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CHAPTER 3

Movements of the Earth from Within

EARTH-SURFACE MOVEMENTS DUE TO PLATE TECTONICS

TO UNDERSTAND THE CHANGES AND MOVEMENTS affecting the appearance of the landscape on large scales we need to review current understanding of some geological systems, especially plate tectonics. Many of the widespread changes that have created landscapes over long periods of time can now be understood using this discovery.

Knowledge of the processes causing the movement of large areas of the Earth’s surface (10–1000 km length scale) has been revolutionised by scientific advances made over the last 50 years. During this time, scientists have become convinced that the whole of the Earth’s surface consists of an outer shell of interlocking tectonic plates (Fig. 12). The word tectonic refers to processes that have built features of the Earth’s crust (Greek: tekt, a builder). The worldwide plate pattern is confusingly irregular – particularly when seen on a flat map – and it is easier to visualise the plates in terms of an interlocking arrangement of panels on the Earth’s spherical surface, broadly like the panels forming the skin of a traditional leather football.

Tectonic plates are features of the lithosphere, the name given to the ~125 km- thick outer shell of the Earth, distinguished from the material below by the strength of its materials (Greek: lithos, stone). The strength depends upon the composition of the material and also upon its temperature and pressure, both of which tend to increase with depth below the Earth’s surface. In contrast to the mechanically strong lithosphere, the underlying material is weaker and known as the asthenosphere (Greek: asthenos, no-strength). Note that in Figure 13 the crustal and outer mantle layers are shown with exaggerated thickness, so that they are visible.

FIG 12. World map showing the present pattern of the largest lithosphere plates.

FIG 13. Diagram of the internal structure of the Earth.

Much of the strength difference between the lithosphere and the asthenosphere depends on the temperature difference between them. The lithosphere plates are cooler than the underlying material, so they behave in a more rigid way when forces are generated within the Earth. The asthenosphere is hotter and behaves in a more plastic way, capable of deforming without fracturing and, to some extent, of ‘flowing’. Because of this difference in mechanical properties and the complex internal forces present, the lithosphere plates can move relative to the material below. To visualise the motion of the plates, we can use the idea of lithospheric plates floating on top of the asthenosphere.

The pattern of earthquake activity and actively unstable mountain belts corresponds very well with the pattern of the tectonic plates now recognised. The largest plates (Fig. 12) clearly mark relatively rigid and stable areas of the lithosphere, with interiors that do not experience as much disturbance as their edges. Plates move relative to each other along plate boundaries, in various ways that will be described below. The plate patterns have been located by investigating distinctive markers within the plates and at their edges, allowing the relative rates of movement between neighbouring plates to be calculated. These rates are very slow, rarely exceeding a few centimetres per year, but over the millions of years of geological time they can account for thousands of kilometres of relative movement.

It has proved much easier to measure plate movements than to work out what has been causing them. However, the general belief today is that the plates move in response to a number of different forces. Circulation (convection) within the mantle is driven by temperature and density differences, but other forces are also at play. Where plates diverge, warm material rises from within the Earth to fill the surface gap, and, being warmer, it may also be elevated above the rest of the plate, providing a pushing force to move the plate across the surface of the Earth. At convergent boundaries, cold, older material sinks into the asthenosphere, providing a pulling force that drags the rest of the plate along behind it. Deep within the Earth, the sinking material melts and is ultimately recycled and brought back to the surface to continue the process.

Knowledge of how tectonic plates interact provides the key to understanding the movement history of the Earth’s crust. However, most people are much more familiar with the geographical patterns of land and sea, which do not coincide with the distribution of tectonic plates (Fig. 12). From the point of view of landscapes and scenery, coastlines are always going to be key features because they define the limits of the land; we make no attempt in this book to consider submarine scenery in detail.

The upper part of the lithosphere is called the crust (Fig. 13). Whereas the distinction between the lithosphere and the asthenosphere is based upon mechanical properties related to temperature and pressure, the distinction between the crust and the lower part of the lithosphere is based upon composition. Broadly speaking, there are two types of crust that can form the upper part of the lithosphere: continental and oceanic. An individual tectonic plate may include just one or both kinds of crust.

Continental crust underlies land areas and also many of the areas covered by shallow seas. Geophysical work shows that this crust is typically about 30 km thick, but may be 80–90 km thick below some high plateaus and mountain ranges. The highest mountains in Britain are barely noticeable on a scale diagram comparing crustal thicknesses (Fig. 14). Continental crust is made of rather less dense materials than the oceanic crust, or the mantle, and this lightness is the reason why land surfaces and shallow sea floors are elevated compared to the deep oceans. Much of the continental crust is very old (up to 3–4 billion years), having formed early in the Earth’s life when lighter material separated from denser materials within the Earth and rose to the surface.

Oceanic crust forms the floors of the deep oceans, typically 4 or 5 km below sea level. It is generally 5–10 km thick and is distinctly more dense than continental crust. Oceanic crust only forms land where volcanic material has been supplied to it in great quantity (as in the case of Iceland), or where other important local forces in the crust have caused it to rise (as is the case in parts of Cyprus). Oceanic crust is generally relatively young (only 0–200 million years old), because its greater density and lower elevation ensures that it is generally subducted and destroyed at plate boundaries that are convergent.

FIG 14. Scale diagram comparing average thicknesses of oceanic and continental crust and lithosphere.

Figure 12 shows the major pattern of tectonic plates on the Earth today. The Mercator projection of this map distorts shapes, particularly in polar regions, but we can see that there are seven very large plates, identified by the main areas located on their surfaces. The Pacific plate lacks continental crust entirely, whereas the other six main plates each contain a large continent (Eurasia, North America, Australia, South America, Africa and Antarctica) as well as oceanic crust. There are a number of other middle-sized plates (e.g. Arabia and India) and large numbers of micro-plates, not shown on the world map.

Figures 12 and 15 also identify the different types of plate boundary, which are distinguished according to the relative motion between the two plates. Convergent plate boundaries involve movement of the plates from each side towards the suture (or central zone) of the boundary. Because the plates are moving towards each other, they become squashed together in the boundary zone. Sometimes one plate moves below the other in a process called subduction, which often results in a deep ocean trench and a zone of mountains and/or volcanoes, as well as earthquake activity (Fig. 15). The earthquakes that happened off Indonesia in December 2004 and off Japan in March 2011 were two of the strongest known since records began. Both seized world attention because of the horrifying loss of life cause by the tsunami waves they generated. Both were the result of sudden lithosphere movements of several metres on faults in the convergent subduction zones where the Australian and Pacific plates have been moving under the Eurasian plate (Fig. 12).

FIG 15. Diagram (not to scale) illustrating the movement processes of plates.

In other cases the plate boundary is divergent, where the neighbouring plates move apart and new material from deeper within the Earth rises to fill the space created. New oceanic crust is created by the arrival and cooling of hot volcanic material from below. The Mid-Atlantic Ridge running through Iceland is one of the examples nearest to Britain of this sort of plate boundary, and volcanic ash-cloud activity there caused widespread disruption to air transport during 2010.

Other plate boundaries, sometimes called transform boundaries, mainly involve movement parallel to the plate edges. The Californian coast zone is the classic example but there are many others, such as the transform boundary between the African and Antarctic plates. In some areas, plate movement is at an oblique angle to the suture and there are components of divergence or convergence as well as movement parallel to the boundary.

Britain today sits in the stable interior of the western Eurasian plate, almost equidistant from the divergent Mid-Atlantic Ridge boundary to the west and the complex convergent boundary to the south where Spain and northwest Africa are colliding. In its earlier history the crust of Britain has been subjected to very direct plate boundary activity. The results of convergent activity in Devonian and Carboniferous times (between 416 and 299 million years ago) are visible at the surface in southwest England, and in Ordovician to Devonian times (between 490 and 360 million years ago) in Wales, northwest England and Scotland (see Chapter 4).

Present-day plate boundaries are often picked out by the location of earthquakes, as described above. Mention should also be made at this point of the importance of volcanoes and igneous bedrock in providing information about movements within the upper levels of the Earth. Highly sophisticated analytical work has illuminated the whole subject of the chemical and mineral evolution of igneous material as it evolves and moves in the crust. For the purposes of this book, a very simple twofold division of igneous rocks into felsic and mafic will be sufficient.

Felsic igneous rocks tend to be light-coloured and of relatively low density, containing the minerals quartz and feldspar. Typical types are granite, syenite (coarsely crystalline) and rhyolite (finely crystalline). Continental crust consists of felsic and mafic igneous rocks, as well as sedimentary and metamorphic rocks.

Mafic igneous rocks tend to be darker-coloured and of relatively high density, containing feldspar and dark minerals rich in magnesium and/or iron, such as augite or hornblende. Typical types are gabbros (coarsely crystalline), andesites and basalts (finely crystalline). Oceanic crust is dominated by mafic igneous materials.

MAKING LOCAL MEASUREMENTS OF EARTH SURFACE MOVEMENTS

We have been considering the large movement systems that originate within the Earth. There are also more local movement systems operating on the Earth’s surface, which are linked to a very variable degree to the large-scale movements of plate tectonics. To explore this complex linkage further, it will be helpful to look now at different processes that may combine to cause particular local movements.

Tectonic plates are defined by their rigidity, so there is relatively little horizontal movement between points within the same plate, compared to the deformation seen in plate boundary zones. This extreme deformation may involve folding and fracturing of the rock materials, addition of new material from below, or absorption of material into the interior during subduction.

Nonetheless, deformation is not restricted solely to plate boundaries and does occur within the plates, although to a lesser extent. In some cases, major structures that originally formed along a plate boundary can become incorporated into the interior of a plate when prolonged collision causes two plates to join. The Caledonian convergent boundary that extended across Scotland (see Chapter 4) provides an excellent example of movements that occurred hundreds of millions of years ago, but also contains many examples of structures formed in later movements. These structures have often been reactivated long after they first formed in order to accommodate forces along the new plate boundary via deformation within the plate. Conversely, changes of internal stress patterns can sometimes lead to the splitting of a plate into two, forming a new, initially divergent plate boundary. Many of the oil- and gas-containing features of the North Sea floor (Fig. 2) originated when a belt of divergent rift faults formed across a previously intact plate.

It needs to be stressed that the patterns of deformation (fracturing and folding) due to these plate motions occur at a wide range of different scales, from centimetres to thousands of kilometres. Sometimes they are visible at the scale of an entire plate boundary, such as the enormous Himalayan mountain chain that marks the collision of India with Asia.

The effects of features as large as plate boundaries on landscapes persist over hundreds of millions of years, long after the most active movement has ceased. For example, parts of southwestern England, Wales and the Scottish Highlands are underlain by bedrocks that were formed in convergent boundary zones of the past. The tin and lead mines of Cornwall owe their existence to a 300-million-year-old convergent plate boundary, where an ocean was destroyed as two plates converged and continents collided. The convergence released molten rock that rose in the crust and gradually cooled to form granite, whilst metals were precipitated in the surrounding crust as ‘lodes’ containing tin and lead (see Chapter 4).

Mapping the patterns of bedrock exposed at the surface often reveals folds and faults that provide key information about the movements that have taken place during the past (Fig. 16). Figure 17 provides a key to some of the terms commonly used to classify these structures, as a step towards understanding the sorts of movement patterns that they represent. In broad terms, folds tend to indicate some form of local convergent movement, though they may be the result of larger movement patterns of a different kind. Normal faults tend to indicate divergent or stretching movements, at least locally, whereas reverse and strike-slip faults tend to indicate convergence. Two broad types of fold are distinguished: synclines are U-shaped downfolds, while anticlines are the opposite – A-shaped upfolds.

Further mapping of folds and faults often reveals complex patterns of changing movements. A complex example is shown in Figure 18. Divergent movements in an area of crust produced plastic deformation in the warmer lower crust, and faulting into a number of discrete blocks in the colder, more brittle upper crust. This was then followed by an episode of convergent movement that resulted in closing up the upper crustal blocks and further flow in the plastic lower crust, causing crustal thickening and mountain building at the surface.

FIG 16. Outcrop in the Atacama Desert, Chile, showing a very regularly bedded succession of mudstones, formed originally as horizontally layered deposits in a lake. Since their deposition the mudstones have been tilted. They have also been fractured during an earthquake, resulting in a step, or normal fault (see Fig. 17), that is particularly clear because it has cut through a white layer in the deposits. An outcrop such as this makes it possible to measure the local movements that have taken place in this material after it was deposited. (© Nicholas Branson)

VERTICAL CRUSTAL MOVEMENTS

The movement of lithospheric plates, as described above, is the main cause of horizontal convergent and divergent movements affecting thousands of kilometres of the Earth’s surface. As shown in Figures 16 to 18, horizontal movements are generally accompanied by vertical movements of local crustal surfaces. Some of these could have produced very large scenic features, such as a mountain belt or a rift valley. In this book we are primarily concerned with scenic features at a more local scale, so we now consider various other processes that may contribute to the creation of vertical crustal movements.

FIG 17. The most important types of folds and faults, and the local patterns of forces responsible.

Vertical crustal movement linked to erosion or deposition

Addition or subtraction of material to the surface of the Earth is happening all the time as sediment is deposited or solid material is eroded. The discipline of sedimentology is concerned with the wide range of different processes that are involved in the erosion, transport and deposition of material, whether the primary agent of movement is water, ice, mud or wind. An important point is that few of these sedimentary processes relate directly to the large tectonic movements of the Earth’s crust that we have discussed above. Landscape is often produced by erosion of thick sedimentary deposits that formed in sedimentary basins where material eroded from the surrounding uplands accumulated. One of the characteristic features of these thick deposits is their layered appearance – as, for example, in the Torridonian Sandstones of northwestern Scotland (see Chapter 4). Layering varies from millimetre-scale laminations produced by very small fluctuations in depositional processes, to sheets hundreds of metres thick that extend across an entire sedimentary basin. These thicker sheets are often so distinctive that they are named and mapped as separate geological units representing significant changes in the local environment at the time they were deposited.

FIG 18. Example of a cross-section through the crust, showing how a divergent movement pattern (A) may be modified by later convergent movements (B and C).

Vertical crustal movements due to loading or unloading

In addition to the direct raising or lowering of the surface by erosion or deposition, there is a secondary effect due to the unloading or loading of the crust that may take some thousands of years to produce significant effects. As mentioned above, we can visualise the lithosphere as ‘floating’ on the asthenosphere like a boat floating in water. Loading or unloading the surface of the Earth by deposition or erosion will therefore lower or raise the scenery, just as a boat will sit lower or higher in the water depending on its load.

An example of such loading has been the build-up of ice sheets during the Ice Age. The weight of these build-ups depressed the Earth’s surface in the areas involved, and when the ice melted the Earth’s surface rose again. Western Scotland provides an example of an area that has been rising because of ‘rebound’ since the ice of the Ice Age disappeared on melting.

A second example of this is the lowering of the area around the Mississippi Delta, loaded by sediment eroded from the more central and northern parts of North America. The Delta region, including New Orleans, is doomed to sink continually as the Mississippi River deposits sediment around its mouth, increasing the crustal load there.

Conversely, unloading of the Earth’s surface will cause it to rise. Recent theoretical work on the River Severn suggests that unloading of the crust by erosion may have played a role in raising the Cotswold Hills to the east and an equivalent range of hills in the Welsh Borders.

Vertical movements due to thermal expansion or contraction

Changes in the temperature of the crust and lithosphere are an inevitable result of many of the processes active within the Earth, because they often involve the transfer of heat. In particular, rising plumes of hot material in the Earth’s mantle, often independent of the plate boundaries, are now widely recognised as an explanation for various areas of intense volcanic activity (for example beneath Iceland today). These plumes are often referred to as ‘hot spots’ (Fig. 15). Heating and cooling leads to expansion or contraction of the lithosphere and can cause the surface to rise or sink, at least locally.

An example of this is the way that Britain was tilted downwards to the east about 60 million years ago. At about this time, eastern North America moved away from western Europe as the North American and Eurasian plates diverged. The divergence resulted in large volumes of hot material from deep within the Earth being brought to the surface and added to the crust of western Britain. It is believed that the heating and expansion of the crustal rocks in the west has elevated them above the rocks to the east, giving an eastward tilt to the rock layers and exposing the oldest rocks in the west and the youngest ones in the east.

THE CHALLENGE OF MEASURING CRUSTAL MOVEMENTS

Having just reviewed some of the processes that may cause movements of the Earth’s surface, it is useful to consider the practical difficulties of how such movements are measured.

For present-day applications, it seems natural to regard sea level as a datum against which vertical landscape movements can be measured, as long as we remember to allow for tidal and storm variations. However, much work has demonstrated that global sea level has changed rapidly and frequently through time, due to climate fluctuations affecting the size of the polar icecaps and changing the total amount of liquid water present in the oceans and seas (see Chapter 5). It has also been shown that plate tectonic movements can have an important effect on global sea level by changing the size and shape of ocean basins.

Attempts have been made to develop charts showing how sea level, generalised for the whole world, has varied through time. However, it has proved very difficult to distinguish a worldwide signal from local variations, and the dating of the changes is often too uncertain to allow confident correlation between areas.

In sedimentary basins, estimates of vertical movements have been made using the thicknesses of sediment layers accumulating over different time intervals in different depths of water. In areas of mountain building, amounts of vertical uplift have been estimated using certain indicator minerals that show the rates of cooling that rocks have experienced as they were brought up to the surface. However, both these approaches are only really possible in areas that have been subjected to movements of the Earth’s crust that are large and continuous enough to dominate completely other possible sources of error.

Local movements are also difficult to estimate, although fold and/or fault patterns may allow a simple measure in some cases. Over short present-day periods of time it has proved to be possible to detect vertical movement patterns using satellite imagery. Movement of sediment across the Earth’s surface by rivers or sea currents can be estimated if mineral grains in the sediment can be tracked back to the areas from which they have come. In the detailed consideration of landscapes in this book, we have to rely on using the widest possible range of types of evidence, carefully distinguishing the times and scales involved. Even then, we are often left with probable movement suggestions rather than certainties.