto was also counted a winner, but in 2006 a controversial decision revoked Pluto’s planet status. The reason was a newly formalized definition of a planet, which stated that an object can only be considered a planet if it is massive enough to have swept its orbit clean of other bodies. Pluto is situated within the icy clutter of the Kuiper belt, so it does not fit this definition.

Pluto’s supporters have argued that Pluto should have been grandfathered in, given that the definition came after Pluto was declared a planet, but to no avail. Pluto has not given up, and on July 13, 2015, it launched an emotional plea with the help of the NASA’s New Horizons probe. New Horizons sent back images of Pluto’s heart (Figure 2.10). On closer inspection, Pluto’s heart was discovered to be broken.

Figure 2.10 Photographs of Pluto. Left: The heart-shaped region called Tombaugh Regio is outlined. This region is named after Pluto’s discoverer Clyde Tombaugh. Right: False-colour images show compositional variations in Tombaugh Regio. Source: Karla Panchuk (2015) CC BY 4.0. Left photo- NASA/APL/SwRI (2015) Public Domain view source, Right photo- NASA/APL/SwRI (2015) Public Domain. view source.

 The Accretion Game and the Solar System Today

The rules and dangers of the planet-forming game help to explain many features of our solar system today.

• Proximity to the sun explains why the terrestrial planets are so much smaller than the gas giant and ice giant planets.
• Mars is smaller than it should be, given the rule that distance from the sun determines how much material a body can accumulate, and this can be explained by its proximity to Jupiter. Jupiter’s immense gravity interfered with Mars’ ability to accrete. Further evidence of Jupiter’s interference is the debris field that forms the asteroid belt. From time to time, Jupiter still flings objects from the asteroid belt out into other parts of the solar system, some of which have collided with Earth to catastrophic effect.
• The Kuiper belt is an icy version of the asteroid belt, consisting of fragments left over from the early solar system. The material in the Kuiper belt is scattered because of Neptune’s gravity. From time to time, Jupiter interferes here as well, flinging Kuiper belt objects toward the sun and into orbit. As these objects approach the sun, the sun causes dust and gas to be blasted from their surface, forming tails. We know these objects as comets.
• Comets may also come from the Oort cloud where gravitational forces from outside of the solar system can hurl objects from the Oort cloud toward the sun.

If you were to get into a time machine and visit Earth shortly after it formed (around 4.5 billion years ago), you would probably regret it. Large patches of Earth’s surface would still be molten, which would make landing your time machine very dangerous indeed. If you happened to have one of the newer time-machine models with hovering capabilities and heat shields, you would still face the inconvenience of having nothing to breathe but a tenuous wisp of hydrogen and helium gas, and depending on how much volcanic activity was going on, volcanic gases such as water vapour and carbon dioxide. Some ammonia and methane might be thrown in just to make it interesting, but there would be no oxygen. Assuming you had the foresight to purchase the artificial atmosphere upgrade for your time machine, it would all be for naught if you materialized just in time to see an asteroid, or worse yet another planet, bearing down on your position. The moral of the story is that early Earth was a nasty place, and a time machine purchase is not something to take lightly.

Why was early Earth so nasty?

Earth Was Hot

Earth’s heat comes from the decay of radioactive elements within Earth, as well as from processes associated with Earth’s formation. Formation processes contributed heat in the following ways:

• Heat came from the thermal energy already contained within the objects that accreted to form Earth.
• Heat came from collisions. When objects hit Earth, some of the energy from their motion went into deforming Earth, and some of it was transformed into heat. Clap your hands vigorously to experience this on a much smaller (and safer!) scale.
• As Earth became larger, its gravitational force became stronger. This increased Earth’s ability to draw objects to it, but it also caused the material making Earth to be compressed, rather like Earth giving itself a giant gravitational hug. Compression causes materials to heat up.

Heating had an important consequence for Earth’s structure. As Earth grew, it collected a mixture of rocky silicate mineral grains as well as iron and nickel. These materials were scattered throughout Earth. That changed when Earth began to heat up: it got so hot that the metals melted and trickled down through the rocky silicate material toward Earth’s centre, becoming Earth’s core. The silicate material became Earth’s crust and mantle. In other words, Earth unmixed itself. The separation of silicate minerals and metals into a rocky outer layer and a metallic core, respectively, is called differentiation. Friction from metal melts moving through Earth caused it to heat up even more.

Earth’s high temperature early in its history also means that early tectonic processes were accelerated compared to today, and Earth’s surface was more geologically active.

Earth Was Bombarded by Objects from Space

Although Earth had swept up a substantial amount of the material in its orbit as it was accreting, unrest within the solar system caused by changes in the orbits of Saturn and Jupiter was still sending many large objects on cataclysmic collision courses with Earth. The energy from these collisions repeatedly melted and even vaporized minerals in the crust, and blasted gases out of Earth’s atmosphere. Very old scars from these collisions are still detectable, although we have to look carefully to see them. For example, the oldest impact site discovered is the 3 billion year old Maniitsoq “crater” in west Greenland, although there is no crater to see. What is visible are rocks that were 20 km to 25 km below Earth’s surface at the time of the impact, but which nevertheless display evidence of deformation that could only be produced by intense, sudden shock.

The evidence of the very worst collision that Earth experienced is not subtle at all. In fact, you have probably looked directly at it hundreds of times already, perhaps without realizing what it is. That collision was with a planet named Theia, which was approximately the size of Mars. Not long after Earth formed, Theia struck Earth (Figure 2.11). When Theia slammed into Earth, Theia’s metal core merged with Earth’s core, and debris from the outer silicate layers was cast into space, forming a ring of rubble around Earth. The material within the ring coalesced into a new body in orbit around Earth, giving us our moon. Remarkably, the debris may have coalesced in 10 years or fewer! This scenario for the formation of the moon is called the giant impact hypothesis.

Figure 2.11 Artist’s impression of a collision between planets. A similar collision between Earth and the planet Theia might have given us our moon. Source: NASA/ JPL-Caltech (2009) Public Domain. view source.

Today’s Atmosphere Took a Long Time to Develop

Earth’s first experiment with having an atmosphere did not succeed. It started out with a thin veil of hydrogen and helium gases that came with the material it accreted. But hydrogen and helium are very light gases, and they bled off into space.

Earth’s second experiment with having an atmosphere went much better. Volcanic eruptions built up the atmosphere by releasing gases. The most common volcanic gases are water vapour and carbon dioxide (CO₂), but volcanoes release a wide variety of gases. Other important contributions include sulphur dioxide (SO₂), carbon monoxide (CO), hydrogen sulphide (H₂S), hydrogen gas, and methane (CH₄). Meteorites and comets also brought substantial amounts of water and nitrogen to Earth. It is not clear what the exact composition of the atmosphere was after Earth’s second experiment, but carbon dioxide, water vapour, and nitrogen were likely the three most abundant components.

One thing we can say for sure about Earth’s second experiment is that there was effectively no free oxygen (O₂, the form of oxygen that we breathe) in the atmosphere. We know this in part because prior to 2 billion years ago, there were no rocks stained red from oxidized iron minerals. Iron minerals were present, but not in oxidized form. At that time, O₂ was produced in the atmosphere when the sun’s ultraviolet rays split water molecules apart. However, chemical reactions removed the oxygen as quickly as it was produced.

It wasn’t until well into Earth’s third experiment — life — that the atmosphere became oxygenated. Photosynthetic organisms used the abundant CO₂ in the atmosphere to manufacture their food, and released O₂ as a by-product. At first all of the oxygen was consumed by chemical reactions as before, but eventually the organisms released so much O₂ that it overwhelmed the chemical reactions. Oxygen began to accumulate in the atmosphere, although present levels of 21% oxygen didn’t occur until about 350 million years ago. Today the part of our atmosphere that isn’t oxygen consists largely of nitrogen (78%).

The oxygen-rich atmosphere on our planet is life’s signature. If geologic processes were the only ones controlling our atmosphere, it would consist mostly of carbon dioxide, like the atmosphere of Venus. It is an interesting notion (or a disconcerting one, depending on your point of view) that for the last 2 billion years the light reflected from our planet has been beaming a bar code out to the universe, similar to the ones in Figure 2.4, except ours says “oxygen.” For 2 billion years, our planet has been sending out a signal that could cause an observer from another world to say, “That’s odd… I wonder what’s going on over there.”

As of January 2018, 6,355 possible exoplanets– extrasolar planets, or planets outside of our solar system- have been detected by preliminary tests. Further tests have confirmed that 3,726 of those candidates are indeed planets. If “other Earths” are defined as planets where we could walk out of a spaceship with no equipment other than a picnic basket, and enjoy a pleasant afternoon on a grassy slope near a stream, then it remains to be seen whether any of these planets fit the description. On the other hand, if “other Earths” refers to rocky worlds approximately Earth’s size, and orbiting within their star’s habitable zone (the zone in which liquid water, and potentially life, can exist), then there is cautious optimism that we have found at least 53 such worlds.

Part of the uncertainty about the 53 possible Earth-like worlds is related to their composition. We don’t yet know their composition; however, it is tempting to conclude that they are rocky because they are similar in size to Earth. Remember the rules of the accretion game: you can only begin to collect gas once you are a certain size, and how much matter you collect depends on how far away from the sun you are. Given how large our gas giant and ice giant planets are compared to Earth, and how far away they are from the sun, we would expect that a planet similar in size to Earth, and a similar distance from its star, should be rocky.

But it isn’t quite as simple as that. We are finding that the rules to the accretion game can result in planetary systems very different from our own. For example, in the planetary systems we have observed, it is common to have planets larger than Earth orbiting closer to their star than Mercury does to the sun. Planets as large as Jupiter are rare, and where large planets do exist, they are much closer to their star than Jupiter is to the sun. To summarize, we need to be cautious about drawing conclusions from our own solar system, just in case we are basing those conclusions on something truly unusual.

On the other hand, the seemingly unique features of our solar system would make planetary systems like ours difficult to spot. One of the ways exoplanets are detected is by measuring the brightness of stars, and looking for tiny variations in brightness that could be caused by a planet passing between the star it orbits and the instrument observing the star. Small planets are harder to detect because they block less of a star’s light than larger planets. Larger planets farther from a star, like our gas giant planets, are difficult to spot because they don’t go past the star as frequently. For example, Jupiter goes around the sun once every 12 years. If someone were observing our solar system, they might have to watch for 12 years to see Jupiter go past the sun once. For Saturn, they might have to watch for 30 years.

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

2.1 Starting With a Big Bang

The universe began 13.8 billion years ago when energy, matter, and space expanded from a single point. Evidence for the big bang is the cosmic “afterglow” from when the universe was still very dense, and red-shifted light from distant galaxies, which tell us the universe is still expanding.

2.2 Forming Planets from the Remnants of Exploding Stars

The big bang produced hydrogen, helium, but heavier elements come from nuclear fusion reactions in stars. Large stars make elements such as silicon, iron, and magnesium, which are important in forming terrestrial planets. Large stars explode as supernovae and scatter the elements into space.

2.3 How to Build a Solar System

Solar systems begin with the collapse of a cloud of gas and dust. Material drawn to the centre forms a star, and the remainder forms a disk around the star. Material within the disk clumps together to form planets. In our solar system, rocky planets are closer to the sun, and ice and gas giants are farther away. This is because temperatures near the sun were too high for ice to form, but silicate minerals and metals could solidify.

2.4 Earth’s First 2 Billion Years

Early Earth was heated by radioactive decay, collisions with bodies from space, and gravitational compression. Heating caused molten metal to sink to Earth’s centre and form a core, and silicate minerals to form the mantle and crust. A collision with a planet the size of Mars knocked debris into orbit around Earth, and the debris coalesced into the moon. Earth’s atmosphere is the result of volcanic degassing, contributions by comets and meteorites, and photosynthesis.

2.5 Are There Other Earths?

The search for exoplanets has identified 53 planets that are similar in size to Earth and within the habitable zone of their stars. These are thought to be rocky worlds like Earth, but the compositions of these planets are not known for certain.

Review Questions

1. How can astronomers view events that happened in the universe’s distant past?

2. In this image of three spectra, one is from the sun, and the other two are from galaxies. One of the galaxies is the Andromeda galaxy. Which spectrum is from Andromeda?

Spectra for the sun and two galaxies. Source: Karla Panchuk (2015) CC BY 4.0.

3. Astronomers looking for some of the earliest stars in the universe were surprised to find a planetary system called HIP 11952, which existed 12.8 billion years ago. This was very early in the universe’s history, when stars still consisted largely of hydrogen and helium. Do you think there were terrestrial planets in this system? Why or why not?

4. Summarize the trends in size and composition of objects in the solar system.

5. What is the frost line, and what does it help to explain?

6.  Why is Pluto not considered a planet?

7. What is differentiation?

8. The exoplanet Kepler-452b is within the habitable zone of its star. In our solar system, planets a similar distance from the Sun are terrestrial planets. Why can we not say for certain that Kepler-452b’s distance from its star means it is a terrestrial planet?

9. Of the planetary systems discovered thus far, none are exactly like our solar system. Does this mean our solar system is unique in the universe?

1. To see an event, light from that event must reach our eyes. Light travels very quickly (about 300,000,000 m/s), but the universe is very, very large. Depending on how far away the event was, it could take billions of years for light to travel from the event to our eyes so we can see it. Astronomers take advantage of this fact to view the universe’s past.

2. B is the spectrum from the Andromeda galaxy. We know that one spectrum represents the sun, which is not moving toward or away from us. (Our orbit is not perfectly circular, but the small eccentricity is not a factor in this comparison.) We know that the Andromeda galaxy is on a collision course with us, so it is the exception to the rule that galaxies are moving away from us, and their light is red-shifted. That means the spectrum B which is shifted furthest to the left (blue-shifted) is Andromeda, and spectrum A which is furthest to the right (red-shifted) is a galaxy moving away from us. That means C is the sun.

Spectra for the sun and two galaxies. Source: Karla Panchuk (2015) CC BY 4.0

3. The planetary system consisted of two Jupiter-sized gas giant planets. Gas giant planets contain large amounts of hydrogen, and hydrogen was plentiful in the early universe. In contrast, terrestrial planets have heavier elements, especially silica, iron, magnesium, and nickel, that had yet to be manufactured by stars. Those elements were not present in sufficient abundance to form terrestrial planets until much later.

4. Closest to the sun we find the small, rocky, terrestrial planets with metal cores. Further out are the gas giant planets, which are the largest in the solar system. They consist mostly of hydrogen, and have cores of rock and ice. Beyond the gas giant planets are the ice giant planets, which are next largest. They have a mantle of ice (not just water ice but ammonia and methane ice), and a rocky core. Smaller objects in the solar system include rocky bodies within the asteroid belt between Mars and Jupiter, and bodies of ice and dust in the Kuiper belt and Oort cloud beyond Neptune.

5. The frost line marks the distance from the sun beyond which temperatures were cool enough to allow ice to form. This helps to explain why the terrestrial planets are closer to the sun, and the Jovian and ice giant planets farther away. Mineral grains could solidify and begin to accrete closer to the sun, forming terrestrial planets, because they have higher melting points. In contrast, water vapour, methane, and ammonia had to be farther from the sun before they could freeze and begin to accrete.

6. Planets are defined as having cleared their orbits of debris. Pluto is located within the Kuiper belt, so it shares its orbit with other objects. There are two other criteria in the definition of a planet: planets in our solar system must orbit the sun, and they must have a spherical shape. Pluto satisfies both these criteria, but sadly the people deciding whether or not Pluto should be a planet are not amenable to a “best two out of three” compromise.

7. Differentiation is the separation of materials within a planet such that dense materials sink to the core. In Earth’s case, the denser materials are iron and nickel.

8. The fact that we have terrestrial planets close to the sun makes sense in terms of the frost line, but it does not seem to be a hard-and-fast rule in other planetary systems. Therefore, we can’t conclude from Kepler-452b’s position alone that it is a terrestrial planet.

9. The rules of the accretion game mean that there are many complex interactions, so even a small difference in the starting conditions or in how the game goes in the beginning could have major implications in the end. For that reason, we shouldn’t expect to find a planetary system that matches ours in every minute detail. However, just because we haven’t found a similar planetary system does not mean one does not exist. Our planet-finding methods are biased toward discovering large planets orbiting close to their stars, whereas our solar system has small planets close to the sun and larger ones farther away. That doesn’t mean our methods won’t eventually turn up a system like ours, just that they are more likely to turn up systems that are different.

Earth consists of three main layers: the crust, the mantle, and the core (Figure 3.4).  The core accounts for almost half of Earth’s radius, but it amounts to only 16.1% of Earth’s volume.  Most of Earth’s volume (82.5%) is its mantle, and only a small fraction (1.4%) is its crust.

Figure 3.4 Earth’s interior. Right- crust, mantle, and outer and inner core to scale.  Left- Cutaway showing continental and ocean crust, and upper mantle layers. The lithosphere is the crust plus the uppermost layer of the mantle. Source: Karla Panchuk (2018) CC BY 4.0. Earth photo by NASA (n.d.) Public Domain view source

Crust

The Earth’s outermost layer, its crust, is rocky and rigid. There are two kinds of crust: continental crust, and ocean crust. Continental crust is thicker, and predominantly felsic in composition, meaning that it contains minerals that are richer in silica. The composition is important because it makes continental crust less dense than ocean crust.

Ocean crust is thinner, and predominantly mafic in composition.  Mafic rocks contain minerals with less silica, but more iron and magnesium. Mafic rocks (and therefore ocean crust) are denser than the felsic rocks of continental crust.

The crust floats on the mantle.  Continental crust floats higher in the mantle than ocean crust because of the lower density of continental crust.  An important consequence of the difference in density is that if tectonic plates happen to bring ocean crust and continental crust into collision, the plate with ocean crust will be forced down into the mantle beneath the plate with continental crust.

Mantle

The mantle is almost entirely solid rock, but it is in constant motion, flowing very slowly. It is ultramafic in composition, meaning it has even more iron and magnesium than mafic rocks, and even less silica.  Although the mantle has a similar chemical composition throughout, it has layers with different mineral compositions and different physical properties.  It can have different mineral compositions and still be the same in chemical composition because the increasing pressure deeper in the mantle causes mineral structures to be reconfigured.

Rocks higher in the mantle are typically composed of peridotite, a rock dominated by the minerals olivine and pyroxene. The Tablelands rock in Figure 3.2 is a type of peridotite. Lower in the mantle, extreme pressures transform minerals and create rocks like eclogite (Figure 3.5), which contains garnets.

Figure 3.5 Eclogite from the Swiss-Italian Alps. Reddish brown spots are garnets. Source: James St. John (2014) CC BY 2.0 view source

Lithosphere

The lithosphere can’t be classified neatly as either crust or mantle because it consists of both.  It is formed from the crust as well as the uppermost layer of the mantle which is stuck to the underside of the crust.  Tectonic plates are fragments of lithosphere.

Asthenosphere

Beneath the lithosphere is the asthenosphere.Tiny amounts of melted rock dispersed through the otherwise solid asthenosphere make the asthenosphere weak compared to the lithosphere. The weakness of the asthenosphere is important for plate tectonics because it deforms as fragments of lithosphere move around upon and through it. Without a weak asthenosphere, plates would be locked in place, unable to move as they do now.

D”

The D” (dee double prime) layer is a mysterious layer beginning approximately 200 km above the boundary between the core and mantle.  (This boundary is referred to as the core-mantle boundary.)  We know it exists because of how seismic waves change speed as they move through it, but it isn’t clear why it’s different from the rest of the mantle.  One idea is that it is minerals are undergoing another transition in this region because of pressure and temperature conditions, similar to the transition between the upper and lower mantle. Other ideas are that small pools of melt are present, or that the differences in seismic properties are due to subducted slabs of lithosphere resting on the core-mantle boundary.

Core

The core is primarily composed of iron, with lesser amounts of nickel. Lighter elements such as sulfur, oxygen, or silicon may also be present. The core is extremely hot (~3500° to more than 6000°C). But despite the fact that the boundary between the inner and outer core is approximately as hot as the surface of the sun, only the outer core is liquid. The inner core is solid because the pressure at that depth is so high that it keeps the core from melting.

Seismology is the study of vibrations within Earth. These vibrations are caused by events such as earthquakes, extraterrestrial impacts, explosions, storm waves hitting the shore, and tides. Seismology is applied to the detection and study of earthquakes, but seismic waves also provide important information about Earth’s interior.

Seismic waves travel through different materials at different speeds, and we can apply knowledge of how they interact with different materials to understand Earth’s layers and internal structures. Similar to the way that ultrasound is used to image the human body, we can measure how long it takes for seismic waves to travel from their source to a recording station.

Another feature of seismic waves is that some, called P-waves, can travel rapidly though both liquids and solids, but others, called S-waves, can only travel though solids, and are slower than P-waves. Observing where P-waves travel, and S-waves do not, allows us to identify regions within Earth that are melted.

Seismic Wave Paths

Seismic waves travel in all directions from their source, but it is convenient to imagine the path traced by one point on the wave front, and represent that path as a seismic ray (arrows, Figure 3.6).

Figure 3.6 Seismic waves and seismic rays. The paths of seismic waves can be represented as rays. Seismic ray paths are bent when they enter a rock layer with a different seismic velocity. Source: Karla Panchuk (2018) CC BY 4.0

When seismic waves encounter a different rock layer, some might bounce off the layer, or reflect, as in the bottom layer of Figure 3.6. But some waves will travel through the layer. If the wave travels at a different speed in the new layer, its path will be bent, or refracted, as it crosses into the new layer. If the wave can travel faster in the new layer, it will be bent slightly toward the contact between the two layers. In Figure 3.6, the ray can travel progressively faster in each layer as it goes down through the layers, and it is bent slightly upward each time it crosses into the next layer. The reverse happens if the wave slows down. On the right side of the diagram, the wave is moving upward through slower and slower layers. It is bent away from the faster layer each time, causing it to take a more direct path to the surface.

Seismic velocities are higher in more rigid layers, so broadly speaking, they get faster deeper within Earth, because higher pressures make layers more rigid. They tend to take curved paths through the Earth because refraction bends their path until they are reflected and directed upward again, as in Figure 3.6.

Discoveries with Seismic Waves

The Moho: Where Crust Meets Mantle

One of the first discoveries about Earth’s interior made through seismology was in the early 1900s by Croatian seismologist Andrija Mohorovičić (pronounced Moho-ro-vi-chich).  He noticed that sometimes, seismic waves arrived at seismic stations (measuring locations) farther from an earthquake before they arrived at closer ones.  He reasoned that the waves that traveled farther were faster because they bent down and traveled faster through different rocks (those of the mantle) before being bent upward back into the crust (Figure 3.7).

Figure 3.7 Depiction of seismic waves emanating from an earthquake (red star). Some waves travel through the crust to the seismic station (at ~6 km/s), while others go down into the mantle (where they travel at ~8 km/s) and are bent upward toward the surface, reaching the station before the ones that travelled only through the crust. Source: Steven Earle (2016) CC BY 4.0 view source

The boundary between the crust and the mantle is now known as the Mohorovičić discontinuity (or Moho). Its depth is between 60 – 80 km beneath major mountain ranges, 30 – 50 km beneath most of the continental crust, and 5 – 10 km beneath ocean crust.

The Core-Mantle Boundary

Arguments for a liquid outer core were supported by a distinctive signature in the global distribution of seismic waves from earthquakes. When an earthquake occurs, there is a zone on the opposite side of Earth where S-waves are not measured. This S-wave shadow zone begins 103° on either side of the earthquake, for a total angular distance of 154° (Figure 3.8, left). There is also a P-wave shadow zone on either side of the earthquake, from 103° to 150° (Figure 3.8, right).

Figure 3.8 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. P-waves do travel through the core, but because the waves that enter the core are refracted, there are also P-wave shadow zones. Source: Steven Earle (2016) CC BY 4.0 view source

The S-wave shadow zone occurs because S-waves cannot travel through the liquid outer core. The P-wave shadow zone occurs because seismic velocities are much lower in the liquid outer core than in the overlying mantle, and the P-waves are refracted in a way that leaves a gap. Not only do the shadow zones tell us that the outer core is liquid, the size of the shadow zones allows us to calculate the size of the core, and the location of the core-mantle boundary.

Seismic Portrait of Earth’s Layers

The change seismic wave velocity with depth in Earth (Figure 3.9) has been determined over the past several decades by analyzing seismic signals from large earthquakes all around the world. Earth’s layers are detectable as changes in velocity with depth. The asthenosphere is visible as a low velocity zone within the upper mantle (Figure 3.9, left). There is an abrupt increase in P-wave velocity at 420 km, showing the depth at which minerals transform into structures that are more stable at higher pressures and temperatures.

Figure 3.9 P-wave and S-wave velocity variations with depth from the crust through the upper mantle (left) and from the crust through to the core (right). Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2016) CC BY 4.0 view source left/ right

The boundary between the upper and lower mantle is visible at 660 km as a sudden change from rapidly increasing P- and S-wave velocities to slow or no change in P-wave and S-wave velocities (Figure 3.9, right). The core-mantle boundary (CMB in Figure 3.9) is apparent as a sudden drop in P-wave velocities, where seismic waves move from solid mantle to liquid outer core. The boundary between the outer core and inner core is marked by a sudden increase in P-wave velocity after 5000 km, where seismic waves move from a liquid back into a solid again.

Seismic Images of Plate Tectonic Structures

Using data from many seismometers and hundreds of earthquakes, it is possible to create images from the seismic properties of the mantle. This technique is known as seismic tomography. Tomography can be used to map out slabs of lithosphere that are entering the mantle, or have disappeared within it. Those slabs are cooler, and therefore more rigid than surrounding mantle rocks, so seismic waves travel through them faster. In Figure 3.10, higher-than-average seismic velocities in cool slabs are indicated in dark blue.

Figure 3.10 P-waves and S-waves used to map out the location of the Cocos slab of lithosphere. The slab appears in dark blue, indicating higher than average seismic wave velocities. Left- Tomograms showing seismic wave anomalies for a 1290 km surface. Right- Cross-sections along the transect marked X-Y on the globe. Source: Karla Panchuk (2018) CC BY 4.0, modified after van der Meer et al. (2018) CC BY 4.0 view source

Thanks to the tomograms, we can see that the Cocos plate, which is colliding with Central America, is part of a much larger slab of lithosphere that has already settled onto the mantle. Tomograms representing a surface at 1290 km depth (Figure 3.10, left) show that at that level, the Cocos slab is beneath the Caribbean Sea. The tomograms on the right show a vertical view along the line X-Y marked on the globe. The vertical tomograms show us that the Cocos slab extends all the way down to the core-mantle boundary.

References

van der Meer, D.G., van Hinsbergen, D.J.J., and Spakman, W., (2018). Atlas of the Underworld: slab remnants in the mantle, their sinking history, and a new outlook on lower mantle viscosity. Tectonophysics 723, p. 309-448. https://doi.org/10.1016/j.tecto.2017.10.004

Earth Gets Hotter the Deeper You Go

Earth’s temperature increases with depth, but not at a uniform rate (Figure 3.11). Earth’s geothermal gradient is 15° to 30°C/km within the crust.  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 approximately 1000°C at the base of the crust, around 3500°C at the base of the mantle, and approximately 6,000°C at Earth’s centre.

Figure 3.11 Geothermal gradient (change in temperature with depth). Left- Geothermal gradient in the crust and upper mantle. The geothermal gradient remains below the melting temperature of rock, except in the asthenosphere. There, temperatures are high enough to melt some of the minerals. Right- Geothermal gradient throughout Earth. Rapid changes occur in the uppermost mantle, and at the core-mantle boundary. Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2016) CC BY 4.0 view source left/ right

The temperature gradient within the lithosphere varies depending on the tectonic setting. Gradients are lowest in the central parts of continents, higher where plates collide, and higher still at boundaries where plates are moving away from each other.

In spite of high temperatures within Earth, mantle rocks are almost entirely solid. High pressures keep them from melting. The red dashed line in Figure 3.11 (right) shows the minimum temperature at which dry mantle rocks will melt. Rocks at temperatures to the left of the line will remain solid. In rocks at temperatures to the right of the line, some minerals will begin to melt. Notice that the red dashed line goes further to the right for greater depths, and therefore greater pressures. Now compare the geothermal gradient with the red dashed line. The geothermal gradient is to the left of the red line, except in the asthenosphere, where small amounts of melt are present.

Convection Helps to Move Heat Within Earth

The fact that the temperature gradient is much lower in the main part of the mantle than in the lithosphere has been interpreted as evidence of convection in the mantle. When the mantle convects, heat is transferred through the mantle by physically moving hot rocks. Mantle convection is the result of heat transfer from the core to the base of the lower mantle. As with a pot of soup on a hot stove (Figure 3.12), the material near the heat source (the soup at the bottom of the pot) becomes hot and expands, making it less dense than the material above. Buoyancy causes it to rise, and cooler material flows in from the sides. Of course, convection in the soup pot is much faster than convection in the mantle. Mantle convection occurs at rates of centimetres per year.

Figure 3.12 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. Source: Steven Earle (2015) CC-BY 4.0 view source

Convection carries heat to the surface of the mantle much faster than heating by conduction. Conduction is heat transfer by collisions between molecules, and is how heat is transferred from the stove to the soup pot. A convecting mantle is an essential feature of plate tectonics, because the higher rate of heat transfer is necessary to keep the asthenosphere weak. Earth’s mantle will stop convecting 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. When mantle convection stops, the end of plate tectonics will follow.

Models of Mantle Convection

In the soup pot example, convection moves hot soup from the bottom of the pot to the top. Some geologists think that Earth’s convection works the same way— hot rock from the base of the mantle moves all the way to the top of the mantle before cooling and sinking back down again. This view is referred to as whole-mantle convection (Figure 3.8, left). Other geologists think that the upper and lower mantle are too different to convect as one. They point to slabs of lithosphere that are sinking back into the mantle, some of which seem to perch on the boundary between the upper and lower mantle, rather than sinking straight through. They also note chemical differences in magma originating in different parts of the mantle— differences that are not consistent with the entire mantle being well stirred. They argue that double-layered convection is a better fit with the observations (Figure 3.13, right).  Still others argue that there may be some locations where convection goes from the bottom of the mantle to the top, and some where it doesn’t (Figure 3.13, middle).

Figure 3.13 Models of mantle convection. Left- whole mantle convection. Rocks rise from the core-mantle boundary to the top of the mantle, then sink to the bottom again. Right- Two-layer convection, in which upper and lower mantle convect at different rates. Middle- Convection paths vary depending on the circumstances. Source: Karla Panchuk (2018) CC BY 4.0

Why Is Earth Hot Inside?

The heat of Earth’s interior comes from a variety of sources. These include the heat contained in the objects that accreted to form Earth, and the heat produced when they collided. As Earth grew larger, the increased pressure on Earth’s interior caused it to compress and heat up.  Heat also came from friction when melted material was redistributed within Earth, forming the core and mantle.

A major source of Earth’s heat is radioactivity, the energy released when the unstable atoms decay. The radioactive isotopes uranium-235 (²³⁵U), uranium-238 (²³⁸U), potassium-40 (⁴⁰K), and thorium-232 (²³²Th) in Earth’s mantle are the primary source. Radioactive decay produced more heat early in Earth’s history than it does today, because fewer atoms of those isotopes are left today (Figure 3.14). Heat contributed by radioactivity is now roughly a quarter what it was when Earth formed.

Figure 3.14 Production of heat within the Earth over time by radioactive decay of uranium, thorium, and potassium. Heat production has decreased over time as the abundance of radioactive atoms has decreased. Source: Steven Earle (2015) CC BY 4.0 view source, modified after Arevalo et al. (2009)

References

Arevalo, R., McDonough, W., & Luong, M. (2009). The K/U ratio of Earth: Insights into mantle composition, structure and thermal evolution. Earth and Planetary Science Letters, 278(3-4), 361-369. https://doi.org/10.1016/j.epsl.2008.12.023

Earth’s liquid iron core convects because it is heated from beneath by the inner core. Because iron is a metal and conducts electricity (even when molten), its motion generates a magnetic field.

Earth’s magnetic field is defined by north and south poles representing lines of magnetic force flowing into Earth in the northern hemisphere and out of Earth in the southern hemisphere (Figure 3.15). Because of the shape of the field lines, the magnetic force is oriented at different angles to the surface in different locations. The tilt, or inclination of magnetic field lines is represented by the tilt of compass needles in Figure 3.15. At the north and south poles, the force is vertical. The force is horizontal at the equator. Everywhere in between, the magnetic force is at an intermediate angle to the surface.

Figure 3.15 Earth’s magnetic field depicted as the field of a bar magnet coinciding with the core. The south pole of the magnet points to Earth’s magnetic north pole. The red and white compass needles represent the orientation of the magnetic field at various locations on Earth’s surface. Source: Karla Panchuk (2018) CC BY-SA 4.0, modified after Steven Earle (2015) CC BY-SA 4.0 view source, and T. Stein (2008) CC BY-SA 3.0 view source

Earth’s magnetic field is generated within the outer core by the convective movement of liquid iron, but although convection is continuous, the magnetic field is not stable. Periodically, the magnetic field decays and then becomes re-established. When it does re-establish, the polarity may have reversed (i.e., your compass would point south rather than north). Over the past 250 Ma, there have been hundreds of magnetic field reversals, and their timing has been anything but regular. The shortest ones that geologists have been able to identify lasted only a few thousand years, and the longest one was more than 30 million years, during the Cretaceous Period (Figure 3.16).

Figure 3.16 Magnetic field reversal chronology for the past 170 Ma. Black stripes mark times when the magnetic field was oriented the same as today. Source: Steven Earle (2015) CC BY 4.0 view source, modified after AnomieX (2010) Public Domain view source

Changes in Earth’s magnetic field have been studied using mathematical models that simulate convection in the outer core (Figure 3.17).  Reversals happened spontaneously when the model was run to simulate a period of several hundred thousand years. Spontaneous reversals can happen because convection does not occur in an orderly way, in spite of what the bar magnet analogy may suggest. Many small-scale variations occur in convection patterns within the inner core, and Earth’s magnetic field over all is the sum of those variations. Magnetic reversals do not happen as frequently as they might, if not for the solid inner core. Magnetic field changes take much longer within the inner core, so reversals in the outer core do not always coincide with reversals in the inner core. Both are required in order for Earth’s magnetic field to flip.

Figure 3.17 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. Source: NASA (2007) Public Domain view source

References

British Geological Survey, Natural Environment Research Council (n.d.). Reversals: Magnetic Flip. Visit website

Glatzmaier, G. A. (n.d.) The Geodynamo. Visit website

Lithospheric Plates Float on the Mantle

The mantle is able to convect because it can deform by flowing over very long timescales. This means that tectonic plates are floating on the mantle, like a raft floating in the water, rather than resting on the mantle like a raft sitting on the ground. How high the lithosphere floats will depend on the balance between gravity pulling the lithosphere down, and the force of buoyancy as the mantle resists the downward motion of the lithosphere. Isostasy is the state in which the force of gravity pulling the plate toward Earth’s centre is balanced by the resistance of the mantle to letting the plate sink.

To see how isostasy works, consider the rafts in Figure 3.18. The raft on the right is sitting on solid concrete. The raft will remain at the same elevation whether there are two people on it, or four, because the concrete is too strong to deform. In contrast, isostasy is in play for the rafts on the left, which are floating in 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. Peanut butter, rather than water, is used in this example because the viscosity of peanut butter (its stiffness or resistance to flowing) more closely represents the relationship between the tectonic plates and the mantle. Although peanut butter has a similar density to water, it’s higher viscosity means that if a person is added to a raft, it will take longer for the raft to settle lower into the peanut butter that it would take the raft to sink into water.

Figure 3.18 Illustration of isostatic relationships between rafts and peanut butter (left), and a non-isostatic relationship between a raft and solid ground (right). Source: Steven Earle (2015) CC BY 4.0 view source

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 g/cm³, is less dense than the mantle (average density of ~3.4 g/cm³ near the surface, but more at depth), and so it is floating on the mantle. When weight is added to the crust through the process of mountain building, the crust slowly sinks deeper into the mantle, and the mantle material that was there is pushed aside (Figure 3.19, left). When erosion removes material from the mountains over tens of millions of years, decreasing the weight, the crust rebounds and the mantle rock flows back (Figure 3.19, right).

Figure 3.19 Isostatic relationship between the crust and the mantle. Mountain building adds mass to the crust, and the thickened crust sinks down into the mantle (left). As the mountain chain is eroded, the crust rebounds (right). Green arrows represent slow mantle flow. Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2016) CC BY 4.0 view source

Isostasy and Glacial Rebound

The crust and mantle respond in a similar way to glaciation. Thick accumulations of glacial ice add weight to the crust, and the crust subsides, pushing the mantle out of the way. The Greenland ice sheet, at over 2,500 m thick, has depressed the crust below sea level (Figure 3.20a). When the ice eventually melts, the crust and mantle will slowly rebound (Figure 3.20b), but full rebound will likely take more than 10,000 years (3.20c).

Figure 3.20 Cross-section through the crust in the northern part of Greenland. a) Up to 2,500 m of ice depresses the crust downward (red arrows) and below sea level. b) After complete melting. Isostatic rebound would be slower than the rate of melting, leaving central Greenland at or below sea level for thousands of years. c) Complete rebound after ~10,000 years raises central Greenland above sea level again. Source: Steven Earle (2015) CC BY 4.0 view source a/ b/ c

Large parts of Canada are still rebounding as a result of the loss of glacial ice over the past 12,000 years, as are other parts of the world (Figure 3.21). The highest rate of uplift is in a large area west of Hudson Bay, where the Laurentide Ice Sheet was the thickest, at over 3,000 m. Ice finally left this region around 8,000 years ago, and the crust is currently rebounding at nearly 2 cm/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.

Glacial rebound in one location means subsidence in surrounding areas (Figure 3.21, yellow through red regions). Regions surrounding the former Laurentide and Fenno-Scandian Ice Sheets that were lifted up when mantle rock was forced aside and beneath them are now subsiding as the mantle rock flows back.

Figure 3.21 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. Source: Steven Earle (2015) CC BY 4.0 view source, modified after Erik Ivins, JPL (2010) Public Domain view source

How Can the Mantle Be Both Solid and Plastic?

You might be wondering how it is possible that Earth’s mantle is solid, rigid rock, 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 non-Newtonian behaviour is the deformation of Silly Putty, which can bounce when it is compressed rapidly when dropped, and will break if you pull on it sharply. But will deform in a liquid manner if stress is applied slowly. The force of gravity applied over a period of hours can cause it to deform like a liquid, dripping through a hole in a glass tabletop (Figure 3.22). Similarly, the mantle will flow when placed under the slow but steady stress of a growing (or melting) ice sheet.

Figure 3.22 Silly Putty exhibiting plastic behavior when acted upon by gravity over several hours. Source: Erik Skiff (2006) CC BY-SA 2006 view source

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

3.1 Earth’s Layers

Earth is divided into a rocky crust and mantle, and a core consisting largely of iron. The crust and the uppermost mantle form the lithosphere, which is broken into tectonic plates.  The next layer, the asthenosphere, allows the plates to move because it deforms by flowing.

3.2 Imaging Earth’s Interior

Seismic waves that travel through Earth are either P-waves or S-waves. P-waves are faster than S-waves, and can pass through fluids. Earth’s layers can be identified by looking at changes in the velocity of seismic waves. Seismic wave shadow zones contributed to knowledge of the depth of the core-mantle boundary, and the knowledge that the outer core is liquid.  Plate tectonic structures within Earth can also be mapped using the seismic waves generated by earthquakes.

3.3 Earth’s Interior Heat

Earth’s temperature increases with depth (to around 6000°C at the centre), but the rate of increase is not the same everywhere. In the lithosphere, thickness and plate tectonic setting are are factors. Deeper within the mantle, convection currents are more important.

3.4 Earth’s Magnetic Field

Earth’s magnetic field is generated by convection of the liquid outer core. The magnetic field is similar to that of a bar magnet, and has force directions that vary with latitude. The polarity of the field is not constant, meaning that the positions of the north and south magnetic poles have flipped from “normal” (as it is now) to reversed and back many times in Earth’s history.

3.5 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 and heavier because of mountain building, it pushes farther down into the mantle. Oceanic crust, being denser than continental crust, floats lower on the mantle than continental crust.

Questions for Review

1. What parts of Earth are most closely represented by typical stony meteorites and typical iron meteorites?
2. Draw a simple diagram of Earth’s layers, and label the approximate locations of the following boundaries: crust/mantle, mantle/core, outer core/inner core.
3. How do P-waves and S-waves differ?
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 causes mantle convection?
7. 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. British Columbia is still experiencing weak post-glacial isostatic uplift, especially in the interior, but also along the coast (see Figure 3.21). Meanwhile, offshore areas are experiencing weak isostatic subsidence. Why?

1. Stony meteorites are similar in composition to Earth’s mantle, while iron meteorites are similar to the core.

2. Compare your answer to Figure 3.4.

3. P-waves can pass through a liquid, and travel approximately twice as fast as S-waves (which cannot pass through a liquid).

4. P-wave velocity decreases at the core-mantle boundary because the outer core is liquid.

5. The mantle gets increasingly dense and strong with depth because of the increasing pressure. This difference affects both P-wave and S-wave velocities, and they are refracted toward the lower density mantle material (meaning they are bent out toward Earth’s surface).

6. The key evidence for mantle convection is that the rate of temperature increase with depth within the mantle is less than expected. This can only be explained by a mantle that is mixing by convection. The mechanism for convection is the transfer of heat from the core to the mantle, causing the to mantle flow.

7. Earth’s magnetic field is generated within the liquid outer core because liquid metal is convecting.

8. The last two reversals of Earth’s magnetic field were at the beginning of the present Brunhes normal chron (0.78 Ma), and at the end of the Jaramillo normal subchron (0.90 Ma). 

9. The isostatic relationship between the crust and the mantle is dependent on the fact that over very long timescales, the mantle deforms by flowing.

10. In the area of the Rocky Mountains the crust is thickened and pushed down into the mantle. In Saskatchewan the crust is thinner and does not extend as far into the mantle.

11. During the Pleistocene glaciation, British Columbia was pushed down by glacial ice. Mantle rock flowed slowly out from under the weighted-down crust and toward the ocean floor. Now that the land area is rebounding, that mantle rock is flowing back and the offshore areas are subsiding.

Alfred Wegener (1880-1930; Figure 4.2) earned a PhD in astronomy at the University of Berlin in 1904, but had a keen interest in geophysics and meteorology, and focused on meteorology for much of his academic career.

Figure 4.2 Alfred Wegener during a 1912-1913 expedition to Greenland. Source: Alfred Wegener Institute (2008) Public Domain view source

In 1911 Wegener happened upon a scientific publication that described matching Permian-aged terrestrial fossils in various parts of South America, Africa, India, Antarctica, and Australia.  He concluded that because these organisms could not have crossed the oceans to get from one continent to the next, the continents must have been joined in the past, permitting the animals to move from one to the other (Figure 4.3).  Wegener envisioned a supercontinent made up of all the present day continents, and named it Pangea (meaning “all land”). He described the motion of the continents reconfiguring themselves as continental drift.

Figure 4.3 The distribution of several Permian terrestrial fossils that are present in various parts of continents now separated by oceans. During the Permian, the supercontinent Pangea included the supercontinent Gondwana, shown here, along with North America and Eurasia. Source: J.M. Watson, USGS (1999) Public Domain view source

Wegener pursued his idea with determination, combing libraries, consulting with colleagues, and making observations in an effort to find evidence in support of it. He relied heavily on matching geological patterns across oceans, such as sedimentary strata in South America matching those in Africa, North American coalfields matching those in Europe, the mountains of Atlantic Canada matching those of northern Britain—both in structure and rock type—and comparisons of rocks in the Canadian Arctic with those of Greenland (Figure 4.4).

Figure 4.4 Diagram from Alfred Wegener’s book Die Entstehung der Kontinente und Ozeane comparing rock types on Canadian Arctic Islands and Greenland. Source: Karla Panchuk (2018) CC BY 4.0. Click the image for more attributions.

Wegener also called upon evidence for the Carboniferous and Permian (~300 Ma) Karoo Glaciation from South America, Africa, India, Antarctica, and Australia (Figure 4.5). 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 m per year, although he admitted that the measurements were not accurate. (The separation rate is actually about 2.5 cm per year.)

Figure 4.5 Carboniferous and Permian Karoo Glaciation in the southern hemisphere. Paleogeographic reconstruction for 306 million years ago. Source: Cropped from C. R. Scotese, PALEOMAP Project (www.scotese.com) view source. Click the image for terms of use.

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.

The main criticism of Wegener’s idea was that he could not explain how continents could move. Remember that, as far as anyone was concerned, Earth’s crust was continuous, not broken into plates. Thus, any mechanism Wegener could think of would have to fit with that model of Earth’s structure. Geologists at the time were aware that continents were made of different rocks than the ocean crust, and that the material making up the continents was less dense, so Wegener proposed that the continents were like icebergs floating on the heavier ocean crust.  He suggested that the continents were moved by the effect of Earth’s rotation pushing objects toward the equator, and by the lunar and solar tidal forces, which tend to push objects toward the west. However, 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 a small minority of geologists, and firmly rejected by most. But within a few decades that was all to change.

Resources

On The Shoulders of Giants: Alfred Wegener

References

Wegener, A. (1920). Die Entstehung der Kontinente und Ozeane. Braunschweig, Germany: Friedr. Vieweg & Sohn. Full text at Project Gutenberg

The untimely death of Alfred Wegener did not 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, 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 continue, 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. Oceans formed above parts of former continents that had settled downward and become submerged.

While this hypothesis helped to address the dilemma of the terrestrial fossils by explaining how continents once connected could now be separated by oceans, it came with its own set of problems.  One problem was that Earth wasn’t cooling fast enough to create the necessary amount of shrinking.  Another problem was the principle of isostasy (already understood for several decades; see Section 3.5 for a review of isostasy), which wouldn’t allow blocks of continental crust to sink in the way necessary for oceans to form in accordance with contractionist theory.

Another widely held view was permanentism, the idea that the continents and oceans have always been generally the same as they are today. This view incorporated a mechanism for creation of mountain chains known as the geosyncline theory. A geosyncline is a thick (potentially 1000s of metres) deposit of sediments and sedimentary rocks, typically situated along the edge of a continent, and derived from continental weathering (Figure 4.6).

Figure 4.6 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 rocks.) Source: Steven Earle (2015) CC BY 4.0 view source

The idea that geosynclines developed into fold-belt mountains originated in the middle of the 19th century. It was first proposed by James Hall and later elaborated upon 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 believed to involve compression by forces pushing from either side, causing sedimentary layers within the geosyncline to fold up. In 1937, Philip Kuenen published a paper of experiments with layers of paraffin wax to test how this might work. He was able to cause layers within a geosyncline to fold up as the geosyncline deepened and became more tightly folded during the experiment (Figure 4.7).

Figure 4.7 Simulation of mountain building within a geosyncline using layers of wax. Left- A sequence of photographs showing deformation in the wax layers as pistons apply increasing amounts of compression from the side. Right- Close-up view of slices through the wax layers at the end of the experiment, showing that stiffer white layers of wax folded in a way that resembled the folds in mountain belts. Source: Karla Panchuk (2018) CC BY 4.0. Photographs from Kuenen (1937) Public Domain view source.

The problem with the geosynclinal hypothesis for mountain building is that the lateral forces required to cause the compression were never adequately explained. Kuenen compressed the wax layers in his experiment by using pistons that pushed horizontally from either side, but described that mechanism as unrealistic. He explained that the pistons were just to get the process started within the experiment, and that in nature the main force was likely that of gravity pulling the geosyncline downward, drawing the sides together as it folded. When the sides were drawn together, this provided the compression to fold the sediments within the geosyncline. He couldn’t specify how the process got started in nature, but suggested that there could be a variety of reasons for an irregularity in the crust to respond to forces in such a way that would trigger downward sagging and folding.

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 explanation offered was that land bridges had once linked the continents, permitting animals and plants to 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 1918-20 edition of the Encyclopedia Americana:

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.

A problem with the land-bridge hypothesis is that there is no evidence of land bridges that could account for the fossil distribution patterns. The world’s oceans are approximately 4 km deep on average, so the underwater slopes leading up to a land bridge would have to have been at least 10s of km wide in most places, and many times that in others. Even if flooded, a land bridge of that size would still be visible in the shape of ocean-floor terrain.  Isostasy would not permit such a land bridge to sink down without leaving a trace.

We do know the locations of some past land bridges, but they were very different from the ones that would be required for this hypothesis. They are bridges such as the flooded Bering Strait land bridge, which is beneath only 30 to 50 m of water, and was exposed when sea levels were much lower because of water being locked in polar ice caps during the last major glaciation event.  The narrowest point of the Bering Strait is 82 km wide. The shortest distance between South America and Africa is more than 2800 km.

Exercise: Fitting the Continents Together

The main continents around the Atlantic Ocean are shown in Figure 4.8 with 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. Refer to a map of Pangea to help you make the fit. Note that the fit of the continents is even better than this, as distortions are introduced when rendering Earth onto a flat map. A better fit could be accomplished if you were to do this exercise upon the surface of a globe.

Figure 4.8 Mesozoic continent shapes. Source: Steven Earle (2015) CC BY 4.0 view source

References

Ingersoll, E. (1919). Land-Bridges Across the Oceans. In The Encyclopedia Americana (Vol. XVI, pp. 692-694). New York, NY: Encyclopedia Americana Corporation. Full text

Kuenen, P. H. (1937) The negative isostatic anomalies in the East Indies (with Experiments). Leidse Geologische Mededelingen 8(2), 169-214. Full text

Two key areas of research ultimately led to the acceptance of continental drift, and the formulation of plate tectonic theory.  One was the study of paleomagnetism, the record of Earth’s magnetic field through time.  The other was exploration of the ocean floor.

Paleomagnetism (Remnant Magnetism)

Figure 4.9 Rock layers recording remnant magnetism. The red arrows represent the direction of the vertical component of Earth’s magnetic field. The oldest rock has a magnetic dip characteristic of the southern hemisphere, but over time the dip changes, indicating that the rocks moved toward magnetic north. Source: Steven Earle (2015) CC BY 4.0 view source

When rocks form, some of the minerals that make them up can become aligned with the Earth’s magnetic field, just like a compass needle pointing to north.  This happens to the mineral magnetite (Fe₃O₄) when it crystallizes from magma.  Once the rock cools the crystals are locked in place.  This means that if the rock moves, the crystals can’t realign themselves, and they retain a remnant magnetism. This would be like jamming your compass needle so that if you turned away from north, the needle would turn with you rather than continuing to point north.

Rocks like basalt, which cool from a high temperature and commonly have relatively high levels of magnetite, are particularly susceptible to being magnetized in this way.  However, even sediments and sedimentary rocks can take on remnant magnetism as long as they have small amounts of magnetic minerals, because the magnetic grains can gradually become lined up with Earth’s magnetic field as the sediments are deposited.

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.  Remember that the vertical component of the magnetic field points more sharply downward the closer it is to the magnetic north pole.  Figure 4.9 shows the vertical component of remnant magnetism in a sequence of rocks.  Notice that the arrow starts out at 500 Ma pointing slightly upward.  This means that the rocks were in the southern hemisphere.  As the rocks get younger, the arrow tilts toward horizontal, and then points downward.  This indicates that the rocks were getting progressively closer to the north magnetic pole.

Apparent Polar Wandering Paths

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 BC, and did 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. Using an analysis similar to that in Figure 4.9, they noticed that rocks of different ages sampled from the same general area showed very different magnetic pole positions (the green line in Figure 4.10). They assumed this meant that Earth’s magnetic pole had moved around significantly over time along polar wandering paths, rather than staying close to the geographic north pole as it does today.  At the time, geophysical models suggested that the magnetic poles did not need to be aligned with the rotational poles, so this wasn’t an unreasonable conclusion, given what was known.

Figure 4.10 Apparent polar-wandering paths (APWP) for Eurasia and North America. The view is from the geographic north pole (black dot) looking down. Dots along each path show the location of magnetic north as determined from paleomagnetic data. Left- Data from Eurasia and North America agree on the location of magnetic north today (time 0), but not at any time in the past. Right- Once continent motion has been accounted for, there is agreement in data from Eurasia and North America on the location of magnetic north over the past 500 million years. Source: Steven Earle (2015) CC BY 4.0 view source

Runcorn and colleagues extended their paleomagnetic studies to North America, and began to realize that their initial conclusion had a problem.  Notice that on the right of Figure 4.10 the polar wandering path for North America (in red) does not match the path for Eurasia (in green).  For example, data from North America suggest that 200 Ma ago, magnetic north was somewhere in China, whereas data from Europe said it was in the Pacific Ocean.  There could only have been one magnetic north pole position at 200 Ma, therefore the only way to explain the discrepancy was if Europe and North America moved along different paths during this time while the pole stayed in more or less the same location.

The polar wandering paths were not actually records of the pole moving, they just looked that way, so the paths are now referred to as apparent polar wandering paths (APWP).  Subsequent paleomagnetic work showed that unique apparent polar wandering paths can be derived from rocks in South America, Africa, India, and Australia. In 1956, Runcorn became a proponent of continental drift.  There was simply no other way to explain the data.

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, this type of evidence was not sufficiently convincing to get them to change their views.

Ocean Basin Geology and Geography

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 (the hills and valleys of the ocean floor) 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 beneath them, and almost as much information about the geophysical nature of ocean rocks as of continental rocks.

Acoustic Depth Sounding

Up until 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 (Figure 4.11) in the 1920s, the number of depth readings increased by many orders of magnitude, and by the 1930s there was no doubt that major mountain chains ran through 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 4.12).

Figure 4.11 A ship-borne acoustic depth sounder. The instrument emits sound (black arcs) that reflects off the sea floor and returns to the surface (white arcs). The time interval between emitting the sound and detecting it on receivers on the ship is proportional to the water depth. Source: Steven Earle (2015) CC BY 4.0 view source

Figure 4.12 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. Source: Steven Earle (2015) CC BY 4.0 view source; Basemap after NOAA (2006) Public Domain view source

The important physical features of the ocean floor are:

• Extensive linear ridges (commonly in the central parts of the oceans) at depths of 2,000 to 3,000 m (Figure 4.12, a)
• Fracture zones perpendicular to the ridges (Figure 4.12, a)
• Deep-ocean plains at depths of 4,000 to 5,000 m (Figure 4.12, a and d)
• Relatively flat and shallow continental shelves with depths under 500 m (Figure 4.12, b)
• Deep trenches (up to 11,000 m deep), most near the continents (Figure 4.12, c)
• Seamounts and chains of seamounts (Figure 4.12, d)

Seismic Reflection Sounding

Seismic reflection sounding involves transmitting high-energy sound bursts and then measuring the echoes with a series of receivers called 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. This allowed sediment thicknesses to be mapped (Figure 4.13).

Figure 4.13 Map of global sediment thickness. Source: NOAA (2003) Public Domain view source

It was soon discovered that although the sediments were up to several 1000s of m thick near the continents, they were relatively thin — or even non-existent — along ocean ridges (Figure 4.14). 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 4.14 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. Source: Steven Earle (2015) CC BY 4.0 view source

Heat Flow Rates

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 higher than average heat-flow rates along the ridges, and lower than average rates in trenches. These data were interpreted as evidence of mantle convection, with areas of high heat flow corresponding to upward convection of hot mantle material, and areas of low heat flow corresponding to downward convection.

Earthquake Belts

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. A remarkable correspondence was observed between earthquake locations 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 4.15).

Figure 4.15 Aleutian Island subduction zone earthquakes. Left- Map view with earthquakes marked as dots. Red dots are the shallowest earthquakes and blue are the deepest. Quakes get deeper further inland from the trench. Right- Cross-section through a-b. Coloured dots show the depth of earthquakes. Colours correspond to dots in the left figure. Earthquake depth is related to the position of the Pacific plate as it travels beneath the North American plate. Source: Steven Earle (2015) CC BY 4.0 view source

Magnetic Stripes on the Sea Floor

In the 1950s, scientists from the Scripps Oceanographic Institute in California persuaded the United States 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 coasts of BC and Washington State. This survey revealed a bewildering pattern of low and high magnetic field intensity in sea-floor rocks (Figure 4.16). When the data were first plotted on a map in 1961, nobody understood them — not even the scientists who collected them. Many 1000s of km of magnetic surveys were conducted over the next several years.

Figure 4.16 Pattern of sea-floor magnetic field intensity off the west coast of British Columbia and Washington. Black regions have higher than average magnetic field instensity, and white regions have lower than average intensity. Source: Steven Earle (2015) CC BY-SA 4.0, modified after U. S. Geological Survey (n.d.) Public Domain view source (adapted from Raff and Mason, 1961).

The wealth of new data from the oceans began to significantly influence geological thinking in the 1960s. In 1960, Harold Hess from Princeton University, advanced a hypothesis 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 4.17). He also suggested that the less-dense continental crust did not descend with oceanic crust into trenches, but that colliding landmasses were thrust up to form mountains.

Hess’s hypotheses formed the basis for our ideas on sea-floor spreading and continental drift, but did not go so far as to claim that the crust is made up of separate plates. The Hess model was not roundly criticized, but also not widely accepted, partly because evidence was still lacking.

Figure 4.17 A representation of Harold Hess’s model for sea-floor spreading and subduction. Source: Steven Earle (2015) CC BY 4.0 view source

Collection of magnetic data from the oceans continued in the early 1960s, but the striped patterns remained unexplained. Some assumed that, as with continental crust, the stripes were related to compositional variations in rock, such as variations in the amount of magnetite. 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 him, noted that the magnetic patterns were symmetrical on either side of the 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 3, Earth’s magnetic field periodically weakens, then becomes 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. The present “normal” event, known as the Bruhnes magnetic chron, has persisted for about 780,000 years. It 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 3.16).

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 would produce a positive anomaly (a black stripe on the sea-floor magnetic map). Oceanic crust created during a reversed event would have polarity opposite Earth’s present field and thus 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, Morley’s papers submitted earlier in 1963 to Nature and The Journal of Geophysical Research were rejected. The idea is sometimes referred to 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 durations 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 known timeline of Earth’s magnetic field reversals.

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 Hawai’ian and Emperor Seamount island chains in the Pacific Ocean (Figure 4.18). The volcanic rock making up these islands becomes progressively younger toward the southeast, culminating in the island of Hawai’i itself, which consists of rock that is almost all younger than 1 Ma.

Figure 4.18 The ages of the Hawai’ian Islands and the Emperor Seamounts in relation to the location of the Hawai’ian mantle plume. Source: Steven Earle (2015) CC BY 4.0 view source; Base map from the National Geophysical Data Centre/USGS (2005) Public Domain view source

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 makes a pronounced change in direction, from northwest-southeast for the Hawai’ian Islands, to nearly north-south for the Emperor Seamounts. This change has been ascribed to a change in direction of the Pacific Plate moving over the stationary mantle plume. An alternative hypothesis is that rather than the Pacific Plate having undergone a sudden change in motion, the plume itself has moved at least 2,000 km south over the period between 81 and 45 Ma (Tarduno et al., 2003).

There is evidence of many such mantle plumes around the world (Figure 4.19). Most are within ocean basins, including places like Hawai’i, 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, and possibly for hundreds of millions of years in some cases.

Figure 4.19 Mantle plume locations.  Selected Mantle plumes: 1: Azores, 3: Bowie, 5: Cobb, 8: Eifel, 10: Galapagos, 12: Hawai’i, 14: Iceland, 17: Cameroon, 18: Canary, 19: Cape Verde, 35: Samoa, 38: Tahiti, 42: Tristan, 44: Yellowstone, 45: Anahim. Source: Ingo Wölbern (2007) Public Domain view source

Oceanic spreading ridges appear to be curved features on Earth’s surface, but the ridges are in fact composed of a series of straight-line segments, offset at intervals by faults perpendicular to the ridge (Figure 4.20). In a paper published in 1965, Tuzo Wilson termed these features transform faults.

Figure 4.20 Part of the Mid-Atlantic ridge near the equator. Transform faults (red lines) are in between the ridge segments (double white lines), where the yellow arrows (indicating relative plate movement) point in opposite directions. Solid white lines are fracture zones. Source: Steven Earle (2015) CC BY 4.0 view source

In the same 1965 paper, Wilson introduced the idea that the crust can be divided into a series of rigid plates, and is thus responsible for the term plate tectonics.

Exercise: Paper Transform Fault Model

J. Tuzo Wilson used a paper model similar to the one in Figure 4.21 to explain transform faults to his colleagues. To use this model, print Figure 4.21, 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. When you’re done, you should have two folds of paper extending downward as in Figure 4.22.

Figure 4.21 Transform fault model. Source: Steven Earle (2015) CC BY 4.0 view source, modified after Stewart (1990).

 

Figure 4.22 Use of the transform fault model. Source: Steven Earle (2015) CC BY 4.0 view source

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 will 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 will also see that the oceanic crust is being magnetized as it forms at the ridge. The magnetic patterns represent the last 2.5 Ma of geological time.

There are other versions of this model available at https://web.viu.ca/earle/transform-model/. For more information see Earle (2004).

References

Earle, S. (2004). A simple paper model of a transform fault at a spreading ridge. Journal of Geoscience Education 52, 391-392.

Raff, A., & Mason, R. (1961) Magnetic survey off the west coast of North America, 40˚ N to 52˚ N latitude. Geological Society of America Bulletin 72, 267-270.

Stewart, J. A. (1990). Drifting continents and colliding paradigms. Bloomington IN: Indiana University Press.

Tarduno, J. A., Duncan, R. A., Scholl, D. W., Cottrell, R. D., Steinberger, B., Thordarson, T., Kerr, B. C., Neal, C. R., Frey, F. A., Torii, M., and Carvallo, C. (2003). The Emperor Seamounts: Southward Motion of the Hawaiian Hotspot Plume in Earth’s Mantle. Science 301(5636), 1064–1069. DOI: 10.1126/science.1086442

Wilson, J. T. (1965). A new class of faults and their bearing on continental drift. Nature 207, 343-347.

The ideas of continental drift and sea-floor spreading became widely accepted by 1965, and more geologists started thinking in these terms. By the end of 1967, Earth’s surface had been mapped into a series of plates (Figure 4.23). The major plates are Eurasian, Pacific, Indian, Australian, North American, South American, African, and Antarctic plates. There are also numerous small plates (e.g., Juan de Fuca, Nazca, Scotia, Philippine, Caribbean), and many very small plates or sub-plates. The Juan de Fuca Plate is actually three separate plates (Gorda, Juan de Fuca, and Explorer), all moving in the same general direction but at slightly different rates.

Figure 4.23 A detailed map of Earth’s tectonic plates. Click on the map to enlarge. Source: Paul Lowman and Jacob Yates, NASA Goddard Space Flight Center (2002) Public Domain view source

Plate motions can be tracked using Global Positioning System (GPS) data from different locations on Earth’s surface. Rates of motions of the major plates range from less than 1 cm/y to more than 10 cm/y. The Pacific Plate is the fastest, moving at more than 10 cm/y in some areas, followed by the Australian and Nazca Plates. The North American Plate is one of the slowest, averaging ~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 rotate as they move; the North American Plate, for example, rotates counter-clockwise, while the Eurasian Plate rotates clockwise.

Boundaries between the plates are of three types: divergent (moving apart), convergent (moving together), and transform (moving side by side). The plates are made up of crust and lithospheric mantle (Figure 4.24). Even though the plates are in constant motion, and move in different directions, there is never a significant amount of space between them. Plates move along the lithosphere-asthenosphere boundary, because the asthenosphere is relatively weak. It deforms as the plates move, rather than locking them in place.

Figure 4.24 The crust and upper mantle. Tectonic plates consist of lithosphere, which includes the crust and the lithospheric (rigid) part of the mantle. Source: Steven Earle (2015) CC BY 4.0 view source

At spreading centres, the lithospheric mantle is relatively thin. The upward convective motion of hot mantle material generates temperatures that are too high for the existence of a significant thickness of rigid lithosphere at the same time that the plates are falling away from each other (Figure 4.17).

The fact that plates include both crustal material and lithospheric mantle material makes it possible for a single plate to be include both oceanic and continental crust. 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. The partial melting happens when hot mantle rock is moved from deep within Earth where pressures are too high for it to be liquid, to shallower depths where the pressure is much lower (Figure 4.25, bottom left).

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 once the melt escapes from the rock in which it formed, and ascends. Most divergent boundaries are located in the oceans, and the crustal material created at a spreading boundary is always oceanic in character; in other words, it is mafic igneous rock (basalt or gabbro, with minerals rich in iron and magnesium). Spreading rates vary considerably, from 1 cm/y to 3 cm/y in the Atlantic, to between 6 cm/y and 10 cm/y in the Pacific. Some of the processes taking place in this setting include (Figure 4.25, top):

• Melted rock (magma) from the mantle rising up to fill the voids left by divergence of the two plates
• Pillow lavas forming where melted rock emerges on the ocean floor and is cooled by seawater (Figure 4.25, bottom right)
• Vertical sheeted dykes intruding into cracks resulting from the spreading
• Magma cooling more slowly in the lower part of the new crust, forming bodies of gabbro

Figure 4.25 Divergent boundary. Lower left- General processes taking place along divergent boundaries. Top- Expanded view of the white box showing divergent boundary processes and materials. Bottom right- Pillow basalts from the ocean floor of Hawai’i. Source: Lower left- Steven Earle (2015) CC BY 4.0 view source; Top- Steven Earle (2015) CC BY 4.0 view source modified after Sinton and Detrick (1992); Lower right- NOAA (1988) Public Domain view source

Spreading is thought to start with lithosphere being warped upward into a dome by buoyant material from an underlying mantle plume or series of mantle plumes. The buoyancy of the mantle plume causes the dome to fracture in a radial pattern, with three arms spaced at approximately 120° (Figure 4.26).

Figure 4.26 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) Source: Steven Earle (2015) CC BY 4.0 view source

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. This type of valley may eventually develop 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 the Atlantic Ocean mantle plume locations in Figure 4.19).

Convergent Boundaries

Convergent boundaries, where two plates are moving toward each other, are of three types, depending on whether ocean or continental crust is present on either side of the boundary. The types are ocean-ocean, ocean-continent, and continent-continent.

Ocean-Ocean Convergent Boundaries

At an ocean-ocean convergent boundary, a plate margin consisting of oceanic crust and lithospheric mantle is subducted, or travels beneath, the margin of the plate with which it is colliding (Figure 4.27). Often it is the older and colder plate that is denser and subducts beneath the younger and hotter plate. Ocean trenches commonly form along these boundaries.

Figure 4.27 Configuration and processes of an ocean-ocean convergent boundary Source: Steven Earle (2015) CC BY 4.0 view source

As the subducting crust is heated and the pressure increases, water is released from within the subducting material. This water comes primarily from alteration of the minerals pyroxene and olivine to serpentine near the spreading ridge shortly after the rock’s formation. The water mixes with the overlying mantle, which lowers the melting point of mantle rocks, causing magma to form. This process is called flux melting or fluid-induced melting.

The newly produced magma, which is lighter than the surrounding mantle rocks, rises through the mantle and sometimes through 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. The largest earthquakes occur near the surface where the subducting plate is still cold and strong.

Examples of ocean-ocean convergent zones are subduction of the Pacific Plate south of Alaska (Aleutian Islands) and west of the Philippines, subduction of the Indian Plate south of Indonesia, and subduction of the Atlantic Plate beneath the Caribbean Plate.

Ocean-Continent Convergent Boundaries

At an ocean-continent convergent boundary, the oceanic plate is subducted beneath the continental plate in the same manner as at an ocean-ocean boundary. Rocks and sediment on the continental slope are thrust up into an accretionary wedge, and compression leads to faults forming within the continental plate (Figure 4.28). 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 4.28 Configuration and processes of an ocean-continent convergent boundary Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2015) CC BY 4.0 view source

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).

Continent-Continent Convergent Boundary

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 4.29). The colliding continental material will not be subducted because it is not dense enough, 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, as well as from any sediments that had accumulated along the shores of both continental masses, and commonly also from some ocean crust and upper mantle material.

Figure 4.29 Configuration and processes of a continent-continent convergent boundary Source: Steven Earle (2015) CC BY 4.0 view source

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.

When a subduction zone is jammed shut by a continent-continent collision, plate tectonic stresses that are still present can sometimes cause a new subduction zone to develop outboard of the colliding plate.

Transform Boundaries

Transform boundaries exist where one plate slides past another without producing or destroying crust, except in the special case where the transform boundary has bends and jogs. There will be collisions and divergence on a small scale as the jogs crash into the bends, or open up small windows to deeper crust.

Most transform faults connect segments of mid-ocean ridges and are thus ocean-ocean plate boundaries (Figure 4.20). Some transform faults connect continental parts of plates. An example is the San Andreas Fault, which connects the southern end of the Juan de Fuca Ridge with the northern end of the East Pacific Rise (a ridge) in the Gulf of California (Figures 4.30 and 4.31). The part of California west of the San Andreas Fault and all of Baja California are on the Pacific Plate. But transform faults do not just connect divergent boundaries; 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.

Figure 4.30 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. Source: Steven Earle (2015) CC BY 4.0 view source

Figure 4.31 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 slide on its foundation. Source: Steven Earle (2015) CC BY 4.0 view source

Exercise: A Different Type of Transform Fault

Figure 4.32 shows the Juan de Fuca (JDF) and Explorer plates off the coast of Vancouver Island. The JDF Plate is moving toward the North American Plate at ~ 4.5 cm/y. 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 (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.

Figure 4.32 Juan de Fuca and Explorer plates separated by the Nootka Fault (marked with red stars). Source: Steven Earle (2015) CC BY 4.0 view source

Plate Tectonics and Supercontinent Cycles

The present continents were once all part of a supercontinent that Alfred Wegener named Pangea (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 (Figure 4.33). 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, and most of Europe had separated from North America. By 50 Ma, Australia had separated from Antarctica, and shortly after that, India collided with Asia.

Figure 4.33 Sequence of paleogeographic reconstructions showing the breakup of Pangea. Source: Karla Panchuk (2017) CC BY-NC-SA 4.0. Maps from C. R. Scotese, PALEOMAP  Project (www.scotese.com). Click the image for map sources and terms of use.

Within the past few million years, rifting has occurred 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.

Pangea was not the first supercontinent. It was preceded by Pannotia (600 to 540 Ma), Rodinia (1,100 to 750 Ma), and by others before that. In fact, in 1966, Tuzo Wilson proposed that supercontinents are part of an on-going cycle, which we now refer to as a Wilson cycle. In a Wilson cycle, continents break up, and fragments drift apart only to collide again and make a new continent.

At present we are in the stages of a Wilson cycle where fragments are drifting and changing their configuration. 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. Because the oceanic crust formed by spreading along 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.

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, it will be pushed farther and farther into the mantle, and eventually the oceanic lithosphere may break away from the continental lithosphere and begin to subduct (Figure 4.34).

Figure 4.34 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) may no longer be pushed to the west and could start to move east because the rate of spreading in the Pacific basin is faster than along the Mid-Atlantic Ridge. Source: Steven Earle (2015) CC BY 4.0 view source

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 again with Europe and Africa. If this continues without changing for another few hundred million years, we will be back to where we started, with one supercontinent (Figure 4.35).

Figure 4.35 Sequence of reconstructions showing the possible future configuration of land masses on Earth at 50, 150, and 250 million years from now. Movements culminate in the formation of a new supercontinent called Pangea Ultima. Source: Karla Panchuk (2017) CC BY-NC-SA 4.0. Maps from C. R. Scotese, PALEOMAP  Project (www.scotese.com). Click the image for map sources and terms of use.

There is strong evidence around the margins of the Atlantic Ocean that this process has taken place before. There are roots of ancient mountain belts along the eastern margin of North America, the western margin of Europe, and the north-western margin of Africa, which show that these landmasses once collided with each other to form a mountain chain. The mountain chain might have been 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 north-eastern 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.

It is probably no coincidence that the Atlantic Ocean rift may have occurred in approximately the same place during two separate events several hundred million years apart. 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 4.36).

Figure 4.36 A scenario for the Wilson cycle. The cycle starts with continental rifting above a series of mantle plumes (red dots, A). The continents separate (B), and then re-converge some time later, forming a fold-belt mountain chain. Eventually rifting is repeated, possibly because of the same set of mantle plumes (D), but this time the rift is in a different place. Source: Steven Earle (2015) CC BY 4.0 view source

References

Sinton, J. M., and Detrick, R. S. (1992). Mid-Ocean Ridge Magma Chambers. Journal of Geophysical Research 97(B1), 197-216.

Mantle convection is often said to be critical to plate tectonics. While this is almost certainly so, there is still debate about the actual forces that make the plates move. One side of the argument holds that the plates are only moved by the traction caused by mantle convection, and that friction between the asthenosphere and lithosphere pulls the lithosphere along as the mantle convects. The other side holds that traction plays only a minor role and that ridge-push and slab-pull are more important (Figure 4.37).

Ridge-push refers to gravity causing lithosphere to slide downhill away from the elevated mid-ocean ridges. Slab-pull refers to the weight of subducting slabs dragging the rest of the plate down into the mantle.

Figure 4.37 Models for plate motion mechanisms. Source: Steven Earle (2015) CC BY 4.0 view source

Kearey and Vine (1996) have listed some compelling arguments in favour of the ridge-push/slab-pull model:

• 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.

• 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. Such high rates of convection are not supported by geophysical models.

• Although large plates have the potential for much higher convection traction, plate velocity is not related to plate area.

References

Kearey, P., & Vine, F. (1996). Global Tectonics (2nd E.). Oxford: Blackwell Science Ltd.

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

4.1 Alfred Wegener’s Arguments for 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.

4.2 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.

4.3 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.

4.4 Plates, 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 to greater than 10 cm/y. The three types of plate boundaries are divergent (plates moving apart and new crust forming), convergent (plates moving together and one possibly 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 can 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.

4.5 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.

Review Questions

1. List some of the evidence used by Wegener to support his idea of moving continents.
2. What was the primary weakness in Wegener’s continental drift hypothesis?
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. How did we learn about the topography of the sea floor in the early part of the 20th century?
6. What evidence from paleomagnetic studies provided support for continental drift?
7. Which parts of the oceans are the deepest?
8. Why is there less sediment along ocean ridges than on other parts of the sea floor?
9. How were the oceanic heat-flow data related to mantle convection?
10. Describe the spatial distribution of earthquakes at ocean ridges and ocean trenches.
11. In the model for ocean basins developed by Harold Hess, what happens at oceanic ridges and what happens at oceanic trenches?
12. What aspect of plate tectonics was not included in the Hess model?
13. What is a mantle plume and what is its expected lifespan?
14. Describe the nature of movement at an ocean ridge transform fault (a) between the ridge segments, and (b) outside of ridge segments.
15. How is it possible for a plate to include both oceanic and continental crust?
16. What is the likely relationship between mantle plumes and the development of a continental rift?
17. Why does subduction not take place at a continent-continent convergence zone?
18. Where are Earth’s most recent sites of continental rifting and creation of new ocean floor?
19. 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?

1. The evidence used by Wegener to support his idea of moving continents included matching continental shapes and geological features on either side of the Atlantic; common terrestrial fossils in South America, Africa, Australia, and India; and data on the rate of separation between Greenland and Europe.
2. The primary weakness of Wegener’s theory was that he had no realistic mechanism for making continents move.
3. Contractionists believed that mountains formed because the crust wrinkled into mountains as Earth cooled and contracted. Permanentists believed that mountains formed by the geosynclinal process.
4. In the late 19th century the trans-Atlantic paleontological matchups were explained by assuming that there must have been land bridges between the continents at some time in the past.
5. Prior to 1920, ocean depths were measured by dropping a weighted line over the side of ship. Echo sounding techniques were developed at around that time and greatly facilitated the measurement of ocean depths.
6. Paleomagnetic studies showed that old rocks on the continents indicated different locations for magnetic north than the position of magnetic north today. They also showed that the difference in pole position from data on different continents increased progressively for older and older rocks. This implied that either Earth had more than one magnetic pole moving around, or that the continents had moved.
7. Trenches associated with subduction zones are the deepest parts of the oceans.
8. The ocean ridge areas are the youngest parts of the sea floor and thus there hasn’t been time for much sediment to accumulate.
9. It was (and still is) thought that high heat flow exists where mantle convection cells are moving hot rock from the lower mantle toward the surface, and that low heat flow exists where there is downward movement of mantle rock.
10. Earthquakes are consistently shallow and relatively small at ocean ridges. At ocean trenches earthquakes become increasingly deep in the direction that the subducting plate is moving.
11. In the Hess model new crust was formed at ocean ridges. Crust was recycled back into the mantle at the trenches.
12. Hess’s model did not include the concept of tectonic plates.
13. A mantle plume is a column of hot rock (not magma) that ascends toward the surface from the lower mantle. Mantle plumes last tens of millions of years to hundreds of millions of years.
14. (a) Between the ridge segments there is movement in opposite directions along a transform fault. (b) Outside of the ridge segments the two plates are moving in the same direction and likely at about the same rate. These regions are known as fracture zones.
15. Tectonic plates are made up of crust and the lithospheric (rigid) part of the underlying mantle. The mantle part ensures that the very different oceanic and continental crust sections of a plate can act as one unit.
16. A mantle plume beneath a continent can cause the crust to form a dome that might eventually split open. Several mantle plumes along a line within a continent could lead to rifting.
17. Subduction does not take place at a continent-continent convergent zone because neither plate is dense enough to sink into the mantle.
18. Continental rifting is taking place along the East Africa Rift, and sea floor has recently been created in the Red Sea and also in the Gulf of California.
19. The accumulation of sediment at a passive ocean-continent boundary will lead to the depression of the lithosphere and could eventually result in the separation of the oceanic and continental parts of the plate and the beginning of subduction.

Protons Are What Make Elements Distinct

All matter, including mineral crystals, is made up of atoms.  All atoms are made up of three main particles: protons, neutrons, and electrons. Protons have a positive charge, neutrons have no charge, and electrons have a negative charge. Protons and neutrons have approximately the same mass, but electrons have a mass that is 10,000 times smaller.

The element hydrogen (H) has the simplest atoms. Most hydrogen atoms have just one proton and one electron. The proton forms the nucleus (the centre of the atom), while the electron orbits around it (Figure 5.5, left). All other elements have more than one proton in their nucleus. Protons repel each other because they are positively charged, but it is possible to have more than one proton in a nucleus because neutrons hold them together. The next most complex atom, helium (He) has two protons and two neutrons in its nucleus, in its most common form. Some atoms of the same element can have different numbers of neutrons. For example, forms of hydrogen exist with one and two neutrons, and a tiny fraction of He atoms have only one neutron. Forms of an element with different numbers of neutrons are called isotopes.

Figure 5.5 Atomic structure of hydrogen and helium showing protons (p+), neutrons (n), and electrons (e-). Source: Bruce Blaus (2014) CC BY 3.0 view source

The number of protons in an atom determines what element it will be, so the number of protons is called the atomic number of that element. The total number of protons and neutrons in the nucleus is the mass number. The mass number distinguishes between isotopes of an element. Isotopes of an element are denoted by putting the mass number as a subscript in front of the symbol for that element. For example, the isotopes of hydrogen are ¹H (1 proton), ²H (1 proton + 1 neutron), and ³H (1 proton + 2 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. This is because the more protons that are concentrated in a small space, the more neutrons are needed to keep the nucleus together.  The most common isotope of uranium (U), for example, is ²³⁸U. It has 92 protons, but requires 146 neutrons to keep them together.  The neutrons are only partly successful.  Uranium is radioactive, meaning that its nucleus will eventually split apart and release energy. What remains of the nucleus has fewer protons, so after decay the atom is a different element.

Electrons Are What Control How Atoms Interact

Electrons orbiting around the nucleus of an atom are arranged in shells (also called energy levels). The first shell can hold only two electrons (as in H and He in Figure 5.5), but the next shell holds up to eight electrons. An atom can have many shells of electrons, but there are never more than 8 outermost electrons interacting with surrounding atoms.

The outermost electrons determine how atoms can be bonded together. Elements that have a full outer shell (e.g., neon, Figure 5.6 right)  are inert because they do not react with other elements to form compounds. These are the noble gases (including helium, argon, krypton, and radon, in addition to neon) in the far-right column of the periodic table. For elements that do not have a full outer shell (e.g., lithium, Figure 5.6 left), the outermost electrons can interact with the outermost electrons of nearby atoms to create chemical bonds.

Figure 5.6 The number of electrons in an atom’s outermost shell (or energy level) determine whether it will bond to other atoms, and how it will bond. Right- Neon has a completely filled outer shell with 8 electrons. It does not bond with other atoms. Left- Lithium has only one electron in its outer shell. It bonds with other atoms. Source: Bruce Blaus (2014) CC BY 3.0 view source

The electron shell configurations for 29 of the first 36 elements are listed in Table 5.1. Note that some of the shells in the table below have more than 8 electrons.  This is because they contain subshells.  For example, the third shell can hold up to 18 electrons because it contains one subshell that can hold 2 electrons, and two subshells that can hold 8 electrons each.

Atoms seek to have a full outer shell. For hydrogen and helium, a full outer shell means two electons. For other elements, it means 8 electrons. Filling the outer shell is accomplished by transferring or sharing electrons with other atoms in chemical bonds.  The type of chemical bond is important for the study of minerals because the type of bond will determine many of a mineral’s physical and chemical properties.

Ionic Bonds

Consider the example of halite again, which is made up of sodium (Na) and chlorine (Cl).  Na has 11 electrons: two in the first shell, eight in the second, and one in the third (Figure 5.7, top). Na readily gives up the third shell electron so it can have the second shell with 8 electrons as its outermost shell.  When it loses the electron, the total charge from the electrons is -10, but the total charge from the protons is +11, so it is left with a +1 charge over all.

Figure 5.7 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). Source: Steven Earle (2015) CC BY 4.0 view source

Chlorine has 17 electrons: two in the first shell, eight in the second, and seven in the third. Cl readily accepts an eighth electron to fill its third shell, and therefore becomes negatively charged because it has a total charge of -18 from electrons, and a total charge of +17 from protons.

In changing their number of electrons, these atoms become ions — the sodium loses an electron to become a positive ion or cation,You can remember that a cation is positive by remembering that a cat has paws (paws sounds like "pos" in "positive"). You could also think of the "t" in "cation" as a plus sign. and the chlorine gains an electron to become a negative ion or anion (Figure 5.7, bottom). Because negative and positive charges attract, sodium and chlorine ions stick together, creating an ionic bond. In an ionic bond, electrons can be thought of as having transferred from one atom to another.

Exercise: Cation or Anion?

A number of elements are listed below along with their atomic numbers (the number of protons, and therefore also the number of electrons in the atom). Assuming that the first electron shell can hold two electrons and subsequent electron shells can hold eight electrons, sketch the electron configurations for these elements, as in the example for fluorine (Figure 5.8). If you fill a shell and have electrons left over, draw another shell around the atom. Predict whether the element is likely to form a cation or an anion, and what charge it would have (e.g., +1, +2, –1).

Figure 5.8 How to draw the electron configuration for fluorine, with an atomic number of 9. Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2015) CC BY 4.0 view source

1. Lithium (3)	5. Beryllium (4)
2. Magnesium (12)	6. Oxygen (8)
3. Argon (18)	7. Sodium (11)
4. Chlorine (17)

 

Covalent Bonds

An element like chlorine can also form bonds without forming ions. For example, two chlorine atoms can each complete their outer shells by sharing electrons.  Chlorine gas (Cl₂, Figure 5.9) is formed when two chlorine atoms form a covalent bond.

Figure 5.9 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. Source: Steven Earle (2015) CC BY 4.0 view source

Carbon is another atom that participates in covalent bonding.  An uncharged carbon atom has six protons and six electrons. Two of the electrons are in the inner shell and four are in the outer shell (Figure 5.10, left). Carbon would need to gain or lose four electrons to have a filled outer shell, and this would create too great a charge imbalance. Instead, carbon atoms share electrons to create covalent bonds (Figure 5.10, right).

Figure 5.10 The electron configuration of carbon (left) and the sharing of electrons in covalent C bonding (right). The electrons shown in blue are shared between adjacent C atoms. Source: Steven Earle (2015) CC BY 4.0 view source

In the mineral diamond (Figure 5.11, left), the carbon atoms are linked together in a three-dimensional framework, where one carbon atom is bonded to four other carbon atoms, and every bond is a very strong covalent bond.

Figure 5.11 Covalently bonded structures. Left: Diamond with three-dimensional structure of covalently bonded carbon. Right: Graphite with covalently bonded sheets of carbon. Sheets are held together by weaker van der Waals forces. Source: Karla Panchuk (2018) CC BY 4.0, modified after Materialscientist (2009) CC BY-SA 3.0 view source

Other Types of Bonds

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. Consider the mineral graphite (Figure 5.11, right): the carbon atoms are linked together in sheets or layers in which each carbon atom is covalently bonded to three others. Graphite-based compounds are strong because of the covalent bonding between carbon atoms within each layer, which is why they are used in high-end sports equipment such as ultralight racing bicycles. Graphite itself is soft, however, because the layers themselves are held together by relatively weak Van der Waals forces. 

Van der Waals forces, like hydrogen bonds, work because molecules can be electrostatically neutral, but still have an end that is slightly more positive and an end that is slightly more negative. In water molecules (Figure 5.12, left), the bent shape puts the hydrogen atoms on one side of the molecule, and the oxygen atom, with more electrons, on the other. The charge is distributed asymmetrically across the water molecule. Contrast this with the straight carbon dioxide (Figure 5.12, right) molecule. The slightly more negative oxygen atoms on the ends are distributed symmetrically on either side of the carbon atom.  

Figure 5.12 Hydrogen bonding. Water molecules (left) are polar molecules (their charge is distributed asymmetrically). Slightly negative parts of the molecule are attracted to slightly positive parts of other water molecules. Carbon dioxide (right) is a non-polar molecule. The slightly negative oxygen atoms are distributed symmetrically on either side of the carbon atom. Source: Karla Panchuk (2018) CC BY-SA 4.0. Modified after Querter (2011) CC BY-SA 3.0 view source and Jynto (2011) CC0 1.0 view source

Metallic bonding occurs in metallic elements because they 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, the dissociated electrons can move freely from one atom to another. This feature accounts for two very important properties of metals: their electrical conductivity and their malleability (they can be deformed and shaped).

Figure 5.13 Metallic bonding. Dissociated electrons (grey dots) move between metal atoms. Source: Karla Panchuk (2018) CC BY-SA 4.0. Nucleus by Fornax (2010) CC BY-SA 3.0 view source

Minerals are organized according to the anion or anion group (a group of atoms with a net negative charge, e.g., SO₄^2–) they contain, because the anion or anion group has the biggest effect on the properties of the mineral.  Silicates, with the anion group SiO₄^4-, are by far the most abundant group in the crust and mantle. (They will be discussed in Section 5.4). The different mineral groups along with some examples of minerals in each group are summarized below.

Oxide Minerals: O^2- Anion

Oxide minerals (Figure 5.14) have oxygen (O^2–) as their anion.  They don’t include anion groups with other elements, such as  the carbonate (CO₃^2–), sulphate (SO₄^2–), and silicate (SiO₄^4–) anion groups. The iron oxides hematite and magnetite are two examples that are important ores of iron. Corundum is 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 aluminum, respectively.

Figure 5.14 Oxide minerals include metal ore minerals, industrial minerals, and gemstones. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Photos by R. Weller/ Cochise College. Click the image for photo sources.

Sulphide Minerals: S^2- Anion

Sulphide minerals (Figure 5.15) include galena, sphalerite, chalcopyrite, and molybdenite, which are the most important ores of lead, zinc, copper, and molybdenum, respectively. Some other sulphide minerals are pyrite, bornite, stibnite, and arsenopyrite. Sulphide minerals tend to have a metallic sheen.

Figure 5.15 Sulphide minerals often have a metallic lustre and include metal ores. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Photos by R. Weller/ Cochise College. Click the image for photo sources.

Sulphate Minerals: SO₄^2- Anion Group

Many sulphate minerals form when sulphate-bearing water evaporates. A deposit of sulphate minerals may indicate that a lake or sea has dried up at that location.  Sulphates with calcium include anhydrite, and gypsum (Figure 5.16). Sulphates with barium and strontium are barite and celestite, respectively. In all of these minerals, the cation has a +2 charge, which balances the –2 charge on the sulphate ion.

Figure 5.16 Sulphate minerals. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Click the image for photo sources.

Halide Minerals: Anions from the Halogen Group

The anions in halides are the halogen elements including chlorine, fluorine, and bromine. Examples of halide minerals are cryolite, fluorite, and halite (Figure 5.17).  Halide minerals are made of ionic bonds. Like the sulphates, some halides also form when mineral-rich water evaporates.

Figure 5.17 Halide minerals. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Click the image for photo sources.

Carbonate Minerals: CO₃^2- Anion Group

The carbonate anion group combines with +2 cations to form minerals such as calcite, magnesite, dolomite, and siderite (Figure 5.18). The copper minerals malachite and azurite are also carbonates.  The carbonate mineral calcite is the main component of rocks formed in ancient seas by organisms such as corals and algae.

Figure 5.18 Carbonate minerals. Source: Karla Panchuk (2018) CC BY-SA 4.0. Photos by Rob Lavinsky, iRocks.com, CC BY-SA 3.0. Click the image for photo sources.

Phosphate Minerals: PO₄^3- Anion

The apatite group of phosphate minerals (Figure 5.19, left) includes hydroxyapatite, which makes up the enamel of your teeth. Turquoise is also a phosphate mineral (Figure 5.19, right).

Figure 5.19 Phosphate minerals. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Photos by R. Weller/ Cochise College. Click the image for photo sources.

Silicates (SiO₄^4–)

The silicate minerals include the elements silicon and oxygen in varying proportions . These are discussed at length in Section 5.4.

Native Element Minerals

These are minerals made of a single element, such as gold, copper, silver, or sulphur (Figure 5.20).

Figure 5.20 Native element minerals are made up of a single element. Source: Karla Panchuk (2018) CC BY-SA 4.0. Click the image for photo sources.

Silicon and oxygen bond covalently to create a silicate tetrahedron (SiO₄^4-), which is a four-sided pyramid shape with oxygen at each corner and silicon in the middle (Figure 5.21). This structure is the building block of many important minerals in the crust and mantle. Silicon has a charge of +4, and oxygen has a charge of -2, so the total charge of the silicate anion is -4.

Figure 5.21 The silica tetrahedron is the building block of all silicate minerals. Source: Karla Panchuk (2018) CC BY-SA 4.0. Modified after Helgi (2013) CC BY-SA 3.0 view source

In silicate minerals, these tetrahedra are arranged and linked together in a variety of ways, from single units to chains, rings, and more complex frameworks.  In the rest of this section we will discuss the structures of the most common silicate minerals in Earth’s crust and mantle.

Exercise: Make a Tetrahedron

Download this PDF file with the tetrahedron pattern below. 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’re feeling ambitious, make several tetrahedra and and use toothpicks through the corners to make the configurations discussed below.

Figure 5.22 Pattern for a tetrahedron. Source: Steven Earle (2015) CC BY 4.0 view source

Isolated Tetrahedra

The simplest silicate structure, that of the mineral olivine (Figure 5.23), is composed of isolated tetrahedra bonded to iron and/or magnesium ions (Figure 5.23 left). In olivine, the –4 charge of each silica tetrahedron is balanced by two iron or magnesium cations, each with a charge of +2. Olivine can be pure Mg₂SiO₄ or Fe₂SiO₄, or a combination of the two, written as (Mg,Fe)₂SiO₄. Magnesium and iron can substitute for each other because they both have a charge of +2, and they are similar in size. Magnesium cations have a radius of 0.73 Å, and iron cations have a radius of 0.62 Å Å stands for Ångstrom, a unit commonly used to express atomic-scale dimensions. One angstrom is 10^–10 m or 0.0000000001 m..

Figure 5.23 Olivine is a silicate mineral made of isolated silica tetrahedra bonded to Fe and Mg ions (left). Olivine crystals (centre) can often be found in the volcanic igneous rock called basalt (right). Source: Karla Panchuk (2018) CC BY-SA 4.0. Left- modified after Steven Earle (2015) CC BY 4.0 view source. Click the image for photo sources.

Although the iron and magnesium ions are similar in size, allowing them to substitute for each other in some silicate minerals, the common ions in silicate minerals have a wide range of sizes (Figure 5.24).  Ionic radii are critical to the composition of silicate minerals, because the structure of the silicate mineral will determine the size of spaces available.

Figure 5.24 The ionic radii in angstroms of some of the common ions in silicate minerals. Radii shown to scale. Notice that iron appears twice with two different radii. This is because iron can exist as a +2 ion (if it loses two electrons when it becomes an ion) or a +3 ion (if it loses three). Fe²⁺ is known as ferrous iron. Fe³⁺ is known as ferric iron. Source: Karla Panchuk (2017) CC BY 4.0. Modified after Steven Earle (2015) CC BY 4.0 view source

Chain Silicates

Pyroxene (Figure 5.25 bottom left) is an example of a single-chain silicate.  The structure of chain silicates is shown in Figure 5.25 (top). In pyroxene, silica tetrahedra form a chain because one oxygen from each tetrahedron is shared with the adjacent tetrahedron. This means there are fewer oxygens in the structure. This can be expressed as an oxygen-to-silicon ratio (O:Si). The O:Si is lower than in olivine (3:1 instead of 4:1), and the net charge per silicon atom is less (–2 instead of –4), because fewer cations are necessary to balance that charge.

Figure 5.25 Chain silicate minerals. Top: Arrangement of silica tetrahedra in single and double chains. Bottom left: Pyroxene crystals (dark crystals) of the variety aegirine (acmite). Bottom right: Amphibole crystal (dark) of the variety hornblende. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Top left- modified after Steven Earle (2015) CC BY 4.0. Top right- modified after Klein & Hurlbut (1993). Photos by R. Weller/ Cochise College. Click the image for sources.

Pyroxene compositions have the silica tetrahedra represented as SiO₃ (e.g., MgSiO₃, FeSiO₃, and CaSiO₃The variation in composition 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).  The structure of pyroxene is more “permissive” than that of olivine, meaning cations with a wider range of ionic radii can fit into it. That’s why pyroxenes can have calcium cations (radius 1.00 Å) substitute for iron (0.63 Å) and magnesium (0.72 Å) .

Exercise: Oxygen to Silicon Ratio

Figure 5.26 shows single chain and double chain structures. Count the number of tetrahedra versus the number of oxygen ions (yellow spheres) for each. Each tetrahedron has one silicon atom.

1. Confirm for yourself that the ratio of silicon to oxygen in the single chain is 1:3.
2. What is the O:Si for the double chain?

Figure 5.26 Single and double chains of tetrahedra. Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2015) CC BY 4.0 single chain/ double chain

Sheet Silicates

In mica structures the silica tetrahedra are arranged in continuous sheets (Figure 5.27), where each tetrahedron shares three oxygen anions with adjacent tetrahedra. Because even more oxygens are shared between adjacent tetrahedra, fewer charge-balancing cations are needed for sheet silicate minerals. Bonding between sheets is relatively weak, and this accounts for the tendency of mica minerals to split apart in sheets (Figure 5.27 bottom right). Two common micas in silicate rocks are biotite (Figure 5.27 bottom left), which contains iron and/or magnesium, making it a dark mineral; and muscovite (Figure 5.27 right), which contains aluminum and potassium, and is light in colour. All of the sheet silicate minerals have water in their structure, in the form of the hydroxyl (OH-) anion.

Figure 5.27 Micas are sheet silicates and split easily into thin layers along planes parallel to the sheets. Biotite mica (lower left) is has Fe and Mg cations. Muscovite mica (lower right) has Al and K instead. The muscovite mica shows how thin layers can split away in a sheet silicate. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Top left- modified after Steven Earle (2015) CC BY 4.0. Top right- modified after Klein & Hurlbut (1993). Photos by R. Weller/ Cochise College. Click the image for sources.

Some sheet silicates typically occur in clay-sized fragments (i.e., less than 0.004 mm). These include the clay minerals kaolinite, illite, and smectite, which are important components of rocks and especially of soils.

Framework Silicates

In framework silicates, tetrahedra are connected to each other in three-dimensional structures rather than in two-dimensional chains and sheets.

Feldspar

Feldspars are a group of very abundant framework silicates in Earth’s crust. They include alumina tetrahedra as well as silicate tetrahedra. In alumina tetrahedra, there is an aluminum cation at the centre instead of a silicon cation.

Feldspars are classified using a ternary (3-fold) system with three end-members (“pure” feldspars). This system is illustrated with a triangular diagram that has each end-member at one corner (Figure 5.28). The distance along a side of the diagram represents the relative abundance of the composition of each end-member.

Figure 5.28 Ternary diagram showing the feldspar group of framework silicate minerals. Alkali feldspars are those with compositions ranging between albite (with a Na cation) and orthoclase and its polymorphs (with a K cation. Plagioclase feldspars are those with compositions ranging between albite and anorthite (with a Ca cation). Source: Karla Panchuk (2018) CC BY-SA 4.0. Ternary diagram modified after Klein & Hurlbut (1993). Click the image for photo sources and a ternary diagram without mineral images.

One end-member is potassium feldspar (also referred to as K-feldspar), which has the composition KAlSi₃O₈. Depending on the temperature and rate of cooling, K-feldspar can occur as one of three polymorphs: orthoclase, sanidine, or microcline.  Another end member is albite, which has sodium instead of potassium (formula NaAlSi₃O₈). As is the case for iron and magnesium in olivine, there is a continuous range of compositions (referred to as a solid-solution series) between albite and orthoclase. Feldspars in this series are referred to as alkali feldspars. Potassium cations are much larger than sodium cations (1.37 Å versus 0.99 Å, respectively), so high temperatures are required to form alkali feldspars with intermediate compositions.

The third end-member is anorthite and it has calcium instead of potassium or sodium (formula CaAl₂Si₃O₈). Feldspars in the solid-solution series between albite and anorthite are called plagioclase feldspars. Calcium and sodium cations are nearly the same size (1.00 Å and 0.99 Å, respectively), so from that perspective it makes sense that they substitute readily for each other, and that any intermediate compositions between CaAl₂Si₃O₈ and NaAlSi₃O₈ can exist. However, calcium and sodium ions don’t have the same charge (Ca²⁺ versus Na+), making it surprising that they substitute so easily. The difference in charge is accommodated by substituting some Al³⁺ for Si⁴⁺.  Albite has one Al and three Si in its formula, while anorthite is has two Al and two Si.  Plagioclase feldspars of intermediate composition also have intermediate proportions of Al and Si.

Quartz

Quartz (SiO₂; Figure 5.29) contains only silica tetrahedra. In quartz, each silica tetrahedron is bonded to four other tetrahedra (with an oxygen shared at every corner of each tetrahedron), making a three-dimensional framework.  As a result, the ratio of silicon to oxygen is 1:2. Because 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 to add cations to balance the charge. The hardness of quartz and the fact that it breaks irregularly (notice the bottom of the crystal in Figure 5.29 right) and not along smooth planes result from the strong covalent/ionic bonds characteristic of the silica tetrahedron.

Figure 5.29 Quartz is another silicate mineral with a three-dimensional framework of silica tetrahedra. Sometimes quartz occurs as well-developed crystals (left), but it also occurs in common rocks such as granite (right). In addition to quartz, the granite contains potassium feldspar, albite, and amphibole. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Photos by R. Weller/ Cochise College. Click the image for photo sources.

References

Klein, C. & Hurlbut, C. S., Jr. (1993). Manual of Mineralogy (after J. D. Dana). New York, NY: John Wiley & Sons, Inc.

The following criteria are required for mineral crystals to grow:

• The elements needed to make the mineral crystals must be present in sufficient abundance and appropriate proportions.
• The physical and chemical conditions must be favourable.
• There must be sufficient time for the atoms to become arranged into a lattice.

Physical and chemical conditions include factors such as temperature, pressure, amount of oxygen available, pH, and the presence of water. The presence of water makes it easier for ions to move to where there are needed, and can lead to the formation of larger crystals over shorter time periods, as with the gypsum crystals at the beginning of this chapter. Time is one of the most important factors because it takes time for atoms to line themselves up into an orderly structure. If time is limited, the mineral grains may remain very small.

Most of the minerals that make up the rocks in the crust and mantle 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. When cooling is rapid and many crystals form at once, only small mineral grains will form before the rock becomes solid. The resulting rock will be fine-grained (i.e., crystals less than 1 mm). When cooling is slow, or when few crystals are growing at a time, relatively large crystals will develop.

Minerals can also form in several other ways:

• Precipitation from a solution (e.g., from hot water flowing underground, or when evaporation concentrates ions in a lake or inland sea)
• Precipitation from a gas (e.g., from vents releasing volcanic gases)
• Metamorphism:  solid minerals react with each other under high pressures and temperatures, and new minerals are formed.
• Weathering: minerals unstable at Earth’s surface are chemically altered by surface processes.
• Organic formation: organisms build shells (primarily of calcite or aragonite), and teeth and bones (primarily of apatite).

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, because they can use the same properties to identify minerals anywhere. That doesn’t mean that it’s easy, however. Identification of minerals takes practice. Some of the mineral properties that are useful for identification are colour, streak, lustre, hardness, habit, cleavage or fracture, and density.

Colour

Some minerals have distinctive colours that useful as diagnostic criteria. The mineral sulphur (Figure 5.30 left) is always a characteristic bright yellow. For other minerals, colour might vary. Hematite is an example of a mineral for which colour is not necessarily diagnostic. In some forms hematite is a deep dull red (a fairly unique colour), but in others it is a metallic silvery black (5.30, right).

Figure 5.30 Colour is a useful diagnostic property for sulphur (left) and for some types of hematite (right) because the yellow and dark red colours are unique to those minerals. In contrast, silvery metallic forms of hematite are similar in appearance to many other minerals. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Photos by R. Weller/ Cochise College. Click the image for photo sources.

For other minerals, the problem is that a single mineral can have a wide range of colours. The colour variations can be the result of varying proportions of trace elements within the mineral, or structural defects within the crystal lattice. In the case of quartz (Figure 5.31), milky quartz gets its white colour from millions of tiny fluid-filled cavities. Smoky quartz gets its grey colour from structural damage caused by natural radiation. Amethyst and citrine get their colours from trace amounts of iron, and rose quartz gets its pink hue from manganese.

Figure 5.31 The many colours of quartz.Quartz can be colourless, milky, a greyish smoky colour, purple, yellow, and pink. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Photos by R. Weller/ Cochise College. Click the image for photo sources.

Streak

The colour of a mineral is what you see when light reflects off the surface of the sample. One reason that colour can be so variable is that the surface textureis variable. A way to get around this problem is to grind a small amount of the sample to a powder and observe the colour of the powder. This colour is the mineral’s streak. The mineral can be powdered by scraping the sample across a piece of unglazed porcelain called a streak plate (Figure 5.32). In Figure 5.32, two samples of hematite have been scraped across the streak plate. Even though one sample is metallic and the other is deep red, both have a similar reddish-brown streak.

Figure 5.32 Hematite leaves a distinctive reddish-brown streak whether the sample is metallic or deep red. Source: Karla Panchuk (2015) CC BY 4.0

Streak is an especially helpful property when minerals look similar. In Figure 5.33 all of the minerals are dark in colour, with varying degrees of metallic sheen. The streaks of these minerals are much more distinctive.

Figure 5.33 Similar dark-grey minerals with varying degrees of metallic sheen leave different colours of streaks. The minerals are from upper left clockwise: hematite, magnetite, sphalerite, and galena. Source: Karla Panchuk (2015) CC BY 4.0

Lustre

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 lustre. Light does not pass through metals, and that is the main reason they look metallic (e.g., the hematite on left of Figure 5.32). Even a thin sheet of metal — such as aluminum foil — will be not permit light to pass through it. Many non-metallic minerals may look as if light will not pass through them, but if you look closely at a thin slice of the mineral you will see that the mineral is translucent or transparent.

If a non-metallic mineral has a shiny, reflective surface, it is said to have a glassy lustre.  The quartz crystals in Figure 5.31 are examples of minerals with glassy lustre. If the mineral surface is dull and non-reflective, it has an earthy lustre (like the hematite on the right of Figure 5.32). Other types of non-metallic lustres are silky, pearly, and resinous. Lustre is a good diagnostic property because most minerals will always appear either metallic or non-metallic, although as Figure 5.31 shows, there are exceptions.

Hardness

One of the most important diagnostic properties of a mineral is its hardness. In practical terms, hardness determines whether or not a mineral can be scratched by a particular material.

In 1812 German mineralogist Friedrich Mohs came up with a list of 10 minerals representing a wide range of hardness, and numbered them 1 through 10 in order of increasing hardness (Figure 5.34, horizontal axis). While each mineral on the list is harder than the one before it, the measured hardness (vertical axis) is not linear. Notice that apatite is about three times harder than fluorite, and diamond is three times harder than corundum.

Figure 5.34 Minerals and reference materials in the Mohs scale of hardness. The measured hardness values are Vickers Hardness numbers. Source: Steven Earle (2015) CC BY 4.0 view source

Some commonly available reference materials are also shown on this diagram, including a typical fingernailNote that artificial fingernails may be much harder than natural fingernails. Some materials used for artificial nails are harder than quartz. (2.5), a piece of copper wire (3.5), a knife blade or 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: 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. The minerals themselves can be used to test other minerals.

Crystal Habit

When minerals form within rocks, there is a possibility that they will form in distinctive crystal shapes if they are not crowded out by other pre-existing minerals. Every mineral has one or more distinctive crystal habits determined by their atomic structure, although it is not that common in ordinary rocks for the shapes to be obvious.

Quartz, for instance, will form six-sided prisms with pointed ends (Figure 5.35 left), 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 5.35 centre), but can also form crystals with 12 faces, known as dodecahedra. The mineral garnet also forms many-sided crystals with an over-all rounded shape (Figure 5.35 right).

Figure 5.35 Hexagonal prisms of quartz (left), intergrown cubic crystals of pyrite (centre), and 24-sided crystals of garnet (right). Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Photos by R. Weller/ Cochise College. Click the image for photo sources.

Some of the terms that are used to describe habit include bladed, botryoidal (grape-like), dendritic (branched), drusy (an encrustation of crystals), equant (similar size in all dimensions), fibrous, platy, prismatic (long and thin), and stubby.

Cleavage and Fracture

Cleavage and fracture describe how a mineral breaks. These characteristics 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 plane or planes, while fracture is an irregular break. Some minerals tend to cleave along planes at various fixed orientations. Some, like quartz, do not cleave at all, only fracture. Minerals that have cleavage can also fracture along surfaces that are not parallel to their cleavage planes.

The way minerals break is determined by the arrangement of atoms within them, and more specifically by the orientation of weaknesses within their crystal lattice. Graphite and mica break off in parallel sheets (Figure 5.36).

Figure 5.36 One direction of cleavage (basal cleavage). Left: Schematic of basal cleavage. Right: Muscovite showing basal cleavage. The white dashed line marks the edge of the cleavage plane. Source: Karla Panchuk (2018) CC BY-SA 4.0. Cleavage diagram modified after M.C. Rygel (2010) CC BY-SA 3.0 view source

Other minerals have two directions of cleavage, classified as two directions at 90° (Figure 5.37 top) and two directions not at 90° (Figure 5.37 bottom). While the diagrams of planes on the left of Figure 5.37 make this difference clear, it may be less obvious in practice. The minerals in Figure 5.37 both have two planes of cleavage that are very close to 90°.  The white dashed lines mark the edges of the planes, as with Figure 5.36.  See if you can find the planes repeated in the images.  The images are close-up views of the minerals, only a few cm across. Sometimes you must look very carefully to find cleavage planes.

Figure 5.37 Two directions of cleavage. Top: Two directions at 90° in pyroxene. Bottom: two directions not at 90° in plagioclase feldspar. Edges of cleavage planes marked with dashed lines. Source: Karla Panchuk (2018) CC BY-SA 4.0. Cleavage diagrams modified after M.C. Rygel (2010) CC BY-SA 3.0 view source

Some minerals have many directions of cleavage.  Figure 5.38 shows  minerals with three directions of cleavage.  Halite (Figure 5.38 top) has three directions at 90° and calcite (Figure 5.38 bottom) has three directions not at 90°.

Figure 5.38 Three directions of cleavage. Top: Three directions at 90° in halite. Bottom: Three directions not at 90° in calcite. Source: Karla Panchuk (2018) CC BY-SA 4.0. Cleavage diagrams modified after M.C. Rygel (2010) CC BY-SA 3.0 view source

There are a few common difficulties that students encounter when learning to recognize and describe cleavage.  One is that it might be necessary to look very closely at a sample to see mineral cleavage.  The key features in Figure 5.37 are only cm or mm in scale.  If crystals are very small, it may not be possible to see cleavage at all. Another issue is that sometimes cleavage is present, but it is poor, meaning the cleavage surface isn’t perfectly flat. Finally it can be difficult to know whether a flat surface on a crystal is a cleavage plane, a crystal face, or simply a surface that happens to be flat. Cleavage planes tend to repeat themselves at different depths throughout the mineral, so if you are unsure whether the surface you are looking at is a cleavage plane, try rotating the mineral in bright light. If cleavage is present, you will generally find that, for a given cleavage direction, all of the cleavage surfaces will glint in the light simultaneously. Crystal faces will also glint in light, but they do not repeat themselves at depth throughout the mineral. The best way to overcome all of these problems is to look at lots of examples.  It’s worth it to be able to identify cleavage and fracture, because cleavage is a reliable diagnostic property for most minerals.

Density

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, are of average density (2.6 to 3.0 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. If you picked up a sample of one of these minerals, it would feel much heavier compared to a similarly sized sample of a mineral with average density. 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 with other minerals in it.

Other properties

Several other properties are useful for identification of some minerals. Some of these are:

• Calcite reacts with dilute acid and will give off bubbles of carbon dioxide.
• Magnetite is strongly magnetic, and some other minerals are weakly magnetic.
• Sphalerite ((Zn,Fe)S) gives off a smell of sulphur when drawn across a streak plate.
• Halite tastes salty.
• Talc feels soapy to the touch.
• Plagioclase feldspar has striations (parallel razor-thin lines etched on the surface) and some varieties show a play of colours when light hits them at the right angle (see the labradorite in Figure 5.28).

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

5.1 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. Atoms strive to have eight electrons in their outermost shell (or two for H and He). Atoms gain, lose, or share electrons to achieve this. In so doing they become either positively charged cations (if they lose electrons) or negatively charged anions (if they gain them).

5.2 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 weak Van der Waals forces. Minerals form in three-dimensional lattices. The configuration of the lattices and the type of bonding within help determine mineral properties.

5.3 Mineral Groups

Minerals are grouped according to the anion part of their formula. Some common types are: oxides, sulphides, sulphates, halides, carbonates, phosphates, silicates, and native minerals.

5.4 Silicate Minerals

Silicate minerals are the most common minerals in Earth’s crust and mantle. They all have silica tetrahedra (four oxygens surrounding a single silicon atom) arranged in different structures (chains, sheets, etc).

5.5 How Minerals Form

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. More rarely, minerals precipitate directly from a gas, such as at a volcanic vent.

5.6 Mineral Properties

Some of the important properties for mineral identification include hardness, cleavage/fracture, density, lustre, colour, and streak colour.

Review Questions

1. What is the electrical charge of 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. How do cations differ from anions?

6. What chemical feature is used in the classification of minerals into groups?

7. Name the mineral group for the following minerals:

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?

1. Charges: proton: +1, neutron: 0, electron: -1, Masses: proton: 1, neutron: 1, electron: almost 0.

2. The element’s atomic number will determine the extent to which its outer layers are populated with electrons. If the outer shell is not quite full, the atom may gain electrons to fill them and become an anion (negative charge). If the outer shell has only a few electrons, it may lose them and become a cation (positive charge). Cations and anions attract each other to form molecules with ionic bonding.

3. Helium and neon (and the other noble gases) have complete outer shells and therefore no tendency to form ionic bonds.

4. Electrons are transferred from one atom to another to form an ionic bond. Electrons are shared between atoms to form a covalent bond.

5. An anion has a negative charge and a cation has a positive charge.

6. Minerals are classified into groups based on their anion or anion group.

7. Name the mineral group for the following minerals:

8. An unbonded silica tetrahedron has one Si ion (+4 charge) and 4 oxygens (-2 charge each) so the overall charge is 4 – 8 = -4 for SiO₄^-4

9. Magnesium can substitute freely for iron in olivine and several other minerals because they have similar charges (+2) and similar ionic radii.

10. Pyroxene is made up of single chains of tetrahedra while amphibole is made up of double chains.

11. The two end-members of the plagioclase series are Albite (NaAlSi₃O₈) and Anorthite (CaAl₂Si₂O₈)

12. In quartz each silica tetrahedron is bonded to four other tetrahedra, and because oxygens are shared at each bond the overall ratio is silicon (+4) to two oxygens (2 x -2 = -4), which is balanced.

13. Some minerals have distinctive colours, but many have a wide range of colours due to differing impurities.

14. Glass has a Mohs hardness of about 5.5 while porcelain is close to 6.5. The mineral is between these two, so it must be close to 6.

A rock is a solid mass of geological materials. Geological materials include individual mineral crystals, inorganic non-mineral solids like glass, pieces broken from other rocks, and even fossils. The geological materials in rocks may be inorganic, but they can also include organic materials such as the partially decomposed plant matter preserved in coal. A rock can be composed of only one type of geological material or mineral, but many are composed of several types. Figure 6.2 shows a rock made of three different kinds of minerals.

Rocks are grouped into three main categories based on how they form. Igneous rocks form when melted rock cools and solidifies. Sedimentary rocks form when fragments of other rocks are buried, compressed, and cemented together; or when minerals precipitate from solution, either directly or with the help of an organism. Metamorphic rocks form when heat and pressure alter a pre-existing rock. Although temperatures can be very high, metamorphism does not involve melting of the rock.

Figure 6.2 This close-up view of the igneous rock pegmatite shows black biotite crystals, colourless quartz crystals, and pink potassium feldspar crystals. Crystals are mm to cm in scale. Source: R. Weller/ Cochise College (2011) Permission for non-commercial educational use. (labels added) view source

The rock components of the crust are slowly but constantly being changed from one form to another. The processes involved are summarized in the rock cycle (Figure 6.3). The rock cycle is driven by two forces:

1. Earth’s internal heat, which causes material to move around in the core and mantle, driving plate tectonics.
2. The hydrological cycle– movement of water, ice, and air at the surface. The hydrological cycle is powered by the sun.

Figure 6.3 The rock cycle describes processes that form the three types of rock: igneous, sedimentary, and metamorphic. These same processes can turn one type of rock into another. Source: Karla Panchuk (2017) CC BY-SA 4.0. Click the image for more attributions.

The rock cycle is still active on Earth because our core is hot enough to keep the mantle moving, the atmosphere is relatively thick, and there is liquid water. On some other planets or their satellites (e.g., Mercury), 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.

We can start anywhere we like to describe the rock cycle, but it’s convenient to start with magma. Magma is melted rock located within the Earth.  Rock can melt at between about 800 °C and 1300 °C, depending on the minerals in the rock, and the pressure the rock is under.  If it cools slowly within the Earth (over centuries to millions of years), magma forms intrusive igneous rocks.  If magma erupts onto the surface, we refer to it as lava.  Lava cools rapidly on Earth’s surface (within seconds to years) and forms extrusive igneous rocks (Figure 6.4).Remember the difference between intrusive and extrusive igneous rocks by recalling that INtrusive rocks form withIN the Earth, and EXtrusive rocks form when lava EXits the Earth's crust.

Figure 6.4 Lava flowing from Kīlauea Volcano, Hawai`i. Source: J. D. Griggs, U. S. Geological Survey (1985) Public Domain view source

Mountain building lifts rocks upward where they are acted upon by weathering. Weathering includes chemical processes that break rocks apart, as well as physical processes. Figure 6.5 shows the result of rocks in mountains being broken apart when water gets into cracks, freezes, and forces the cracks wider. Uplift through mountain building is how rocks once buried deep within Earth can be exposed at Earth’s surface.

Figure 6.5 Mountains being broken apart by the wedging action of ice near La Madaleta Glacier, Spain. Source: Luis Paquito (2006) CC BY-SA 2.0 view source

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. Figure 6.6 shows transportation of fine-grained sediment particles by wind during the Great Depression in the 1930s.

Figure 6.6 Wind transports sediment in a dust storm near Okotoks, Alberta, Canada in July of 1933. Source: Glenbow Museum Archives, File Number NA-2199-1 (1933) Public Domain view source

Sediments are deposited in stream channels, lakes, deserts, and the ocean. Some depositional settings result in characteristic sedimentary structures, such as the ripples that formed when flowing water moved sand along the bottom of the South Saskatchewan River (Figure 6.7).

Figure 6.7 Sand ripples along the South Saskatchewan River, near Saskatoon SK (dog for scale). Source: Karla Panchuk (2008) CC BY-SA 4.0 view source

Unless they are re-eroded and moved along, sediments will eventually be buried by more sediments. At depths of hundreds of metres or more, the sediments become compressed, forcing particles closer together. Mineral crystals grow around and between the particles, binding them together (cementing them). The hardened cemented sediments are sedimentary rock. Figure 6.8 shows an example of an ancient sedimentary rock in which ripple structures are preserved, and visible in cross-section as wavy lines.

Figure 6.8 Ripples preserved in 1.2 Ga old sandstone. Notice the wavy lines above the coin. This is a side view of the ripples. Source: Anne Burgess (2008) CC BY-SA 2.0 view source

Rocks that are buried very deeply within the crust can reach pressures and temperatures much higher than those at which sedimentary rocks form. Existing rocks that are heated up and squeezed under those extreme conditions are transformed into metamorphic rocks (Figure 6.9). The transformation to a metamorphic rock can happen through physical changes, such as when the minerals making up an existing rock re-form into larger crystals of the same mineral. It can also happen through chemical changes, when minerals within the rock react to form new minerals.

Figure 6.9 Limestone, a sedimentary rock formed in marine waters, has been altered by metamorphism into this marble visible on Quadra Island, BC. Source: Steven Earle (2015) CC BY 4.0 view source 

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

6.1 What Is a Rock?

A rock is a solid mass of geological materials. Geological materials include individual mineral crystals, inorganic non-mineral solids like glass, pieces broken from other rocks, and even fossils.

6.2 The Rock Cycle

There are three main types of rock. Igneous rocks form when melted rock cools and solidifies. Sedimentary rock forms from fragments of other rocks, or when crystals precipitate from solution. Metamorphic rocks form when existing rocks are altered by heat, pressure, and/or chemical reactions. The rock cycle summarizes the processes that contribute to transformation of rock from one type to another. The rock cycle is driven by Earth’s internal heat, and by processes happening at the surface that are driven by solar energy.

Review Questions

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?

1. The rock must be exposed at surface. This means uplift and removal of overlying rocks and sediments is required. Once exposed, chemical and/or physical weathering can reduce the rock to smaller loose fragments (sediments). The sediments can be eroded and then transported by a variety of mechanisms.
2. Sediments are buried beneath other sediments, where pressure compacts the sediments, forcing grains closer together. Mineral cement forms around the grains, binding them to each other and into solid rock.
3. Rock is buried deeply in the crust and exposed to very high temperatures and pressures. Under those conditions, a new type of rock is formed when minerals undergo physical changes and chemical reactions.

Igneous rocks form when melted rock cools. Melted rock originates within Earth as magma.  Magma compositions vary, but will have eight main elements in different proportions. The most abundant elements are oxygen and silicon, followed by aluminum, iron, calcium, sodium, magnesium, and potassium. These eight elements are also the most abundant in Earth’s crust (Figure 7.2).  All magmas have varying proportions of lighter elements such as hydrogen, carbon, and sulphur. Lighter elements are converted into gases like water vapour, carbon dioxide, hydrogen sulphide, and sulphur dioxide as the magma cools.

Figure 7.2 Average composition of Earth’s crust by mass. Source: Steven Earle (2016) CC BY 4.0 view source

Magma composition depends on the composition of the rocks that melted to form the magma, and on the conditions under which the melting happened. Most igneous rock in Earth’s crust comes from magmas that formed through partial melting of existing rock, either in the upper mantle or the crust. During partial melting, only some of the minerals within a rock melt. This happens because different minerals have different melting temperatures. The melt is less dense than the surrounding rock, and will percolate upward without the source rock having melted completely. The result is magma with a different composition than the original rock. Partial melting produces melt that has more silica than the original rock, because minerals higher in silica have lower melting points.

To see how partial melting works, consider the mix of materials in Figure 7.3a. It contains white blocks of candle wax, black plastic pipe, green beach glass, and pieces of aluminum wire. When the mixture is heated to 50 °C in a warm oven, the wax melts into a clear liquid (Figure 7.3b), but the other materials remain solid. This is partial melting.

Figure 7.3 An experiment to illustrate partial melting. (a) The original components are 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 poured off and allowed to cool, making a solid with a different overall composition from the original mixture. Source: Steven Earle (2015) CC BY 4.0 view source

When the mixture is heated to 120 °C, the plastic melts and mixes with the wax, but the aluminum and glass still remain solid (Figure 7.3c). This is still considered partial melting because solid materials remain. When the plastic and wax mixture is poured into a separate container and allowed to cool, the resulting solid has a very different composition from the original mixture (Figure 7.3d). The plastic and wax are analogous to more silica-rich minerals with relatively lower melting points than other minerals in the same rock.

Of course, partial melting in the real world isn’t as simple as the example in Figure 7.3. Many rocks are much more complex than the four-component system used here. Some mineral components of rocks may have similar melting temperatures, and begin to melt at the same time. The melting temperature of a mineral may change in the presence of other minerals. Also, when rocks melt, the process can take millions of years, unlike the 90 minutes required to melt the pipe and wax in the experiment in Figure 7.3.

Why Rocks Melt

The magma that is produced by partial melting is less dense than the surrounding rock. Magma from partial melting of mantle rocks rises upward through the mantle, and may pool at the base of the crust, or rise through the crust. Moving magma carries heat with it, and some of that heat is transferred to surrounding rocks. If the melting temperature of a rock is less than the temperature of the magma, the rock will begin to melt, and the composition of the magma may change to reflect a mixture of sources. But adding heat is not the only way to trigger melting.

Decompression Melting

Earth’s mantle is almost entirely solid rock, in spite of temperatures that would cause rock at Earth’s surface to melt. Mantle rock remains solid at those temperatures because the rock is under high pressure. This means that melting can be triggered without adding heat if the rock is already hot enough, and the pressure is reduced (Figure 7.4, left, white dashed boxes). Melting triggered by a reduction in pressure is called decompression melting.

Figure 7.4 Melting triggers. Left- Decompression melting occurs when rock rises or the overlying crust thins. Right- Flux-induced melting occurs when volatile compounds such as water are added. Source: Karla Panchuk (2018) CC BY 4.0. Modified after Steven Earle (2016) CC BY 4.0 view source

Pressure is reduced when mantle rocks move upward due to convection, or rise as a plume within the mantle. Pressure is also reduced where the crust thins, such as along rift zones.

Flux-induced Melting

When a substance such as water is added to hot rocks, the melting points of the minerals within those rocks decreases. If a rock is already close to its melting point, the effect of adding water can be enough to trigger partial melting. The added water is a flux, and this type of melting is called flux-induced melting. In Figure 7.4 (right), the rock (represented by the dashed box) is not hot enough to be right of the line where dry mantle rocks melt, but it is to the right of the line where wet mantle rocks melt.

Flux-induced partial melting of rock happens in subduction zones. Minerals are transformed by chemical reactions under high pressures and temperatures, and a by-product of those transformations is water. Relatively little water is required to trigger partial melting. In laboratory studies of the conditions of partial melting in the Japanese volcanic arc, rocks with only 0.2% of their weight consisting of water melted by up to 25%.

Cooling Magma Becomes More Viscous

Viscosity refers to the ease with which a substance flows. A substance with low viscosity is runnier than a substance with high viscosity. At temperatures over 1300°C, most magma is entirely liquid because there is too much energy for the atoms to bond together. As magma loses heat to the surrounding rocks and its temperature drops, things start to change. Silicon and oxygen combine to form silica tetrahedra.  With further cooling, the tetrahedra start to link together into chains, or polymerize. These silica chains make the magma more viscous. Magma viscosity has important implications for the characteristics of volcanic eruptions.

Exercise: Making Magma Viscous

This is a quick and easy 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 1/2 cup (125 mL) of water in a saucepan over medium heat. Add 2 teaspoons (10 mL) of white flour and stir while continuing to heat the mixture until boiling. The white flour represents silica. The mixture should thicken like gravy because the gluten in the flour becomes polymerized into chains during this process.

Now 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 4 teaspoons (20 mL) of water in a cup. Add 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 (Figure 7.5) because there is more gluten, and more chains have formed.

Figure 7.5 Thick mixture of flour and water. Source: Steven Earle (2016) CC BY 4.0 view source

References

Kushiro, I. (2007). Origins of magmas in subduction zones: a review of experimental studies. Proceedings of the Japan Academy, Series B, Physical and Biological Sciences 83(1), 1-15. Read the paper

The minerals that make up igneous rocks crystallize (solidify, freeze) at a range of different temperatures. This explains why cooling magma can have some crystals within it and yet remain predominantly liquid. The sequence in which minerals crystallize from a magma as it cools is known as Bowen’s reaction series (Figure 7.6).

Figure 7.6 Bowen’s reaction series describe the sequence in which minerals form as magma cools. Source: Steven Earle (2016) CC BY 4.0 view source

How Did We Get Bowen’s Reaction Series?

Understanding how the reaction series was derived is key to understanding what it means.

Figure 7.7 Norman Bowen in his laboratory. Source: University of Chicago Photographic Archive, apf1-00841, Special Collections Research Center, University of Chicago Library. Click the image for source and terms of use.

Norman Levi Bowen (Figure 7.7) was born in Kingston Ontario. He studied geology at Queen’s University and then at Massachusetts Institute of Technology. In 1912 he joined the Carnegie Institution in Washington, D.C., where he carried out ground-breaking experiments into how magma cools.

Working mostly with mafic magmas (magmas rich in iron and magnesium), he determined the order of crystallization of minerals as the temperature drops. First, he melted the rock completely in a specially made kiln. Then he allowed it to cool slowly to a specific temperature before quenching (cooling it quickly) so that no new minerals could form. The rocks that formed were studied under the microscope and analyzed chemically. This was done over and over, each time allowing the magma to cool to a lower temperature before quenching.

The result of these experiments was the reaction series which, even a century later, is still an important basis for our understanding of igneous rocks.

Discontinuous and Continuous Series

Bowen’s reaction series (Figure 7.6) has two pathways for minerals to form as magma cools: on the left is the discontinuous series. This refers to the fact that one mineral is transformed into a different mineral through chemical reactions. On the right is the continuous series, where plagioclase feldspar goes from being rich in calcium to being rich in sodium.

Continuous Series

At about the point where pyroxene begins to crystallize, plagioclase feldspar also begins to crystallize. At that temperature, the plagioclase is calcium-rich (toward the anorthite end-member). As the temperature drops, and providing that there is sodium left in the magma, the plagioclase that forms is a more sodium-rich variety (toward the albite end-member). The series is continuous because the mineral is always plagioclase feldspar, but the series involves a transition from calcium-rich to sodium-rich.

When cooling happens relatively quickly, instead of getting crystals which are of uniform composition, individual plagioclase crystals can be zoned from calcium-rich in the centre to more sodium-rich around the outside (Figure 7.8). This occurs because calcium-rich early-forming plagioclase crystals become coated with progressively more sodium-rich plagioclase as the magma cools.

Figure 7.8 Plagioclase crystal exhibiting compositional zones. Source: Akademia Górniczo-Hutnicza w Krakowie Otwartych Zasobów Edukacyjnych (n.d.) CC BY-NC-SA view source/ view context

Discontinuous Series

Olivine begins to form at just below 1300°C, but as the temperature drops, olivine becomes unstable. The early-forming olivine crystals react with silica in the remaining liquid and are converted into pyroxene, something like this:

As long as there is silica remaining and the rate of cooling is slow, this process continues down the discontinuous branch: olivine reacts to form pyroxene, and the pyroxene reacts to form amphibole. Under the right conditions amphibole will form to biotite. Finally, if the magma is quite silica-rich to begin with, there will still be some left at around 750 °C to 800 °C, and from this last magma, potassium feldspar, quartz, and maybe muscovite mica will form.

Notice that the sequence of minerals that form goes from isolated tetrahedra (olivine) toward increasingly complex arrangements of silica tetrahedra. Pyroxene consists of single chains, amphibole has double chains, mica has sheets of tetrahedra, and potassium feldspar and quartz at the bottom of the series have tetrahedra connected to each other in three dimensions.

If the magma cools enough, the first minerals to form will be completely used up in later chemical reactions.  This is why igneous rocks do not normally have both olivine (at the top of the series) and quartz (at the bottom). Exceptions can occur when rocks that crystallized early in the series come into contact with magmas representing compositions later in the series, such as with the dark green olivine-rich xenoliths included within the quartz- and feldspar-rich rock in Figure 7.9. The dark line around the xenoliths is amphibole, which formed as the olivine reacted with the melt. In some of the smaller xenoliths within this boulder, the olivine has been completely transformed into amphibole.

Figure 7.9 Boulder with olivine-rich xenoliths surrounded by silica-rich rock. Black rims on the xenoliths are where the olivine has reacted with the silica-rich melt, forming amphibole. Right- Enlarged view of the amphibole reaction rim. Source: Karla Panchuk (2018) CC BY 4.0

Magma Composition: Mafic, Intermediate, and Felsic

The composition of the original magma determines how far the reaction process can continue before all of the magma is used up. In other words, it determines which minerals will form. Magma compositions are reported in terms of the fraction of mass of oxides (e.g., Al₂O₃ rather than just Al; Figure 7.10). On average, mafic"Mafic" combines the words MAgnesium and FerrIC (containing iron). magma (Figure 7.10, left) is approximately half SiO₂ by mass, and more than 25% iron, magnesium, and calcium oxides by mass. Average felsic"Felsic" combines the words FELdspar and SIliCa. magmas (Figure 7.10, right) are closer to 75% SiO₂ by mass, and have approximately 5% iron, magnesium, and calcium oxides. Sodium and potassium oxides account for approximately 10% of felsic magmas by mass, but only 5% of mafic magmas. Magmas that fall between mafic and felsic magmas have an intermediate composition (Figure 7.10, centre).

Figure 7.10 Chemical compositions of typical mafic, intermediate, and felsic magmas. Source: Karla Panchuk (2018) CC BY 4.0 modified after Steven Earle (2016) CC BY 4.0 view source

What Determines the Composition of Magma?

Why Is There No Ultramafic Magma Anymore?

Partial Melting Makes Magma That Is Richer in Silica

In partial melting, some components of a mixture melt before others do.  In the case of mafic magma, it is produced when ultramafic rocks undergo partial melting. In general, silicate minerals with more silica will melt before those with less silica.  This means the partial melt will have more silica than the rock as a whole.

Fractional Crystallization Also Makes Magma Richer In Silica

A number of processes that take place within a magma chamber can affect the types of rocks that form once magma cools and crystallizes. If the magma has a low viscosity— which is likely if the magma is mafic—the crystals that form early, such as olivine (Figure 7.11a), may slowly settle toward the bottom of the magma chamber (Figure 7.11b). This process is called fractional crystallization.

Figure 7.11 Formation of a zoned magma chamber. a- Olivine crystals form. b- Olivine crystals settle to the base of the magma chamber, leaving the upper part of the chamber richer in silica. c- Olivine crystals remelt, making magma at the base of the chamber more mafic. Source: Steven Earle (2015) CC BY 4.0 view source

The formation of olivine removes iron- and magnesium-rich components, leaving the overall composition of the magma near the top of the magma chamber more felsic. The crystals that settle might either form an olivine-rich layer near the bottom of the magma chamber. Or, because the lower part of the magma chamber is likely to be hotter than the upper part, the crystals might remelt. If remelting happens, crystal settling will make the magma at the bottom of the chamber more mafic than it was to begin with (Figure 7.11c).

Magma Composition Also Changes When Other Rocks Are Melted And Mixed In

Magma chambers aren’t isolated from their surroundings.  If the rock in which the magma chamber is located (called the country rock) is more felsic than the magma, the country rock may also melt, adding to the magma already in the magma chamber (Figure 7.12).  Sometimes magma carries fragments of unmelted rock, called xenoliths, within it.  Melting of xenoliths can also alter the composition of magma, as can re-melting of crystals that have settled out of the magma.

Figure 7.12 The composition of magma in a magma chamber is affected by fractional crystallization within the magma chamber, but it can also be affected by partial melting of the rock surrounding the magma chamber, melting of xenoliths within the magma, or re-melting of crystals that have settled to the bottom of the magma chamber. Source: Steven Earle (2015) CC BY 4.0 view source

Classification By Mineral Abundance

Igneous rocks can be divided into four categories based on their chemical composition: felsic, intermediate, mafic, and ultramafic. The diagram of Bowen’s reaction series (Figure 7.6) shows that differences in chemical composition correspond to differences in the types of minerals within an igneous rock.  Igneous rocks are given names based on the proportion of different minerals they contain.  Figure 7.13 is a diagram with the minerals from Bowen’s reaction series, and is used to decide which name to give an igneous rock.

Figure 7.13 Classification diagram for igneous rocks. Igneous rocks are classified according to the relative abundances of minerals they contain. A given rock is represented by a vertical line in the diagram. In the mafic field, the arrows represent a rock containing 48% pyroxene and 52% plagioclase feldspar. The name an igneous rock gets depends not only on composition, but on whether it is intrusive or extrusive. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0, modified after Steven Earle (2015) CC BY 4.0 view source and others, with photos by R. Weller/Cochise College. Click the image for links to photos and notes on image construction. High-resolution version.

To see how Figure 7.13 works, first notice the scale in percent along the vertical axis.  The interval between each tick mark represents 10% of the minerals within a rock by volume.  An igneous rock can be represented as a vertical line drawn through the diagram, and the vertical scale used to break down the proportion of each mineral it contains.  For example, the arrows in the mafic field of the diagram represent a rock containing 48% pyroxene and 52% plagioclase feldspar. An igneous rock at the boundary between the mafic and ultramafic fields (marked with a vertical dashed line) would have approximately 20% olivine, 50% pyroxene, and 30% Ca-rich plagioclase feldspar by volume.

Classification By Grain Size

The name an igneous rock gets also depends on whether it cools within Earth (an intrusive or plutonic igneous rock), or whether it cools on the Earth’s surface after erupting from a volcano (an extrusive or volcanic igneous rock). For example, a felsic intrusive rock is called granite, whereas a felsic extrusive rock is called rhyolite. Granite and rhyolite have the same mineral composition, but their grain size gives each a distinct appearance.

The key difference between intrusive and extrusive igneous rocks—the size of crystals making them up—is related to how rapidly melted rock cools. The longer melted rock has to cool, the larger the crystals within it can become.  Magma cools much slower within Earth than on Earth’s surface because magma within Earth is insulated by surrounding rock.  Notice that in Figure 7.13, the intrusive rocks have crystals large enough that you can see individual crystals—either by identifying their boundaries, or seeing light reflecting from a crystal face.  A rock with individual crystals that are visible to the unaided eye has a phaneritic or coarse-grained texture. The extrusive rocks in the second row have much smaller crystals.  The crystals are so small that individual crystals cannot be distinguished, and the rock looks like a dull mass. A rock with crystals that are too small to see with the unaided eye has an aphanitic or fine-grained texture.  Table 7.1 summarizes the key differences between intrusive and extrusive igneous rocks.

What this means is that two igneous rocks comprised of exactly the same minerals, and in the exactly the same proportions, can have different names.  A rock of intermediate composition is diorite if it is course-grained, and andesite if it is fine-grained.  A mafic rock is gabbro if it is course-grained, and basalt if fine-grained. The course-grained version of an ultramafic rock is peridotite, and the fine-grained version is komatiite. It makes sense to use different names because rocks of different grain sizes form in different ways and in different geological settings.

Does This Mean an Igneous Rock Can Only Have One Grain Size?

No. Something interesting happens when there is a change in the rate at which melted rock is cooling.  If magma is cooling in a magma chamber, some minerals will begin to crystallize before others do.  If cooling is slow enough, those crystals can become quite large.

Now imagine the magma is suddenly heaved out of the magma chamber and erupted from a volcano.  The larger crystals will flow out with the lava. The lava will then cool rapidly, and the larger crystals will be surrounded by much smaller ones.  An igneous rock with crystals of distinctly different size (Figure 7.14) is said to have a porphyritic texture, or might be referred to as a porphyry.  The larger crystals are called phenocrysts, and the smaller ones are referred to as the groundmass.

Figure 7.14 Porphyritic rhyolite with quartz and potassium feldspar phenocrysts within a dark groundmass. Porphyritic texture (when different crystal sizes are present) is an indication that melted rock did not cool at a constant rate. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Photo by R. Weller/Cochise College (2011) view source

Exercise: Which Mineral Will the Phenocryst Be?

As a magma cools below 1300°C, minerals start to crystallize within it. If the magma is then erupted, 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 the diagram shown here, predict what phenocrysts might be present where the magma cooled as far as line a. Which would be present where magma cooled to line b?

Figure 7.15 Bowen’s reaction series. Source: Steven Earle (2015) CC BY 4.0 view source

Classifying Igneous Rocks According to the Proportion of Dark Minerals

If you unsure of which minerals are present in an intrusive igneous rock, there is a quick way to approximate the composition of that rock.  In general, igneous rocks have an increasing proportion of dark minerals as they become more mafic (Figure 7.16).

Figure 7.16 Simplified igneous rock classification according to the proportion of light and dark (or ferromagnesian) minerals. Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2015) CC BY 4.0 view source

The dark-coloured minerals are those higher in iron and magnesium (e.g., olivine, pyroxene, amphibole, biotite), and for that reason they are sometimes referred to collectively as ferromagnesian minerals. By estimating the proportion of light minerals to dark minerals in a sample, it is possible to place that sample in Figure 7.16.  Graphical scales are used to help visualize the proportions of light and dark minerals (Figure 7.17).

Figure 7.17 A guide for estimating the proportion of dark minerals in an igneous rock. Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2015) view source

It is important to note that estimating the proportion of dark minerals is only approximate as a means for identifying igneous rocks. One problem is that plagioclase feldspar is light-coloured when it is sodium-rich, but can appear darker if it is calcium-rich. Plagioclase feldspar is not ferromagnesian, so it falls in the non-ferromagnesian (light minerals) region in Figure 7.16 even when it has a darker colour.

Exercise: Classifying Igneous Rocks by the Proportion of Dark Minerals

The four igneous rocks shown below have differing proportions of ferromagnesian silicates (dark minerals). Estimate the proportion of dark minerals using the guide in Figure 7.17, and then use Figure 7.16 to determine the likely rock name for each one.

Figure 7.18 Identify these rocks by estimating the proportion of dark minerals in each sample. Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2015) CC BY 4.0 view source

 

Classifying Igneous Rocks When Individual Crystals Are Not Visible

The method of estimating the percentage of minerals works well for phaneritic igneous rocks, in which individual crystals are visible with little to no magnification. If an igneous rock is porphyritic but otherwise aphanitic (e.g., Figure 7.14), the minerals present as phenocrysts give clues to the identity of the rock. However, there are cases where mineral composition cannot be determined by looking at visible crystals. These include volcanic rocks without phenocrysts, and glassy igneous rocks.

Volcanic Rocks Without Phenocrysts

In the absence of visible crystals or phenocrysts, volcanic rocks are be classified on the basis of colour and other textural features. As you may have noticed in Figure 7.13, the colour of volcanic rocks goes from light to dark as the composition goes from felsic to mafic. Rhyolite is often a tan or pinkish colour, andesite is often grey, and basalt ranges from brown to dark green to black (Figure 7.19).

Figure 7.19 In volcanic igneous rocks, individual crystals are not visible. Colours change from light to dark as the composition of the rocks go from felsic to mafic. Vesicles and amygdules are common characteristics of basalt. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Photos by R. Weller/ Cochise College. Click the image for links to photos.

Basalt often shows textural features related to lava freezing around gas bubbles. When magma is underground, pressure keeps gases dissolved, but once magma has erupted, the pressure is much lower. Gases dissolved in the lava are released, and bubbles can develop. When lava freezes around the bubbles, vesicles are formed (circular inset in 7.19). If the vesicles are later filled by other minerals, the filled vesicles are called amygdules (box inset in Figure 7.19).

Glassy Volcanic Rocks

Crystal size is a function of cooling rate. The faster magma or lava cools, the smaller the crystals it contains. It is possible for lava to cool so rapidly that no crystals can form. The result is called volcanic glass. Volcanic glass can be smooth like obsidian or vesicular like scoria (mafic) and pumice (felsic; Figure 7.20). Pumice can float on water because of its low-density felsic composition and enclosed vesicles.

Figure 7.20 Glassy volcanic rocks. Obsidian has a glassy lustre, but scoria and pumice are highly vesicular. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Photos by R. Weller/Cochise College. Click the image for links to photos.

In most cases, a body of hot magma is less dense than the rock surrounding it, so it has a tendency to creep upward toward the surface. It does so in a few different ways:

• Filling and widening existing cracks
• Melting the surrounding rock
• Pushing the rock aside (where the rock is hot enough and under enough pressure to deform without breaking)
• Breaking the rock

When magma forces itself into cracks, breaks off pieces of rock, and then envelops them, this is called stoping.  The resulting fragments are xenolithsFrom the Greek words xenos, meaning "foreigner" or "stranger," and lithos for "stone.". Xenoliths may appear as dark patches within a rock (Figure 7.21).

Figure 7.21 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. Source: Steven Earle (2015) CC BY 4.0 view source

Some of the magma may reach the surface, resulting in volcanic eruptions, but most cools within the crust. The resulting body of rock is called a pluton.After Pluto was demoted from planet status, astronomers tried to come up with a name for objects like Pluto. For a while they considered "pluton" however geologists rightly objected that they had first claim on the word. In the end the International Astronomical Union settled on "dwarf planet" instead. Plutons can have different shapes and different relationships with the surrounding country rock (Figure 7.22). These characteristics determine what name the pluton is given.

Large, irregularly shaped plutons are called stocks or batholiths, depending upon their size. Tabular plutons are called dikes if they cut across existing structures, and sills if they are parallel to existing structures. Laccoliths are like sills, except they have caused the overlying rocks to bulge upward. Pipes are cylindrical conduits.

Figure 7.22 Plutons can have a variety of shapes, and be positioned in a variety of ways relative to the surrounding rocks. They are named according to these characteristics. Source: Karla Panchuk (2018) CC BY 4.0

Types of Plutons

Stocks and Batholiths

Large irregular-shaped plutons are called either stocks or batholiths, depending on their area. If an irregularly shaped body has an area greater than 100 km², then it’s a batholith, otherwise it’s a stock. Note that our knowledge of the size of a body can be limited to what we see at the surface. A body with an area of less than 100 km²  exposed at the surface might in fact be much larger at depth. It might be classified as a stock initially, until someone is able to map out its true extent.

Batholiths are typically formed 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 ComplexAlso referred to as the Coast Range Batholith, which extends all the way from the Vancouver region to southeastern Alaska (Figure 7.23).

Figure 7.23 The Coast Range Plutonic Complex (also called the Coast Range Batholith) is the largest in the world. It is part of a chain of batholiths along the western coast of North America. Source: Karla Panchuk (2018) CC BY 4.0. Modified after Bally (1989).

Tabular Intrusions

Tabular (sheet-like) plutons are classified according to whether or not they are concordant with (parallel to) existing layering (e.g., sedimentary bedding or metamorphic foliationSedimentary bedding refers to the layers in which sedimentary rocks form. Metamorphic foliation refers to the way minerals or other elements in a rock are aligned as a result of being deformed by heat and pressure. Bedding and foliation will be discussed in more detail in later chapters.) in the country rock. A sill is concordant with existing layering, and a dikeAlso spelled dyke. is discordant. If the country rock has no bedding or foliation, then any tabular body within it is a dike. Note that the sill-versus-dike designation is not determined simply by the orientation of the feature. A dike could be horizontal and a sill could be vertical- it all depends on the orientation of features in the surrounding rocks.

A laccolith is a sill-like body that has expanded upward by deforming the overlying rock. If a sill forms, but magma pools and sags downward, it creates a lopolith.

Pipes

A pipe, as the name suggests, is a cylindrical body with a circular, elliptical, or even irregular cross-section, that serves as a conduit (or pipeline) for the movement of magma from one location to another. Pipes may feed volcanoes, but pipes can also connect plutons.

Chilled Margins

As discussed already, plutons can interact with the rocks into which they are intruded. Partial melting of the country rock may occur, or stoping may form xenoliths. The heat from magma can even cause causing mineralogical and textural changes in country rock. However, country rock can also affect the magma.

The most obvious effect that country rock can have on magma is a chilled margin along the edges of the pluton (Figure 7.24). The country rock is much cooler than the magma, so magma that comes into contact with the country rock cools faster than magma toward the interior of the pluton. Rapid cooling leads to smaller crystals, so the texture along the edges of the pluton is different from that of the interior of the pluton, and the colour may be darker.

Figure 7.24 A mafic dike with chilled margins within basalt at Nanoose, B.C. The coin is 24 mm in diameter. The dike is about 25 cm across and the chilled margins are 2 cm wide. Source: Steven Earle (2015) CC BY 4.0 view source

Exercise: Pluton Problems

The diagram below is a cross-section through part of the crust showing a variety of intrusive igneous rocks. Indicate whether each of the plutons labelled a to e on the diagram below is a dike, a sill, a stock, or a batholith. (Note the trees for scale.)

Figure 7.25 A variety of igneous intrusions. Source: Steven Earle (2015) CC BY 4.0 view source

 References

Bally, A. W. (1989). Plate 10. Selected distribution maps, rate of accumulation maps, and lithofacies maps—Phanerozoic, North America. In A. W. Bally & A. R. Palmer (Eds.), The Geology of North America—An Overview: Volume A. Boulder: Geological Society of America.

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

7.1 Magma and How It Forms

Magma is molten rock, and in most cases, it forms from partial melting of existing rock. The chemistry of magma depends on the source rock that is melting, as well as the degree of partial melting that occurs. Magma forms by decompression melting, flux-induced melting, and heat transfer. Magmas range in composition from ultramafic to felsic. Mafic rocks are rich in iron, magnesium, and calcium, and contain approximately 50% silica. Felsic rocks are richer in silica (~70%) and have lower levels of iron, magnesium, and calcium, and higher levels of sodium and potassium than mafic rocks.

7.2 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. Bowen’s 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), by mixing in material from the surrounding rocks by partial melting, and by mixing with magmas of differing chemistry.

7.3 Classification of Igneous Rocks

Igneous rocks are classified based on their mineral composition and texture. Felsic igneous rocks have less than 20% dark minerals (ferromagnesian silicates including amphibole and/or biotite) with varying amounts of quartz, both potassium and plagioclase feldspars, and sometimes muscovite. Mafic igneous rocks have more than 50% dark minerals (primarily pyroxene) plus plagioclase feldspar. Most intrusive igneous rocks are phaneritic (individual crystals are visible unmagnified). If there were two stages of cooling (slow then fast), the texture may be porphyritic (large crystals in a matrix of smaller crystals).

7.4 Intrusive Igneous Bodies

Magma intrudes into country rock by pushing it aside or melting through it. Intrusive igneous bodies tend to be irregular (stocks and batholiths), tabular (dikes and sills), or pipe-like. Batholiths have areas of 100 km² or greater, while stocks are smaller. Sills are parallel to existing layering in the country rock, while dikes cut across layering. A pluton that intruded into cold rock is likely to have a chilled margin.

Review Questions

1. What is the significance of the term reaction in Bowen’s reaction series?

2. Why is it common for plagioclase crystals to be zoned from relatively calcium-rich in the middle to more sodium-rich toward the edge?

3. What must happen within a magma chamber for fractional crystallization to take place?

4. Explain the difference between aphanitic and phaneritic textures.

5. Name the following rocks:
(a) An extrusive rock with 40% Ca-rich plagioclase and 60% pyroxene
(b) An intrusive rock with 65% plagioclase, 25% amphibole, and 10% pyroxene
(c) An intrusive rock with 25% quartz, 20% potassium feldspar, 50% plagioclase feldspar, and minor amounts of biotite

6. What is the difference between a concordant tabular intrusion and a discordant tabular intrusion?

7. Why do dikes commonly have fine-grained margins?

8. What is the difference between a batholith and a stock?

9. Describe two ways in which batholiths intrude into existing rock.

10. Why is compositional layering a common feature of mafic plutons but not of felsic plutons?

1. As the temperature decreases minerals that formed early (e.g., olivine) may react with the remaining magma to form new minerals (e.g., pyroxene).

2. Calcium-rich plagioclase forms early on in the cooling process of a magma, but as the temperature drops, a more sodium-rich variety forms around the existing crystals.

3. Some minerals must begin to form while melt is still present. Early-forming minerals, which are typically quite dense (e.g., olivine), will sink to the bottom of the magma chamber if the magma is not too viscous, thus becoming separated from the rest of the magma. The composition of the remaining magma will be more felsic than before.

4. Phaneritic texture means that individual crystals can be distinguished by the naked eye. Aphanitic texture means that individual crystals cannot be distinguished without a microscope. The dividing line is somewhere between 0.1 and 1 mm, depending on the minerals.

5. a) basalt; b) diorite; c) granite

6. A concordant body (a sill) is parallel to any pre-existing layering (e.g., bedding or foliation) in the country rock is. A discordant body (a dike) cuts across any pre-existing layering, or has intruded in country rock without layering (e.g., in granite).

7. When the hot magma intrudes into cold country rock, the margins cool quickly and small crystals form, whereas magma that is not in contact with the cool country rock will cool more slowly, and larger crystals will form. The chilled margin is the band of small crystals along the edge of the dike.

8. A batholith has an area of 100 km² or greater, whereas a stock is smaller.

9. Batholiths (or stocks) intrude into existing rock by (a) melting through the country rock, or (b) causing the country rock to break and fall into the magma (stoping), or (c) pushing the country rock aside.

10. Compositional layering forms when early-crystallizing minerals sink toward the bottom of a magma chamber. This can only happen in non-viscous magma. Mafic magma is typically much less viscous than felsic magma.

Intrusive igneous rocks form at depths of 100s of metres to 10s of kilometres. Most metamorphic rocks are formed at depths of kilometres to 10s of kilometres. Sediments are turned into sedimentary rocks only when they are buried by other sediments to depths in excess of several 100s of metres. Weathering cannot happen until these rocks are revealed at Earth’s surface by uplift and the erosion of overlying materials. Once the rock is exposed at the surface as an outcrop, weathering begins.

The agents of mechanical weathering can be broadly classified into two groups: those that cause the outer layers of a rock to expand, and those that act like wedges to force the rock apart.

Mechanical Weathering By Expansion

Some processes at Earth’s surface can cause a thin outer layer of a rock to expand. Deeper than the thin outer layer, the rock does not expand. The difference is accommodated by a crack developing between the outer and inner layers, breaking the outer layer off in slabs (Figures 8.2 and 8.3). When layers break off a rock in slabs or sheets, it is referred to as exfoliation.

Figure 8.2 Close-up view of exfoliation of a granite dome in the Enchanted Rock State Natural Area, Texas, USA. Source: Wing-Chi Poon (2005) CC BY-SA 2.5 view source

Figure 8.3 View of exfoliation at a distance (centre of image) in granite exposed on the west side of the Coquihalla Highway north of Hope, B.C. Source: Steven Earle (2015) CC BY view source

Granite tends to exfoliate parallel to the exposed surface because it does not have planes of weakness to determine how it breaks. In contrast, sedimentary rocks tend to exfoliate along the contacts between different sedimentary layers, and metamorphic rocks tend to exfoliate parallel to aligned minerals.

Reasons Rocks Expand

A rock within the Earth has pressure exerted upon it by other rocks sitting above it. This is called confining pressure. When the overlying mass is removed by weathering, the confining pressure decreases, allowing the rock to expand. The cracking that results is sometimes called pressure-release cracking.

Heating a rock can also cause it to expand. If the rock is heated rapidly, as during a wildfire, cracks can form. If it goes through large daily temperature swings (e.g., in the desert where it is very hot during the day but cold at night), cracking can also eventually result as the rock is weakened.

Mechanical Weathering by Wedging

In wedging, a pre-existing crack in a rock is made larger by forcing it open.

Frost Wedging

Frost wedging (or ice wedging) happens when water seeps into cracks, then expands upon freezing. The expansion enlarges the cracks (Figure 8.4). The effectiveness of frost wedging depends on how often freezing and thawing occur.  Frost wedging won’t be as important 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.

Figure 8.4 A rock broken by ice wedging sits in a stream in Mount Revelstoke National Park, Canada. Rocks break apart when ice expands in pre-existing cracks. Source: Karla Panchuk (2018) CC BY 4.0.

Frost wedging is most effective in Canada’s climate, where for at least part of the year temperatures oscillate between warm and freezing. In many parts of Canada, the temperature swings between freezing at night and thawing in the day tens to hundreds of times a year. Even in warm coastal areas of southern British Columbia, 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 8.5).

Figure 8.5 An area with very effective frost wedging near Keremeos, BC. The fragments that were wedged away from the cliffs above have accumulated in a talus deposit at the base of the slope. The rocks in this area are variable in colour, which is reflected in the colours of the talus. Source: Steven Earle (2015) CC BY 4.0 view source

Salt Wedging

Salt wedging happens when saltwater seeps into rocks and then evaporates on a hot sunny day. Salt crystals grow within cracks and pores in the rock, and the growth of these crystals 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 8.6). The honeycomb structure of rounded holes, called tafoni, is related to the original roughness of the surface. Low spots collect salt water, causing the effect to be accentuated around existing holes.

Figure 8.6 Tafoni (Honeycomb weathering) in sandstone on Gabriola Island, British Columbia. The holes are caused by crystallization of salt within rock pores. Source: Steven Earle (2015) CC BY 4.0 view source

Plant and Animal Activity

The effects of plants are significant in mechanical weathering. Roots can force their way into even the tiniest cracks. They exert tremendous pressure on the rocks as they grow, widening the cracks and breaking the rock. This is called root wedging (Figure 8.7).

Figure 8.7 Root wedging along a quarry wall. Left: Rocks beneath the thick red beds have been split into sheets by tree roots. Right: A closer examination reveals that tree roots are working into vertical cracks as well. Source: Karla Panchuk (2018) CC BY 4.0

Although most 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. Humans modify vast tracts of land by excavation, and have a profound effect on accelerating mechanical weathering.

Exercise: Mechanical Weathering

What mechanical weathering processes do you think take place on this mountain in British Columbia?

Figure 8.8 Granite at the top of Siy’ám’ Smánit (also known as Stawamus Chief Mountain), near Squamish, British Columbia. Source: Steven Earle (2015) CC BY 4.0 view source

 Erosion

Mechanical weathering is greatly facilitated by erosion.  Erosion is the removal of weathering products, such as fragments of rock. This exposes more rock to weathering, accelerating the process. A good example of weathering and erosion working together is the talus shown in Figure 8.5. The rock fragments forming the talus piles were broken off the steep rock faces at the top of the cliff by ice wedging, and then removed by gravity.

Gravity does not always work alone to remove weathering products. Other agents of erosion include water in streams, ice in glaciers, and waves on coasts.

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 specific to the mineral and the environmental conditions. Some minerals, like quartz, are virtually unaffected by chemical weathering. Others, like feldspar, are easily altered.

Types of Chemical Weathering Reactions

Dissolution

Dissolution reactions produce ions, but no minerals, and are reversible if the solvent is removed. A household example would be dissolving a teaspoon of table salt (the mineral halite) in a glass of water. The halite will separate into Na⁺ and Cl^– ions. If the water in the glass is allowed to evaporate, there will not be enough water molecules to hold the Na⁺ and Cl^– ions apart, and the ions will come together again to form halite. Gypsum and anhydrite are other minerals that will dissolve in water alone.

Other minerals, such as calcite, will dissolve in acidic water. Acidic water is common in nature, because carbon dioxide (CO₂) in the atmosphere reacts with water vapour in the atmosphere, and with water on land and in the oceans to produce carbonic acid (Figure 8.9).

Figure 8.9 Calcite weathering by dissolution. Top: Carbon dioxide reacts with water to make acid. Bottom: Acid reacts with calcite and produces ions. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Modified after What-When-How. Molecules from JMSE Molecular Editor, Bienfait and Ertl (2013), with permission for CC BY-NC-SA use.

While rainwater and atmospheric CO₂ can combine to create carbonic acid, the amount of CO₂ in the air is enough to make only very weak carbonic acid. In contrast, biological processes acting in soil can result in a much higher concentration of CO₂ within soil, as well as adding organic acids. Water that percolates through the soil can become significantly more acidic.

Calcite is a major component of the sedimentary rock called limestone (typically more than 95%). In the presence of acidic groundwater, limestone can dissolve underground.  Over time the dissolution can remove enough of the calcite to form caves.

If dissolution of limestone or other materials removes enough rock to undermine support near the surface, the surface may collapse, creating a sinkhole such as the one in Figure 8.10, downstream of the Mosul Dam in Iraq.

Figure 8.10 Sinkhole downstream of the Mosul Dam in Iraq. The sinkhole is a result of dissolution of gypsum and anhydrite layers. Source: U. S. Army Corps of Engineers (2007) Public Domain view source

Although the sinkhole in Figure 8.9 might appear minor, it indicates a serious problem. The dam itself is constructed on limestone supported by beds of gypsum and anhydrite. Gypsum and anhydrite are soluble in water, and the gypsum and anhydrite beneath the dam are rapidly dissolving away. This was the case prior to construction of the dam. However, once the dam was filled, the increased water pressure began to force water through the formations much faster, accelerating dissolution. Ongoing measures to fill gaps with grout are required, or else there is a grave risk of catastrophic failure, placing nearly 1.5 million people at risk.

Hydrolysis

The term hydrolysis combines the prefix hydro, referring to water, with lysis, which is derived from a Greek word meaning to loosen or dissolve. Thus, you can think of hydrolysis as a chemical reaction where water loosens the chemical bonds within a mineral. This might sound the same as dissolution but the difference is that hydrolysis produces a different mineral in addition to ions. An example of hydrolysis is when water reacts with potassium feldspar to produce clay minerals and ions. The results can be seen by comparing weathered and unweathered surfaces of the same sample of granite (Figure 8.11). On the recently broken unweathered surface (Figure 8.11, left) feldspar is visible as bright white crystals. On a weathered surface (right) the feldspar has been altered to the chalky-looking clay mineral kaolinite.

Figure 8.11 A piece of granite with unweathered (left) and weathered (right) surfaces. On the unweathered surfaces the feldspars are still fresh and glassy looking. On the weathered surface there are chalky white patches where feldspar has been altered to the clay mineral kaolinite. Source: Karla Panchuk (2018) CC BY 4.0. Photos by Steven Earle (2015) CC-BY 4.0 view source

Silicate minerals other than feldspar can undergo hydrolysis, but with different end results. For example, pyroxene can be converted to the clay minerals chlorite or smectite. Olivine can be converted to the clay mineral serpentine.

Hydration

Hydration reactions involve water being added to the chemical structure of a mineral. An example of a hydration reaction is when anhydrite (CaSO₄) is transformed into gypsum (CaSO₄·2H₂O). A consequence of hydration is that the resulting mineral has a greater volume than the original mineral. In the case of the Mosul Dam, hydration of anhydrite has important consequences. The increase in volume applied force to an overlying limestone layer, breaking it into pieces. While unbroken limestone is a strong enough material upon which to build a foundation, broken limestone is too weak to provide a safe foundation.

Oxidation

Oxidation happens when free oxygen (i.e., oxygen not bound up in molecules with other elements) is involved in chemical reactions. Oxidation reactions provide valuable insight into Earth’s early surface conditions because there is a clear transition in the rock record from rocks containing no minerals that are products of oxidation reactions, to rocks containing abundant minerals produced by oxidation. This reflects a transition from an oxygen-free atmosphere to an oxygenated one.

In iron-rich minerals such as olivine, the oxidation reaction begins with taking iron out of the mineral and putting it into solution as an ion. Olivine reacts with carbonic acid, leaving dissolved iron, bicarbonate, and silicic acid:

Fe₂SiO₄ + 4H₂CO₃ → 2Fe^₂+ +  4HCO₃^– +  H₄SiO₄

Iron and oxygen dissolved in water react in the presence of bicarbonate to produce hematite and carbonic acid:

2Fe^₂+  + ½ O₂ + 2H₂O + 4HCO₃^–  → Fe₂O₃  + 4H₂CO₃

When the olivine in basalt is oxidized, the basalt takes on a reddish colour that is distinct from the dark grey or black of unweathered basalt (Figure 8.12).

Figure 8.12 Basalt pillows in Andalusia, Spain, with reddish weathered surfaces. Where parts of the pillows have broken away, darker unweathered basalt is visible. Source:Ignacio Benvenuty Cabral (2011) CC BY-NC-SA view source

The oxidation reaction would be similar for other iron-containing silicate minerals such as pyroxene, amphibole, and biotite. Iron in sulphide minerals such as pyrite (FeS₂) can also be oxidized in this way.

Hematite is not the only mineral that can result from oxidation. In fact, a wide range of iron oxide minerals that can form in this way, In granite, for example, biotite and amphibole can be altered to form the iron oxide and iron hydroxyoxide minerals that are referred to in combination as limonite (orange material in Figure 8.13).

Figure 8.13 Biotite and amphibole in this granite have been altered by oxidation to limonite (orange-yellow coating), which is a mixture of iron oxide and iron hydroxyoxide minerals. Source: Steven Earle (2015) CC-BY 4.0 view source

Oxidation Reactions and Acid Rock Drainage

Oxidation reactions can pose an environmental problem in areas where rocks have elevated levels of sulphide minerals such as pyrite. This is because when oxygen and water react with pyrite, sulphuric acid is produced:

2FeS₂ + 7O₂ + 2H₂O → 2FeSO₄ + 2H₂SO₄

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. In these cases the problem is referred to as acid mine drainage. One example is the Mt. Washington Mine near Courtenay on Vancouver Island (Figure 8.12), but there are many similar sites across Canada and around the world.

Figure 8.14 Acid mine drainage. Left: Mine waste where exposed rocks undergo oxidation reactions and generate acid at the Washington Mine, BC. Right: An example of acid drainage downstream from the mine site. Source: Steven Earle (2015) CC BY 4.0 view source

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 easily dissolve in water, which can be toxic to aquatic life 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.

References

Bienfait, B., & Ertl P. (2013). JSME: a free molecule editor in JavaScript. Journal of Cheminformatics 5(24). https://doi.org/10.1186/1758-2946-5-24

Weathering does not happen at the same rate in all environments. The same types of weathering do not happen in all environments. There are a variety of factors that determine what kinds of weathering will occur, and how fast the processes will proceed.

Climate

Water and temperature are key factors controlling both weathering rates and the types of weathering that occur:

• Water is required for chemical weathering reactions to occur.
• Water must be present for ice wedging to happen.
• Higher temperatures speed up chemical reactions.
• Climate will determine whether water is present mostly in liquid form, solid form (ice), or both.
• Climate will determine what plant life is available to force rocks apart with their roots, and to contribute organic acids to soils to aid in chemical weathering.

This means, for example, that chemical weathering will be faster in a tropical rainforest than in the Arctic, a cold desert. It means physical weathering will be the predominant form of weathering in the Arctic.

Oxygen and Carbon Dioxide

The presence and abundance of oxygen and carbon dioxide affect chemical weathering rates. Surface environments on Earth almost all have some free oxygen available, permitting oxidation reactions to take place. Exceptions are in settings such as deep lakes or swamps where oxygen cannot easily mix into the water, and where biological processes consume the oxygen rapidly.

Carbon dioxide, which acidifies water and contributes to chemical weathering, is more concentrated in some settings than others. For example, because of the activities of organisms, soils can have very high concentrations of carbon dioxide, whereas carbon dioxide concentrations will be lower on surfaces free of soils and exposed to the atmosphere.

Minerals

The minerals making up a rock will determine what kinds of chemical weathering reactions are possible, and how rapidly chemical weathering reactions occur. Under the same conditions, dissolution of the calcite making up limestone will occur more rapidly than hydrolysis reactions happening to feldspar in granite. Quartz is very resilient to chemical weathering, and will remain long after calcite and feldspar have been weathered away. A rock with grains cemented by calcite will weather faster than a rock with grains cemented by quartz.

In general, differences in the rates of chemical weathering among minerals can be broken down as follows:

• Minerals that weather by dissolution (e.g., halite, gypsum, calcite) are the easiest to weather.
• Silicate minerals with lower silica to oxygen ratios (e.g., silicates made of isolated silica tetrahedra or single chains) are easier to weather than silicate minerals with higher ratios (e.g., those made of silica tetrahedra arranged sheets or frameworks).
• Minerals that are by-products of chemical weathering are some of the most resistant to further chemical weathering, although they may be more prone to physical weathering (e.g., clay minerals).

Weathering Makes Weathering Go Faster

Weathering accelerates weathering. Physical weathering forms cracks and breaks rocks apart into smaller pieces. The smaller the pieces, the greater the surface area exposed to chemical weathering. When the newly exposed surfaces are exposed to chemical weathering, it weakens the rock even further, making it more susceptible to physical weathering processes.

Differential Weathering

When rocks in an outcrop weather at different rates, the result is called differential weathering. Differential weathering causes some beds in an outcrop to be recessed relative to the others, because beds that are slow to weather will take longer to recede than weaker beds (Figure 8.15).

Figure 8.15 Differential weathering in an outcrop along the Blaeberry River near Golden BC. The recessed beds within the outcrop are weathering faster than the surrounding beds. Source: Karla Panchuk (2009) CC BY 4.0

The visible products of weathering and erosion are the unconsolidated materials that we find around us on slopes, beneath glaciers, in stream valleys, on beaches, and in deserts. The loose collection of material is referred to as sediment, and the individual pieces that make it up are called clasts.  Clasts can be of any size: sand-sized and smaller (in which case they might be referred to as particles or grains), or larger than a house.

Clasts can range widely in size and shape (Figure 8.16) depending on the processes involved in making and transporting them. If and when deposits like these are turned into sedimentary rocks, the mineralogy and textures of these rocks will vary significantly. Importantly, when we describe sedimentary rocks that formed millions of years in the past, we can study the mineralogy and textures to make inferences about the conditions that existed during the deposition of the sediment, and the later burial and formation of sedimentary rock.  The properties we look at are composition, grain size, sorting, rounding, and sphericity.

Figure 8.16 Products of weathering and erosion formed under different conditions. Source: Steven Earle (2016) CC BY 4.0 view source

Composition

Composition refers to the mineral or minerals making up the clast. Small clasts might be single mineral grains, but larger ones can have several different mineral grains, or even several different pieces of rock within them.  The composition can tell us about what rock the sediments came from, and about the geological setting from which the sediment was derived.

Not all minerals have the same hardness and resistance to weathering, so as weathering and erosion proceed, some minerals become more abundant than others within sediments. Quartz is one example of a mineral that is more abundant. It is highly resistant to weathering by weak acids or reaction with oxygen. This makes it unique among the minerals that are common in igneous rocks. Quartz is also very hard, so it is resistant to mechanical weathering.

In contrast, feldspar and iron- and magnesium-bearing minerals are not as resistant to weathering. As weathering proceeds, they are likely to be broken into small pieces and converted into clay minerals and dissolved ions. Ultimately this means that quartz, clay minerals, iron oxides, aluminum oxides, and dissolved ions are the most common products of weathering.

Grain Size

Whether a grain is large or small tells us about its journey from its source to where it was deposited.  Mechanical weathering can break off large pieces from rock.  Large pieces carried along by streams will bump into each other, causing smaller pieces to break off.  Over time the grains get smaller and smaller still. If we find grains that are very small, we can conclude that they travelled over a long distance.

Geologists have a specific set of definitions to describe the size of grains (Figure 8.17).

Figure 8.17 Classification of grain sizes. Silt and clay are considered fine-grained particles, sand is medium-grained, and particles larger than sand are considered coarse-grained.. Source: Karla Panchuk (2016) CC BY 4.0 Click the image for a text version.

The scale has some of the grain sizes listed in microns (µm). There are 1000 µm in 1 mm. The particles classified as sand are what you would intuitively think of as being sand-sized, so an easy way to remember the scale is that anything smaller than sand is fine-grained, and anything larger is coarse-grained. Fine sand grains are still easily discernible with the naked eye. Silt grains are barely discernible in rocks, and silty rocks feel gritty when rubbed. Clay grains are invisible to the naked eye, and rocks comprised of clay feel smooth when rubbed.

One other thing to notice about this scale is that the finest-grained particle is referred to as clay. While a clay-sized particle could be composed of clay minerals (and often they are), it doesn’t have to be. Any particle of that size would be referred to as clay.

Grain Size and Transportation

The grain size of sediments is not just for purposes of description. It’s also a valuable clue to the processes that have acted on those sediments, because the size of the clast determines how much energy is required to move it.

Whether or not a medium such as water or air has the ability to move a clast of a particular size and keep it moving depends on the velocity of the flow.  For the most part, the faster the medium flows, the larger the clasts that can be moved. Figure 8.18 shows a stream bed that now contains only a trickle of water—barely enough to move particles of sand or cool puppy feet. But the velocity of water in the stream changes from season to season, as does the volume of water.  All of the clasts in the stream bed were transported there by water at some point.

Figure 8.18 Ruby looks upstream in a channel near Golden BC. For much of the year the only water in the stream is the trickle in which Ruby stands, but in the spring the water can flow rapidly enough to carry boulders. Source: Karla Panchuk (2009) CC BY 4.0

Very fine-grained particles are the exception to rule that the larger the clasts, the faster the water that is required to transport them. Clay and silt grains stick together, requiring higher water velocities to pick them up and move them than some larger particles. Water that flows fast enough to pick up sand would not be fast enough to pick up clay.

Sorting

Weathering can break off large fragments of rock, and erosion and transport can break these fragments down to smaller and smaller sizes. The extent to which the grains in sediment differ in size is described by sorting (Figure 8.19, top).

Figure 8.19 Top: Sorting of grains, ranging from well sorted where the grains are similar in size, to poorly sorted, where the grains vary greatly in size. Bottom: Rounding refers to how smooth or rough the edges of a clast are.  Clasts with sharp edges and corners are angular. Clasts with smooth surfaces are rounded.  Clasts that fall in between are sub-angular or sub-rounded. Source: Reagan et al. (2015) CC BY 3.0 view source

If the grains in a sample of sediment are the same size or very nearly so, the sediment is said to be well sorted.  If the grains vary substantially in size, the sediment is poorly sorted. Because grains become progressively smaller as they are transported, sorting improves the further the sediments are from their source.

Rounding

Rounding refers to whether clasts have sharp edges and corners or not (Figure 8.19, bottom). If the grains are rough, with lots of edges and corners, then they are referred to as angular. Grains with smooth surfaces are rounded. Grains in between can be sub-angular or sub-rounded. The farther sediments are transported, the rounder they become.

Sphericity

Sphericity describes whether a grain is elongate or not.  Grains that are longer than they are wide (like an ellipse) have low sphericity, whereas grains that have the same diameter no matter where you measure it (like a sphere) have high sphericity. In the bottom row of boxes in Figure 8.19 the grains at the top of each box exhibit high sphericity, and the grains at the bottom exhibit low sphericity.  Notice that a grain can be angular but still have high sphericity.  It can be rounded, but still have low sphericity. Sphericity also increases the further the sediments are from their source.

Exercise: Looking at Sand

Three samples of sand are shown below. Read the descriptions of what they contain and where they are from, then describe each sample in terms of grain size, sorting, rounding, and sphericity.

Figure 8.20 Three examples of sand grains. Source: Steven Earle (2016) CC BY 4.0 View sources: Sample A, Sample B, Sample C.

 

Sample A: Fragments of red coral, algae plates, and urchin needles from a shallow water area (~2 m depth) near a reef in Belize. The grains are between 0.1 and 1 mm across.

 

Sample B: Quartz and rock fragments from a glacial stream deposit near Osoyoos, BC. The grains are between 0.25 and 0.5 mm in diameter.

 

Sample C: Grains of olivine (green) and volcanic glass (black) from a beach on the big island of Hawai’i. The grains are approximately 1 mm in diameter.

References

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 the soil!

Many people refer to any loose material on Earth’s surface as soil, but to scientists soil is the material that includes organic matter, forms within the top few tens of centimetres of the surface, and is important for sustaining plant growth.

Soil is a complex mixture of minerals (~45%), organic matter (~5%), and empty space (~50%, filled to varying degrees with air and water). The mineral content of soil varies, 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 (Figure 8.21). Sand and silt components are dominated by quartz, with lesser amounts of feldspar and rock fragments. The clay component is dominated by clay minerals.

Figure 8.21 Soil texture classification diagram. Textures are determined by the proportions of sand-, silt-, and clay-sized grains. Source: Mike Norton (2011) CC BY-SA 3.0 view source

Factors Affecting How Soil Forms

Soil forms through the mechanical and chemical weathering of rocks and sediments, and the accumulation and decay of organic matter. 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 of vegetation
• The parent rock or sediment that was weathered to make the soil
• The slope of the surface where soil is accumulating
• How long soil has been forming at a location

Climate

Both the mechanical breakup of rocks and the chemical weathering of minerals contribute to soil formation.  The downward percolation of water brings dissolved ions and also facilitates chemical reactions. Soil forms most readily under temperate to tropical conditions, and moderate precipitation.  Temperature matters because chemical weathering reactions and those facilitated by organisms proceed fastest under warm conditions, and plant growth is enhanced in warm climates. Where the climate is cooler, the rates of chemical weathering reactions decrease, and when water is frozen, may cease entirely.

Although water is needed for chemical weathering to take place, too much water can lead to soils that lack nutrients. In rain forests, for example, high rainfall contributes so much water that important nutrients are leached away, and acidic soils are left behind. In humid and poorly drained regions, swampy conditions may prevail, producing soil that is dominated by organic matter, but low in inorganic nutrients.

Too little water (e.g., in deserts and semi-deserts) limits the rate of downward chemical transport, and it also means that salts and carbonate ions dissolved in upward-moving groundwater can precipitate and build up in sediments, hindering organic activity. These soils also lack organic matter (Figure 8.22).

Figure 8.22 Soil consisting of wind-blown silt (loess) and little organic matter in an arid part of north-eastern Washington state. Source: Steven Earle (2016) CC BY 4.0 view source

Parent Material

Parent material for soils can be any type of bedrock, and any type of unconsolidated sediment, 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. This doesn’t mean that the soils themselves have been transported, but that the soil developed on unconsolidated material rather than on bedrock.

Sandy soils develop from quartz-rich parent material, such as granite, sandstone, or loose sand. 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, which is a source of the important soil nutrient phosphorus. Basaltic parent material tends to generate very fertile soils because, in addition to phosphorus, it provides significant amounts of iron, magnesium, and calcium.  The iron, magnesium, and calcium come from minerals such as olivine ((Mg,Fe)₂SiO₄) and plagioclase feldspar (CaAl₂Si₂O₈) in the basalt.

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.

Slope

Soil can only develop where surface materials remain in place and are not frequently washed away or lost to mass wasting (landslides). Soils cannot develop where the rate of soil formation is lower than the rate of erosion, so steep slopes tend to have little or no soil.

Time

Even under ideal conditions, soil takes thousands of years to develop. Virtually all of southern Canada was covered with glaciers up until 14,000 years ago, and most of the central and northern parts of BC, the prairies, Ontario, and Quebec were still glaciated at 12,000 years ago. Glaciers remained in the central and northern parts of Canada until around 10,000 years ago, so conditions were still not ideal for soil development even in the southern regions. This means that soils in Canada, particularly 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, in areas of mass wasting, or where an area has been resurfaced by volcanic deposits.

Because soil takes so long to form, human activities that damage soils have long-term consequences for ecosystems, and for the utility of the soil for food production.

Soil Horizons

When soils form, the downward movement of clay, water, and dissolved ions can lead to the development of chemically and texturally distinct layers known as soil horizons. In temperate climates, common soil horizons that develop are the following (Figure 8.23):

• O horizon— A layer of organic matter
• A horizon— Partially decayed organic matter mixed with mineral material
• E horizon— The eluviated (leached) layer from which some of the clay and iron have been removed to create a pale layer that may be sandier than the other layers
• B horizon— Where clay, iron, and other elements from the overlying soil accumulate
• C horizon— Contains broken fragments of rock. Weathering of the underlying bedrock or sediments is not yet complete.

Figure 8.23 Typical horizons in a temperate soil, from Wales. Source: Karla Panchuk (2018) CC BY 4.0. Photograph: Richard Hartnup (2005) Public Domain view source

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.

How Soil Is Lost

Like all geological materials, soil is subject to erosion.  Under natural conditions on gentle slopes, the rate of soil formation either balances or exceeds the rate of erosion. However, human practices 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, this protection is lost. When soil is not protected, wind and water can easily erode it away.

Water erosion is accentuated on sloped surfaces because fast-flowing water has greater eroding power than still water. Raindrops can disaggregate exposed soil particles, putting clay into suspension in the water. Sheetwash—unchannelled flow across a surface—carries suspended material away, and channels erode right through the soil layer, removing both fine and coarse material (Figure 8.24).

Figure 8.24 Soil erosion by rain and unchanneled runoff in a field in Alberta. Source:Alberta Agriculture and Rural Development. Click the image for source information and terms of use.

Wind erosion is exacerbated by the removal of trees that act as wind breaks, and by agricultural practices that leave bare soil exposed (Figure 8.25).

Figure 8.25 Soil erosion by wind in Alberta. Source:Alberta Agriculture and Rural Development. Click the image for source information and terms of use.

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.

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 (Table 8.1), and you can explore the distribution of soils using Agriculture and Agri-Food Canada’s interactive map (Figure 8.26). See the resources section at the bottom of the page for additional sources of information on Canadian soils, including videos.

Figure 8.26 Distribution of soil orders in Canada. Click here to go to the interactive map. Source: Agrifood and Agriculture Canada. Contains information licensed under the Open Government License – Canada. Click the image for terms of use.

Processes of soil formation include downward transport of solid and dissolved materials, and the nature of those processes depends in large part on the climate. In Canada’s predominantly cool and humid climate—characteristic of most places other than the far north— podzolization is the norm. This involves downward transportation of hydrogen, iron, and aluminum from the upper part of the soil profile, and accumulation of clay, iron, and aluminum in the B-horizon. Most of the podzols, luvisols, and brunisols of Canada form through various types of podzolization.

In the grasslands of the dry southern parts of the prairie provinces and in some of the drier parts of southern BC, dark brown organic-rich chernozem soils are dominant. In some cases, weak calcification takes place when calcium is leached from the upper layers and accumulates in the B-horizon. Development of caliche layers is rare in Canada.

Organic soils form in areas with poor drainage and a rich supply of organic matter, such as in swamps. 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 slow. The soils are called cryosols (the cryo prefix is used to indicate extreme cold). In permafrost areas, the freeze-thaw process churns the soil, resulting in limited soil horizon development.

Resources

Soils of Canada (University of Saskatchewan)

Soil Classification: Soil Orders of Canada Watch videos about each soil order.

Carbon cycling on Earth operates on different timescales depending on the components of the Earth system that are involved. Over the short term, biological processes are important. In particular, living organisms — mostly plants — consume carbon dioxide from the atmosphere to make their tissues.  After they die, the carbon is released back into the atmosphere over years to decades as the plant matter decays.

Over the longer term, geological processes drive the carbon cycle. Geological carbon-cycle processes operate very slowly, but they affect much more of Earth’s carbon than the biological component. Carbon can move from the biological cycle to the geological cycle if it is buried in sedimentary rocks. The biological carbon could be fragments of plant material or organic molecules that are preserved as coal or in organic-rich shale. It could also be calcium carbonate body parts of marine organisms that are preserved in limestone.

The geological component of the carbon cycle is shown in Figure 8.27. The various steps in the process (not necessarily in this order) are as follows:

Figure 8.27 The geological component of the carbon cycle includes: (a) organic carbon in peat, coal and permafrost, (b) weathering of silicate minerals converts atmospheric carbon dioxide to dissolved bicarbonate, (c) marine organisms convert dissolved carbon to calcium carbonate, (d) carbon compounds are stored in sediments, (e) carbon-bearing sediments are transferred to longer-term storage in the mantle, and (f) carbon dioxide is released back to atmosphere during volcanic eruptions. Source: Steven Earle (2016) CC BY 4.0 view source

At some times in Earth’s history, the geological carbon cycle has been balanced, with carbon being released to the atmosphere by some processes at approximately the same rate as other processes store it. Under these conditions, the climate can remain relatively stable.

At other times, the balance is upset. Prolonged periods of greater than average volcanism can cause an imbalance. The eruption of the Siberian Traps at around 250 Ma warmed the climate significantly over a few million years, leading to a mass extinction.

Mountain-building events may also cause an imbalance. The formation of the Himalaya range between about 40 Ma and 10 Ma ago exposed rocks to weathering over a large region. The over-all rate of weathering on Earth increased because the mountains were so high, and the range was so extensive. The weathering of these rocks — most importantly the hydrolysis of feldspar — consumed atmospheric carbon dioxide and transferred carbon to the oceans and to ocean-floor carbonate minerals. Decreasing carbon dioxide levels contributed to climate cooling that culminated in the Pleistocene glaciations.

Today, burning fossil fuels is causing an imbalance in the carbon cycle. Burning coal, oil, and gas releases in a geological instant carbon that was stored by the biological carbon cycle over hundreds of millions of years. Scientists who study Earth’s past climate tell us that today carbon dioxide is being added to the atmosphere faster than during some of the most extreme climate change events in Earth history. Eventually, higher carbon dioxide levels will accelerate chemical weathering, and that will help to remove some of the carbon dioxide from the atmosphere. However, weathering is part of the geological carbon cycle, and operates over long timescales. If humans stopped burning all fossil fuels today, it could still take thousands of years for balance to be restored.

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

8.1 Mechanical Weathering

Rocks weather when they are exposed to surface conditions. In most cases, conditions at Earth’s surface are very different from the conditions under which the rocks formed. Mechanical weathering processes include exfoliation, freeze-thaw, salt crystallization, and the wedging effects of plant growth.

8.2 Chemical Weathering

Chemical weathering takes place when minerals within rocks are not chemically stable in their existing environment. Chemical weathering processes include 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.

8.3 Controls on Weathering Processes and Rates

Chemical weathering is faster when temperatures are warmer and moisture is present. Physical weathering is more important in regions with frequent freeze-thaw cycles. Weathering rates can depend on the abundance oxygen and carbon, and will vary with the mineral composition of a rock. Weathering accelerates weathering by exposing more surface area to chemical reactions.

8.4 Weathering and Erosion Produce Sediments

Quartz grains are one of main products of weathering and erosion because quartz is resistant to chemical and mechanical weathering. Clay minerals, iron oxide and iron hydroxide minerals, aluminum hydroxide minerals, and ions in solution are common products of chemical weathering. Particles produced by weathering can be described in terms of their composition, grain size, sorting, rounding, and sphericity.

8.5 Weathering and Soil Formation

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.

8.6 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.

8.7 Weathering and Climate Change

The geological component of the carbon cycle affects Earth’s climate over the long term by changing atmospheric carbon dioxide levels. Carbon is added to the atmosphere during volcanic eruptions. It is extracted from the atmosphere when silicate minerals are weathered, and when it is transformed into organic matter by plants. Organic matter can be stored in soil, permafrost, and rocks. Burning of fossil fuels involves moving carbon from geological reservoirs to the atmosphere on timescales much faster than the geological carbon cycle operates.

Review Questions

1. What must happen to a body of rock before exfoliation can occur?
2. Saskatchewan’s climate is consistently cold in the winter and consistently warm in the summer. What times of year would frost wedging to be an important weathering mechanism?
3. What are the products of the hydrolysis of the feldspar albite (NaAlSi₃O₈)?
4. Oxidation weathering of  pyrite (FeS₂) can lead to acid rock drainage (ARD). What are the environmental impacts of ARD?
5. Imagine that you and a friend encounter an old graveyard on a walk through the forest. You see a granite tombstone with the date 1705 carved into it. The tombstone next to it is too badly weathered to read the date. Your friend looks at the badly weathered stone and declares, “Look how weathered this stone is! It must be from way before 1705.” Do you agree?
6. Many sand deposits are dominated by quartz, with very little feldspar. What weathering and erosion conditions are required to get feldspar-rich sand?
7. What ultimately happens to most of the clay that forms during the hydrolysis of silicate minerals?
8. Why are the slope and the parent materials important factors in soil formation?
9. Which soil constituents move downward to produce the B-horizon of a soil?
10. What are the main processes that lead to the erosion of soils in Canada?
11. Where in Canada would you expect to find a chernozemic soil? What characteristics of this region produce this type of soil?
12. Where are brunisolic soils found in Saskatchewan?
13. 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?

1. Rock must be exposed at surface. It has to be uplifted from where it formed deep in the crust, and the material on top has to be eroded.
2. Frost wedging is most effective when temperatures swing between freezing and thawing from day to day. In Saskatchewan that happens consistently in the early spring and late fall.
3. The feldspar albite (NaAlSi₃O₈) will be converted to a clay (such as kaolinite) and sodium ions in solution.
4. Acid rock drainage (ARD) creates acidic stream runoff. It also increases the solubility of a wide range of metals, some of which are toxic to wildlife and humans.
5. If the stones are both granite, then it would be reasonable to conclude that the badly weathered tombstone is much older than the other, because it has been exposed to weathering for much longer. On the other hand, if the badly weathered stone is a rock that is less resistant to weathering, like limestone, then the badly weathered stone could be the same age, or even younger than the granite one.
6. Feldspar-rich sand is formed where granitic rocks are being weathered and where mechanical weathering predominates over chemical weathering. For a deposit of feldspar-rich sand to be preserved, the sand must be deposited close to its source to limit the opportunities for chemical weathering.
7. Most of the clay that forms during hydrolysis of silicate minerals ends up in rivers and is washed out to the oceans. There it eventually settles to the sea floor.
8. On a steep slope, gravity will remove materials, making it unlikely for soils to accumulate. The mineral composition of the parent rock or sediment will influence the composition of the resulting soil.
9. Clay minerals and iron move downward to produce the B horizon of a soil.
10. Wind and water are the main processes of soil erosion in Canada. Removal of vegetation makes it easier for erosion to happen.
11. Chernozemic soils are common in the southern prairies, where organic matter from grasslands is added to soils
12. Brunisolic soils are found in the northern half of Saskatchewan where forest cover is common.
13. The weathering of feldspar to clay involves the conversion of atmospheric carbon dioxide to dissolved bicarbonate, which ends up in the ocean.

How Clastic Sediments Become Sedimentary Rocks

Lithification (Figure 9.3) is the process of converting sediments into solid rock. Compaction is the first step. Sediments that have been deposited are buried when more and more sediments accumulate above them. The weight of the overlying sediments pushes the clasts together, closing up some of the pore spaces (the gaps between grains) and forcing them together. Pore spaces often contain water (although they can also contain air or even hydrocarbons), so the water is squeezed out.

Figure 9.3 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

Cementation is the next step. Groundwater flowing through the remaining pore spaces contains ions, and these ions may precipitate, leaving behind minerals in the pore spaces. These minerals bind the grains together, and are referred to collectively as cement. Quartz and calcite are common cement minerals, but depending on pressure, temperature, and chemical conditions, cement might also include other minerals such as hematite and clay.

Figure 9.4 shows sandstone viewed under a microscope. The grains are all quartz but they appear different shades of grey because they are being viewed through cross-polarized light. It is difficult to tell the grains from the cement in this case because both are made of quartz, but in the image on the right the more obvious grain boundaries are marked with dashed lines.  Some of the cement is marked with blue shading. Using the image on the right, see if you can pick out the grain boundaries in the image on the left.

Figure 9.4 Sandstone under a microscope. Grains and cement are quartz. Left- Original image. Right- Visible grain boundaries are marked with dashed lines, and some of the cement is shaded in blue. Source: Karla Panchuk (2018) CC BY 4.0 modified after Woudloper, Public Domain view source

Types of Clastic Sedimentary Rocks

Clastic sedimentary rocks are named according to the characteristics of clasts (rock and mineral fragments) that comprise them. These characteristics include grain size, shape, and sorting. The different types of clastic sedimentary rocks are summarized in Figure 9.5.

Figure 9.5 Types of clastic sedimentary rocks. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0, Photos by James St. John and  R. Weller/ Cochise College. Click the image for more attributions.

Coarse-Grained Clastic Rocks

Clastic sedimentary rocks in which a significant proportion of the clasts are larger than 2 mm are known as conglomerate if the clasts are well rounded, or breccia if they are angular (Figure 9.5, top row). Conglomerates form in high-energy environments, such as fast-flowing rivers, where the particles can become rounded as they bump into each other while being carried along.  Breccias typically form where the particles are not transported a significant distance, such as in alluvial fans and talus slopes.

Medium-Grained Clastic Rocks

Sandstone (Figure 9.5, middle row) is a very common sedimentary rock, and there are many different kinds of sandstone. It is worth knowing something about the different types because they are organized according to characteristics that are useful for the detective work of figuring out what conditions led to the formation of a particular sandstone. Broadly, sandstones can be divided into two groups: arenite and wacke (rhymes with tacky).

Arenite is “clean” sandstone consisting mostly of sand-sized grains and cement, with less than 15% of fine-grained silt and clay in the matrix (the material between the sand-sized grains). Arenites are subdivided according to what the sand-sized grains are made of (Figure 9.6). If 90% or more of the grains are quartz, then the sandstone is called a quartz arenite (also called a quartz sandstone). If more than 10% of the grains are feldspar and more of the grains are feldspar than fragments of other rocks (lithic“Lithic” means “rock.” Lithic clasts are rock fragments (multimineralic fragments), as opposed to single-mineral fragments. fragments) then the sandstone is called an arkosic arenite, or just arkose. If the rock has more than 10% rock fragments, and more rock fragments than feldspar, it is lithic arenite.

Figure 9.6 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. Source: Steven Earle (2015) CC BY 4.0 view source

Wacke is a “dirty” sandstone, containing 15-75% fine-grained particles (clay, silt) in its matrix.  A wacke can have more fine-grained particles than cement in its matrix, making for a crumbly rock.  Wackes are subdivided in the same way that arenites are: quartz wacke, feldspathic wacke, and lithic wacke. Another name for a lithic wacke is greywacke.

Figure 9.7 shows thin sectionsThin sections are slivers of rock sliced thinly enough so that light can pass through them, and they can be examined under a microscope. (microscopic views) of quartz arenite, arkose, and lithic wacke. In the images, quartz grains are marked Q, feldspar grains are marked F, and lithic fragments are marked L. Notice the relative abundances of each component in the three types of rocks.

Figure 9.7 Photos of thin sections of three types of sandstone. 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. Source: Steven Earle (2016) CC BY 4.0 view source

Fine-Grained Clastic Rocks

Rock composed of at least 75% silt- and clay-sized clasts is called mudrock (Figure 9.5, bottom row). If a mudrock shows evidence of fine layers (laminations) and breaks into sheets, it is called shale. Otherwise, it is siltstone (dominated by silt), mudstone (a mix of silt and clay), or claystone (dominated by clay). The fine-grained nature of mudrocks tells us that they form in very low energy environments, such as lakes, flood plains, and the deep ocean.

Exercise: Classifying Sandstones

Use Figures 9.6 and 9.7 to give the appropriate name to the sandstone in each of the magnified thin sections shown below.

Figure 9.8a Sandstone 1. Source: Steven Earle (2015) CC BY 4.0 view source	Figure 9.8b Sandstone 2. Source: Steven Earle (2015) CC BY 4.0 view source
Sandstone 1. Rounded sand-sized grains are approximately 99% quartz and 1% feldspar. Silt and clay make up less than 2% of the rock.	Sandstone 2. Angular sand-sized grains are approximately 70% quartz, 20% lithic fragments, and 10% feldspar. Silt and clay make up ~20% of the rock.

Clastic sediments are deposited in a wide range of environments, including from melting glaciers, slope failures, rivers (both fast and slow flowing), lakes, deltas, and ocean environments (both shallow and deep). Depending on the grain size in particular, they may eventually form into rocks ranging from mudstone to breccia and conglomerate. By examining clastic sedimentary rocks it is possible to translate the classification you have just learned into an interpretation of the environment in which the rocks were deposited.

Sediment Maturity

Maturity in sediments refers to the extent to which sediment characteristics reflect prolonged weathering and transport. Prolonged weathering and transport cause clasts to become smaller, rounder, and more well-sorted. It removes minerals that are more susceptible to weathering, such as feldspar, leaving a sediment consisting predominantly of quartz or clay. On the spectrum of sediment maturity, quartz sandstone or shale would be mature sedimentary rocks, and wacke or conglomerate would be an immature rocks.

Clastic sedimentary rocks are dominated by components that have been transported as solid clasts (clay, silt, sand, etc.). In contrast, chemical and biochemical sedimentary rocks are dominated by components that have been transported as ions in solution (e.g., Na⁺, Ca²⁺, HCO_3^–, 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. The difference between chemical and biochemical sedimentary rocks is that in biochemical sedimentary rocks, organisms play a role in turning the ions into sediment. This means the presence and nature of biochemical sedimentary rocks are linked to the life requirements of the organisms that comprise them. In chemical sedimentary rocks, the process is inorganic, often resulting from a body of water evaporating and concentrating the ions.  It is possible for one type of sedimentary rock to form from both chemical (inorganic) and biochemical (organically mediated) processes.

Chemical and biochemical sedimentary rocks are classified based on the minerals they contain, and are frequently dominated by a single mineral. It is true that some clastic sedimentary rocks, such as quartz arenite, can also be dominated by a single mineral, but the reasons for this are different. A clastic sedimentary rock can contain whatever minerals were present in the parent rock.  The minerals the clastic rock ends up containing will depend on how much “processing” the sediments undergo by physical and chemical weathering, and transport, before the sediment was cemented. On the other hand, chemical sedimentary rocks are limited largely to those minerals that are highly soluble in water.  Because mineral content is a defining characteristic of chemical and biochemical sedimentary rocks, we will use it to organize our discussion of these rocks.

Carbonate Rocks

Carbonate rocks are those in which the dominant mineral contains the carbonate anion (CO₃^2-).  The main carbonate minerals are calcite and aragonite. Both minerals have the formula CaCO₃ but they have different crystal structures.  A less common carbonate mineral that is still important for forming carbonate rocks is dolomite, which has the formula CaMg(CO₃)₂. It is similar to calcite and aragonite, except that some of the calcium is replaced with magnesium. Dolomite is more common as a replacement mineral, which has replaced calcite in carbonate rocks.

Limestone

Limestone is comprised of calcite and aragonite. It can occur as a chemical sedimentary rock, forming inorganically due to precipitation, but most limestone is biochemical in origin.  In fact, limestone is by far the most common biochemical sedimentary rock.

Almost all limestone forms in marine (i.e., oceans or salty seas) environments, and most of that forms on the shallow continental shelves, especially in tropical regions with coral reefs. Today continental shelves are relatively narrow zones along the margins of continents, but for large parts of geologic history sea-level was much higher, and large parts of the interiors of continents were flooded.

Reefs are highly productive ecosystems populated by a wide range of organisms, many of which use calcium and bicarbonate ions from seawater to make carbonate minerals (especially calcite) for their shells and other structures. These include corals as well as green and red algae, urchins, sponges, molluscs, and crustaceans. Some of micro-organisms use CaCO₃ to build tiny tests (shells) which accumulate on the ocean floor when these organisms die. Erosion can break all of these carbonate materials apart, scattering fragments throughout surrounding region (Figure 9.9).

Figure 9.9 Various corals and green algae on a reef at Ambergris, Belize. The light-coloured sand consists of carbonate fragments eroded from the reef organisms. Source: Steven Earle (2015) CC BY 4.0 view source

Figure 9.10 shows a cross-section through a typical reef environment in a tropical region (normally between 40° N and 40° S). Reefs tend to form 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. Reef-derived sediments are dominated by reef-type carbonate fragments of all sizes, including mud.

Figure 9.10 Cross-section through a typical tropical reef. Source: Steven Earle (2015) CC BY 4.0 view source

In many such areas, carbonate-rich sediments also accumulate in quiet lagoons, where mud and mollusc-shell fragments predominate (Figure 9.11, left) or in offshore areas with strong currents, where either foraminifera tests accumulate (Figure 9.11, middle) or calcite crystallizes inorganically to form ooids – spheres of calcite that form in shallow tropical ocean water with strong currents (Figure 9.11, right).

Figure 9.11 Carbonate rocks and sediments. Left- Mollusc-rich limestone formed in a lagoon area at Ambergris, Belize. Middle- Foraminifera-rich sediment from a submerged carbonate sandbar near to Ambergris, Belize . Right- Ooids from a beach at Joulters Cay, Bahamas. Sources: Left, Middle- Steven Earle (2015) CC BY 4.0 view source; Right- Wilson44691 (2010) Public Domain view source

Limestone also accumulates in deeper water, from the steady settling out of the carbonate shells of tiny organisms that lived near the ocean surface. Processes on the ocean floor cause the water in the deepest parts of the ocean to become more acidic. This puts a lower limit on how deep in the ocean calcite and aragonite can accumulate, because they dissolve under acidic conditions.

Tufa and Travertine

Calcite can form chemical sedimentary rocks on land in a number of environments. Tufa forms at springs. The tufa towers in Figure 9.12 formed where spring water discharged into lake water.

Figure 9.12 Tufa towers (made of calcium carbonate) in Mono Lake, California. Evaporation keeps the concentration of ions in the lake very high, allowing the calcium carbonate to precipitate. Source: Brocken Inaglory (2006) CC BY-SA 3.0 view source

Travertine (which is less porous) forms at hot springs. Similar material precipitates within limestone caves to form speleothems (mineral deposits in caves, Figure 9.13) such as stalactites and stalagmites.

Figure 9.13 Speleothems in Cave Nefza in Tunisia Source: Badreddine Besbes (2015) CC BY-SA 3.0 view source

Dolostone

Dolostone (also referred to as dolomite) is the carbonate rock made of the mineral dolomite (CaMg(CO₃)₂). Dolostone is quite common (there’s a whole Italian mountain range named after it), which is surprising because marine organisms do not precipitate dolomite. Dolomite forms through dolomitization, a process thought to involve chemical reactions between magnesium-rich water percolating through rocks, and sediments containing calcite.

Calcite and dolomite can be distinguished from one another by applying a drop of weak acid to the rock; calcite will react with weak acid, whereas dolomite will not. Also, when dolomite weathers, it tends to turn buff (tan) in colour, whereas calcite tends toward grey and white.

Chert

Chert is made of silica (SiO₂). It has the same chemical formula as quartz, but is cryptocrystalline, meaning that the quartz crystals comprising chert are so small that it is difficult to see them even under a microscope.  Chert can be a chemical sedimentary rock, often forming as beds within limestone (Figure 9.14), or as irregular lenses or blobs (nodules). It can also be biochemical. Some tiny marine organisms (e.g., diatoms and radiolaria) make their tests from silica. When they die their tiny shells (or tests) settle slowly to the bottom of the lake or ocean, where they accumulate and are transformed into chert.

Figure 9.14 Chert (brown layers) interbedded with limestone, Triassic Quatsino Fm, Quadra Island, BC. All of the layers have been folded, and the chert, being more resistant to weathering than limestone, stands out. Source: Steven Earle (2015) CC BY 4.0 view source

Banded Iron Formations (BIFs)

Some ancient chert beds — most dating to between 1800 and 2400 Ma — are also part of a rock known as a  banded iron formation (BIF). It is a deep sea-floor deposit of iron oxide that is a common ore of iron. These rocks consist of alternating layers of dark iron oxide minerals (magnetite and hematite) and chert stained red by hematite (Figure 9.15).

Figure 9.15 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. Source: Andre Karwath CC BY-SA (2005) view source

Figure 9.16 Spotted Lake, near Osoyoos, BC. 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. Source: Steven Earle (2015) CC BY 4.0 view source

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 as potash at numerous locations across Saskatchewan from evaporites that formed during the Devonian (~385 Ma) when an inland sea occupied much of the region.

 

Exercise: Making Evaporite

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. It may look a little like the Figure 9.17. These crystals are up to ~3 mm in diameter.

Figure 9.17 Salt crystals up to ~ 3 mm across. Source: Steven Earle (2015) CC BY 4.0 view source

 

Organic sedimentary rocks are those containing large quantities of organic molecules. Organic molecules contain carbon, but in this context we are referring specifically to molecules with carbon-hydrogen bonds, such as materials from the soft tissues of plants and animals. In other words, the carbon in calcite- CaCO₃ wouldn’t make calcite an organic mineral because it isn’t bonded to hydrogen.

An important organic sedimentary rock is coal. Most coal forms in swampy land adjacent to rivers and within deltas, and where climates are humid and tropical to temperate. The vigorous growth of vegetation leads to an abundance of organic matter that accumulates within stagnant, acidic water. This limits decay and oxidation of the organic material. If this situation—where the dead organic matter is submerged in oxygen-poor water—is maintained for centuries to millennia, a thick layer of material can accumulate. Limited decay will transform this layer into peat (Figure 9.18a, Figure 9.19 upper left).

Figure 9.18 Formation of coal. (a) Accumulation of organic matter within a swampy area forms a layer of peat; (b) The organic matter is buried under sediment and is compressed; (c) With greater burial, lignite coal forms; (d) At even greater depths, bituminous and eventually anthracite coal form. Source: Steven Earle (2015) CC BY 4.0 view source

At some point the swamp deposit is covered with more sediment — typically because a river changes its course or sea level rises (Figure 9.18b). As more sediments are added, the organic matter is compressed and heated as temperatures increase with depth. This has the effect of concentrating the carbon within the coal. The amount of heating will determine how far this process progresses.

The further the process does progress, the more the coal will go from having obvious pieces of plant material within it, to being a black, shiny mass.  Low-grade lignite coal forms at depths between 100 m to 1,500 m and temperatures up to ~50°C (Figure 9.18c). This is still a relatively early stage in the coal formation process, so the lignite commonly displays plant fossils that have not yet been destroyed in the process of coalification (Figure 9.19 upper right).

At between 1,000 m to 5,000 m depth and temperatures up to 150°C m, bituminous coal forms (Figure 9.18d, 9.19 lower right). At depths beyond 5,000 m and temperatures over 150°C, anthracite coal forms (Figure 9.19 lower left). In fact, as temperatures rise, the lower-grade forms of coal are actually being transformed from sedimentary to metamorphic rocks.

Figure 9.19 The formation of coal begins when plant matter is prevented from decaying by accumulating in low-oxygen, acidic water. A layer of peat forms. Heating and compression of peat form lignite, bituminous coal, and finally anthracite, as pressure and temperature increases. Source: Karla Panchuk (2017) CC BY-NC-SA 4.0. Photos by R. Weller/ Cochise College and U. S. Geological Survey. Click the image for more attributions and terms of use.

The transition from peat to anthracite results in a progressive increase in the carbon concentration, in hardness, and in the amount of energy available to be released upon combustion.

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 9.20.

Figure 9.20 Some of the important depositional environments for sediments and sedimentary rocks. Source: Steven Earle (2015) CC BY 4.0 view source, modified after Mike Norton (2008) CC BY-SA 2.0 view source

Tables 9.1 and 9.2 provide a summary of the processes and sediment types that pertain to the various depositional environments illustrated in Figure 9.19. The types of sediments that accumulate in these environments are examined in more detail in the last section of this chapter.

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. This is because they have only been deposited relatively recently — perhaps a few centuries or millennia ago — and will 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 in which sediments can be preserved for that long. Most such basins are formed by plate tectonic processes (Figure 9.21).

Figure 9.21 Some types of tectonically produced basins: (a) trench basin, (b) forearc basin, (c) foreland basin, and (d) rift basin. Source: Steven Earle (2015) CC BY 4.0 view source

Trench basins form where a subducting oceanic plate dips beneath the overriding continental or oceanic lithosphere. They can be several kilometres deep, and in many cases, host thick sequences of sediments from nearby 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, the channel between Vancouver Island and the BC mainland, is a forearc basin.

A foreland basin is caused by the mass of a mountain range depressing the crust. A rift basin forms where continental crust is being pulled apart, and the crust on both sides the rift subsides. If rifting continues this will eventually becomes a narrow sea, and then an ocean basin. The East African rift basin represents an early stage in this process.

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 that can be summarized as follows:

• The principle of original horizontality states that sediments accumulate in essentially horizontal layers. The implication is that tilted sedimentary layers observed today must have been subjected to tectonic forces that tilted them.
• The principle of superposition states that sedimentary layers are deposited in sequence, and that unless the entire sequence has been turned over by tectonic processes, the layers at the bottom are older than those at the top.
• The principle of inclusions states that any rock fragments in a sedimentary layer must be older than the layer that contains them. For example, the cobbles in a conglomerate must have been formed before the conglomerate.
• The principle of faunal succession states that there is a well-defined order in which organisms have evolved through geological time, and therefore the identification of specific fossils in a rock can be used to determine its age.

These and other principles are discussed in more detail in Chapter 19.

In addition to these principles that apply to all sedimentary rocks, a number of other important characteristics of sedimentary processes lead to 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 and depositional environment that ultimately resulted in the rocks that we are studying.

Bedding refers to sedimentary layers that can be distinguished from one another on the basis of characteristics such as texture, composition, colour, or weathering characteristics (Figure 9.22). They may also be similar layers separated by partings, narrow regions marking weaker surfaces where erosion is enhanced. 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. Bedding can form in almost any depositional environment.

Figure 9.22 Beds in the Triassic Sulphur Mt. Formation near Exshaw, Alberta. Bedding is defined by differences in colour and texture, and also by partings (darker lines) between beds that may otherwise appear to be similar. Source: Steven Earle (2015) CC BY 4.0 view source

Cross-bedding is bedding that contains angled layers. It forms when sediments are deposited by flowing water or wind (Figure 9.23). Cross-beds in streams tend to be on the scale of cm to tens of cm, while those in aeolian (wind deposited) sediments can be on the scale of metres.

Figure 9.23 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 wind direction from right to left during deposition. One bed dips in the opposite direction, implying a different wind direction. Source: Steven Earle (2015) CC BY 4.0 view source

Cross-beds form as sediments are deposited on the leading edge of an advancing ripple or dune. Each layer is related to a different ripple that advances in the flow direction, and is partially eroded by the following ripple (Figure 9.23). Cross-bedding is a very important sedimentary structure to recognize because it can provide information on 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 9.24 Formation of cross-beds as a series of ripples or dunes that migrate with the flow. Each ripple advances forward (right to left in this view) as more sediment is deposited on its leading face. Source: Steven Earle (2015) CC BY 4.0 view source

Ripples, which are associated with the formation of cross-bedding under unidirectional flow, may be preserved on the surfaces of sedimentary beds. Ripples formed in flowing water can also help to determine flow direction because they tend to have their steepest surface facing down-flow. Ripples can also form from back-and-forth flows, like at a beach, but these do not leave cross-beds, and are symmetrical, without one side steeper than the other.

Graded bedding is characterized by a change 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 (Figure 9.25), a product of deposition from a slowing current. Some graded beds are reversed (coarser at the top), and this normally results from deposition by a fast-moving debris flow. Most graded beds form in a submarine fan environment, where sediment-rich flows descend periodically from a shallow marine shelf down a slope and onto the deeper sea floor.

Figure 9.25 Graded bedding going from pebbles at the bottom to sand at the top. Source: Cropped from James St. John (2018) CC BY 2.0 view source

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 9.26).

Figure 9.26 Imbrication of clasts in a fluvial environment. Source: Steven Earle (2015) CC BY 4.0 view source

Mud cracks form when a shallow body of water (e.g., a tidal flat or pond), into which muddy sediments have been deposited, dries up and cracks (Figure 9.27). This happens because the clay in the upper mud layers shrinks upon drying.

Figure 9.27 Mud cracks in a tidal flat in England. Source: Alan Parkinson (2000) CC BY-SA 2.0 view source

The various structures described above are critical to understanding and interpreting the formation of sedimentary rocks. In addition to these structures, 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.

A Note About Fossils

Fossils are not covered in detail in this book, but they are extremely important for understanding sedimentary rocks. Fossils can be used to date sedimentary rocks, but just as importantly, they tell us a great deal about the depositional environment of the sediments and the climate at the time: they can help to differentiate marine, aquatic, and terrestrial environments; estimate the depth of the water; detect the existence of currents; and estimate average temperature and precipitation.

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) (http://www.stratigraphy.org/) has established a set of conventions for grouping, describing, and naming sedimentary rock units.

The main stratigraphic unit is a formation. 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 on 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 millions to 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 (rock type). For example, a formation that includes both shale and sandstone might be divided into members, one of which is shale, and the other sandstone. In some areas, where even more detail is required, 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.

The sedimentary rocks of the Nanaimo Group on Vancouver Island 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 BC mainland (Figure 9.28). Nanaimo Group comprises a 5000 m thick sequence of conglomerate, sandstone, and mudstone layers. Coal was mined from the Nanaimo Group rocks from around 1850 to 1950 in the Nanaimo region, and is still being mined in the Campbell River area.

Figure 9.28 The distribution of the Upper Cretaceous Nanaimo Group rocks on Vancouver Island, the Gulf Islands, and in the Vancouver area. Source: Steven Earle (2015) CC BY 4.0 view source, modified after Mustard (1994).

The Nanaimo Group is divided into 11 formations (Table 9.3). In general, the boundaries between formations are based on major lithological differences. 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. The differences in the depositional environments are probably a product of variations in tectonic-related uplift over time.

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. With the exception of the Haslam formation, they were divided into members, because that was useful for understanding the rocks in the areas where coal was mined.

There is much variety in the Nanaimo Group rocks, and it would take hundreds of photographs to illustrate all of the different types of rocks. Nevertheless, a few representative examples are shown in Figures 9.30-9.32.

Figure 9.30 Nanaimo Group, Spray Formation. Turbidite layers on Gabriola Island. Each turbidite set consists of a lower sandstone layer (light colour) that grades upward into siltstone, and then into mudstone. Source: Steven Earle (2015) CC BY 4.0 view source

Figure 9.31 Nanaimo Group, Pender Formation. Two separate layers of fluvial sandstone with a thin (approx. 75 cm) coal seam in between. Source: Steven Earle (2015) CC BY 4.0 view source

Figure 9.32 Nanaimo Group, Comox Formation. The metal object is the end of a rock hammer that is 3 cm 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. Source: Steven Earle (2015) CC BY 4.0 view source

References

Mustard, P. (1994). The Upper Cretaceous Nanaimo Group, Georgia Basin. In J. Monger (Ed.), Geology and Geological Hazards of the Vancouver Region. Geological Survey of Canada Bulletin 481, 27-95.

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

9.1 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.

9.2 Chemical and Biochemical Sedimentary Rocks

Chemical and biochemical sedimentary rocks form from ions that were transported in solution, and then converted into minerals by chemical and/or biological processes. The most common biochemical rock, limestone, typically forms in shallow tropical marine environments, where biological activity is a very important factor. Chert and banded iron formations can be from deep-ocean environments. Evaporites form where the waters of lakes and inland seas become supersaturated due to evaporation.

9.3 Organic Sedimentary Rocks

Organic sedimentary rocks contain abundant organic carbon molecules (molecules with carbon-hydrogen bonds). An example is coal, which forms when dead plant material is preserved in stagnant swamp water, and later compressed and heated.

9.4 Depositional Environments and Sedimentary Basins

There is a wide range of depositional environments, both on land (including glaciers, lakes, and rivers) and in the ocean (including deltas, reefs, shelves, and the deep-ocean floor). In order to be preserved, sediments must accumulate in sedimentary basins, many of which form through plate-tectonic processes.

9.5 Sedimentary Structures and Fossils

The deposition of sedimentary rocks can be described in terms of a series of principles, including original horizontality, superposition, and faunal succession. Sedimentary rocks can also have distinctive structures 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.

9.6 Groups, Formations, and Members

Sedimentary sequences are classified into groups, formations, and members so that they can be mapped easily and without confusion.

Review Questions

1. What are the minimum and maximum sizes of sand grains?
2. The material that makes up a rock such as conglomerate cannot be deposited by a slow-flowing river. Why?
3. Describe the two main processes of lithification.
4. What is the difference between a lithic arenite and a lithic wacke?
5. How does a feldspathic arenite differ from a quartz arenite?
6. 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?
7. What is the original source of the carbon that is present within carbonate deposits such as limestone?
8. What long-term environmental change on Earth led to the deposition of banded iron formations?
9. Name two important terrestrial depositional environments and two important marine ones.
10. What is the origin of a foreland basin, and how does it differ from a forearc basin?
11. Explain the origins of  (a) bedding, (b) cross-bedding, (c) graded bedding, and (d) mud cracks.
12. Under what conditions will reverse-graded bedding form?
13. What are the criteria for the application of a formation name to a series of sedimentary rocks?
14. Explain why some of the Nanaimo Group formations have been divided into members, while others have not.

1. Sand grains range in size from 1/16 mm to 2 mm.

2. Conglomerate cannot be deposited by a slow-flowing river because clasts larger than 2 mm are not transported by slow-moving water.

3. Sediments are buried beneath other sediments where, because of the increased pressure, they become compacted and water is forced out from between the grains. With additional burial they are warmed to the point where cementing minerals can form between the grains (less than 200˚C).

4. Lithic arenite has less than 15% silt- and clay-sized particles, while a lithic wacke has more than 15%. Both have more than 10% rock fragments and more rock fragments than feldspar.

5. Feldspathic arenite has more than 10% feldspar and more feldspar than rock fragments. Quartz arenite has less than 10% feldspar and less than 10% rock fragments. Both have less than 15% silt and clay.

6. Source area lithology: rock that contains quartz (such as granite or sandstone). Strong weathering is required to remove feldspar, and long fluvial transportation to round the grains.

7. The carbon within carbonate deposits such as limestone ultimately comes from the atmosphere.

8. Most of Earth’s banded iron formations formed during the initial oxygenation of the atmosphere between 2.4 and 1.8 Ga because iron that had been soluble in the anoxic oceans became insoluble in the oxidized oceans.

9. Terrestrial depositional environments: rivers, lakes, deltas, deserts, glaciers. Marine depositional environments: continental shelves, continental slopes, deep ocean.

10. A foreland basin forms in the vicinity of a large range of mountains where the weight of the mountains depresses the crust on either side. A forearc basin lies between a subduction zone and the related volcanic arc.

11. (a) Bedding forms where there is an interruption or change in the depositional process, or a change in the composition of the material being deposited. (b) Cross-bedding forms in fluvial or aeolian environments where sand-sized sediments are being moved and ripples or dunes are present. (c) Graded bedding forms when transport energy decreases, depositing finer and finer particles. (d) Mud cracks form where fine-grained sediments (silt or clay) are allowed to dry.

12. Reverse-graded bedding forms during gravity flows, such as debris flows.

13. A formation is a series of beds that is distinct from other beds above and below it, and is thick enough to be shown on the geological maps that are widely used within the area in question.

14. The Nanaimo Group was actively mined for coal for many decades. During that time the names were given to members and individual beds that were important to the coal miners.

The main factors that control metamorphic processes are:

• The chemical composition of the parent rock
• The temperature at which metamorphism takes place
• The pressure applied, and whether the pressure is equal in all directions or not
• The amount and type of fluid (mostly water) that is present during metamorphism
• The amount of time over which metamorphic conditions are sustained

Mineral composition

Parent rocks can be from any of the three rock types: sedimentary, igneous, or metamorphic.  The critical feature of the parent rock is its mineral composition.  This is because the stability of minerals—how they are influenced by changing conditions—is what determines which minerals form as metamorphism takes place. When a rock is subjected to increased temperatures and pressures, some minerals will undergo chemical reactions and turn into new minerals, while others might just change their size and shape.

Temperature

The temperature under which metamorphism occurs is a key variable in determining which metamorphic reactions happen. Mineral stability depends on temperature, pressure, and the presence of fluids. Minerals are stable over a specific range of temperatures. Quartz, for example, is stable from surface temperatures up to approximately 1800°C. If the pressure is higher, that upper limit will also be higher. If there is water present, it will be lower. Most other common minerals have upper limits between 150°C and 1000°C.

Some minerals will change their crystal structure depending on the temperature and pressure. Quartz has different polymorphs that are stable between 0°C and 1800°C. The minerals kyanite, andalusite, and sillimanite are polymorphs with the composition Al₂SiO₅. The fact that they are stable at different pressures and temperatures means that their presence can be used to determine the pressures and temperatures experienced by a metamorphic rock that contains one or more of the polymorphs (Figure 10.3).

Figure 10.3 The Al₂SiO₅ polymorphs andalusite, kyanite, and sillimanite, and their stability fields. Source: Karla Panchuk (2018) CC BY-SA 4.0. Photos by Rob Lavinsky/ iRocks.com (pre-2010) CC BY-SA 3.0. View source for andalusite/ kyanite/ sillimanite.

Pressure

Pressure has implications for mineral stability, and therefore the mineral content of metamorphic rocks, but it also determines the texture of metamorphic rocks. When directed pressure (or directed stress) acts on a rock, it means the stress on the rock is much greater in one direction than another. In an experiment with cylinders of modeling clay stacked in a block (Figure 10.4, top), pushing down on the clay from above resulted in higher directed pressure in the up-down direction (larger arrows; downward from pushing on the clay, and upward from the force of the table beneath the clay) than in the sideways direction, where only air pressure was acting (small arrows). The clay cylinders became elongated in the direction of least pressure.

Figure 10.4 Modelling clay experiments showing the effects of pressure on textures. Top: Directed pressure- clay was set on a flat surface and pushed down on from above (large arrows). Cylinders making up the clay block became elongated in the direction of least stress. Bottom: Shear stress applied to the top and bottom of a block of clay caused the interior to stretch. Note white dashed reference circles and elongated ellipses. Source: Karla Panchuk (2018) CC BY 4.0

Rocks undergo shear stress when forces act parallel to surfaces. In another modelling-clay experiment, applying oppositely directed forces to the top and bottom of a block of clay (Figure 10.4, bottom) caused diagonal stretching within the block. Note the change in shape of the dashed white reference circles.

In both experiments, parts of the clay became elongated in a particular direction. When mineral grains within a rock become aligned like this, it produces a fabric called foliation. Foliation is described in more detail later in this chapter.

Fluids

Water is the main fluid present within rocks of the crust, and the only one considered 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. This speeds the process up so metamorphism might occur more rapidly, 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 substances, making it an important medium for moving ions from one place to another within the crust. Processes facilitated by hot water are called hydrothermal processes (hydro refers to water, and thermal refers to heat).

Time

Most metamorphic reactions occur very slowly. Estimates of the growth rates of new minerals within a rock during metamorphism suggest that new material is added to the outside of mineral crystals at a rate of approximately 1 mm per million years. Very slow reaction rates make it 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 there is a good chance that metamorphic reactions will be completed. For example, an important setting for metamorphism 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: How Long Did It Take?

The large reddish crystals in Figure 10.5 are garnet, and the surrounding light coloured rock is dominated by muscovite mica. The Euro coin is 23 mm in diameter. Assume that the diameters of the garnets increased at a rate of 1 mm per million years. Based on the approximate average diameter of the garnets visible, estimate how long this metamorphic process might have taken.

Figure 10.5 Garnet-mica schist from the Greek island of Syros. Source: Graeme Churchard (2005) CC BY 2.0 view source

 

How Foliation Develops

When a rock is acted upon by pressure that is not the same in all directions, or by shear stress (forces acting to “smear” the rock), minerals can become elongated in the direction perpendicular to the main stress. The pattern of aligned crystals that results is called foliation.

Foliation can develop in a number of ways. Minerals can deform when they are squeezed (Figure 10.6), becoming narrower in one direction and longer in another.

Figure 10.6 Foliation that develops when minerals are squeezed and deform by lengthening in the direction perpendicular to the greatest stress (indicated by black arrows). Left- before squeezing. Right- after squeezing. Source: Steven Earle (2015) CC BY 4.0 view source

If a rock is both heated and squeezed during metamorphism, and the temperature change is enough for new minerals to form from existing ones, the new minerals can be forced to grow longer perpendicular to the direction of squeezing (Figure 10.7). If the original rock had bedding (represented by diagonal lines in Figure 10.7, right), foliation may obscure the bedding.

Figure 10.7 Effects of squeezing and aligned mineral growth during metamorphism. Left: Protolith with diagonal bedding. Right: Metamorphic rock derived from the protolith. Elongated mica crystals grew perpendicular to the main stress direction. The original bedding is obscured. Source: Steven Earle (2015) CC BY 4.0 view source

This is not always the case, however. The large boulder in Figure 10.8 in has strong foliation, oriented nearly horizontally in this view, but it also has bedding still visible as dark and light bands sloping steeply down to the right.

Figure 10.8 A geologists sits on a rock that has foliation (marked by the dashed line that is nearly horizontal), and still retains evidence of the original bedding (steeply dipping dashed line). The rock has undergone a relatively low degree of metamorphism, which is why the bedding is still visible. Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2015) CC BY 4.0 view source

Foliation and Crystal Habit

Most foliation develops when new minerals are forced to grow perpendicular to the direction of greatest stress. This effect is especially strong if the new minerals grow in platy or elongated shapes. The rock in the upper left of Figure 10.9 is foliated, and the microscopic structure of the same type of foliated rock is shown in the photograph beneath it. Over all, the photomicrograph shows that the rock is dominated by elongated crystals aligned in bands running from the upper left to the lower right. The stress that produced this pattern was greatest in the direction indicated by the black arrows, at a right angle to the orientation of the minerals. The aligned minerals are mostly mica, which has a platy crystal habit, with plates stacked together like pages in a book.

Figure 10.9 A foliated metamorphic rock called phyllite (upper left). The satin sheen comes from the alignment of minerals. Lower left- a view of the same kind of rock under a microscope showing mica crystals (colourful under polarized light) aligned in bands. The region outlined in a red dashed line shows a lens of quartz crystals that do not display alignment. Upper right- stacks of platy mica crystals. Lower right- a blocky quartz crystal. Source: Karla Panchuk (2018) CC BY-SA 4.0. Click the image for photo sources.

The zone in the photomicrograph outlined with the red dashed line is different from the rest of the rock. Not only is the mineral composition different—it is quartz, not mica—but the crystals are not aligned. The quartz crystals were subjected to the same stress as the mica crystals, but because quartz grows in blocky shapes rather than elongated ones, the crystals could not be aligned in any one direction.

Even though the quartz crystals themselves are not aligned, the mass of quartz crystals forms a lens that does follow the general trend of alignment within the rock. This happens because the stress can cause some parts of the quartz crystals to dissolve, and the resulting ions flow away at right angles to the greatest stress before forming crystals again.

The effects of recrystallization in Figure 10.9 would not be visible with the unaided eye, but when larger crystals or large clasts are involved, the effects can be visible as “shadows” or “wings” around crystals and clasts. The rock in Figure 10.10 had a quartz-rich conglomerate as a parent rock. Differential stress has caused quartz pebbles within the rock to become elongated, and it has also caused wings to form around some of the pebbles (see the pebble in the dashed ellipse). The location of the wings depends on the distribution of stress on the rock (Figure 10.10, upper right).

Figure 10.10 Metaconglomerate with elongated of quartz pebbles. The pebbles have developed “wings” to varying degrees (e.g., white dashed ellipse). These are the result of quartz dissolving where stress is applied, and flowing away from the direction of maximum stress before recrystallizing (upper right sketch). Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Photo by R. Weller/ Cochise College view source. Click the image to view terms of use.

Foliation Controls How Rocks Break

Foliated metamorphic rocks have elongated crystals that are oriented in a preferred direction. This forms planes of weakness, and when these rocks break, they tend to break along surfaces that parallel the orientation of the aligned minerals (Figure 10.11). Breaks along planes of weakness within a rock that are caused by foliation are referred to as rock cleavage, or just cleavage.  This is distinct from cleavage in minerals because mineral cleavage happens between atoms within a mineral, but rock cleavage happens between minerals.

Figure 10.11 Close-up view of a metamorphic rock with aligned elongated crystals. The crystals control the shape of the break in the rock (black gap), resulting in breaks occurring along parallel surfaces. Source: Karla Panchuk (2018) CC BY 4.0

The mineral alignment in the metamorphic rock called slate is what causes it to break into flat pieces (Figure 10.12, left), and is why slate has been used as a roofing material (Figure 10.12, right). The tendency of slate to break into flat pieces is called slaty cleavage.

Figure 10.12 Rock cleavage in the fine-grained metamorphic rock called slate results in breaks along relatively flat surfaces (left). This is why slate has been used for roofing material (right). Source: Left- Roger Kidd (2008) CC BY-SA 2.0 view source; Right- Michael C. Rygel (2007) CC BY-SA 3.0 view source

Rock cleavage is what caused the boulder in Figure 10.8 to split from bedrock in a way that left the flat upper surface upon which the geologist is sitting.

Metamorphic rocks are broadly classified as foliated or non-foliated. Non-foliated metamorphic rocks do not have aligned mineral crystals. Non-foliated rocks form when pressure is uniform, or near the surface where pressure is very low. They can also form when the parent rock consists of blocky minerals such as quartz and calcite, in which individual crystals do not align because they aren’t longer in any one dimension. This distinction breaks down in zones of intense deformation, where even minerals like quartz can be squeezed into long stringers, much like squeezing toothpaste out of a tube (Figure 10.13).

Figure 10.13 Rocks from the Western Carpathians mountain range without deformation (left) and after deformation (right). Scale bar: 1 mm. Left- An undeformed granitic rock containing the mica mineral biotite (Bt), plagioclase feldspar (Pl), potassium feldspar (Kfs), and quartz (Qtz). Right- A metamorphic rock (mylonite) resulting from extreme deformation of granitic rocks. Quartz crystals have been flattened and deformed. The other minerals have been crushed and deformed into a fine-grained matrix (Mtx). Source: Farkašovský et al. (2016) CC BY-NC-ND. Click the image to view the original figure captions and access the full text.

Types of Foliated Metamorphic Rocks

Four common types of foliated metamorphic rocks, listed in order of metamorphic grade or intensity of metamorphism are slate, phyllite, schist (pronounced “shist”), and gneiss (pronounced “nice”). Each of these has a characteristic type of foliation

Slate

Slate (Figure 10.14) forms from the low-grade metamorphism of shale. Slate has microscopic clay and mica crystals that have grown perpendicular to the maximum stress direction. Slate tends to break into flat sheets or plates, a property described as slaty cleavage.

Figure 10.14 Slate, a low-grade foliated metamorphic rock. Left- Slate fragments resulting from rock cleavage. Right- The same rock type in outcrop. Source: Karla Panchuk (2018) CC BY-SA 4.0. Photos: Left- Vincent Anciaux (2005) CC BY-SA 3.0 view source; Right- Gretarsson (2006) CC BY-SA 3.0 view source

Phyllite

Phyllite (Figure 10.15) is similar to slate, but has typically been heated to a higher temperature. As a result, the micas have grown larger.  They still are not visible as individual crystals, but the larger size leads to a satiny sheen on the surface.  The cleavage of phyllite is slightly wavy compared to that of slate.

Figure 10.15 Phyllite, a fine-grained foliated metamorphic rock. Left- A hand sample showing a satin texture. Right- The same rock type in outcrop in the city of Sopron, Hungary. Source: Karla Panchuk (2018) CC BY-SA 4.0. Photos: Left- Chadmull (2006) Public Domain view source; Right- Laszlovszky András (2008) CC BY-SA 2.5 view source

Schist

Schist (Figure 10.16) forms at higher temperatures and pressures and exhibits mica crystals that are large enough to see without magnification. Individual crystal faces may flash when the sample is turned in the light, making the rock appear to sparkle. Other minerals such as garnet might also be visible, but it is not unusual to find that schist consists predominantly of a single mineral.

Figure 10.16 Schist, a medium- to high-grade foliated metamorphic rock. Top- Hand sample showing light reflecting off of mica crystals. Bottom- Close-up view of mica crystals and garnet. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Photos by R. Weller/ Cochise College. Click the image for photo sources and terms of use.

Gneiss

Gneiss (Figure 10.17) forms at the highest pressures and temperatures, and has crystals large enough to see with the unaided eye. Gneiss features minerals that have separated into bands of different colours. The bands of colours are what define foliation within gneiss. Sometimes the bands are very obvious and continuous (Figure 10.17, upper right), but sometimes they are more like lenses (upper left). 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.

Figure 10.17 Gneiss, a coarse-grained, high grade metamorphic rock, is characterized by colour bands. Top- Hand samples showing that colour bands can be continuous (left) or less so (right). Bottom- Gneiss in outcrop at Belteviga Bay, Norway. Notice the light and dark stripes on the rock. Source: Karla Panchuk (2018) CC BY-SA 4.0. Click the image for more attributions.

While slate and phyllite typically form only from mudrock protoliths, 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: 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, gneiss that originated as basalt and is dominated by amphibole, is an amphibole gneiss or amphibolite (Figure 10.18).

Figure 10.18 Amphibolite in thin section (2mm field of view), derived from metamorphism of a mafic igneous rock. Green crystals are the amphibole hornblende, and colourless crystals are plagioclase feldspar. Note horizontal crystal alignment. Source: D.J. Waters, University of Oxford view source/ view context. Click the image for original figure caption and terms of use.

Types of Non-foliated Metamorphic Rocks

Metamorphic rocks that form under low-pressure conditions or under the effects confining pressure, which is equal in all directions, do not become foliated. In most cases, this is because they are not buried deeply enough, and the heat for the metamorphism comes from a body of magma that has moved into the upper part of the crust. Metamorphism that happens because of proximity to magma is called contact metamorphism. Some examples of non-foliated metamorphic rocks are marble, quartzite, and hornfels.

Marble

Marble (Figure 10.19) is metamorphosed limestone. When it forms, the calcite crystals recrystallize (re-form into larger blocky calcite crystals), and any sedimentary textures and fossils that might have been present are destroyed. If the original limestone is pure calcite, then the marble will be white.  On the other hand, if it has impurities such as clay, silica, or magnesium, the marble could be “marbled” in appearance (Figure 10.19, bottom).

Figure 10.19 Marble is a non-foliated metamorphic rock with a limestone protolith. Left- Marble made of pure calcite is white. Upper right- microscope view of calcite crystals within marble that are blocky and not aligned. Lower right- A quarry wall showing the “marbling” that results when limestone contains components other than calcite. Source: Karla Panchuk (2018) CC BY-NC-SA. Click the image for more attributions.

Quartzite

Quartzite (Figure 10.20) is metamorphosed sandstone. It is dominated by quartz, and in many cases, the original quartz grains of the sandstone are welded together with additional silica. Sandstone often contains some clay minerals, feldspar or lithic fragments, so quartzite can also contain impurities.

Figure 10.20 Quartzite is a non-foliated metamorphic rock with a sandstone protolith. Left- Quartzite from the Baraboo Range, Wisconsin. Right- Photomicrograph showing quartz grains in quartzite from the Southern Appalachians. In the upper left half of the image, blocky quartz crystals show some evidence of alignment running from the upper right to the lower left. Source: Karla Panchuk (2018) CC BY-SA 4.0. Photomicrograph: Geologian (2011) CC BY-SA 3.0 view source

Even if formed under directed pressure, quartzite is generally not foliated because quartz crystals do not normally 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.

Hornfels

Hornfels is another non-foliated metamorphic rock that normally forms during contact metamorphism of fine-grained rocks like mudstone or volcanic rocks. Hornfels have different elongated or platy minerals (e.g., micas, pyroxene, amphibole, and others) depending on the exact conditions and the parent rock, yet because the pressure wasn’t substantially higher in any particular direction, these crystals remain randomly oriented.

The hornfels in Figure 10.21 (left) appears to have gneiss-like bands, but these actually reflect the beds of alternating sandstone and shale that were in the protolith. They are not related to alignment of crystals due to metamorphism. On the right of Figure 10.21 is a microscopic view of another sample of hornfels, also from a sedimentary protolith. The dark band at the top is from the original bedding.  Here you can see that the brown mica crystals (biotite) are not aligned.

Figure 10.21 Hornfels, a non-foliated metamorphic rock formed from a fine-grained protolith. Left- Hornfels from the Novosibirsk region of Russia from a sedimentary protolith. Dark and light bands preserve the bedding of the original sedimentary rock. The rock has been recrystallized during contact metamorphism and does not display foliation. (scale in cm). Right- Hornfels in thin section from a sedimentary protolith. Note that the brown mica crystals are not aligned. The dark band at the top reflects the layering within the sedimentary parent rock, similar to the way those layers are preserved in the sample on the left. Source: Left- Fedor (2006) Public Domain view source; Right- D.J. Waters, University of Oxford view source/ view context. Click the image for terms of use.

What Happens When Different Rocks Undergo Metamorphism?

The nature of the parent rock controls the types of metamorphic rocks that can form from it under differing metamorphic conditions (temperature, pressure, fluids). The kinds of rocks that can be expected to form at different metamorphic grades from various parent rocks are listed in Table 10.1.

Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2015) CC BY 4.0 view source. Click the table for a text version.

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. Sandstone and limestone don’t change much either because their metamorphic forms (quartzite and marble, respectively) have the same mineral composition, but re-formed larger crystals.

On the other hand, some rocks can change substantially.  Mudrock (e.g., shale, mudstone) can start out as slate, then progress through phyllite, schist, and gneiss, with a variety of different minerals forming along the way.  Schist and gneiss can also form from sandstone, conglomerate, and a range of both volcanic and intrusive igneous rocks.

Migmatite: Both Metamorphic and Igneous

If a metamorphic rock is heated enough, it can begin to undergo partial melting in the same way that igneous rocks do.  The more felsic minerals (feldspar, quartz) will melt, while the more mafic minerals (biotite, hornblende) do not.  When the melt crystallizes again, the result is light-coloured igneous rock interspersed with dark-coloured metamorphic rock.  This mixed rock is called migmatite (Figure 10.22). Note that the foliation present in the metamorphic rock is no longer present in the igneous rock. Liquids cannot support a differential stress, so when the melt crystallizes, the foliation is gone.

Figure 10.22 Migmatite photographed near Geirangerfjord in Norway. Source: Siim Sepp (2006) CC BY-SA 3.0 view source

A fascinating characteristic of migmatites is ptygmatic (pronounced “tigmatic“) folding. These are folds look like they should be impossible because they are enveloped by rock which does not display the same complex deformation (Figure 10.23).  How could those wiggly folds get in there without the rest of the rock being folded in the same way?

Figure 10.23 Ptygmatic folding from Broken Hill, New South Wales, Australia. Ptygmatic folding happens when a stiff layer within a rock is surrounded by weaker layers. Folding causes the stiff layer to crinkle while the weaker layers deform around it. Source: Roberto Weinberg (http://users.monash.edu.au/~weinberg) view source. Click the image for terms of use.

The answer to the ptygmatic fold mystery is that the folded layer is much stiffer than the surrounding layers.  When squeezing forces act on the rock, the stiff layer buckles but the surrounding rock flows rather than buckling, because it isn’t strong enough to buckle.

References

Farkašovský, R., Bónová, K., & Košuth, M. (2016). Microstructural, modal and geochemical changes as a result of granodiorite mylonitisation – a case study from the Rolovská shear zone (Čierna hora Mts, Western Carpathians, Slovakia). Geologos 22(3), 171-190. doi: 10.1515/logos-2016-0019 View full text

The outcome of metamorphism depends on pressure, temperature, and the abundance of fluid involved, and there are many settings with unique combinations of these factors. Some types of metamorphism are characteristic of specific plate tectonic settings, but others are not.

Burial Metamorphism

Burial metamorphism occurs when sediments are buried deeply enough that the heat and pressure cause minerals to begin to recrystallize and new minerals to grow, but does not leave the rock with a foliated appearance. As metamorphic processes go, burial metamorphism takes place at relatively low temperatures (up to ~300 °C) and pressures (100s of m depth). To the unaided eye, metamorphic changes may not be apparent at all. Contrast the rock known commercially as Black Marinace Gold Granite (Figure 10.24)—but which is in fact a metaconglomerate—with the metaconglomerate in Figure 10.10. The metaconglomerate formed through burial metamorphism does not display any of the foliation that has developed in the metaconglomerate in Figure 10.10.

Figure 10.24 Metaconglomerate formed through burial metamorphism. The pebbles in this sample are not aligned and elongated as in the metaconglomerate in Figure 10.10. Source: James St. John (2014) CC BY 2.0 view source

Regional Metamorphism

Regional metamorphism refers to large-scale metamorphism, such as what happens to continental crust along convergent tectonic margins (where plates collide).  The collisions result in the formation of long mountain ranges, like those along the western coast of North America.  The force of the collision causes rocks to be folded, broken, and stacked on each other, so not only is there the squeezing force from the collision, but from the weight of stacked rocks. The deeper rocks are within the stack, the higher the pressures and temperatures, and the higher the grade of metamorphism that occurs. Rocks that form from regional metamorphism are likely to be foliated because of the strong directional pressure of converging plates.

The Himalaya range is an example of where regional metamorphism is happening because two continents are colliding (Figure 10.25). Sedimentary rocks have been both thrust up to great heights—nearly 9 km 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 in the crust, rock buried to 9 km below sea level in this situation could be close to 18 km below the surface of the ground, and it is reasonable to expect temperatures up to 500°C. Notice the sequence of rocks that from, beginning with slate higher up where pressures and temperatures are lower, and ending in migmatite at the bottom where temperatures are so high that some of the minerals start to melt. These rocks are all foliated because of the strong compressing force of the converging plates.

Figure 10.25 Regional metamorphism beneath a mountain range resulting from continent-continent collision. Arrows show the forces due to the collision. Dashed lines represent temperatures that would exist given a geothermal gradient of 30 ºC/km. A sequence of foliated metamorphic rocks of increasing metamorphic grade forms at increasing depths within the mountains. Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2015) CC BY 4.0 view source

Seafloor (Hydrothermal) Metamorphism

At an oceanic spreading ridge, recently formed oceanic crust of gabbro and basalt is slowly moving away from the plate boundary (Figure 10.26). Water within the crust is forced to rise in the area close to the source of volcanic heat, drawing in more water from further away. This eventually creates a convective system where cold seawater is drawn into the crust, heated to 200 °C to 300 °C as it passes through the crust, and then released again onto the seafloor near the ridge.

Figure 10.26 Seafloor (hydrothermal) metamorphism of ocean crustal rock on either side of a spreading ridge. Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2015) CC BY 4.0 view source

The passage of this water through the oceanic crust at these temperatuers promotes metamorphic reactions that change the original olivine and pyroxene minerals in the rock to chlorite ((Mg₅Al)(AlSi₃)O₁₀(OH)₈) and serpentine ((Mg, Fe)₃Si₂O₅(OH)₄). Chlorite and serpentine are both hydrated minerals, containing water in the form of OH in their crystal structures. 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 melting.

The low-grade metamorphism occurring at these relatively low pressures and temperatures can turn mafic igneous rocks in ocean crust into greenstone (Figure 10.27), a non-foliated metamorphic rock.

Figure 10.27 Greenstone from the metamorphism of seafloor basalt that took place 2.7 billion years ago. The sample is from the Upper Peninsula of Michigan, USA. Source: James St. John (2012) CC BY 2.0 view source

Subduction Zone Metamorphism

At subduction zones, where ocean lithosphere is forced down into the hot mantle, there is a unique combination of relatively low temperatures and very high pressures.  The high pressures are to be expected, given the force of collision between tectonic plates, and the increasing lithostatic pressure as the subducting slab is forced deeper and deeper into the mantle. The lower temperatures exist because even though the mantle is very hot, ocean lithosphere is relatively cool, and a poor conductor of heat. That means it will take a long time to heat up, can be several hundreds of degrees cooler than the surrounding mantle. In Figure 10.28, notice that the isotherms (lines of equal temperature, dashed lines) plunge deep into the mantle along with the subducting slab, showing that regions of relatively low temperature exist deeper in the mantle.

Figure 10.28 Regional metamorphism of oceanic crust at a subduction zone occurs at high pressure but relatively low temperatures. Source: Steven Earle (2015) CC BY 4.0 view source

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).  Glaucophane is blue, and the major component of a rock known as blueschist. If you have never seen or even heard of blueschist, that not surprising. What is surprising is that anyone has seen it! Most of the blueschist that forms in subduction zones continues to be subducted. It turns into eclogite at about 35 km depth, and then eventually sinks deep into the mantle, never to be seen again. In only a few places in the world, the subduction process was interrupted, and partially subducted blueschist returned to the surface. One such place is the area around San Francisco. The blueschist at this location is part of a set of rocks known as the Franciscan Complex (Figure 10.29).

Figure 10.29 Franciscan Complex blueschist exposed north of San Francisco. The blue colour of the rock is due to the presence of the amphibole mineral glaucophane. Source: Steven Earle (2015) CC BY 4.0 view source

Contact Metamorphism

Contact metamorphism happens when a body of magma intrudes into the upper part of the crust. Heat is important in contact metamorphism, but pressure is not a key factor, so contact metamorphism produces non-foliated metamorphic rocks such as hornfels, marble, and quartzite.

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 that develops around the magma body, will depend on a number of factors, including the type of country rock, the temperature of the intruding body, the size of the body, and the volatile compounds within the body (Figure 10.30). 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. This will allow the heat to extend farther into the country rock, creating a larger aureole. Volatiles may exsolve from the intruding melt and travel into the country rock, facilitating heating and carrying chemical constituents from the melt into the rock. Thus, aureoles that form around “wet” intrusions tend to be larger than those forming around their dry counterparts.

Figure 10.30 Schematic cross-section of the middle and upper crust showing two magma bodies. The upper body, which has intruded into cool unmetamorphosed rock, 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. Source: Steven Earle (2015) CC BY 4.0 view source

Contact metamorphic aureoles are typically quite small, from just a few centimetres around small dykes and sills, to as much as 100 m around a large stock. Contact metamorphism can take place over a wide range of temperatures—from around 300 °C to over 800 °C. Different minerals will form depending on the exact temperature and the nature of the country rock.

Although bodies of magma can form in a variety of settings, one place magma is produced in abundance, and where contact metamorphism can take place, is along convergent boundaries with subduction zones, where volcanic arcs form (Figure 10.31). Regional metamorphism also takes place in this setting, and because of the extra heat associated with the magmatic activity, the geothermal gradient is typically steeper in these settings (between ~40 and 50 °C/km). Under these conditions, higher grades of metamorphism can take place closer to surface than is the case in other areas.

Figure 10.31 Contact metamorphism (yellow rind) around a high-level crustal magma chamber, and regional metamorphism in a volcanic-arc related mountain range. Dashed lines show isotherms. Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2015) CC BY 4.0 view source

Shock Metamorphism

When extraterrestrial objects hit Earth, the result is a shock wave.  Where the object hits, pressures and temperatures become very high in a fraction of a second.  A “gentle” impact can hit with 40 GPa and raise temperatures up to 500 °C. Pressures in the lower mantle start at 24 GPa (GigaPascals), and climb to 136 GPa at the core-mantle boundary, so the impact is like plunging the rock deep into the mantle and releasing it again within seconds.  The sudden change associated with shock metamorphism makes it very different from other types of metamorphism that can develop over hundreds of millions of years, starting and stopping as tectonic conditions change.

Two features of shock metamorphism are shocked quartz, and shatter cones.  Shocked quartz (Figure 10.32 left) refers to quartz crystals that display damage in the form of parallel lines throughout a crystal.  The quartz crystal in Figure 10.32 has two sets of these lines.  The lines are small amounts of glassy material within the quartz, formed from almost instantaneous melting and resolidification when the crystal was hit by a shock wave. Shatter cones are cone-shaped fractures within the rocks, also the result of a shock wave (Figure 10.32 right).  The fractures are nested together like a stack of ice-cream cones.

Figure 10.32 Shock metamorphism features. Left- Shocked quartz with lines of glassy material, from the Suvasvesi South impact structure in Finland. Right- Shatter cones from the Wells Creek impact crater in the USA. Sources: Left- Martin Schmieder CC BY 3.0 view source. Right- Zamphuor (2007) Public Domain view source.

Dynamic Metamorphism

Dynamic metamorphism is the result of very high shear stress, such as occurs along fault zones. Dynamic metamorphism occurs at relatively low temperatures compared to other types of metamorphism, and consists predominantly of the physical changes that happen to a rock experiencing shear stress. It affects a narrow region near the fault, and rocks nearby may appear unaffected.

At lower pressures and temperatures, dynamic metamorphism will have the effect of breaking and grinding rock, creating cataclastic rocks such as fault breccia (Figure 10.33). At higher pressures and temperatures, grains and crystals in the rock may deform without breaking into pieces (Figure 10.34, left). The outcome of prolonged dynamic metamorphism under these conditions is a rock called mylonite, in which crystals have been stretched into thin ribbons (Figure 10.34, right).

Figure 10.33 Fault breccia, created when shear stress along a fault breaks up rocks. Left- close-up view of fault breccia clearly showing dark angular fragments. Right- A fault-zone containing fragments broken from the adjacent walls (dashed lines). Note that the deformation does not extend far past the margins of the fault zone. Source: Karla Panchuk (2018) CC BY 4.0. Click the image for more attributions.

Figure 10.34 Mylonite, a rock formed by dynamic metamorphism. Left- An outcrop showing the early stages of mylonite development, called protomylonite. Notice that the deformation does not extend to the rock at the bottom of the photograph. Middle- Mylonite showing ribbons formed of drawn-out crystals. Right- Microscope view of mylonite with mica (colourful crystals) and quartz (grey and black crystals). This is a case where the shape of quartz crystals is controlled more by stress than by crystal habit. Source: Karla Panchuk (2018) CC BY-SA 4.0. Click the image for more attributions.

References

Bucher, K., & Grapes, R. (2011) Petrogenesis of Metamorphic Rocks, 8th Edition. Springer.

French, B.M. (1998). Traces of Catastrophe: A Handbook of Shock-Metamorphic Effects in Terrestrial Meteorite Impact Structures. Houston, TX: Lunar and Planetary Institute  Read full text

Metamorphic Facies

In any given metamorphic setting there can be a variety of protolith types exposed to metamorphism.  While these rocks will be exposed to the same range of pressure and temperatures conditions within that setting, the metamorphic rock that results will depend on the protolith. A convenient way to indicate the range of possible metamorphic rocks in a particular setting is to group those possibilities into metamorphic facies. In other words, a given metamorphic facies groups together metamorphic rocks that form under the same pressure and temperature conditions, but which have different protoliths.

Figure 10.35 shows the different metamorphic facies as patches of different colours. The axes on the diagram are temperature and depth; the depth within the Earth will determine how much pressure a rock is under, so the vertical depth axis is also a pressure axis. Each patch of colour represents a range of temperature and pressure conditions where particular types of metamorphic rocks will form. Metamorphic facies are named for rocks that form under specific conditions (e.g., eclogite facies, amphibolite facies etc.), but those names don’t mean that the facies is limited to that one rock type.

Figure 10.35 Metamorphic facies and types of metamorphism shown in the context of depth and temperature. The metamorphic rocks formed from a mudrock protolith under regional metamorphism with a typical geothermal gradient are listed. Letters correspond to the types of metamorphism shown in Figure 10.36. Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2016) CC BY 4.0 view source

Another feature to notice in the diagram are the many dashed lines. The yellow, green, and blue dashed lines represent the geothermal gradients in different environments. Recall that the geothermal gradient describes how rapidly the temperature increases with depth in Earth. In most areas (green dashed line), the rate of increase in temperature with depth is 30 °C/km. In other words, if you go 1,000 m down into a mine, the temperature will be roughly 30 °C warmer than the average temperature at the surface.  In volcanic areas (yellow dashed line), the geothermal gradient is more like 40 to 50 °C/km, so the temperature rises much faster as you go down. Along subduction zones (blue dashed line), the cold ocean lithosphere keeps temperatures low, so the gradient is typically less than 10 °C/km.

The yellow, green, and blue dashed lines in Figure 10.35 tell you what metamorphic facies you will encounter for rocks from a given depth in that particular environment. A depth of 15 km in a volcanic region falls in the amphibolite facies.  Under more typical conditions, this is the greenschist facies, and in a subduction zone it is the blueschist facies. You can make the connection more directly between the metamorphic facies and the types of metamorphism discussed in the previous section by matching up the letters a through e in Figure 10.35 with the labels in Figure 10.36.

Figure 10.36 Environments of metamorphism in the context of plate tectonics: (a) regional metamorphism related to mountain building at a continent-continent convergent boundary, (b) seafloor (hydrothermal) metamorphism of oceanic crust in the area on either side of a spreading ridge, (c) 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. Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2015) CC BY 4.0 view source

One other line to notice in Figure 10.35 is the red dashed line on the right-hand side of the figure. This line represents temperatures and pressures where granite will begin to melt if there is water present. Migmatite is to the right of the line because it forms when some of the minerals in a metamorphic rock begin to melt, and then cool and crystallize again.

Index Minerals

Some common minerals in metamorphic rocks are shown in Figure 10.37, arranged in order of the temperature ranges within 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 10.37 Metamorphic index minerals and approximate temperature ranges. Source: Steven Earle (2015) CC BY 4.0 view source

Even though the limits of the stability ranges are vague, the stability range of each mineral is still small enough that the minerals can be used as markers for those metamorphic conditions. Minerals that make good markers of specific ranges of metamorphic conditions are called index minerals.

The Meguma Terrane of Nova Scotia: An Example of How to Use Index Minerals

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 Terrane (Figure 10.38, top )—collided with the existing eastern margin of North America. The clastic sedimentary rocks within this terrane were variably metamorphosed.

Figure 10.38 Regional metamorphic zones in the Meguma Terrane of southwestern Nova Scotia. Top- Map of metamorphic zones. Bottom- Stability ranges for minerals within the Meguma Terrane. Source: Karla Panchuk (2017) CC BY 4.0. Modified after Steven Earle (2015) CC BY 4.0 view source, Keppie & Muecke (1979) and White & Barr (2012).

Index minerals have been used to map areas of higher or lower metamorphic intensity, called metamorphic zones. A metamorphic zone is a region bounded by the first appearance of an index mineral. In the Meguma Terrane, the biotite zone (Figure 10.38, darker green) begins in the east and north with the first appearance of biotite. The biotite zone ends toward the south and west where garnet first appears. Because index minerals can have overlapping stability conditions, a lower-intensity index mineral can still be present in a higher-intensity metamorphic zone.

Knowledge of metamorphic zones makes it possible to draw conclusions about the geological conditions in which metamorphic rocks formed. The highest-intensity metamorphism—the sillimanite zone—is in the southwest. Progressively lower grades of metamorphism exist toward the east and north. The rocks of the sillimanite zone were likely heated to over 700 °C, and therefore must have been buried to depths between 20 km and 25 km. The surrounding lower-grade rocks were not buried as deeply, and the rocks within the peripheral chlorite zone were likely not buried to more than about 5 km depth.

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. The collision caused rocks to be folded, and to be faulted and stacked on top of each other. These mountain-building processes thickened Earth’s crust, and increased its mass locally as the mountains grew. The increased mass of the growing mountains caused the lithosphere to float lower in the mantle (Figure 10.39, left). As the mountains were eventually eroded over tens of millions of years, the crust floated higher and higher in the mantle, and erosion exposed metamorphic rocks that were deep within the mountains (Figure 10.39, right).

Figure 10.39 Schematic cross-section through the Meguma Terrane. Left- Metamorphic zones and temperatures when mountain-building processes thickened the crust. Right- The mountains have been eroded, exposing metamorphic rocks that formed deep within the mountains. Source: Karla Panchuk (2018) CC BY 4.0, modified after Steven Earle (2015) CC BY 4.0 view source left/ right.

Building a narrative for the metamorphism in Nova Scotia’s Meguma Terrane is just one example of how index minerals can be used.

Exercise: Scottish Metamorphic Zones

The map in Figure 10.40 shows 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.

1. Label the three coloured areas of the map with the appropriate zone names (garnet, chlorite, and biotite).
2. 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 metamorphism.

Figure 10.40 Metamorphic zones in Barrovian metamorphism. Source: Steven Earle (2015) CC BY 4.0 view source

References

Keppie, D., & Muecke, G. (1979). Metamorphic map of Nova Scotia. (Nova Scotia Department of Mines and Energy, Map 1979-006).

White, C. E., & Barr, S. M. (2012) Meguma Terrane Revisted: Stratigraphy, Metamorphism, Paleontology and Provenance. Geoscience Canada 39(1). Full text

The heat from a body of magma in the upper crust can create a very dynamic situation with geologically interesting and economically important implications. In the simplest cases, water does not play a big role, and the main process is heat transfer from the pluton to the surrounding rock, creating a zone of contact metamorphism (Figure 10.41a). In many cases, however, water is released from the magma body as crystallization takes place, and this water is dispersed along fractures in the surrounding rock (Figure 10.41b). The water released from a magma chamber is typically rich in dissolved minerals. As this water cools, it interacts with the surrounding rocks, changing both the chemistry of the water and the chemistry of the rocks. This can cause minerals to precipitate from the water. Minerals can also precipitate if the water boils because of a drop in pressure. The precipitated minerals form veins within fractures in the surrounding rock. Quartz veins are commonly formed in this situation, and can include other minerals such as pyrite, hematite, calcite, and even silver and gold.

Figure 10.41 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. Source: Steven Earle (2015) CC BY 4.0 view source

Heat from the magma body will cause surrounding groundwater 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 thousands of years, resulting in the circulation of millions of tonnes of groundwater from the surrounding region past the pluton.

Hot water circulating through the rocks and interacting chemically with them 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 10.42). Chemical change in rocks due to interaction with hot water is called hydrothermal alteration.

Figure 10.42 White veins of calcite in limestone of the Comox Formation, Nanaimo BC. Quarter for scale. Source: Steven Earle (2016) CC BY 4.0 view source

Metamorphic reactions involve the release of fluids as minerals change, and chemical reactions with locally-derived fluids. However, if a large amount of externally-derived fluid—such that supplied by magma—is flushed through the system at the high pressures and temperatures characteristic of metamorphism, it can substantially alter the chemical composition of the rock. This type of hydrothermal alteration is called metasomatism.

A special type of metasomatism takes place where a hot pluton intrudes into carbonate rock such as limestone. Magmatic fluids rich in silica, calcium, magnesium, iron, and other elements can dramatically change the chemistry of the limestone, forming minerals that would not normally exist in either the igneous rock or limestone. A rock called skarn results, containing minerals such as garnet, epidote, magnetite, and pyroxene, among others (Figure 10.43).

Figure 10.43 Skarn from Mount Monzoni, Northern Italy, with recrystallized calcite (blue), garnet (brown), and pyroxene (green). The rock is 6 cm across. Source: Siim Sepp (2012) CC BY-SA 3.0 view source

Exercise: Contact Metamorphism and Metasomatism

A pluton that has intruded into a series of sedimentary rocks, including sandstone, mudstone, and limestone (Figure 10.44). What types of metamorphic rocks would you expect to see at locations a, b, and c?

Figure 10.44 Contact metamorphism and metasomatism of sedimentary rocks. Source: Steven Earle (2015) CC BY 4.0 view source

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

10.1 Controls on 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 direction of pressure, the volumes and compositions of fluids that are present, and the amount of time available for metamorphic reactions to take place.

10.2 Foliation and Rock Cleavage

When the pressure acting on a rock is not uniform in all directions, foliation can develop. Foliation may occur in the form of platy or elongated mineral crystals that have grown at right angles to the maximum pressure, or it may develop when crystals or clasts within a rock are deformed. Foliation causes crystals or clasts within a rock to become aligned. When metamorphic rocks break parallel to the direction of foliation, rock cleavage results.

10.3 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.

10.4 Types of Metamorphism and Where They Occur

Almost all regions that experience metamorphism are being acted upon by plate-tectonic processes. Oceanic crustal rock can be metamorphosed near the spreading ridge where it was formed. Regional metamorphism takes place in areas where mountain ranges are forming, which are most common at convergent boundaries. Contact metamorphism takes place around magma bodies in the crust, which are also most common above convergent boundaries. Shock metamorphism happens when extraterrestrial bodies impact Earth, and is unusual among metamorphic processes because it occurs in seconds or minutes, rather than taking millions of years. Dynamic metamorphism occurs when shear stress is applied to rocks, such as along faults.

10.5 Metamorphic Facies and Index Minerals

Metamorphic facies are groups of metamorphic rocks that form under the same range of pressure and temperature conditions, but from different parent rocks. Geologists use index minerals such as chlorite, garnet, andalusite, and sillimanite to identify metamorphic zones. Index minerals tell us about the pressure and temperature conditions under which metamorphic rocks formed.

10.6 Metamorphic Hydrothermal Processes and Metasomatism

Contact metamorphism takes place around magma bodies that have intruded into cool rocks in the crust. Heat from magma is transferred to the surrounding country rock, resulting in mineralogical and textural changes. Hot water from a cooling body of magma, or from convection of groundwater driven by the heat of the pluton, can lead to hydrothermal alteration. When large volumes of fluid are flushed through rocks experiencing metamorphic pressures and temperatures, metasomatism results. Metasomatism can cause valuable metals to accumulate in the surrounding rocks.

Review Questions

1. What are the two main agents of metamorphism, and what are their respective roles in producing metamorphic rocks?
2. What types of metamorphic rocks will form if a mudrock experiences very low, low, medium, and high-grade metamorphism?
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 particular temperature and pressure characteristics of this geological setting?
8. Rearrange the following minerals in order of increasing metamorphic grade: biotite, garnet, sillimanite, chlorite.
9. What is the role of magmatic fluids in the metamorphism that takes place adjacent to a pluton?
10. How does metasomatism differ from regional metamorphism?
11. How does the presence of a hot pluton contribute to metasomatism?
12. What determines whether metasomatism will produce skarn?
13. For each of the following metamorphic rocks, indicate the likely parent rock and the grade and/or type of metamorphism: chlorite schist, slate, mica-garnet schist, amphibolite, marble.

1. Heat and pressure are the main agents of metamorphism. Heat leads to mineralogical changes in the rock. Pressure also influences those mineralogical changes, while directed pressure (greater pressure in one direction) leads to foliation.
2. Very low grade: slate; low grade: phyllite; medium grade: schist; high grade: gneiss.
3. Granite remains largely unchanged at lower metamorphic grades because its minerals are still stable at those lower temperatures.
4. Foliation develops in schist when new platy minerals grow with their longest dimension at a right angle to the direction of greatest pressure.
5. At a spreading ridge the heat from volcanism leads to the development of a groundwater convection system in the rock of the oceanic crust. Heated water rises in the hot regions and is expelled into the ocean, while cold ocean water is drawn into the crust to replace it. The heated water leads to the conversion of olivine and pyroxene into chlorite and serpentine.
6. The geothermal gradient varies as a function of tectonic setting, being greatest in volcanic regions and lowest along subduction zones. As a result the depth at which specific metamorphic grades is achieved will vary: the depth will be greater when the gradient is lower.
7. The geothermal gradient is low within subduction zones because the cold subducting oceanic crust takes a long time to heat up. Pressure increases with depth at the normal rate, but temperature does not.
8. Order of increasing metamorphic grade: chlorite, biotite, garnet, sillimanite.
9. Water from any source facilitates metamorphism. Magmatic fluids typically contain dissolved ions at higher concentrations than in regular groundwater (especially copper, zinc, silver, gold, lithium, beryllium, boron and fluorine), leading to the formation of a unique set of minerals.
10. Metasomatism involves fluids from magmatic or groundwater sources that play an important role in transporting ions into the system, and leading to the formation of new minerals. Regional metamorphism takes place over a larger area, depends more on plate tectonic conditions, and does not involve flushing the system with large amounts of fluid.
11. A hot pluton heats the surrounding water, causing groundwater to convect. This can result a great deal of water, in some cases with elevated levels of specific ions, passing through the rock. Water from magma within the pluton also contributes to metasomatism.
12. Limestone must be present to produce skarn.
13. Parent rocks and metamorphic grades and types:
    

    Metamorphic Rock	Likely Parent Rock	Grade and/or Type of Metamorphism
    Chlorite schist	A rock enriched in ferromagnesian minerals, such as basalt	Low-grade regional metamorphism
    Slate	Mudrock (shale, mudstone)	Very low grade regional metamorphism
    Mica-garnet schist	A rock that is rich in aluminum, which includes most clay-bearing rocks	Medium-grade regional metamorphism
    Amphibolite	A rock enriched in ferromagnesian minerals, such as basalt	Medium- to high-grade regional metamorphism
    Marble	Limestone or dolomite	Regional or contact metamorphism

Volcanoes Are Where Magma Emerges

A volcano is a location where molten rock flows out, or erupts, onto Earth’s surface as lava. Volcanic eruptions can happen on land or underwater. Some volcanic eruptions flow from mountains (such as Mount Garibaldi in Figure 11.1), but others do not. Fissure eruptions (Figure 11.3) are volcanic eruptions flowing from long cracks in the Earth.

Figure 11.3 Kamoamoa fissure eruption on the flanks of the Hawai’ian Kīlauea Volcano in March of 2011. Source: Karla Panchuk (2017) CC BY-SA 4.0. Photograph: U. S. Geological Survey (2011) Public Domain view source. Click the image for more attributions.

Volcano Anatomy

The main parts of a volcano are shown in Figure 11.4. When volcanoes erupt, magma moves upward from a magma chamber and into a vent or conduit. It flows out from a crater at the top, or sometimes emerges at a secondary site on the side of the volcano resulting in a flank eruption. Erupted materials accumulate around the vent forming a volcanic mountain. The accumulated material might consist of layers of solidified lava, called lava flows, but it might also include fragments of various sizes that have been thrown from the volcano.

Figure 11.4 The parts of a volcano (not to scale). Source: Karla Panchuk (2017) CC BY 4.0

Crater or Caldera?

A crater is the basin above a volcano’s vent. Craters have diameters on the scale of 10s to 100s of metres. A caldera is a bowl-shaped structure that resembles a crater, but it is much larger (km in scale) and forms when a volcano collapses in on itself. The process is illustrated in Figure 11.5, going from left to right. It begins when an eruption occurs, and the magma chamber beneath the volcano is drained. If a significant part of a volcano’s mass is supported by magma within the chamber, then depleting the magma also reduces the support for the volcano. The loss of support causes part of the volcano to collapse into the void in the magma chamber, leaving behind a broad basin rimmed by the remnants of the volcano. Over time, the basin can fill with water. If there is still activity within the magma chamber, magma may force its way upward again, causing the floor of the caldera to be lifted, or erupting to form a new volcano within the caldera.

Figure 11.5 Formation of a caldera. Calderas are the result of a volcano collapsing into a drained magma chamber. Source: Karla Panchuk (2017) CC BY 4.0. Modified after U. S. Geological Survey (2002) Public Domain view source

The island of Santorini (Figure 11.6) is an example of a caldera. The island itself is the rim of the caldera, and the bay is the flooded basin. The two small islands in the middle of the bay formed from magma refilling the chamber that feeds the volcano, as in the far right of Figure 11.5. The caldera formed after an enormous eruption between 1627 and 1600 BCEFriedrich, W. L., Kromer, B., Friedrich, M., Heinemeier, J., Pfeiffer, T., & Talamo, S. (2006). Santorini Eruption Radiocarbon Dated 1627-1600 B.C. Science (312)5773, 548. doi: 10.1126/science.1125087. The eruption is thought to have contributed to the downfall of the Minoan civilization, and some speculate that it might also be the source of the myth of Atlantis, a story about a lost continent that sank beneath the sea after a natural disaster.

Figure 11.6 The Greek Island of Santorini. Left: Aerial view of the island forming a ring around a flooded caldera. Right: A view from the rim of the caldera. The other side of the rim is visible in the distance. Source: Karla Panchuk (2017) CC BY-SA 4.0. Satellite image: NASA/GSFC/MITI/ERSDAC/JAROS, and U.S./Japan ASTER Science Team (2000) Public Domain view source; Caldera photograph: Klearchos Kapoutsis (2010) CC BY 2.0 view source Click the image for more attributions.

Volcanic eruptions produce three types of materials: gas, lava, and fragmented debris called tephra.

Volcanic Gas

Magma contains gas. At high pressures, the gases are dissolved within magma. However, if the pressure decreases, the gas comes out of solution, forming bubbles. This process is analogous to what happens when a pop bottle is opened. Pop is bottled under pressure, forcing carbon dioxide gas to dissolve into the fluid. As a result, a bottle of pop that you find on the supermarket shelf will have few to no bubbles. If you open the bottle, you decrease the pressure within it. The pop will begin to fizz as carbon dioxide gas comes out of solution and forms bubbles.

The main component of volcanic gas emissions is water vapour, followed by carbon dioxide (CO₂), sulphur dioxide (SO₂), and hydrogen sulphide (H₂S).

Volcanoes release gases when erupt, and through openings called fumaroles (Figure 11.7). They can also release gas into soil and groundwater.

Figure 11.7 A fumarole at Puʻu ʻŌʻō Crater, Hawaii. The yellow crust along the margin of the fumarole is made of sulphur crystals. The crystals form when sulphur vapour cools as it is released from the fumarole. Source: U. S. Geological Survey (2016) Public Domain View source

Lava

The ease with which lava flows and the structures it forms depend on how much silica and gas the lava contains. The more silica, the more polymerization (formation of long molecules) occurs, stiffening the lava. The stiffness of lava is described in terms of viscosity– lava that flows easily has low viscosity, and lava that is sticky and stiff has high viscosity.

In general, high-silica lava contains more gas than low-silica lava. When the gas forms into bubbles, viscosity increases further. Consider the pop analogy again. If you were to shake the bottle vigorously and then open it, the pop would come gushing out in a thick, frothy flow. In contrast, if you took care to not shake the bottle before opening it, you could pour out a thin stream of fluid.

Chemical Composition Affects the Thickness and Shape of Lava Flows

The thickness and shape of a lava flow depends on its viscosity. The greater the viscosity, the thicker the flow, and the shorter the distance it travels before solidifying. Highly viscous lava might not flow very far at all, and simply accumulate as a bulge, called a lava dome, in a volcano’s crater. Figure 11.8 shows a dome formed from rhyolitic lava in the crater of Mt. St. Helens.

Figure 11.8 Lava dome in the crater of Mt. St. Helens. Source: Terry Feuerborn (2011) CC BY-NC 2.0 view source

Less viscous rhyolitic lava can travel further, as with the thick flow in Figure 11.9 (right). The left of Figure 11.9 shows thin streams of freely-flowing, low-silica, low-viscosity basaltic lava.

Figure 11.9 Lava flows. Left: A geologist collects a sample from a basaltic lava flow in Hawaii. Right: an andesitic lava flow from Kanaga Volcano in the Aleutian Islands. Source: Left- U. S. Geological Survey (2014) Public Domain view source; Right- Michelle Combs, U. S. Geological Survey (2015) Public Domain view source

Low-viscosity basaltic lava flows may travel extended distances if they move through conduits called lava tubes. These are tunnels within older solidified lava flows. Figure 11.10 (top) shows a view into a lava tube through a hole in the overlying rock, called a skylight. Figure 11.10 (bottom) shows the interior of a lava tube, with a person for scale. Lava tubes form naturally and readily because flowing mafic lava preferentially cools near its margins, forming solid lava levées that eventually close over the top of the flow. Lava within tubes can flow for 10s of km because the tubes insulate the lava from the atmosphere and slow the rate at which the lava cools. The Hawai’ian volcanoes are riddled with thousands of old, drained lava tubes, some as long as 50 km.

Figure 11.10 Lava tubes. Top: An opening in the roof of a lava tube (called a skylight) permitting a view of lava flowing through the tube (Puʻu ʻŌʻō crater, Kīlauea). The opening is approximately 6 m across. Bottom: Inside a lava tube that channelled lava away from Mt. St. Helens in an eruption 1,895 years ago. Sources: Top: U. S. Geological Survey (2016) Public Domain. view source Bottom: Thomas Shahan (2013) CC BY-NC 2.0 view source

Lava Structures

Pahoehoe

Lava flowing on the surface can take on different shapes as it cools. Basaltic lava with an unfragmented surface, like that in Figure 11.9 (right), is called pahoehoe. (pronounced pa-hoy-hoy). Pahoehoe can be smooth and billowy. It can also develop a wrinkled texture, called ropy lava, as shown in Figure 11.11. Ropy lava forms when the outermost layer of the lava cools and develops a skin (visible as a dark layer in Figure 11.11, left), but the skin is still hot and thin enough to be flexible. The skin is stiffer than the lava beneath it, and is dragged by flowing lava and folded up into wrinkles. Figure 11.11 (right) is a close-up view after a cut has been made to show the internal structure of a wrinkled lava flow. Notice the many holes, or vesicles, within the lava, formed when the lava solidified around gas bubbles.

Figure 11.11 Ropy lava from Hawaii. Left: Ropy texture forming as a thin surface layer of lava cools and is wrinkled by the motion of lava flowing beneath it. Right: Cross-section view of ropy lava. Sources: Left: Z. T. Jackson (2005) CC BY NC-ND 2.0 view source; Right: Fiddledydee (2011) CC BY-NC 2.0 view source.

A’a and Blocky Lava

If the outer layer of the lava flow cannot accommodate the motion of lava beneath by deforming smoothly, the outer layer will break into fragments as lava moves beneath it. This could happen if the lava flow develops a thicker, more brittle outer layer, or if it moves faster. The result is a sharp and splintery rubble-like lava flow called a’a (pronounced like “lava” but without the l and v). Figure 11.12 (left) shows a close-up view of the advancing front of an a’a lava flow (the flow is moving toward the viewer). Figure 11.12 (right) shows an a’a lava flow viewed from the side. Compare the texture of the a’a flow with the texture of the lighter-grey pahoehoe lava in the foreground of the picture.

Figure 11.12 Aa lava flows. Left: Close-up view of a’a forming during an eruption of Pacaya Volcano in Guatemala. Field of view approximately 1 m across. Right: Rubbly reddish-brown a’a lava flow viewed from Chain of Craters Road, Hawai’i Volcanoes National Park. Pahoehoe is visible in lighter grey in the foreground. Sources: Photo of Hawaiian aa and pahoehoe: Roy Luck (2009) CC BY 2.0 view source; Pacaya aa: Greg Willis (2008) CC BY-SA 2.0 (labels added) view source.

Higher viscosity andesitic lava flows also develop a fragmented surface, called blocky lava. This is visible in the toe of the andesitic lava flow from Figure 11.9 (right). The difference between a’a and the andesitic blocky lava is that the blocky lava has fragments with smoother surfaces and fewer vesicles.

Lava Pillows

When lava flows into water, the outside of the lava cools quickly, making a tube (Figure 11.13 (top left)). Blobs of lava develop at the end of the tube (Figure 11.13 (top right)), forming pillows. The bottom left of Figure 10.13 shows pillows covering the sea floor, and the bottom right shows the distinctive rounded shape of pillows in outcrop. Because pillows always form underwater, finding them in the rock record gives us information that the environment was underwater.

Figure 11.13 Pillow lavas. Top left: A tube of lava extruding underwater. Hot lava can be seen through cracks in the wall of the tube. The image is approximately 1 m across. (Pacific Ocean, near Fiji). Top right: The rounded end of a tube with cracks showing the lava within. (Pacific Ocean, near Fiji). Bottom left: sea floor near the Galápagos Islands covered with pillow lavas. Bottom right: A boulder made of 2.7 billion year old pillows derived from the Ely Greenstone in north-eastern Minnesota. Sources: Top left- NSF and NOAA (2010) CC BY 2.0 view source; Top right- NSF and NOAA (2010) CC BY 2.0 view source; Bottom left- NOAA Okeanos Explorer Program, Galápagos Rift Expedition 2011 (2011) CC BY 2.0 view source; Bottom right- James St. John (2015) CC BY 2.0 view source.

Columnar Joints

When lava flows cool and solidify, they shrink. Long vertical cracks, or joints, form within the brittle rock to allow for the shrinkage. Viewed from above, the joints form polygons with 5, 6, or 7- sides, and angles of approximately 120º between sides (Figure 11.14).

Figure 11.14 Columnar joints viewed from above, Giant’s Causeway, Northern Ireland. Source: Meg Stewart (2012) CC BY-SA 2.0 view source

Figure 11.15 shows a side view of columnar joints in a basaltic lava flow in Iceland.

Figure 11.15 Columnar joints in a basaltic lava flow, Svartifoss (Black Fall) Vatnajökull National Park, Iceland. Source: Ron Kroetz (2015) CC BY-ND 2.0. view source

Pyroclastic Materials

The pop bottle analogy illustrates another key point about gas bubbles in fluid, which is that the bubbles can propel fluid. In the same way that shaking a pop bottle to make more bubbles will cause pop to gush out when the bottle is opened, gas bubbles can violently propel lava and other materials from a volcano, creating an explosive eruption.

Collectively, loose material thrown from a volcano is referred to as tephra. Individual fragments are referred to in general terms as pyroclasts, so sometimes tephra is also referred to as pyroclastic debris. Pyroclasts are classified according to size.

Volcanic Ash

Particles less than 2 mm in diameter are called volcanic ash. Volcanic ash consists of small mineral grains and glass. Figure 11.16 shows volcanic ash on three scales: in the upper left is ash from the 2010 eruption of Eyjafjallajökull in Iceland. The image was taken with a scanning electron microscope at approximately 1000 times magnification. In the upper right is ash from the 1980 eruption of Mt. St. Helens, collected in Yakima, Washington, about 137 km northeast of Mt. St. Helens. Individual particles are under 1 mm in size. Figure 11.16 (bottom) shows a village near Mt. Merapi in Indonesia dusted in ash after an eruption 2010.

Figure 11.16 Volcanic ash. Upper left: Ash from 2010 eruption of Eyjafjallajökull in Iceland, magnified approximately 1000x. Upper right- Ash from the 1980 eruption of Mt. St. Helens, collected at Yakima, Washington. Bottom: Indonesian village after the eruption of Mt. Merapi in 2010. Sources: Upper left: Birgit Hartinger, AEC (2010) CC BY-NC-ND 2.0. view source Upper right: James St. John (2014) CC BY 2.0 (scale added) view source Bottom: AusAID/Jeong Park (2010) CC BY 2.0. view source  

Lapilli

Fragments with dimensions between 2 mm and 64 mm are classified as lapilli. Figure 11.17 (upper left) shows lapilli at the ancient city of Pompeii, which was buried when Mt. Vesuvius erupted in 79 C.E. Figure 11.17 (lower left) is a form of lapilli called Pele’s tears, named after the Hawai’ian diety Pele. Pele’s tears form when droplets of lava cool quickly as they are flung through the air. Rapidly moving through the air may draw the Pele’s tears out into long threads called Pele’s hair (Figure 11.17, right). The dark masses in Figure 11.17 (right) within the Pele’s hair are Pele’s tears.

Figure 11.17 Lapilli are pyroclasts ranging between 2 mm and 64 mm in size. Upper left: lapilli from the site of the ancient city of Pompeii. Lower left: Pele’s tears, a type of lapilli that forms when droplets of lava fly through the air. Right: Pele’s hair, which form when Pele’s tears are drawn out into thin threads as they fly. Sources: Upper left: Pauline (2009) CC BY-NC-ND 2.0 view source; Lower left: James St. John (2014) CC BY 2.0 (scale added) view source; Right: James St. John (2009) CC BY 2.0 (scale added) view source.

Blocks and Bombs

Fragments larger than 64 mm are classified as blocks or bombs, depending on their origin. Blocks are solid fragments of the volcano that form when an explosive eruption shatters the pre-existing rocks. Figure 11.18 shows one of many blocks from an explosive eruption at the Halema‘uma‘u crater at Kīlauea Volcano in May of 1924. The block has a mass of approximately 7 tonnes and landed 1 km from the crater.

Figure 11.18 Volcanic block weighing approximately 7 tonnes thrown 1 km from the Halema‘uma‘u crater at Kīlauea Volcano on May 18, 1924. Source: U. S. Geological Survey (1924) Public Domain view source

Bombs form when lava is thrown from the volcano and cools as it travels through the air. Traveling through the air may cause the lava to take on a streamlined shape, as with the example in Figure 11.19.

Figure 11.19 Volcanic bomb with a streamlined shape. Source: James St. John (2016) CC BY 2.0 (scale added) view source

Effects of Gas on Lapilli and Bombs

The presence of gas in erupting lava can cause lapilli and bombs to take on distinctive forms as the lava freezes around the gas bubbles, giving the rocks a vesicular (hole-filled) texture. Pumice (Figure 11.20) forms from gas-filled felsic lava. Figure 11.20 (right), shows a magnified view of the sample on the left. The dark patches in the photograph are mineral crystals that formed in the magma chamber before the lava erupted. Pumice floats on water because some of the holes are completely enclosed, and air-filled.

Figure 11.20 Lapilli-sized pumice collected from the shores of Lake Atitlán in Guatemala by H. Herrmann. The lake is a flooded caldera, and is surrounded by active volcanoes. Right: Magnified view showing vesicular structure and amphibole crystals (dark patches). Source: Karla Panchuk (2017) CC BY 4.0

The mafic counterpart to pumice is scoria (Figure 11.21, left). Mafic lava can also form reticulite (Figure 11.21, right), a rare and fragile rock in which the walls surrounding the bubbles have all burst, leaving behind a delicate network of glass.

Figure 11.21 Mafic lapilli with vesicular textures. Left: Scoria from Mount Fuji, Japan. Scoria is the denser mafic counterpart to pumice. Right: Reticulite from Kīlauea Volcano. Reticulite is a delicate network of volcanic glass that forms when the walls separating gas bubbles pop. Sources: Left- James St. John (2014) CC BY 2.0 (scale added) view source; Right- James St. John (2014) CC BY 4.0 (scale added) view source.

References

U. S. Geological Survey (2013) Mt. St. Helens National Volcanic Monument. Retrieved on 11 June 2017. Visit website

The products of volcanism that build volcanoes and leave lasting marks on the landscape include lava flows that vary in viscosity and gas content, and tephra ranging in size from less than a mm to blocks with masses of many tonnes. Individual volcanoes vary in the volcanic materials they produce, and this affects the size, shape, and structure of the volcano.

There are three types of volcanoes: cinder cones (also called spatter cones), composite volcanoes (also called stratovolcanoes), and shield volcanoes. Figure 11.22 illustrates the size and shape differences amongst these volcanoes.

Shield volcanoes, which get their name from their broad rounded shape, are the largest. Figure 11.22 shows the largest of all shield volcanoes- in fact, the largest of all volcanoes on Earth- Mauna Loa, which makes up a substantial part of the Island of Hawai‘i and has a diameter of nearly 200 km. The summit of Mauna Loa is presently 4,169 m above sea level, but this represents only a small part of the volcano. It rises up from the ocean floor at a depth of approximately 5,000 m. Furthermore, the great mass of the volcano has caused it to sag downward into the mantle by an additional 8,000 m. In total, Mauna Loa is a 17,170 m thick accumulation of rock.

Figure 11.22 Comparison of volcano sizes and shapes. Broad, rounded shield volcanoes are the largest, followed by cone-shaped composite volcanoes. Straight-sided cinder cones are the smallest, and barely visible in the scale of the diagram. Source: Karla Panchuk (2017) CC BY 4.0 modified after Steven Earle (2015) CC BY 4.0 view original

Kīlauea Volcano is also a shield volcano, albeit a much flatter one. Kīlauea Volcano rises only 18 m about the surrounding terrain, and is almost not visible in the scale of the diagram, however it still stretches over a distance of 125 km along the eastern side of the Island of Hawai‘i.

Composite volcanoes are the next largest. Mt. St. Helens is shown on the left of Figure 11.22. It rises 1,356 m above the surrounding terrain in the Cascade Range of the western United States, and has a diameter of approximately 6 km. Composite volcanoes tend to be no more than 10 km in diameter. Unlike shield volcanoes, composite volcanoes have a distinctly conical shape, with sides that steepen toward the summit.

Cinder cones are the smallest, and almost too small to see next to a volcano like Mauna Loa. Eve Cone is a cinder cone on the flanks of Mt. Edziza in northwestern British Columbia. It rises 172 m above the landscape, and has a diameter of under 500 m. Cinder cones have straight sides, unlike upward-steepening composite volcanoes, or rounded shield volcanoes.

Volcano Structure

Shield Volcanoes

Shield volcanoes, like the Sierra Negra volcano in the Galápagos Islands (Figure 11.23, top), get their gentle hill-like shape because they are built of successive flows of low-viscosity basaltic lava (Figure 11.23, bottom). The low viscosity of the lava means that it can flow for long distances, resulting in the greater size of shield volcanoes compared to composite volcanoes or cinder cones.

Figure 11.23 Shield volcano. Top: The Sierra Negra volcano in the Galápagos Islands exhibits the low, rounded shape characteristic of shield volcanoes. Bottom: Diagram of a shield volcano island, showing the build up of basaltic lava flows. Sources: Top- BRJ INC. (2012) CC BY-NC-ND 2.0 view source. Bottom- Karla Panchuk (2017) CC BY 4.0

Composite Volcanoes (Stratovolcanoes)

Composite volcanoes, like Cotopaxi in Figure 11.24 (top), consist of layers of lava alternating with layers of tephra (blocks, bombs, lapilli, and ash; Figure 11.24, bottom). The layers (strata) is where the alternative name, stratovolcano comes from. Cotopaxi displays the characteristic shape of composite volcanoes, which have slopes that get steeper near the top of the volcano. The change in the slope reflects the accumulation of tephra fragments near the volcano’s vent. Composite volcanoes typically erupt higher viscosity andesitic and rhyolitic lavas, which do not travel as far from the vent as basaltic lavas do. This results in volcanoes of smaller diameter than shield volcanoes. A notable exception is Mt. Fuji in Japan, which erupts basaltic lava.

Figure 11.24 Composite volcano. Top: Cotopaxi in Ecuador exhibits the upward-steepening cone characteristic of composite volcanoes. Bottom: Diagram of a composite volcano showing alternating layers of lava and tephra. Sources: Karla Panchuk (2017) CC BY 4.0; Top photo by Simon Matzinger (2014) CC BY 2.0 view source. Click the image for more attributions. 

From a geological perspective, composite volcanoes tend to form relatively quickly and do not last very long. If 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.

Cinder Cones (Spatter Cones)

Cinder cones, like Mt. Capulin in Figure 11.25, have straight sides and are typically less than 200 m high. Most are made up of fragments of scoria (vesicular rock from basaltic lava) that were expelled from the volcano as gas-rich magma erupted. Because cinder cones are made up almost exclusively of loose fragments, they have very little strength. They can be eroded away easily, and relatively quickly.

Figure 11.25 Cinder cone. These small, straight-sided volcanoes are made of volcanic fragments ejected when gas-rich basaltic lava erupts. Sources: Karla Panchuk (2017) CC BY 4.0, with photograph by R. D. Miller, U. S. Geological Survey (1980) Public Domain view source. Click the image for more attributions.

References

Rubin, K. (n.d.) Mauna Loa Volcano. Retrieved 23 August 2017. Visit website

Volcanoes produce a variety of materials when they erupt. The characteristics of the eruptions themselves also vary from one volcano to the next, and sometimes from one eruption to the next for the same volcano. Eruptions are classified according how explosive they are, and the height of their eruption column– how high they blast material into the air.

Both the explosiveness of an eruption and the height of the eruption column are related in part to the composition of magma and the amount of gas it contains. In particular, magmas with more silica erupt more explosively. The higher the silica content, the greater the viscosity of the magma. This means more pressure can build up before the magma is forced out of the volcano. Magma with more silica also tends to contain more dissolved gas. The gas helps to propel magma out of the volcano, in the same way that the bubbles in a shaken bottle of pop cause the pop to foam out when the lid is removed.

There are four types of eruptions with properties determined mostly by the silica content of magma, and the amount of gas it contains. In order of increasing explosiveness, these are Hawai’ian, Strombolian, Vulcanian, and Plinian eruptions. Any composition of magma can have an explosive eruption if the magma suddenly encounters water. Hot magma contacting groundwater or seawater causes the water to flash to steam. Explosive eruptions driven by water are called hydrovolcanic (or phreatic) eruptions.

Hawai‘ian Eruptions

Hawai‘ian eruptions are named after the characteristic eruptions of volcanoes of the Hawai‘ian islands. Hawai‘ian eruptions are effusive (flowing) rather than explosive because they erupt low-viscosity basaltic lava. Hawai‘ian eruptions form shield volcanoes and can also take the form of fissure eruptions. Fissure eruptions occur when lava erupts from long cracks in the ground rather than from a central vent.

Figure 11.26 shows examples from two eruptions on of Hawai‘i. In the upper left and right are images from the November 1959 eruption of Kīlauea Iki Crater. The upper left shows a fissure eruption and effusive flow of lava. Burning trees appear as bright spots toward the bottom of the photo. Figure 11.26 (right) shows a lava fountain reaching 425 m above Kīlauea Iki Crater. U. S. Geological Survey scientists reported that volcanic bombs up to 60 cm across smashed the guard rail and dented the asphalt on the road. Figure 11.26 (lower left) shows Hawaiian Volcano Observatory (HVO) scientists making a quick getaway, with lava fountains from Mauna Loa Volcano in the background.

Figure 11.26 Hawai‘ian eruptions. Top left: Fissure eruption at Kīlauea Iki Crater in November of 1959. Bottom left: Lava fountains from an eruption of Mauna Loa Volcano in 1984. Right: Lava fountain from Kīlauea Iki Crater eruption in November of 1959. Sources: Top left- U. S. Geological Survey (1959) Public Domain. view source Bottom left: R. B. Moore, U. S. Geological Survey (1984) Public Domain. view source Right- U. S. Geological Survey (1959) Public Domain. view source    

The photographs of the Kīlauea Iki Crater and Mauna Loa Volcano eruptions make the point that while Hawai‘ian eruptions are considered “gentle” eruptions, this is a relative term. “Gentle” eruptions range from lava flows that can be safely sampled by trained personnel, as in Figure 11.5, to lava fountains that soar hundreds of metres above the tree tops and rain large and dangerous rocks upon the surroundings.

Strombolian Eruptions

Strombolian eruptions, named for Mt. Stromboli in Italy, occur when basaltic lava has higher viscosity and higher gas content. The sticky lava is ejected in loud, violent, but short-lived spattery eruptions. Clumps of gas-rich lava thrown 10s to 100s of metres in the air accumulate as scoria in a pile around the vent, forming cinder cones. Figure 11.27 shows a strombolian eruption in the crater of Mt. Etna. A smaller cinder cone is forming around the vent as lava sputters out of it.

Figure 11.27 Strombolian eruption of Mt. Etna. Sputtering lava forms a smaller cinder cone around a vent within the crater of Etna. Source: Karla Panchuk (2017) CC BY-SA 4.0. Photograph- Robin Wylie (2012) CC BY 2.0. view source Click the image for more attributions.  

Vulcanian Eruptions

Vulcanian eruptions get their name from the volcanic Italian island of Vulcano, which itself takes the name of the Roman god of fire, Vulcan. In Roman mythology, Vulcan was the maker of armour and weaponry for the gods, and volcanic eruptions were attributed to him working in his forge.

Vulcanian eruptions are far more explosive than Strombolian eruptions, and can blast tephra and gas to a height of 5 to 10 km. The explosiveness is related to a build-up of pressure as the higher viscosity of intermediate silica content lava restricts the escape of gas. Vulcanian eruptions produce large quantities of ash in addition to blocks and bombs.

The Vulcanian eruption of Mt. Pelée on the island of Martinique in 1902 resulted in the first detailed documentation by geologists of a devastating phenomenon that is now referred to as a pyroclastic flow (Figure 11.28). Volcanic debris from the collapse of a lava dome on Mt. Pelée combined with hot gas to form a searing avalanche that raced down the mountain, over the city of St. Pierre, and into the harbour.

Figure 11.28 A series of photos taken by Alfred Lacroix during the eruption of Mt. Pelée on May 8, 1902 showing the development of the pyroclastic flow that destroyed the city of St. Pierre and nearly 30,000 inhabitants. Source: Karla Panchuk (2017) CC BY 4.0. Photograph: Alfred Lacroix (1902) Public Domain. view original Click the image for more attributions.

The French geologist Alfred Lacroix described what he saw as a “nuée ardente,” or thick fiery cloud. The following first-hand account was published in Cosmopolitan Magazine in July of 1902, attributed to Ellery S. Scott, a sailor on the steamship Roraima:

Figure 11.29 Two stereographs of the ruins of St. Pierre, published in 1902. Stereographs are viewed with a stereoscope to make an image appear three dimensional. Top- “St. Pierre, ‘the city of dead,’ Mt. Pelee smoking, Martinique”; Bottom- “Overlooking the mud-filled Roxelane River bed, and ash-covered ruins, to Mont Pelée, St. Pierre, Martinique.” Source: Top- Boston Public Library (2013) CC BY 2.0 view source; Bottom- Boston Public Library (2013) CC BY 2.0 view source

Plinian Eruptions

Plinian eruptions are explosive eruptions of intermediate to felsic lava, and can form eruptive columns up to 45 km high. The origin of the name is the eruption of Vesuvius in 79 CE, which buried the towns of Pompeii and Herculaneum. The Roman admiral Gaius Plinius Secundus, also known as Pliny the Elder, attempted a rescue mission when he saw the column of ash and debris above Vesuvius, but died of unknown causes without being able to reach Herculaneum.

A more recent Plinian eruption was that of Mt. Redoubt on April 21, 1990, shown in Figure 11.30. Pyroclastic flows resulted, as did lahars, landslides that formed when glaciers melted and turned volcanic ash into mud. The shape of the eruptive column, with parts of the column appearing to spread out in flat layers at different levels, reflects differences in atmospheric characteristics.

Figure 11.30 Plinian eruption of Mt. Redoubt in Alaska on April 21, 1990. Source: Karla Panchuk (2017) CC BY 4.0. Photograph: R. Clucas, U. S. Geological Survey (1990) Public Domain. view source Click the image for more attributions.

Hydrovolcanic (Phreatic) Eruptions

Hydrovolcanic eruptions can be far more explosive than Plinian eruptions. They occur when water in the form of groundwater, seawater, or even melting glacial ice or snow comes into contact with magma. Heat from the magma changes water suddenly to steam, which can expand to more than a thousand times the original volume of water. The sudden expansion results in an explosive force that can blast a volcano to pieces and create large amounts of volcanic ash.

In April of 2010, activity by the Icelandic volcano Eyjafjallajökull (Figure 11.31) melted the glacier above it, releasing large quantities of water and triggering a hydrovolcanic eruption. Ash rose in a plume 10 km high, and was blown westward and into the skies over Europe. Volcanic ash can damage or destroy aircraft engines, so the precaution was taken to prohibit air travel for a 5-day period. The enormous economic impact of stopping flights has led to numerous studies about the best way to deal with similar events with volcanic ash in the future.

Figure 11.31 Hydrovolcanic eruption of Eyjafjallajökull in April of 2010. Left- Eruptive column with volcanic lightning. Volcanic lightning is caused by the static electricity generated by volcanic ash particles rubbing together. Right- Another view of the ash cloud, with westward winds carrying ash toward Europe where it would disrupt air traffic. Source: Karla Panchuk (2017) CC BY-SA 4.0. Left photograph: Terje Sørgjerd (2010) CC BY-SA 3.0 view source Right photograph: Henrik Thorburn (2010) CC BY 3.0 view source Click the image for more attributions.

References

Bressan, D (2012). Geology Scene Investigation: Death by Volcanic Fire. Visit website

British Geological Survey (n.d.). Eyjafjallajökull eruption, Iceland | April/May 2010. Visit website  

Digital History Project (2011). “Eyewitness Account to Eruption of Mont Pelee Matinique St Pierre Fort de France” By Ellery S. Scott. Visit website 

Rosen, J. (2015). Benchmarks: May 8, 1902: The deadly eruption of Mount Pelée. Visit website

U. S. Geological Survey (1997). Pyroclastic flows. Visit website

Thus far volcanoes have been discussed in terms of the kinds of volcanic mountains they form, the materials they produce, and the style of eruption they have. All of these characteristics can be tied together into a big picture by considering the plate tectonic settings in which magma forms (Figure 11.32). The vast majority of volcanoes are present along plate tectonic boundaries.

Figure 11.32 Plate tectonic settings of volcanism. Volcanoes along subduction zones are the result of flux melting (lowering the melting point by adding water). Decompression melting produces volcanoes along divergent margins (ocean spreading centres and continental rift zones), as well as above mantle plumes. Contact between hot mafic partial melts and felsic rocks can trigger partial melting of the felsic rocks (melting from conduction). Source: Karla Panchuk (2017) CC BY 4.0. Modified after Steven Earle (2015) CC BY 4.0 view original and U. S. Geological Survey (1999) Public Domain view original

There are four main scenarios to consider:

• Divergent boundaries where melting is triggered by decompression
• Subduction zones (ocean-ocean and ocean-continent convergent boundaries) where flux melting occurs as water is released from subducting ocean crust
• Hot spots where plumes of hot mantle material rise up, then melt as a result of decompression.
• Melting by conduction when magma transfers heat to rocks having a lower melting temperature.

Decompression Causes Volcanism Along Spreading Centres and Rift Zones

At an ocean spreading ridge (centre of Figure 11.32), convection moves hot mantle rock slowly upward at rates of cm per year. At roughly 60 km below the surface, the mantle rocks have decompressed is enough to permit partial melting of approximately 10% of the ultramafic rock. Mafic magma is produced, and it moves up toward the surface. Magma fills vertical fractures produced by the spreading and spills out onto the sea floor making pillow lavas and lava flows. Spreading-ridge volcanism is taking place approximately 200 km offshore from the west coast of Vancouver Island.

In continental rift zones where continental crust is thinning (far right in Figure 11.32), a similar decompression process occurs, triggering partial melting of ultramafic mantle rocks. However, if the continental crust above the region where melting occurs has a lower melting temperature than the mafic melt that is produced, the continental crust will also melt.

Continental rift zones can have a range of volcano types. If mafic magma erupts, shield volcanoes, broad lava flows, and cinder cones result. However, if rocks of other compositions are melted and added in, or the mafic magma undergoes fractional crystallization before erupting, then composite volcanoes will also form.

Water Causes Partial Melting Along Subduction Zones

At an ocean-continent convergent boundary (Figure 11.32, right) or ocean-ocean convergent boundary (Figure 11.32, left), oceanic crust is pushed down into the mantle. Although temperatures are high, the slab is kept from melting by high pressures. However, under these conditions minerals in the slab release water from within their crystal structures. The water lowers the melting point of rock above the slab, and partial melting is triggered within the mantle. Mafic magma rises through the mantle to the base of the crust. There it contributes to partial melting of crustal rock, and more felsic material is added to the magma. The magma, now intermediate in composition, continues to rise and assimilate crustal material. In the upper part of the crust, it accumulates into plutons. Over time, fractional crystallization of magma within the pluton can make it even more silica-rich. From time to time, the magma from the plutons rises toward surface, leading to volcanic eruptions. 

Composite volcanoes with Vulcanian or Plinian eruption styles are characteristic of the volcanic arcs that form in subduction zones, although in the Trans-Mexico Volcanic Belt, Strombolian eruptions produce short-lived cinder cones. Where two margins of oceanic crust collide, the volcanic arc will be a chain of volcanic islands. Where continental and oceanic crust collide, there will be a volcanic arc on the continental crust.

Mt. St. Helens: A Composite Volcano in the Cascades Range Continental Volcanic Arc

On May 18, 1980 at 8:32 a.m., a M5.1 earthquake shook Mt. St. Helens, and marked the start of a 9-hour Plinian eruption (Figure 11.33) with a 24 km high eruption column and multiple pyroclastic flows. By the time the eruption was over, a large part of the volcano had been blasted away.

Figure 11.33 Eruption of composite subduction-zone volcano Mt. St. Helens on May 18, 1980. Top- Plinian eruption column. Bottom left- Mt. St. Helens before the eruption. Bottom right- The remains of Mt. St. Helens after the eruption. Sources: Top- Karla Panchuk (2017) CC BY 4.0; Top- Photograph by NOAA (1980) Public Domain view source. Bottom left- R. Hoblitt, U. S. Geological Survey, Cascades Volcano Observatory (1979) Public Domain (label added) view source. Bottom right- Steven Earle (2015) CC BY 4.0 (label added) view source. Click the image for more attributions.

 

The explosive eruption was driven by gas-rich rhyolitic magma, however not all of Mt. St. Helens’ eruptions have been of felsic or intermediate material. The lava tube in Figure 11.10 (bottom) is from a time when Mt. St. Helens erupted basaltic lava. Data from the iMUSH (Imaging Magma Under St. Helens) project show that a magma chamber is present beneath Mt. St. Helens at between 5 and 14 km depth, but that a much larger magma chamber is present below it, which extends down to the mantle (Figure 11.34). Earthquakes in the 24 hours after the 1980 eruption suggested movement of magma within the smaller chamber (yellow arrows in Figure 11.34), but earthquakes from 1980 to 2005 indicate movement of magma within the deeper chamber as well (black arrows).

Figure 11.34 Magma chambers beneath Mt. St. Helens and Indian Heaven Volcanic Field, sketched from iMUSH (Imaging Magma Under St. Helens) project results. In a 24 hour period after the May 18, 1980 eruption, earthquakes in and around the smaller magma chamber suggested migration of magma (yellow arrows). Earthquakes recorded between 1980 and 2005 suggest migration of magma within a larger chamber that extends to the mantle (black arrows). The larger magma chamber might feed another smaller chamber beneath the Indian Heaven Volcanic Field. Source: Karla Panchuk (2017) CC BY 4.0, based on Kiser et al. (2016), Figure 4B.

 

The complex history of Mt. St. Helens could reflect changes in the composition of magma within the small chamber over time, as fractionation proceeds, and the magma becomes more silica rich. However, movement of more mafic magma from the larger chamber could also contribute to eruptions with different chemical compositions. The larger magma chamber may be connected to a chamber feeding the nearby Indian Heaven Volcanic Field, which contains shield volcanoes and cinder cones, and for which basalt makes up 80% of erupted materials.

Mantle Plumes Can Cause Volcanism Away from Plate Boundaries

Mantle plumes are rising columns of hot solid rock. The column may be kilometres to 10s of kilometres across, but near the surface it spreads out to create a mushroom-like head that is 10s to over 100 kilometres across. Mantle plumes are different from the convection that normally occurs beneath ocean spreading centres: plumes rise approximately 10 times faster than mantle convection normally occurs, and may originate deep in the mantle, possibly just above the core-mantle boundary.

When the mantle plume rises to the base of the lithosphere, the pressure is low enough to permit partial melting of the plume material, producing mafic magma. Heat carried by the mantle plume may also melt rock adjacent to the plume. The magma rises and feeds hotspot volcanoes. The lithospheric plate above the mantle plume is moving across the plume, so a chain of hotspot volcanoes can result as existing hotspot volcanoes are slowly moved away from the mantle plume, and new volcanoes form in the lithosphere.

Many shield volcanoes are associated with mantle plumes, including those that make up the Hawai’ian islands. All of the Hawai’ian volcanoes are related to the mantle plume that currently lies beneath Mauna Loa, Kilauea, and Lōʻihi (Figure 11.35, top). There is evidence of crustal magma chambers beneath all three active Hawai’ian volcanoes. At Kīlauea, the magma chamber appears to be several kilometres in diameter, and is situated between 8 km and 11 km below surface (Lin et al., 2014). In this area, the Pacific Plate is moving northwest at a rate of about 7 cm/year. This means that the earlier formed — and now extinct — volcanoes have now moved well away from the mantle plume. The hotspot has in fact been present for at least 85 million years (Regelous et al., 2003), as evidenced by the long chain of eroded and submerged mountains stretching to the Aleutian Trench (Figure 11.35, bottom).

Figure 11.35 Hawai’ian hotspot volcanoes and volcanic chain. Top- A mantle plume beneath Hawai’i supplies magma to Mauna Loa Volcano, Kīlauea Volcano, and Lōʻihi Seamount. Volcanoes to the northwest are no longer active because they have moved away from the plume. Bottom- Bathymetric (depth) map showing the chain of islands stretching toward the Aleutian Trench, and marking the progress of the Pacific Plate over the mantle plume. Source: Top- J. E. Robinson, U. S. Geological Survey (2006) Public Domain view source. Bottom- National Geophysical Data Center/ U. S. Geological Survey (2006) Public Domain (labels added) view source.

Kīlauea Volcano is approximately 300 ka old, while neighbouring Mauna Loa Volcano is over 700 ka and Mauna Kea Volcano is over 1 Ma. If volcanism continues above the Hawaii mantle plume in the same manner that it has for the past 85 Ma, it is likely that Kīlauea Volcano will continue to erupt for at least another 500,000 years. By that time, its neighbour, Lōʻihi Seamount, 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.

Large Igneous Provinces (LIPs)

While the Hawaii mantle plume has produced a relatively low volume of magma for approximately 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 (LIPs) 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, which extends across Washington, Oregon, and Idaho in the United States (Figure 11.36). This volcanism, which covered an area of about 160,000 km² with basaltic rock up to several hundred metres thick, took place between 17 and 14 Ma.

Figure 11.36 Part of the Columbia River Basalt Group at Frenchman Coulee, eastern Washington, United States. All of the flows visible here have formed large (up to two metres in diameter) columnar basalts, a result of relatively slow cooling of flows that are tens of m 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. Source: Steven Earle (2015) CC BY 4.0 view source

The mantle plume that is assumed to be responsible for the Columbia River LIP 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 km³ of felsic magma. This is approximately 900 times the volume of the 1980 eruption of Mt. St. Helens, but only 5% of the volume of mafic magma in the Columbia River LIP.

Most other LIP eruptions are much bigger. The Siberian Traps (also basalt), which erupted at the end of the Permian period at 251 Ma, are estimated to have produced approximately 40 times as much lava as the Columbia River LIP. The largest known LIP is the Ontong Java Plateau, located in the southwest Pacific Ocean. It formed at around 122 Ma, and presently covers 1,500,000 km2 and has a volume of 5,000,000 km³. But this is only a small fraction of its original size. The majority of it has been subducted, and it may have been split into pieces that have been classified as separate LIPs.

Kimberlites

Kimberlite pipes are carrot-shaped cones of ultramafic rock. They form from the explosive eruption of mantle plumes originating at depths of 150 to 450 km in the mantle. The plume makes its way to the surface quickly (over hours to days), having little interaction with the surrounding rocks, and thus preserving a sample of the ultramafic mantle. As the plume nears the surface, a build-up of gas causes it to pick up speed, and by the time it reaches the surface it may be travelling faster than the speed of sound. The explosiveness of kimberlite eruptions means that they do not form volcanic mountains on the surface, but leave circular holes in the ground.

Kimberlite eruptions that originate at depths greater than 200 km beneath old, thick, continental crust travel through the region of the mantle where diamond is stable. In some cases, such as in Saskatchewan and the Northwest Territories, kimberlites bring diamond-bearing material to the surface. All of Earth’s diamond deposits are thought to have originated in this way.

Diamond mines in kimberlites, such as the Ekati Mine in the Northwest Territories, are easy to spot by the characteristic circular hole that develops as miners excavate the cone-shaped structure (Figure 11.37). The kimberlites at Ekati erupted between 45 and 60 Ma. Many kimberlites are older, and 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 date to approximately 30 Ma.

Figure 11.37 The Ekati diamond mine in the Northwest Territories, part of the Lac de Gras kimberlite field. Source: Karla Panchuk (2017) CC BY-SA 4.0; Photograph by J. Pineau (2010) CC BY-SA 3.0 view source. Click the image for more attributions.

References

Kiser, E., Palomeras, I., Levander, A., Zelt, C., Harder, S., Schmandt, B., Hansen, S., Creager, K., & Ulberg, C. (2016). Magma reservoirs from the upper crust to the Moho inferred from high-resolution Vp and Vs models beneath Mount St. Helens, Washington State, USA. Geology (44)6, 411-414. DOI: 10.1130/G37591.1

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, 42(3), 187-190. DOI: 10.1130/G35001.1

Regelous, M., Hofmann, A. W., Abouchami, W., & Galer, S. J. G. (2003) Geochemistry of lavas from the Emperor Seamounts, and the geochemical evolution of Hawaiian magmatism from 85 to 42 Ma. Journal of Petrology 44(1), 113-140. DOI: 10.1093/petrology/44.1.113 view PDF

U. S. Geological Survey, Volcano Hazards Program (n.d.). Indian Heaven Volcanic Field visit website

U. S. Geological Survey, Volcano Hazards Program (n.d.). Mount St. Helens: 1980 Cataclysmic Eruption visit website

The basaltic lava flows produced by volcanoes on the island of Hawai’i are responsible for extensive damage to homes, infrastructure, and habitats. Figure 11.38 shows lava flows (in black) from the Puʻu ʻŌʻō crater of Kīlauea Volcano. The lava flow destroyed the house, and is encroaching on the transfer station. Smoke in the background marks locations where additional flows have broken out and are burning vegetation.

Figure 11.38 Lava flow from Kīlauea’s Puʻu ʻŌʻō crater. Lava (in black) has destroyed a house and threatens a transfer station. Source: U. S. Geological Survey (2014) Public Domain view source

In spite of the damage that lava flows can cause, they are not the volcanic hazard with the greatest impact on lives and infrastructure. Even the relatively free-flowing Hawai’ian basaltic lava moves slowly enough that it can be escaped on foot. Far more dangerous hazards are related to gases and volcanic debris. However, the largest impact and the greatest suffering are caused not by the immediate effects of volcanic eruptions, but by large-scale changes to climate and environments caused by volcanism. Indirect effects resulting in respiratory distress, toxicity, famine, and habitat destruction 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.

Volcanic Gas and Tephra Emissions

Large volumes of rock and gases are emitted during major Plinian eruptions at composite volcanoes, and a large volume of gas is released during some very high-volume effusive eruptions. Gases and fine particles of volcanic ash can cause respiratory distress and poisoning, and ash poses a risk for aircraft.

Most of the tephra from large explosive eruptions 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) fall closer to the volcano, and the accumulation of tephra from large eruptions can cause serious damage and casualties. When the large eruption of Mt. Pinatubo in the Philippines occurred in 1991, tens of centimetres of ash accumulated in fields and on rooftops in the surrounding populated region. Heavy typhoon rains hit the island at the same time and added to the weight of the tephra. The weight was too much for roofs to bear, and thousands of structures collapsed, causing at least 300 of the 700 deaths attributed to the eruption.

One of the long-term effects of adding volcanic particles and gases to the atmosphere is cooling. 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 sulphur dioxide combined with water to make sulphate aerosols, which block incoming solar energy. The accumulation of sulphate aerosols over that 8 months 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, and 50% of cattle. The ensuing famine, along with HF poisoning, resulted in more than 10,000 human deaths, about 25% of the population.

Pyroclastic Flows

In a typical explosive eruption at a composite volcano, the tephra and gases are ejected with explosive force and sent high up into the atmosphere. As the eruption proceeds, and the amount of gas in the rising magma starts to decrease, and less gas is supplied to the eruption column. Parts of the column will become denser than air, leading the column to collapse and flow downward along the flanks of the volcano (Figure 11.39), picking up speed as it cools.

Figure 11.39 The Plinian eruption of Mt. Mayon, Philippines in 1984. Although most of the eruption column is ascending into the atmosphere, pyroclastic flows are traveling down the sides of the volcano in several places. Warnings were issued in time to evacuate 73,000 people. Source: C. G. Newhall, U. S. Geological Survey (1984) Public Domain view source

Pyroclastic flows can travel over water, in some cases for many kilometres. In 1902 the pyroclastic flow from the eruption of Mt. Pelée traveled out into the harbour and destroyed several wooden ships anchored there. The pyroclastic flow from the 1883 eruption of Krakatau traveled 80 km across the Sunda Straits and claimed victims on the southwest coast of Sumatra. It also triggered a tsunami.

One of the most famous pyroclastic flows occurred when Mt. Vesuvius erupted in 79 CE. It buried the cities of Pompeii and Herculaneum, killing an estimated 18,000 people.

Lahar

A lahar is any mudflow or debris flow that is related to a volcano (Figure 11.40). 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 when the volcano Nevado del Ruiz caused the ice dam on a glacial lake to fail. The resulting lahar killed 23,000 in Armero, about 50 km from the volcano.

Figure 11.40 Mud left behind from the lahar after the May 18, 1980 eruption of Mt. St. Helens. The lahar carried the boulder to its present location. Source: L. Topinka, U. S. Geological Survey (1980) Public Domain view source

Lahars can also happen when there is no volcanic eruption, because 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. Fatalities were largely because of mudflows and debris flows triggered by intense rainfall — some regions received almost 2 m of rain over a few days.

At Casita Volcano in Nicaragua, 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. It struck the towns of El Porvenir and Rolando Rodriguez killing more than 2,000 people. 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.

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 Mt. 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 (Figure 11.41) as magma moved from depth into a magma chamber within the mountain itself. Early on the morning of May 18, a moderate earthquake struck and destabilized the bulge, leading to Earth’s largest observed landslide in historical times. The failure of this part of the volcano exposed the underlying magma chamber, causing it to explode sideways. This in turn exposed the conduit leading to the magma chamber below, resulting in a Plinian eruption lasting nine hours.

Figure 11.41 Bulge forming on the north side of Mt. St. Helens, April 27 1980. Source: P. Lipman, U. S. Geological Survey (1980) Public Domain view source

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

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. Bulging — the deformation of part of the volcano, indicating that a magma chamber at depth is swelling or becoming more pressurized
3. Seismicity — 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. Seismicity ceases — a sudden decrease in the rate of earthquake activity. This may indicate that magma has stalled, and that\ something is about to give way
5. Big bump — a pronounced bulge on the side of the volcano (like the one at Mt. St. Helens in 1980), which may indicate that magma has moved close to surface
6. Steam — steam eruptions ( 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, it is possible to determine the necessary equipment to have and actions to take to monitor a volcano and predict when it might erupt. 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, instruments that detect vibration. 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 an early warning that something is changing beneath one of the volcanoes. There are currently enough seismometers in the Lower Mainland and on Vancouver Island to provide this information. You can view a map of Canadian National Seismograph Network here.

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 11.42). 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 11.42 A seismometer installed in 2007 in the vicinity of the Nazco Cone, BC. Source: Cathie Hickson (n.d.) used with permission.

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, or by collecting samples (Figure 11.43) and analyzing them in a lab.

Figure 11.43 A geologist collects a gas sample from Sherman Crater, Mt. Baker, Washington. Gas is drawn through a titanium tube inserted in a fumarole, and collected in a glass vacuum flask. Source: D. Tucker, U. S. Geological Survey (2006) Public Domain view source

Measuring Deformation

There are two main ways to measure ground deformation at a volcano. 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 11.27). GPS is more effective than a tiltmeter because it provides information on how far the ground has actually moved — east-west, north-south, and up-down.

Figure 11.44 A GPS unit installed at Hualālai 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. Source: U. S. Geological Survey (n.d.) Public Domain view source

Putting It All Together

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.

Canada’s volcanically active regions are located in British Columbia and the Yukon Territory (Figure 11.45). At least 49 eruptions have occurred within these regions in the last 10,000 years. There are five volcanic regions associated with three types of plate tectonic settings: a subduction zone, a mantle plume, and a continental rift zone.

Figure 11.45 Canada’s volcanic regions are located in British Columbia and the Yukon Territory. Volcanism is associated with three tectonic settings: the subduction zone along the west coast (Garibaldi Volcanic Belt, Wrangell Volcanic Belt), a continental rift zone (Wells Gray-Clearwater Volcanic Field, Stikine Volcanic Belt), and a mantle plume (Anahim Volcanic Belt). Source: Volcanoes Canada, Canadian Hazards Information Service, Natural Resources Canada (n.d.) view source Click the image for copyright information.

Subduction Zone Volcanism: Wrangell and Garibaldi Volcanic Belts

The Wrangell Volcanic Belt is the result of subduction beneath the North American Plate. Volcanoes in the Canadian part of the Wrangell Volcanic Belt erupted between 17.8 and 10.4 million years ago. They were fed by lava that seeped up along a leaky transform fault.

Southwestern British Columbia is at the northern end of the Juan de Fuca subduction zone, and part of the Cascade Volcanic Arc that extends south through Washington and Oregon. The Canadian part of the Cascade Arc has had a lower rate and volume of volcanism than U. S. portions. One reason is that the northern part of the Juan de Fuca Plate is subducting more slowly than the rest of the plate, or else has stalled.

The Garibaldi Volcanic Belt has several volcanic centres , or regions where volcanism has caused multiple volcanoes to develop (Figure 11.46).

Figure 11.46 Volcanic centres within the Garibaldi Volcanic Belt. The most recent eruption was 2,350 years ago at Mt. Meager. Source: Sémhur (2007) CC BY-SA 4.0 view source. Click the image to enlarge.

The most recent volcanic activity in this area was 2,350 years ago at Mt. Meager. An explosive eruption similar in magnitude to that of Mt. St. Helens in 1980 spread ash as far east as Alberta. There was also significant volcanic activity at Mt. Price and Mt. Garibaldi approximately 10,000 years ago as glacial ice receded. In both cases, lava and tephra built up against glacial ice. The western side of Mt. Garibaldi failed by sector collapse when the ice melted, leaving rocks unsupported. Eruption beneath glacial ice resulted in the formation of a tuya– a steep-sided, flat-topped volcano- called The Table near Mt. Garibaldi (Figure 11.47).

Figure 11.47 The Table, a tuya near Mt. Garibaldi. Tuyas form when volcanoes erupt beneath ice, and their shape is determined by rapid cooling beneath the ice sheet. Source: Andre Charland (2004) CC BY 2.0 view source

Mantle Plume Volcanism: Anahim Volcanic Belt

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 (Figure 11.48). The North American 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.

Figure 11.48 Anahim Volcanic Belt, the result of a mantle plume beneath the North American Plate. Source: Sémhur (2007) CC BY-SA 4.0 view source Click the image to enlarge.

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 11.49). The name “Rainbow” refers to the bright colours displayed by some of the volcanic rocks as they weather.

Figure 11.49 Tsitsutl, the “painted mountain” within the Rainbow Range of the Anahim Volcanic Belt. The vibrant colours of the Rainbow Range are the result of chemical weathering. Source: Drew Brayshaw (2015) CC BY-NC 2.0 view source

Rift-Related Volcanism: Wells Gray-Clearwater Volcanic Field and Stikine Volcanic Belt

While British Columbia 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 (Figure 11.50), and the Stikine Volcanic Belt (also called the Northern Cordillera Volcanic Province), which ranges across the northwestern corner of the province.

Figure 11.50 Wells Gray-Clearwater Volcanic Field is the result of extension in the crust. Source: Sémhur (2007) CC BY-SA 4.0. view source Click the image to enlarge.

The Stikine Volcanic Belt 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, lava overran a village on the Nass River, and 2,000 people were lost. The region is now part of the Anhluut’ukwsim Laxmihl Angwinga’asanskwhl Nisga’a (Nisga’a Memorial Lava Bed Park).

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, Mt. Edziza itself (Figure 11.51) 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 11.51 Mount Edziza, in the Stikine Volcanic Belt, BC, with Eve Cone in the foreground. Source: NASS5518 (2008) CC BY 2.0 view source

References

Geological Survey of Canada (n.d.) Catalog of Canadian Volcanoes: Anahim volcanic belt Visit website

Geological Survey of Canada (n.d.) Catalog of Canadian Volcanoes: Garibaldi volcanic belt: Garibaldi Lake volcanic field Visit website

Skulski, T., Francis, D., & Ludden, J. (1991) Arc-transform magmatism in the Wrangell volcanic belt. Geology (19)1, 11-14. doi:10.1130/0091-7613(1991)019<0011:ATMITW>2.3.CO;2

Trop, J. M., Hart, W. K., Snyder, D., & Idleman, B. (2012). Miocene basin development and volcanism along a strike-slip to flat-slab subduction transition: Stratigraphy, geochemistry, and geochronology of the central Wrangell volcanic belt, Yakutat-North America collision zone. Geosphere (8)4, 805-834. doi:10.1130/GES00762.1

Volcanoes Canada, Canadian Hazards Information System, Natural Resources Canada (n.d.). Where Are Canada’s Volcanoes? Visit website

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

11.1 What Is A Volcano?

Volcanoes are places where molten rock escapes to Earth’s surface. Some volcanoes are cone-shaped or hill-shaped mountains, and some eruptions happen along fissures. Eruptions are fed by a magma chamber beneath the volcano. Sometimes a volcano collapses into empty space in the magma chamber beneath, forming a caldera.

11.2 Materials Produced by Volcanic Eruptions

Volcanoes produce gas, lava flows, and debris called tephra. The characteristics of a lava flows depend on whether the lava is thin and runny (mafic with low gas content) or thick and sticky (felsic with high gas content). Tephra is classified according to size. Ash is less than 2 mm in diameter, lapilli is between 2 mm and 64 mm, and blocks and bombs are larger than 64 mm.

11.3 Types of Volcanoes

Cinder cones are relatively small straight-sided volcanoes that are composed mostly of mafic rock fragments. Composite volcanoes consist of alternating layers of lava flows and tephra. The tend to be intermediate to felsic in composition, and get steeper toward the top. Shield volcanoes are broad, low, hill-like volcanoes that form from layers of low-viscosity mafic lava.

11.4 Types of Volcanic Eruptions

Volcanic eruptions can be classified according to how explosive they are, and how high into the atmosphere they blast material. Hawai’ian eruptions are relatively gentle effusive eruptions of low-viscosity mafic lava, and form shield volcanoes. Strombolian eruptions are more vigorous eruptions of mafic tephra. They blast material hundreds of metres into the air. The tephra falls out of the atmosphere to form a cinder cone. Vulcanian eruptions are explosive eruptions of intermediate composition lava, producing pyroclastic flows and eruptive columns from 5 to 10 km high. Plinian eruptions are highly explosive eruptions of felsic lava, and can produce eruption columns up to 45 km high. Both Vulcanian and Plinian eruptions are associated with composite volcanoes. Hydrovolcanic eruptions are the explosive result of magma or lava interacting with water, and rapidly changing the water to steam.

11.5 Plate Tectonics and Volcanism

Volcanism is closely related to plate tectonics. Most volcanoes are associated with convergent plate boundaries (at subduction zones), but a great deal of volcanic activity also occurs 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.

11.6 Volcanic Hazards

Most direct volcanic hazards are related to volcanoes that erupt explosively, especially composite volcanoes. Pyroclastic flows, some as hot as 1000 ˚C, can move at hundreds of km/h and will kill anything in the way. Lahars, volcano-related mudflows, can be large enough to destroy entire towns.  Lava flows are also destructive, but tend to move slowly enough to permit people to get to safety. Indirect hazards claim far more lives than direct hazards, and include famine related to volcanically-induced climate cooling.

11.7 Monitoring Volcanoes and Predicting Eruptions

Clues that a volcanic eruption might soon occur include earthquakes, a change in the type and amount of gases released, and changes in the shape of the volcano as magma moves within it. Volcanoes are monitored using seismometers to detect earthquakes, volcanic gases are sampled and analyzed, and instruments are used to detect deformation of the volcano. These tools make it possible to assess the hazard posed by a given volcano, and the risk of eruption.

11.8 Volcanoes in British Columbia

British Columbia and the Yukon Territory include examples of volcanoes that form as a result of fluid-induced melting along a subduction zone (the Wrangell and Garibaldi volcanic belts) , as a result of decompression where the crust is thinning and stretching (Stikine Volcanic Belt and Wells Gray-Clearwater Volcanic Field), and because of mantle plume activity (Anahim Volcanic Belt).

Review Questions

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 and why do pillow lavas form?
6. What two kinds of volcanic materials make up 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. Why do shield volcanoes have gentle slopes?
10. Which type of volcanic mountain would last longest: a shield volcano, a cinder cone, or a composite 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 for activity?
13. What is the likely geological origin of the Nazko Cone?
14. What might be the explanation for southwestern B.C. having much less subduction-related volcanism than adjacent Washington and Oregon?

1. The three main tectonic settings for volcanism are (1) subduction zones at convergent plate boundaries, (2) divergent plate boundaries, and (3) mantle plumes (a.k.a. hot spots).
2. The primary mechanism for partial melting at a convergent plate boundary is the addition of water to hot mantle rock. The water reduces the melting temperature of the rock (flux melting).
3. The explosiveness of a volcanic eruption depends on the pressure of the magma. Gases create that pressure, and if the magma is viscous those gases cannot escape easily. Felsic and intermediate magmas tend to have more gas than mafic magmas, and are also more viscous, trapping the gas in.
4. When magma is deep within the crust the pressure is too high for the gases to bubble out of solution.
5. Pillow lavas form where mafic lava erupts in water. When the magma oozes out into the water the outside cools first forming a hard skin that maintains the pillow shape.
6. Composite volcanoes are formed of layers of lava flows and tephra (volcanic fragments ranging from fine ash to blocks and bombs) from explosive eruptions.
7. A lahar is a mud flow or debris flow on a volcano. Lahars are common on composite volcanoes because they are steeper than shield volcanoes, they typically have ice and snow, and they are not as strong as shield volcanoes.
8. Some lahars form during an eruption when snow and ice melt quickly, while others may form from heavy rain.
9. The lava that forms shield volcanoes is typically low viscosity. It can flow easily and also tends to form lava tubes. As a result, it is able to travel a long way from the vent, forming a low broad shield.
10. Cinder cones erode rapidly because they are mostly piles of tephra. Composite volcanoes are more resistant to weathering and erosion than cinder cones because lava flows help hold together tephra, but composite volcanoes still don’t last as long as shield volcanoes. Shield volcanoes are stronger because they consist more of lava flows than tephra.
11. Weak seismic activity is associated with all stages of a volcanic eruption. In the early stages magma is moving at depth and pushing rock aside, creating small earthquakes. The flow of magma can also produce special type of seismic response known as a harmonic tremor.
12. GPS technology is used to determine if there is any slow deformation of the flanks of a volcano related to movement of magma toward the surface.
13. The Nazko Cone is thought to be related to a mantle plume.
14. One hypothesis to explain the lower rate of volcanism in British Columbia than in adjacent Washington and Oregon is that the northern part of the Juan de Fuca Plateis not subducting as quickly as the rest of the plate.

Earthquake Shaking Comes from Elastic Deformation

Earthquakes occur when rock ruptures (breaks), causing rocks on one side of a fault to move relative to the rocks on the other side. Although motion along a fault is part of what happens when an earthquake occurs, rocks grinding past each other is not what creates the shaking. In fact, it could be said that the earthquake happens after rocks have undergone most of the displacement. Consider this: if rocks slide a few centimetres or even metres along a fault, would that motion alone explain the incredible damage caused by some earthquakes? If you were in a car that suddenly accelerated then stopped, you would feel a jolt. But earthquakes are not a single jolt. Buildings can swing back and forth until they shake themselves to pieces, train tracks can buckle and twist into s-shapes, and roads can roll up and down like waves on the ocean. During an earthquake, rock is not only slipping. It is also vibrating like a plucked guitar string.

Rocks might seem rigid, but when stress is applied they may stretch. If there hasn’t been too much stretching, a rock will snap back to its original shape once the stress is removed. Deformation that is reversible is called elastic deformation. Rocks that are stressed beyond their ability to stretch can rupture, allowing the rest of the rock to snap back to its original shape. The snapping back of the rock returning to its original shape causes the rock to vibrate, and this is what causes the shaking during an earthquake. The snapping back is called elastic rebound.

Figure 12.3 (top) shows this sequence of events. Stress is applied to a rock and deforms it. The deformed rock ruptures, forming a fault. After rupturing, the rock above and below the fault snaps back to the shape it had before deformation.

Figure 12.3 Elastic deformation, rupture, and elastic rebound. Top: Stress applied to a rock causes it to deform by stretching. When the stress becomes too much for the rock, it ruptures, forming a fault. The rock snaps back to its original shape in a process called elastic rebound. Bottom: On an existing fault, asperities keep rocks on either side of the fault from sliding. Stress deforms the rock until the asperities break, releasing the stress, and causing the rocks to spring back to their original shape. Source: Karla Panchuk (2017) CC BY 4.0. Modified after Steven Earle (2015) CC BY 4.0 view original.

Ruptures can also occur along pre-existing faults (Figure 12.3, bottom). The rocks on either side of the fault are locked together because bumps along the fault, called asperities, prevent the rocks from moving relative to each other. When the stress is great enough to break the asperities, the rocks on either side of the fault can slide again. While the rocks are locked together, stress can cause elastic deformation. When asperities break and release the stress, the rocks undergo elastic rebound and return to their original shape.

Rupture Surfaces Are Where the Action Happens

Images like 12.3 are useful for illustrating elastic deformation and rupture, but they can be misleading.  The rupture that happens doesn’t occur as in 12.3, with the block being ruptured through and through.  The rupture and displacement only happen along a subsection of a fault, called the rupture surface.  In Figure 12.4, the rupture surface is the dark pink patch.  It takes up only a part of the fault plane (lighter pink). The fault plane represents the surface where the fault exists, and where ruptures have happened in the past.  Although the fault plane is drawn as being flat in Figure 12.4, faults are not actually perfectly flat.

The location on the fault plane where the rupture happens is called the hypocentre or focus of the earthquake (Figure 12.4, right). The location on Earth’s surface immediately above the hypocentre is the epicentre of the earthquake.

Figure 12.4 Rupture surface (dark pink), on a fault plane (light pink). The diagram represents a part of the crust that may be tens or hundreds of kilometres long. The rupture surface is the part of the fault plane along which displacement occurred. Left: 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. Coloured arrows represent propagation of failure on a rupture surface. In this case, the failure starts at the dark blue heavy arrow and propagates outward, reaching the left side first (green arrows) and the right side last (yellow arrows). Right: An earthquake’s location can be described in terms of its hypocentre (or focus), the location on the fault plane where the rupture happens, or in terms of its epicentre (red star), the location above the hypocentre. Source: Left: Steven Earle (2015) CC BY 4.0 view source. Right: Karla Panchuk (2017) CC BY 4.0.

Within the rupture surface, the amount of displacement varies. In Figure 12.4, the larger arrows indicate where there has been more displacement, and the smaller arrows where there has been less.  Beyond the edge of the rupture surface there is no displacement at all.  Notice that this particular rupture surface doesn’t even extend to the land surface of the diagram.

The size of a rupture surface and the amount of displacement along it will depend on a number of factors, including the type and strength of the rock, and the degree to which the rock was stressed beforehand. The magnitude of an earthquake will depend on the size of the rupture surface and the amount of displacement.

A rupture doesn’t occur all at once along a rupture surface. It starts at a single point and spreads rapidly from there. Figure 12.4 illustrates a case where rupturing starts at the heavy blue arrow in the middle, then continues through the lighter blue arrows. The rupture spreads to the left side (green arrows), then the right (yellow arrows).

Depending on the extent of the rupture surface, the propagation of failures (incremental ruptures contributing to making the final rupture surface) from the point of initiation is typically completed within seconds to several tens of seconds. The initiation point isn’t necessarily in the centre of the rupture surface; it may be close to one end, near the top, or near the bottom.

Shifting Stress Causes Foreshocks and Aftershocks

Earthquakes don’t usually occur in isolation. There is often a sequence in which smaller earthquakes occur prior to a larger one, and then progressively smaller earthquakes occur after.  The largest earthquake in the series is the mainshock.  The smaller ones that come before are foreshocks, and the smaller ones that come after are aftershocks. These descriptions are relative, so it can be necessary to reclassify an earthquake.  For example, the strongest earthquake in a series is classified as the mainshock, but if another even bigger one comes after it, the bigger one is called the mainshock, and the earlier one is reclassified as an aftershock.

A rupture surface does not fail all at once.  A rupture in one place leads to another, which leads to another.  Aftershocks and foreshocks represent the same thing, except on a much larger scale. The rupture illustrated in Figure 12.4 reduced stress in one area, but in doing so, transferred stress to others (Figure 12.5). Imagine a frayed rope breaking strand by strand.  When a strand breaks, the tension on that strand is released, but the remaining strands must still hold up the same amount of weight. If another strand breaks under the increased burden, the remaining strands have an even greater burden than before.  In the same way that the stress causes one strand after another to fail, a rupture can trigger subsequent ruptures nearby.

Figure 12.5 Stress changes related to an earthquake. Stress decreases in the area of the rupture surface, but increases on adjacent parts of the fault. Source: Steven Earle (2015) CC BY 4.0 view source.

Numerous aftershocks were associated with the magnitude 7.8 earthquake that struck Haida Gwaii in October of 2012 (Figure 12.6; mainshock in red, aftershocks in white). Some of the stress released by the mainshock was transferred to other nearby parts of the fault, and contributed to a cascade of smaller ruptures. But stress transfer need not be 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 aftershocks from the Haida Gwaii earthquake are scattered rather than located only on the main faults.

Figure 12.6 Magnitude 7.8 Haida Gwaii earthquake and aftershocks. Mainshock (red circle marks the epicentre) occurred on October 28th, 2012. Aftershocks are for the period from October 28th to November 10th of 2012. Although the epicentre is near a transform boundary, the rupture was influenced more by compression related to the subduction zone. Source: Karla Panchuk (2017) CC BY 4.0. Base map with epicentres from the U. S. Geological Survey Latest Earthquakes tool view interactive map. Subduction zone after Wang et al. (2015). Click the image for more attributions.

The effects of stress transfer may not show immediately. Aftershocks can be delayed for hours, days, weeks, or even years. Because stress transfer affects a region, not just a single fault, and because there can be delays between the event that transferred stress and the one that was triggered by the transfer, it can sometimes be hard to be know whether one earthquake is actually associated with another, and whether a foreshock or aftershock should be assigned to a particular mainshock.

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 (observed as rapid seismic vibrations on instruments). It was first discovered on the Vancouver Island part of the Cascadia subduction zone by Geological Survey of Canada geologists Herb Dragert and Gary Rogers.Rogers, G. and Dragert, H. (2003). Episodic tremor and slip on the Cascadia subduction zone: The chatter of silent slip. Science, 300, 1942-1943.

The boundary between the subducting Juan de Fuca plate and the North America plate can be divided into three segments (Figure 12.7). The cold upper part of the boundary is the locked zone. There the plates are stuck together for long periods of time. When slip does occur, it generates very large earthquakes. The last time the locked zone along Canada’s west coast slipped was January 26, 1700. It caused an earthquake of magnitude 9. The warm lower part of the boundary, called the continuous slip zone, is sliding continuously because the warm rock is weaker. The central part of the boundary, the ETS zone, 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 12.7 Episodic tremor and slip along the Cascadia subduction zone. The Juan de Fuca plate is locked to the North American plate at the top of the subduction zone, but lower down it is slipping continuously. In the intermediate (ETS) zone, the plate alternately sticks and slips on a regular schedule. Source: Steven Earle (2015) CC BY 4.0 view source

It might seem that periodic slip along this part of the plate helps to reduce tension, and thus reduce the risk of a large earthquake. In fact, the opposite is likely the case. The movement along the ETS part of the plate boundary transfers stress to the adjacent locked part of the plate. During the two-week ETS period, the transfer of stress means an increased chance of a large earthquake.

Since 2003, ETS processes have also been observed in subduction zones in Mexico and Japan.

Additional Resources

IRIS Teachable Moment slides for the October 2012 Haida Gwaii earthquake

References

Wang, K., Jiangheng, H., Schulzeck, F., Hyndman, R. D., and Riedel, M. (2015). Thermal Condition of the 27 October 2012 Mw 7.8 Haida Gwaii Subduction Earthquake at the Obliquely Convergent Queen Charlotte Margin. Bulletin of the Seismological Society of America, 105(2B), 1290–1300. doi: 10.1785/0120140183

The shaking from an earthquake travels away from the rupture in the form of seismic waves. Seismic waves are measured to determine the location of the earthquake, and to estimate the amount of energy released by the earthquake (its magnitude).

Types of Seismic Waves

Seismic waves are classified according to where they travel, and how they move particles.

Body Waves

Seismic waves that travel through Earth’s interior are called body waves. P-waves are body waves that move by alternately compressing and stretching materials in the direction the wave moves. For this reason, P-waves are also called compression waves. The “P” in P-wave stands for primary, because P-waves are the fastest of the seismic waves. They are the first to be detected when an earthquake happens.

A P-wave can be simulated by fixing one end of a spring to a solid surface, then giving the other end a sharp push toward the surface (Figure 12.8, top).  The compression will propagate (travel) along the length of the spring. Some parts of the spring will be stretched, and others compressed. Any one point on the spring will jiggle forward and backward as the compression travels along the spring.

Figure 12.8 Seismic waves simulated using a spring and rope attached to a fixed surface. Top: P-waves travel as pulses of compression. Bottom: S-waves move particles at right angles to the direction of motion. Source: Karla Panchuk (2018) CC BY 4.0 modified after Steven Earle (2015) CC BY 4.0 view original.

S-waves are body waves that move with a shearing motion, shaking particles from side to side. S-waves can be simulated by fixing one end of a rope to a solid surface, then giving the other end a flick (Figure 12.8, bottom). Any one point on the rope will move from side to side at a right angle to the direction in which the snaking motion is traveling. The “S” in S-wave stands for secondary, because S-waves are slower than P-waves, and are detected after the P-waves are measured. S-waves cannot travel through liquids.

P-waves and S-waves can travel rapidly through geological materials, at speeds many times the speed of sound in air.

Surface Waves

When body waves reach Earth’s surface, some of their energy is transformed into surface waves, which travel along Earth’s surface. Two types of surface waves are Rayleigh waves and Love waves (Figure 12.9). Rayleigh waves (R-waves) are characterized by vertical motion of the ground surface, like waves rolling on water. Love waves (L-waves) are characterized by side-to-side motion. Notice that the effects of both kinds of surface waves diminish with depth in Figure 12.9.

Surface waves are slower than body waves, and are detected after the body waves. Surface waves typically cause more ground motion than body waves, and therefore do more damage than body waves.

Figure 12.9 Surface waves travel along Earth’s surface and have a diminished impact with depth. Rayleigh waves (left) cause a rolling motion, and Love waves (right) cause the ground to shift from side to side. Source: Steven Earle (2015) CC BY 4.0 view source. Click the image for more attributions.

Recording Seismic Waves Using a Seismograph

A seismometer is an instrument that detects seismic waves. An instrument that combines a seismometer with a device for recording the waves is called a seismograph. The graphical output from a seismograph is called a seismogram. Figure 12.10 (right) shows how a seismograph works. The instrument consists of a frame or housing that is firmly anchored to the ground. A mass is suspended from the housing, and can move freely on a spring. When the ground shakes, the housing shakes with it, but the mass remains fixed. A pen attached to the mass moves up and down on a rotating drum of paper, drawing a wavy line, the seismogram. The seismograph in Figure 12.10 (right) is oriented to measure vertical ground motion. The photo on the left shows a seismograph oriented to record horizontal ground motion.

Figure 12.10 How a seismograph records earthquakes. Source: Left- Karla Panchuk (2018) CC BY-NC-SA 4.0 modified after IRIS (2012) “How Does a Seismometer Work?” view source; Right: Karla Panchuk (2018) CC BY-SA 4.0, photo by Z22 (2014) CC BY-SA 3.0 view source. Click the image for more attributions.

The pen and drum of a mechanical seismograph record the motion of the ground relative to the mass. However, unless an earthquake causes a large amount of ground motion directly beneath the seismograph, the height of the wave recorded on paper might be very small, making the seismogram difficult to analyze. The seismograph on the right has a device to amplify the ground motion, drawing larger waves that are easier to study.

Modern seismographs record shaking as electrical signals, and are able to transmit those signals. This means seismologists need not return to the instrument to collect recordings before the records can be examined.

Finding The Location of an Earthquake

P-waves travel faster than S-waves. As the waves travel away from the location of an earthquake, the P-wave gets farther and father ahead of the S-wave. Therefore, the farther a seismograph is from the location of an earthquake, the longer the delay between when the P-wave arrival is recorded, and the S-wave arrival is recorded. The delay between the P-wave and S-wave arrival appears as a widening gap in a diagram of P-wave and S-wave travel times (Figure 12.11, grey lines).

P-wave and S-wave arrival times can be identified on seismograms. In the three seismograms in Figure 12.11, the arrivals of the P-waves and S-waves are marked with arrows, and the interval in minutes between the P-wave and S-wave arrivals are noted. The seismograms were recorded at three different seismic stations (earthquake monitoring locations equipped with seismographs). The distance of each station from the earthquake is determined by finding the distance along the graph where the gap between the P-wave and S-wave travel-time curves matches the delay between P-wave and S-wave arrivals on the seismogram.

Figure 12.11 Using P-wave and S-wave travel times to determine how far seismic waves have travelled. Grey curves show the distance travelled by P-waves and S-waves after an earthquake occurs. P-waves are faster than S-waves, and the gap between them increases with time and distance. The delays between P-wave and S-wave arrivals on seismograms are matched to the curve to find the distances of seismic stations from the source of the seismic waves. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0 modified after IRIS (n.d.) “How Are Earthquakes Located?” view source

The delay between the P-wave and S-wave arrival at a seismic station can indicate how far the station is from the source of the earthquake, but not the direction from which the seismic waves travelled. The possible locations of the earthquake can be represented on a map by drawing a circle around the seismic station, with the radius of the circle being the distance determined from the P-wave and S-wave travel times (Figure 12.12). If this is done for at least three seismic stations, the circles will intersect at the origin of the earthquake.

Figure 12.12 Locating earthquakes by drawing three circles with radii of lengths determined from P-wave and S-wave travel times. Station names (SOCO, TEIG, SSPA) correspond to seismograms in Figure 12.11. Source: IRIS (n.d.) “How Are Earthquakes Located?” view source Click the image for terms of use.

How Big Was It?

Earthquakes can be described in terms of their magnitude, which reflects the amount of energy released by the shaking. They can also be described in terms of intensity, which characterizes the impact of the shaking on people and their surroundings.

Earthquake Magnitude

Earthquake magnitudes are determined by measuring the amplitudes of seismic waves. The amplitude is the height of the wave relative to the baseline (Figure 12.13). Wave amplitude depends on the amount of energy carried by the wave. The amplitudes of seismic waves reflect the amount of energy released by earthquakes.

Figure 12.13 Seismogram for a small earthquake that occurred near Vancouver Island in 1997. The maximum amplitude of the S-wave is indicated. Source: Karla Panchuk (2018) CC BY 4.0 modified after Steven Earle (2015) CC BY 4.0 view source

The Richter magnitude scale uses the amplitudes of S-waves, and corrects for the decrease in amplitude that happens as the waves travel away from their source. The correction depends on how seismic waves interact with the specific rock types through which they travel, and therefore on local conditions, so the Richter magnitude is also referred to as the local magnitude.

While news reports about earthquakes might still refer to the “Richter scale” when describing magnitudes, the number they report is most likely the moment magnitude. The moment magnitude is calculated from the seismic moment of an earthquake. The seismic moment takes into account the surface area of the region that ruptured, how much displacement occurred, and the stiffness of the rocks. Moment magnitude can capture the difference between short earthquakes and longer ones resulting from larger ruptures, even of both types of earthquakes generate the same amplitude of waves. The moment magnitude scale is also better for earthquakes that are far from the seismic station. Seismic wave measurements are still used to determine the moment magnitude, however different waves are used than for the local magnitude scale.

The magnitude scale is a logarithmic one rather than a linear one- an increase of one unit of magnitude corresponds to a 32 times increase in energy release (Figure 12.14). There are far more low-magnitude earthquakes than high-magnitude earthquakes. In 2017 there were 7 earthquakes of M7 (magnitude 7) or greater, but millions of tiny earthquakes.

Figure 12.14 Earthquake magnitude and corresponding energy release. Energy release increases by approximately 32 times for each unit change in magnitude. Source: IRIS (n.d.) “How Often Do Earthquakes Occur?” view source

Earthquake Intensity

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 now call the Modified Mercalli Intensity Scale (Table 12.1). To determine the intensity of an earthquake, reports are collected about what people felt and how much damage was done. The reports are then used to assign intensity ratings to regions where the earthquake was felt.

Intensity values are assigned to locations, rather than to the earthquake itself. This means that intensity can vary for a given earthquake, depending on the proximity to the epicentre and local conditions. For the 1946 M7.3 Vancouver Island earthquake, intensity was greatest in the central island region (Figure 12.15). In some communities within this region, 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, but with less intensity.

Figure 12.15 Intensity map for the M7.3 Vancouver Island earthquake on June 23, 1946. Source: Earthquakes Canada, Natural Resources Canada (2016) view source. Click the image for terms of use.

Intensity estimates are important as a way to identify regions that are especially prone to strong shaking. A key factor is the nature of the underlying geological materials. The weaker the underlying materials, the more likely it is that there will be strong shaking. Areas underlain by strong solid bedrock tend to experience far less shaking than those underlain by unconsolidated river or lake sediments.

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 near the epicentre, but 350 km away in heavily populated Mexico City there was tremendous damage and approximately 5,000 deaths. The reason 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. Consequently, the shaking was amplified.  Survivors of the disaster recounted that the ground in some areas moved up and down by approximately 20 cm 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.

References

Ammon, C. J. (2001). Earthquake Size Visit website

Bands of earthquakes trace out plate boundaries (coloured dots, Figure 12.16). The depths of earthquakes, and the width of the band, depend on the type of plate boundary. Mid-ocean ridges and transform margins have shallow earthquakes (usually less than 30 km deep), in narrow bands close to plate margins. Subduction zones have earthquakes at a range of depths, including some more than 700 km deep. Bands of earthquakes are wider along subduction zones because they take place throughout the subducting slab that extends beneath the opposing plate. Wide swaths of scattered earthquakes may correspond to continent-continent collision zones, such as between the Eurasian plate and the African, Arabian, and Indian plates to the south. Wide swaths of scattered earthquakes may also correspond to continental rift zones, such as in eastern Africa.

Figure 12.16 Earthquakes greater than magnitude 5, from 2000 to 2008. Bands of earthquakes mark tectonic plates. Narrow bands with shallow earthquakes (marked in red) indicate transform boundaries or mid-ocean ridge divergent boundaries. Wider bands with earthquakes at a range of depths are subduction zones. Wide bands of scattered earthquakes mark continent-continent convergent margins (e.g., between the Indian and Eurasian plates), or continental rift zones (e.g., in eastern Africa). Source: Lisa Christiansen, Caltech Tectonics Observatory (2008) view source. Plate and ocean basin labels added. Click the image for terms of use.

Earthquakes are also relatively common at a few locations away from plate boundaries. 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. Locations include the Great Rift Valley area of Africa, the Lake Baikal area of Russia, and Tibet.

Earthquakes at Divergent and Transform Plate Boundaries

Earthquakes along divergent and transform plate margins are shallow (usually less than 30 km deep) because below those depths, rock is too hot and weak to avoid being permanently deformed by the stresses in those settings. If deformation is permanent, then removing the stress does not result in the rocks snapping back to their original shape. No snapping back means no shaking.

Mid-ocean ridge divergent plate margins are offset by numerous transform faults (Figure 12.17). The locations of earthquakes along mid-ocean ridges, and the mechanisms for causing them, depend on how rapidly the mid-ocean ridges are spreading. The Pacific-Antarctic Ridge (left) is spreading relatively rapidly at 42 to 94 mm/year, depending on the location along the ridge.  Rapid spreading causes rocks near the axis of the spreading centre to be hot and weak. As a result, most of the earthquakes (white dots) are located along transform faults, where rocks are cooler and stronger. Along rapidly spreading ridges, new ocean crust is bent upward into wide, high ridges. As spreading proceeds and crust moves away from the ridge, the bend is relaxed, and the crust stretches and breaks. This triggers earthquakes many kilometres away from the ridge.

Figure 12.17 Locations of earthquakes of magnitude 4 and greater from 1990 to 2010 along two mid-ocean ridges. Plate boundaries are marked in red. Arrows show the direction of plate motion. Left: Rapidly spreading Pacific-Antarctic ridge with earthquakes concentrated along transform faults. Right: Slowly spreading Southwest Indian Ridge, with earthquakes along both spreading segments and transform faults. Source: Karla Panchuk (2017) CC BY 4.0. Base maps with epicentres generated using the U. S. Geological Survey Latest Earthquakes website. Visit Latest Earthquakes

The Southwest Indian Ridge (right) spreads very slowly, at approximately 14 mm/year. Rocks are cooler and stronger along the slowly spreading ridge than along the rapidly spreading one. In the slow-spreading environment, earthquakes are generated when rocks along the ridge axis stretch and break. Earthquakes are more evenly distributed between divergent and transform segments of the boundary than they are along fast-spreading ridges.

Earthquakes in continental rift zones are also shallow, but scattered more broadly than those along mid-ocean ridges. Lake Baikal (Figure 12.28), the world’s oldest, deepest, and largest freshwater lake, formed 25 million years ago because of continental rifting. Note the scale in Figure 12.18, and compare how widely the earthquakes (blue dots) are spread in the Lake Baikal region, versus along the mid-ocean ridges in Figure 12.17.

Figure 12.18 Blue circles mark the locations of earthquakes of M4 and greater from 1990 to 2010 along the Lake Baikal rift zone. White lines show some of the faults in the region. White lines with tick marks are normal faults related to spreading. Arrows show the direction of spreading. White lines without tick marks are transform faults. The Siberian Craton (shaded region) is strong 2 billion year old crust. Source: Karla Panchuk (2017) CC BY 4.0. Faults after U. S. Geological Survey (see references). Base maps (inset and rift views) with epicentres generated using the U. S. Geological Survey Latest Earthquakes website. Visit Latest Earthquakes

One reason for the difference in earthquake distribution in continental rift zones is that the rifts are only beginning to form. Faulting is “disorganized” within the continental crust. There is no well-established spreading centre, unlike mid-ocean ridges. Another reason is that the locations of faults, and thus earthquakes, in continental rift zones are affected by pre-existing geological structures within continental crust. In the case of the Lake Baikal rift, the strong, ancient crust of the Siberian Craton influences the orientation of the faults forming the rift. Faults run parallel to the craton near Lake Baikal. As rifting extends to the east, the part of the craton in the upper right of Figure 12.18 may deflect rifting southward.

Earthquakes at Convergent Boundaries

Subduction Zones

Along convergent plate margins with subduction zones, earthquakes range from shallow to depths of up to 700 km. Earthquakes occur where the two plates are in contact, as well as in zones of deformation on the overriding plate, and along the subducting slab deeper within the mantle. The result is that epicentres of earthquakes farther to the interior of the overriding plate will correspond to increasingly deep earthquakes.

Where the Pacific plate subducts beneath the North American plate, forming the Aleutian volcanic arc (Figure 12.19), earthquakes increase in depth moving northward, following the Pacific plate into the mantle. Earthquakes between 0 and 33 km deep (red circles) occur closest to the subduction zone (red line; teeth point in the direction of the subducting slab). While there is some overlap, earthquakes between 33 and 70 km deep (white circles) occur in a band that reaches farther the north. Farthest north are the epicentres for earthquakes between 70 and 300 km deep (green dots). The deepest earthquake during the seven year interval shown in Figure 12.19 is represented by the large green dot farthest to the north. It occurred at a depth of 265 km.

Figure 12.19 Earthquakes of M4.5 and greater from 2010 to 2017 along the Aleutian Trench subduction zone (red line; teeth point in the direction of the subducting slab). White arrows show the directions of plate movement. Circle colours indicate the depths of earthquakes (see legend, lower left). Earthquakes become deeper moving north from the subduction zone. Source: Karla Panchuk (2017) CC BY 4.0. Base maps with epicentres generated using the U. S. Geological Survey Latest Earthquakes website. Visit Latest Earthquakes

Earthquakes occur in subduction zones for a variety of reasons. Stresses associated with the collision of two plates cause deformation in the overriding plate, and thus shallow earthquakes. Shallow earthquakes also happen on the subducting slab when a locked zone (orange line, Figure 12.20) ruptures. The locked zone is where the largest earthquakes on Earth, called megathrust earthquakes, occur. There is the potential for a wider rupture zone on a gently dipping subduction zone boundary compared to other boundaries.

Figure 12.20 Factors contributing to earthquakes in subduction zones. Not all factors shown here are present in all subduction zones. Stress from bending, flexing, and stretching may cause ruptures. Changes in the mechanical properties of the mantle may affect how subucting slabs move, contributing additional stresses. The histogram at right shows the global average number of earthquakes at depth, for earthquakes greater than M5. The increase in earthquakes beneath 480 km may be caused in part by weakening as olivine transforms into high pressure/temperature polymorphs. Source: Karla Panchuk (2017) CC BY 4.0. Modified after Green et al. (2010).

If subduction is rapid, the subducting plate will bend more as it enters the mantle (slab A in Figure 12.20), causing the upper edge of the plate to stretch, and the interior and lower edge to be compressed. Stress from bending can cause shallow to intermediate earthquakes on these plates. Even without bending, the subducting slab can become stretched by its own weight as it falls into the mantle.

The 410 km and 660 km discontinuities in Figure 12.20 mark boundaries where minerals transform into other, denser minerals that are stable at higher pressures and temperatures. When the subducting slab reaches the 660 km discontinuity (the top of the lower mantle), the increase in density in the surrounding mantle may slow down the leading edge of the sinking slab. Earthquakes can be generated when the slab is compressed by the lower mantle resisting its motion at the same time that the upper part of the slab continues to fall.

Slower rates of subduction mean that the subducting slab will enter the mantle at a lower angle (slab B in Figure 12.20). These slabs might not have earthquakes from being bent downward into the mantle, as with slab A, but earthquakes may be triggered by changes in stress if the plate relaxes and unbends.

The bar chart on the right of Figure 12.20 shows global average number of earthquakes that occur at different depths. Earthquakes are most abundant at the surface, and then fall to a minimum at 300 km. The number of earthquakes remains low until almost 500 km depth, and reaches a second peak around 600 km depth. The second peak might be explained by interactions between the subducting plate and the dense mantle beneath the 660 km discontinuity, but another hypothesis is that it is related to delayed mineral transformations. The subducting slab warms as it goes deeper into the mantle, but the warming is not uniform. The outer edges of the slab will warm before the interior does. The 410 km discontinuity is where olivine is transformed into the minerals wadsleyite and ringwoodite under normal mantle pressure and temperature conditions. However, if the interior of the subducting slab is still too cool at that depth, olivine will be retained to depths below 410 km. Olivine weakens prior to transforming into the high pressure minerals, and the weakening may make it easier for the slab to rupture.

Continent-Continent Convergence Zones

Where continents collide, earthquakes are scattered over a much wider area compared to earthquakes along mid-ocean ridges, transform margins, or subduction zones. An example is where the Indian plate collides with the Eurasian plate (Figure 12.21). At one time, India was a separate continent, and ocean crust separated India from the Eurasian plate. For a time, a subduction zone existed where ocean lithosphere from the Indian plate subducted beneath the Eurasian plate. But when the two land masses finally met, they became locked together and the subduction zone was closed off. Today the Indian plate is still pushing against the Eurasian plate in the regions indicated by the red arrows in Figure 12.21. The collision is accommodated by transform boundaries along the Indian plate. Regions of overall transform motion are indicated in Figure 12.21 with blue arrows.

Figure 12.21 Earthquakes of M4.5 and greater from 1990 to 2017 along the collision zone between the Indian and Eurasian plates. Red lines- plate boundaries; red arrows- collision zones; blue arrows- transform zones. Source: Karla Panchuk (2017) CC BY 4.0. Base maps with epicentres generated using the U. S. Geological Survey Latest Earthquakes website. Visit Latest Earthquakes

The majority of earthquakes in Figure 12.21 occur at depths less than 70 km, however they are still abundant down to 150 km, and extend to more than 300 km depth at some locations. Deeper earthquakes may be caused by continued northwestward subduction of part of the Indian plate beneath the Eurasian plate in this area. Even though the area is no longer a subduction zone, the subducted slab still remains, and is subject to stresses that can trigger earthquakes.

Some of the earthquakes in Figure 12.21 are related to the transform faults on either side of the Indian plate, and most of the others are related to the squeezing caused by the continued convergence of the Indian and Eurasian plates. That squeezing has caused the Eurasian plate to be thrust over the Indian plate, building the Himalayas and the Tibet Plateau to enormous heights. Most of the earthquakes of Figure 12.21 are related to the thrust faults shown in Figure 12.22 (and to hundreds of other similar ones that cannot be shown at this scale). The southernmost thrust fault in Figure 12.22 (the Main Boundary Fault) is equivalent to the convergent boundary in Figure 12.21.

Figure 12.22 Schematic diagram of the India-Asia convergent boundary, showing examples of the types of faults along which earthquakes are focused. Source: Steven Earle (2015) CC BY 4.0 view source after D. Vuichard (Figure 2.3) in Ives and Messerli (1989).

Intraplate Earthquakes

Intraplate earthquakes (within-plate earthquakes) are those that occur away from plate boundaries. Some intraplate earthquakes are related to human activities. When humans trigger earthquakes it is referred to as induced seismicity. In Saskatchewan there have been 20 earthquakes since 1985 (all less than magnitude 4), and the majority occurred near potash mines. Excavation changes the stress in surrounding rocks, so earthquakes may occur in the rocks above excavated parts of the mine. In Alberta, induced seismicity is triggered by hydraulic fracturing operations when water pressure increases along existing faults, causing them to slip.

Intraplate earthquakes not related to human activities often occur along ancient rift zones. In eastern Canada, the Charlevoix seismic zone (approximately 100 km northeast of Québec City; Figure 12.23), is associated with a rift-zone faults that developed when an ancient ocean basin began to form more than 500 million years ago. Coincidentally, the rocks of the Charlevoix Seismic Zone are also fragmented because of a meteorite impact (the crater margin is indicated by the blue circle in Figure 12.23), weakening them further. While the Charlevoix zone is far from any boundary of the North American plate, tectonic forces acting on plate boundaries are still transmitted to the interior of the continent, contributing to the stress that causes the faults along the rift zone to rupture.

Figure 12.23 Charlevoix seismic zone, site of intraplate earthquakes. The location of the Charlevoix seismic zone is indicated by the star on the map of Canada. Dots indicate earthquake epicentres. The size of the dots indicates magnitude. White lines indicate fault segments. The dashed circle marks the edge of a crater formed by a meteorite impact 342 million years ago. Source: Karla Panchuk (2017) CC BY-SA 4.0. Epicentre data from Earthquakes Canada. Click the image for more attributions and data sources.

Intraplate earthquakes can be large earthquakes. The Charlevoix seismic zone has had five earthquakes of magnitudes between 6 and 7 since 1663. The New Madrid seismic zone in the Mississippi River Valley had a series of four earthquakes with magnitudes between 7 and 8 in the winter of 1811-1812. The population of the region was sparse at the time, but today there are major cities in the New Madrid seismic zone, including Memphis, Tennessee, and St. Louis, Missouri.

References

Buck, W. R., Lavier, L. L., & Poliakov, A. N. B. (2005). Modes of faulting at mid-ocean ridges. Nature 434(7034), 719-23. doi: 10.1038/nature03358

Coastal and Marine Geology Program, U. S. Geological Survey (n.d.). Lake Baikal – A Touchstone for Global Change and Rift Studies. Visit website

DeMets, C., Gordon, R. G., & Argus, D. F. (2010). Geologically current plate motions. Geophysical Journal International 181, 1-80. doi: 10.1111/j.1365-246X.2009.04491.x

Earthquakes Canada, Natural Resources Canada (2016) Earthquake zones in Eastern Canada Visit website

Earthquakes can have direct impacts, such as structural damage to buildings from shaking, and secondary impacts, such as triggering landslides, fires, and tsunami. The types and extent of impacts will depend on local conditions where the earthquake strikes. The geological materials in the area matter, as does the type of terrain, and whether the region is near the coast or not. The extent of impact and type of damage will depend on whether the area is predominantly urban or rural, densely or sparsely populated, highly developed or underdeveloped. It will depend on whether the infrastructure has been designed to withstand shaking.

Damage to Structures from Shaking

As with the example of the 1985 Mexico earthquake, the geological foundations on which structures are built will affect the amount of shaking that occurs. Earthquakes produce seismic waves that vibrate at different rates, or frequencies. Waves with rapid vibrations have a high frequency, and waves with slower vibrations have lower frequencies.

The energy of the higher frequency waves tends to be absorbed by solid rock. Lower frequency waves pass through solid rock without being absorbed, but are 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. During the 1989 Loma Prieta earthquake, parts of a two-layer highway in the Oakland area near San Francisco collapsed where they were built on soft sediments (Figure 12.24).

Figure 12.24 A collapsed section of the Cypress Freeway in Oakland California. Source: U. S. Geological Survey (1989) Public Domain view source

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, but even though Turkey had a relatively strong building code in the 1990s, adherence to the code was poor. Builders did whatever they could to save costs, including using inappropriate materials in concrete, and reducing the amount of steel reinforcing. The result was more than 17,000 deaths in the 1999 M7.6 Izmit earthquake (Figure 12.25). 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 12.25 Damage from the 1999 M7.6 Izmit, Turkey earthquake. Source: Left; USGS (1999) Public Domain view source; Right: USGS (1999) Public Domain view source

Structures underlain by sediments may be at risk of another hazard, called liquefaction, in which sediment is transformed into a fluid. When water-saturated sediments are shaken, the grains may lose contact with each other, and no longer support one another. Water between the grains holds them apart, causing the sediment to turn to mud and flow. The loss of support can lead to the collapse of buildings or other structures that might otherwise have sustained little damage. During the 1964 M7.6 earthquake in Niigata, Japan, liquefaction caused buildings to sink into the sediments (Figure 12.26).

Figure 12.26 Collapsed apartment buildings in the Niigata area of Japan. The material beneath the buildings was liquefied by the 1964 earthquake. Source: DOC/NOAA/NESDIS/NCEI (1964) Public Domain view source

Parts of the Fraser River delta are also prone to liquefaction-related damage. The region is characterized by a 2 m to 3 m thick layer of fluvial silt and clay above a layer of water-saturated fluvial sand that is at least 10 m thick (Figure 12.27). Under these conditions, seismic shaking can be amplified, and the sandy sediments will liquefy. This could lead to subsidence and tilting of buildings. 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. Current building-code regulations in the Fraser delta area require that measures be taken to strengthen the ground beneath multi-story buildings prior to construction.

Figure 12.27 Recent unconsolidated sedimentary layers in the Fraser River delta area (top) and the potential consequences in the event of a damaging earthquake. Source: Steven Earle (2015) CC BY 4.0 view source

Experiments with Liquefaction

To see liquefaction for yourself, go to a sandy beach and find a place near the water’s edge where the sand is wet. While standing in one place on a wet part of the beach, start moving your feet up and down at a frequency of about once per second. Within a few seconds the previously firm sand will start to lose strength, and you’ll gradually sink in up to your ankles.

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

Figure 12.28 Fine, damp sand before shaking (left) and after (right). Source: Steven Earle (2015) CC BY 4.0 view source

As you were moving your feet up and down or thumping the container, it’s likely that you soon discovered the most effective rate for getting the sand to liquefy. 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.

Slope Failure

Ground shaking during an earthquake can be enough to weaken rock and loose materials to the point of failure. Earthquakes can also trigger failures on slopes that are already weak. In January of 2011 a M7.6 earthquake offshore of El Salvador triggered slope failures that killed nearly 600 people (Figure 12.29).

Figure 12.29 A slope gives way in a suburb of San Salvador after the January 2001 earthquake offshore of El Salvador. This is one of hundreds of slope failures resulting from the earthquake. Source: U. S. Geological Survey (2001) Public Domain view source

Fires

Fires are commonly associated with earthquakes because gas lines rupture and electrical lines are damaged when the ground shakes. Most of the damage in the great 1906 San Francisco earthquake was caused by massive fires in the downtown area of the city (Figure 12.30). Some 25,000 buildings were destroyed by those fires, which were fueled by gas leaking from broken pipes. Fighting the fires was difficult because water mains had also ruptured. Today the risk of fires can be reduced through P-wave early warning systems if utility operators can decrease pipeline pressure and break electrical circuits.

Figure 12.30 Fires in San Francisco following the 1906 earthquake. Source: Pillsbury Picture Co. (1906) Public Domain, courtesy of the Library of Congress Prints and Photographs Division view source

Tsunami

Large earthquakes that take place beneath the ocean have the potential to displace large volumes of water. In a subduction zone, for example, the overriding plate becomes distorted by elastic deformation. It is squeezed laterally and pushed up (Figure 12.31 top). When an earthquake happens, the plate rebounds over an area of thousands of square kilometres, generating waves- a tsunami (Figure 12.31, middle). The waves spread across the ocean at velocities of several hundred kilometres per hour. Tsunami can make it to the far side of an ocean in about the same time as a passenger jet.

Figure 12.31 Tsunami triggered along a subduction zone. Top: The overriding crust is deformed because it is locked to the subducting slab. Middle: When the locked zone ruptures, the crust rebounds, and waves are created. Bottom: Tsunami waves have small amplitudes in the deep ocean, but once in shallow water, they slow down, causing the waves to become taller and closer together. Source: Karla Panchuk (2018) CC BY-NC-SA 4.0. Top and middle modified after Steven Earle (2015) CC BY 4.0 view source. Bottom modified after COMET/UCAR (1997-2017) view source. Click the image for COMET/UCAR attribution and terms of use.

Tsunami waves gain their height as they travel through shallower waters. In the deep ocean, the waves may be so small as to go undetected by ships, but when they are slowed down by interacting with the ocean floor, they can become much taller (Figure 12.31, bottom). In the tsunami following the 2004 Sumatra earthquake, the tallest waves were more than 30 m high.

Subduction earthquakes must be large to generate significant tsunami. 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 can have an impact across an entire ocean basin. They spread across the ocean at velocities of several hundred kilometres per hour, and can make it to the far side of an ocean in about the same time as a passenger jet. In 1700 a rupture along the Cascadia thrust fault running from Vancouver Island to northern California resulted in a M9 earthquake. It generated a tsunami that travelled across the Pacific Ocean, and was recorded in Japan nine hours later. A computer simulation of the tsunami (Figure 12.32) shows how long it took tsunami waves from the Cascadia earthquake to travel across the Pacific Ocean, and how high the waves were. Notice that over all, the waves decrease in height moving away from the rupture, but they increase in height again as they reach the opposite side of the Pacific Ocean.

Figure 12.32 Computer simulation of the tsunami from the 1700 M9 Cascadia earthquake. Colours show open-ocean wave heights. Contours show travel time in hours. Tsunami wave heights increase as the tsunami reached the western margin of the Pacific ocean. Source: NOAA/PMEL/Center for Tsunami Research (2011) Public Domain view source / view context

References

Earthquakes Canada, Natural Resources Canada (2016). The M9 Cascadia Megathrust Earthquake of January 26, 1700. Visit website

University Corporation for Atmospheric Research (2010). Propagation. Tsunamis. Visit website

U. S. Geological Survey (2014). Indian Ocean Tsunami Remembered — Scientists reflect on the 2004 Indian Ocean that killed thousands. Visit website

It has long been a dream of seismologists, geologists, and public safety officials to be able to accurately predict the location, magnitude, and timing of earthquakes on time scales that would be useful for minimizing danger to the public and damage to infrastructure (e.g., weeks, days, hours). Many methods of prediction that have been explored. These include using observations of warning foreshocks, changes in magnetic fields, episodic tremor and slip, changing groundwater levels, strange animal behaviour, patterns in the timing between earthquakes, and how stress is transferred after a rupture. So far, none of these has provided a reliable method. Although there are some reports of successful earthquake predictions, they are rare, and many are surrounded by doubtful circumstances.

The Parkfield Prediction Experiment

There was great hope for earthquake predictions late in the 1980s when attention was focused on part of the San Andreas Fault at Parkfield, approximately 200 km south of San Francisco, California. Between 1881 and 1965 there were five earthquakes at Parkfield. They were spaced at approximately 20-year intervals, all confined to the same 20 km-long segment of the fault, and all very close to M6 (Figure 12.33). Both the 1934 and 1966 earthquakes were preceded by small foreshocks exactly 17 minutes before the main quake.

Figure 12.33 Earthquakes on the Parkfield segment of the San Andreas Fault between 1881 and 2004. Source: Steven Earle (2015) CC BY 4.0 view source

The U. S. Geological Survey recognized this as an excellent opportunity to understand earthquakes and earthquake prediction, so they armed the Parkfield area with a huge array of geophysical instruments and waited. The next earthquake was expected to happen around 1987, but nothing happened! The “1987 Parkfield earthquake” finally struck in September 2004. Fortunately, all of the equipment was still there to record the earthquake, but it was no help from the perspective of earthquake prediction. There were no significant precursors to the 2004 Parkfield earthquake in any of the parameters measured, including tremors, changes in rock deformation, the magnetic field, the electrical conductivity of the rock, and creep (motion along the fault that is not accompanied by earthquakes). There was no foreshock. In other words, even though every available technique was used to monitor it, the 2004 earthquake came with no warning whatsoever.

Earthquake Probabilities

To be useful to the public and governments, earthquake predictions must be accurate most of the time, not just some of the time. If a prediction method is only accurate 10% of the time (and even that isn’t possible with the current state of knowledge), the public will lose faith in the process very quickly, and then will ignore all of the predictions. The hope for earthquake prediction is not dead, but it was hit hard by the Parkfield experiment.

Today the focus of efforts in earthquake-prone regions is to provide forecasts of earthquake probability. Earthquake probabilities express the likelihood that an earthquake of a given magnitude will occur at a location within a given period of time. An example of this approach for the San Francisco Bay region of California is shown in Figure 12.34. Based on a wide range of information, including past earthquake history, accumulated stress from plate movement, and known stress transfer, seismologists and geologists have predicted the likelihood of a M6.7 or greater earthquake on each of eight major faults that cut through the region. The greatest probabilities are on the San Andreas, Rogers Creek/Hayward faults, and Calaveras/Paicines faults. There is a 72% chance that a major and damaging earthquake will take place somewhere in the region prior to 2043.

Figure 12.34 Earthquake outlook for 2014-2047. Probabilities for individual faults are marked on the faults. Source: U. S. Geological Survey (2014) Public Domain view report

Using Earthquake Probability Information

Decision makers can use forecasts of earthquake probability to assist with educating the public about earthquake risks, and to determine what action is necessary to make infrastructure earthquake-safe. Building safe infrastructure requires strong building codes, and enforcement of those codes. Building code compliance is robust in most developed countries, but is inadequate in many developing countries.

New buildings are not the only ones requiring attention. Existing buildings — especially schools and hospitals — and other structures such as bridges and dams, must also be made safe. British Columbia began a multi-billion-dollar program in 2004 to make public schools safer for students. The program is focused on older public schools, because, according to the government, those built since 1992 already comply with modern seismic codes. Some schools would require too much work to make upgrading economically feasible and they are replaced. Where upgrading is feasible, the school is assessed carefully before any upgrade work is initiated. The seismic mitigation program identified 346 schools as being at high risk and in need of upgrades. As of December 2017, upgrades were completed at 168 schools, 28 schools were under construction or had approval to proceed with construction, and 150 did not yet have plans in place for upgrades.

The program in British Columbia illustrates a challenge with seismic upgrades of public buildings, which is that governments must make adequate funds available for the upgrades to be done in a timely manner. The priority allocated to funding those projects will depend on how urgent the need for upgrades is considered to be, given other demands on public funds.

Earthquake Preparedness

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

References

Finn, W. D. L., & Dexter, A. (2012). Risk Management Plan for School Seismic Mitigation Program. View report

Province of British Columbia (n.d.). Seismic Mitigation Program. Visit website

Province of British Columbia (2017). Seismic Mitigation Program Progress Report. View report

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

12.1 What Is an Earthquake?

An earthquake is the shaking that results when a deformed body of rock snaps back to its original shape. The rupture is initiated at a point but quickly spreads across the area of a fault, with aftershocks initiated by stress transfer. Episodic tremor and slip is a periodic slow movement, accompanied by harmonic tremors, along the middle part of a subduction zone boundary.

12.2 Measuring Earthquakes

Earthquakes produce seismic waves that can be measured by a seismograph. The amplitudes of seismic waves are used to determine the amount of energy released by an earthquake- its magnitude. For the moment magnitude scale used today, the amount of energy released by an earthquake is proportional to the size of the rupture surface, the amount of displacement, and the strength of the rock. Intensity is a measure of the amount of shaking that occurs, and damage done at locations that experience an earthquake. Intensity will vary depending on the distance to the epicentre, the depth of the earthquake, and the type of geological materials present.

12.3 Earthquakes and Plate Tectonics

Most earthquakes happen at or near plate boundaries. Along divergent and transform boundaries earthquakes are shallow (less than 30 km depth), but at convergent boundaries they can be hundreds of kilometers beneath the surface. The largest earthquakes happen when a broad segment of the locked zone of a subduction zone ruptures. Intraplate earthquakes happen away from plate boundaries. They can be caused by human activities, or renewed motion on ancient faults.

12.4 The Impacts of Earthquakes

Damage to buildings is the most serious consequence of most large earthquakes. The amount of damage is related to the type and size of buildings, how they are constructed, and the nature of the material on which they are built. Other important consequences are fires, damage to bridges and highways, slope failures, liquefaction, and tsunami.

12.5 Forecasting Earthquakes and Minimizing Impacts

There is no reliable technology for predicting earthquakes, but the probability of one happening within a certain time period can be forecast. We can minimize earthquake impacts by ensuring that the public is aware of the risk, that building codes are enforced, that existing buildings like schools and hospitals are seismically sound, and that both public and personal emergency plans are in place.

Review Questions

1. What causes the shaking during an earthquake?
2. What is a rupture surface, and how does the area of a rupture surface relate to earthquake magnitude?
3. What is an aftershock and how are aftershocks related to stress transfer?
4. Episodic slip on the middle part of the Cascadia subduction zone is thought to increase the stress on the locked zone. Why?
5. What is the difference between the magnitude of an earthquake and its intensity?
6. How much more energy is released by a magnitude 7 earthquake compared to a magnitude 5 earthquake?
7. The images below show earthquake locations for three regions of ocean lithosphere. The colour scheme indicates the depths of earthquakes. a) The images show a subduction zone, a slowly spreading mid-ocean ridge, and a rapidly spreading mid-ocean ridge. Which is which? b) In the image with the subduction zone, which side is the subducting plate, and which is the overriding plate?
   
   

   Figure 12.35 Earthquakes along plate margins. Dots are colour coded according to depth. Source: Map details from Lisa Christiansen, Caltech Tectonics Observatory (2008) view source. Click the image for terms of use.
8. Why is earthquake damage likely to be more severe for buildings built on unconsolidated sediments as opposed to on solid rock?
9. Why are fires common during earthquakes?
10. What type of earthquake is likely to lead to a tsunami?
11. What did we learn about earthquake prediction from the 2004 Parkfield earthquake?
12. What are some of the things we should know about an area in order to help minimize the impacts of an earthquake?
13. What is the difference between earthquake prediction and forecasting earthquake probability?

1. Rocks under stress can deform elastically. When rocks break, or a rupture occurs along an existing fault, the deformed rocks snap back to their original shape, causing vibrations. The vibrations are the shaking felt during earthquakes.

2. The rupture surface is the surface over which there is displacement of rock during an earthquake. The magnitude of an earthquake is proportional to the area of the rupture surface.

3. An aftershock is an earthquake that is triggered when a rupture from a previous earthquake has transferred too much stress to rocks, causing further ruptures.

4. Episodic slip on the middle part of the Cascadia subduction zone decreases stress within that area, but some of that stress is transferred to the locked zone along the plate boundary. This increases the amount of stress on the locked part.

5. Magnitude is the amount of energy released by an earthquake. Each earthquake has only one magnitude, although there are different ways of measuring it, and they may give slightly different results. Intensity is a measure of the amount of damage done or what people felt. Intensity varies depending on the distance to the epicentre and the type of rock or sediment underlying an area.

6. Each unit increase in magnitude corresponds to a 32x increase in energy. The difference between M5 and M7 is two units of magnitude. An earthquake that is two magnitude units larger would release 32 x 32 = 1,024 times as much energy.

7.  a) A: Slowly spreading mid-ocean ridge- earthquakes are shallow and evenly distributed; B: Subduction zone- earthquakes range in depth; C: Rapidly-spreading mid-ocean ridge- earthquakes are shallow, and occur in patches, corresponding to a greater likelihood of earthquakes where transform faults offset ridges.

b) The plate to the right of the chain of earthquakes is the subducting plate. The earthquakes increase in depth moving from right to left, indicating that the subducting slab is getting deeper farther to the left.

8. Unconsolidated sediments, especially if they are saturated with water, can lose strength when subjected to earthquake shaking. This can cause buildings to subside or tilt. Unconsolidated sediments can also amplify the vibrations of an earthquake.

9. Gas lines and electrical transmission wires are typically damaged during an earthquake, and this can lead to serious fires.

10. A large subduction-zone earthquake (greater than M7.5) can generate a tsunami because those earthquakes result in sufficient vertical displacement of the sea floor.

11. The 2004 Parkfield earthquake showed that we cannot rely on foreshocks to predict earthquakes, or on any of the many other parameters that were being carefully measured around Parkfield in the years leading up to the quake.

12. We should know about the history of past large earthquakes, the typical locations of small earthquakes, the types of geological materials beneath the surface (especially soft water-saturated sediments), the types of infrastructure that is present, and the various ways that people can be evacuated from an area or assistance can be brought in.

13. Forecasting involves estimating the risk of an earthquake happening in a region within a period of time. Prediction involves stating that an earthquake is likely to happen at a certain location on a specific day or month or year in the future. With our current state of knowledge of earthquakes, prediction is not possible.