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Tides are related to very long-wavelength but low-amplitude waves on the ocean surface (and to a much lesser extent on very large lakes) that are caused by variations in the gravitational effects of the Sun and Moon. Tide amplitudes in shoreline areas vary quite dramatically from place to place. On the west coast of Canada, the tidal range is relatively high, in some areas as much as 6 m, while on most of the east coast the range is lower, typically around 2 m. A major exception is the Bay of Fundy between Nova Scotia and New Brunswick, where the daily range can be as great as 16 m. Anomalous tides like that are related to the shape and size of bays and inlets, which can significantly enhance the amplitude of the tidal surge. The Bay of Fundy has a natural oscillation cycle of 12.5 hours, and that matches the frequency of the rise and fall of the tides in the adjacent Atlantic Ocean. Ungava Bay, on Quebec’s north coast, has a similarly high tidal range. As the tides rise and fall they push and pull a large volume of water in and out of bays and inlets and around islands. They do not have as significant an impact on coastal erosion and deposition as wind waves do, but they have an important influence on the formation of features within the intertidal zone, as we’ll see in the following sections.

When waves approach an irregular shore, they are slowed down to varying degrees, depending on differences in the water depth, and as they slow, they are bent or refracted. In Figure 17.11, wave energy is represented by the red arrows. That energy is evenly spaced out in the deep water, but because of refraction, the energy of the waves — which moves perpendicular to the wave crests — is being focused on the headlands (Frank Island and Cox Point in this case). On irregular coasts, the headlands receive much more wave energy than the intervening bays, and thus they are more strongly eroded. The result of this is coastal straightening. An irregular coast, like the west coast of Vancouver Island, will eventually become straightened, although that process will take millions of years. Figure 17.11 The approach of waves (white lines) in the Cox Bay area of Long Beach, Vancouver Island. The red arrows represent wave energy; most of that energy is focused on the headlands of Frank Island and Cox Point. [SE] Wave erosion is greatest in the surf zone, where the wave base is impinging strongly on the sea floor and where the waves are breaking. The result is that the substrate in the surf zone is typically eroded to a flat surface known as a wave-cut platform (or wave-cut terrace) (Figure 17.12). A wave-cut platform extends across the intertidal zone.

Relatively resistant rock that does not get completely eroded during the formation of a wave-cut platform will remain behind to form a stack. An example from the Juan de Fuca Trail of southwestern Vancouver Island is shown in Figure 17.13. Here the different layers of the sedimentary rock have different resistance to erosion. The upper part of this stack is made up of rock that resisted erosion, and that rock has protected a small pedestal of underlying softer rock. The softer rock will eventually be eroded and the big rock will become just another boulder on the beach.

Arches and sea caves are related to stacks because they all form as a result of the erosion of relatively non-resistant rock. An arch in the Barachois River area of western Newfoundland is shown in Figure 17.14. This feature started out as a sea cave, and then, after being eroded from both sides, became an arch. During the winter of 2012/2013, the arch collapsed, leaving a small stack at the end of the point. If you look carefully at the upper photograph you can see that the hole that makes the arch developed within a layer of relatively soft and weak rock. Figure 17.14 Top: An arch in tilted sedimentary rock at the mouth of the Barachois River, Newfoundland, July 2012. Bottom: The same location in June 2013. The arch has collapsed and a small stack remains. [Photo: Dr. David Murphy, used with permission] Figure 17.15 summarizes the process of transformation of an irregular coast, initially produced by tectonic uplift, into a straightened coast with sea cliffs (wave-eroded escarpments) and the remnants of stacks, arches, and wave-cut platforms. The next stages of this process would be the continued landward erosion of the sea cliffs and the complete erosion of the stacks and wave-cut platforms in favour of a continuous and nearly straight sandy beach.

Some coastal areas are dominated by erosion, an example being the Pacific coast of Canada and the United States, while others are dominated by deposition, examples being the Atlantic and Caribbean coasts of the United States. But on almost all coasts, both deposition and erosion are happening to varying degrees most of the time, although in different places. This is clearly evident in the Tofino area of Vancouver Island (Figure 17.1), where erosion is the predominant process on the rocky headlands, while depositional processes predominate within the bays. On deposition-dominant coasts, the coastal sediments are still being eroded from some areas and deposited in others. The main factor in determining if a coast is dominated by erosion or deposition is its history of tectonic activity. A coast like that of British Columbia is tectonically active, and compression and uplift have been going on for tens of millions of years. This coast has also been uplifted during the past 15,000 years by isostatic rebound due to deglaciation. The coasts of the United States along the Atlantic and the Gulf of Mexico have not seen significant tectonic activity in a few hundred million years, and except in the northeast, have not experienced post-glacial uplift. These areas have relatively little topographic relief, and there is now minimal erosion of coastal bedrock. On coasts that are dominated by depositional processes, most of the sediment being deposited typically comes from large rivers. An obvious example is where the Mississippi River flows into the Gulf of Mexico at New Orleans; another is the Fraser River at Vancouver. There are no large rivers bringing sandy sediments to the west coast of Vancouver Island, but there are still long and wide sandy beaches there. In this area, most of the sand comes from glaciofluvial sand deposits situated along the shore behind the beach, and some comes from the erosion of the rocks on the headlands. The components of a typical beach are shown in Figure 17.16. On a sandy marine beach, the beach face is the area between the low and high tide levels. A berm is a flatter region beyond the reach of high tides; this area stays dry except during large storms.

Most beaches go through a seasonal cycle because conditions change from summer to winter. In summer, sea conditions are relatively calm with long-wavelength, low-amplitude waves generated by distant winds. Winter conditions are rougher, with shorter-wavelength, higher-amplitude waves caused by strong local winds. As shown in Figure 17.17, the heavy seas of winter gradually erode sand from 475 Physical Geology 476 beaches, moving it to an underwater sandbar offshore from the beach. The gentler waves of summer gradually push this sand back toward the shore, creating a wider and flatter beach. Figure 17.17 The differences between summer and winter on beaches in areas where the winter conditions are rougher and waves have a shorter wavelength but higher energy. In winter, sand from the beach is stored offshore. [SE] The evolution of sandy depositional features on sea coasts is primarily influenced by waves and currents, especially longshore currents. As sediment is transported along a shore, either it is deposited on beaches, or it creates other depositional features. A spit, for example is an elongated sandy deposit that extends out into open water in the direction of a longshore current. A good example is Goose Spit at Comox on Vancouver Island (Figure 17.18). At this location, the longshore current typically flows toward the southwest, and the sand eroded from a 60 m high cliff of Pleistocene glaciofluvial Quadra Sand is pushed in that direction and then out into Comox Harbour. Figure 17.18 The formation of Goose Spit at Comox on Vancouver Island. The sand that makes up Goose Spit is derived from the erosion of Pleistocene Quadra Sand (a thick glaciofluvial sand deposit, as illustrated in the photo on the right). [SE] The Quadra Sand at Comox is visible in Figure 17.19. There are numerous homes built at the top of the cliff, and the property owners have gone to considerable expense to reinforce the base of the cliff with large angular rocks (rip-rap) and concrete barriers so as to limit further erosion of their properties. One result of this will be to starve Goose Spit of sediments and eventually contribute to its erosion. Of course the rocks and concrete barriers are only temporary; they will be eroded by strong winter storms over the next few decades and the Quadra Sand will once again contribute to the maintenance of Goose Spit.

Figure 17.19 The Quadra Sand cliff at Comox, and the extensive concrete and rip-rap barrier that has been constructed to reduce erosion. Note that the waves (dashed lines) are approaching the shore at an angle, contributing to the longshore current. [SE] A spit that extends across a bay to the extent of closing, or almost closing it off, is known as a baymouth bar. Most bays have streams flowing into them, and since this water has to get out, it is rare that a baymouth bar will completely close the entrance to a bay. In areas where there is sufficient sediment being transported, and there are near-shore islands, a tombolo may form (Figure 17.20).

Tombolos are common around the southern part of the coast of British Columbia, where islands are abundant, and they typically form where there is a wave shadow behind a nearshore island (Figure 17.21). This becomes an area with reduced energy, and so the longshore current slows and sediments accumulate. Eventually enough sediments accumulate to connect the island to the mainland with a tombolo. There is a good example of a tombolo in Figure 17.1, and another in Figure 17.22.

In areas where coastal sediments are abundant and coastal relief is low (because there has been little or no recent coastal uplift), it is common for barrier islands to form (Figure 17.23). Barrier islands are elongated islands composed of sand that form a few kilometres away from the mainland. They are common along the U.S. Gulf Coast from Texas to Florida, and along the U.S. Atlantic Coast from Florida to Massachusetts. North of Boston, the coast becomes rocky, partly because that area has been affected by post-glacial crustal rebound.

Figure 17.23 Assateague Island on the Maryland coast, U.S. This barrier island is about 60 km long and only 1 km to 2 km wide. The open Atlantic Ocean is to the right and the lagoon is to the left. This part of Assateague Island has recently been eroded by a tropical storm, which pushed massive amounts of sand into the lagoon. DelmarvaAssateague_aerial_ViewCV.jpg]

Some coasts in tropical regions (between 30° S and 30° N) are characterized by carbonate reefs. Reefs form in relatively shallow marine water within a few hundred to a few thousand metres of shore in areas where there is little or no input of clastic sediments from streams, and marine organisms such as corals, algae, and shelled organisms can thrive. The associated biological processes are enhanced where upwelling currents bring chemical nutrients from deeper water (but not so deep that the water is cooler than about 25°C) (Figure 17.24). Sediments that form in the back reef (shore side) and fore reef (ocean side) are typically dominated by carbonate fragments eroded from the reef and from organisms that thrive in the back-reef area that is protected from wave energy by the reef.

Sea-level change has been a feature on Earth for billions of years, and it has important implications for coastal processes and both erosional and depositional features. There are three main mechanisms of sea- level change, as described below. Eustatic sea-level changes are global sea-level changes related either to changes in the volume of glacial ice on land or to changes in the shape of the sea floor caused by plate tectonic processes. For example, changes in the rate of mid-ocean spreading will change the shape of the sea floor near the ridges, and this affects sea level. Over the past 20,000 years, there has been approximately 125 m of eustatic sea-level rise due to glacial melting. Most of that took place between 15,000 and 7,500 years ago during the major melting phase of the North American and Eurasian Ice Sheets (Figure 17.25). At around 7,500 years ago, the rate of glacial melting and sea-level rise decreased dramatically, and since that time, the average rate has been in the order of 0.7 mm/year. Anthropogenic climate change led to accelerating sea-level rise starting around 1870. Since that time, the average rate has been 1.1 mm/year, but it has been gradually increasing. Since 1992, the average rate has been 3.2 mm/year. Figure 17.25 Eustatic sea-level curve for the past 24 ka (sea-level rise resulting from the melting of glacial ice). Sea-level rise is global; the locations listed in the caption are the places where data were acquired to create this diagram. [https://en.wikipedia.org/wiki/ Sea_level_rise#/media/File:Post-Glacial_Sea_Level.png]

Almost all of Canada and parts of the northern United States were covered in thick ice sheets at the peak of the last glaciation. Following the melting of this ice, there has been an isostatic rebound of continental

Physical Geology 482 crust in many areas. This ranges from several hundred metres of rebound in the central part of the Laurentide Ice Sheet (around Hudson Bay) to 100 m to 200 m in the peripheral parts of the Laurentide and Cordilleran Ice Sheets — in places such as Vancouver Island and the mainland coast of B.C. In other words, although global sea level was about 130 m lower during the last glaciation, the glaciated regions were depressed at least that much in most places, and more than that in places where the ice was thickest.

Figure 17.26 This stream is on the southwest coast of Vancouver Island near Sooke. Like many other streams along this coast, it used to flow directly into the ocean, but the land has been uplifted by post-glacial isostatic rebound. [SE] Tectonic sea-level changes are local changes caused by tectonic processes. The subduction of the Juan de Fuca Plate beneath British Columbia is creating tectonic uplift (about 1 mm/year) along the western edge of Vancouver Island, although much of this uplift is likely to be reversed when the next large subduction-zone earthquake strikes. Coastlines in areas where there has been net sea-level rise in the geologically recent past are commonly characterized by estuaries and fiords. Howe Sound, north of Vancouver, is an example of a fiord (Figure 17.27). This valley was filled with ice during the last glaciation, and there has been a net rise in sea level here since that time. Coastlines in areas where there has been net sea-level drop in the geologically recent past are characterized by uplifted wave-cut platforms (or stream valleys as shown in Figure 17.26). Uplifted beach lines are another product of relative sea-level drop, although these are difficult to recognize in areas with vigorous vegetation. They are relatively common in Canada’s far north.

There are various modifications that we make in an attempt to influence beach processes for our own purposes. Sometimes these changes are effective, and may appear to be beneficial, although in most cases there are unintended negative consequences that we don’t recognize until much later. An example is at the beach near Comox (described above), which has been armoured with rip-rap and concrete blocks in an attempt to limit the natural erosion that is threatening the properties at the top of the cliff (Figure 17.19). As already noted, the unintended effect of this installation will be to starve Goose Spit of sediment. As long as the armour remains in place, which might be several decades, there is a risk that the spit will start to erode, which will affect many of the organisms that use that area as their habitat, and many of the people who go there for recreation. Seawalls, like the one around Vancouver’s Stanley Park (Figure 17.28), also help to limit erosion and can be very pleasant amenities for the public, but they have geological and ecological costs. When a shoreline is “hardened” in this way, important marine habitat is lost and sediment production is reduced, and that can affect beaches elsewhere. Seawalls also affect the behaviour of waves and longshore currents, sometimes with negative results.

Another example is at Sunset Beach in Vancouver. As shown in Figure 17.29, a series of rip-rap breakwaters (structures parallel to the shore) were built in the 1990s and sand has accumulated behind them to form the beach. The breakwaters have acted as islands and the sand has been deposited in the low-energy water behind them, in the same way that a tombolo forms. This can be seen from a photograph taken from the Burrard Bridge in 2015 (Figure 17.30). The two benefits of this project are that a pleasant beach has been created, and some of the sediment that previously would have been moved into False Creek, and could have blocked its entrance, has been trapped in English Bay. The negative impacts are probably not well understood, but have likely involved loss of marine animal habitat.

Groynes (or groins in the U.S.) have an effect that is similar to that of breakwaters, although groynes are constructed perpendicular to the beach (Figure 17.31), and they trap sediment by slowing the longshore current.

Most of the sediment that forms beaches along our coasts comes from rivers, so if we want to take care of beaches, we have to take care of rivers. When a river is dammed, its sediment load is deposited in the resulting reservoir, and for the century or two while the reservoir is filling up, that sediment cannot get to the sea. During that time, beaches (including spits, baymouth bars, and tombolos) within tens of kilometres of the river’s mouth (or more in some cases) are at risk of erosion.

Referring to Table 17.1, approximately what size of waves (amplitude and wavelength) would you expect with a 65 km/h wind blowing for 40 hours over 1,000 km of sea? If the average wavelength of a series of waves is 100 m, at what depth of water will the waves start to feel the bottom, and how will that change their behaviour? What is the difference between a longshore current and longshore drift? On this diagram, the waves (dashed blue lines) are approaching an irregular coast. The red arrows represent the energy of those waves, and one has been extended to show where that energy would hit the shore. Extend the other “energy lines” in a similar way, and comment on how this relates to erosion of this coastline. Explain the origins of a wave-cut platform. How do we define the limits of the beach face, and what are some other terms used to describe this zone? A spit is really just a beach that is only attached to the shore at one end. What conditions are necessary for the formation of a spit? Barrier islands are common along the Atlantic coast of the U.S. as far north as Massachusetts. Why are there almost none in the northeastern U.S. or along the coasts of New Brunswick, Nova Scotia, and Newfoundland? This diagram represents an island on the central coast of B.C. that has experienced 140 m of isostatic rebound since deglaciation, and has also been affected by the global eustatic sea-level rise over the same period. The dashed line marks sea level during glaciation. How much higher or lower should that line be now?

Describe the origins of the major topographic features of the sea floor, including continental shelves and slopes, spreading ridges, seamount chains and isolated seamounts, and deep submarine canyons Describe the various components of oceanic crust: pillow basalts, sheeted dykes, gabbro bodies, layered gabbro, and layered ultramafic rock Describe the age distribution of oceanic crust, and explain why all of it is relatively young Summarize the types of sediments and sedimentary rocks that accumulate on the sea floor, and explain why different types of sediment are present in different areas Explain the origins of sea-floor methane hydrates Describe and explain regional variations in the salinity and temperature of ocean water Describe the general nature of major ocean-surface currents and the origins of deep-ocean circulation patterns Explain the importance of ocean currents to our climate

Oceans cover 71% of Earth’s surface and hold 97% of Earth’s water. The water they contain is critical to plate tectonics, to volcanism, and of course, to life on Earth. It is said that we know more about the surface of the Moon than the floor of the oceans. Whether this is true or not, the important point is that the ocean floor is covered with an average of nearly 4,000 m of water, and it’s pitch black below a few hundred metres so it’s not easy to discover what is down there. We know a lot more about the oceans than we used to, but there is still a great deal more to discover. Earth has had oceans for a very long time, dating back to the point where the surface had cooled enough to allow liquid water, only a few hundred million years after Earth’s formation. At that time there were no continental rocks, so the water that was here was likely spread out over the surface in one giant (but relatively shallow) ocean.

We examined the topography of the sea floor from the perspective of plate tectonics in Chapter 10, but here we are going to take another look at the important features from an oceanographic perspective. The topography of the northern Atlantic Ocean is shown in Figure 18.2. The important features are the extensive continental shelves less than 250 m deep (pink); the vast deep ocean plains between 4,000 and 6,000 m deep (light and dark blue); the mid-Atlantic ridge, in many areas shallower than 3,000 m; and the deep ocean trench north of Puerto Rico (8,600 m). Figure 18.2 The topography of the Atlantic Ocean sea floor between 0° and 50° north. Red and yellow colours indicate less than 2,000 m depth; green less than 3,000 m; blue 4,000 m to 5,000 m; and purple greater than 6,000 m. [from NASA/CNES at: marine_topo/jpg_images/topo8.jpg] A topographic profile of the Pacific Ocean floor between Japan and British Columbia is shown in Figure 18.3. Be careful when interpreting this diagram (and others like it), because in order to show the various features clearly the vertical axis is exaggerated, in this case by about 200 times. The floor of the Pacific, like those of the other oceans, is actually very flat, even in areas with seamounts or deep trenches. The vast sediment-covered abyssal plains of the oceans are much flatter than any similar-sized areas on the continents. The main features of the Pacific Ocean floor are the continental slopes, which drop from about 200 m to several thousand metres over a distance of a few hundred kilometres; the abyssal plains — exceedingly flat and from 4,000 m to 6,000 m deep; volcanic seamounts and islands; and trenches at subduction zones that are up to 11,000 m deep.

The ocean floor is almost entirely underlain by mafic oceanic crust (mostly basalt and gabbro, as described in more detail below), while the continental slopes are underlain by felsic continental crust (mostly granitic and sedimentary rocks). And, as you’ll remember from Chapter 10, the heavier oceanic crust floats lower on the mantle than continental crust does, and that’s why oceans are oceans. The continental shelf and slope offshore from Nova Scotia is shown in Figure 18.4. In this passive- margin area (no subduction zone), the shelf is over 150 km wide. On the Pacific coast of Canada, the shelf is less than half as wide. Continental shelves are typically less than 200 m in depth; 200 m is also the limit of the photic zone, the maximum depth to which sufficient light penetrates to allow photosynthesis to take place. As a result of that photosynthesis, the photic zone is oxygenated, and therefore suitable for animal life. Approximately 90% of marine life is restricted to the photic zone. The photic zone is also known as the epipelagic zone. The mesopelagic zone extends from 200 m to 1,000 m; the bathypelagic zone from 1,000 m to 4,000 m; and abyssalpelagic zone is deeper than 4,000 m. (Pelagic refers to the open ocean, and thus excludes areas that are near to the shores or the ocean floor.) Although the temperature of the ocean surface varies widely, from a few degrees either side of freezing in polar regions to over 25°C in the tropics, in most parts of the ocean, the water temperature is around 10°C at 1,000 m depth and about 4°C from 2,000 m depth all the way to the bottom. Figure 18.4 The generalized topography of the Atlantic Ocean floor within 300 km of Nova Scotia. The vertical exaggeration is approximately 25 times. The panel at the bottom shows the same profile without vertical exaggeration. [SE] 493 Chapter 18 Geology of the Oceans The deepest parts of the ocean are within the subduction trenches, and the deepest of these is the Marianas Trench in the southwestern Pacific (near Guam) at 11,000 m (Figure 18.5). There are other trenches in the southwestern Pacific that are over 10,000 m deep; the Japan Trench is over 9,000 m deep; and the Puerto Rico and Chile-Peru Trenches are over 8,000 m deep. Trenches that are relatively shallow tend to be that way because they have significant sediment infill. There is no recognizable trench along the subduction zone of the Juan de Fuca Plate because it has been filled with sediments from the Fraser and Columbia Rivers (or their ancient equivalents). Figure 18.5 The generalized topography of the Pacific Ocean floor in the area of the Marianas Trench, near Guam. The dashed grey line represents the subduction of the Pacific Plate (to the right) beneath the Philippine Plate (to the left). [SE]

As we discussed in Chapter 10, oceanic crust is formed at sea-floor spreading ridges from magma generated by decompression melting of hot upward-moving mantle rock (Figure 10.18). About 10% of the mantle rock melts under these conditions, producing mafic magma. This magma oozes out onto the sea floor to form pillow basalts (Figure 18.1), breccias (fragmented basaltic rock), and flows, interbedded in some cases with limestone or chert. Beneath the volcanic rock are layers with gabbroic sheeted dykes (which sometimes extend up into the pillow layer), gabbroic stocks, and finally layered peridotite (ultramafic rock) at the base. The ultramafic rock of the mantle lies below that. Over time, the igneous rock of the oceanic crust gets covered with layers of sediment, which eventually become sedimentary rock, including limestone, mudstone, chert, and turbidites. The lithologies of the layers of the oceanic crust are shown in Figure 18.6.

The age of the oceanic crust has been determined by systematic mapping variations in the strength of the Earth’s magnetic field across the sea floor and comparing the results with our understanding of the record of Earth’s magnetic field reversal chronology for the past few hundred million years. The ages of different parts of the crust are shown in Figure 18.7. The oldest oceanic crust is around 280 Ma in the eastern Mediterranean, and the oldest parts of the open ocean are around 180 Ma on either side of the north Atlantic. It may be surprising, considering that parts of the continental crust are close to 4,000 Ma old, that the oldest sea floor is less than 300 Ma. Of course, the reason for this is that all sea floor older

Physical Geology 496 than that has been either subducted or pushed up to become part of the continental crust. For example, there are fragments of sea floor in British Columbia that date back to around 380 and 220 Ma, and there are similar rocks in the Canadian Shield that are older than 3 Ga. As one would expect, the oceanic crust is very young near the spreading ridges (Figure 18.7), and there are obvious differences in the rate of sea-floor spreading along different ridges. The ridges in the Pacific and southeastern Indian Oceans have wide age bands, indicating rapid spreading (approaching 10 cm/y on each side in some areas), while those in the Atlantic and western Indian Oceans are spreading much more slowly (less than 2 cm/y on each side in some areas).

As is evident from Figures 18.2 and 18.3, the sea floor is dotted with chains of seamounts, isolated seamounts, and ocean islands. Almost all of these features are volcanoes, and most are much younger than the oceanic crust on which they formed. Some seamounts and ocean islands are formed above mantle plumes, the best example being Hawaii. The oldest of the Hawaiian/Emperor seamounts is dated at around 80 Ma; it is situated on oceanic crust aged around 90 to 100 Ma. The youngest of the Hawaiian lavas — at Kilauea Volcano on the island of Hawaii — is just a few hours old (or less!) and the island is surrounded by oceanic crust that is around 85 Ma old. All of the mantle-plume-derived volcanic islands are dominated by mafic rocks. Many seamounts are related to subduction along ocean-ocean convergent boundaries. These include the Aleutians, extending from Alaska to Russia, and the Lesser Antilles in the eastern part of the Caribbean. Some of the linear belts of volcanoes in the Pacific Ocean do not show age-distance relationships like the volcanoes of the Hawaii-Emperor chain or the Galapagos Islands. For example, the Line Islands, which spread out over more than 1,000 km south of the Hawaiian chain, were all formed between 70 and 85 Ma and are interpreted to be related to rifting. Most tropical islands have associated carbonate reefs, in some cases, as fringes right around the island, and in some cases, as barriers some distance away. In many cases, the reef is there, but the island that is assumed to have led to its formation is gone. The formation of fringing reefs, barrier reefs, and atolls is illustrated in Figure 18.8.

The key factor in this process is sea-level change, either because of post-glacial sea-level rise, or because of subsidence of a volcano — as it is moved away from a spreading ridge — or both. If the rate of sea- level change is slow enough (e.g., less than 1 cm/year), a reef can keep up and maintain its position at sea level long after its parent volcanic island has disappeared beneath the waves.

Except within a few kilometres of a ridge crest, where the volcanic rock is still relatively young, most parts of the sea floor are covered in sediments. This material comes from several different sources and is highly variable in composition, depending on proximity to a continent, water depth, ocean currents, biological activity, and climate. Sea-floor sediments (and sedimentary rocks) can range in thickness from a few millimetres to several tens of kilometres. Near the surface, the sea-floor sediments remain unconsolidated, but at depths of hundreds to thousands of metres (depending on the type of sediment and other factors) the sediment becomes lithified.

Terrigenous sediment is derived from continental sources transported by rivers, wind, ocean currents, and glaciers. It is dominated by quartz, feldspar, clay minerals, iron oxides, and terrestrial organic matter. Pelagic carbonate sediment is derived from organisms (e.g., foraminifera) living in the ocean water (at various depths, but mostly near surface) that make their shells (a.k.a. tests) out of carbonate minerals such as calcite. Pelagic silica sediment is derived from marine organisms (e.g., diatoms and radiolaria) that make their tests out of silica (microcrystalline quartz). Volcanic ash and other volcanic materials are derived from both terrestrial and submarine eruptions. Iron and manganese nodules form as direct precipitates from ocean-bottom water. The distributions of some of these materials around the seas are shown in Figure 18.9. Terrigenous sediments predominate near the continents and within inland seas and large lakes. These sediments tend to be relatively coarse, typically containing sand and silt, but in some cases even pebbles and cobbles. Clay settles slowly in nearshore environments, but much of the clay is dispersed far from its source areas by ocean currents. Clay minerals are predominant over wide areas in the deepest parts of the ocean, and most of this clay is terrestrial in origin. Siliceous oozes (derived from radiolaria and diatoms) are common in the south polar region, along the equator in the Pacific, south of the Aleutian Islands, and within large parts of the Indian Ocean. Carbonate oozes are widely distributed in all of the oceans within equatorial and mid-latitude regions. In fact, clay settles everywhere in the oceans, but in areas where silica- and carbonate-producing organisms are prolific, they produce enough silica or carbonate sediment to dominate over clay.

Carbonate sediments are derived from a wide range of near-surface pelagic organisms that make their shells out of carbonate (Figure 18.10). These tiny shells, and the even tinier fragments that form when they break into pieces, settle slowly through the water column, but they don’t necessarily make it to the bottom. While calcite is insoluble in surface water, its solubility increases with depth (and pressure) and at around 4,000 m, the carbonate fragments dissolve. This depth, which varies with latitude and water temperature, is known as the carbonate compensation depth, or CCD. As a result, carbonate oozes are absent from the deepest parts of the ocean (deeper than 4,000 m), but they are common in shallower areas such as the mid-Atlantic ridge, the East Pacific Rise (west of South America), along the trend of the Hawaiian/Emperor Seamounts (in the northern Pacific), and on the tops of many isolated seamounts.

All terrestrial erosion products include a small proportion of organic matter derived mostly from terrestrial plants. Tiny fragments of this material plus other organic matter from marine plants and animals accumulate in terrigenous sediments, especially within a few hundred kilometres of shore. As the sediments pile up, the deeper parts start to warm up (from geothermal heat), and bacteria get to work breaking down the contained organic matter. Because this is happening in the absence of oxygen (a.k.a. anaerobic conditions), the by-product of this metabolism is the gas methane (CH4). Methane released by the bacteria slowly bubbles upward through the sediment toward the sea floor. At water depths of 500 m to 1,000 m, and at the low temperatures typical of the sea floor (close to 4°C), water and methane combine to create a substance known as methane hydrate. Within a few metres to hundreds of metres of the sea floor, the temperature is low enough for methane hydrate to be stable and hydrates accumulate within the sediment (Figure 18.11). Methane hydrate is flammable because when it is heated, the methane is released as a gas (Figure 18.11). The methane within sea- floor sediments represents an enormous reservoir of fossil fuel energy. Although energy corporations and governments are anxious to develop ways to produce and sell this methane, anyone that understands the climate-change implications of its extraction and use can see that this would be folly. As we’ll see in the discussion of climate change in Chapter 19, sea-floor methane hydrates have had significant impacts on the climate in the distant past. Figure 18.11 Left: Methane hydrate within muddy sea-floor sediment from an area offshore from Oregon. [https://upload.wikimedia.org/wikipedia/ commons/4/49/Gashydrat_im_Sediment.JPG] Right: Methane hydrate on fire features/files/2012/01/New-Image.jpg]

As everyone knows, seawater is salty. It is that way because the river water that flows into the oceans contains small amounts of dissolved ions, and for the most part, the water that comes out of the oceans is the pure water that evaporates from the surface. The salts of the ocean (dominated by sodium, chlorine, and sulphur) (Figure 18.12) are there because they are very soluble and they aren’t consumed by biological processes (most of the calcium, for example, is used by organisms to make carbonate minerals). If salts are always going into the ocean, and never coming out, one might assume that the oceans have been continuously getting saltier over geological time. In fact this appears not to be the case. There is geological evidence that Earth’s oceans became salty early during the Archaean, and that at times in the past, they have been at least half again as salty as they are now. This implies that there must be a mechanism to remove salt from the oceans, and that mechanism is the isolation of some parts of the ocean into seas (such as the Mediterranean) and the eventual evaporation of those seas to create salt beds that become part of the crust. The Middle Devonian Prairie Evaporite Formation of Saskatchewan and Manitoba is a good example of this.

The average salinity of the oceans is 35 g of salt per litre of water, but there are significant regional variations in this value, as shown in Figure 18.13. Ocean water is least salty (around 31 g/L) in the Arctic, and also in several places where large rivers flow in (e.g., the Ganges/Brahmaputra and Mekong Rivers in southeast Asia, and the Yellow and Yangtze Rivers in China). Ocean water is most salty (over 37 g/L) in some restricted seas in hot dry regions, such as the Mediterranean and Red Seas. You might be surprised to know that, in spite of some massive rivers flowing into it (such as the Nile and the Danube), water does not flow out of the Mediterranean Sea into the Atlantic. There is so much evaporation happening in the Mediterranean basin that water flows into it from the Atlantic, through the Strait of Gibraltar. In the open ocean, salinities are elevated at lower latitudes because this is where most evaporation takes place. The highest salinities are in the subtropical parts of the Atlantic, especially north of the equator. The northern Atlantic is much more saline than the north Pacific because the Gulf Stream current brings a massive amount of salty water from the tropical Atlantic and the Caribbean to the region around Britain, Iceland, and Scandinavia. The salinity in the Norwegian Sea (between Norway and Iceland) is substantially higher than that in other polar areas.

How salty is the sea? If you’ve ever swum in the ocean, you’ve probably tasted it. To understand how salty the sea is, start with 250 mL of water (1 cup). There is 35 g of salt in 1 L of seawater so in 250 mL (1/4 litre) there is 35/4 = 8.75 or ~9 g of salt. This is just short of 2 teaspoons, so it would be close enough to add 2 level teaspoons of salt to the cup of water. Then stir until it’s dissolved. Have a taste! Of course, if you used normal refined table salt, then what you added was almost pure NaCl. To get the real taste of seawater you would want to use some evaporated seawater salt (a.k.a. sea salt), which has a few percent of magnesium, sulphur, and calcium plus some trace elements. Not unexpectedly, the oceans are warmest near the equator — typically 25° to 30°C — and coldest near the poles — around 0°C (Figure 18.14). (Sea water will remain unfrozen down to about -2°C.) At southern Canadian latitudes, average annual water temperatures are in the 10° to 15°C range on the west coast and in the 5° to 10°C range on the east coast. Variations in sea-surface temperatures (SST) are related to redistribution of water by ocean currents, as we’ll see below. A good example of that is the plume of warm Gulf Stream water that extends across the northern Atlantic. St. John’s, Newfoundland, and Brittany in France are at about the same latitude (47.5° N), but the average SST in St. John’s is a frigid 3°C, while that in Brittany is a reasonably comfortable 15°C.

Currents in the open ocean are created by wind moving across the water and by density differences related to temperature and salinity. An overview of the main ocean currents is shown in Figure 18.15. As you can see, the northern hemisphere currents form circular patterns (gyres) that rotate clockwise, while the southern hemisphere gyres are counter-clockwise. This happens for the same reason that the water in your northern hemisphere sink rotates in a clockwise direction as it flows down the drain; this is caused by the Coriolis effect.

Because the ocean basins aren’t like bathroom basins, not all ocean currents behave the way we would expect. In the North Pacific, for example, the main current flows clockwise, but there is a secondary current in the area adjacent to our coast — the Alaska Current — that flows counter-clockwise, bringing relatively warm water from California, past Oregon, Washington, and B.C. to Alaska. On Canada’s eastern coast, the cold Labrador Current flows south past Newfoundland, bringing a stream of icebergs past the harbour at St. John’s (Figure 18.16). This current helps to deflect the Gulf Stream toward the northeast, ensuring that Newfoundland stays cool, and western Europe stays warm. Figure 18.15 Overview of the main open-ocean currents. Red arrows represent warm water moving toward colder regions. Blue arrows represent cold water moving toward warmer regions. Black arrows represent currents that don’t involve significant temperature changes. [From: https://upload.wikimedia.org/wikipedia/commons/9/9b/ Corrientes-oceanicas.png]

The currents shown in Figure 18.15 are all surface currents, and they only involve the upper few hundred metres of the oceans. But there is much more going on underneath. The Gulf Stream, for example, which is warm and saline, flows past Britain and Iceland into the Norwegian Sea (where it becomes the Norwegian Current). As it cools down, it becomes denser, and because of its high salinity, which also contributes to its density, it starts to sink beneath the surrounding water (Figure 18.17). At this point, it is known as North Atlantic Deep Water (NADW), and it flows to significant depth in the Atlantic as it heads back south. Meanwhile, at the southern extreme of the Atlantic, very cold water adjacent to Antarctica also sinks to the bottom to become Antarctic Bottom Water (AABW) which flows to the north, underneath the NADW.

The descent of the dense NADW is just one part of a global system of seawater circulation, both at surface and at depth, as illustrated in Figure 18.18. The water that sinks in the areas of deep water formation in the Norwegian Sea and adjacent to Antarctica moves very slowly at depth. It eventually resurfaces in the Indian Ocean between Africa and India, and in the Pacific Ocean, north of the equator.

Figure 18.18 The thermohaline circulation system, also known as the Global Ocean Conveyor [from NASA at: https://en.wikipedia.org/wiki/ Thermohaline_circulation#/media/ File:Thermohaline_Circulation_2.png] The thermohaline circulation is critically important to the transfer of heat on Earth. It brings warm water from the tropics to the poles, and cold water from the poles to the tropics, thus keeping polar regions from getting too cold and tropical regions from getting too hot. A reduction in the rate of thermohaline circulation would lead to colder conditions and enhanced formation of sea ice at the poles. This would start a positive feedback process that could result in significant global cooling. There is compelling evidence to indicate that there were major changes in thermohaline circulation, corresponding with climate changes, during the Pleistocene Glaciation.

Explain why relatively coarse terrigenous sediments (e.g., sand) tend to accumulate close to the continents, while terrigenous clay is dispersed all across the ocean floor. Although clay is widely dispersed in the oceans, in some areas, deep-sea sediments are dominated by clay, while in others they are dominated by carbonate or silica ooze. Why do these differences exist? Explain why carbonate sediments are absent from the deepest parts of the oceans. What is the source of the carbon that is present in sea-floor methane hydrate deposits? Where are the saltiest parts of the oceans? Why? Explain why sea-surface water with the greatest density is found in the north Atlantic, as shown on this map. [SE after: https://upload.wikimedia.org/wikipedia/en/3/31/SeaSurfaceDensity.jpg] What type of ocean currents result from the relatively dense water in the north Atlantic? How do the open-ocean currents affect the overall climate patterns on Earth?

Summarize the properties of greenhouse gases and their role in controlling the climate Explain the difference between climate forcing and climate feedbacks Describe the mechanisms of climate forcing related to solar evolution, continental drift, continental collisions, volcanism, Earth and Sun orbital variations, and changing ocean currents Describe the significance of albedo to climate and how the melting of ice or snow and forestry affect albedo Explain the roles of the melting of permafrost, breakdown of methane hydrates, and temperature- related solubility of CO2 as positive feedbacks Describe some of the ways that our extraction and use of fossil fuels contribute to climate change Explain how food production contributes to climate change List some of the steps that we can take as individuals to limit our personal contribution to climate change Describe the role of climate change in sea-level rise, and why we are already committed to more than a metre of additional sea-level rise Explain the link between climate change and the distribution of diseases and pests

Figure 19.1 Core from Ocean Drilling Program hole 1220b (southeast of Hawaii) showing the boundary between the Paleocene and the Eocene (at 55.8 Ma). Marine life was decimated during the 100,000 years of the Paleocene-Eocene thermal maximum, and the dark part of the core represents the absence of carbonate sediment from planktonic organisms. The scale is in centimetres. [SE, after Ocean Drilling Program, used with permission] If one thing has been constant about Earth’s climate over geological time, it is its constant change. In the geological record, we can see this in the evidence of glaciations in the distant past (see section 16.1 in Chapter 16), and we can also detect periods of extreme warmth by looking at the isotope composition of sea-floor sediments, such as those in the core shown in Figure 19.1. Not only has the climate changed frequently, the temperature fluctuations have been very significant. Today’s mean global temperature is about 15°C. During Snowball Earth times, the global mean was as cold as -50°C, while at various times during the Paleozoic and Mesozoic and during the Paleocene-Eocene thermal maximum, it was close to +30°C. But in spite of these dramatic climate changes, Earth has been habitable from very early in its history — as soon as liquid water was present — right through to the present day. That continuous habitability is perhaps a little more surprising than you might think, as we’ll see below. A significant part of this chapter is about the natural processes of climate change and how they work. It’s critically important to be aware of those natural climate change processes if we want to understand anthropogenic climate change. First, this awareness helps us to understand why our activities are causing the present-day climate to change, and second, it allows us to distinguish between natural and anthropogenic processes in the climate record of the past 250 years.

There are two parts to climate change, the first one is known as climate forcing, which is when conditions change to give the climate a little nudge in one direction or the other. The second part of climate change, and the one that typically does most of the work, is what we call a feedback. When a climate forcing changes the climate a little, a whole series of environmental changes take place, many of which either exaggerate the initial change (positive feedbacks), or suppress the change (negative feedbacks). An example of a climate-forcing mechanism is the increase in the amount of carbon dioxide (CO2) in the atmosphere that results from our use of fossil fuels. CO2 traps heat in the atmosphere and leads to climate warming. Warming changes vegetation patterns; contributes to the melting of snow, ice, and permafrost; causes sea level to rise; reduces the solubility of CO2 in sea water; and has a number of other minor effects. Most of these changes contribute to more warming. Melting of permafrost, for example, is a strong positive feedback because frozen soil contains trapped organic matter that is converted to CO2 and methane (CH4) when the soil thaws. Both these gases accumulate in the atmosphere and add to the warming effect. On the other hand, if warming causes more vegetation growth, that vegetation should absorb CO2, thus reducing the warming effect, which would be a negative feedback. Under our current conditions — a planet that still has lots of glacial ice and permafrost — most of the feedbacks that result from a warming climate are positive feedbacks and so the climate changes that we cause get naturally amplified by natural processes.

Throughout this chapter we’ll be talking about the role of greenhouse gases (GHGs) in controlling the climate, so it’s important to understand what greenhouse gases are and how they work. As you know, the dominant gases of the atmosphere are nitrogen (as N2) and oxygen (as O2). These gas molecules have only two atoms each and are not GHGs. Some of the other important gases of the atmosphere are water Physical Geology 512 vapour (H2O), carbon dioxide (CO2), and methane (CH4). All of these have more than two atoms, and they are GHGs. All molecules vibrate at various frequencies and in various ways, and some of those vibrations take place at frequencies within the range of the infrared (IR) radiation that is emitted by Earth’s surface. Gases with two atoms, such as O2, can only vibrate by stretching (back and forth), and those vibrations are much faster than the IR radiation. Gases with three or more atoms (such as CO2) vibrate by stretching as well, but they can also vibrate in other ways, such as by bending. Those vibrations are slower and match IR radiation frequencies. When IR radiation interacts with CO2 or with one of the other GHGs, the molecular vibrations are enhanced because there is a match between the wavelength of the light and the vibrational frequency of the molecule. This makes the molecule vibrate more vigorously, heating the surrounding air in the process. These molecules also emit IR radiation in all directions, some of which reaches Earth’s surface and causes the greenhouse effect.

The longest-term natural forcing variation is related to the evolution of the Sun. Like most other stars of a similar mass, our Sun is evolving. For the past 4.57 billion years, its rate of nuclear fusion has been increasing, and it is now emitting about 40% more energy (as light) than it did at the beginning of geological time (Figure 19.2). A difference of 40% is big, so it’s a little surprising that the temperature on Earth has remained at a reasonable and habitable temperature for all of this time. The mechanism for that relative climate stability has been the evolution of our atmosphere from one that was dominated by CO2, and also had significant levels of CH4 — both GHGs — to one with only a few hundred parts per million of CO2 and just under 1 part per million of CH4. Those changes to our atmosphere have been no accident; over geological time, life and its metabolic processes have evolved and changed the

The Gaia hypothesis, developed by British scientist and environmentalist James Lovelock in the 1960s, is the theory that organisms evolve in ways that contribute to ensuring that their environment remains habitable. It does not include any sort of coordination of effort among organisms or any consciousness of a need to make changes. Gaia is not a superorganism. A way of understanding Gaia is through Lovelock’s simple Daisyworld model. A planet with a warming star is populated only by two types of daisies, white ones and black ones. The black ones contribute to warming because they absorb solar energy, while the white ones reflect light and contribute to cooling. As the star’s luminosity gradually increases, the white daisies have better outcomes because their reflectivity cools their local environment, while the black daisies, suffering from the heat, do not reproduce as well. Over time white daisies gradually dominate the population, but eventually the star becomes so bright that even white daisies cannot compensate, and all of the daisies perish. Obviously Earth is not Daisyworld, but similar processes — such as the evolution of photosynthetic bacteria that consume CO2 — have taken place that influence the atmosphere and moderate the climate.

Plate tectonic processes contribute to climate forcing in several different ways, and on time scales ranging from tens of millions to hundreds of millions of years. One mechanism is related to continental position. For example, we know that Gondwana (South America + Africa + Antarctica + Australia) was positioned over the South Pole between about 450 and 250 Ma, during which time there were two major glaciations (Andean-Saharan and Karoo) affecting the South polar regions (Figure 16.2) and cooling the rest of the planet at the same time. Another mechanism is related to continental collisions. As described in Chapter 16, the collision between India and Asia, which started at around 50 Ma, resulted in massive tectonic uplift. The consequent accelerated weathering of this rugged terrain consumed CO2 from the atmosphere and contributed to gradual cooling over the remainder of the Cenozoic. Also, as described in Chapter 16, the opening of the Drake Passage — due to plate-tectonic separation of South America from Antarctica — led to the development of the Antarctic Circumpolar Current, which isolated Antarctica from the warmer water in the rest of the ocean and thus contributed to Antarctic glaciation starting at around 35 Ma. As we discussed in Chapter 4, volcanic eruptions don’t just involve lava flows and exploding rock fragments; various particulates and gases are also released, the important ones being sulphur dioxide and CO2. Sulphur dioxide is an aerosol that reflects incoming solar radiation and has a net cooling effect that is short lived (a few years in most cases, as the particulates settle out of the atmosphere within a couple of years), and doesn’t typically contribute to longer-term climate change. Volcanic CO2 emissions can contribute to climate warming but only if a greater-than-average level of volcanism is sustained over a long time (at least tens of thousands of years). It is widely believed that the catastrophic end-Permian

Earth’s orbit around the Sun is nearly circular, but like all physical systems, it has natural oscillations. First, the shape of the orbit changes on a regular time scale — close to 100,000 years — from being close to circular to being very slightly elliptical. But the circularity of the orbit is not what matters; it is the fact that as the orbit becomes more elliptical, the position of the Sun within that ellipse becomes less central or more eccentric (Figure 19.3a). Eccentricity is important because when it is high, the Earth- Sun distance varies more from season to season than it does when eccentricity is low.

Figure 19.3 The cycles of Earth’s orbit and rotation [a: SE after https://upload.wikimedia.org/wikipedia/ commons/thumb/d/da/Eccentricity_zero.svg/ 1163px-Eccentricity_zero.svg.png], b: https://upload.wikimedia.org/wikipedia/commons/ thumb/a/ae/Earth_obliquity_range.svg/ 2000px-Earth_obliquity_range.svg.png] c: https://upload.wikimedia.org/wikipedia/commons/ thumb/4/43/ Earth_precession.svg/ 2000px-Earth_precession.svg.png] Second, Earth rotates around an axis through the North and South Poles, and that axis is at an angle to the plane of Earth’s orbit around the Sun (Figure 19.3b). The angle of tilt (also known as obliquity) varies on a time scale of 41,000 years. When the angle is at its maximum (24.5°), Earth’s seasonal differences are accentuated. When the angle is at its minimum (22.1°), seasonal differences are minimized. The current hypothesis is that glaciation is favoured at low seasonal differences as summers would be cooler and snow would be less likely to melt and more likely to accumulate from year to year. Third, the direction in which Earth’s rotational axis points also varies, on a time scale of about 20,000 years (Figure 19.3c). This variation, known as precession, means that although the North Pole is presently pointing to the star Polaris (the pole star), in 10,000 years it will point to the star Vega. The importance of eccentricity, tilt, and precession to Earth’s climate cycles (now known as Milankovitch Cycles) was first pointed out by Yugoslavian engineer and mathematician Milutin Milankovitch in the early 1900s. Milankovitch recognized that although the variations in the orbital cycles did not affect the total amount of insolation (light energy from the Sun) that Earth received, it did affect where on Earth that energy was strongest. Glaciations are most sensitive to the insolation received at latitudes of around 65°, and with the current configuration of continents, it would have to be 65° north (because there is almost no land at 65° south). The most important issues are whether the northern hemisphere is pointing toward the Sun at its closest or farthest approach, and how eccentric the Sun’s position is in Earth’s orbit. Two opposing situations are illustrated in Figure 19.4. In the upper panel, the northern hemisphere is at it farthest distance from the Sun during summer, which means cooler summers. In the lower panel, the northern hemisphere is at its closest distance to the Sun during summer, which means hotter summers. Cool summers — as opposed to cold winters — are the key factor in the accumulation of glacial ice, so the upper scenario in Figure 19.4 is the one that promotes glaciation. This factor is greatest when eccentricity is high.

Figure 19.4 The effect of precession on insolation in the northern hemisphere summers. In (a) the northern hemisphere summer takes place at greatest Earth-Sun distance, so summers are cooler. In (b) (10,000 years or one-half precession cycle later) the opposite is the case, so summers are hotter. The red dashed line represents Earth’s path around the Sun. Data for tilt, eccentricity, and precession over the past 400,000 years have been used to determine the insolation levels at 65° north, as shown in Figure 19.5. Also shown in Figure 19.5 are Antarctic ice-core temperatures from the same time period. The correlation between the two is clear, and it shows up in the Antarctic record because when insolation changes lead to growth of glaciers in the northern hemisphere, southern-hemisphere temperatures are also affected. Figure 19.5 Insolation at 65° N in July compared with Antarctic ice-core temperatures [By SE, using data from Valerie Masson-Delmotte, EPICA Dome C ice core 800KYr deuterium data and temperature estimates WDCA Contribution Series Number : 2007 -091 NOAA/NCDC Paleoclimatology Program, Boulder CO, USA. Retrieved from: ftp://ftp.ncdc.noaa.gov/pub/data/paleo/icecore/ antarctica/epica_domec/edc3deuttemp2007.txt and from Berger, A. and Loutre, M.F. (1991). Insolation values for the climate of the last 10 million years. Quaternary Science Reviews, 10, 297-317.] Ocean currents are important to climate, and currents also have a tendency to oscillate. Glacial ice cores show clear evidence of changes in the Gulf Stream (and other parts of the thermohaline circulation system) that affected global climate on a time scale of about 1,500 years during the last glaciation. The Physical Geology 518 east-west changes in sea-surface temperature and surface pressure in the equatorial Pacific Ocean — known as the El Niño Southern Oscillation or ENSO — varies on a much shorter time scale of between two and seven years. These variations tend to garner the attention of the public because they have significant climate implications in many parts of the world. The past 65 years of ENSO index values are shown in Figure 19.6. The strongest El Niños in recent decades were in 1983 and 1998, and those were both very warm years from a global perspective. During a strong El Niño, the equatorial Pacific sea-surface temperatures are warmer than normal and heat the atmosphere above the ocean, which leads to warmer-than-average global temperatures.

As already stated, climate feedbacks are critically important in amplifying weak climate forcings into full-blown climate changes. When Milankovitch published his theory in 1924, it was widely ignored, partly because it was evident to climate scientists that the forcing produced by the orbital variations was not strong enough to drive the significant climate changes of the glacial cycles. Those scientists did not recognize the power of positive feedbacks. It wasn’t until 1973, 15 years after Milankovitch’s death, that sufficiently high-resolution data were available to show that the Pleistocene glaciations were indeed driven by the orbital cycles, and it became evident that the orbital cycles were just the forcing that initiated a range of feedback mechanisms that made the climate change. Since Earth still has a very large volume of ice — mostly in the continental ice sheets of Antarctica and Greenland, but also in alpine glaciers and permafrost — melting is one of the key feedback mechanisms. Melting of ice and snow leads to several different types of feedbacks, an important one being a change in albedo. Albedo is a measure of the reflectivity of a surface. Earth’s various surfaces have widely differing albedos, expressed as the percentage of light that reflects off a given material. This is important because most solar energy that hits a very reflective surface is not absorbed and therefore does little to warm Earth. Water in the oceans or on a lake is one of the darkest surfaces, reflecting less than 10% of the incident light, while clouds and snow or ice are among the brightest surfaces, reflecting 70% to 90% of the incident light (Figure 19.7).

When sea ice melts, as it has done in the Arctic Ocean at a disturbing rate over the past decade, the albedo of the area affected changes dramatically, from around 80% down to less than 10%. Much more solar energy is absorbed by the water than by the pre-existing ice, and the temperature increase is amplified. The same applies to ice and snow on land, but the difference in albedo is not as great. When ice and snow on land melt, sea level rises. (Sea level is also rising because the oceans are warming and that increases their volume.) A higher sea level means a larger proportion of the planet is covered with water, and since water has a lower albedo than land, more heat is absorbed and the temperature goes up a little more. Since the last glaciation, sea-level rise has been about 125 m; a huge area that used to be land is now flooded by heat-absorbent seawater. During the current period of anthropogenic climate change, sea level has risen only about 20 cm, and although that doesn’t make a big change to albedo, sea-level rise is accelerating. Physical Geology 520 Most of northern Canada has a layer of permafrost that ranges from a few centimetres to hundreds of metres in thickness; the same applies in Alaska, Russia, and Scandinavia. Permafrost is a mixture of soil and ice (Figure 19.8), and it also contains a significant amount of trapped organic carbon that is released as CO2 and CH4 when the permafrost breaks down. Because the amount of carbon stored in permafrost is in the same order of magnitude as the amount released by burning fossil fuels, this is a feedback mechanism that has the potential to equal or surpass the forcing that has unleashed it.

In some polar regions, including northern Canada, permafrost includes methane hydrate (see section 18.3), a highly concentrated form of CH4 trapped in solid form. Breakdown of permafrost releases this CH4. Even larger reserves of methane hydrate exist on the sea floor, and while it would take significant warming of ocean water down to a depth of hundreds of metres, this too is likely to happen in the future if we don’t limit our impact on the climate. There is strong isotopic evidence that the Paleocene- Eocene thermal maximum (see Figure 19.1) was caused, at least in part, by a massive release of sea-floor methane hydrate. There is about 45 times as much carbon in the ocean (as dissolved bicarbonate ions, HCO3-) as there is in the atmosphere (as CO2), and there is a steady exchange of carbon between the two reservoirs. But the solubility of CO2 in water decreases as the temperature goes up. In other words, the warmer it gets, the more of that oceanic bicarbonate gets transferred to the atmosphere as CO2. That makes CO2 solubility another positive feedback mechanism. Vegetation growth responds positively to both increased temperatures and elevated CO2 levels, and so in general, it represents a negative feedback to climate change because the more the vegetation grows, the more CO2 is taken from the atmosphere. But it’s not quite that simple because when trees grow bigger and more vigorously, forests become darker (they have lower albedo) so they absorb more heat. Furthermore, climate warming isn’t necessarily good for vegetation growth; some areas have become too hot, too dry, or even too wet to support the plant community that was growing there, and it might take centuries for something to replace it successfully. All of these positive (and negative) feedbacks work both ways. For example, during climate cooling, growth of glaciers leads to higher albedos, and formation of permafrost results in storage of carbon that would otherwise have returned quickly to the atmosphere.

When we talk about anthropogenic climate change, we are generally thinking of the industrial era, which really got going when we started using fossil fuels (coal to begin with) to drive machinery and trains. That was around the middle of the 18th century. The issue with fossil fuels is that they involve burning carbon that was naturally stored in the crust over hundreds of millions of years as part of Earth’s process of counteracting the warming Sun. Some climate scientists argue that anthropogenic climate change actually goes back much further than the industrial era, and that humans began to impact the climate by clearing land to grow grains in Europe and the Middle East around 8,000 years BCE and by creating wetlands to grow rice in Asia around 5,000 years BCE. Clearing forests for crops is a type of climate-forcing because the CO2 storage capacity of the crops is generally lower than that of the trees they replace, and creating wetlands is a type of climate forcing because the anaerobic bacterial decay of organic matter within wetlands produces CH4. In fact, whether anthropogenic climate change started with the agricultural revolution or the industrial revolution is not important, because the really significant climate changes didn’t start until the early part of the 20th century, and although our activities are a major part of the problem, our increasing numbers are a big issue as well. Figure 19.9 shows the growth of the world population from around 5 million, when we first started growing crops, to about 18 million when wetland rice cultivation began, to over 800 million at the start of the industrial revolution, to over 7,300 million today. A big part of the incredible growth in our population is related to the availability of the cheap and abundant energy embodied in fossil fuels, which we use for transportation, heating and cooling, industry, and food production. It will be hard to support a population of this size without fossil fuels, but we have to find a way to do it. Figure 19.9 World population growth over the past 12,000 years [by SE from data at: WorldPopulationAnnual12000years_interpolated_HYDEandUN/ WorldPopulationAnnual12000years_interpolated_HYDEandUN.csv]

Physical Geology 522 and an increasing dependence on fossil fuels have driven the anthropogenic climate change of the past century. The trend of mean global temperatures since 1880 is shown in Figure 19.10. For approximately the past 55 years, the temperature has increased at a relatively steady and disturbingly rapid rate, especially compared to past changes. The average temperature now is approximately 0.8°C higher than before industrialization, and two-thirds of this warming has occurred since 1975. One of the driving factors of the recent increase in the rate of climate change has been the migration of North Americans from city centres to the suburbs, and the resulting need for virtually every household to own at least one car, when previously they were able to get around on foot or public transit.

The Intergovernmental Panel on Climate Change (IPCC), established by the United Nations in 1988, is responsible for reviewing the scientific literature on climate change and issuing periodic reports on several topics, including the scientific basis for understanding climate change, our vulnerability to observed and predicted climate changes, and what we can do to limit climate change and minimize its impacts. Figure 19.11, from the fifth report of the IPCC, issued in 2014, shows the relative contributions of various GHGs and other factors to current climate forcing, based on the changes from levels that existed in 1750. Figure 19.12 shows the IPCC’s projections for temperature increases over the next 100 years. The biggest anthropogenic contributor to warming is the emission of CO2, which accounts for 50% of positive forcing. CH4 and its atmospheric derivatives (CO2, H2O, and O3) account for 29%, and the halocarbon gases (mostly leaked from air-conditioning appliances) and nitrous oxide (N2O) (from burning fossils fuels) account for 5% each. Carbon monoxide (CO) (also produced by burning fossil fuels) accounts for 7%, and the volatile organic compounds other than methane (NMVOC) account for 3%. CO2 emissions come mostly from coal- and gas-fired power stations, motorized vehicles (cars, trucks, and aircraft), and industrial operations (e.g., smelting), and indirectly from forestry. CH4 emissions come from production of fossil fuels (escape from coal mining and from gas and oil production), livestock farming (mostly beef), landfills, and wetland rice farming. N2O and CO come mostly from the combustion of fossil fuels. In summary, close to 70% of our current GHG emissions come from fossil fuel production and use, while most of the rest comes from agriculture and landfills.

Exercise 19.3 What Does Radiative Forcing Tell Us? The bottom part of Figure 19.11 shows the total radiative forcing levels for 2011, 1980, and 1950, expressed relative to the forcing that existed in 1750. This forcing is measured in radiance at Earth’s surface in watts per square metre. For reference, the daily average irradiance for Earth is approximately 240 W/m2, so compared with 1750, we’ve increased that by 2.29 W/m2, or a little under 1%. We can use radiative forcing numbers to estimate the impact on Earth’s surface temperature by applying the following simple equation: ΔT = ΔF * 0.8, where ΔT is the expected change in average surface temperature and ΔF is the change in radiative forcing. Applying this to the value for 2011, we get ΔT = 0.8 * 2.29 = 1.8°C. From Figure 19.10, you can see that the global temperature difference between 1880 and 2011 is 0.8 – (-0.6) = 1.4°C. The temperature change between 1750 and 1880 could have been close to 0.4°C, so that puts us in about the right range. Use the ΔT = ΔF * 0.8 equation to estimate the temperature differences for 1950 and 1980, and see how those compare with the actual temperatures from Figure 19.10.

We’ve all experienced the effects of climate change over the past decade. However, it’s not straightforward for climatologists to make the connection between a warming climate and specific weather events, and most are justifiably reluctant to ascribe any specific event to climate change. In this respect, the best measures of climate change are those that we can detect over several decades, such as the temperature changes shown in Figure 19.10, or the sea-level rise shown in Figure 19.13. As already stated, sea level has risen approximately 20 cm since 1750, and that rise is attributed to both warming (and therefore expanding) seawater and melting glaciers and other land-based snow and ice (melting of sea ice does not contribute directly to sea-level rise as it is already floating in the ocean).

Projections for sea-level rise to the end of this century vary widely. This is in large part because we do not know which of the above climate change scenarios (Figure 19.12) we will most closely follow, but many are in the range from 0.5 m to 2.0 m. One of the problems in predicting sea-level rise is that we do not have a strong understanding of how large ice sheets, such as Greenland and Antarctica, will respond to future warming. Another issue is that the oceans don’t respond immediately to warming. For example, with the current amount of warming, we are already committed to a future sea-level rise of between 1.3 m and 1.9 m, even if we could stop climate change today. This is because it takes decades to centuries for the existing warming of the atmosphere to be transmitted to depth within the oceans and to exert its full impact on large glaciers. Most of that committed rise would take place over the next century, but some would be delayed longer. And for every decade that the current rates of climate change continue, that number increases by another 0.3 m. In other words, if we don’t make changes quickly, by the end of this century we’ll be locked into 3 m of future sea-level rise.

Physical Geology 526 by 2070 approximately 150 million people living in coastal areas could be at risk of flooding due to the combined effects of sea-level rise, increased storm intensity, and land subsidence. The assets at risk (buildings, roads, bridges, ports, etc.) are in the order of $35 trillion ($35,000,000,000,000). Countries with the greatest exposure of population to flooding are China, India, Bangladesh, Vietnam, U.S.A., Japan, and Thailand. Some of the major cities at risk include Shanghai, Guangzhou, Mumbai, Kolkata, Dhaka, Ho Chi Minh City, Tokyo, Miami, and New York. One of the other risks for coastal populations, besides sea-level rise, is that climate warming is also associated with an increase in the intensity of tropical storms (e.g., hurricanes or typhoons), which almost always bring serious flooding from intense rain and storm surges. Some recent examples are New Orleans in 2005 with Hurricane Katrina, and New Jersey and New York in 2012 with Hurricane Sandy (Figure 19.14).

Tropical storms get their energy from the evaporation of warm seawater in tropical regions. In the Atlantic Ocean, this takes place between 8° and 20° N in the summer. Figure 19.15 shows the variations in the sea-surface temperature (SST) of the tropical Atlantic Ocean (in blue) versus the amount of power represented by Atlantic hurricanes between 1950 and 2008 (in red). Not only has the overall intensity of Atlantic hurricanes increased with the warming since 1975, but the correlation between hurricanes and sea-surface temperatures is very strong over that time period.

A similar trend is evident for British Columbia, based on weather data from 1945 to 2005 for 29 stations distributed around the province (Figure 19.17). Of those stations, 19 show an increase in precipitation and 10 show a decrease; although the decreases are all less than 12%, some of the increases are greater than 48%. Based on the data from these stations, it is estimated that approximately 60 mm/year more precipitation fell on British Columbia in 2005 compared with 1945. That is equivalent to about six months of the average flow of the Fraser River.

While the overall amount of precipitation (total volume of rain plus snow) increased at 19 out of 29 stations between 1945 and 2005, the amount of snowfall decreased at every single station. This is a disturbing trend for operators of winter resorts and hydroelectric dams, the Wildfire Management Branch, people who drink water from reservoirs that are replenished by snow, and people who eat food that is grown across western Canada and is irrigated with water derived from melting snow.

The geographical ranges of diseases and pests, especially those caused or transmitted by insects, have been shown to extend toward temperate regions because of climate change. West Nile virus and Lyme disease are two examples that already directly affect Canadians, while dengue fever could be an issue in the future. Canadians are also indirectly affected by the increase in populations of pests such as the mountain pine beetle (Figure 19.18).

A summary of the impacts of climate change on natural disasters is given in Figure 19.19. The major types of disasters related to climate are floods and storms, but the health implications of extreme temperatures are also becoming a great concern. In the decade 1971 to 1980, extreme temperatures were the fifth most common natural disasters; by 2001 to 2010, they were the third most common.

For several weeks in July and August of 2010, a massive heat wave affected western Russia, especially the area southeast of Moscow, and scientists have stated that climate change was a contributing factor. Temperatures soared to over 40°C, as much as 12°C above normal over a wide area, and wildfires raged in many parts of the country (Figure 19.20). Over 55,000 deaths are attributed to the heat and to respiratory problems associated with the fires.

What property of greenhouse gases allows them to absorb infrared radiation and thus trap heat within the atmosphere? Explain why the emission of CO2 from fossil fuel use is a climate forcing, while the solubility of CO2 in seawater is a climate feedback. Explain how the positioning of Gondwana at the South Pole contributed to glaciation during the Paleozoic. Most volcanic eruptions lead to short-term cooling, but long-term sustained volcanism can lead to warming. Describe the mechanisms for these two different consequences. Using the orbital information on eccentricity, tilt, and precession, we could calculate variations in insolation for any latitude on Earth and for any month of the year. Why is it useful to choose the latitude of 65° as opposed to something like 30°? Why north instead of south? Why July instead of January? If the major currents in the oceans were to slow down or stop, how would that affect the distribution of heat on Earth, and what effect might that have on glaciation? Explain the climate implications of the melting and breakdown of permafrost.

Describe the importance of geological resources to our way of life Summarize the types of materials mined in Canada and explain some of the processes involved in the formation of metal deposits Explain how a metal deposit is developed into a mine Define acid rock drainage (ARD) and discuss why some mines can lead to ARD and contamination of the environment by metals Summarize some of the important industrial materials extracted in Canada and describe what they are used for Describe the processes that lead to the formation of coal deposits Explain the processes that lead to the formation of oil and gas, the distinction between source rocks and reservoir rocks, and the importance of traps Describe the origins and recovery of some of the unconventional fossil fuels Describe the origins, discovery, and extraction of diamonds in Canada

535 Steven Earle It has been said that “if you can’t grow it, you have to mine it,” meaning that anything we can’t grow we have to extract from Earth in one way or another. This includes water, of course, our most important resource, but it also includes all the other materials that we need to construct things like roads, dams, and bridges, or manufacture things like plates, toasters, and telephones. Even most of our energy resources come from Earth, including uranium and fossil fuels, and much of the infrastructure of this electrical age depends on copper (Figure 20.1). Virtually everything we use every day is made from resources from Earth. For example, let’s look at a tablet computer (Figure 20.2). Most of the case is made of a plastic known as ABS, which is made from either gas or petroleum. Some tablets have a case made from aluminum. The glass of a touch screen is made mostly from quartz combined with smaller amounts of sodium oxide (Na2O), sodium carbonate (Na2CO3), and calcium oxide (CaO). To make it work as a touch screen, the upper surface is coated with indium tin oxide. When you touch the screen you’re actually pushing a thin layer of polycarbonate plastic (made from petroleum) against the coated glass — completing an electrical circuit. The computer is then able to figure out exactly where you touched the screen. Computer processors are made from silica wafers (more quartz) and also include a significant amount of copper and gold. Gold is used because it is a better conductor than copper and doesn’t tarnish the way silver or copper does. Most computers have nickel-metal-hydride (NiMH) batteries, which contain nickel, of course, along with cadmium, cobalt, manganese, aluminum, and the rare-earth elements lanthanum, cerium, neodymium, and praseodymium. The processor and other electronic components are secured to a circuit board, which is a thin layer of fibreglass sandwiched between copper sheets coated with small amounts of tin and lead. Various parts are put together with steel screws that are made of iron and molybdenum.

That’s not everything that goes into a tablet computer, but to make just those components we need a pure-silica sand deposit, a salt mine for sodium, a rock quarry for calcium, an oil well, a gas well, an aluminum mine, an iron mine, a manganese mine, a copper-molybdenum-gold mine, a cobalt-nickel mine, a rare-earth element and indium mine, and a source of energy to transport all of the materials, process them, put them together, and finally transport the computer to your house or the store where you bought it.

Mining has always been a major part of Canada’s economy. Canada has some of the largest mining districts and deposits in the world, and for the past 150 years, we have been one of the most important suppliers of metals. Extraction of Earth’s resources goes back a long way in Canada. For example, the First Nations of British Columbia extracted obsidian from volcanic regions for tools and traded it up and down the coast. In the 1850s, gold was discovered in central British Columbia, and in the 1890s, even more gold was discovered in the Klondike area of Yukon. These two events were critical to the early development of British Columbia, Yukon, and Alaska. Canada’s mining sector had revenues in the order of $37 billion in 2013. The majority of that was split roughly equally among gold, iron, copper, and potash, with important but lesser amounts from nickel and diamonds (Figure 20.3). Revenues from the petroleum sector are significantly higher, at over $100 billion per year.

A metal deposit is a body of rock in which one or more metals have been concentrated to the point of being economically viable for recovery. Some background levels of important metals in average rocks are shown on Table 20.1, along with the typical grades necessary to make a viable deposit, and the corresponding concentration factors. Looking at copper, for example, we can see that while average rock has around 40 ppm (parts per million) of copper, a grade of around 10,000 ppm or 1% is necessary to make a viable copper deposit. In other words, copper ore has about 250 times as much copper as typical rock. For the other elements in the list, the concentration factors are much higher. For gold, it’s 2,000 times and for silver it’s around 10,000 times.

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

A magmatic deposit is one in which the metal concentration takes place primarily at the same time as the formation and emplacement of the magma. Most of the nickel mined in Canada comes from magmatic deposits such as those in Sudbury (Ontario), Thompson (Manitoba) (Figure 20.4), and Voisey’s Bay (Labrador). The magmas from which these deposits form are of mafic or ultramafic composition (derived from the mantle), and therefore they have relatively high nickel and copper contents to begin with (as much as 100 times more than normal rocks in the case of nickel). These elements may be further concentrated within the magma as a result of the addition of sulphur from partial melting of the surrounding rocks. The heavy nickel and copper sulphide minerals are then concentrated further still by gravity segregation (i.e., crystals settling toward the bottom of the magma chamber). In some cases, there are significant concentrations of platinum-bearing minerals. Most of these types of deposits around the world are Precambrian in age — probably because the mantle was significantly hotter at that time, and the necessary mafic and ultramafic magmas were more likely to be emplaced in the continental crust.

Much of the copper, zinc, lead, silver, and gold mined in Canada is mined from volcanic–hosted massive sulphide (VHMS) deposits associated with submarine volcanism (VMS deposits). Examples are the deposits at Kidd Creek, Ontario, Flin Flon on the Manitoba-Saskatchewan border, Britannia on Howe Sound, and Myra Falls (within Strathcona Park) on Vancouver Island. VMS deposits are formed from the water discharged at high temperature (250° to 300°C) at ocean-floor hydrothermal vents, primarily in areas of subduction-zone volcanism. The environment is comparable to that of modern-day black smokers (Figure 20.5), which form where hot metal- and sulphide-rich water issues from the sea floor. They are called massive sulphide deposits because the sulphide minerals (including pyrite (FeS2) , sphalerite (ZnS), chalcopyrite (CuFeS2), and galena (PbS)) are generally present in very high concentrations (making up the majority of the rock in some cases). The metals and the sulphur are leached out of the sea-floor rocks by convecting groundwater driven by the volcanic heat, and then quickly precipitated where that hot water enters the cold seawater, causing it to cool suddenly and change chemically. The volcanic rock that hosts the deposits is formed in the same area and at the same general time as the accumulation of the ore minerals.

Figure 20.5 Left: A black smoker on the Juan de Fuca Ridge off the west coast of Vancouver Island. Right: A model of the formation of a volcanogenic massive sulphide deposit on the sea floor. [left: NOAA at: explorations/10index/background/plumes/media/ black_smoker.html, right: SE]

Porphyry deposits are the most important source of copper and molybdenum in British Columbia, the western United States, and Central and South America. Most porphyry deposits also host some gold, which may be, in rare cases, the primary commodity. B.C. examples include several large deposits within the Highland Valley mine (Figure 20.1) and numerous other deposits scattered around the central part of the province. A porphyry deposit forms around a cooling felsicstock in the upper part of the crust. They are called “porphyry” because upper crustal stocks are typically porphyritic in texture, the result of a two-stage cooling process. Metal enrichment results in part from convection of groundwater related to the heat of the stock, and also from metal-rich hot water expelled by the cooling magma (Figure 20.6). The host rocks, which commonly include the stock itself and the surrounding country rocks, are normally highly fractured and brecciated. During the ore-forming process, some of the original minerals in these rocks are altered to potassium feldspar, biotite, epidote, and various clay minerals. The important ore minerals include chalcopyrite (CuFeS2), bornite (Cu5FeS4), and pyrite in copper porphyry deposits, or molybdenite (MoS2) and pyrite in molybdenum porphyry deposits. Gold is present as minute flakes of native gold. This type of environment (i.e., around and above an intrusive body) is also favourable for the formation of other types of deposits — particularly vein-type gold deposits (a.k.a. epithermal deposits). Some of the gold deposits of British Columbia (such as in the Eskay Creek area adjacent to the Alaska panhandle), and many of the other gold deposits situated along the western edge of both South and North America are of the vein type shown in Figure 20.6, and are related to nearby magma bodies.

Most of the world’s major iron deposits are of the banded iron formation type, and most of these formed during the initial oxygenation of Earth’s atmosphere between 2,400 and 1,800 Ma. At that time, iron that was present in dissolved form in the ocean (as Fe2+) became oxidized to its insoluble form (Fe3+) and accumulated on the sea floor, mostly as hematite interbedded with chert (Figure 20.7). Unlike many other metals, which are economically viable at grades of around 1% or even much less, iron deposits are only viable if the grades are in the order of 50% iron and if they are very large.

Figure 20.7 Banded iron formation from an unknown location in North America on display at a museum in Germany. The rock is about 2 m across. The dark grey layers are magnetite and the red layers are hematite. Chert is also present. [https://upload.wikimedia.org/ wikipedia/commons/5/5f/Black-band_ironstone_%28aka%29.jpg]

There are several different types of uranium deposits, but some of the largest and richest are those within the Athabasca Basin of northern Saskatchewan. These are called unconformity-type uranium deposits because they are all situated very close to the unconformity between the Proterozoic Athabasca Group sandstone and the much older Archean sedimentary, volcanic, and intrusive igneous rock (Figure 20.8). The origin of unconformity-type U deposits is not perfectly understood, but it is thought that two 543 Chapter 20 Geological Resources particular features are important: (1) the relative permeability of the Athabasca Group sandstone, and (2) the presence of graphitic schist within the underlying Archean rocks. The permeability of the sandstone allowed groundwater to flow through it and leach out small amounts of U, which stayed in solution in the oxidized form U6+. The graphite (C) created a reducing environment (non-oxidizing) that converted the U from U6+ to insoluble U4+, at which point it was precipitated as the mineral uraninite (UO2).

Metal deposits are mined in a variety of different ways depending on their depth, shape, size, and grade. Relatively large deposits that are quite close to the surface and somewhat regular in shape are mined using open-pit mine methods (Figure 20.1). Creating a giant hole in the ground is generally cheaper than making an underground mine, but it is also less precise, so it is necessary to mine a lot of waste rock along with the ore. Relatively deep deposits or those with elongated or irregular shapes are typically mined from underground with deep vertical shafts, declines (sloped tunnels), and levels (horizontal tunnels) (Figures 20.09 and 20.11). In this way, it is possible to focus the mining on the ore body itself. However, with relatively large ore bodies, it may be necessary to leave some pillars to hold up the roof.

A typical metal deposit might contain a few percent of ore minerals (e.g., chalcopyrite or sphalerite), mixed with the minerals of the original rock (e.g., quartz or feldspar). Other sulphide minerals are commonly present within the ore, especially pyrite. When ore is processed (typically very close to the mine), it is ground to a fine powder and the ore minerals are physically separated from the rest of the rock to make a concentrate. At a molybdenum mine, for example, this concentrate may be almost pure molybdenite (MoS2). The rest of the rock is known as tailings. It comes out of the concentrator as a wet slurry and must be stored near the mine, in most cases, in a tailings pond. The tailings pond at the Myra Falls Mine on Vancouver Island is shown in Figure 20.12, and the settling ponds for waste water from the concentrator are shown in Figure 20.13. The tailings are contained by an embankment. Also visible in the foreground of Figure 20.12 is a pile of waste rock, which is non- ore rock that was mined in order to access the ore. Although this waste rock contains little or no ore minerals, at many mines it contains up to a few percent pyrite. The tailings and the waste rock at most mines are an environmental liability because they contain pyrite plus small amounts of ore minerals. When pyrite is exposed to oxygen and water, it generates sulphuric acid — also known as acid rock drainage (ARD). Acidity itself is a problem to the environment, but because the ore elements, such as copper or lead, are more soluble in acidic water than neutral water, ARD is also typically quite rich in metals, many of which are toxic.

Figure 20.12 The tailings pond at the Myra Falls Mine on Vancouver Island. The dry rock in the middle of the image is waste rock. The structure on the right is the headframe for the mine shaft. Myra Creek flows between the tailings pond and the headframe. [SE]

Tailings ponds and waste-rock storage piles must be carefully maintained to ensure their integrity and monitored to ensure that acidic and metal-rich water is not leaking out. In August 2014, the tailings pond at the Mt. Polley Mine in central B.C. failed and 10 million cubic metres of waste water along with 4.5 million cubic metres of tailings slurry was released into Polley Lake, Hazeltine Creek, and Quesnel Lake (Figure 20.14, a and b). As of July 2015, the environmental implications of this event are still not fully understood.

Figure 20.14a The Mt. Polley Mine area prior to the dam breach of August 2014. The tailings were stored in the area labelled “retention basin.” [https://en.wikipedia.org/wiki/Mount_Polley_mine_disaster] Figure 20.14b The Mt. Polley Mine area after the tailings dam breach of August 2014. The water and tailings released flowed into Hazeltine Creek, and Polley and Quesnel Lakes. [https://en.wikipedia.org/wiki/ Mount_Polley_mine_disaster] Most mines have concentrators on site because it is relatively simple to separate ore minerals from non-ore minerals and thus significantly reduce the costs and other implications of transportation. But separation of ore minerals is only the preliminary stage of metal refinement, for most metals the second stage involves separating the actual elements within the ore minerals. For example, the most common ore of copper is chalcopyrite (CuFeS2). The copper needs to be separated from the iron and sulphur to make copper metal and that involves complicated and very energy-intensive processes that are done at

There are several metal refineries (including smelters) in Canada; some examples are the aluminum refinery in Kitimat, B.C. (which uses ore from overseas); the lead-zinc smelter in Trail, B.C.; the nickel smelter at Thompson, Manitoba; numerous steel smelters in Ontario, along with several other refining operations for nickel, copper, zinc, and uranium; aluminum refineries in Quebec; and a lead smelter in New Brunswick.

Metals are critical for our technological age, but there are a lot of other not-so-shiny materials that are needed to facilitate our way of life. For everything made out of concrete or asphalt, we need sand and gravel. To make the cement that holds concrete together, we also need limestone. For the glass in our computer screens and for glass-sided buildings, we need silica sand plus sodium oxide (Na2O), sodium carbonate (Na2CO3), and calcium oxide (CaO). Potassium is an essential nutrient for farming in many areas, and for a wide range of applications (e.g., ceramics and many industrial processes), we also need various types of clay. The best types of aggregate (sand and gravel) resources are those that have been sorted by streams, and in Canada the most abundant and accessible fluvial deposits are associated with glaciation. That doesn’t include till of course, because it has too much silt and clay, but it does include glaciofluvial outwash, which is present in thick deposits in many parts of the country, similar to the one shown in Figure 20.15. In a typical gravel pit, these materials are graded on-site according to size and then used in a wide range of applications from constructing huge concrete dams to filling children’s sandboxes. Sand is also used to make glass, but for most types of glass, it has to be at least 95% quartz (which the sandy layers shown in Figure 20.15 are definitely not), and for high-purity glass and the silicon wafers used for electronics, the source sand has to be over 98% quartz.

Approximately 80 million tonnes of concrete are used in Canada each year — a little over 2 tonnes per person. The cement used for concrete is made from approximately 80% calcite (CaCO3) and 20% clay. This mixture is heated to 1450°C to produce the required calcium silicate compounds (e.g., Ca2SiO4). The calcite typically comes from limestone quarries like the one on Texada Island, B.C. (Figure 20.16). Limestone is also used as the source material for many other products that require calcium compounds, including steel and glass, pulp and paper, and plaster products for construction.

Sodium is required for a wide range of industrial processes, and the most convenient source is sodium chloride (rock salt), which is mined from evaporite beds in various parts of Canada. The largest salt mine in the world is at Goderich, Ontario, where salt is recovered from the 100 m thick Silurian Salina Formation. The same formation is mined in the Windsor area. Rock salt is also used as a source of sodium and chlorine in the chemical industry to melt ice on roads, as part of the process of softening water, and as a seasoning. Under certain conditions, the mineral sylvite (KCl) accumulates in evaporite beds, and this rock is called potash. This happened across the Canadian prairies during the Devonian, creating the Prairie evaporite formation (Figure 6.16). Potassium is used as a crop fertilizer, and Canada is the world’s leading supplier, with most of that production coming from Saskatchewan. Another evaporite mineral, gypsum (CaSO4.2H20), is the main component of plasterboard (drywall) that is widely used in the construction industry. One of the main mining areas for gypsum in Canada is in the Milford Station area of Nova Scotia, site of the world’s largest gypsum mine. Rocks are quarried or mined for many different uses, such as building facades (Figure 20.17), countertops, stone floors, and headstones. In most of these cases, the favoured rock types are granitic rocks, slate, and marble. Quarried rock is also used in some applications where rounded gravel isn’t suitable, such as the ballast (road bed) for railways, where crushed angular rock is needed.

There are numerous types of fossil fuels, but all of them involve the storage of organic matter in sediments or sedimentary rocks. Fossil fuels are rich in carbon and almost all of that carbon ultimately originates from CO2 taken out of the atmosphere during photosynthesis. That process, driven by solar energy, involves reduction (the opposite of oxidation) of the carbon, resulting in it being combined with hydrogen instead of oxygen. The resulting organic matter is made up of complex and varied carbohydrate molecules. Most organic matter is oxidized back to CO2 relatively quickly (within weeks or years in most cases), but any of it that gets isolated from the oxygen of the atmosphere (for example, deep in the ocean or in a stagnant bog) may last long enough to be buried by sediments and, if so, may be preserved for tens to hundreds of millions of years. Under natural conditions, that means it will be stored until those rocks are eventually exposed at the surface and weathered. In this section, we’ll discuss the origins and extraction of the important fossils fuels, including coal, oil, and gas. Coal, the first fossil fuel to be widely used, forms mostly on land in swampy areas adjacent to rivers and deltas in areas with humid tropical to temperate climates. The vigorous growth of vegetation leads to an abundance of organic matter that accumulates within stagnant water, and thus does not decay and oxidize. This situation, where the dead organic matter is submerged in oxygen-poor water, must be maintained for centuries to millennia in order for enough material to accumulate to form a thick layer (Figure 20.18a). At some point, the swamp deposit is covered with more sediment — typically because a river changes its course or sea level rises (Figure 20.18b). As more sediments are added, the organic matter starts to become compressed and heated. Low-grade lignite coal forms at depths between a few 100 m and 1,500 m and temperatures up to about 50°C (Figure 20.18c). At between 1,000 m to 5,000 m depth and temperatures up to 150°C m, bituminous coal forms (Figure 20.18d). At depths beyond 5,000 m and temperatures over 150°C, anthracite coal forms. Figure 20.18 Formation of coal: (a) accumulation of organic matter within a swampy area; (b) the organic matter is covered and compressed by deposition of a new layer of clastic sediments; (c) with greater burial, lignite coal forms; and (d) at even greater depths, bituminous and eventually anthracite coal form. [SE]

There are significant coal deposits in many parts of Canada, including the Maritimes, Ontario, Saskatchewan, Alberta, and British Columbia. In Alberta and Saskatchewan, much of the coal is used for electricity generation. Coal from the Highvale Mine (Figure 20.19), Canada’s largest, is used to feed the Sundance and Keephills power stations west of Edmonton. Almost all of the coal mined in British Columbia is exported for use in manufacturing steel.

While almost all coal forms on land from terrestrial vegetation, most oil and gas is derived primarily from marine micro-organisms that accumulate within sea-floor sediments. In areas where marine productivity is high, dead organic matter is delivered to the sea floor fast enough that some of it escapes oxidation. This material accumulates in the muddy sediments, which become buried to significant depth beneath other sediments. As the depth of burial increases, so does the temperature — due to the geothermal gradient — and gradually the organic matter within the sediments is converted to hydrocarbons (Figure 20.20). The first stage is the biological production (involving anaerobic bacteria) of methane. Most of this escapes back to the surface, but some is trapped in methane hydrates near the sea floor. At depths beyond about 2 km, and at temperatures ranging from 60° to 120°C, the organic matter is converted by chemical processes to oil. This depth and temperature range is known as the oil window. Beyond 120°C most of the organic matter is chemically converted to methane.

The organic matter-bearing rock within which the formation of gas and oil takes place is known to petroleum geologists as the source rock. Both liquid oil and gaseous methane are lighter than water, so as liquids and gases form, they tend to move slowly toward the surface, out of the source rock and into reservoir rocks. Reservoir rocks are typically relatively permeable because that allows migration of the fluids from the source rocks, and also facilitates recovery of the oil or gas. In some cases, the liquids and gases make it all the way to the surface, where they are oxidized, and the carbon is returned to the atmosphere. But in other cases, they are contained by overlying impermeable rocks (e.g., mudrock) in situations where anticlines, faults, stratigraphy changes, and reefs or salt domes create traps (Figure 20.21).

The liquids and gases that are trapped within reservoirs become separated into layers based on their density, with gas rising to the top, oil below it, and water underneath the oil. The proportions of oil and gas depend primarily on the temperature in the source rocks. Some petroleum fields, such as many of

Figure 20.22 Seismic section through the East Breaks Field in the Gulf of Mexico. The dashed red line marks the approximate boundary between deformed rocks and younger undeformed rocks. The wiggly arrows are interpreted migration paths. The total thickness of this section is approximately 5 km. [SE after File:Sedimentary-basin-analysis_fig4-55.png] In general, petroleum fields are not visible from the surface, and their discovery involves the search for structures in the subsurface that have the potential to form traps. Seismic surveys are the most commonly used tool for early-stage petroleum exploration, as they can reveal important information about the stratigraphy and structural geology of subsurface sedimentary rocks. An example from the Gulf of Mexico south of Texas is shown in Figure 20.22. In this area, a thick evaporite deposit (“salt”) has formed domes because salt is lighter than other sediments and tends to rise slowly toward the surface; this has created traps. The sequence of deformed rocks is capped with a layer of undeformed rock.

The type of oil and gas reservoirs illustrated in Figures 20.21 and 20.22 are described as conventional reserves. Some unconventional types of oil and gas include oil sands, shale gas, and coal-bed methane. Oil sands are important because the reserves in Alberta are so large (the largest single reserve of oil in the world), but they are very controversial from an environmental and social perspective. They are “unconventional” because the oil is exposed near the surface and is highly viscous because of microbial changes that have taken place at the surface. The hydrocarbons that form this reserve originated in deeply buried Paleozoic rocks adjacent to the Rocky Mountains and migrated up and toward the east (Figure 20.23). The oil sands are controversial primarily because of the environmental cost of their extraction. Since the oil is so viscous, it requires heat to make it sufficiently liquid to process. This energy comes from gas; approximately 25 m3 of gas is used to produce 0.16 m3 (one barrel) of oil. (That’s not quite as bad as it sounds, as the energy equivalent of the required gas is about 20% of the energy embodied in the produced oil.) The other environmental cost of oil sands production is the devastation of vast areas of land where strip-mining is taking place and tailings ponds are constructed, and the unavoidable release of contaminants into the groundwater and rivers of the region. At present, most oil recovery from oil sands is achieved by mining the sand and processing it on site. Exploitation of oil sand that is not exposed at the surface depends on in situ processes, an example being the injection of steam into the oil-sand layer to reduce the viscosity of the oil so that it can be pumped to the surface.

Shale gas is gas that is trapped within rock that is too impermeable for the gas to escape under normal conditions, and it can only be extracted by fracturing the reservoir rock using water and chemicals under extremely high pressure. This procedure is known as hydraulic fracturing or “fracking.” Fracking is controversial because of the volume of water used, and because, in some jurisdictions, the fracking companies are not required to disclose the nature of the chemicals used. Although fracking is typically done at significant depths, there is always the risk that overlying water-supply aquifers could be contaminated (Figure 20.24). Fracking also induces low-level seismicity. During the process that converts organic matter to coal, some methane is produced, which is stored within the pores of the coal. When coal is mined, methane is released into the mine where it can become a serious explosion hazard. Modern coal-mining machines have methane detectors on them and actually stop operating if the methane levels are dangerous. It is possible to extract the methane from coal beds without mining the coal; gas recovered this way is known as coal-bed methane. Figure 20.24 Depiction of the process of directional drilling and fracking to recover gas from impermeable rocks. The light blue arrows represent the potential for release of fracking chemicals to aquifers. [by SE, after https://en.wikipedia.org/wiki/ Hydraulic_fracturing#/media/File:HydroFrac2.svg]

Although Canada’s diamond mining industry didn’t get started until 1998, diamonds are currently the sixth most valuable product mined in the country (Figure 20.3), and Canada ranks sixth in the world in diamond production. Diamonds form deep in the mantle (approximately 200 km to 250 km depth) under very specific pressure and temperature conditions, from carbon that is naturally present in mantle rock (not from coal). The diamond-bearing rock is brought to the surface coincidentally via a type of volcanism that is extremely rare (the most recent kimberlite eruption is thought to have taken place 10,000 years ago and prior to that at around 30 Ma). There is more on the volcanology of kimberlites in section 4.3. All of the world’s kimberlite diamond deposits are situated within ancient shield areas (cratons) in Africa, Australia, Russia, South America, and North America. It has long been known that diamonds could exist within the Canadian Shield, but up until 1991, exploration efforts had been unsuccessful. In 1980 two geologists, Chuck Fipke and Stu Blusson, started searching in the Northwest Territories by sampling glacial sediments looking for some of the minerals that are normally quite abundant within kimberlites: chromium-bearing garnet, chromium- bearing pyroxene, chromite (Cr2O3), and ilmenite (FeTiO3). These distinctive minerals are used for this type of exploration because they are many times more abundant in kimberlite than diamond is. After more than a decade of exploration, Fipke and Blusson finally focused their search on an area 250 km northeast of Yellowknife, and, in 1991, they announced the discovery of a diamond-bearing kimberlite body at Lac de Gras. That discovery is now the Diavik Mine, and there is another diamond mine — Ekati — 25 km to the northwest (Figure 20.25). There are two separate mines at Diavik accessing three different kimberlite bodies, and there are five at Ekati. See Figure 4.22 for a close-up view of the Ekati Mine. There are six operating diamond mines in Canada, four in the Northwest Territories (including Diavik and Ekati), and one each in Nunavut and Ontario. Figure 20.25 Diamond mines in the Lac de Gras region, Nunavut. The twin pits of the Diavik Mine are visible in the lower right on an island within Lac de Gras. The five pits of the Ekati mine are also visible, on the left and the upper right. The two main mine centres are 25 km apart. view.php?id=84085&src=eoa-iotd]