Chapter 4 Volcanism
As noted in the previous section, the types of magma produced in the various volcanic settings can differ significantly. At divergent boundaries and oceanic mantle plumes, where there is little interaction with crustal materials the magma tends to be consistently mafic. At subduction zones, where the magma ascends through significant thicknesses of crust, interaction between the magma and the crustal rock—some of which is quite felsic—leads to increases in the felsic character of the magma.
As illustrated in Figure 4.2.1, several processes can make magma that is stored in a chamber within the crust more felsic than it was to begin with, and can also contribute to development of vertical zonation from more mafic at the bottom to more felsic at the top. Partial melting of country rock and country-rock xenoliths increases the overall felsic character of the magma; first, because the country rocks tends to be more felsic than the magma, and second, because the more felsic components of the country rock melt preferentially. Settling of ferromagnesian crystals from the upper part of the magma, and possible remelting of those crystals in the lower part can both contribute to the vertical zonation from relatively mafic at the bottom to more felsic at the top.
From the perspective of volcanism there are some important differences between felsic and mafic magmas. First, as we’ve already discussed, felsic magmas tend to be more viscous because they have more silica, and hence more polymerization. Second, felsic magmas tend to have higher levels of volatiles; that is, components that behave as gases during volcanic eruptions. The most abundant volatile in magma is water (H2O), followed typically by carbon dioxide (CO2), and then by sulphur dioxide (SO2).
The general relationship between the SiO2 content of magma and the amount of volatiles is shown in Figure 4.2.2. Although there are many exceptions to this trend, mafic magmas typically have 1% to 3% volatiles, intermediate magmas have 3% to 4% volatiles, and felsic magmas have 4% to 7% volatiles.
Differences in viscosity and volatile levels have significant implications for the nature of volcanic eruptions. When magma is deep beneath the surface and under high pressure from the surrounding rocks, the gases remain dissolved. As magma approaches the surface, the pressure exerted on it decreases. Gas bubbles start to form, and the more gas there is in the magma, the more bubbles form. If the magma is runny enough for gases to rise up through it and escape to surface, the pressure will not become excessive. Assuming that it can break through to the surface, the magma will flow out relatively gently. An eruption that involves a steady non-violent flow of magma is called .
A good analogy for a magma chamber in the upper crust is a plastic bottle of pop on the supermarket shelf. Go to a supermarket and pick one up off the shelf (something not too dark). You’ll find that the bottle is hard because it was bottled under pressure, and you should be able to see that there are no gas bubbles inside.
Buy a small bottle of pop (you don’t have to drink it!) and open it. The bottle will become soft because the pressure is released, and small bubbles will start forming. If you put the lid back on and shake the bottle (best to do this outside!), you’ll enhance the processes of bubble formation, and when you open the lid, the pop will come gushing out, just like an explosive volcanic eruption.
A pop bottle is a better analogue for a volcano than the old baking soda and vinegar experiment that you did in elementary school, because pop bottles—like volcanoes—come pre-charged with gas pressure. All we need to do is release the confining pressure and the gases come bubbling out, forcing the pop with them.
If the magma is felsic, and therefore too viscous for gases to escape easily, or if it has a particularly high gas content, it is likely to be under high pressure. Viscous magma doesn’t flow easily, so even if there is a conduit for it to move towards surface, it may not flow out. Under these circumstances pressure will continue to build as more magma moves up from beneath and gases continue to exsolve. Eventually some part of the volcano will break and then all of that pent-up pressure will lead to an explosive eruption.
Mantle plume and spreading-ridge magmas tend to be consistently mafic, so effusive eruptions are the norm. At subduction zones, the average magma composition is likely to be close to intermediate, but as we’ve seen, magma chambers can become zoned and so compositions ranging from felsic to mafic are possible—even likely. Eruption styles can be correspondingly variable.
- Loss of olivine or pyroxene (by crystal settling) makes the upper magma more felsic.
- Partial melting of country rock makes the magma more felsic.
- Partial or complete melting of xenoliths makes the magma more felsic.
- Possible re-melting of olivine or pyroxene can make the lower magma more mafic.
- Figure 4.2.1: © Steven Earle. CC BY.
- Figure 4.2.2: Original image by Schminke, 2004. (Schminke, H-U., 2004, Volcanism, Springer-Verlag, Heidelberg). Modified by Steven Earle.
- Figure 4.2.3: Champagne uncorking photographed with a high speed air-gap flash © Niels Noordhoek. CC BY-SA.
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 10.4.3). 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 18.0.1), 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 18.2.1.
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 18.2.2. 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 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 18.2.2), 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 centimetres per year (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).
This map shows the magnetic patterns on the Juan de Fuca plate. The coloured bands represent periods of normal magnetism, while the white bands represent reversed magnetism. A magnetic-reversal time scale is also shown.
- How old is the oldest part of the Juan de Fuca Plate that is subducting along the Cascadia subduction boundary?
- How old is the youngest part of the Juan de Fuca Plate that is subducting?
The magnetic patterns and chronology shown here have been colour-coded to make them easy to interpret, but on most such maps the magnetic patterns are shown only as black and white stripes, making it much more difficult to interpret the ages of the sea floor. Magnetic-reversal patterns that have no context (such as the 0 age along the spreading ridge in this case) are very difficult to interpret.
As is evident from Figures 18.1.1 and 18.1.2, 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 now more than a year old (last eruption was April 30th 2018). 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 kilometres 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 18.2.4.
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.
Figure 18.2.3 image description: The Juan de Fuca plate lies between the Pacific Plate and the North America Plate along the west coast of Vancouver Island and Washington State. The Juan de Fuca Plate is subducting under the North America Plate along the Cascadia subduction boundary. The Juan de Fuca Plate is youngest along the Juan de Fuca ridge at the Pacific Plate and is older as it moves east. The magnetic time scales shows periods of magnetic reversal, and the ages of the parts of the Juan de Fuca plate that are subducting range from just over 0 Ma in the northwest corner of the plate to over 8 Ma in the southeast corner of the plate. [Return to Figure 18.2.3]
Figure 18.2.4 image description: A volcanic island forms and a fringing reef develops around it in the water. It becomes a barrier reef as the volcanic island subsides and water is able to pool between the island and the reef to form a lagoon. An atoll is formed when the volcano subsides enough that it no longer breaches the ocean surface but the reef remains to form a pool. [Return to Figure 18.2.4]
- Figure 18.8: "Age of oceanic lithosphere" © National Oceanic and Atmospheric Administration. Adapted by Steven Earle. CC BY-SA.