World Organization of Volcano Observatories (WOVO)
The World Organization of Volcano Observatories was established as the result of a meeting of representatives from world-wide volcano observatories, held in Guadeloupe in 1981. WOVO became International Association of Volcanology and Chemistry of the Earth’s Interior Commission in the following year.
The principal aims of the World Organization of Volcano Observatories are:
1. To stimulate cooperation between scientists working in observatories and to create or improve ties between observatories and institutions directly involved in volcano monitoring.
2. To facilitate an exchange of views and experience in volcano monitoring by convening periodic meetings including field-based ones, by periodic newsletters, and by promoting a specific e-mail observatories network service.
3. To maintain an up-to-date inventory (Directory of Volcano Observatories) of networks and of instrumentation and manpower that could be made available to any of the member institutions if a situation arises that requires scientific support.
4. To supply technical support at observatories in developing countries and to create a WOVO fund to be used by these observatories (activities of WOVO generally speaking are focused on dangerous volcanoes in developing countries).
5. To organise an international task force and to promote funding from international organisations that could help defray travel and related expenses of scientific support teams.

Volcano Monitoring Techniques
Geologists have developed several methods to monitor changes in active volcanoes. These methods allow geologists to forecast and, in some cases, predict, the onset of an eruption. A forecast indicates that the volcano is “ready” to erupt. A prediction states that a volcano will erupt within a specified number of hours or days (Wright and Pierson, 1992). Several methods developed by and currently used at the Hawaiian Volcano Observatory are introduced in the following paragraphs. The Teaching Suggestions provide additional description and activities.

How do Volcanologists predict volcanic eruptions?
The prediction of volcanic eruptions is difficult because, to be of practical use, they must be made before eruptions! Its a lot easier to see patterns in monitoring data after an eruption has occurred. But great progress has been made because of the lessons learned over many years at Kilauea volcano in Hawaii, and applied and modified at Mt. St. Helens before and during in eruption sequence in the 1980s.
Meaningful prediction requires careful monitoring of a volcano’s vital signs. Seismometers can be used to pinpoint earthquakes which track the rise of magma and its movement along fissures. Measurements of the tilt of the entire mountain provide additional information about the “breathing” of the volcano as magma moves inside it. Instruments that sniff SO2, CO2 and other gases also can signal changes in the volcano. At some volcanoes the seismic information seems most reliable, at others the tilt tells the story. But the best predictions come from the combination of all of these methods into a volcano monitoring and prediction system.
You must remember that each volcano is unique. The pattern of events that signifies an eruption at one volcano may not occur before an eruption at a different volcano. And the same volcano may change its eruptive behavior at any time! The good news is that general trends in precursor behavior our being observed at a variety of volcanoes around the world so that volcanologists are getting better at predicting eruptions.
How does a thermocouple measure the temperature of lava without melting?
A thermocouple works on the principle that the electrical resistance at the point where two wires of different composition join, is very sensitive to the temperature. So…a thermocouple consists of two wires joined to an electrical source. Current passes through the wires and they only touch eachother out at the tip of the thermocouple probe. The temperature of that tip controls the resistance of the current passing through the place where they join, and this resistance can be measured and converted to temperature information. These little wires are usually sheathed in a ceramic insulation with an outer sheath of stainless steel. The melting points of stainless steel, ceramic, and the two wires are all higher than the temperature of lava, so none of them melt.


There are three types of lava and lava flows: pillow, pahoehoe, and aa. Pillow lavas are volumetrically the most abundant type because they are erupted at mid-ocean ridges and because they make up the submarine portion of seamounts and large intraplate volcanoes, like the Hawaii-Emperor seamount chain. Pahoehoe is the second most abundant type of lava flow.
Eruptions under water or ice produce pillow lava. James Moore of the U.S. Geological Survey made the first underwater observations of the formation of pillow lavas (Moore and others, 1973). Moore and his coworkers studied lava from the Mauna Ulu eruption They described pillow lavas as elongate, interconnected flow lobes that are elliptical or circular in cross-section (Moore, 1975). This photo shows pillow lava forming off the south coast of Kilauea volcano, Hawaii. Photograph by Gordon Tribble and courtesy of U.S. Geological Survey.
Lava flows near the coast tend to spread out laterally and enter the ocean over a large front. Several lava tubes may be active across the flow front. Larger lava tubes feed down slope and maintain pressure in the growing lobes that extend into the water. Flow lobes develop most readily on steep slopes and result in stacks of pillow “tongues” that parallel the offshore slope. This offshore stack of lava produces beds of rubble parallel to the offshore slope. Geologists call these steeply dipping beds foreset beds.

On November 8, 1992, lava entered the ocean just east of the archeological sites at Kamoamoa on the south flank of Kilauea volcano. By the end of the month, nearly all of Kamoamoa was buried under lava. By May 1993, a lava delta had extended the coastline about 1,600 feet (500 m) to the south. This aerial view is to the west. Photograph by Christina Heliker, U.S. Geological Survey, May 12, 1993.
The foreset beds are part of a lava delta that grows towards the ocean and provides a platform on which subaerial lavas can extend, thus adding new land to the island. However, some lobes are not continuous and break apart to form pillow sacks, thus adding rubble to the flank of the volcano. The active front of a lava delta is called a bench. Like the delta, the bench is built on rubble. However, the bench is relatively unstable and can collapse, falling into the ocean (Mattox, 1993).
A bench collapsed at the Lae Apuki entry in April 1993, removing a block 690 feet (210 m) long, 46 feet (14 m) wide, and 26 feet (8 m) thick from the front of the delta. A visitor, standing on the bench at the time of the collapse, disappeared into the ocean. This aerial view shows an active bench adding new land out from the old sea cliff (bottom of photo). Dark areas are black sand. Photograph by Christina Heliker, U.S. Geological Survey, August 25, 1994.
Pillow lavas are also found near the summit of Mauna Kea These pillow lavas were produced by a subglacial eruption that occurred 10,000 years ago. The pillow is about 3 feet (1 m) in diameter and has a glassy rim. Figure 21.11 from Porter, 1987.
Pillow lavas can also form when flows enter a river or lake. The pillows in the photo formed in the Wailuku River above Hilo, Hawaii about 3,500 years ago. The round cobbles were transported by the river. Photograph by Jack Lockwood, U.S. Geological Survey, June 14, 1982.
Pahoehoe lava is characterized by a smooth, billowy, or ropy surface. A ropy surface develops when a thin skin of cooler lava at the surface of the flow is pushed into folds by the faster moving, fluid lava just below the surface.
Pahoehoe flows can evolve into lava tubes. One way that tubes form is by the crusting over of channelized lava flows. As the crust on a flow becomes thicker, it insulates the lava in the interior of the flow. The lava drains down slope and feeds the advancing front or flows into the ocean. When the eruption stops or the vent is abandoned, the lava drains from the tube. Thurston lava tube is an excellent example of a lava tube. The rock surrounding a lava tube serves as a insulator to keep the lava hot and fluid. Because the lava remains hot, it can travel great distances from the vent. For example, tube-fed pahoehoe lava traveled 7.3 miles (11.7 km) from the Kupaianaha vent to the village of Kalapana (Mattox and others, 1993). Photo of volcanologist looking through a skylight to see inside of a lava tube. Photograph by J.D. Griggs, U.S. Geological Survey, February 2, 1989.
Pahoehoe flows tend to be relatively thin, from a few inches to a few feet thick. Road cuts and craters expose stacks of lava flows that make the volcano.
Aa is characterized by a rough, jagged, spinose, and generally clinkery surface. Aa flows advance much like the tread of a bulldozer. This photo is looking across an aa channel. Note the character of the aa that makes the wall of the channel. Photo by Steve Mattox, October 12, 1990.
The interior of the flow is molten and several feet thick. The molten core grades upwards and downwards into rough clinkers. As the flow advances, clinker on the surface is carried forward relative to the molten interior. The clinkers continue to move forward until they roll down the steep front. The clinkers are then overridden by the molten core. Aa lava flows tend to be relatively thick compared to pahoehoe flows. Aa flows from the 1984 Mauna Loa eruption ranged from 6 to 23 feet (2-7 m) in thickness.
During the early episodes of the current eruption, aa flows up to 36 feet (11 m) thick surged through the Royal Gardens subdivision at rates as great as 108 ft/min (33 m/min)(Neal and Decker, 1983). The molten core of the flow is exposed. Note vehicles for scale. Photograph by R.W. Decker, U.S. Geological Survey, July 2, 1983.
Since the mid-1800s, geologists have tried to explain the causes of pahoehoe and aa lava. Several factors were offered to explain the transition: impediment of flow by obstacles, flowage during cooling, quantity of lava, conditions under the flow, and deep cooling. In the last few decades, with careful observations of numerous lava flows, geologists have reached a better understanding of the transition of pahoehoe to aa. One important influence is the viscosity of the lava. Viscosity is the resistance of a fluid to flow. For example, molasses has a higher viscosity that water. The following paragraphs outline some of the more important observations on the transition of pahoehoe to aa.

How is lava formed?
First, there is a definition we need to make. Just to keep things straight, geologists use the word “magma” for molten rock that is still underground, and the word “lava” when it has reached the surface.
Rocks in the mantle and the core are still hot from the formation of the Earth about 4.6 billion years ago. When the Earth formed, material collided at high speeds. These collisions generated heat (try clapping your hands together – they get hot) that heat became trapped in the Earth. There is also heat within the earth produced by radioactive decay of naturally-occurring radioactive elements. It is the same process that allows a nuclear reactor to generate heat, but in the earth, the radioactive material is much less concentrated. However, because the earth is so much bigger than a nuclear power plant it can produce a lot of heat. Rocks are good insulators so the heat has been slow to dissipate.
This heat is enough to partially melt some rocks in the upper mantle, about 50-100 km below the surface. We say partially melt because the rocks don’t completely melt. Most rocks are made up of more than one mineral, and these different minerals have different melting temperatures. This means that when the rock starts to melt, some of the minerals get melted to a much greater degree than others. The main reason this is important is that the liquid (magma) that is generated is not just the molten equivalent of the starting rock, but something different.
You could think of making a “rock” out of sugar, butter, and shaved ice. Pretend that they are mixed equally so that your rock is 1/3 sugar, 1/3 butter, and 1/3 shave ice. If you start melting this “rock”, however, the “magma” that is generated will be highly concentrated in the things that melt more easily, namely the ice (now water) and butter. There will be a little bit of molten sugar in your magma, but not much, most of it will still be crystalline.
The most common type of magma produced is basalt (the stuff that is erupted at mid-ocean ridges to make up the ocean floors, as well as the stuff that is erupted in Hawai’i). Soon after they’re formed, little drops of basaltic magma start to work their way upward (their density is slightly less than that of the solid rock), and pretty soon they join with other drops and eventually there is a good flow of basaltic magma towards the surface. If it makes it to the surface it will erupt as basaltic lava.
What is lava made of?
Lava is made up of crystals, volcanic glass, and bubbles(volcanic gases). As magma gets closer to the surface and cools, it begins to crystallize minerals like olivine and form bubbles of volcanic gases. When lava erupts it is made up of a slush of crystals, liquid, and bubbles. The liquid “freezes” to form volcanic glass.
Chemically lava is made of the elements silicon, oxygen, aluminum, iron, magnesium, calcium, sodium, potassium, phosphorus, and titanium (plus other elements in very small concentrations. Have a look at the background information in Minerals, Magma, and Volcanic Rocks.

Most rocks are made of the following elements:
Element (Symbol)
Weight percent
Oxygen (O) 46.6
Silicon (Si) 27.7
Aluminum (Al) 8.1
Iron (Fe) 5.0
Calcium (Ca) 3.6
Sodium (Na) 2.8
Potassium (K) 2.6
Magnesium (Mg) 2.1
Total: 98.5
The amount of these different elements in a rock can be determined. Geologists report the values as oxides (the element combined with oxygen). For example, a rock might have 6 weight percent FeO (iron oxide).
Different types of lava have different chemical compositions.
How does lava change its composition?
Geochemists use several elements to track the history of a batch of magma. For our example we’ll start with magnesium (and think of it as magnesium oxide, MgO, like the geochemists do). Beneath Hawaii, when the mantle melts it produces a magma with about 17 weight percent MgO. Volcanologists call magma with this composition a picrite. As the magma rises crystals begin to form. The first crystal to form is olivine, which contains the elements magnesium, silicon, and oxygen in the proportions 2 to 1 to 4. The crystal is more dense than the surrounding magma and it begins to settle. It is estimated that an than the surrounding magma and it begins to settle. It is estimated that an olivine crystal 1 mm in diameter falls through the magma at a rate of 20 meters per hour. By settling the crystals are being removed from the magma, causing the chemical composition of system to change. As more and more olivine crystals settle, the magma has less and less magnesium oxide and more and more silica. For most volcanoes, by the time the original magma from the mantle is erupted its composition has changed from a picrite to a basalt.
If the magma is not erupted, the same process continues but more and different minerals become involved. In general, the removal of these crystals drives the composition away from that of a basalt and towards that of a rhyolite. The amount of magnesium oxide and iron oxide continues to decrease and the amount of silica, sodium oxide, and potassium oxide increases.
Other processes can become involved. The magma may melt and incorporate crustal rocks that tend to contain more silica. This drives the composition towards rhyolite. The rising magma may also intersect and mix with a magma that has evolved differently. This will also change its composition.
These processes (crystal settling, assimilation, and magma mixing) influence the composition of the magma. Geochemists study the minerals and the chemistry of the lava to determine which processes were involved and how big of a role each one played.
How hot is lava?
The temperatures of lavas vary depending on their chemical composition. Hawaiian lava (basalt) is usually around 1100 C. Volcanoes such as Mt. St. Helens erupt lava that are around 800 C.
Here is a list of temperatures for the common types of lava:
Rock type
Temperature (C) Temperature (F)
Rhyolite 700-900 1292-1652
Dacite 800-1100 1472-2012
Andesite 950-1200 1742-2192
Basalt 1000-1250 1832-2282

How long does it take lava to cool?

Lava cools very quickly at first and forms a thin crust that insulates the interior of the lava flow. As a result, basaltic lava flows can form crusts that are thick enough to walk on in 10-15 minutes but the flow itself can take several months to cool! Because of the insulating properties of lava, it cools slower and slower over time. Thick stacks of lava flows (30 m or 100 ft thick) can take years to cool completely. An extreme example is a lava flow that was erupted in 1959 and partly filled a pit crater (Kilauea Iki). The “ponded” flow was about 85 meters thick (about 280 ft thick). It was drilled in 1988, and there was still some mushy, not-quite-solid stuff down near the bottom, 29 years after it erupted!

How far can lava flow?
The longest recent flow in Hawai’i was erupted from Mauna Loa in 1859. It is about 51 kilometers long from the vent to the ocean (we don’t know how much longer it went out under water, but probably not too much farther). There is a lava flow at Undara in Queensland, Australia is 100 miles (160 km) long.
How fast does lava flow?
In Hawaii the fastest flows we’ve recorded were those of the 1950 Mauna Loa eruption. These were going about 6 miles (10 kilometers) per hour through thick forest. That was the velocity of the flow front. Once the lava flows became established and good channels developed, the lava in the channels was going at more like 60 km/hour!
On January 10,1977, a lava lake at Nyiragongo drained in less than one hour. The lava erupted from fissures on the flank of the volcano and moved at speeds up to 40 miles per hour (60 km/hr). About 70 people were killed.
How close can I get to lava and will it hurt or kill me?

How close you can get depends on what kind of lava flow it is, and whether you are upwind or downwind. For example, the most approachable lava is pahoehoe. This is because each toe forms an insulating skin seconds after emerging on the surface. This skin is at first flexible and then hardens, but even when flexible it is a good insulator. This serves to keep the interior of an active pahoehoe toe hot and fluid but also prevents you from getting burned by the radiant heat. If the wind is at your back, you can easily approach long enough and close enough to get a sample with a hammer. It is still hot, and unless yo! are well-protected you can only be that close for a minute or so. You also notice that as soon as you peel the skin off to get at the molten interior, th! heat goes way up. This is heat that you can’t stand, you have to get back otherwise blisters start to form. It is hot enough that you can’t accidentally step on active lava.
Skylights into lava tubes on pahoehoe flows are quite hot, and have to be approached from upwind. They are so hot that the air shimmers over them so they are hard to miss. They are dangerous not as much because of the radiant heat from the lava inside but because of the super-heated plume of air coming out. You have to be really careful that the wind doesn’t shift, and many a volcanologist has gotten singed skin and hair when the wind changed.
An ‘a’a flow is terrible to work near. Instead of a relatively continuous skin! ‘a’a flows have discontinuous layers of clinker, and a huge amount of radiant heat escapes from between the clinker. ‘A’a flows also move faster so you really have to be quick on your feet if you want a sample. Additionally, ‘a’a flows tend to form open channels rather than lava tubes. The channels can sometimes have completely incandescent surfaces because they are flowing so fast that any skin that forms is immediately torn or sunk. We remember once flying over a large channel in a helicopter. We must have been at least 200-40! meters above the flow, but as soon as we were over the channel we could immediately feel the radiant heat through the windows! We would bet that nobod! has been downwind of an active ‘a’a channel.
Lava won’t kill you if it briefly touches you. You would get a nasty burn, but unless you fell in and couldn’t get out, you wouldn’t die. With prolonged contact, the amount of lava “coverage” and the length of time it was in contac! with your skin would be important factors in how severe your injuries would be! The health of the individual, the amount of time before care can be given and the quality of that care would also be important. In fact there have been 2 cases at the Hawaiian Volcano Observatory where a geologist fell into lava. Fortunately in both instances the lava was not very deep and they were able to get out quickly. Both ended up in the hospital and it was a scary and painful experience. Both recovered fine. Blong (1984) points out that little research has been done on injuries caused by lava. People have been killed by very fast moving lava flows. A recent example was the 1977 eruption at Nyiragongo.

Why is lava different colors?
The color of lava depends on its temperature. It starts out bright orange (1000-1150 C). As it cools the color changes to bright red (800-1000 C), then do dark red (650-800 C), and to brownish red (500-650 C). Solid lava is black (but can still be very hot).
It is the same reason that electric stove elements turn colors. First red then orange when turned on high (don’t do that experiment by yourself). As things get hotter they start to glow, first red, then orange, then yellow, then green, then blue, then violet. The sun’s color, for example, is sort of yellow-green, but you can’t tell by looking at it (and remember, you should NEVER look at the sun).
Why is lava so hot?
Lava is hot for two reasons:
1. It’s hot deep in the Earth (about 100 km down) where rocks melt to make magma.
2. The rock around the magma is a good insulator, so the magma doesn’t lose much heat on the way to the surface.

How do volcanoes affect people?
Painting of Mount Vesuvius
Volcanoes affect people in many ways, some are good, some are not. Some of the bad ways are that houses, buildings, roads, and fields can get covered with ash. As long as you can get the ash off (especially if it is wet), your house may not collapse, but often the people leave because of the ash and are not around to continually clean off their roofs. If the ashfall is really heavy it can make it impossible to breathe.
Lava flows are almost always too slow to run over people, but they can certainly run over houses, roads, and any other structures.
Pyroclastic flows are mixtures of hot gas and ash, and they travel very quickly down the slopes of volcanoes. They are so hot and choking that if you are caught in one it will kill you. They are also so fast (100-200 km/hour) that you cannot out-run them. If a volcano that is known for producing pyroclastic flows is looking like it may erupt soon, the best thing is for you to leave before it does.
Some of the good ways that volcanoes affect people include producing spectacular scenery, and producing very rich soils for farming.

Water vapor, the most common gas released by volcanoes, causes few problems. Sulfur dioxide, carbon dioxide and hydrogen are released in smaller amounts. Carbon monoxide, hydrogen sulfide, and hydrogen fluoride are also released but typically less than 1 percent by volume.
Gases pose the greatest hazard close to the vent where concentrations are greatest. Away from the vent the gases quickly become diluted by air. For most people even a brief visit to a vent is not a health hazard. However, it can be dangerous for people with respiratory problems.
The continuous eruption at Kilauea presents some new problems. Long term exposure to volcanic fumes may aggravate existing respiratory problems. It may also cause headaches and fatigue in regularly healthy people. The gases also limit visibility, especially on the leeward side of the island where they become trapped by atmospheric conditions.
The biggest eruption
The 1815 explosive eruption of Tambora volcano in Indonesia and the subsequent caldera collapse produced 9.5 cubic miles (40 cubic kilometers) of ash. The eruption killed 10,000 people. An additional 80,000 people died from crop loss and famine.

To put it mildly, ash is bad for jet aircraft engines. Apparently the problem is much more severe for modern jet engines which burn hotter than the older ones. Parts of these engines operate at temperatures that are high enough to melt ash that is ingested. Essentially you end up with tiny blobs of lava inside the engine. This is then forced back into other parts where the temperatures are lower and the stuff solidifies. As you can imagine this is pretty bad. One problem that I heard about is that pilots start losing power and apply the throttle, causing the engine to be even hotter and melt more ash.
Added to this is the fact that ash is actually tiny particles of glass plus small mineral shards–pretty abrasive stuff. You can imagine that dumping a whole bunch of abrasive powder into a jet engine is not good for the engine. This has been a pretty non-scientific explanation of the problem. I just found an article that describes the problem a little more technically.
“The ash erodes sharp blades in the compressor, reducing its efficiency. The ash melts in the combustion chamber to form molten glass. The ash then solidifies on turbine blades, blocking air flow and causing the engine to stall.

Safe distance
The distance you have to evacuate depends entirely on what kind of eruption is going on. For example, Pinatubo, one of the largest recent eruptions sent pyroclastic flows at least 18 km down its flanks, and pumice falls were hot and heavy even beyond that. For example, pumice 7 cm across fell at Clark Air base which is 25 km from the volcano! A 7 cm pumice won’t necessarily kill you but it does mean that there is a lot of pumice falling, and if you don’t get out and continuously sweep off your roof it may fall in and you’ll get squashed.
On the other hand, the current eruption at Ruapehu is relatively small. In fact, there were skiers up on the slopes when the eruptions commenced, and even though they were only 1-2 km from the vent they managed to escape. The volcanologists routinely go up on the higher slopes of Ruapehu during these ongoing eruptions to collect ash and take photographs.
So you see, you need to know something about what you think the volcano is going to do before you decide how far to run away. I guess if you have no idea of what the volcano is planning, and have no idea of what it has done in the past, you might want to be at least 25-30 km away, make sure you have a good escape route to get even farther away if necessary, and by all means stay out of low-lying areas!

Cities and Towns
Mount Mayon , in the Philippines, is a classic example of a stratovolcano. Photograph copyrighted and provided by Steve O’Meara of Nature.Stock.
The effect an eruption will have on a nearby city could vary from none at all to catastrophic. For example, atmospheric conditions might carry ash away from the city or topography might direct lahars and pyroclastic flows to unpopulated areas. In contrast, under certain atmospheric, eruption and/or topographic conditions, lahars, pyroclastic flows, and/or ash fall could enter the city causing death and destruction.
This scenario brings up several interesting problems. How do you evacuate a large population if there is little warning before the eruption? Where do these people go? If an eruption is highly likely yet hasn’t happened yet how long can people be kept away from their homes and businesses?
I should point out that in most volcanic crises geologists advise local civil defense authorities. The civil defense authorities decide what to do concerning evacuations, etc.
The IAVCEI has a program to promote research on “Decade” Volcanoes. Decade volcanoes are likely to erupt in the near future and are near large population centers. Mount Rainier in Washington and Mauna Loa in Hawaii are two Decade volcanoes in the U.S. Other Decade volcanoes include Santa Maria, Stromboli, Pinatubo, and Unzen.
What happens to the towns around a volcano when it erupts depends on many things. It depends of the size and type of eruption and the size and location of the town. A few examples might help. The 1984 eruption of Mauna Loa in Hawaii sent lava towards Hilo but the eruption stopped before the flows reached the town. The 1973 eruption of Heimaey in Iceland buried much of the nearby town of Heimaey under lava and cinder. The 1960 eruption of Kilauea in Hawaii buried all of the nearby town of Kapoho under lava and cinder. In 1980, ash from Mount St. Helens fell on many towns in Washington and Oregon. The 1902 eruption of Mount Pelee on the island of Martinique destroyed the town of Saint Pierre with pyroclastic flows. In 1985, the town of Armero was partially buried by lahars generated on Ruiz. For more examples see Decker and Decker (1989).

How do volcanoes affect the atmosphere and climate?

There are two things to think about. The first is how the weather near an erupting volcano is being affected. The second is how large eruptions will affect the weather/climate around the world. I think more people are worried about the second issue than the first.
As far as I know, the main effect on weather right near a volcano is that there is often a lot of rain, lightning, and thunder during an eruption. This is because all the ash particles that are thrown up into the atmosphere are good at attracting/collecting water droplets. We don’t quite know how the lightning is caused but it probably involves the particles moving through the air and separating positively and negatively charged particles.
Another problem that we are having here in Hawai’i involves the formation of vog, or volcanic fog. The ongoing eruption is very quiet, with lava flowing through lava tubes and then into the ocean. Up at the vent is an almost constant plume of volcanic fume that contains a lot of sulfur dioxide. This SO2 combines with water in the atmosphere to form sulfuric acid droplets that get carried in the trade winds around to the leeward side of the Big Island. The air quality there has been really poor since the eruption started in 1983 and they are getting pretty tired of it.
As for the world-wide affects of volcanic eruptions this only happens when there are large explosive eruptions that throw material into the stratosphere. If it only gets into the troposphere it gets flushed out by rain. The effects on the climate haven’t been completely figured out. It seems to depend on the size of the particles (again mostly droplets of sulfuric acid). If they are big then they let sunlight in but don’t let heat radiated from the Earth’s surface out, and the net result is a warmer Earth (the famous Greenhouse effect). If the particles are smaller than about 2 microns then they block some of the incoming energy from the Sun and the Earth cools off a little. That seems to have been the effect of the Pinatubo eruption where about a 1/2 degree of cooling was noticed around the world. Of course that doesn’t just mean that things are cooler, but there are all kinds of effects on the wind circulation and where storms occur. Some folks think that large eruptions can cause the weather phenomena called “El Nino” to start. This is a huge disruption of the Earth’s atmospheric circulation. The connection hasn’t been accepted by everybody though.
An even more controversial connection involves whether or not volcanic activity on the East Pacific Rise (a mid-ocean spreading center) can cause warmer water at the surface of the East Pacific, and in that way generate an El Nino. Dr. Dan Walker here at the University of Hawai’i has noticed a strong correlation between seismic activity on the East Pacific Rise (which he presumes indicates an eruption) and El Nino cycles over the past ~25 years.

How do volcanoes affect plants and animals?

Lava flows covering the Kamoamoa area of Hawai`i Volcanoes National Park. Photograph by Steve Mattox, November 14, 1992.
Plants are destroyed over a wide area, during an eruption. The good thing is that volcanic soil is very rich, so once everything cools off, plants can make a big comeback!
Livestock and other mammals have been killed by lava flows, pyroclastic flows, tephra falls, atmospheric effects, gases, and tsunami. They can also die from famine, forest fires, and earthquakes caused by or related to eruptions.
Mount St. Helens provides an example. The Washington Department of Game estimated that 11,000 hares, 6,000 deer, 5,200 elk, 1,400 coyotes, 300 bobcats, 200 black bears, and 15 mountain lions died from the pyroclastic flows of the 1980 eruption.
Aquatic life can be affected by an increase in acidity, increased turbidity, change in temperature, and/or change in food supply. These factors can damage or kill fish.
Eruptions can influence bird migration, roosting, flying ability, and feeding activity.
The impact of eruptions on insects depends on the size of the eruption and the stage of growth of the insect. For example, ash can be very abrasive to wings.

How quickly do plants begin to grow back?
The answer is that it depends on how much rain falls in the particular area. For example, on the rainy side of the island of Hawai’i, flows that are only 2 years old already have ferns and small trees growing on them. Probably in 10 years they’ll be covered by a low forest. On the dry side of Hawai’i there are flows a couple hundred years old with hardly a tuft of grass in sight. This means that when you are looking at old lava flows and trying to determine how old they are based on the amount of vegetation, you have to take the climate into effect as well.

What are some good things that volcanoes do?
Aso viewed from the visitors center. Small plume above Aso during a period of mild Strombolian eruptions, December 30, 1991. Photograph by Mike Lyvers.
That’s a good question. I guess the main good effect that volcanoes have on the environment is to provide nutrients to the surrounding soil. Volcanic ash often contains minerals that are beneficial to plants, and if it is very fine ash it is able to break down quickly and get mixed into the soil.
Perhaps the best place to look for more information about this would be to look up references about some of the countries where lots of people live in close proximity to volcanoes and make use of the rich soils on volcanic flanks. These would include Indonesia, The Philippines, Japan, Italy, etc.
I suppose another benefit might be the fact that volcanic slopes are often rather inaccessible, especially if they are steep. Thus they can provide refuges for rare plants and animals from the ravages of humans and livestock.
Finally, on a very fundamental scale, volcanic gases are the source of all the water (and most of the atmosphere) that we have today. The process of adding to the water and atmosphere is pretty slow, but if it hadn’t been going on for the past 4.5 billion years or so we’d be pretty miserable.

Volcanoes have done wonderful things for the Earth. They helped cool off the earth removing heat from its interior. Volcanic emissions have produced the atmosphere and the water of the oceans. Volcanoes make islands and add to the continents.
Volcanic deposits are also used as building materials. In the 1960’s Robert Bates published Geology of the Industrial Rocks and Minerals. He noted that basalt and diabase are quarried in the northeastern and northwest states. Most of the basalt and diabase is used for crushed stone: concrete aggregate, road metal, railroad ballast, roofing granules, and riprap. High-denisity basalt and diabase aggregate is used in the concrete shields of nuclear reactors. Some diabase is used for dimension stone (“black granite”).
Pumice, volcanic ash, and perlite are mined in the west. Pumice and volcanic ash are used as abrasives, mostly in hand soaps and household cleaners. The finest grades are used to finish silverware, polish metal parts before electroplating, and for woodworking. Bates reports that in ancient Rome lime and volcanic ash were mixed to make cement. In modern times pumice and volcanic ash have been used to make cement for major construction projects (dams) in California and Oklahoma. Pumice and volcanic ash continue to be used as lightweight aggregate in concrete, especially precast concrete blocks. Crushed and ground pumice are also used for loose-fill insulation, filter aids, poultry litter, soil conditioner, sweeping compound, insecticide carrier, and blacktop highway dressing. Perlite is volcanic glass (made of rhyolite) that has incorporated 2-5% water. Perlite expands rapidly when heated. Perlite is used mostly as aggregate in plaster. Some perlite is used as aggregate in concrete, especially in precast walls.

How were the volcanoes of Hawaii created?
About 95% of the world’s volcanoes are located near the boundaries of tectonic plates. The other 5% are thought to be associated with mantle plumes and hot spots. Mantle plumes are areas where heat and/or rocks in the mantle are rising towards the surface. A hot spot is the surface expression of the mantle plume. Geologists think the hot spots are stationary and the tectonic plates are mobile. The hot spot provides magma which supplies volcanoes. The movement of the plate carries the volcano off the hot spot and it gradually becomes extinct and usually subsides below sea level. A new volcano begins to form on the sea floor and grow towards the surface. Some of these volcanoes rise above sea level to make islands. The classic example is the Hawaii-Emperor volcanic chain, a line of volcanoes and seamounts that extends from Hawaii to Daikakuji, 2,200 miles (3,500 km) to the northwest. Suiko seamount is 65 million years old, the oldest seamount associated with the Hawaiian hot spot. The hot spot is probably older but any volcanoes it produced have been destroyed by subduction.
At first glance the shape and topography of the Hawaii-Emperor volcanic chain might remind you of an island arc. However, volcanoes of the Hawaii-Emperor volcanic chain are progressively younger towards Hawaii and made of basalt. Island arcs are produced at convergent plate boundaries where an ocean plate is subducted beneath an adjacent ocean plate. The Aleutians of Alaska is an example. There is no age progression and andesite is the most common rock.

What is the history of Mount. St. Helens prior to the May 1980 eruption?
According to Volcanoes of the World, by Simkin and Siebert (1994, Geoscience Press, P.O. Box 42948, Tucson, AZ 85733-2948), Mt. St. Helens erupted 23 times prior to 1831, based on charcoal dates. After that (and before the 1980 eruption) it erupted in 1831, 1835, 1842, 1847, 1848, 1849, 1853, 1854, and 1857.

Eruptions began at Mount St. Helens about 40,000 years ago. The deposits are air-fall tephra and pyroclastic flows, the type of material produced by explosive eruptions. One pumiceous tephra deposit produced during this episode had a volume as great as any subsequent tephra eruption at Mount Saint Helens. I think it would qualify as a big eruption.
A recent paper suggests that eruptions at Mount Saint Helens began as long ago as 80,000 years. Berger and Busacca (1995) dated a loess (wind-blown slit-sized material) deposit just below a tephra layer in eastern Washington that is known to be from Mount St. Helens. The loess is about 80,000 years old. The tephra is thought to be slightly younger.

What happened when Mount St. Helens erupted on May 18, 1980?
Chronologies and summaries of the May 18, 1980 eruption For more information on what happen during the May 18, 1980 eruption of Mount St. Helens click here and here. The USGS’s Cascade Volcano Observatory has a very detailed account of the eruption, with a list of references.
The climatic eruption began at 08:32 PDT on May 18, 1980.

57 people were killed directly by the eruption. There was also a plane crash, a traffic accident, and shoveling ash which killed a total of 7 more.

Height of Mount St. Helens
The summit elevation was 9,760 feet (2,975 m) before the eruption. After the eruption the elevation of the new summit was about 8,525 feet (2,600 m).
About 0.25 cubic kilometers of new volcanic rock was erupted on May 18, 1980. This would make a cube about 600 meters (~2000 ft) on a side! But much more material was moved by the eruption: the entire northern side of the volcano collapsed and flowed down hill. The volume of this collapse was about 2.5 cubic kilometers – ten times bigger than the new lava.

The eruption began at 8:45 a.m. At noon, the ash plume (in the upper troposphere and lower stratosphere) had reached Moscow, Idaho. By about 3 p.m. it was near Missoula, Montana and starting to spread south. By 6 pm it was eas! of Pocatello, Idaho. At the end of the day, about 16 hours after the eruption started, the ash plume was near central Colorado.
A huge volume of ash was created by the various 1980 eruptions of Mount St. Helens. Every community affected had its own ways of dealing with the ash. Tons of ash was probably washed down storm drains and into sewer systems as people cleaned roofs and sidewalks. Local landfills received ash. Many tons of ash came down rivers and streams into the Columbia River. The river had to be dredged to allow shipping to pass, and the sand dredged from the bottom was deposited in large dikes along the Columbia. These dikes are now covered in grass and trees, but if a person was to dig down a few feet they would find the ash.
Ash can still be found on the floor of forested areas all around the mountain, and the blast zone is still heavily covered by ash. Souvenir “ash” trays and mugs made from Mount St. Helens ash can be purchased at gift shops near the mountain.
Will Mt. St. Helens erupt again?
Mt. St. Helens still emits steam pretty much every day, but it hasn’t had any eruptions since 1985. We can not say if it will erupt again this century, but studies indicate that there is still hot magma under the mountain, so it will probably erupt sometime. Fortunately, Mt. St. Helens is still carefully monitored by volcanologists, so that if or when the mountain begins to become active again, we should have early warning.

It is likely that Mt. St. Helens will have one or more small eruptions during your lifetime, but probably unlikely that it will have another big one. At the time of the 1980 eruption there was the weight of the whole volcanic cone available to keep the magma pressure from erupting. This allowed a large amount of pressure to build up and consequently the eruption was large. Now the whole top of the mountain is gone so there is a good deal less weight available as a counter balance to the pressure. This means that eruptions occur after lesser amounts of pressure have built up and the eruptions are therefore smaller. Eventually, of course, Mt. St. Helens will recover its cone-shaped shape by filling in the amphitheater with lava domes and ash. Once this occurs (probably over a period of a few hundred years) it’ll be ready to suffer another large 1980-size eruption.
Since the famous 1980 eruption there has been a lava dome growing within the caldera, and every once in a while part of the dome either collapses or is blasted away by either gas or steam explosions. These eruptions seem to be getting less and less frequent but have probably not stopped.
Mount St. Helens erupted at least 10 times in the 200 years before the 1980 eruption. It is considered to be the most active volcano in the Cascades. Volcanologists that study Mount St. Helens believe it is likely to erupt again within a few decades or a century at most.
It can be difficult to know when a volcano is dead. Many cinder cones, are monogenetic, erupting only once before the volcano is dead. Stratovolcanoes can erupt many times in only a few centuries, like Llaima in Chile, which has erupted at least 40 times in the last 350 years. Stratovolcanoes can also go many centuries without an eruption, like Newberry volcano in Oregon, which has not erupted in 1,300 years but will probably erupt again. Large volcanic systems, like Yellowstone, can go hundreds of thousands of years between eruptions.
What is the current status of Mt. St. Helens?
The U.S. Geological Survey’s Cascade Volcano Observatory reports that the number of small magnitude earthquakes (less than magnitude 1) beneath Mount St. Helens has increased slowly and steadily from less than 10 events per month in January (1995) to about 100 events per month in September (1995). They point out that the volcano has been quiet during this period and that no explosion or emission of gas or ash has occurred from the lava dome. Similar seismic activity was observed prior to and during a series of small gas explosions from the dome in 1989-1991. Although these explosions were relatively small, they did throw dome rocks 1 foot (30 cm) in diameter 0.5 miles (0.8 km) from the dome and sent ash plumes as high as 20,000 feet (6,100 m) above the sea level. Because these events can happen without warning some trails in the area have been closed.
What were the effects on people when Mt St Helens erupted?
The eruption killed 57 people, in the lateral blast, ashfall, and lahars. The causes to death included asphyxiation, thermal injuries, and trauma. Four indirect death were caused by a cropduster hitting powerlines during the ashfall, a traffic accident during poor visibilty, and two heart attacks from shoveling ash.
The Toutle River was flooded by melting snow and ice from the mountain. About 12 million board feet of stockpiled lumber were sweep in the river. Eight bridges were destroyed. 200 homes were destroyed or damaged. Debris dams were added to help control sediment in the rivers.
Thirty logging trucks, 22 transport vehicles, and 39 railcars were damaged or destroyed along with 4.7 billion board feet of timber.
Shipping was stopped on the Columbia River and some vessels were stranded. In eastern Washington, falling ash stranded 5,000 motorist. Ash had to be cleared from runways and highways.
For a limited time, some people living near the eruption suffered from post traumatic stress syndrome: depression, troubled sleep, irritability, ans a sense of powerlessness.
From 1980-1990, 74 research projects were funded by the National Science Foundation at a total cost of just less than $5 million. The Mount St. Helens Visitors Center at Castle Rock cost $5.5 million to construct. Trails, roads in the park, and interpretive centers cost another $42.3 million. New highway and bridges from the Toutle River to Johnston Ridge cost $145 million. facilities along this road will cost another $25 million.
What On June 23, 1994, a lava lake formed in Nyiragongo volcano in eastern Zaire. Activity has continued at the lava lake. Nyiragongo is 3,459 m tall and at 1.5S, 29.3E.
Erta Ale has been in continuous eruption since 1967.
Was the most recent volcanic eruption in Africa?
On June 23, 1994, a lava lake formed in Nyiragongo volcano in eastern Zaire. Activity has continued at the lava lake. Nyiragongo is 3,459 m tall and at 1.5S, 29.3E.
Erta Ale has been in continuous eruption since 1967.

Nyiragongo (1.5S, 29.3E) is associated with the East African Rift. The volcano is in Zaire, not far from the border with Rwanda, in Virunga National Park. The town of Goma is 11 miles (18 km) south of the summit of Nyiragongo and on the shore of Lake Kivu. Goma served as an encampment for nearly a million refugees from the civil war in Rwanda.
Nyiragongo (elevation: 11,365 feet, 3,465 m) is a stratovolcano. A crater at the summit of Nyiragongo contained a lava lake from 1894 to 1977. On January 10, 1977, the lava lake drained in less than one hour. The lava erupted from fissures on the flank of the volcano and moved at speeds up to 40 miles per hour (60 km/hr). About 70 people were killed. The fluid lava reached within 2,000 feet (600 m) of the Goma airport. From June 1982 to early 1982 the volcano was active with a lava lake in the crater and phreatic explosions and lava fountaining.
The most recent activity at Nyiragongo began in June of 1994. A lava lake once again filled part of the crater. The lava lake was approximately 130 feet (40 m) in diameter and sent lava flows onto the floor of the 2,600 foot (800 m) diameter crater. The surface of the lava lake was about 500 feet (150 m) below the level of the lake when it drained in 1977.
The lava at Nyiragongo is a nephelinite. Nephelinite is a type of alkaline (high concentration of alkali elements, Na and K) lava. It tends to be aphanitic (no crystals visible to the unaided eye). With a microscope crystals of the following minerals might be seen: nepheline, clinopyroxene, olivine, iron-titanium oxides, feldspar and possibly leucite.

Volcanoes in Cameroon are part of the Cameroon line, a chain of volcanoes extending from Annobon Island in the Atlantic Ocean northeastward through Cameroon. The oldest rocks have been dated at 70 million years old. Nine volcanoes along the line are active. A fissure eruption occurred at Mt. Cameroon in 1982.
Volcanism along the Cameroon line is related to rifting – where a continent breaks into two pieces. About 110 million years ago a giant rift broke apart what became Africa and South America and the South Atlantic Ocean began to form. A smaller rift formed within the African continent. This older rift, called the Benue Trough, is north of and parallel to the Cameroon line. About 80 million years ago, during a reorganization of plate boundaries, the African plate rotated counter-clockwise. Then a new rift formed that failed to split Africa but apparently did form conduits that allowed magma to ultimately reach the surface and form the volcanoes of the Cameroon line.
1986 In 1986, a cloud of gas burst out of Lake Nyos because a landslide disturbed the stratification with the lake. The stratification (layering) is caused by different amounts of CO2 dissolved in the water. The layers are stable (they don’t mix or changes position). The alternative mechanism, that the burst was caused by a phreatic eruption (involving lava into the lake), is not generally accepted.
Dr. Niels Oskarsson proposed reducing the amount of CO2 in the lake by continuously draining the bottom waters from the lake. Water from rainfall would displace an equal volume of water from the bottom of the lake. Over time this would establish low CO2 levels in the lake.
The 1987 conference recommended continuous telemetered monitoring of the deep water temperature, pH, conductivity, and alkalinity of the lake. Unfortunately, no funding or plans for implementation were developed (in 1987).

I think there is a pretty good consensus as to what happened at Lake Nyos (in 1986, by the way). A few folks disagree on the details, but the main event was the sudden release of a large cloud of carbon dioxide gas from the lake. Lake Nyos is a water-filled throat of an old volcano and it is deep and funnel-shaped. Although no longer erupting, there is still gas being released by the old plumbing system under the lake. Carbon dioxide gas was released directly into the deepest waters of the lake, where it could remain in solution (the way that carbon dioxide stays in solution in an un-opened soda or beer). In this situation the lake could build up a large amount of carbon dioxide dissolved in the deeper water. This was a stable situation. The carbon-dioxide charged water was slightly denser than the normal water in the upper levels of the lake, and the weight of the overlying water kept the carbon dioxide in solution in the deeper parts of the lake.
However, nature decided to unbalance the situation. This is where the disagreement among volcanologists comes in. It is agreed that somehow some of that carbon dioxide-rich water was displaced upward into shallower depths to the point where the overlying water pressure was lower and carbon dioxide bubbles could start to form (like when you lower the pressure on a soda by opening the bottle and suddenly bubbles start to form). At Lake Nyos, once these bubbles started to form they wanted to rise to the top, this brought up more carbon dioxide-rich water which then also started to develop bubbles, and pretty soon there was a big rush of carbon dioxide bubbles to the surface. What people don’t agree on is what the trigger for this unbalancing event was. Most people, I think, feel that there was some sort of landslide into the lake that stirred up the water. There are a few volcanologists who think there was some type of eruption in the deeper part of the lake, but they are in the minority.
Once all this carbon dioxide reached the surface, it splashed some lake water out of the lake, like a big bubble bursting. Carbon dioxide is denser than air, so it hugged the ground and flowed down the stream valley that leads away from the lake. Unfortunately many homes and at least one town are also along this valley and the inhabitants were caught by this cloud of ground-hugging gas. Carbon dioxide usually kills people by displacing the air that they need to breathe, but in high-enough concentrations it is poisonous as well.
So you see that although volcanologists might disagree on the tectonic details underlying Cameroon and might even disagree on the triggering mechanism for the 1986 disaster, but the danger is pretty well understood. Obviously one way to minimize the chances of this happening again is to prevent the deep lake waters from becoming gas-charged. A program was started to have a pump running that brought up the deep water (in a small controlled way), where it was pumped into the air like a fountain. This allowed smaller amounts of deep water to lose their carbon dioxide gradually rather than having the potential of a big bubble occurring again. Another worry is that the lake walls themselves are not very strong (they are constructed of tuff, partially solidified ash). The problem is that there is this lake at an elevation higher than the main towns nearby, and if for any reason the walls of the lake were breached there would be a flood of water that could be just as dangerous as a flood of carbon dioxide. I’m not sure but I think there were plans to pump water out of the lake to try and keep the level and pressure down.

Mt. Erebus.
The most recent eruption of Mount Erebus began in 1972 and stopped in 1992. It shares some similarities with both Kilauea and Mount St. Helens but also has some significant differences. Like Kilauea current eruption, the vent was a lava lake that produced some lava flows. Unlike Kilauea, Mount Erebus is a stratovolcano, the same type of volcano as Mount St. Helens. Eruptive activity at Mount Erebus tends to be strombolian , a little more explosive than Kilauea but far less than Mount St. Helens. The composition of the volcanic products at each volcano is different. The silica content of lava from Kilauea and Mount Erebus are about 50 weight percent. However, Mount Erebus rocks have greater amounts of alkali elements (sodium and potassium), about 9-10 weight percent compared to the 3 weight percent for rocks from Kilauea. Ash and lava from Mount St. Helens are called dacite because they contain about 64 weight percent silica, much more than the other two volcanoes.
The tectonics around Erebus are not too clearly understood, but as summarized by Tom Simkin and Lee Seibert in “Volcanoes of the World”, there is a large continental rift that is cutting through the W part of Antarctica (I guess it is this splitting-apart that has formed the Ross Sea). The most famous continental rift is the E. African Rift, and it too is associated with volcanism.
There are chapters on Mt. Erebus in “Volcanoes of the Antarctic Plate and Southern Oceans” by WE LeMasurier and JW Thomson, editors. In the summary chapter on the Mt. Erebus volcanic province, PR Kyle writes “Mt Erebus is an active volcano…and contains a persistent convecting lava lake of anorthoclase phonolite magma. Small Strombolian eruptions occur on a daily basis…often ejecting anorthoclase phonolite bombs onto the crater rim. From September to December 1984, larger Strombolian eruptions occurred more frequently, ejecting bombs up to 2 km from the crater and sending small eruption columns to over 2 km high.,,”
The lava lake in the summit crater counts as an ongoing eruption but it is not a particularly active eruption. The fact that the volcano is composed of layers of lava and ash erupted mainly from the summit means that it does erupt in a bigger way, it is just that no humans have seen it happen.
What happened at Ruapehu in 1995?
Since the activity in 1995, the volcano has quieted down and the only threat is to people in the immediate vicinity of the summit. The nearest large town is Taupo, with a population that probably varies between 20,000 and 60,000 depending on whether it is summer or winter. Taupo is ~100 km NE of Ruapehu. Other smaller but closer towns are Turangi, Waiouru, Ohakune, and National Park. Even tinier places are Whakapapa, Iwikau, Tukino, and Turoa Villages, all within~10 km of the summit.
As far as I know there hasn’t been any damage from the 1995 eruption other than the minor inconvenience of slight dustings of ash. There have also been the inconvenience to lots of skiers hoping to enjoy the last month of ski season and being kept off the slopes. The first day of the eruption 4 volcanologists were injured in a plane crash. Two received only minor scrapes and bruises but the other two were more seriously hurt and are still recovering. I have forwarded your question to a friend who is working at the volcano observatory there and hopefully she can provide some more answers.
The current research involves studying both the activity that led up to the 1995 eruption and the activity during the most recent eruption. The former is an attempt to try and understand what the precursors were and can the be better identified prior to the next eruption. The latter is to try and understand the mechanisms of explosions, ashfall, and lahar generation and emplacement.

Ruapehu has had 50 historic eruptions, more than any other crater lake in the world. Most of these eruptions have been phreatic. The last major eruption was in March of 1992. In this century the longest period of repose between eruptions at Ruapehu has been 12 years (1906-1908). The average period of repose was about 4 years.
Eruptions at Mount Ruapehu are caused by the interaction of lava and water. Volcanologists call this a phreatic eruption. Because the lava erupts into the bottom of a crater lake, a lot of heat energy converts water to steam. The steam expands violently, throwing water and sometimes ash into the air. The recent increase in activity was caused by a greater volume of lava being erupted into the lake. Similar, but usually smaller, phreatic eruptions have been occurring at Ruapehu since 1889.
Has there ever been any volcanic activity in Australia?
There has not been any eruptions in Australia in this century. The most recent eruption in Australia was at Mt. Gambier, a shield volcano in the Newer Volcanic Province, Victoria. The Newer Volcanics Province in Victoria Australia is made of four shield volcanoes and associated vents: Red Rock, Mt. Napier, Mt. Schank, and Mt. Gambier. They last erupted between 5850 and 2900 B.C. The eruptions were explosive and some generated lava flows. It is impossible to say if the volcanoes will erupt again. However, there have been rare earthquakes in the area, most recently in 1976. There are numerous volcanic islands north and east of Australia including North Island, New Zealand, the islands of Vanuatu, the Solomon Islands, New Britian, and Indonesia. There are numerous interesting volcanic provinces in Australia. There are flood basalts of Cambrian age (about 650 million years old) northeast of Halls Crossing in northern Australia. Volcanism commenced about 70 million years ago at volcanic centers in southeast Queensland and northeast South Wales. Compositions range from basalt to rhyolite and includes shields, plugs, and domes. In north Queensland there are some very long basaltic lava flows. For example, at Undara a flow is 100 miles (160 km) long. You might have a look at Johnson and others (1989).
The 1883 eruption of Krakatau.
Krakatau erupted in 1883, in one of the largest eruptions in recent time. Krakatau is an island volcano along the Indonesian arc, between the much larger islands of Sumatra and Java (each of which has many volcanoes also along the arc). There is a very fine book about the Krakatau eruption by Tom Simkin and Richard Fiske (Simkin, T., and Fiske, R.S., Krakatau 1883: The volcanic eruption and its effects: Smithsonian Institution Press: Washington, D.C., 464 p.), so if you really want to know about the eruption you should go to the nearest bookstore or library to find that.
Here are some highlights from their summary of effects:
1. The explosions were heard on Rodriguez Island, 4653 km distant across the Indian Ocean, and over 1/13th of the earth’s surface.
2. Ash fell on Singapore 840 km to the north, Cocos (Keeling) Island 1155 km to the SW, and ships as far as 6076 km west-northwest. Darkness covered the Sunda Straits from 11 a.m. on the 27th until dawn the next day.
3. Giant waves reached heights of 40 m above sea level, devastating everything in their path and hurling ashore coral blocks weighing as much as 600 tons.
4. At least 36,417 people were killed, most by the giant sea waves, and 165 coastal villages were destroyed.
5. When the eruption ended only 1/3 of Krakatau, formerly 5×9 km, remained above sea level, and new islands of steaming pumice and ash lay to the north where the sea had been 36 m deep.
6. Every recording barograph in the world documented the passage of the atmospheric pressure wave, some as many as 7 times as the wave bounced back and forth between the eruption site and its antipodes for 5 days after the explosion.
7. Tide gauges also recorded the sea wave’s passage far from Krakatau. The wave “reached Aden in 12 hours, a distance of 3800 nautical miles, usually traversed by a good steamer in 12 days”.
8. Blue and green suns were observed as fine ash and aerosol, erupted perhaps 50 km into the stratosphere, circled the equator in 13 days.
9. Three months after the eruption these products had spread to higher latitudes causing such vivid red sunset afterglow that fire engines were called out in New York, Poughkeepsie, and New Haven to quench the apparent conflagration. Unusual sunsets continued for 3 years.
10. Rafts of floating pumice-locally thick enough to support men, trees, and no doubt other biological passengers-crossed the Indian Ocean in 10 months. Others reached Melanesia, and were still afloat two years after the eruption.
11. The volcanic dust veil that created such spectacular atmospheric effects also acted as a solar radiation filter, lowering global temperatures as much as 1.2 degree C in the year after the eruption. Temperatures did not return to normal until 1888. The book is full of many more amazing bits of information. Hopefully these small excerpts will be useful to you.

Recent activity
Krakatau is still active. The presently-active vent has formed a small island in the middle of the ocean-filled caldera that developed during the famous big eruption of 1883. The island is called Anak Krakatau, which means child-of-Krakatau. It is pretty much erupting all the time at a low level, but once or twice a year it has slightly larger eruptions that people notice and sometimes report in the news. Of course none of these are anywhere near the size of the famous 1883 eruption.
Krakatau is following a pattern that is pretty common for volcanoes. This pattern involves hundreds to thousands of years of small eruptions to build up the volcano followed by 1 or more huge eruptions that causes the volcano to collapse into a caldera, and then the cycle starts over again.
I think the chances of a huge 1883-style eruption are very small for the time being. However, it is certainly dangerous to go onto Anak Krakatau, especially if it is one of its more agitated moods. It is probably not even very smart to spend too much time on the small islands that form the remnants of what was once the main Krakatau island. This is because even a small collapse of Anak Krakatau could generate a small tsunami that could sweep towards these islands. Since they are so close to Anak Krakatau there wouldn’t be very much time for a warning.
The 1991 eruption of Mt. Pinatubo was the third largest volcanic eruption this century (after Santa Maria, Guatemala in 1902 and Novarupta, Alaska in 1912). Because it was so large – and it strongly effected both greenhouse warming and the ozone hole – there have been lots of scientific publications about it.
When is the last time Mt. Fuji erupted?
Mt. Fuji is a beautiful example of a stratovolcano, and is almost a perfect symmetric cone (at least when viewed from far away). It is mostly basalt, which is a little bit unusual for stratovolcanoes. Most stratovolcanoes are constructed of andesite or dacite compositions. The fact that it is a stratovolcano means that it is composed of layers of both lava and ash. The fact that it is such a beautiful cone probably indicates that it hasn’t recently suffered a big eruption.
“Volcanoes of the World” by Tom Simkin and Lee Seibert lists 63 eruptions of Mt. Fuji since about 9000 years ago. Obviously most of these have been determined by using carbon-14 dating rather than accounts by witnesses. However, the most recent 22 eruptions are listed as having been recorded by people. The most recent eruption was in 1709 but probably the most famous was in 1707.

How were the Deccan Flood basalts formed?
The currently popular theory is that once in a while great teardrop-shaped blobs of magma work their way to the surface from near the core-mantle boundary. These huge blobs result in extensive volcanism (that we call flood basalts). Following behind the teardrop shaped blob is a long skinny tail that persists for millions of years, and results in what we call hotspot volcanism. Some of the evidence for this is that most flood basalt provinces can be tied to a hotspot trace – in the Deccan Traps example the hotspot is now under Reunion Island. On the Indian Ocean floor you can trace a line of now-extinct volcanoes that follow back to NW India.
Are there any volcanoes erupting in Costa Rica?
November 1996 activity at Rincon de la Vieja Rincon de la Vieja began erupting volcanic ash and water vapor on November 6, 1995. Rincon de la Vieja is an active composite volcano located 30 miles (50 km) south of Lake Nicaragua. Local authorities have evacuated 300 families. Scientists are on their way to the remote volcano. Great amounts of ash have been reported in local rivers. Heavy cloud formations have prevented satellite observations.
In 1982, J. Bruce Gemmel climbed the volcano and provided a description to McClelland and others (1989). Collapse craters trend east-northeast to west-southwest across the summit of the volcano. The main cone is heavily vegetated with the exception of three craters to the west. At that time, the most recently active crater (diameter 800 feet, 250 m) was 0.6 mile (1 km) northwest of the main cone. A lake covered the crater floor.
The most recent confirmed eruption of Rincon de la Vieja was in early 1992. Rincon de la Vieja has erupted at least 16 times since 1851. Most eruptions are phreatic and include the emission of gas and ash. Lahars are often generated by displacement of the crater lake.
What type of volcanic damage prevention is there in Central America?
Unfortunately most of the Central American countries are too poor to do much with damage prevention. I remember being told by a Guatemalan geologist that they can’t keep their seismometers going because the people come along and steal the batteries for their cars. A dedicated engineering geologist would go a long way towards lowering the problems, by determining the places where the danger from lahars is the greatest, by determining the most unstable slopes, by determining which roads are the best for evacuation routes

What’s the most recent eruption of Vesuvius and will it erupt again?
Vesuvius has erupted about three dozen times since 79 A.D., most recently from 1913-1944. The 1913-1944 eruption is thought to be the end of an eruptive cycle that began in 1631. It has not erupted since then. Vesuvius is an active volcano, it will erupt again.

The oldest dated rock at Mt Vesuvius is about 300,000 years old. It was collected from a well drilled near the volcano.
Vesuvius erupted catastrophically in 79 A.D., burying the towns of Herculaneum and Pompeii. The Somma Rim, a caldera-like structure formed by the collapse of a stratovolcano about 17,000 year ago, flanks Vesuvius to the east.
The 79 A.D. eruption of Vesuvius was the first volcanic eruption ever to be described in detail. From 18 miles (30 km) west of the volcano, Pliny the Younger, witnessed the eruption and later recorded his observations in two letters.
Volcanologists now refer to sustained explosive eruptions which generate high-altitude eruption columns and blanket large areas with ash as plinian eruptions. It is estimated that at times during the eruption the column of ash was 20 miles (32 km) tall. About 1 cubic mi (4 cubic km) of ash was erupted in about 19 hours. Vesuvius has erupted about three dozen times since 79 A.D., including a large, explosive eruption in 1631 that killed 4,000 people. The most recent eruption was from 1913-1944.
An excellent source of information on Vesuvius is an article titled “The Eruption of Vesuvius in A.D. 79” by Sigurdsson and others (Sigurdsson, H., Carey, S., Cornell, W., and Pescatore, T., 1985, The eruption of Vesuvius in A.D. 79:
What happened at Pompeii?
Probably one of the most detailed studies of a large explosive eruption anywhere was that done on the AD 79 Vesuvius eruption by Sigurdsson et al (1985) in the National Geographic Research vol. 1, no. 3, pp. 332-387. It should be required reading for any students of Vesuvius and Pompeii. It is a long article so I won’t re-type more than just a short part that deals with the people being killed.
The citizens of Pompeii became directly aware of Vesuvius’ eruption sometime in the early afternoon of 24 August as coarse pumice-fall began to plunge the city into darkness. No trace remains of the initial fine ashfall that affected areas to the north and east farther up the slopes of the volcano. The initial explosive phase of the eruption (A-1) may have been witnessed from Pompeii — an excellent view of the summit — but without direct consequences; it probably only generated curiosity.
Pumice and lithics rained continuously from early afternoon on 24 August to early the next morning, according to accounts of Pliny the Younger. During this time 130-140 cm of white pumice (A-2) accumulated, on top of which another 110 to 130 cm of gray pumice was laid down (A-3 to A-5). Although the pumice layer appears relatively homogeneous, the diameter of lithics and pumice varies throughout (Figure 23), reflecting variations in height of the eruption column during the Plinian phase. For example, relatively large, dense pieces of pumice are concentrated about 10 cm above the base of the gray pumice layer.
Pompeii happened to be located on the secondary thickness maximum of the fallout deposit, and thus received the thickest accumulation of pumice- and lithic-fall. Roofs probably collapsed about halfway through the deposition of white pumice (some 40 cm). With most structures unsafe for habitation, an exodus from the city is likely to have begun. Escape of most of Pompeii’s residents can thus be attributed to the extended yet comparatively innocuous Plinian fallout phase.
The first surge to reach Pompeii swept against the north wall of the city in the early morning of 24 August, depositing dark gray ash near the Herculaneum Gate (S-3). Neither Vesuvius Gate, 200 m to the east, nor other sites in the Pompeii area evidence this surge. Therefore it is possible that the surge cloud did not extend inside the city walls but flowed just west of Pompeii over the villa dei Misteri and Villa Diomede. The surge must have caused alarm and almost unbearable conditions in the city. Evidence from Mount St. Helens in 1980 and El Chichon in 1982 indicates a peripheral zone of high heat associated with the distal ends of surge clouds (Moore & Sisson 1981; Sigurdsson, Carey et al. 1984). The first surges of the eruption (S-1 and S-2), which form a characteristic doublet in the middle of the pumice-fall at Boscoreale and Oplontis, did not reach as far southeast as Pompeii.
Within the city, 3 cm of pumice- and lithic-fall (A-6) accumulated before the next surge (S-4) overwhelmed the city. This surge also extended to Bottaro, 1 km south of Pompeii (Figure 8), and to Tricino, 3 km east of Pompeii (Figure 10).
The majority of human remains discovered in the excavations have been found on top of the pumice-fall layer, lying within surge deposits S-4 and S-5, but principally buried by the thick S-6 surge (Figure 24). Because of their fine-grained, silty nature, the surges have preserved accurate molds of the victims including details of facial expressions and sometimes clothing. With time, the soft tissues have decayed, leaving only bones in the hollow cavities. In 1860, Giuseppe Fiorelli developed the ingenious technique of making plaster casts of these impressions before the surrounding surge is disturbed (figure 25). Many hundreds of casts of the dead in Pompeii have since been made in this manner. In 1966, for example, casts of 13 victims were made in the Garden of the Fugitives, where they fell in various groups of adults and children on top of the pumice-fall deposit. Thus evidence is compelling that the S-4 surge was the lethal event in Pompeii. Since detailed stratigraphic studies of the deposits have not yet been feasible inside Pompeii, the extent of building damage resulting from the S-4 surge cannot be judged. By this time the ground-floor levels of buildings had already been buried and only the upper stories protruded from the pumice blanket.
As I said, this is a very comprehensive study, and you should get a copy or two to read. It seems as if the people survived the fall of over a meter of pumice but that this pumice did start to collapse walls. When pyroclastic surges (fast-moving, horizontally-directed, ash and rock-laden blasts) occurred, the remaining people were killed, either by the heat, by lack of oxygen in the thick dusty cloud, or by being struck by larger blocks. They were buried by a later surge. Most of what you see as the “bodies” are apparently plaster casts rather than preserved bodies.
Why are there volcanoes in Iceland?
Iceland is an island made of numerous overlapping volcanoes. The volcanoes are associated with a divergent plate boundary, the Mid-Atlantic Ridge, where new oceanic plate is added to the North American and Eurasian plates and with a hot-spot that underly. No wonder there is such prolific volcanic activity in Iceland!

Surtsey erupted from November of 1963 to June of 1967. At the end of the eruption the new island was about 2,500 feet (800 m) in diameter. The island has not grown since the last eruption. The volcano is made of lava flows and tephra on a submarine (under water) base of pillow lava. The subaerial (under air) lava flows serve as a protective cap on part of the island. The part of the island that is made of only tephra is prone to wave erosion. As tephra is removed by waves the island will shrink in size.
Surtsey is featured in a number of volcano books because it popped to the surface of the Atlantic Ocean and was able to be studied pretty closely almost from the first moments.
Prior to the eruption that location was known to be shallow so it wasn’t a huge surprise; the shallow water was actually the summit of an undersea volcano that wasn’t quite at the surface yet. The first eruptions were very explosive because of the mixture of hot lava and ocean water, and even after the island built above sea level, water was still able to seep through the ash to make more explosions. Eventually, however, the vent got sealed off from the ocean and the eruptions became non-explosive Hawaiian-style lava fountains instead of steam explosions. In fact, explosive water-lava eruptions are now called “Surtseyan” after Surtsey. A number of times water would suddenly gain access to the vent and instantly the activity would go from quiet Hawaiian-style fountaining to explosive Surtseyan explosions. This created quite a hazard to those studying the volcano.
Surtsey is still around. Volcanoes of the World (by Tom Simkin and Lee Seibert) lists its last activity as 1967. It is now being attacked by waves, and if it does not erupt again it may eventually disappear. Many times during the early days of its eruption it was nearly washed away by the waves. That is because it was originally mostly ash (formed by the explosive interaction of erupting lava and ocean water). As soon as the cone had built to the point that water could no longer gain access to the vent, the eruption style changed to more Hawaiian style, with lava fountains producing lava flows. These hardened flows are much more resistant than the earlier-produced ash, and they guarantee that the island won’t be washed away very quickly.
What is Etna doing?
Etna (location: 37.7N, 15.0E) is a 10,791 foot (3,290 m) tall stratovolcano on the northeastern edge of Sicily. It certainly deserves to be on a list of famous volcanoes because it has the longest documented record of volcanism in the world, erupts frequently, and is the largest volcano in Europe. Etna probably does not belong on a list of great (individual) volcanic eruptions. In mythology, Etna was identified as the location of the forge of Volcan, home of the Cyclopses, and where the giant Enceladus laid (eruptions being his breath and earthquakes his motion). Etna has erupted nearly 160 times since the first recorded eruption in 1500 B.C. Most eruptions consist of lava flows. Small to moderate explosive eruptions occur near the summit less often. Etna is not known for any eruptions that caused great loss of life. Known violent eruptions occurred in 1169, 1669, 1752-1754, 1893-1899, 1917, 1940, and 1945. Known fatal eruptions occurred in 141 B.C., 1329 A.D., 1536 A.D., 1832 A.D., 1843 A.D., 1928 A.D., 1979 A.D., and 1987 A.D. I could not find information on the number and cause of fatalities during most of these eruptions. Nine people were killed and 23 were injured (150 tourist were in the area) on September 12, 1979, by a 30-second explosion that threw large blocks near the crater rim. Blocks 10 inches (25 cm) in diameter fell 1,300 feet (400 m) away. Two people were killed and 7 others injured by falling volcanic material 1,600 feet (500 m) from the crater in the 1987 phreatic eruption.
Eruption mechanism
The general accept explanation for volcanism at Mt. Etna, Vulcano, and Stromboli is the subduction of part of the northward-moving African Plate beneath the Eurasian Plate. Subduction may be in a late stage of evolution or may have ceased. Mt. Etna is unusual because it is adjacent to but just outside of the subduction zone. It has been suggested that material associated with the subduction zone migrates into the lithosphere adjacent to the Aeolian arc. Lava from Etna shares a few chemical similarities to hot spot volcanoes. However, these may result from modification at the source or during ascent of the magma.

What’s the cause of volcanism in Yellowstone and is it likely to erupt again?
The U.S. Geological Survey experts on Yellowstone have an excellent homepage that describes the volcanic system. They state that Yellowstone will erupt again but at the present level of activity there is no need for immediate concern about eruptions.
Yellowstone is the result of three very large eruptions. The first was about 2.1 million years ago and erupted 600 cubic miles (2,500 cubic km) of ash. It created the Island Park caldera about 45 miles (75 km) long. The second eruption was about 1.3 million years ago. It produced the Henrys Fork caldera in the west end of the older (2.1 million year) caldera. About 600,000 years ago an eruption produced about 250 cubic miles (1,000 cubic km) of ash and the present-day Yellowstone caldera. Since then there have been large volume (250 cubic miles; 1,000 cubic km) rhyolitic lava flows between 150,000 and 70,000 years ago. Yellowstone is still an active volcano. It sits above a hot spot and will erupt again. To learn more about hot spots.
North Dakota’s volcanic future
Yellowstone is not coming to us. We are going towards Yellowstone (no offense to North Dakotans but this sounds good to me). The location of the Yellowstone hot spot is fixed within the Earth. The North American plate is moving about 2.3 cm to the southwest each year. Assuming plate motions remain constant, the North Dakota-Montana state line will be over the hotspot in 22 million years. The town of Bismarck will be over the hotspot in 33 million years. Fargo will be over the hotspot in 44 million years. Duluth will have to wait 59 million years before it basks in the warmth of geothermal splendor.

What’s going on at Mammoth Lakes?
The surface of part of the Long Valley caldera has been doming upward gradually for some time. Volcanologists do not believe any material is working its way to the surface. Have a look at Current Activity and Eruptions. Past eruptions at Long Valley caldera were very large. Future eruptions might also be large but the recent flurry of earthquakes is nothing to be concerned about, so far.
Activity at Long Valley caldera has returned to background levels after an increase in the number and size of earthquakes prompted a change in alert status over the last few days.
A swarm of earthquakes occurred beneath the south moat of Long Valley caldera on Thursday and Friday. A series of earthquakes with magnitude greater than 3 occurred beneath the south moat on Friday. Earthquake activity beneath the south moat declined after a magnitude 3.6 earthquake on the late afternoon on Friday, February 16.
Two other earthquakes (magnitudes 3.7 and 2.3) occurred near Red Slate Mountain 6 miles (4 km) south of the caldera and 1 mile (0.6 km) south-southwest of Mammoth Lakes on Sunday evening and Monday morning, respectively.
This level of unrest occurs at Long Valley caldera every few weeks or months.
Current activity is consistent with a NORMAL STATUS.
Normal activity in Long Valley caldera includes as many as 10 to 20 earthquakes of magnitude 2 of less per day with occasional swarms of small earthquakes. The earthquakes are accompanied by steady uplift of the resurgent dome at a rate of about 1 inch (2 to 3 cm) per year.
Why is there volcanic activity there?
The Long Valley Caldera is situated along the east edge of the Sierra Nevada Mountains and the western edge of the Basin and Range Province. Earthquakes, extension, and faulting indicate the area is tectonically active. You are correct that the area is not associated with subduction. The volcanism is probably related to extension. As the crust thins, the asthenosphere rises closer to the surface and it may melt, producing magma. The magma probably rises into the crust where it heats and melts rocks. The melted crustal rocks are the rhyolite magmas that are erupted during caldera collapse.
Is Mt. Rainier a dangerous volcano?
Mount Rainier is potentially the most dangerous volcano in the Cascades because it is very steep, covered in large amounts of ice and snow, and near a large population that lives in lowland drainages. Numerous debris avalanches start on the volcano. The largest debris avalanche traveled more than 60 miles (100 km) to Puget Sound. The most recent eruption was about 2,200 years ago and covered the eastern half of the park with up to one foot (30 cm) of lapilli, blocks, and bombs.
Are there volcanoes in Canada?
There are more than 100 volcanic centers in British Columbia and the Yukon. Many are monogenetic cones but there are also large shield and dome complexes. Eruptions began a few million years ago. The most recent eruption was about 200-250 years ago Stikine Volcanic Belt.
The volcanic rocks are divided into five groups with diverse types of volcanoes and tectonic settings. In southern British Columbia, the Pemberton and Garibaldi volcanic belts and the Chilcotin Group plateau are related to the subduction of the Juan de Fuca and Explorer plates beneath the North American continent. The Anahim Volcanic Belt trends easterly across central British Columbia and is probably related to a mantle hot spot. The Stikine Volcanic Belt forms a broad zone of volcanoes in northwestern British Columbia and the southern Yukon. These volcanoes are probably related to shear along the Queen Charolette transform fault to the west. The Wrangell Volcanic Belt is an arc of continental volcanoes associated with the subduction of the Pacific plate beneath the North American plate. Volcanism has also occurred at the Clearwater-Quesnel and McConnell Creek area and at Alert Bay.
Why are there no active volcanoes in the Great Plains and Eastern U.S.?
Most subaerial eruptions are associated with subduction zones, places where oceanic plates are forced under overriding, less dense plates. In this setting, heat and water in the mantle cause rocks to melt. If the melt (magma) makes it to the surface there is an eruption. There are no active subduction zones on the east coast (or the Great Plains).
Why are there so many volcanoes in the Pacific Northwest and Alaska?
The distribution of volcanoes in the northwest and Alaska is the result of plate tectonics. In the northwest, the oceanic Farallon Plate is being pushed beneath (subducted) the continental margin of the North American Plate. When the subducted plate comes in contact with the hot asthenosphere beneath the continental plate conditions are right for the rocks in the asthenosphere to melt. The melt, called magma, rises to the surface to build volcanoes. Since the subduction zone is a long curvi-linear feature it produces a similar line of volcanoes, called an arc, on the continent. Alaskan volcanoes are the result of the subduction of the Pacific plate under the North American plate.

If new oceanic lithosphere is created at mid-ocean ridges, where does it go? Geologists had the answer to this question before Vine and Matthews presented their hypothesis. In 1935, K. Wadati, a Japanese seismologist, showed that earthquakes occurred at greater depths towards the interior of the Asian continent. Earthquakes beneath the Pacific Ocean occurred at shallow depths. Earthquakes beneath Siberia and China occurred at greater depths. After World War II, H. Benioff observed the same distribution of earthquakes but could not offer a plausible explanation.
The movement of oceanic lithosphere away from mid-ocean ridges provides an explanation. Convection cells in the mantle help carry the lithosphere away from the ridge. The lithosphere arrives at the edge of a continent, where it is subducted or sinks into the asthenosphere. Thus, oceanic lithosphere is created at mid-ocean ridges and consumed at subduction zones, areas where the lithosphere sinks into the asthenosphere. Earthquakes are generated in the rigid plate as it is subducted into the mantle. The dip of the plate under the continent accounts for the distribution of the earthquakes.
How many volcanoes are there in North America?
We get so many questions about how many volcanoes are in Washington or Oregon or Canada, we thought we’d break it down and give you a listing of volcanoes by state (and country). If your state is not listed it means there are no volcanoes there (as described below). This list is compiled from Volcanoes of North America: United States and Canada, by Wood and Kienle. They list all volcanoes that are younger that 5 million years old and are morphologically distinct. There are about 262 volcanoes and volcanic fields in North America.

Number of Volcanoes
and Volcanic Fields
Canada* 21
New Mexico*

Why were there many active volcanoes in southern California about 10,000 years ago, but not now?
I will give you two hypothesis.
The first one is that perhaps the volcanoes in that region (which part of California you’re talking about, I don’t know) aren’t really extinct. There are a number of volcanoes that have average repose periods between eruptions that are measured in thousands of years. To a human it might seem as if the volcano is dead, but in fact it is just between eruptions.
The second idea is that magma is required to produce volcanoes and the tectonic setting of California has changed from one that produces magma to one that doesn’t. Tectonic changes such as these take millions of years to occur though, so I don’t think this second hypothesis is very likely.
Which volcano killed the most people?
The eruption of Tambora in Indonesia in 1815 killed the most people. It was a huge eruption that sent ash into the stratosphere that then spread around the world. World climate was noticeably cooler the following year, and in places it was called “the year without a summer”. Closer to the eruption itself thousands of people were killed, and due to the destruction of crops, disease, contamination of water, etc., tens of thousands more died in the years that followed. Overall about 92,000 people died as a result of the eruption. Back in 1815 there wasn’t much news coming out of Indonesia to the western world but if it were to happen today it would definitely have been a big deal.

How many act ive volcanoes are there?
The absolute number of volcanoes that exists depends on your definition: active only, active, dormant plus extinct volcanoes? And even if we decide on a definition, nobody has really counted all of the volcanoes, especially the tens on thousands on the sea floor. The best guess is 1511 volcanoes have erupted in the last 10,000 years and should be considered active. This number is from the new Smithsonian Institution book, “Volcanoes of the World: Second Edition” compiled by Tom Simkin and Lee Siebert.

What’s the biggest volcano in the Solar System?
So far, the largest volcano in our Solar System is Olympus Mons on Mars. It is about 17 miles (27 km) tall. That’s a long hike for some future explorer. Mount Everest is about 6 miles (10 km) tall

What’s the tallest volcano in the world?
The highest volcano is Ojos del Salado in Chile. It is 22,589 feet (6,887 m) tall.
From base to summit, the tallest would be Mauna Kea, which, when measured from its base on the ocean floor, is more than 30,000 feet high.

What is the largest eruption ever?
The biggest eruption I know of was at Yellowstone about 2.2 million years ago. An explosive eruption produced 2,500 cubic kilometers of ash! That’s about 2,500 times more ash than Mount St. Helens erupted!

Why do volcanoes erupt?
Volcanoes erupt because of density and pressure. The lower density of the magma relative to the surrounding rocks causes it to rise (like air bubbles in syrup). It will rise to the surface or to a depth that is determined by the density of the magma and the weight of the rocks above it. As the magma rises, bubbles start to form from the gas dissolved in the magma. The gas bubbles exert tremendous pressure. This pressure helps to bring the magma to the surface and forces it in the air, sometimes to great heights.
It’s sort of like the bubbles of gas in a bottle of soda. Before you open the soda you don’t see many bubbles because the pressure in the bottle keeps the gas dissolved in the soda. When you open the bottle the pressure is released and the gas bubbles leave the soda. If you shake up the bottle first, the soda gets pushed out by the bubbles of gas as they rush out. You might try this, but do it outside.
What are the s igns that a volcano is about to erupt?
Short answer
Several things happen when a volcano is about to erupt, some of the most obvious are listed here:
1. The number and size of earthquakes increase in and around the volcano.
2. The ground deforms or “bulges” at the eruption site.
3. A lot more gas comes out of the volcano.
Longer answer
There are lots of signs that are examined, depending on how closely monitored the particular volcano is. Probably the most common type of monitoring is by seismicity. Even one seismometer can tell if there is an increase of seismic activity on a usually seismically-quiet volcano. If you have at least 3 seismometers, and they are strategically placed, you can triangulate on earthquakes to see if they are occurring in a place that indicates perhaps magma movement. By examining the seismic data over a period of time you may be able to determine if the earthquakes are migrating towards the surface (suggesting that magma is also migrating towards the surface since the earthquakes are probably being generated as magma breaks rocks that are in its way).
Another type of data that is used is the study of ground deformation. When magma moves up into the shallow plumbing of a volcano, it takes up space and pushes the surrounding rock outward. This also causes the surface of the volcano to deform. Some points move upward and any two points will move farther apart. By using very accurate leveling and distance-measuring techniques, these surface changes can be measured. Usually the changes are a few mm over a distance of a few hundred meters, but sometimes they are dramatic. For example prior to many eruptions at Kilauea, the summit bulges 1-2 meters upward. In the last few days prior to the big Mt. St. Helens eruption the northern flank was bulging outward at a few meters per day!
Some people like to monitor volcanoes by constantly monitoring gases that come out of fumaroles. Most active volcanoes have fumaroles where volcanic gases escape to the surface. It is relatively easy to monitor the temperatures of these gases, and an anomalous increase in temperature might be a sign that magma has moved closer to the surface. Monitoring the composition of the gases is more difficult to do, and changes in the composition are way more difficult to interpret. Many times just visual changes to fumarole areas are indications of impending activity. If the area of active degassing gets larger, if the plants nearby die suddenly, if the color of any lakes or ponds nearby changes…Many volcanoes have summit lakes through which heat and gases rise to the surface and escape. Many of these lakes have strange colors due to all the dissolved minerals in them, and many of the colored ones change color, pH, temperature, etc. These too, are signs of change below but are often difficult to interpret.
A number of people are studying ways in which to use satellite data to monitor volcanoes. It is possible to obtain thermal images of volcanic areas, and by comparing images on a monthly or bi-weekly basis, increases or decreases in temperatures can be detected. Additionally, some new technologies have allowed for the determination of very accurate topography from satellite data. This technology may someday allow for the remote monitoring of surface deformation associated with sub-surface magma movement. This process is still being developed. It usually takes too long to get satellite data processed for this technique to be useful in a rapidly-escalating crisis so it would be used over the long term, in the years to months prior to an eruption rather than the hours prior.

Why are some eruptions gentle and others violent?
Volcanic eruptions could be thought as a continuum between two end members. At one extreme is the gentle effusion of lava. Most Hawaiian eruptions would be a examples of this type of eruption. At the other extreme is the explosive ejection of ash from a vent. The May 18, 1980 Mount St. Helens eruption would be an example of this type of eruption.
The two main factors that influence how a volcano will erupt are viscosity and gas content. Both are related to the composition of the magma. Hawaiian volcanoes tend to erupt basalt, which is low in viscosity and low in gas content (about 0.5 weight percent). The gas that is present can readily escape and little pressure builds up in the magma. At the other extreme, rhyolite magmas are very viscous and can contain a lot of gas (up to 7-8 weight present). As the magma moves into the vent and the pressure drops, the gas wants to escape. The magma is very sticky and resists the expansion of the gas bubbles. Ultimately, enough bubbles grow and expand to blow the magma into ash size fragments and eject them violently into the atmosphere.

How high can explosive eruptions go and how far can the ash be spread?
Well, that depends on how big the eruption is and how big the debris is that you are concerned about. As you might imagine a big eruption will send material farther. Additionally, the big material from any eruption doesn’t get thrown as far as the finer stuff.
Volcanologists go out into the field to figure out the distribution of erupted pyroclastic material. They will go to numerous sites around the volcano and measure (in general) 3 things: 1) the total thickness of the pyroclastic deposit at each location; 2) the average size of the 10 largest pumice at each location; and 3) the average of the 10 largest lithic clasts at each location (a lithic is a pre-existing rock that is blown apart in the explosive eruption). They then draw contours around the data that they have collected. In some cases, if the geologists are studying a very old eruption, they may not even know where the vent was. The contours of the thickness and size measurements should close around the vent so that its location can be determined.
Data from a very extensive study of the AD 79 Vesuvius eruption by Haroldur Sigurdsson shows that for each of two particularly strong blasts during the eruption, the pumice layer was 100 cm thick up to ~20 km downwind, 50 cm thick out to about 50 km, 25 cm thick out at ~60 km, and so on. Pumice 15 cm in diameter made it out ~6 km downwind, 10-cm pumice made it ~7 km, and 5-cm pumice even further.
The finest dust often gets carried hundreds or even thousands of km downwind.
Explosive eruption plumes such as those generated at volcanoes like Mt. St. Helens or Pinatubo can reach high into the atmosphere. The highest Mt. St. Helens plume on May 18, 1980 reached about 31 km (101,700 feet), and the highest Pinatubo plume got as far as 45 km (147,600 feet).

Here are some highlights very fine book about the 1883 Krakatau eruption by Tom Simkin and Richard Fiske (Simkin, T., and Fiske, R.S., Krakatau 1883: The volcanic eruption and its effects: Smithsonian Institution Press: Washington, D.C., 464 p.). They should give you and idea about how far ash can travel during a large eruption.
-Ash fell on Singapore 840 km to the north, Cocos (Keeling) Island 1155 km to the SW, and ships as far as 6076 km west-northwest. Darkness covered the Sunda Straits from 11 a.m. on the 27th until dawn the next day.
-Blue and green suns were observed as fine ash and aerosol, erupted perhaps 50 km into the stratosphere, circled the equator in 13 days.
-Three months after the eruption these products had spread to higher latitudes causing such vivid red sunset afterglow that fire engines were called out in New York, Poughkeepsie, and New Haven to quench the apparent conflagration. Unusual sunsets continued for 3 years.
-The volcanic dust veil that created such spectacular atmospheric effects also acted as a solar radiation filter, lowering global temperatures as much as 1.2 degree C in the year after the eruption. Temperatures did not return to normal until 1888.
How long do eruptions last?
Historic eruptions have lasted less than a day to thousands of years. In 1977, the lava lake at Nyiragongo drained in less than one hour. In contrast, Stromboli has had a low-level of activity since 450 BC (about 2,400 years).
Percent of Volcanoes Duration
9% < 1 day
16% < 2 days
24% < 1 week
30% < 2 weeks
43% < 1 month
53% < 2 months
17% > 1 year
7% > 3 years
0.5% > 30 years
Durations of eruption based on 3,211 historic eruptions. Data from Simkin and Siebert (1994).
The median duration of historic eruptions is 7 weeks.
Simkin and Siebert (1994) make several important observations:
1. the paroxysmal phase of an eruption can occur during any interval of a volcanoes eruption, for example the 1980 eruption of Mount St. Helens and the 1883 eruption of Krakatau were preceded by months of low-level activity;
2. some eruptions can reach their paroxysmal phase within an hour after the eruption starts, for example the 1886 eruption at Tarawera and the 1977 eruption of Usu;
3. to uninstrumented observers, unrest prior to some eruptions can be too short to provide a warning of an impeding eruption – highlighting the need for careful instrumental monitoring of active volcanoes.

What is the approxiamate temperature of an eruption cloud?
The cloud that rises above a volcano is surprisingly cold, probably close to freezing! This is because the gasses are expanding so rapidly.
The temperature of the ash flows that can be erupted, on the other hand, are on the order of hundreds of degrees Celsius.

Could a series of eruptions possibly have caused the extinction of the dinosaurs due to the ash depoited in the atmosphere?
Lots of people have definitely proposed that idea, but it is still not decided. For one thing there would have to be eruptions such as we’ve never seen before. About the only possibility would be a flood basalt eruption (and one did indeed occur around the 65 million years-ago time). The only problem is that flood basalt eruptions aren’t explosive. They are persistent though so they might have filled the atmosphere with enough sulfuric acid aerosols (vog) to make life difficult. There’s no doubt that a big meteorite hit around 65 million years also. One idea I’ve heard lately is that the flood basalt eruption (the Deccan Traps in NW India) had stressed life on the planet, and the big meteorite impact was the last straw.
There have been other global extinctions, and if you go by the percentage of Earth’s life that has disappeared in them, the one 65 million years ago is not the most drastic. What I wrote above would require that a meteorite hits during a time of flood basalt volcanism more than once. Is that too much of a coincidence to occur multiple times? I don’t know.

Why are there volcanoes?
Volcanoes are a natural way that the Earth and other planets have of cooling off. Planets are warm in their mantles. Heat inside planets escapes towards their surfaces. Heat sometimes melts rocks, which then rise buoyantly toward the planet’s surface. When the hot rocks – called magma – and included gases break through the crust, an eruption occurs. The buildup of ash and lava flows around the eruption hole (or vent) makes a volcano. Some volcanoes erupt for only a short time – a few days to weeks and never erupt again. Large volcanoes such as stratovolcanoes and shields erupt many thousands of times throughout their lifetimes of hundreds of thousands to a few million years.

The Earth has volcanoes because it is hot inside. In some places it is hot enough to turn solid rock into liquid rock. Geologists call the liquid rock magma. The magma rises towards the surface because it is less dense than the surrounding rock (like a hot air balloon rising through the cooler air). If the magma reaches the surface it is called lava and lava accumulates to make a volcano.
The Earth’s interior is hot, and, as all hot things will do, the interior is trying to lose its heat and cool. It cools mostly by releasing its heat through volcanic eruptions. That is why there are volcanoes.

What are the different types of volcanoes?
There are pretty much 3 types of volcanoes, but as with all kinds of classifications you can find exceptions. The three types that I am thinking of are: 1) shield volcanoes (such as we have here in Hawai’i); 2) stratovolcanoes (such as Mt. St. Helens and Pinatubo); and 3) large rhyolite complexes (such as Yellowstone and Taupo). You might also want to add the mid-ocean ridges, flood basalts, and monogenetic fields to make that 6 types in all. “Volcanoes of the World” by Tom Simkin and Lee Seibert, lists 26 different types, but that’s probably kind of extreme.
Here are some brief descriptions:
Shield volcanoes–the largest of all volcanoes on Earth (not counting flood basalt flows). The Hawaiian volcanoes are the most famous examples. These volcanoes are mostly made up of basalt, a type of lava that is very fluid when erupted. For this reason these volcanoes are not steep (you can’t pile up a fluid that easily runs downhill). These volcanoes are only explosive if water somehow gets into the vent, otherwise they are characterized by low-explosivity fountaining that forms cinder cones and spatter cones at the vent, however, 95% of the volcano is lava rather than pyroclastic material. Shield volcanoes are the common product of hotspot volcanism but they can also be found along subduction-related volcanic arcs and out by themselves as well.
Stratovolcanoes–making up the largest percentage (~60%) of the Earth’s volcanoes, these are characterized by eruptions of cooler and more viscous lavas than basalt. The usual lavas that erupt from stratovolcanoes are andesite, dacite, and occasionally rhyolite. These more viscous lavas allow gas pressures to build up to high levels (they are effective “plugs” in the plumbing), therefore these volcanoes often suffer explosive eruptions. They are usually about 50/50 lava and pyroclastic material, and the layering of these products gives them their other common name of composite volcanoes. Stratovolcanoes are commonly found along subduction-related volcanic arcs.
Large rhyolite caldera complexes–the most explosive of Earth’s volcanoes. These are volcanoes that often don’t even look like volcanoes. They are usually so explosive when they erupt that they end up collapsing in on themselves rather than building any tall structure. The collapsed depressions are called calderas, and they indicate that the magma chambers associated with the eruptions are huge. Fortunately we haven’t had to live through one of these since 83 AD when Taupo erupted. Yellowstone is the most famous U.S. example of one of these. Their origin is still not well-understood. Many folks think that Yellowstone is associated with a hotspot, however, a hotspot association with most other rhyolite calderas doesn’t work.
Monogenetic fields. These also don’t look like a “volcano”, rather they are a collection of sometimes hundreds to thousands of separate vents and flows. These are the product of very low supply rates of magma. The supply rate is so slow and spread out that between the times of eruptions the plumbing doesn’t stay hot so the next batch of magma doesn’t have any preferred pathway to the surface and it makes its own path. A monogenetic field is kind of like taking a single volcano and spreading all its separate eruptions over a large area. There are a number of monogenetic fields in the American southwest, and there is a famous one in Mexico called the Michoacan-Guanajuato field.
Flood basalt provinces–another strange type of “volcano”. Some parts of the world are covered by thousands of square kilometers of thick basalt lava flows–some flows are more than 50 meters thick, and individual flows extend for hundreds of kilometers. The old idea was that these flows went whooshing over the countryside at incredible velocities. The new idea is that these flows are emplaced more like pahoehoe flows–slow moving, with most of the great thickness being accomplished by injecting lava into the interior of an initially thin flow. The most famous U.S. example of a flood basalt province is the Columbia River Basalts, covering most of SE Washington State, and extending all the way to the Pacific and into Oregon. The Deccan Traps of northwest India are a much larger flood basalt province.
Mid-ocean ridge volcanism occurs at plate margins where oceanic plates are created. There is a system of mid-ocean ridges more than 70,000 km long that stretches through all the ocean basins–some folks consider this the largest volcano on Earth. Here, the plates are pulled apart by convection in the upper mantle, and basalt lava intrudes to the surface to fill in the space. Or, the basalt intrudes to the surface and pushes the plates apart. Or, better yet, it is a combination of these two processes. Either way, this is how the oceanic plates are created. A recent mid-ocean ridge eruption took place along the Gorda Rise–the mid-ocean ridge that separates the Juan de Fuca plate from the northern part of the Pacific plate.
What comes out of volcanoes?
Lava, pyroclasts (broken pieces of lava that fly through the air), and gas come out of volcanoes. We have described lava before. Pyroclasts are classified by size. Ash is the smallest pyroclast. Blocks and bombs are the largest. Cinder and pumice are also types of pyroclasts. Water vapor, carbon dioxide, and sulfur dioxide are the most common volcanic gases.
Specific examples of different kinds of eruptions:
Lots of different things come out of a volcano when it erupts depending on what kind of eruption it is. If it is a shield volcano like we have here in Hawai’i, then there is usually a fountain of molten lava that reaches anywhere from 10 to 500 meters into the air. This fountain builds a spatter cone or cinder cone around the vent. Meanwhile, if enough lava is falling from the fountain, a lava flow can develop. If the amount of lava feeding the flow is high, then the flow will move rapidly downhill away from the vent. Rapid-moving flows continually disrupt their surfaces and are constantly exposing more red-hot lava to the atmosphere. This means that the flow is losing a lot of heat and consequently its viscosity increases. As the lava continues to flow rapidly, but now with a high viscosity it starts to get torn into jagged pieces rather than flow nicely. This is how an ‘a’a flow develops.
In some eruptions there is almost no fountaining and the lava just flows slowly away from the vent. In these cases the surface of the lava is not disrupted and can solidify even while the inside is still molten. This is how pahoehoe flow move. If these pahoehoe flows go on long enough then lava tubes can develop within the flow. These lava tubes allow lava to reach the flow front from the vent without losing much heat so it is still pretty fluid even 10’s of kilometers from the vent.
At more explosive volcanoes eruptions are very different. The main difference is that the viscosity of the magma (how fluid or how pasty it is) is much higher. This really viscous magma acts as an effective plug on the vent and allows gas pressures to build to very high. Eventually the gas pressure is higher than even the viscous lava can stand, and an explosive eruption occurs. These explosions remove the cap of viscous lava that was plugging the vent so that the pressure is now lower. With the new low pressure, more gas bubbles can expand and push more lava out of the vent, and on and on and on. Once one of these explosive eruptions starts it pretty much continues until the available magma is used up. These big explosions reach 10’s of km into the atmosphere sometimes, and spread fine ash over huge areas.
Sometimes instead of going up, the hot mixture of gas and ash flows out of the vent and hugs the ground. These fast-moving hot mixtures are called pyroclastic flows and they are very dangerous. Because they are mostly gas, they can move quickly, up to 200 km/hour. They are sometimes up to 600 degrees centigrade. With this combination of speed and heat they are the most dangerous phenomenon that a volcano can produce. They may leave only a thin layer of ash after they pass through, but for those few moments while the pyroclastic flow is passing through nothing can live. Pyroclastic flows killed about 25,000 people in the town of St. Pierre in 1902. This disaster prompted Thomas A. Jaggar to dedicate his life to studying volcanoes, and he went on to found the Hawaiian Volcano Observatory.
What is a hot spot?
Mantle plumes are areas of hot, upwelling mantle. A hot spot develops above the plume. Magma generated by the hot spot rises through the rigid plates of the lithosphere and produces active volcanoes at the Earth’s surface. As oceanic volcanoes move away from the hot spot, they cool and subside, producing older islands, atolls, and seamounts. As continental volcanoes move away from the hot spot, they cool, subside, and become extinct.
Hot spots are places within the mantle where rocks melt to generate magma. The presence of a hot spot is inferred by anomalous volcanism (i.e. not at a plate boundary), such as the Hawaiian volcanoes within the Pacific Plate. The Hawaiian hot spot has been active at least 70 million years, producing a volcanic chain that extends 3,750 miles (6,000 km) across the northwest Pacific Ocean. Hot spots also develop beneath continents. The Yellowstone hot spot has been active at least 15 million years, producing a chain of calderas and volcanic features along the Snake River Plain that extends 400 miles (650 km) westward from northwest Wyoming to the Idaho-Oregon border.

Can volcanoes form just anywhere?
I guess you can think of 3 main places where volcanoes originate, hot spots, divergent plate boundaries(such as rifts and mid-ocean ridges), and convergent plate boundaries(subduction zones).
The origin of the magma for hot spots is not well known. We do know that the magma comes from partial melting within the upper mantle, probably from depths not too much greater than 100 km. The actual source of the heat that causes the partial melting (the actual hotspot itself) is almost certainly much deeper than that, but we really don’t know how deep or even exactly what a hotspot is!
At a divergent margin, two tectonic plates are moving apart, and magma that is generated in the upper mantle flows upward to fill in the space. This magma is probably generated at depths that are shallower than those for hotspot magmas. People argue about whether the magma forcing its way to the surface causes the plates to move apart or whether the plates move apart and the magma just reacts to that and fills in the space. Perhaps it is a combination of these two. The most extensive example of this type of volcanism is the system of mid-ocean ridges. Continental examples include the East African Rift, the West Antarctic Rift, and the Basin and Range Province in the southwestern US.
The final major place where volcanism originates is at convergent margins (subduction zones)–where an oceanic plate dives under either another oceanic plate or perhaps a continental plate. As the plate gets pushed further and further it starts to give off its volatiles (mostly water), and these migrate upwards into the mantle just under the overriding plate. The addition of these volatiles to this overriding mantle probably lowers the melting point of that mantle so that magma is generated. Part of the magma may also be generated by the downgoing plate actually starting to melt as it gets into the hotter and hotter interior.

How are volcanoes and earthquakes related?
Some, but not all, earthquakes are related to volcanoes. For example, most earthquakes are along the edges of tectonic plates. This is where most volcanoes are too. However, most earthquakes are caused by the interaction of the plates not the movement of magma.
Most earthquakes directly beneath a volcano are caused by the movement of magma. The magma exerts pressure on the rocks until it cracks the rock. Then the magma squirts into the crack and starts building pressure again. Every time the rock cracks it makes a small earthquake. These earthquakes are usually too weak to be felt but can be detected and recorded by sensitive instruments. Once the plumbing system of the volcano is open and magma is flowing through it, constant earthquake waves, called harmonic tremor, are recorded (but not felt).
There is a general relationship, in that volcanic eruptions and earthquakes occur along the same major areas of the world. . .because both occur most often at plate boundaries. . .such as around the edge of the Pacific Ocean.

What is a volcano?
A volcano is a vent in the surface of the Earth through which magma and associated gases and ash erupt; also, the form or structure, usually conical, that is produced by the ejected material.

What are the different parts of a volcano?
The basic parts are the magma supply system (which may or may not include a magma chamber), the plumbing system that gets magma from the magma chamber to the surface, the layers of lava and/or ash, and perhaps a summit crater or caldera.
How many active volcanoes are there in the world?
The Smithsonian Institution’s list of historical volcanic eruptions lists about 800 active or dormant volcanoes
What is the difference between active, dormant, and extinct volcanoes?
An active volcano is one that has erupted sometime during the last few hundred years.
A dormant volcano is one that has not erupted during the last few hundred years, but it has erupted during the last several thousand years.
An extinct volcano is one that has not erupted during the last several thousand years.
How do you make a model of a volcano?
Volcano world has an excellent section that for those interested in building volcano models . It has everything from simple volcanoes to ones that actually erupt! They are all pretty easy and cheap to make.

How did hot spots get formed?
Nobody really knows the answer. One thought is that they represent bumps on the surface that represents the boundary between the Earth’s outer core and its lower mantle. these bumps cause upward streaming of material that we call hotspots. The honest answer is that lots of folks are working on it but haven’t come up with the answer yet.
What part of the world do scientists call the “ring of fire”?
The “ring of fire” refers to the rim of the Pacific ocean, from southern Chile all the way around to New Zealand.

The earth is a dynamic planet. Its rigid outer surface layer is broken into several tectonic plates which are in constant motion relative to one another. As demonstrated in the world map below, most of the ~550 active volcanoes on earth are located along the margins of adjacent plates.

Tectonic plates are composed of lithosphere, the rigid outer portion of the earth. With a thickness of about 100 km, the lithosphere is composed of an upper layer of crust (~7 km thick under the oceans, and ~35 km thick under the continents) and a lower, denser layer of the earth’s upper mantle. The lithosphere is underlain by the asthenosphere, a hot, mobile layer of partially molten rock lying within the earth’s upper mantle. (For detailed information, click the Earth’s Interior.)
Volcanic eruptions above these lithospheric plates are driven by the ascent of magma (molten rock) from deep beneath the surface. The various magma types are described in Physicochemical Controls on Eruption Style. They vary from mafic, intermediate, to felsic as their silica (SiO2) content increases. Mafic (basaltic) magmas are generated directly from the mantle, either within the asthenosphere or within the overlying mantle lithosphere. Many mafic-to-intermediate (basaltic-to-andesitic) magmas appear to be derived from the melting of hydrated lithospheric mantle. More differentiated, intermediate-to-felsic magmas, on the other hand, are partly derived from the melting of continental crust by hot, mafic magmas that either pond at the crust-mantle boundary, or intrude into the overlying continents where they reside in magma chambers located at various crustal levels.
Volcanism is typically widespread along plate boundaries. Although volcanism in the interior of plates is less common, these intraplate regions can also generate voluminous eruptive products. The regional volcano-tectonic processes associated with plate-boundary environments and intraplate environments are described in more detail below.

Plate boundaries mark the sites where two plates are either moving away from one another, moving toward one another, or sliding past one another. Adjacent plates are delineated by three types of boundaries defined by this relative motion:
• Divergent plate boundaries — Plates diverge from one another at the site of thermally buoyant mid-oceanic ridges. Oceanic crust is created at divergent plate boundaries.
• Convergent plate boundaries — Plates converge on one another at the site of deep oceanic trenches. Oceanic crust is destroyed at convergent plate boundaries.
• Transform plate boundaries — Plates slide past one another.
Although volcanism is abundant at divergent and convergent plate boundaries, there is a distinct lack of significant volcanism associated with transform plate boundaries. Spreading center volcanism occurs at divergent plate margins, and subduction zone volcanism occurs at convergent plate margins. Intraplate volcanism describes volcanic eruptions within tectonic plates.

The most volcanically active belt on Earth is known as the Ring of Fire, a region of subduction zone volcanism surrounding the Pacific Ocean. Subduction zone volcanism occurs where two plates are converging on one another. One plate containing oceanic lithosphere descends beneath the adjacent plate, thus consuming the oceanic lithosphere into the earth’s mantle. This on-going process is called subduction. As the descending plate bends downward at the surface, it creates a large linear depression called an oceanic trench. These trenches are the deepest topographic features on the earth’s surface. The deepest, 11 kilometers below sealevel, is the Mariana trench, which lies along the western margin of the Ring of Fire. Another example, forming the northern rim of the Ring of Fire, is the Aleutian trench.
The crustal portion of the subducting slab contains a significant amount of surface water, as well as water contained in hydrated minerals within the seafloor basalt. As the subducting slab descends to greater and greater depths, it progressively encounters greater temperatures and greater pressures which cause the slab to release water into the mantle wedge overlying the descending plate. Water has the effect of lowering the melting temperature of the mantle, thus causing it to melt. The magma produced by this mechanism varies from basalt to andesite in composition. It rises upward to produce a linear belt of volcanoes parallel to the oceanic trench, as exemplified in the above image of the Aleutian Island chain. The chain of volcanoes is called an island arc. If the oceanic lithosphere subducts beneath an adjacent plate of continental lithosphere, then a similar belt of volcanoes will be generated on continental crust. This belt is then called a volcanic arc, examples of which include the Cascade volcanic arc of the U.S. Pacific northwest, and the Andes volcanic arc of South America.

Although most volcanic rocks are generated at plate boundaries, there are a few exceptionally active sites of volcanism within the plate interiors. These intraplate regions of voluminous volcanism are called hotspots. Twenty-four selected hotspots are shown on the adjacent map. Most hotspots are thought to be underlain by a large plume of anomalously hot mantle. These mantle plumes appear to be generated in the lower mantle and rise slowly through the mantle by convection. Experimental data suggests that they rise as a plastically deforming mass that has a bulbous plume head fed by a long, narrow plume tail. As the head impinges on the base of the lithosphere, it spreads outward into a mushroom shape. Such plume heads are thought to have diameters between ~500 to ~1000 km.

Many scientists believe that mantle plumes may be derived from near the core-mantle boundary, as demonstrated in this computer simulation from the Minnesota supercomputing lab. Note the bulbous plume heads, the narrow plume tails, and the flattened plume heads as they impinge on the outer sphere representing the base of the lithosphere.

Decompressional melting of this hot mantle source can generate huge volumes of basalt magma. It is thought that the massive flood basalt provinces on earth are produced above mantle hotspots. Although most geologists accept the hotspot concept, the number of hotspots worldwide is still a matter of controversy.

The Pacific plate contains several linear belts of extinct submarine volcanoes, called seamounts, an example of which is the Foundation seamount chain
The formation of at least some of these intraplate seamount chains can be attributed to volcanism above a mantle hotspot to form a linear, age-progressive hotspot track. Mantle plumes appear to be largely unaffected by plate motions. As lithospheric plates move across stationary hotspots, volcanism will generate volcanic islands that are active above the mantle plume, but become inactive and progressively older as they move away from the mantle plume in the direction of plate movement. Thus, a linear belt of inactive volcanic islands and seamounts will be produced. A classic example of this mechanism is demonstrated by the Hawaiian and Emperor seamount chains.
The “Big Island” of Hawaii lies above the mantle plume. It is the only island that is currently volcanically active. The seven Hawaiian Islands become progressively older to the northwest. The main phase of volcanism on Oahu ceased about 3 million years ago, and on Kauai about 5 million years ago. This trend continues beyond the Hawaiian Islands, as demonstrated by a string of seamounts (the Hawaiian chain) that becomes progressively older toward Midway Island. Midway is composed of lavas that are ~27 million years old. Northwest of Midway, the volcanic belt bends to the north-northwest to form the Emperor seamount chain. Here, the seamounts become progressively older until they terminate against the Aleutian trench. The oldest of these seamounts near the trench is ~70 million years old. This implies that the mantle plume currently generating basaltic lavas on the Big Island has been in existence for at least 70 million years!
The Hawaiians were very good at recognizing the difference in the older, eroded volcanic islands and newer islands to the southeast, where volcanic features are more pristine. Legend has it that Pele, the Hawaiian goddess of fire, was forced from island to island as she was chased by vairous gods. Her journey is marked by volcanic eruptions, as she progressed from the island of Kaua’i to her current home on the Big Island. The legend corresponds well with the modern scientific notion of the age progression of these volcanic islands.

The Earth’s internal heat source provides the energy for our dynamic planet, supplying it with the driving force for plate-tectonic motion, and for on-going catastrophic events such as earthquakes and volcanic eruptions. This internal heat energy was much greater in the early stages of the Earth than it is today, having accumulated rapidly by heat conversion associated with three separate processes, all of which were most intense during the first few hundred thousand years of the Earth’s history: (1) extraterrestrial impacts, (2) gravitational contraction of the Earth’s interior, and (3) the radioactive decay of unstable isotopes.

Most scientists believe that our solar system evolved from the accretion of solid particles derived from a large nebular cloud – the so-called Nebular Hypothesis. Under this scenario, proto-planet Earth would have grown over time from a barrage of extraterrestrial impacts, increasing its mass with each bombardment. As the proto-planet grew in size its increased gravitational field would have attracted even more objects its surface. The composition of these colliding bodies would have included metal-rich fragments (i.e.., iron meteorites), rocky fragments (i.e., stony meteorites), and icy fragments (i.e., comets). Although accretion was much more prevalent in the early stages of the Earth’s history, these extraterrestrial collisions are still occurring today, exemplified by shooting stars and fireballs in the night sky, and by the occasional impact of larger bodies on the Earth’s surface.
Such particles travel at great velocities, typically ~30,000–50,000 km/hr, similar to that of the Earth as it rotates around the Sun. The very large amount of kinetic energy inherent in these moving bodies is instantly converted to heat energy upon impact, thus providing a component to the Earth’s internal heat source.

In the early stages of planetary accretion, the earth was much less compact than it is today. The accretionary process led to an increasingly greater gravitational attraction, forcing the Earth to contract into a smaller volume. Increased compaction resulted in the conversion of gravitational energy into heat energy, much like a bicycle pump heats up due to the compression of air inside it. Heat conducts very slowly through rock, so that the rapid build up of this heat source within the Earth was not accommodated by an equally rapid loss of heat through the surface.

Radioactive elements are inherently unstable, breaking down over time to more stable forms. The unstable isotope Uranium-238, for example, will slowly decay to Lead-206. All such radioactive decay processes release heat as a by product of the on-going reaction. In its early stages of formation, the young earth had a greater complement of radioactive elements, but many of these (e.g., aluminum-26) are short-lived and have decayed to near extinction. Others with a more lengthy rate of decay and are still undergoing this radioactive process, thus still releasing heat energy. The greater complement of unstable elements in the early Earth thus generated a greater amount of heat energy in its initial stages of formation.

The heat buildup inside earth reached a maxim early in the Earth’s history and has declined significantly since. The greater heat content of the early Earth was the product of (1) a greater abundance of radioactive elements, (2) a greater number of impacts, and (3) the early gravitational crowding. The initial accretion of particles resulted in a rather homogeneous sphere composed of a loose amalgam of metallic fragments (iron meteorites), rocky fragments (stony meteorites), and icy fragments (comets). However, the increased heat content of the early Earth resulted in melting of the Earth’s interior, so that the young planet became density stratified with the heavier (metallic) materials sinking to the center of the earth, and the lighter (rocky) materials floating upward toward the surface of the earth. The very lightest volatile materials (derived from comets) were easily melted or vaporized, rising beyond the earth’s rocky surface to form the early oceans and the atmosphere. We now have a differentiated earth due to melting and mobilization of materials driven by the earth’s internal heat engine. This has resulted in the development of a series of concentric layers that are both density and compositionally stratified. This demonstrated in the diagram below, courtesy of the USGS.
These layers include (1) the dense inner core composed largely of solid Fe and subordinate Ni, with radius of about 1200 km, (2) the molten outer core composed largely of liquid Fe, with subordinate sulfur, with a radius of about 2250 km, (3) the mantle, composed of relatively dense rocky materials, with radius of about 2800 km thick, and (4) the crust which comprises the thin relatively light outer skin of the earth, is divisible into two types: the oceanic crust (~7 km thick) and the continental crust (about 35 km thick). Whereas oceanic crust is composed of basaltic rock, the less dense continental crust is composed of a great variety of rock types having an overall average composition akin to granite.
Within the mantle exists the asthenosphere (Grk. asthenos = weak), between about 100 km and 350 km, which is a special zone composed of hot, weak material that is capable of gradual flow. The layer above the asthenosphere is the lithosphere (Grk. lithos = rock), the rigid and relatively cool outer layer of the earth, composed of both crust and a portion of the upper mantle.
Lying above the lithosphere is (1) the liquid hydrosphere, comprising 71% of the Earth’s surface, and (2) that the still lighter gaseous atmosphere, both of which were ultimately derived from the accretion of comets. The occurrence of these volatile components along the outermost portion of the Earth is a product of volcanic outgassing during the differentiation event.


There is a great range in the explosivity of volcanic eruptions. Many eruptions are relatively quiescent and are characterized by the calm, nonviolent extrusion of lava flows on the earth’s surface. Other eruptions, however, are highly explosive and are characterized by the violent ejection of fragmented volcanic debris, called tephra, which can extend tens of kilometers into the atmosphere above the volcano.
Whether or not an eruption falls into one of these end-member types depends on a variety of factors, which are ultimately linked to the composition of the magma (molten rock) underlying the volcano. Magma composition is discussed below, followed by a description of the controlling factors on explosivity — viscosity, temperature, and the amount of dissolved gases in the magma.

Only ten elements make up the bulk of most magmas: oxygen (O), silicon (Si), aluminum (Al), iron (Fe), magnesium (Mg), titanium (Ti) calcium (Ca), sodium (Na), potassium (K), and phosphorous (P). Because oxygen and silicon are by far the two most abundant elements in magma, it is convenient to describe the different magma types in terms of their silica content (SiO2). The magma types vary from mafic magmas, which have relatively low silica and high Fe and Mg contents, to felsic magmas, which have relatively high silica and low Fe and Mg contents. Mafic magma will cool and crystallize to produce the volcanic rock basalt, whereas felsic magma will crystallize to produce dacite and rhyolite. Intermediate-composition magmas will crystallize to produce the rock andesite. Because the mafic rocks are enriched in Fe and Mg, they tend to be darker colored than the felsic rock types.
~50% Mafic Basalt
~60% Intermediate Andesite
~65% Felsic (low Si) Dacite
~70% Felsic (high Si) Rhyolite
There also exists more unusual magmas that erupt less commonly on the Earth’s surface as ultramafic, carbonatite, and strongly alkaline lavas.

One way to classify lavas is by their alkali content, reflected in their weight percent of Na2O + K2O. In contrast to the common lava types (basalt, andesite, dacite, and rhyolite), there exists less common lavas that define mildly alkaline trends (e.g., with increasing silica content: alkali basalt, trachybasalt, trachyandesite, trachyte, and comendite), and strongly alkaline trends (e.g., with increasing silica content: tephrite, phonotephrite, tephriphonolite, and phonolite). Although these lavas can occur in a variety of tectonic settings, they are typically found in either (1) continental or oceanic intraplate settings, where there is often a lack of significant tectonic control, (2) continental rift zones, and (3) the back-arc setting of subduction zones.
Alkali-rich lavas are often charateristic of the waning stages of volcanism, as demonstrated, for example, in the late-stage parasitic cones and flows found on Hawaiian shield volcanoes. Differentiated trachytic and phonolitic lavas typically erupt as low-volume flows with high aspect ratios, as demonstrated above in the phonolite coulées of western Saudi Arabia. However, extensive sheetflows of similar lavas have been recognized in continental rift zones, as exemplified in the voluminous phonolitic lavas of the Ethiopian rift system.

Carbonatites are perhaps the most unusual of all lavas. They are defined, when crystalline, by having more than 50% carbonate (CO3-bearing) minerals, and typically they are composed of less than 10% SiO2. There are only 330 known carbonatite localities on Earth, most of which are shallow intrusive bodies of calcite-rich igneous rock in the form of volcanic necks, dikes, and cone-sheets. These generally occur in association with larger intrusions of alkali-rich silicate igneous rocks. Extrusive carbonatites are particularly rare, and appear to be restricted to a few continental rift zones, such as the Rhine valley and the East African rift system.
Most carbonatite lavas have low eruption temperatures, between 500 and 600 degrees Centigrate (compared with >1100 degrees Centigrate for basaltic lavas). They typically have low viscosities due to their lack of silica polymerization. Thus, carbonatite flows are generally only a few centimeters thick, with surface textures that vary from a’a to pahoehoe. Although they often resemble flowing lobes of black mud, they are hot enough to display glowing, deep-red colors when seen at night. Some active carbonatite flows are enriched in alkalies (Na and K) – these are called natrocarbonatites. Soon after their eruption, dark natrocarbonatite flows will cool and rapidly turn white due to reaction with atmospheric water.

The East African Rift System contains a vast array of igneous rocks, dominated by voluminous floods of basaltic lava. Its highest peaks, Mt. Kilamanjoro and Mt. Kenya, are active volcanoes located along the eastern rift of Lake Victoria. Lying between these impressive peaks is the only know active carbonatite volcano in the world – Oldoinyo Lengai. Rising over 2200 meters from the valley floor, the bulk of this volcanic cone is composed phonolitic tephra. However, its upper portion is dominated by natrocarbonatite lava flows. Historic eruptions of natrocarbonatite have filled much of the summit crater, shown here courtesy of Marco Fulle – stromboli.net. Note the tall conical hornitos protruding from the crater floor.

A history of controversy has surrouned the origin of carbonatites, and whether or not calcite carbonatites are primary or secondary lavas. Some scientists believe that the Ca-carbonatites are generated by fractional melting of crustal carbonate rocks. However, others believe that the Ca-carbonatites are not primary magmas at all, but rather derived from the alteration of Na-rich natrocarbonatite lavas. Virtually, all natrocarbonatites on Earth are found in historic eruptions, and none in ancient rocks. Many have attributed this to the very high solubility of sodium carbonate. The argument being that all carbonatites were originally natrocarbonatites that have had sodium removed by hydrothermal solutions and rainwater.
The natrocarbonatites themselves appear to have clear isotopic signatures indicating their origin from a mantle source. The very high Na2O and very low MgO compositions of the Oldoinyo Lengai natrocarbonatites suggest that they are highly differentiated magmas. The common association of carbonatites with alkali-rich silicate rocks (e.g., nephelinite) suggests further that liquid immiscibility between silicate and carbonate magmas may have played a role in their origin.

Named from their type locality along the Komati River in South Africa, komatiites are ultramafic volcanic rocks, having very low silica contents (~40-45%) and very high MgO contents (~18%). These lavas are exceptional not only for their compositions, but also for their very old, restricted ages. These lavas have no modern analogs. The youngest komatiites (from Gorgona Island, Columbia) have been dated at about 90 million years; however, all other komatiites are about three billion years old or older. These ancient lava flows erupted at a time when the Earth’s internal heat was much greater than today, thus generating exceptionally hot, fluid lavas with calculated eruption temperatures in excess of 1,600 degrees C (2,900 degrees F). In comparison, typical basaltic lavas erupting today have eruption temperatures of about 1,100 degrees C.
A spectauclar identifying trait of komatiites is their spinifex texture, which resembles a lacey mesh of acicular (needle-like) olivine crystals, typically surrounded by lighter-colored, interstitial minerals such as plagioclase, tremolite, and/or chlorite. The elongated nature of the komatiite olivine crystals is quite distinct when compared with the equant to tabular olivine crystals seen in most basaltic rocks.
Since we have never observed a komatiite eruption, we have had to deduce the fluid flow character and eruption style of these lavas from the properties and textures of the ancient rocks. Their high eruption temperatures, for example, are calculated from their olivine-rich compositions. The charateristic spinifex textures on the otherhand, indicate that these lavas cooled very rapidly. Combined, the high eruption temperatures and the low silica contents indicate that komatiites erupted as very fluid lavas, having exceptionally low viscosities and low aspect ratios. It is believed that these hot, fluid lavas would have been turbulent, and therefore capable of a significant amount of both mechanical and thermal erosion. Indeed, many scientists believe that the sinuous rilles exposed on the moon’s surface are erosive valleys produced by komatiite lava flows, contemporaneous with the ancient eruption of similar lavas on Earth.

Volumetrically, most lava is of basaltic composition. Basaltic melts have overall lower gas contents and are more fluid than their andesitic-to-rhyolitic counterparts. Their higher fluidity (lower viscosity) is a product of their lower SiO2 (silica) contents. When gases exsolve from basaltic melts they are allowed to rise unimpeded through the fluid magma without a significant build up of gas pressure. This results in relatively calm, nonexplosive eruptions, and a preponderance of lava. In contrast, when gases exsolve from felsic magmas, their upward mobility is impeded by the high viscosity of the melt. This results in the buildup of gas pressure, which generates explosive eruptions associated with a preponderance of pyroclastic ejecta. The low viscosity of basaltic lavas allows them to be extruded over great distances, often producing high-volume lava flows with low aspect ratios (ratio of thickness to area). Under the right conditions, de-gassed felsic magmas can also erupt lava in a nonviolent manner. However, felsic lavas tend to be much thicker than basaltic lavas and have much higher aspect ratios.

The volume of magma generated over a given amount of time is known as the effusion rate. The effusion rates for historical eruptions of basaltic lava are highly variable, from 0.5 to 5000 m3/sec. Historic flows from Mt. Etna average about 0.5 m3/sec, whereas the fissure-generated flows associated with the Icelandic Laki eruption in 1783 were released at a rate of ~5000 m3/sec. The effusion rates for andesite and dacite are much lower (~10 to 0.05 m3/sec) due to their higher viscosities.

Lava is highly variable in composition and it therefore varies greatly in flow character. Once crystallized, this compositional diversity is reflected in variations in rock texture, mineralogy, and volcanic landforms. Click on the links below for a more detailed description of the compositional lava types:

Basaltic lava flows erupt primarily from shield volcanoes, fissure systems, scoria cones, and spatter cones. These fluid lava flows can be subdivided into two end-member structural types, based primarily on the nature of lava flow surfaces:
• Pahoehoe Lava — Surfaces are smooth, billowy, or ropy.
• A’a lava — Surfaces are fragmented, rough, and spiny, with a “cindery” appearance.

Pahoehoe can take several different forms. As the smooth lava surface cools to turns to a dark gray color and becomes less fluid and more viscous, behaving more like a plastic substance than a truly liquid substance. As lava continues to flow underneath this plastic skin, the surface can bunch up or wrinkle into a form that resembles coiled rope. Such a surface is called ropy pahoehoe. In addition to these ropy surfaces, solidified basalt flows can also display shelly to slabby surfaces. Shelly pahoehoe contains a billowy flow top with a frothy vesicular surface skin, only a few centimeters thick, overlying large cavities, generally 5-30 centimeters thick. These shelly surfaces often collapse when walking on the top of the flow. Slabby pahoehoe contains a series of closely spaced slabs, a few meters across and a few centimeters thick, broken and tilted by mass movement, or drainage, of the underlying lava. Slabby pahoehoe is often gradational to a’a lava.

Pahoehoe lavas are typically the first to erupt from a vent. They are relatively thin (1-2 m) and very fluid with low viscosities. They advance downslope in a sort of smooth “rolling motion.” The front of the flow usually advances as a thin (< 20 cm) glowing lobe that will chill and crust over after 1-2 meters of flow. It will slow and be overrun by a new lobe that propagates downslope until it also chills, and in turn is overrun by another flow. Overriding lavas and breakouts on the flow top and sides thus produce compound flows composed of several lobes cooling against one another. Slower moving pahoehoe flows will advance through the protrusion of small bulbous appendages at the flow front, called pahoehoe toes. The image above shows a breakout and the advancement of pahoehoe toes along the sides of a ropy pahoehoe lava flow. As the lava surface cools and thin skin becomes more viscous, progressive breakouts will occur, thus advancing the flow forward, as demonstrated below. Where pahoehoe toes advance rapidly, usually down steeper slopes, an elongated protrusions may emerge, called entrail pahoehoe.

High effusion rates may result in the development of inflated pahoehoe sheetflows. These large-volume pahoehoe flows are emplaced initially as thin sheets, 20-30 cm thick. Cooling of these flow sheets will produce a smooth pahoehoe crust which initially behaves in a plastic fashion. However, after attaining a thickness of 2-5 cm, the crust behaves more rigidly and develops strength. As the underlying liquid core of the flow increases in size due to sustained lava injection, the hydrostatic head of the flow is distributed evenly throughout. This can result in flow inflation and uniform uplift of the entire flow sheet. Such flows will inflate and thicken by as much as 4 meters. Many of these inflated flows will crystallize in place; however, some will deflate after emplacement as the fluid core drains beneath the solidified crust. Deflation is evident in many Hawaiian flows which contain tree molds that currently stand 1-2 meters above the current surface of the pahoehoe sheet flows. An example of deflation on Kilauea Volcano, Hawaii, is demonstrated to the left by solidified lava around high-standing tree trunks to produce hollow tree molds standing above the deflated pahoehoe surface.

The a’a flows shown below are advancing over older pahoehoe surfaces. Although these lava flows are often more viscous, and typically thicker, than pahoehoe lavas they tend to advance at greater rates. Their flow fronts can vary from two meters to as much as twenty meters thick. Their forward motion is similar to the movement of a tractor tread. A jumbled mass of debris steepens at the flow front until a section breaks off and tumbles forward. Inward from the flow front, the flow usually contains an upper rubbly flow top, and a lower massive part of viscous lava insulated from the overlying rubble. The jagged cinder blocks that break off the front are then overridden by the massive lava core of the flow which pushes forward.

Pahoehoe is often converted to a’a as lava advances downslope, away from the volcano. Conversion of a’a to pahoehoe, on the other hand, never takes place. The pahoehoe-to-a’a conversion can be caused by either an increase in flow viscosity, or an increase in the rate of shear. Cooling, gas loss, and crystallization of the lava will cause it to become increasingly more polymerized and viscous as it advances farther downslope. Although the downslope increase in viscosity can convert pahoehoe to a’a, a higher rate of shear can also cause the conversion. Shear increases in flows that have higher effusion rates. A’a begins to form when effusion rates are >5-10 cubic meters per second. The rate of shear can also increase as lava advances down increasingly steeper slopes. This is demonstrated, for example, by the conversion of pahoehoe to a’a in recent lava flows that have advanced down the Hilina Pali escarpment during the ongoing eruption of the Pu’u O’o Volcano, Hawaii.

Whereas basalt forms a’a and pahoehoe surface forms, andesite generally produces blocky lava. Here, the surface contains smooth-sided, angular fragments (blocks) that are not as splintery or vesicular as a’a lava fragments. The blocky nature of these flows is attributed to the higher viscosity of andesite. These viscous lavas have relatively high aspect ratios (thickness/area), generally > 1/100, and some are thick enough to form as lava domes. Andesite commonly erupts from stratovolcanoes, where they form small-volume flows that typically advance only short distances down the flanks of a volcano. The two examples shown here are short andesite flows advancing down the slope of the Lascar volcano in Chile, and the Colima volcano in Mexico.
Although detached blocks occur on the tops of andesite flows, the flow interior is composed of massive lava which grades downward into an autobrecciated (self-fragmented) basal layer. The flow moves by the injection of lobes of lava into the collar of blocky rubble that comprises the flow front. The flow-front rubble is then continually overridden by the massive lava core to form the fragmented basal layer of the flow.

With increasing silica (SiO2) content and polymerization, the viscosities of these felsic lavas will increase accordingly. Although dacitic-to-rhyolitic lavas typically erupt from stratovolcanoes, they are not as abundant as andesite lava. Instead, felsic eruptions from stratovolcanoes are more commonly explosive and associated with the generation of tephra and pyroclastic flows. These explosive eruptions are a function of the high viscosities and high gas contents of dacitic and rhyolitic magmas. Such eruptions, however, will often deplete the magma source in dissolved gases. De-gassed magma can then rise to the surface and extrude in a less violent fashion, as dacite to rhyolite lava. These lavas are volumetrically smaller than their pyroclastic counterparts, and typically form after major eruptive events. Viscous dacitic-to-rhyolitic lavas generally ooze out of the volcano’s central vent to form symmetrical lava domes. In some cases, however, asymmetric flow down one side of the vent can produce an elongated extrusion, called a coulée or a dome flow. These viscous flows vary from 50 to 500 meters in thickness, and are typically only a few kilometers long. However, some coulées may flow for several kilometers, probably due to high eruption temperatures, which would lower their viscosities (also see western Arabian phonolite coulées – one and two).

As demonstrated above, felsic magmas will either (1) erupt explosively to produce extensive deposits of tephra, or (2) nonexplosively to produce degassed, viscous lava (domes, coulees, or obsidian flows) which advance only short distances from their vents. There has been a significant amount of controversy, therefore, over rare rhyolite lavas that appear to occur as large-volume flows (10-100 cubic kilometers).
Most such flows occur near continental hotspots. The best known examples are those associated with (1) the Yellowstone hotspot track near the Idaho-Oregon border, and (2) the Ethiopian hotspot in northeastern Africa. These large-volume felsic volcanic rocks have outcrop, hand specimen, and thin section characteristics typical of lava flows. However, many volcanologists suspect that they are not lava flows at all, but rather rheomorphic ignimbrites. These are densely welded pyroclastic flows of pumice and ash, which were thick and hot enough to flow downslope and obliterate primary pyroclastic structures. They suggest that the original pumice and ash fragments have been streaked out like toffee strands so that the pyroclastic nature of the flow becomes unrecognizable.
Still, other workers are convinced that these voluminous rhyolites are in fact lava flows, and that their extensive distribution can be attributed to their unique position above a continental hotspot. They suggest that the widespread melting of continental crust above the hotspot could generate superheated, low-viscosity melts that would erupt in a fluid fashion to produce voluminous sheets of rhyolitic lava. The controversy is still unresolved.

The single most important factor on the form of a lava flow is probably the effusion rate, or rate of discharge, measured in cubic meters per second. The higher the effusion rate, the greater the distance traveled before cooling lowers the viscosity and impedes flow. The effusion rates for andesite to dacite lavas (10 to 0.05 cubic meters per second) are a few orders of magnitude lower than those for basalt (0.5 to 5000 cubic meters per second). Thus, basalt lavas are generally much more extensive than the more felsic rock types. Basaltic lavas extruded at relatively low effusion rates generally produce a number of small flow units extruded closely in time so that they overrun each other in anastomising channels. Upon cooling, these flow units produce compound flows. Basalt lavas with higher effusion rates, however, produce extensive flows, often extruded as large-volume flowsheets. These flows are composed of a single cooling unit and are called simple flows. The high effusion rates associated with flood basalt provinces suggest that they contain large-volume simple flows.

As a lava flow cools and crystallizes into a coherent rock, it begins to contract. Shrinkage of the rock mass results in the development of numerous cracks, called joints. Although lava flows can display a variety of jointing patterns, many of the thicker simple flows exhibit a three-tiered character: from bottom to top, a lower colonade, a middle entablature, and an upper colonade. The lower colonade is composed of columnar joints, which develop perpendicular to the cooling surface, at the base of the flow, where they intersect to form a series of columns. The columns may vary in length from one to five meters, with diameters that are generally less than one meter. Each column is polygonal in cross-section (typically hexagonal) and bounded by 4-to-8 joints. If present, the entablature is composed of an array of closely spaced, subvertical joints, which forms a more convoluted pattern of cracks known as hackly jointing. The upper colonade is generally poorly develped or non-existent. If present, the crude columnar joints are vesicular and much shorter than those displayed in the lower colonade.

Although highly effusive eruptions may advance downslope as massive sheets of basaltic lava, such flows are rare in the historic record. Fluid basalt will more commonly move down slope by creating its own channelways above gently sloping terrains, or by flowing down in pre-existing stream channels. A well-developed lava channel is shown here near a series of eruptive vents on the northeast rift zone of the Mauna Loa Volcano, Hawaii. Lava channels can develop in both pahoehoe and a’a basalt. Once established, lava channels will often develop lava levees along their lengths. In pahoehoe flows, natural levees will be constructed as the channel periodically overflows, thus allowing the lava to congeal and cool along the banks of the channel. In a’a flows, levee buildup can also occur by the bulldozing effect of the moving lava, which will push a’a blocks onto the lava banks. Although lava channels may initially be restricted to a topographic valley, as levees build up over time, they may eventually rise above the surrounding ground surface. Sometimes several small lava channels will flow around subdued hills or high ground to produce a kipuka, an Hawaiian term for “island.” In Hawaii, these features are generally noted by their mature vegetation, which stands in contrast to the stark lack of vegetation in the younger, surrounding basalt flows.
One characteristic of a’a channels is the occurrence of nearly spherical masses of lava with diameters of a few inches to over ten feet. These features are called accretionary lava balls. When broken apart, they reveal a spiral structure which appears to form when a small fragment of solidified lava rolls along the surface of an a’a flow. The lava then adheres to the rolling mass, much like snow sticks to a rolling snowball.

It is not unusual for lava to accumulate in volcanic craters, filling the craters to a high level to generate lava lakes. If the effusion rate is constant and high, the lava may rise above the crater rim, breaching the crater and spilling out onto the adjacent surface. One example of this is the Kupaianaha lava lake in Hawii, which overtopped it’s crater in 1986. A much older example shown in the image to the left. This is the Jabal Hil scoria cone, located on the Harrat Kishb lava field, about 150 kilometers northeast of Mecca in western Saudi Arabia. The Jabal Hil flow field is Post-Neolithic in age and composed of anastomosing pahoehoe and a’a lava flows that have cascaded down from the crater rim due to overspilling of a crater lake.
On flat or gentle slopes, the surfaces of pahoehoe flows are sometimes marked by elliptical domed structures, from 2 to 10 meters high, called tumuli. These are best developed on the surface of flows that are ponded in depressions, like craters or calderas. A tumulus develops when slow-moving lava beneath a solidified crust wells upward. The brittle crust usually buckles to accommodate the inflating core of the flow, thus creating a central crack along the length of the tumulus. These structures sometimes grade into elongate forms called pressure ridges, which commonly develop subparallel to the flow direction at flow margins, but perpendicular to the flow direction in the central portions of the flow. Localized upheaval of the underlying lava to produce tumuli can result from a variety of mechanisms, including (a) obstruction of lateral flow downstream behind a slow-moving lava front, (b) diversion of flow upward against an irregular topographic surface, (c) local upheaval from the volatilization of groundwater, and (d) deflation of an inflated pahoehoe sheet flow leading to differential sagging and crustal buckling.

Rootless eruptions are not connected at depth to a magma chamber, but rather result from surface eruptions on pahoehoe surfaces. When the pahoehoe crust thickens and the underlying lava becomes cool, viscous, and gas-depleted, pasty lava can squeeze up through the axial fracture of the tumulus like cold toothpaste. This process creates a protrusion of cheesy basalt commonly called a squeeze-up. Squeeze-ups can occur as bulbous mounds on pahoehoe surfaces, or as linear wedges in the axial core of tumuli. It the lava remains hot and liquid, local extrusions through the cracked pahoehoe surfaces can produce rootless cones of spatter, called hornitos, or driblet cones. Such features can erupt through the axial cracks of tumuli as well as through cracks and openings overlying active lava tubes.

The chilling and crystallization of basaltic lava around the sides, bottom, and top of lava channels produces a rock-encased conduit called a lava tube. The surrounding crystalline basalt remains relatively hot and insulates the interior lava in the lava tube from further crystallization. Although lava tubes are generally restricted to pahoehoe lava flows, they can form in some thick a’a flows, like those that occur on Mt. Etna. The tube morphology provides a very efficient mechanism for basaltic lava flows to travel great distances away from their source without significant heat loss. Their lateral extents are highly variable. Although lava tubes are typically < 1 km long, they may exceed several kilometers, as demonstrated, for example, by a lava tube in Queensland, Australia that is > 100 km long. The diameters of lava tubes are also highly variable, from < 1m to as much as 15m.
High lava temperatures of about 1150 degrees Centigrade (~ 2100 degrees Fahrenheit) are maintained throughout the tube system. These high temperatures allow the lava to remain very fluid. The high fluidity of lava in lava tubes and lava channels leads to high transport rates. During the 1984 eruption of Mauna Loa, for example, channel flow was measured by USGS volcanologists at 35 mph. Lava speeds within the current tube system associated with the Pu’u O’o eruption have been measured up to 23 mph. This fast moving, turbulent lava, combined with the high lava temperatures, can result in thermal erosion of the tube interior. Once established, well-insulated lava tubes can thus erode both downward and laterally over time. The turbulence from fluid motion and from constantly escaping gases generates spatter fragments that coat the walls of many lava tubes. Eventually, the eruption will cease and the remaining lava in the tube will drain downslope to expose a hollow, inactive lava tube. Such tubes typically contain a flat floor and exhibit high-lava marks on the conduit walls. A thin coating of agglutinated spatter is often found on the tube walls, together with icicle-shaped lava stalactites hanging from the roof. Mound-like lava stalagmites derived from spatter dripping onto the floor of tubes are less common. The example shown here is the well-known Thurston lava tube near the Kilauea summit caldera in Hawaii Volcanoes National Park (Photo courtesy of the USGS).
Because lava tubes remain buried, they often go unrecognized on the surface. However, many lava tubes are delineated on the surface by a linear or curvilinear series of collapse depressions, called skylights along the axis of the tube. On active lava tubes, skylights provide a unique view of actively flowing incandescent lava in the lava tube system. The example shown here from a skylight developed above an active lava tube associated with the Kupaianaha eruption in 1989 on the east rift system of the larger Kilauea volcano, Hawaii.

The viscosity of a substance is a measure of its consistency. Viscosity is defined as the ability of a substance to resist flow. In a sense, viscosity is the inverse of fluidity. Cold molasses, for example, has a higher viscosity than water because it is less fluid. A magma’s viscosity is largely controlled by its temperature, composition, and gas content (see downloadable programs at the bottom of this page). The effect of temperature on viscosity is intuitive. Like most liquids, the higher the temperature, the more fluid a substance becomes, thus lowering its viscosity.
Composition plays an even greater role in determining a magma’s viscosity. A magma’s resistance to flow is a function of its “internal friction” derived from the generation of chemical bonds within the liquid. Chemical bonds are created between negatively charged and positively charged ions (anions and cations, respectively). Of the ten most abundant elements found in magmas (see above), oxygen is the only anion. Silicon, on the other hand, is the most abundant cation. Thus, the Si-O bond is the single most important factor in determining the degree of a magma’s viscosity. These two elements bond together to form “floating radicals” in the magma, while it is still in its liquid state (i.e., Si-O bonds begin to form well above the crystallization temperature of magma). These floating radicals contain a small silicon atom surrounded by four larger oxygen atoms (SiO4). This atomic configuration is in the shape of a tetrahedron. The radicals are therefore called silicon-oxygen tetrahedra, as shown here.

These floating tetrahedra are electrically charged compounds. As such, they they are electrically attracted to other Si-O tetrahedra. The outer oxygen atoms in each tetrahedron can share electrons with the outer oxygen atoms of other tetrahedra. The sharing of electrons in this manner results in the development of covalent bonds between tetrahedra. In this way Si-O tetrahedra can link together to form a variety shapes: double tetrahedra (shown here, C), chains of tetrahedra, double chains of tetrahedra, and complicated networks of tetrahedra. As the magma cools, more and more bonds are created, which eventually leads to the development of crystals within the liquid medium. Thus, the Si-O tetrahedra form the building blocks to the common silicate minerals found in all igneous rocks. However, while still in the liquid state, the bonding of tetrahedra results in the polymerization of the liquid, which increases the “internal friction” of the magma, so that it more readily resists flow. Magmas that have a high silica content will therefore exhibit greater degrees of polymerization, and have higher viscosities, than those with low-silica contents.
The amount of dissolved gases in the magma can also affect it’s viscosity, but in a more ambiguous way than temperature and silica content. When gases begin to escape (exsolve) from the magma, the effect of gas bubbles on the bulk viscosity is variable. Although the growing gas bubbles will exhibit low viscosity, the viscosity of the residual liquid will increase as gas escapes. The overall bulk viscosity of the bubble-liquid mixture depends on both the size and distribution of the bubbles. Although gas bubbles do have an effect on the viscosity, the more important role of these exsolving volatiles is that they provide the driving force for the eruption. This is discussed in more detail below.

As dissolved gases are released from the magma, bubbles will begin to form. Bubbles frozen in a porous or frothy volcanic rock are called vesicles, and the process of bubble formation is called vesiculation or gas exsolution. The dissolved gases can escape only when the vapor pressure of the magma is greater than the confining pressure of the surrounding rocks. The vapor pressure is largely dependent on the amount of dissolved gases and the temperature of the magma.
Explosive eruptions are initiated by vesiculation, which in turn, can be promoted in two ways: (1) by decompression, which lowers the confining pressure, and (2) by crystallization, which increases the vapor pressure. In the first case, magma rise can lead to decompression and the formation of bubbles, much like the decompression of soda and the formation of CO2 bubbles when the cap is removed. This is sometimes referred to as the first boiling. Alternatively, as magma cools and anhydrous minerals begin to crystallize out of the magma, the residual liquid will become increasingly enriched in gas. In this case, the increased vapor pressure in the residual liquid can also lead to gas exsolution. This is sometimes referred to as second (or retrograde) boiling. Both mechanisms can trigger an explosive volcanic eruption.

The amount of dissolved gas in the magma provides the driving force for explosive eruptions. The viscosity of the magma, however, is also an important factor in determining whether an eruption will be explosive or nonexplosive. A low-viscosity magma, like basalt, will allow the escaping gases to migrate rapidly through the magma and escape to the surface. However, if the magma is viscous, like rhyolite, its high polymerization will impede the upward mobility of the gas bubbles. As gas continues to exsolve from the viscous melt, the bubbles will be prevented from rapid escape, thus increasing the overall pressure on the magma column until the gas ejects explosively from the volcano. As a general rule, therefore, nonexplosive eruptions are typical of basaltic-to-andesitic magmas which have low viscosities and low gas contents, whereas explosive eruptions are typical of andesitic-to-rhyolitic magmas which have high viscosities and high gas contents.
(centigrade) VISCOSITY GAS
~50% mafic ~1100 low low nonexplosive
~60% intermediate ~1000 intermediate intermediate intermediate
~70% felsic ~800 high high explosive
There are, however, two exceptions to this general rule. Andesitic-to-rhyolitic lavas that have been degassed often erupt at the surface nonexplosively as viscous lava domes or obsidian flows. Similarly, many of the so-called hydrovolcanic eruptions involve basaltic-to-andesitic magmas that erupt explosively in the presence of groundwater or surface water..

There are about 550 volanoes on earth that have erupted in historic times. Such volcanoes are considered to be geologically active (see: What is an “active” volcano?). In addition, there are an equivalent number of dormant volcanoes that have not erupted in historic time, but have erupted in the past 10,000 years. Both dormat and “active” volcanoes have the potential to erupt again. On any given day, there are about ten volcanoes that are actively erupting.
Eruptions are highly variable in size and explosiveness. As the table below demonstrates, small eruptions are more frequent than larger eruptions. It takes a greater amount of time to build up the necessary gas pressures needed for the larger eruptions. The global frequency of small eruptions, producing 0.001 to .01 cubic kilometers of volcanic material, is once every few months, whereas the frequency of very large eruptions, producing thousands of cubic kilometers of ash, is about once every 100,000 years. The last truly giant eruption on earth occurred at the Toba volcano on Sumatra 74,000 years ago. It produced ~2,800 cubic kilometers of ash, more than 2000 times the amount generated by the 1980 eruption of Mt. St. Helens!
SIZE (cubic km)
.001-.01 Kilauea, Unzen
several months
.01-.1 Etna
5 years
.1-1.0 St. Helens (1980)
10 years
1-10 Pinatubo, Katmai
100 years
10-100 Krakatau (1883)
1000 years
100-1000 Tambora (1815)
10,000 years
>1000 Yellowstone, Toba
100,000 years

MEASURING EXPLOSIVENESS: The Volcano Explosivity Index (VEI)
It is extremely difficult to assess the magnitude of an eruption in a truly quantitative way. However, an approximation of the explosiveness of an eruption can be obtained from the degree that the airborne volcanic products (tephra) have been fragmented. The greater the explosivity, the greater the fragmentaion of the tephra deposits. Since most eruptions on earth have not been observed by humankind, the degree of tephra fragmentation is the only criterion we can use to determine the explosivity of ancient (non-observed) eruptions.
However, for historic eruptions, that have been observed, we can use additional criteria. Chris Newhall of the U.S. Geological Survey and Steve Self at the Open University(UK) have developed a simple, semiquantitative scheme for estimating the magnitude of historic eruptions, called the Volcanic Explosivity Index (VEI). Historical eruptions can be assigned a VEI number on a scale of 0 to 8, using one or more of the following criteria:
Volume of ejecta
Height of the eruptive column
Qualitative descriptions (“gentle”, “effusive”, “explosive”, “cataclysmic”, etc.)
Style of past activity
Height of spreading of the eruptive plume head (in troposphere or stratosphere)
The volume of ejecta and the plume height of the eruption column are probably the two most reliable criteria to use in giving an eruption a VEI number. The VEI numbers below correspond the the following eruption characteristics:

(See Eruption Types)
0 <100 m 1000s m3 Hawaiian Kilauea

1 100-1000 m 10,000s m3 Hawaiian/Strombolian Stromboli

2 1-5 km 1,000,000s m3 Strombolian/Vulcanian Galeras (1992)
3 3-15 km 10,000,000 m3 Vulcanian Ruiz (1985)

4 10-25 km 100,000,000s m3 Vulcanian/Plinian Galunggung (1982)
5 >25 km 1 km3 Plinian St. Helens (1980)

6 >25 km 10s km3 Plinian/Ultra-Plinian Krakatau (1883)

7 >25 km 100s km3 Ultra-Plinian Tambora (1815)

8 >25 km 1000s km3 Ultra-Plinian Toba (74 ka)

The VEI is similar to the Richter scale for measuring magnitude, in that each interval on the VEI represents an increse in magnitude of about 10 (i.e., it is logarithmic). The VEI has been used by workers of the Smithsonian Institution to assign magnitude to Holocene (<10,000 years before present) volcanoes in the Catalog of Active Volcanoes of the World. No Holocene volcano has been assigned a VEI of 8, although four eruptions, including Tambora (1815) have been assigned a VEI of 7.

The model below depicts a cross-sectional view an explosively erupting volcano, typical of a so-called Plinian eruption. A variety of pressure surfaces exist within the magma column beneath the erupting volcano, and within the eruption column above the volcano. The section below describes the nature of these pressure surfaces and explains the dynamic processes associated with the eruption model. To view a QuickTime movie of this eruption, click movie. To view a 3D flyby beneath the eruption plume of the volcano, click flyby.

THE MAGMA COLUMN — Eruptions are fed from a magma column that exists below the point source of the eruption. The magma column contains two critical pressure surfaces:
(1) The exsolution surface occurs in the magma reservoir beneath the volcano. It separates a zone of magma containing dissolved volatiles from an overlying zone of magma containing exsolved gas bubbles
(2) The fragmentation surface occurs at the top of the magma column. It separates the zone of magma containing exsolved gas from the overlying eruption column. Fragmentation of the magma is generated by rapid gas expansion and bubble explosion.
The exsolution of gas from magma (or “boiling”) is called vesiculation. These gas bubbles (vesicles) begin to form at the exsolution surface. Vesiculation is promoted by decompression of magma as it rises upward (where the confining pressure is less than the dissolved gas pressure). Fragmentation of the bubble walls then begins at the fragmentation surface; here, the gas bubbles grow during ascent until they become unstable and explode. This occurs when the volume of bubbles is about 75% of the total volume of the magma column. Gas release is confined to the diameter of the magma column, and the eruption velocity is controlled mainly by the gas content. The low strength of surface rocks and the high initial exit pressure commonly results in vent erosion, so that a flared vent shape develops which enhances velocity. This marks an upward transition from subsonic to supersonic flow.
THE ERUPTION COLUMN –The fragmentation surface is the point source of the eruption. The region of hot gas and broken pyroclastic particles above the fragmentation surface is called the eruption column. It transports pyroclastic materials from the ground into the atmosphere. Common observed heights for plinian eruptive columns are between 2 and 45 km. The eruption column for the 1994 eruption of the Klyuchevskoi volcano in Kamchatka is shown here.
The physical and mechanical behavior of the erupting column changes as it rises from the the fragmentation surface into the stratosphere. These progressive behavioral changes allow the eruption column to be subdivided into three regions:
(1) the gas thrust region in the lower column, driven by gas expansion,
(2) the convective thrust region in the upper column, driven by the constant release of thermal energy from internal ash, and
(3) the umbrella region at the top of the eruption column. The umbrella region is also known as the “downwind plume”.

The gas thrust region is a region of vertical movement of gas and plastic magma particles that is internally powered by gas expansion (decompression) at the base of the eruptive column. Its overall density depends on the particle-to-gas ratio. Initially, the gas thrust is denser than air because of its incorporation of pyroclastic particles. However, as these particles fall out (as ballistics), the density is reduced so that the hot gases in the column are now less than that of the atmosphere. At this point, the gas thrust gives way to convective uprise and development of the convective thrust region, which comprises about 90% of the eruptive column. The convective plume is driven by the constant release of thermal energy from internal ash. At some point, the bulk density of the column will become equal to that of the atmosphere, so that its upward mobility is only controlled by its momentum. The column thus spreads out in the umbrella region. The bottom of umbrella region is where densities of the plume and the surrounding air are equal. Continued upward mobility towards the top of the umbrella region is controlled by momentum. The umbrella region is often asymmetric due to the effect of high atmospheric winds in the stratosphere.

Dr. Ken Wohletz at the Los Almos National Laboratory has created a magnificent computer program which allows the user to construct a diverse group of volcanic landforms through the simulation of variable types of eruptions. The parameters of each eruption type can be set independently. The eruption software can be downloaded free by clicking Erupt 2.5. A commercial version (Erupt 3.0) will soon be available from RockWare.


Volcanic landscapes contain diverse landforms. The most recognizable of these include volcano edifices, calderas, and lava domes. Each of these landforms can vary markedly in size, shape, composition, and eruptive history.

A volcanic vent is an opening exposed on the earth’s surface where volcanic material is emitted. All volcanoes contain a central vent underlying the summit crater of the volcano. The volcano’s cone-shaped structure, or edifice, is built by the more-or-less symmetrical accumulation of lava and/or pyroclastic material around this central vent system. The central vent is connected at depth to a magma chamber, which is the main storage area for the eruptive material. Because volcano flanks are inherently unstable, they often contain fractures that descend downward toward the central vent, or toward a shallow-level magma chamber. Such fractures may occasionally tap the magma source and act as conduits for flank eruptions along the sides of the volcanic edifice. These eruptions can generate cone-shaped accumulations of volcanic material, called parasitic cones. Fractures can also act as conduits for escaping volcanic gases, which are released at the surface through vent openings called fumaroles.

Although every volcano has a unique eruptive history, most can be grouped into three main types based largely on their eruptive patterns and their general forms. The form and composition of the three main volcano types are summarized.

Straight sides with steep slopes; large summit crater Basalt tephra; occasionally andesitic Strombolian


Very gentle slopes; convex upward Basalt lava flows Hawaiian


Gentle lower slopes, but steep upper slopes; concave upward; small summit crater Highly variable; alternating basaltic to rhyolitic lavas and tephra with an overall andesite composition Plinian

SUBORDINATE VOLCANO TYPES — Lava and tephra can erupt from vents other than these three main volcano types. A fissure eruption, for example, can generate huge volumes of basalt lava; however, this type of eruption is not associated with the construction of a volcanic edifice around a single central vent system. Although point-source eruptions can generate such features as spatter cones and hornitos, these volcanic edifices are typically small, localized, and/or associated with rootless eruptions (i.e., eruptions above the surface of an active lavaflow, unconnected to an overlying magma chamber) . Vent types related to hydrovolcanic processes generate unique volcanic structures, discussed separately under hydrovolcanic eruptions.


Scoria cones, also known as cinder cones, are the most common type of volcano. They are also the smallest type, with heights generally less than 300 meters. They can occur as discrete volcanoes on basaltic lava fields, or as parasitic cones generated by flank eruptions on shield volcanoes and stratovolcanoes. Scoria cones are composed almost wholly of ejected basaltic tephra. The tephra is most commonly of lapilli size, although bomb-size fragments and lava spatter may also be present. The tephra fragments typically contain abundant gas bubbles (vesicles), giving the lapilli and bombs a cindery (or scoriaceous) appearance. The tephra accumulates as scoria-fall deposits which build up around the vent to form the volcanic edifice. The edifice has very steep slopes, up to 35 degrees, although older eroded scoria cones typically have gentler slopes, from 15 to 20 degrees. Unlike the other two main volcano types, scoria cones have straight sides and very large summit craters, with respect to their relatively small edifices. They are often symmetric, although many are asymmetric due to (1) the build up of tephra on the downwind flank of the edifice, (2) elongation of the volcano above an eruptive fissure, or (3) partial rafting of an outer wall of the volcano due to basalt lava oozing outward from beneath the volcano edifice. Where scoria cones have been breached, they typically reveal red-oxidized interiors.
Scoria cones are generated by Strombolian eruptions, which produce eruptive columns of basalt tephra generally only a few hundred meters high. Many scoria cones are monogenetic in that they only erupt once, in contrast to shield volcanoes and stratovolcanoes. An exception is the Cerro Negro volcano in Nicaragua, which is the Earth’s most historically active scoria cone. It is one of several parasitic cones on the northwest flank of Las Pilas volcano. Cerro Negro has erupted more than twenty times since it was born in 1850. Its most recent eruptions were in 1992 and 1995.

Shield volcanoes are broad, low-profile features with basal diameters that vary from a few kilometers to over 100 kilometers (e.g., the Mauna Loa volcano, Hawaii). Their heights are typically about 1/20th of their widths. The lower slopes are often gentle (2-3 degrees), but the middle slopes become steeper (~10 degrees) and then flatten at the summit. This gives shield volcanoes a flank morphology that is convex in an upward direction. Their overall broad shapes result from the extrusion of very fluid (low viscosity) basalt lava that spreads outward from the summit area, in contrast to the vertical accumulation of airfall tephra around scoria-cone vents, and the build-up of viscous lava and tephra around stratovolcanoes. Cross-sections through shield volcanoes reveal numerous thin flow units of pahoehoe basalt, typically < 1 m thick. Pyroclastic deposits are minor (< 1%) and of limited dispersal, generally from flank eruptions associated with parasitic scoria cones, or from rare, localized hydrovolcanic eruptions.
Shield volcanoes are generated by Hawaiian eruptions. However, there is some variability in their eruptive style, which translates into variations in shield morphology and size. The almost perfect symmetry and small volume (~15 km3) of Icelandic shields, for example, stands in marked contrast to the elongation and huge volume (thousands of km3) of Hawaiian shields. These variations are largely attributed to the monogenetic, small-volume, centralized summit eruptions, typical of icelandic shields, and the polygenetic, large-volume, linear fissure eruptions, typical of most hawaiian shields. Still different are the symmetrical Galapagos shields, shown below, which have steep middle slopes (>10 degrees) and flat tops occupied by large and very deep calderas. These shield types appear to be generated by ring-fracture eruptions, which delineate the sides of the caldera and mark the site of caldera collapse.


Stratovolcanoes, also known as composite cones, are the most picturesque and the most deadly of the volcano types. Their lower slopes are gentle, but they rise steeply near the summit to produce an overall morphology that is concave in an upward direction. The summit area typically contains a surprisingly small summit crater. This classic stratovolcano shape is exemplified by many well-known stratovolcanoes, such as Mt. Fuji in Japan, Mt. Mayon in the Philippines, and Mt. Agua in Guatemala.
In detail, however, stratovolcano shapes are more variable than these classic examples, primarily because of wide variations in eruptive style and composition. Some may contain several eruptive centers, a caldera, or perhaps an amphitheater as the result of a lateral blast (e.g., Mt. St. Helens).
Typically, as shown in the image to the left, stratovolcanoes have a layered or stratified appearance with alternating lava flows, airfall tephra, pyroclastic flows, volcanic mudflows (lahars), and/or debris flows. The compositional spectrum of these rock types may vary from basalt to rhyolite in a single volcano; however, the overall average composition of stratovolcanoes is andesitic. Many oceanic stratovolcanoes tend to be more mafic than their continental counterparts. The variability of stratovolcanoes is evident when examining the eruptive history of individual volcanoes. Mt. Fuji and Mt. Etna, for example, are dominanted by basaltic lava flows, whereas Mt. Rainier is dominated by andesitic lava, Mt. St. Helens by andesitic-to-dacitic pyroclastic material, and Mt. Lassen by dacitic lava domes.
Stratovolcanoes typically form at convergent plate margins, where one plate descends beneath an adjacent plate at the site of a subduction zone. Examples of subduction-related stratovolcanoes can be found in many places in the world, but they are particularly abundant along the rim of the Pacific Ocean, a region known as Ring of Fire. In the Americas, the Ring of Fire includes stratovolcanoes forming the Aleutian islands in Alaska, the crest of the Cascade Mountains in the Pacific Northwest, and the high peaks of the Andes Mounains in South America. A satellite view of three stratovolcanoes from the Andes is shown here:
The eruptive history of most stratovolcanoes is delineated by highly explosive Plinian eruptions. These dangerous eruptions are often associated with deadly pyroclastic flows composed of hot volcanic fragments and toxic gases that advance down slopes at hurricane-force speeds. Like shield volcanoes, stratovolcanoes are polygenetic; however, they differ from shield volcanoes in that they erupt infrequently, with typical repose intervals of hundreds of years between eruptions. Most active stratovolcanoes worldwide appear to be < 100,000 years old, although some, like Mt. Rainier, may be more than 1 million years old.

Classifying a volcano as active, dormant, or extinct is a subjective and inexact exercise. A volcano is generally considered active if it has erupted in historic time. This definition, however, is rather ambiguous, because recorded history varies from thousands of years in Europe and the Middle East, to only a few hundred years in other regions of the world, like the Pacific Northwest of the United States. Scientists generally consider a volcano active if it is currently erupting, or exhibiting unrest through earthquakes, uplift, and/or new gas emissions. The Smithsonian Institution’s catalog of active volcanoes, recognizes 539 volcanoes with historic eruptions. In addition, there are 529 volcanoes that have not erupted in historic times, but which exhibit clear evidence of eruption in the past 10,000 years. These latter volcanoes are probably best considered “dormant,” since they have the potential to erupt again.
Whether or not inactive volcanoes are considered truly extinct, or just dormant, depends partly on the average repose interval between eruptions. As noted in eruptive variability, explosive eruptions like those at Toba and Yellowstone have repose intervals of hundreds of thousands of years, whereas non-explosive eruptions have very short repose intervals. Thus, the Yellowstone region, which has not experienced an eruption for 70,000 years, can not be considered extinct. In fact, many scientists consider Yellowstone to be active because of high uplift rates, frequent earthquakes, and a very active geothermal system. Many inactive scoria cones, on the other hand, may be viewed as extinct shortly after they erupt, because such volcanoes are typically monogenetic and only erupt once.

When an erupting volcano empties a shallow-level magma chamber, the edifice of the volcano may collapse into the voided reservoir, thus forming a steep, bowl-shaped depression called a caldera (Spanish for kettle or cauldron). These features are highly variable in size, ranging from 1-100 km in diameter. In most cases, they can be readily differentiated from summit craters, which are generally much smaller and form by explosive erosion of the central vent. Felsic calderas are surrounded by thick blankets of pumice derived from the eruption of voluminous pyroclastic sheet flows.
Variations in form and genesis allow calderas to be subdivided into three types:
• Crater-Lake type calderas associated with the collapse of stratovolcanoes
• Basaltic calderas associated with the summit collapse of shield volcanoes
• Resurgent calderas which lack an association with a single centralized vent

This caldera type is generated after the main phase of a Plinian eruption, during collapse of a stratovolcano into the void of the underlying, depleted magma chamber. Although the waning phase of a Plinian eruption is often associated with the generation of pyroclastic flows, piston-like collapse of the volcanic edifice can generate the additional eruption of voluminous, pumice-dominated sheet flows along ring fractures surrounding the collapsing mass. These sheetflows form thick deposits of ignimbrite, the hallmark of both Crater-Lake type and resurgent calderas.
Crater Lake, Oregon, is the namesake of this caldera type. With a water depth of 600 m, Crater Lake is the deepest fresh-water lake in North America. The caldera walls rise above the lake level an additional 600 m. This large depression formed from the violent eruption and collapse of the ancestral stratovolcano Mt. Mazama about 6850 years ago. The massive eruption generated ~50 times more tephra than the Mt. St. Helens eruption in 1980. About 30 km of pyroclastic material erupted during the main plinian phase, thus depleting the magma chamber and leaving its roof unsupported. As ignimbrites erupted toward the end of the plinian phase, the volcano edifice began to collapse along ring fractures. The collapse generating additional pyroclastic sheet flows and a 10-km-wide caldera (see Mazama ignimbrite). The caldera has since been the site of several small eruptions which have covered parts of the caldera floor with andesitic to rhyolitic lava. The most voluminous of these post-caldera eruptions have built the volcanic cone of Wizard Island on the western side of the lake. These eruptions ceased about 2000 years ago.
The two most recent eruptions associated with the generation of Crater-Lake type calderas are both from Indonesia — the colossal eruption of Tambora in 1815, and the equally spectacular eruption of Krakatau in 1883.

The summit regions of many active shield volcanoes are marked by calderas. Hawaiian examples include the Mokuaweoweo caldera on Mauna Loa and the Kilauea caldera on Kilauea. Others include the Erta Al caldera in Ethiopia, the summit caldera of Piton del la Fournaise on Reunion Island, and the spectacular basaltic calderas on the shield volcanoes of the Galapagos Islands. Most basaltic shield volcano calderas on earth are 1-5 km in diameter. Those observed on Mars, however, are extraordinarily large, the largest being the Olympus Mons caldera, with a diameter of more than 60 km!
Whereas Crater-Lake type calderas are associated with the explosive eruption felsic magma, generating voluminous pyroclastic sheet flows, basaltic calderas are not produced by such catastrophic events. Instead, they subside in increments to produce a nested structure of pits and terraces, as shown in the photos above for the Kilauea caldera and the Erta Al caldera. Basaltic calderas like these are gradually enlarged by episodic collapse, due to the extraction of lava from shallow-level magma chambers underlying the summit areas.

In some cases, the extraction of magma may occur through fractures that feed eruptions along the flanks of the shield volcano. Local collapse along such fractures can generate a linear system of pit craters that emanate away from many basaltic calderas and summit craters. Photo by Vic Camp.

Resurgent calderas are the largest volcanic structures on earth. They are associated with massive eruptions of voluminous pyroclastic sheet flows, on a scale not yet observed in historic times. The youngest of these resurgent calderas is the 74,000-year-old Toba Caldera on the Indonesian Island of Sumatra. The Toba eruption generated 2800 times more pyroclastic material than the moderate Plinian eruption of Mt. St. Helens in 1980! There are three resurgent calderas in the United States less than 1.5 million years old — the Valles Caldera in New Mexico, the Long Valley Caldera in California, and the Yellowstone Caldera in Wyoming.
With diameters ranging from 15 to 100 km, resurgent calderas dwarf those of Crater-Lake type. They are similar to Crater-Lake type calderas in that they are also generated by crustal collapse above shallow magma chambers. Resurgent calderas, however, are too large to have been associated with a Crater-Lake type central volcano, like Mt. Mazama. Apart from their large size, the definitive feature of resurgent calderas is a broad topographic depression with a central elevated mass resulting from post-collapse upheaval of the caldera floor. Uplift is commonly more than one kilometer. The caldera floor is typically filled with rhyolitic lavas, obsidian flows, and domes, and the uplifted centers often contain elongate rifts (graben) along their crests.
The evolution of a resurgent caldera is depicted in the illustration shown here (modified from Calvin J. Hamilton). Caldera formation begins with crustal uplift associated with the arrival of a large plume of gas-rich rhyolitic magma (A). Ring fractures propagate outward from the chamber toward the surface and these are used as conduits for escaping magma. Decompression of the magma results in massive vesiculation and the explosive eruption of tephra into the high atmosphere (B). As the eruption wanes, pyroclastic flows begin to erupt from the ring fractures. The magma chamber is depleted and its crustal roof begins to collapse, generating additional pyroclastic flows (C). As the eruption ends, hundreds of meters of pyroclastic flows have now accumulated within the caldera (inflow facies) and beyond the caldera walls (outflow facies) (D). The floor of the caldera may quickly be occupied by a lake (E). Resurgence of the caldera floor takes about 1000 to 100,000 years to accomplish (F). This period of uplift may be controlled by compression of the remaining magma beneath the collapsed roof of the chamber, or by the arrival of new magma into the chamber.

Valles Caldera, New Mexico
The Valles caldera lies within the Jemez volcanic field in New Mexico, at the northern end of the north-south trending Rio Grand rift. It is one of three resurgent calderas in the U.S. that erupted less than 1.5 million years ago; the other two being the Long Valley caldera of California and the Yellowstone caldera. The Valles caldera developed above the site of an earlier eruption associated with the collapse of the Toledo caldera. Both eruptions took place between 1.4 and 1.0 million years ago. Collapse of the Valles caldera produced a circular depression 23 km (14 miles) in diameter. Eruption and collapse of the Toledo caldera generated pyroclastic flows of the Lower Bandelier tuff, which can be seen along the canyon walls west of the caldera. Eruption and collapse of the Valles caldera generated the overlying pyroclastic flows of the Upper Bandelier tuff.
Valles Caldera — The lowlands in this shaded elevation map of the Valles caldera are shown in blue and the highlands in orange. The circular depression and the resurgent dome of the caldera are evident in the center of the image. The prominent radial valleys along the outer rim of the caldera are incised into the ouflow facies of the Bandelier tuff. Copyright Calvin J. Hamilton.
Cerro Galan Caldera, Argentina
The recently discovered Cerro Galan Caldera, an elliptical structure with a diameter of ~35 km, lies along the crest of the Andes Mountain chain in northwestern Argentina. A magnificently exposed related caldera, Cerro Panizos, lies farther north, along the Argentina-Bolivia border. The Cerro Galan caldera formed 2.2 million years ago with a colossal eruption that generated more than 1000 km3 of pumice-bearing pyroclastic flows (ignimbrites). The outflow facies of these flows radiate outward from caldera for distances of over 100 km. Within the caldera, the ignimbrites accumulated to a thickness of over 1.2 km. Resurgence probably began a few thousand years after its birth. At one time, the caldera enclosed a large lake, the remainder of which is present as a salty remnant on the western side of the caldera, called Laguna Diamante. The caldera floor was uplifted asymmetrically along its eastern side, probably as a consequence of a small intrusion of new magma above the one that cause the eruption.
Cerro Galan Caldera — The caldera walls and resurgent floor of this north-south elliptical structure is evident in this false-colored Thematic Mapper Landsat image. Laguna Diamante lies in the southwest corner of the caldera. Courtesy of Peter Francis.
After resurgence, a thick pile of dacitic coulées erupted non-explosively along the northern caldera ring fracture about 2.1 million years ago. These probably represent the last vestiges of the magma that fed the eruption of pyroclastic flows. The non-explosive nature of these dacitic coulées is due to their low volatile content, which was lost during rapid vesiculation (degassing) of the magma by the caldera-forming eruption.

Yellowstone Caldera, Wyoming
Yellowstone National Park sits on a high plateau supported by a hot mantle plume. This so-called hotspot has been the site of three caldera-forming eruptions in the last two million years. The last eruption, ~630,000 years ago, generated ~1000 km3 of pyroclastic sheet flows and airfall tephra which fell as far away as Louisiana. The Yellowstone calderas are part of a southwest-trending, linear system of calderas that becomes progressively older from Yellowstone through the Snake River Plain of Idaho, to the 17-million-year-old McDermitt Caldera along the Oregon-Idaho border. Most of the calderas in Idaho are covered with younger basaltic lavas associated with the Snake River Plain. This appears to be one of the few continental examples of a hotspot track, similar to the oceanic hotspot track of the Hawaiian and Emporer seamount chains.
The caldera-forming rhyolite eruptions along the hotspot track, and the basalt eruptions along the Snake River Plain, clearly are derived from seperate sources. The basaltic melts are partial melts from the mantle and the rhyolitic melts are partial melts of continental crust. The mantle plume that currently exists beneath Yellowstone generated hot basalt magmas at depth, which ascended into the lower crust where they melted continental crust to produce the rhyolitic magmas. These rhyolite melts then ponded beneath the Yellowstone region, in a shallow magma chamber that generating doming and the development of ring fractures on the Yellowstone surface. Explosive volcanism above this elevated and fractured dome began about 630,000 years ago, and collapse of the structure soon followed.
Yellowstone Caldera — The ascent of hotspot-generated basaltic magmas results in melting of the continental crust to produce rhyolitic magmas. These gas-rich felsic magmas pond at shallow levels and erupt along ring fractures. The caldera is formed due to piston-like collapse of the magma chamber roof after the rhyolitic magma chambers are partially depleted. Courtesy of John Wiley and Sons.
Two resurgent domes currently occupy the caldera floor — the Mallard Lake dome to the west and the Sour Creek dome to the east. Resurgence has been associated with the intermittent extrusion of largely degassed rhyolitic lavas, most of which have erupted over the last 150,000 years.
Although felsic lavas are generally associated with explosive Plinian-type eruptions, they sometimes extrude at the surface as relatively cold, viscous lava domes. Vesiculation during eruption will deplete the magma in volatiles, thus depriving the magma of the very gases needed for a sustained eruption. As the eruption wanes, this degassed magma will rise up the central vent where it extrudes at the summit crater as a massive plug of viscous lava. The process is analagous to a slow-moving plug of toothpaste extruding from a toothpaste tube. In rare cases lava domes can erupt independently through fractures unassociated with a central-vent volcano. Several such domes are present in the Mono Lake-Inyo Craters volcanic chain in east-central California; this chain has erupted four lava domes in the past 700 years.
Three examples of lava domes, from Alaska, Japan, and Saudi Arabia are shown below. For a spectacular 3D simulated view of a lava dome, click Mt. St. Helens lava dome.
The morphology of lava domes is variable. They are typically thick, steep extrusions, but their shapes can vary from circular, low-profile domes (tortas) to cylindral spines with thick talus slopes (Peléean domes). Less sluggish types are gradational to lava flows and are sometimes referred to as dome flows or coulées. Lava domes have compositions ranging from basaltic andesite to rhyolite, although most are composed of crystal-rich dacite. Some lava domes erupt obsidian (volcanic glass), which forms when magma is cooled so quickly that individual minerals do not have time to crystallize from the melt.
Slowly rising lava domes may grow for months or for several years in the aftermath of explosive eruptions. The earth’s longest historical dome-building event has generated the Santiaguito dome in Guatemala. This viscous lava dome began to erupt in 1922 and is still growing today! An empirical example of dome growth through time is shown in the illustration below for the Mt. St. Helens lava dome from 1980 to 1983.
The slow growth of mature lava domes is generally nonexplosive. However, younger domes composed of lava that is not entirely degassed, may occassionally explode. The periodic explosion or gravitational collapse of these viscous masses can sometimes generate deadly pyroclastic flows. These avalanches of hot lava fragments and toxic gases have killed thousands of people throughout history. For a more thorough explaination of dome-collapse pyroclastic flows, link to nuée ardentes.

This section describes the volcanic features and phenomena generated by volcanic eruptions. The eruptive products are highly variaible and largely dependent on the composition, viscosity, and gas content of the erupting magma. Lava flows, for example, are more common in relatively non-explosive basaltic eruptions associated with shield volcanoes, scoria cones, and fissures. On the other hand, pyroclastic flows, lahars and voluminous tephra deposits are more common in explosive andesitic-to-rhyolitic eruptions associated with stratovolcanoes. Gaseous emissions are also examined, as are their harmful effects on both local and global scales.

The injection of basaltic lava into surface water can be quiescent or highly explosive. Pahoehoe tends to pour into water in a passive manner, sometimes with little interaction beyond the boiling of water. More explosive interactions are generally associated with a’a flows. Water has a greater opportunity to penetrate into the cracks of a’a, where it expands explosively and fragments the lava into airborne particles. Littoral cones can develop at the water’s edge from the accumulation of these airborne fragments.

When basalt enters water passively, it forms pillow basalt. These are bulbous bodies with quenched, glassy rinds. They are often spherical, with diameters of 30-100 cm, although many are lobate. During formation, the outer surface of each lobe chills instantaneously. However, the constant injection of lava may cause a single lobe to swell, thus cracking the glassy rind, allowing a new lobe bud forth. Individual pillows can break off the budding lobe and cool as a single mass. As the interior of a pillow cools, it forms crude cooling joints that radiate inward, perpendicular to the cooling surface. Each pillow will settle into gaps between the pillows below it. In continental lakes, pillow accumulations can be several tens of meters thick, whereas those that develop on the ocean floors can be hundreds of meters thick.

Quenching and fragmentation of basaltic lavas produces an accumulation of angular, glassy fragments called hyaloclastite. Basalt fragmentation can occur by the explosive eruption process, or by an essentially nonexplosive process associated with the spalling of pillow basalt rinds by thermal shock or chill shattering. The hyaloclastite fragments are equant, angular shards of black glass, generally between .25 and 2.0 cm in diameter . This water-quenched basalt glass is called sideromelane, a pure variety of glass that that is transparent, and lacks the very small iron-oxide crystals found in the more common opaque variety of basalt glass called tachylite. In hyaloclastite, these glassy fragments are typically surrounded by a matrix of yellow-to-brown palagonite, a wax-like substance that forms from the hydration and alteration of the sideromelane. A related hydrovolcanic deposit, called palagonite tuff, contains a mixture of sideromelane shards and coarser-grained fragments of basaltic rock in a palagonite matrix.
Limu o Pele – Recent hydrovolcanic basalt deposits may preserve paper-thin glassy fragments known as Limu o Pele (seawead of Pele). These form from the bursting of large glass bubbles generated by the rapid expansion of steam as lava interacts with seawater. These are not well-preserved in most ancient deposits.
BASALT INTERACTION WITH WET SEDIMENT: spiracles, clastic dikes, peperites, and invasive lava flows
A variety of features can be generated by basaltic lava flowing over wet sediment. As lava advances over a wet spot, like a puddle, the resulting steam explosion can inject sediment into the interior of the flow. The sediment protrusion produced by this event is called a spiracle. If the steam explosion is large enough, it might generate an explosion tube through the entire flow. Similar explosion tubes are created by CO2 explosions above vegetated areas. Steam action can also produce clastic dikes. These features develop in fractures that act as conduits for the transport of sediment-carrying steam into the body of the lava flow.
In some rare cases, the moving lava may mix with the wet sediment to produce peperite, a heterogeneous deposit composed of dark basalt fragments and light-colored sediment. The formation of peperites is not well understood, but they appear to be associated with the invasion of basalt into wet sediment, without significant explosion, but with at least some fragmentation of the basalt.
Many basaltic rocks in the geologic record can be found sandwiched between two sedimentary rocks. If field evidence shows that the basalt cooled against both the sediment below and above, then it is typically classified as a shallow, near-horizontal intrusion called a sill. However, many workers on the Columbia River Flood Basalt Province have found that some of these sill-like bodies are not intrusions at all, but rather invasive lava flows. Although liquid, the basaltic lava has a higher density than the wet sediment. Thus,a basaltic flow on the surface of the earth will tend to burrow into an accumulation of wet sedimtent rather than flow above it.
As the head of the flow bulldozes into the deposit, the invasion is passive, with little disruption of the sediment. The sediment-lava contact is typically horizontal. Pillow-like structures are scarce, as are appendages of lava that protrude vertically into the overlying sediment. Unlike surface lavas, the upper surface of invasive flows is rarely vesicular, and is typically marked by a glassy selvage due to rapid cooling against the overlying sediment. It may be that the geologic record contains many invasive lavas that have previously interpreted as basaltic sills.

The rapid eruption of expanding gases results in the obliteration and fragmentation of magma and rock. The greater the explosivity, the greater the amount of fragmentation. Individual eruptive fragments are called pyroclasts (“fire fragments”). Tephra (Greek, for ash) is a generic term for any airborne pyroclastic accumulation. Whereas tephra is unconsolidated, a pyroclastic rock is produced from the consolidation of pyroclastic accumulations into a coherent rock type.

CLASSIFICATION OF PYROCLASTS – Tephra is classified on the basis of pyroclast size:
ASH — Very fine-grained fragments (< 2 mm), generally dominated by broken glass shards, but with variable amounts of broken crystal and lithic (rock) fragments. Courtesy of USGS.
LAPILLI — Pea- to walnut-size pyroclasts (2 to 64 mm). They often look like cinders. In water-rich eruptions, the accretion of wet ash may form rounded spheres known as accretionary lapilli (left). Courtesy of USGS.

BLOCKS AND BOMBS — Fragments >64 mm. Bombs are ejected as incandescent lava fragments which were semi-molten when airborne, thus inheriting streamlined, aerodynamic shapes. Blocks (not shown) are ejected as solid fragments with angular shapes. Courtesy of J.P. Lockwood, USGS.
Within this size classification, specific types of tephra can be further defined by physical attributes. For example, lapilli-size fragments of basaltic lava may cool quickly while airborne, to form glassy teardrop-shaped lapilli called Pele’s tears. During strong winds, these molten fragments can be drawn out into fine filaments called Pele’s hair. Nonexplosive Hawaiian-type eruptions often produce lapilli- to bomb-size fragments, called spatter which remain airborne for only a short amount of time so that are still liquid when they hit the ground surface. Some lapilli- to bomb-size pyroclasts maintain a frothy appearance due to escaping volcanic gases and the rapid accumulation of bubbles (vesicles). These highly vesicular rock types include mafic and felsic varieties. Vigorous gas escape in felsic lavas produces pumice, whereas similar gas escape in mafic lavas produces reticulite. The high vesicularity of pumice lowers the density of this rock type so that it can literally float on water. Reticulite has a still lower density, with vesicles occupying up to 98% of the total volume. Unlike pumice, however, most of the bubble walls in reticulite are broken so that it sinks in the presence of water. Reticulite, however, is not as common as scoria, a denser mafic rock containing a smaller abundance of relatively large vesicles. Click the links below to view images of these pyroclast types.

In some cases, unconsolidated pyroclasts can be welded, compacted, or cemented into a coherent pyroclastic rock. The nongenetic classification of pyroclastic rocks is partly based on the relative abundance of the incorporated pyroclast types discussed above. For example:
• Ash tuff – rock dominated by ash; sometimes simply referred to as tuff.
• Lapilli tuff – rock dominated by lapilli.
• Tuff breccia – rock containing 25% to 75% blocks and/or bombs.
• Pyroclastic breccia – rock containing at least 75% blocks and bombs.
• Agglomerate – rock containing at least 75% bombs.
• Agglutinate – rock composed of fused, largely unrecognizable, basalt spatter fragments.

Pyroclastic airfall deposits (tephra) contain pyroclasts that are coarser near the vent (bomb and lapilli size) and become increasingly finer grained away from the vent (lapilli to ash size). The mafic to felsic composition of these deposits will vary with eruption type. Two end-member genetic types of airfall deposits are recognized:
• Scoria-fall deposits — These are derived from Strombolian eruptions of scoria cones. The deposit is composed of basaltic to andesitic vesiculated pyroclasts (scoria) lying near the eruptive vent.
• Pumice-fall deposits — These are derived from Plinian eruptions of stratovolcanoes. The deposit is composed of highly vesiculated dacitic to rhyolitic pyroclasts (pumice) which can be destributed for hundreds of square kilometers away from the vent.
In addition to these two end-member types, some airfall deposits can be very highly fragmented to generate ash-fall deposits. Such fine-grained deposits are common in highly explosive Vulcanian eruptions and hydrovolcanic eruptions.
A pyroclastic flow is a fluidized mixture of solid to semi-solid fragments and hot, expanding gases that flows down the flank of a volcanic edifice. These awesome features are heavier-than-air emulsions that move much like a snow avalanche, except that they are fiercely hot, contain toxic gases, and move at phenomenal, hurricane-force speeds, often over 100 km/hour. They are the most deadly of all volcanic phenomena.

The extraordinary velocity of a pyroclastic flow is partly attributed to its fluidization. A moving pyroclastic flow has properties more like those of a liquid than a mass of solid fragments. Its fluid behavior can only be described as spectacular, as evidenced by the 6000-year-old Koya flow in southern Japan, which traveled more than 60 km from its source, ten of which were over open water! The Koya flow left a deposit that was only two meters thick over its 60 km extent. Such mobility comes from the disappearance of inter-particle friction. A fluidized flow is best described as a dispersion of large fragments in a medium of fluidized fine fragments. A constant stream of hot, expanding gases keeps the smallest of the fragments (ash and lapilli size particles) in constant suspension. This solid-gas mixture can then support larger fragments that float in the matrix. The expanding gas component is derived from a combination of (1) the constant exsolution of volcanic gas emitted by the hot pyroclasts, and (2) from the ingestion, heating, and rapid expansion air during movement of the flow.

The terminology of pyroclastic flows and pyroclastic flow deposits can be complex and confusing. In general, there are two end-member types of flows:
• NUÉE ARDENTES — these contain dense lava fragments derived from the collapse of a growing lava dome or dome flow, and
• PUMICE FLOWS — these contain vesiculated, low-density pumice derived from the collapse of an eruption column.

A nuée ardente deposit is called a block-and-ash deposit (see below), whereas a pumice flow deposit is called an ignimbrite. Each end-member type of pyroclastic flow is discussed below, followed by a description of ignimbrites.

The French geologist Alfred Lacroix attached the name nuée ardente (glowing cloud) to the pyroclastic flow from Mt. Pelée that destroyed the city of St. Pierre in 1902. The flow was generated from the explosive collapse of a growing lava dome at the summit of the volcano, which then swept down on the city. Thus, nuée ardente eruptions are often called Peléen eruptions. However, this term cannot be as narrowly defined as the other eruption types, because nuée ardentes are often linked with both Plinian and Vulcanian activity.
Although the term nuée ardente is now applied to all pyroclastic flows generated by dome collapse, it is somewhat of a misnomer to describe these features as a “glowing cloud.” A more precise term would be glowing avalanche. The bulk of these hot block-and-ash flows hug the ground surface, but are disguised by an overlying cloud of fine ash particles that are winnowed out of the flow by a processes called elutriation. Nuée ardentes, therefore, are composed of two related parts: a pyroclastic avalanche largely hidden from view by an overlying ash cloud, sometimes called a co-ignimbrite ash.
Nuée ardente deposits are composed of dense, non-vesiculated, blocky fragments derived from the collapsed lava dome. They therefore differ significantly from the highly vesiculated ignimbrites which are derived from eruption column collapse. Nuée ardente deposits contain blocks in a fine-grained matrix of ash. The deposits, therefore, are called block-and-ash deposits. They are denser than ignimbrites, and typically are less extensive.

Pumice flows are pumice-rich pyroclastic flows derived from the collapse of an eruption column. The lowermost part of the eruption column is called the gas thrust region. Here, the density of the eruption column is greater than the density of the surrounding air. The column continues to rise, however, because of the thrust provided by the release and rapid expansion of volcanic gas. Occasionally, the gas thrust region may become so chock-full of debris that its high density cannot be supported by the thrust of expanding gases. The column thus collapses downward under gravity as a mass of vesiculating pumice that advances rapidly down the flanks of the volcano.
Although both nuée ardentes and pumice flows are fluidized, pumice flows are more energetic and mobile. This is partly attributed to their lower densities, but also to their greater store of kinetic energy generated by vertical drops up to several kilometers above the volcano’s summit. The further it falls, the greater its kinetic energy, and the further and faster it will travel horizontally.
Pumice flows have a tripartite division. The main body hugs the ground surface and is dominated by pumice fragments in an ash matrix. Like nuée ardentes, this pyroclastic avalanche is overlain by an ash cloud of elutriated fine ash. An additional component of a pumice flow is the ground surge. These are forward-springing jets of incandescent ash that occur in the advancing head of the flow. They advance with a rolling and rapidly puffing movement which is thought to be caused by the ingestion of air in the front of the flow. Air ingestion produces strong fluidation in flow front, and explosive heating of the air causes some of the material to be hurled forward as a low-density, turbulent surge.
A single Plinian-type eruption may generate hundreds of pumice flows which typically flow down valleys radiating outward from the summit of the volcano. Individual flows may vary in length from a few kilometers to tens of kilometers. These are miniscule, however, in comparison to the massive pumice flows generated by caldera collapse. Caldera-generated flows are not restricted to valleys, but rather fill in valleys and adjacent low ridges to produce pumice-dominated pyroclastic sheet flows that can obliterate an area the size of Ohio in a few minutes. These huge eruptions can eject a thousand cubic kilometers of material from ring fractures in just a few hours. The last such eruption on earth took place at Toba, Indonesia, about 74,000 years ago to deposit an ignimbrite with a volume of over 2000 cubic kilometers. Similar eruptions in the United States occurred less than two million years ago at the Long Valley, Valles and Yellowstone calderas.

Lahar is an Indonesian term for a volcanic mudflow. These lethal mixtures of water and tephra have the consistency of wet concrete, yet they can flow down the slopes of volcanoes or down river valleys at rapid speeds, similar to fast-moving streams of water. These mud slurries carry debris ranging in size from ash to lapilli, to boulders more than 10 meters in diameter. Lahars can vary from hot to cold, depending on their mode of genesis. The maximum temperature of a lahar is 100 degrees Centigrade, the boiling temperature of water.

Lahars are generated by a variety of mechanisms. The majority are produced by intense rainfall during or after an eruption. A tragic example of such an event was the 1991 eruption of Mt. Pinatubo in the Philippines, which was contemporaneous with the arrival of a major hurricane. An estimated 700 people died from buiral by the ensuing lahars, together with the collapse of structures benearth the wet ash. As demonstrated by the graph below, lahars can also be generated directly from a volcanic eruption as massive amounts of water are generated either by the rapid melting of ice and snow, or by the disruption of crater lakes.
Pyroclastic flows are particularly efficient at generating lahars because they have the capability to melt large quantities of snow and ice in a just few hours. A tragic example of this mechanism occurred in 1985 when pyroclastic flows erupted at the snow-covered summit of the Nevado del Ruiz volcano in Columbia. In only a few hours, the eruption generated a lahar that killed ~23,000 people in the village of Armero and adjacent towns located several tens of kilometers downslope from the summit. Lahars can also be generated by the basal melting of glaciers by lava flows. Basal melting of glacial ice in Iceland has produced largest historic lahars, in terms of discharge. These water-rich, glacial outburst floods are called jokulhlaups.

MT. RAINIER: The Future Site of a Catastrophic Lahar
The snow-covered peaks of the Cascade volcanoes in Washington, Oregon, and northern California pose a clear threat to surrounding towns and villages. Past events suggest that a catastrophic lahar could lie in the future of Mt. Rainier, the largest of the Cascade volcanoes.
The 4000 m high summit of Mt. Rainier contains the largest system of alpine glaciers in the Cascade Range. The periodic melting of glacier ice from Mt. Rainier has generated at least 50 major lahars over the past 10,000 years. The largest of these mudflow deposits, one of the world’s largest, is the ~5700-year-old Osceola lahar, shown in the adjacent map (courtesy of USGS). The Osceola lahar travelled down the White River, over 112 km from its source. It then spread out at its mouth to cover an area of over 300 square kilometers along the shoreline of Puget Sound. The recent geologic history of Mt. Rainier demonstrates that a major mudflow descends down the White River once every 600 years. The younger 500-year-old Electron lahar (see map) was also generated from Mt. Rainier. It flowed 56 kilometers down the Puyallup River to within 15 kilometers of Tacoma, Washington. More than 300,000 people now live in the area covered by these extensive lahars! Unlike floods, such catastropic mudflows can occur with little or no warning. Some volcanologists have predicted that Mt. Rainier will be the site of the next Cascade eruption. Therefore, the volcano is monitored closely, with the hope that we can warn the local population before the next lahar strikes.

Other than free oxygen, generated by photosynthesis, all atmospheric gases were derived from inside the earth and released by volcanic eruptions. The gaseous portion of magma varies from ~1 to 5% of the total weight. Water vapor constitutes 70-90%. The remaining gases include CO2, SO2, and trace amounts of of N, H, CO, S, Ar, Cl, and F. These subordinate gases can combine with hydrogen and water to produce numerous toxic compounds, such as HCl, HF, H2SO4, H2S, which are typical products of fumarolic activity.
H2O (70-90%)
SO2 N, H, S,
F, Ar,
CO, Cl HCl, HF
A variety of sulfur aerosols may be present and sulfur itself may condenses around the fumarole into a crystalline accumulation called sulfaterra (yellow ground). On some volcanoes, enough sulfur is present to be mined as an economic resource. H2S is sometimes called “sewer gas” because it has a rotten egg odor. It is an insidious poison that irritates the eyes, nose, and throat. SO2 on the other hand, is the biting, chocking gas that you smell right after you’ve lit a kitchen match. When these two gases occur together, they react quickly with each other (within minutes) to produce sulfaterra and water vapor.

Monitoring volcanic gases can be helpful in predicting eruptions. For example, an increase in CO2 and SO2 concentrations emitted from fumaroles may indicate increasing magmatic activity beneath the volcano. The composition and relative volumes of these volatiles can be measured in a variety of ways:
• DIRECT MEASUREMENTS — Gases escaping from fumaroles can be collected in evacuated flasks and analyzed in geochemical laboratories. Although much useful data has been assembled in this way, the technique can be dangerous and there is great difficulty in collecting samples that are uncontaminated by atmospheric air.

• COSPEC MEASUREMENTS — Another technique involves the remote sensing of some gases by a correlation spectrometer, a devise originally developed in the 1970s to monitor SO2 and other gases from factory smokestacks. In volcanically active areas, the mounting of COSPECS on helicopters allows SO2 concentrations to be cataloged and monitored on a daily, or weekly, basis. The most recent COSPECs are infrared spectrometers which measure the amount of infrared light absorbed by CO2 and other gases.

• TOMS MEASUREMENTS — On a more regional scale, the distribution and amount of sulfur dioxide released into the stratosphere by volcanic eruptions can be measured by the total ozone mapping spectrometer (TOMS), which is a sensor originally sent aloft in 1979 aboard the Nimbus-7 and Meteor-3 satellites. Improved TOMS instruments sent into lower orbits in 1997 have dramatically improved our detection capabilities.
The destructive effects of volcanism are usually attributed to voluminous outpourings of lava, or to catastrophic pyroclastic eruptions. However, volcanic gases are also dangerous because they are hot and toxic, and their emission in large quantities can easily kill unaware people. For example, lava tubes on active and dormant volcanoes may be dangerous to walk in because they can trap poisonous gases. In rare cases, thousands of people have died in volcanic gas eruptions. To read about a catastrophic example of gas expulsion, see Lake Nyos. Two areas of current volcanic hazard associated with gas emission are described.

One of America’s largest ski resorts is located at Mammoth Mountain, a large complex dome on the western edge of the Long Valley Caldera, California. The Long Valley Caldera marks the site of one of the largest volcanic eruptions in North America over the last 1 m.y. Renewed activity in the form of increased seismic events and fumarole gas emission began in areas around the caldera in the mid-1980s. Shallow-level earthquakes indicated that magma was rising along the SE margin of the caldera to a depth of 3-4 km.
There is clear evidence that CO2 gas is currently ascending through the rocks surrounding Mammoth Mountain. Carbon dioxide has killed trees in at least four areas around the volcano, The largest area of tree destruction covers about 28 acres. The trees are not dying from voluminous surface emissions that could be harmful to humans, but rather from the buildup of CO2 in the underlying soil, which destroys their root systems. Because this gas is denser than air, however, there is a danger that it can collect in small depressions out of the wind, or in low-lying valleys where people could suffocate if camping overnight. An example of this occurred on March 11, 1990, when Fred Richter, a forest ranger at Mammoth got caught in a blizzard. He found refuge in an old cabin that was barely visible through the snow drifts. When he lowered himself in through a hatch in the top, he thought he was safe. However, he quickly began to experience a feeling of suffocation and his pulse began to race uncontrollably. As he began to check for fumes from the cabin’s fireplace, his legs became weak and his body wanted to rest. Instead he struggled back outside where he quickly revived himself in the fresh air. Richter’s experience was brought on by oxygen deprivation in an enclosed space filled with magmatic carbon dioxide derived from ground seepage into the cabin.

The Pu’u O’o eruption on the Kilauea volcano began in 1983. COSPEC measurements before the eruption showed that the volcano was releasing about 150 tons of SO2 per day, nearly all from the summit caldera. However, during the high fire-fountaining eruptions at Pu’u O’o the amount of SO2 released was as high as 30,000 tons per day (these episodes occurred every 3-4 weeks between 1983 and 1986 and lasted 24 hours or less). The current average emission is about 1,800 tons per day.
Chemical reactions are responsible for two types of toxic gas mixtures that are a particular problem for residents and tourists on the the Big Island of Hawaii. These toxic clouds are called VOG and LAZE.
VOG: Sulfur dioxide released at the summit, the east rift, and along skylights in the tube system reacts chemically with oxygen, dust particles, sunlight and water in the air to form a mixture of sulfate aerosols (tiny particles and droplets), sulfuric acid, and other oxidized sulfate species that is known as “vog”. The pollution clouds of vog are normally carried by the northeasterly trade winds to the southwest, where the winds wrap around the island, sending the vog up the western Kona coast, which is the major tourist location on the island. When the trade winds are moderate, the vog stays on the east side of the island. The vog is toxic and hazardous to humans.

LAZE: When lava enters the ocean, another type of chemical reaction takes place between molten lava and sea water. White-plume clouds are generated by intense heat and vigorous chemical reactions. These misty clouds, called “laze” contain a mixture of hydrochloric acid and seawater vapor. Laze clouds can drop rainwater on people along the coast. It is often more acidic than lemon juice, with a pH of 1.5 to 2.5, but it is more corrosive than lemon juice to the skin and clothing. The HCL is toxic and causes irritation to the throat, lungs, eyes, and nose.

There is considerable debate on the role that humans play in changing global climate through both the burning of fossil fuels and the release of chlorofluorocarbon (CFC) gases. Some argue that human interaction poses less of a threat to our atmosphere than do natural processes, like volcanic eruptions. This places a great deal of importance on understanding the role of volcanic eruptions in affecting global climate change. Whatever the source, it is apparent that compositional changes in the earth’s atmosphere generate three principal climatic effects:
Intense sunlight in the stratosphere (above 12 km) produces bluish colored ozone (O3) by naturally breaking down normal oxygen molecules (O2) into two highly reactive oxygen atoms (O). Each oxygen atom then quickly bonds with an oxygen molecule to form ozone. Ozone absorbs UV radiation, and in the process ozone is changed back into an oxygen molecule and an oxygen atom. A balance exists in ozone destruction and production, so that an equilibrium concentration exists in the stratosphere. This equilibrium has probably existed throughout much of geologic time. Recently, however, an ozone hole has been detected in the stratosphere over Antarctica, presumably due to the atmospheric build up of ozone-destroying CFCs by humans. Ozone depletion has resulted in a greater penetration of ultraviolet radiation on the earth’s surface, which is harmful to life on earth because it damages cellular DNA. The ozone effect does not appear to have a direct influence on global temperatures.
Certain gases, called greenhouse gases (primarily carbon dioxide and water vapor; but also methane, N2O, and CFCs), allow short wavelength radiation from the sun (UV and visible light) to penetrate through the lower atmosphere to the earth’s surface. These same gases, however, absorb long wavelength radiation (infrared), which is the energy the earth reradiates back into space. The trapping of this infrared heat energy by these greenhouse gases results in global warming. Global warming has been evident since the beginning of the Industrial Revolution. Most scientists attribute global warming to the release of greenhouse gases through the burning of fossil fuels.
Suspended particles, such as dust and ash, can block out the earth’s sunlight, thus reducing solar radiation and lowering mean global temperatures. The haze effect often generates exceptionally red sunsets due to the scattering of red wavelengths by submicron-size particles in the stratosphere and upper troposphere.

Volcanic eruptions can enhance all three of these climate effects to variable degrees. They contribute to ozone depletion, as well as to both cooling and warming of the earth’s atmosphere. The role of volcanic eruptions on each climate effect is described below.
The halide acid HCl has been shown to be effective in destroying ozone; however, the latest studies show that most volcanic HCl is confined to the troposphere (below the stratosphere), where it is washed out by rain. Thus, it never has the opportunity to react with ozone. On the other hand, satellite data after the 1991 eruptions of Mt.Pinatubo (the Philippines) and Mt. Hudson (Chile) showed a 15-20% ozone loss at high latitudes, and a greater than 50% loss over the Antarctic! Thus, it appears that volcanic eruptions can play a significant role in reducing ozone levels. However, it is an indirect role, which cannot be directly attributed to volcanic HCl. Eruption-generated particles, or aerosols, appear to provide surfaces upon which chemical reactions take place. The particles themselves do not contribute to ozone destruction, but they interact with chlorine- and bromine-bearing compounds from human-made CFCs. Fortunately, volcanic particles will settle out of the stratosphere in two or three years, so that the effects of volcanic eruptions on ozone depletion are short lived. Although volcanic aerosols provide a catalyst for ozone depletion, the real culprits in destroying ozone are human-generated CFCs. Scientists expect the ozone layer to recover due to restrictions on CFCs and other ozone-depleting chemicals by the United Nations Montreal Protocol on Substances that Deplete the Ozone Layer. However, future volcanic eruptions will cause fluctuations in the recovery process.
Volcanic eruptions can enhance global warming by adding CO2 to the atmosphere. However, a far greater amount of CO2 is contributed to the atmosphere by human activities each year than by volcanic eruptions. Volcanoes contribute about 110 million tons/year, whereas other sources contribute about 10 billion tons/year. The small amount of global warming caused by eruption-generated greenhouse gases is offset by the far greater amount of global cooling caused by eruption-generated particles in the stratosphere (the haze effect). Greenhouse warming of the earth has been particularly evident since 1980. Without the cooling influence of such eruptions as El Chichon (1982) and Mt. Pinatubo (1991), described below, greenhouse warming would have been more pronounced.
Volcanic eruptions enhance the haze effect to a greater extent than the greenhouse effect, and thus they can lower mean global temperatures. It was thought for many years that the greatest volcanic contribution of the haze effect was from the suspended ash particles in the upper atmosphere that would block out solar radiation. However, these ideas changed in the 1982 after the eruption of the Mexican volcano, El Chichon. Although the 1980 eruption of Mt. St. Helens lowered global temperatures by 0.1OC, the much smaller eruption of El Chichon lowered global temperatures three to five times as much. Although the Mt. St. Helens blast emitted a greater amount of ash in the stratosphere, the El Chichon eruption emitted a much greater volume of sulfur-rich gases (40x more). It appears that the volume of pyroclastic debris emitted during a blast is not the best criteria to measure its effects on the atmosphere. The amount of sulfur-rich gases appears to be more important. Sulfur combines with water vapor in the stratosphere to form dense clouds of tiny sulfuric acid droplets. These droplets take several years to settle out and they are capable to decreasing the troposphere temperatures because they absorb solar radiation and scatter it back to space.
Observational evidence shows a clear correlation between historic eruptions and subsequent years of cold climate conditions. Four well-known historic examples are described below.
LAKI (1783) — The eastern U.S. recorded the lowest-ever winter average temperature in 1783-84, about 4.8OC below the 225-year average. Europe also experienced an abnormally severe winter. Benjamin Franklin suggested that these cold conditions resulted from the blocking out of sunlight by dust and gases created by the Iceland Laki eruption in 1783. The Laki eruption was the largest outpouring of basalt lava in historic times. Franklin’s hypothesis is consistent with modern scientific theory, which suggests that large volumes of SO2 are the main culprit in haze-effect global cooling.
TAMBORA (1815) — Thirty years later, in 1815, the eruption of Mt. Tambora, Indonesia, resulted in an extremely cold spring and summer in 1816, which became known as the year without a summer. The Tambora eruption is believed to be the largest of the last ten thousand years. New England and Europe were hit exceptionally hard. Snowfalls and frost occurred in June, July and August and all but the hardiest grains were destroyed. Destruction of the corn crop forced farmers to slaughter their animals. Soup kitchens were opened to feed the hungry. Sea ice migrated across Atlantic shipping lanes, and alpine glaciers advanced down mountain slopes to exceptionally low elevations.
KRAKATAU (1883) — Eruption of the Indonesian volcano Krakatau in August 1883 generated twenty times the volume of tephra released by the 1980 eruption of Mt. St. Helens. Krakatau was the second largest eruption in history, dwarfed only by the eruption of neighboring Tambora in 1815 (see above). For months after the Krakatau eruption, the world experienced unseasonably cool weather, brilliant sunsets, and prolonged twilights due to the spread of aerosols throughout the stratosphere. The brilliant sunsets are typical of atmospheric haze. The unusual and prolonged sunsets generated considerable contemporary debate on their origin.They also provided inspiration for artists who dipicted the vibrant nature of the sunsets in several late 19th-century paintings, two of which are noted here.

In London, the Krakatau sunsets were clearly distinct from the familiar red sunsets seen through the smoke-laden atmosphere of the city. This is demonstrated in the painting shown here of a sunset from the banks of the Thames River, created by artist William Ascroft on November 26, 1883.
The vivid red sky in Edvard Munch’s painting “The Scream” was inspired by the vibrant twilights in Norway, his native land.

PINATUBO (1991) — Mt. Pinatubo erupted in the Philippines on June 15, 1991, and one month later Mt. Hudson in southern Chile also erupted. The Pinatubo eruption produced the largest sulfur oxide cloud this century. The combined aerosol plume of Mt. Pinatubo and Mt. Hudson diffused around the globe in a matter of months. The data collected after these eruptions show that mean world temperatures decreased by about 1 degree Centigrade over the subsequent two years. This cooling effect was welcomed by many scientists who saw it as a counter-balance to global warming.


This section describes the variability of eruption types, from quiescent lava emissions to extremely violent, explosive events. Eruption variability is largely related to magma composition and the amount of water present. The various eruption types are typically associated with particular volcano types. Shield volcanoes, for example, generate low-viscosity basalts associated with calm, effusive eruptions. It is common to find traditional names from classic eruptions to describe other eruptions and volcano forms: Hawaiian, Strombolian, Vulcanian, Surtseyan and Vesuvian (Plinian), for example. Although these names are somewhat poorly defined and subjectively applied, they are widely used in the volcanological literature.
Height of Eruption Column and Degree of Explosivity — Whereas the height of the eruption column can be measured directly in observed eruptions, it can also be estimated in ancient eruptions by measuring the geographic dispersal of the airfall tephra. The degree to which this tephra is fragmented provides a means to measure the explosiveness of the eruption. The diagram here (modified from Cas and Wright, 1988) demonstrates how these parameters will vary with eruption type. For more detailed information on these parameters, click on image. Note that hydrovolcanic eruptions (above dashed line) are generally the most explosive, but do not necessarily generate the highest eruption columns.

Fissure eruptions should not be considered in isolation, because they are also intimately related to Hawaiian eruptions. However, the unique character of fissure eruptions warrants a separate description here.
In contrast to the point-source, centralized eruptions that typify most volcanoes, fissure eruptions are generated at several contemporaneous sites along a linear fracture, or along an en echelon (parallel, but offset) fracture system, such as that shown in the image here. Regional fracture systems can appear where the Earth’s crust is broken and pulled apart by tensional forces. If these regions are underlain by reservoirs of basaltic magma, this low-viscosity melt will utilize the fractures and ascend through the crust to generate a fissure eruption. For example, Mid-oceanic ridges (divergent plate margins) typically extrude basaltic magma from fissure eruptions because these are areas where global-scale extension is coincident with the rise of partially molten asthenosphere. Because Iceland is the subaerial extension of the Mid-Atlantic Ridge, it is one of the world’s most active sites for basaltic fissure eruptions. For this reason, fissure eruptions are also known as Icelandic eruptions. The largest lava flow in recorded history was generated by a fissure eruption in south central Iceland in 1783. Known as the Laki flow, it erupted from a 25-kilometer-long fissure to produce 12 cubic kilometers of lava, filling two deep river valleys and covering an area greater than 500 square kilometers.
Fissure eruptions are also common on the flanks of many large volcanoes and, therefore, they are not restricted to areas undergoing regional extension. Magma-filled fissures radiating from the summit regions of active volcanoes like Mt. Etna, Mauna Loa, and Kilauea propagate outward from the central vent system. Extrusion from these propagating fissures can produce elongate volcano morphologies, such as those that are typical of many Hawaiian shield volcanoes. Note, for example, the axial elongation of the Mauna Loa shield volcano shown in the image to the left. Mauna Loa fissure eruptions are generated along two axial rift zones connected at the Mokuaweoweo summit crater. Each rift zone is underlain by magma-filled fissures. The image here displays several lava flows radiating downslope from these axial rift zones. Most of these erupted in historic times.

ERUPTION STYLE: the “Curtain of Fire”
As fluid, gas-poor basaltic magma rises up through a fissure, it is extruded at the surface as a wall of incandescent, liquid-to-plastic fragments known as a curtain of fire. Two such eruptions are shown below from extrusive events on the Kilauea volcano, Hawaii. Fissure eruptions are quiescent, and the height of the airborne eruptive material is small, often only a few tens of meters. The basaltic fragments in the curtain of fire thus remain largely liquid when they hit the ground. These coherent lumps of hot, fluid lava are called spatter. When they land, they can be hot and fluid enough to fuse together to form an aggregate called agglutinate, or agglutinated spatter. Spatter commonly builds up as banks along the fissure sides to produce spatter ramparts.
When fissures cease to erupt, the remaining magma residing in the fissure will cool and crystallize into an igneous rock intrusion. The resulting rock structure is called a dike. Dikes are tabular in shape, and they cut discordantly across adjacent rock layers. In areas of ancient volcanism, dikes are often delineated as resistant walls standing above more easily eroded rock types.

Massive fissure eruptions in the geological past have generated extraordinarily voluminous lava flows that form large continental flood basalt provinces. Individual provinces can cover hundreds of square kilometers, with average thicknesses of one kilometer. These flood-basalt eruptions are rare in the geologic record. They generate huge volumes of basalt over a very short time intervals, typically in only 1-2 million years. Well-known examples include (1) the Columbia River flood basalts, the bulk of which erupted from 17-14 million years ago in the northwestern United States, (2) the Deccan flood basalts, which erupted about 65 million years ago in western India, and (3) the Siberian flood basalts, which erupted about 245 million years ago in northern Siberia.
Flood-basalt eruptions are often intimately related to rifting or to stretching of the earth’s crust above a region of hot mantle. This process can generate huge volumes of magma that rises through fractures to produce massive fissure eruptions on the surface. Basalt filled fissures on the Columbia Plateau, are currently exposed as dikes. About 14 million years ago, 700 cubic kilometers of basalt erupted from a single such fissure on the Columbia Plateau to form the Roza flow. The Roza flow is typical in volume to many of the larger flows in the Columbia River Basalt Province. These flow volumes dwarf the 12 cubic kilometers of the largest historic basaltic flow (the 1783 Laki flow), by more than an order of magnitude. Whereas the Laki flow advance ~40 kilometers from its source fissure, the largest of the Columbia River Basalt flows travelled up to 500 kilometers west of their source fissures.

PELE: the Hawaiian Goddess of Fire
Many native Hawaiians have a strong religous belief concerning Pele, the Hawaiian goddess of fire. According to legend, her spirit resides in the Halemaumau crater on the Kilauea volcano. At one time, she had a short and violent marriage to Kamapuaa, the god of water. As demonstrated in the painting shown here (courtsey of the artist, Herb Kane), Pele routed Kamapuaa from their Halemaumau home and, in a rage, chased him with streams of lava into the sea. This symbolism accurately portrays the often violent interaction of lava and water associated with explosive hydrovolcanic eruptions. Typically, however, Hawaiian eruptions are much more quiescent. The frequent outpouring of basaltic lava on Kilauea is a fitting reminder to the faithful that Pele is alive and well.
Hawaiian eruptions are the calmest of the eruption types. They are characterized by the effusive emission of highly fluid basalt lavas with low gas contents. The relative volume of ejected pyroclastic material is less than that of all other eruption types. The hallmark of Hawaiian eruptions is steady lava fountaining and the production of thin lava flows that eventually build up into large, broad shield volcanoes. Eruptions are also common in central vents near the summit of shield volcanoes, and along fissures radiating outward from the summit area. Lava advances downslope away from their source vents in lava channels and lava tubes.
Fissure.htmlFissure eruptions are common occurrences on the “Big Island” of Hawaii. They often begin as a line of vents (curatin of fire) that gives way to eruptions concentrated at one or two cental vents lying along the fissure. The Pu’u O’o eruptive series, for example, has been erupting basaltic lava on the Kilauea shield volcano since 1983. These eruptions began on January 3 with a six-kilometer-long curtain of fire on the east rift system of Kilauea. Intermittent fissure eruptions soon gave way to a centralized eruption site on the east rift, about 15 km east of the Kilauea summit caldera, which generated a scoria-and-spatter cone, called the Pu’u O’o volcano. In 1986 the Kupaianaha volcano developed about 3 kilometers farther down rift. It erupted smoothed-surface pahoehoe lava until early 1992. Since that time the main eruption site has been centered at Pu’u O’o.
The Kilauea summit caldera and east rift system are evident on the above map-view and 3D images. The blue-to-purple regions descending down the southeastern slope of Kilauea (far right) are lava flows generated during the Pu’u O’o eruptive series, through early 1994.

Central-vent Hawaiian eruptions are noted for their spectacular jet-like sprays of liquid lava called fire fountains. These incandescent jets ascend hundreds of meters into the air. They can occur in short spurts, or last for hours on end. One of the most spectacular fire fountaining events ever recorded on Kilauea produced a lava spray 580 m high at the Kilauea Iki vent in 1959. However, this is dwarfed by the 1600 m fire fountain generated by an Hawaiian eruption on the Japanese Island of Oshima in 1986. The top of fire fountains are often carried away downwind to produce an airborne curtain of glowing fragments that showers downward. The indivudual liquid-to-plastic fragments (clasts) generally cool quickly by radiating their heat into the atmosphere. Thus, they are chilled and solid by the time they hit the ground, where they accumulate as cindery fragments called scoria. However, during very high eruption rates, the fire fountains become so dense that the clasts can no longer radiate heat freely into the atmosphere. These clasts are kept hot by the heat of surrounding clasts. Under these conditions the molten clasts, spatter, may hit the ground and fuse together to form agglutinated spatter cones and spatter ramparts. If the eruption rates are high enough, spatter-fed flows (clastogenic lavas) may develop as hot spatter fragments blend together on the ground and flow away.
The smallest pyroclasts during fire-fountaining will be carried downwind from near the the top of the eruptive jet. They will chill quickly into small glassy black spheres, dumbells, or teardrop shapes called Pele’s tears. During high winds, the teardrop shapes are sometimes drawn out as long filaments, the tails of which can break off to produce Pele’s hair. During periods of high vesiculation, basalt foam can quench into the glassy rock recitulite, also known as thread-lace scoria, which has the lowest density of any know rock type.

The fluid basalt associated with Hawaiian eruptions sometimes ponds in vents, craters, or broad depressions to produce lava lakes. In some cases, lava may erupt from a vent located within a crater, or surface lava flow may pour into a crater or broad depression. The image shown here is a lava lake that occupied the Kupaianaha vent on the east rift system of Kilauea in 1986. Currently active lava lakes occur in only a few locations: Mt. Erebus in Antarctica, Erta’ Ale in Ethiopia, and Nyiragongo in the Congo. The Kilauea volcano has had an active history of producing lava lakes in its numerous craters. Perhaps longest-lived lava lake in historic times was the near permanent lake that occupied the Halemaumau crater for most of the hundred-year period between 1823 to 1924. This lake was destroyed in 1924 by a massive hydrovolcanic eruption. As lava lakes cool, they produce a grey-silver crust that is usually only a few centimeters thick, as shown here in the image of the Kupaianaha lava lake. Active lava lakes contain young crust that is continually destroyed and regenereated. Convective motion of the underlying lava causes the crust to break into slabs and sink. This then exposes new lava at the surface that cools into a new crustal layer which will again break up into slabs and be recycled into the circulating lava beneath the crust.

Strombolian eruptions are named from the small volcano-island of Stromboli (image), located between Sicily and Italy. This volcano has been erupting almost constantly for hundreds of years. It erupts irregularly every twenty minutes or so to produce an episodic lightshow that gives rise to its nickname, the “Lighthouse of the Mediterranean”.
The term “strombolian” has been used indiscriminately to describe a variety of volcanic eruptions that vary from small volcanic blasts, to kilometer-high eruptive columns. However, true strombolian activity is characterized by short-lived, explosive outbursts of pasty lava ejected a few tens or hundreds of meters into the air. Unlike Hawaiian eruptions, Strombolian eruptions never develop a sustained eruption column. They eject relatively viscous basaltic lava from the throat of the volcano. Build up of the high gas pressures required to fragment this somewhat pasty lava, results in episodic explosions with booming blasts. Although Strombolian eruptions are much noisier than Hawaiian eruptions, they are no more dangerous. As shown in the images above and below, Strombolian explosions eject bomb- and lapilli-sized fragments that travel in parabolic ballistic paths before accumulating around the vent to construct the volcanic edifice. Typically, these eruptions form scoria cones composed of basaltic pryoclasts. However, mafic stratovolcanoes can also exhibit common Strombolian activity, evident for example, at Mt. Eberus in Antarctica and at Stromboli itself.

The small Italian island of Vulcano, provides the family name for all volcanoes. Historic eruptions led the Romans to believe that this island was the forge of Vulcan, son of Jupiter and blacksmith to the Roman gods, as depicted in the image left, with permission from the artist, Jeffrey Hulen. The island of Vulcano also gives its name to a particular eruption style — vulcanian. Vulcanian eruptions initially occur as a series of discrete, canon-like explosions that are short-lived, lasting for only minutes to a few hours, often with high-velocity ejections of bombs and blocks. Once the volcano “clears its throat,” however, the subsequent eruptions can be relatively quiet and sustained. Vulcanian eruptions are more explosive than Strombolian eruptions with eruptive columns commonly between 5 and 10 km heigh. The volume of tephra produced is relatively small (less than one cubic kilometer), but dispersed over a moderately wide area. Periodic on-going vulcanian activity is currently exhibited by the Sakurajima volcano in Japan, and by the Tavurvur volcano in Papua New Guinea, shown in the following images:
In contrast to basaltic strombolian eruptions, vulcanian eruptions are most often associated with andesitic to dacitic magma. The high viscosity of these magmas makes it difficult for the vesiculating gases to escape. This leads to the build up of high gas pressure and explosive eruptions. The ejected lava fragments do not take on the aerodynamic shapes common to Strombolian eruptions. This is partly due to the higher viscosity of the erupting magma, but also because the ejecta often incorporates a high proportion of crystalline material broken away from the rock plugging the throat of the volcano. These eruptions are often associated with growing lava domes, such as that at Mt. Pelée in 1902, and with the genration of pyroclastic flows from dome collapse.
Vulcanian deposits contain large blocks and bombs near the vent. Bread-crust bombs (left) are particularly abundant. These resemble a crusty loaf of bread broken by deep cracks that often expose a frothy interior. These large pyroclastic fragments form when viscous, gas-rich magma is ejected from the vent to produce a bomb whose exterior chills quickly to a glassy or fine-grained crust while in flight. The interior of the bomb, however, continues to vesiculate on the ground, which leads to expansion of the interior and cracking of the brittle outer crust.
Although blocks and bombs are common in proximal deposits, the bulk of Vulcanian deposits is very fine grained and dominated by ash. The abundance of ash indicates a high degree of fragmentation, which can only be generated by magmas with high gas contents. In some cases, these high gas contents are derived from heated meteoric water. It is likely, therefore, that many vulcanian eruptions are at least partially hydrovolcanic. Although the ash-fall deposits generated by volcanian eruptions are highly fragmented, they are only moderately dispersed. This suggests a high degree of explosiveness (high fragmentation) associated with the development of eruptive columns that are of only moderate heights (moderate dispersal).

Plinian (or Vesuvian) eruptions typify the well-known historic eruptions that produce powerful convecting plumes of ash ascending up to 45 kilometers into the stratosphere. These explosive eruption types are named after Pliny the Younger, a Roman statesman who wrote a remarkably objective account of the eruption of Italy’s Mt. Vesuvius (left) in 79 AD. Pliny’s uncle, Caius Plinius (Pliny the Elder), was a much respected naturalist and Admiral in the Roman navy who died during the eruption. To properly record the circumstances of his esteemed uncle’s death, Pliny the Younger wrote two letters to the historian Tactius describing the Mt. Vesuvius eruption. The eruption killed thousands of people and buried the Roman towns of Pompeii and Herculaneum under huge volumes of tephra, pyroclastic flows, and lahars. Pompeii laid buried for over 1700 years until it was rediscovered by accident during the excavation of a water line. Uncovering the remains of Pompeii has not only broadened our understanding of Plinian-type eruptions, but it has also provided a unique understanding of the lives of ordinary people during Roman times.
These spectacularly explosive eruptions are associated with volatile-rich dacitic to rhyolitic lava, which typically erupts from stratovolcanoes. The duration these eruptions is highly variable, from hours to days. The longest eruptions appear to be associated with the most felsic volcanoes. Although Plinian eruptions typically invlove felsic magma, they can occasionally occur in fundamentally basaltic volcanoes where the magma chambers become differentiated and zoned to create a siliceous top. An example of this was the Hekla eruption (Iceland) of 1947-48. Over the past 800 years, Hekla has had a history of generating violent initial eruptions of pumice, lasting a few hours, followed by prolonged extrusion of basaltic lava from the lower part of the chamber.
Rather than producing the discrete explosions that are typical of Vulcanian and Strombolian eruptions, Plinian eruptions generate sustained eruptive columns. Although they differ markedly from nonexplosive Hawaiian eruptions, Plinian eruptions are similar to Hawaiian fire fountaining in that both of these eruption types generate sustained eruption plumes. In both, the eruption plumes are maintained because the growing bubbles rise at about the same rate as the magma moves up through the central vent system.
Plinian eruptions generate large eruptive columns that are powered upward partly by the thrust of expanding gases, and by convective forces with exit velocities of several hundred meters per second. Some reach heights of ~45 km. These eruptive columns produce widespread dispersals of tephra which cover large areas with an even thickness of pumice and ash (see pumice-fall deposits). The region of pyroclastic fall accumulation is generally asymmetric around the volcano as the eruptive column is carried in the direction of the prevailing wind, as shown here in this NASA image of the Klyuchevskaya eruption in 1994.

The Plinian eruption of the Klyuchevskaya volcano on Russia’s Kamchatka Peninsula was captured here by the Space Shuttle Endeavor in 1994. The prevailing wind direction produced a pronounced asymmetry of the plume head. Courtesy of NASA.
The regions surrounding Plinian eruptions are not only subject to large volumes of pumice airfall (from 0.5 to 50 km3), but they are also subject to the most dangerous types of volcanic phenomena: pyroclastic flows and lahars. The occasional collapse of the eruptive column will generate hot, pyroclastic flows that advance down the volcano flanks at hurricane-force speeds. In addition, large volumes of water are often generated by the melting of snow banks and alpine glaciers during the eruption. The mixing of this water with unconsolidated tephra can generate volcanic mudflows (lahars). These features have the consistency of wet concrete, yet they can advance down slopes at the same rate as a rapidly moving stream.
The human devastation associated with the Plinian eruption of Mt. Vesuvius in 79 AD is largely attributed to all of these volcanic phenomena. Pompeii was located to the southeast, on the downwind side of the volcano. Not only was it subjected to the destructive force of several pyroclastic flows, but it was also buried under a huge thickness of airfall tephra. Although the village of Herculaneum was also destroyed by the eruption, it was located west of the volcano, and was not subjected to the same volume of airfall tephra that buried Pompeii. Instead, Herculeaneum was largely buried by pyroclastic flows and massive lahars which advanced down the volcano’s western flank.

Hydrovolcanic eruptions are generated by the intereaction of magma with either groundwater or surface water. Explosive hydrovolcanic eruptions of basaltic lava are sometimes called Surtseyan, after the eruption off Iceland in 1963. Surtseyan eruptions are considered to be the “wet” equivalents of Strombolian-type eruptions, although they are much more explosive. This high explosivity is a hallmark of hydrovolcanic activity. As the water is heated, it flashes to steam and expands explosively, thus fragmenting the magma into exceptionally fine-grained ash. When the volcanic island of Surtsey was born in the Atlantic, the initial hydrovolcanic eruptions were spectacularly explosive. As the volcano grew, however, the rising lava in the central vent interacted with water to a lesser degree, so that the waning stages of the eruption became more Strombolian in character.
Hydrovolcanic eruptions are not restricted to the underwater development of oceanic islands. Many explosive Surtseyan events are generated on land by rising conduits (diapirs) of basaltic magma that interact with water-bearing strata (e.g., aquifers) at shallow levels beneath the surface. The two examples are shown here are from Ukinrek in Alaska and Capelinhos in the Azores.
Note the radial cloud emanating from the base of the Capelinhos eruptive column. This phenomenon is a base surge, a characteristic feature of many Surtseyan-type eruptions. These surges are analogous to the ring-shaped, ground-hugging clouds that travel radially away from the vertical columns generated by nuclear explosions. Base surges are derived from the gravitational collapse of the “wet” eruptive column, which is denser than those associated with “dry” eruptions. Base surge deposits are wedge-shaped, with their thickest end near the vent. Dune-shaped deposits are common near the vent, indicated lateral transport analogous to the lateral movement and deposition of sand grains along the face of a moving sand dune. The bedding is often disrupted by “bomb sags” containing large ejected blocks which followed ballistic trajectories unrelated to the lateral surge. Spherical accumulations of wet, accreted ash, called accretionary lapilli, are also common to many base surge deposits.

Hydrovolcanic explosions generate maars and tuff rings. These vent types are large circular depressions with low rims of ejected debris. Maars are excavated into the substrate, thus exposing older rocks along their inner walls. Tuff rings, however, are built above the substrate. Maars contain a greater proportion of fragmented basement rocks in the ejecta blanket. This suggests that maars are derived from steam blasts (phreatic eruptions) generated well above the diapiric intrusion. Tuff rings, on the other hand, contain a greater proportion of magmatic (juvenile) fragments (see for example, palagonite tuff). Such deposits are consistent with explosions derived from a combination of heated groundwater and vesiculating magma (phreatomagmatic eruptions) from relatively shallower intrusions.
Ciruclar-shaped maars and tuff rings appear to be generated above rising columns of magma (diapirs). However, if the magma rises along a linear fracture zone, the interaction with groundwater may result in the reaming out of the fissure. Such eruptions can be even more destructive than those associated with maars and tuff rings, as exemplified by the highly explosive Tarawera eruption which buried three villages and killed 150 people on New Zealand’s North Island in 1886.

Although littoral cones have a hydrovolcanic origin, they are not true hydrovolcanic vents. They are composed of basalt tephra that accumulate as ejected debris from the explosive interaction of moving lava and seawater. Although pahoehoe flows generally enter the water in a relatively passive manner, a’a flows often enter the sea explosively to produce cone-shaped piles of ejected debris. The irregular cracks bounding cindery blocks of a’a allow water to penetrate into the interior of the hot flow where it then flashes to steam explosively. Basalt exiting from lava tubes can also generate littoral cones from episodic explosions due to disruption of the lava stream by incoming waves or swells. The littoral cones that develop from these explosive ejections are typically open on the seaward side.
The various types of eruptions described in this section can be simulated with an extraordinary computer program developed by Dr. Ken Wohletz. For a free download of this program, double click Erupt 2.5.

There are on average about ten volcanoes erupting on earth on any given day. Most of these events are small eruptions that may be precursors to larger eruptions. There have been many cataclysmic eruptions in historic times. Since 1800, there have been 19 eruptions on earth in which a thousand or more people have died.

OVERVIEW: The Tropical Paradise of St. Pierre
The infamous volcano of Mt. Pelée, shown in this 1987 photo, looms over the village of St. Pierre on the French Caribbean Island of Martinique. This sleepy little village shows little of the grandeur of turn-of-the-century St. Pierre, which was a vibrant colonial city, known to European tourists as the “Paris of the West Indies.” With its red-tiled cottages, rambling streets, and tropical vegetation, this prosperous little city was renowned for its beauty. In the official 1894 census, the population of St. Pierre was around 20,000. Although most were native Martiniquans, the wealth and political power were controlled largely by Creoles and a few French colonial officials and civil servants. No one at the time could have predicted the horror that was to descend on this tropical paradise with the reawakening of Mt. Pelée in the Spring of 1902.
Although in January 1902 Mt. Pelée began to show an abrupt increase in fumarole activity, the public showed little concern. This changed, however, on April 23 when minor explosions began at the summit of the volcano. Over the next few days, St. Pierre was rocked by earth tremors, showered in ash, and enveloped in a thick cloud of choking sulfurous gas. These nightmarish conditions deteriorated further when the city and outlying villages were invaded by ground-dwelling insects and snakes driven from the slopes of Mt. Pelée by the ashfalls and tremors. Horses, pigs, and dogs screamed as red ants and foot-long centipedes crawled up their legs and bit them. Thousands of poisonous snakes joined the fray. An estimated 50 humans, mostly children, died by the snake bites, along with some 200 animals.
As the summit eruptions intensified, water in the Etang Sec crater lake was heated to near boiling. On May 5, the crater rim gave way, sending a torrent of scalding water cascading down the River Blanche. The hot water mixed with loose pyroclastic debris to generate a massive lahar with a downslope speed of nearly 100 kilometers per hour. This large volcanic mudflow buried everything in its path. Near the mouth of the river, north of St. Pierre, it overran a rum distillery, killing 23 workmen. The lahar continued into the sea, where it generated a three-meter-high tsunami which flooded the low-lying areas along the waterfront of St. Pierre.

Living near the volcano became increasingly stressful, leading many to consider leaving St. Pierre for Martinique’s second city, Fort-de-France. On the day of the lahar, however, Governor Louis Mouttet received a report from a committee of civic leaders who climbed the volcano to assess the danger. The only scientist in the group was a local high school teacher. The report stated that “there is nothing in the activity of Mt. Pelée that warrants a departure from St. Pierre.” It concluded that “the safety of St. Pierre is completely assured.” The report eased the public’s fears, and gave hope to city officials who were particularly anxious that voters remain in the city to cast their ballots for an election that was to be held on May 11. The only people with enough money to leave the island were the wealthy, nearly all of which belonged to the Progressive Party of Governor Mouttet. Mouttet convinced the conservative editor of the daily newspaper Les Colonies to downplay the danger of the volcano, and to lead the effort to encourage people to remain. Still, some residents left the city for Fort-de-France. This prompted Governor Mouttet to send in troops to patrol the road to Fort-de-France, with orders to turn back refugees who were trying to leave. Based on the soothing articles that appeared in Les Colonies, many people in the countryside flocked to St. Pierre thinking that it was the safest place to be. The population ballooned to about 28,000, nearly all of which would perish in the cataclysmic eruption of May 8.

The election scheduled for May 11 would not take place. The report issued by the investigating committee on May 5, failed to realize the potential danger of a large V-shaped notch cut through cliffs surrounding the summit crater. The notch was like a colossal gun sight pointing directly at St. Pierre four miles below. At about 7:50 a.m. on May 8, the volcano erupted with a deafening roar. A large black cloud composed of superheated gas, ash and rock rolled headlong down the south flank of Mt. Pelée at more than 100 miles per hour, its path directed by the V-shaped notch at the summit. In less than one minute it struck St. Pierre with hurricane force. The blast was powerful enough to carry a three-ton statue sixteen meters from its mount. One-meter-thick masonary walls were blown into rubble and support girders were mangled into twisted strands of metal. The searing heat of the cloud ignited huge bonfires. Thousands of barrels of rum stored in the city’s warehouses exploded, sending rivers of the flaming liquid through the streets and into the sea. The cloud continued to advanced over the harbor where it destroyed at least twenty ships anchored offshore. The hurricane force of the blast capsized the steamship Grappler, and its scorching heat set ablaze the American sailing ship Roraima, killing most of her passengers and crew. The Roraima had the misfortune of arriving only a few hours before the eruption. Those on on board could only watch in horror as the cloud descended on them after annihilating the city of St. Pierre. Of the ~28,000 people in St. Pierre, there were only two known survivors.
The dynamic cloud of hot gases and incandescent solid particles that destroyed St. Pierre was a pyroclastic flow, a feature that was unknown to science at the time. Subsequent examples observed on Mt. Pelée were described by French volcanologists as nuée ardentes, or glowing clouds.

Although there were only two survivors in St. Pierre, there were other survivors on the outskirts of the town and in some of the ships moored in the harbor. Mercifully, death came quickly to those that perished. Some may have died by the sheer force of the blast, but most died within a few seconds after inhaling the scorching fumes and ash of the pyroclastic flow. The throats and lungs of most of the deceased were seared and their bodies badly burned.
The tales of the two male survivors of St. Pierre are described briefly here, as well as the astonishing story of a young girl who looked straight into the mouth of a volcanic vent just before Mt. Pelée began to erupt.

A young shoemaker, Léon Compere-Léandre, was sitting on his doorstep when the nuée ardente hit. Although he was severely burnt he survived, partly because of his good health, but also because his house was near the edge of the pyroclastic flow. Here is his experience, in his own words:
“I felt a terrible wind blowing, the earth began to tremble, and the sky suddenly became dark. I turned to go into the house, with great difficultuy climbed the three or four steps that separated me from my room, and felt my arms and legs burning, also my body. I dropped upon a table. At this moment four others sought refuge in my room, crying and writhing with pain, although their garmets showed no sign of having been touched by flame. At the end of 10 minutes one of these, the young Delavaud girl, aged about 10 years, fell dead; the others left. I got up and went to another room, where I found the father Delavaud, still clothed and lying on the bed, dead. He was purple and inflated, but the clothing was intact. Crazed and almost overcome, I threw myself on a bed, inert and awaiting death. My senses returned to me in perhaps an hour, when I beheld the roof burning. With sufficient strength left, my legs bleeding and covered with burns, I ran to Fonds-Sait-Denis, six kilometers from St. Pierre.”

The only other known survivor in St. Pierre became a minor celebrity. He was a husky 25-year-old roustabout named Louis-Auguste Cyparis, locally known simply as “Samson”. In early April, Samson was put in jail for wounding one of his friends with a cutlass. Towards the end of his sentence, he escaped from a labouring job in town, danced all night, and then turned himself into the authorities the following morning. For this, he was sentenced to solitary confinement for a week in the prison’s dungeon. On May 8, he was alone in his dungeon with only a small grated opening cut into the wall above the door. While waiting for his breakfast, his cell became dark and he was overcome by intense gusts of hot air mixed with ash that had entered through the grated opening. He held his breathe while experiencing intense pain. After a few moments, the heat subsided. He was severally burned, but managed to survive for four days before he was rescued by people exploring the ruins of St. Pierre. After he recovered, he received a pardon and eventually joined the Barnum & Bailey Circus, where he toured the world billed as the “Lone Survivor of St. Pierre.”

One of the most incredible escapes from Mt. Pelée was that of a young girl named Havivra Da Ifrile. Very early in the morning of May 8, Havivra was on her way to services at the catherdral in St. Pierre when her mother sent her on an errand. She was to walk to her aunt’s pastry shop near a local tourist attraction known as the Corkscrew. The “Corkscrew” was named after a tourist trail that wound down into an ancient crater, or parasitic cone, located halfway up the flank of the volcano. As Havivra approached the Corkscrew, she noticed smoke rising from the crater. After looking into the crater, she described it in this manner: “There I saw the bottom of the pit all red, like boiling, with little blue flames coming from it.” She apparently saw three people trying to run up the Corkscrew before they were engulfed in “. . . a puff of blue smoke . . ” and “. . . fell as if killed.” She fled toward St. Pierre.
“Just as I got to the main street I saw this boiling stuff burst from the top of the Corkscrew and run down the side of the hill. It followed the road first, but then as the stream got bigger, it ate up the houses on both sides of the road. Then I saw that a boiling red river was coming from another part of the hill and cuttung off the escape of the people who were running from their houses.”
Frightened, Havivra ran to the shore and jumped into her brother’s small boat and headed along the shore to a cave that she used to play pirate in with her friends.
“But before I got there I looked back — and the whole side of the mountain which was near the town seemed to open and boil down on the screaming people. I was burned a good deal by the stones and ashes that came flying about the boat, but I got to the cave.”
While in the safety of the cave, she heard a hissing sound as the hot pyroclastic debris entered the water. The last thing she remembered before lapsing into unconsciousness was the water rising rapidly toward the roof of the cave. She was later found by the French cruiser Suchet drifting two miles out to sea in her charred and broken boat.


The Apollo missions to the moon in the 1960’s set the stage for space exploration in the 1970’s, 80’s, and 90’s through the Mariner, Pioneer, Viking, Voyager, Magellan, and Galileo missions. The data gathered by these missions has greatly expanded our knowledge of volcanism on other worlds. Volcanic features comparable to those on earth are particularly well exposed on the Earth’s Moon, Mars, Venus, and the Jovian moon, Io. The most stunning discoveries of the 1990’s are the great variety of spectacular volcanic features displayed by radar imagery of Venus, and the dramatic discovery of perennial eruptions on Jupiter’s satellite Io.

The moon, called Luna by the Romans, is our closest celestial neighbor and our only natural satellite. With a diameter of 3476 km (2086 mi), this rocky mass orbits the earth once every 29.5 days. The image to the left shows the stark contrast between the gray, highly cratered lunar surface in the foreground and our blue planet Earth in the distant background. The average orbital separation between these two bodies is 384,400 km (230,640 mi). The moon was first visited by the Soviet spacecraft Luna 2 in 1959. A decade later, on June 20, 1969, it was first visited by humans. The last of these historic visits, by Apollo Mission astronauts, was in December 1972. The lunar surface was extensively mapped by the spacecraft Clementine in 1994, and again by the Lunar Prospector in 1999.

The moon contains a ~70-km-thick basaltic crust that differentiated from a denser, underlying mantle at least 4.4 billion years ago. The crustal rocks returned by the Apollo Mission range in age from 4.3 to 3.1 billion years old. In contrast, the oldest rocks on earth are rarely more than 3 billion years old. The lunar crust is easily discernible into two primary types of terrains: the older highlands, which comprise about 83% of the lunar surface (light regions in image), and the younger lunar maria, which comprise the remaining 17% (dark regions). The heavily cratered highlands are covered with a layer of regolith, a mixture of fine dust and fragmented debris generated by meteorite impacts. Beneath the regolith, two crustal rocks types dominate the highlands:
(1) breccia — a coherent rock of broken and welded fragments, and
(2) anorthosite — the most abundant highlands rock; it is a feldspar-rich variety of gabbro (which, in turn, is the coarse-grained, equivalent of basalt)

Compared with the highlands, the lunar maria (mare, singular) are much darker and have relatively smooth surfaces. The maria are more abundant on the near side of the moon, where they cover about 30% of the lunar surface; on the far side they cover only about 2%. The reason for this is not well understood. Each mare is composed of vast sheets of basaltic lava which erupted from the broken crust associated with large impact craters. There is no evidence to suggest that the impact events triggered the volcanic eruptions. However, the fractures associated with large impact craters may have provided easy egress for younger, mantle-derived melts to rise to the surface. In a sense, these extensive lava fields are the lunar equivalents of terrestrial flood basalt provinces. The maria range in thickness from 1-4 km.

Mare Basalt — This sample of mare basalt, ~3.7 billion years old, was collected by Apollo 17 astronauts. The small holes in the sample are vesicles, which formed by the degassing of the lava as it cooled.

The highlands anorthosite is dominated by the mineral plagioclase feldspar. The most popular model for crustal evolution suggests that the anorthosite crystallized early in the moon’s formation from a global magma ocean. Under this scenario, the lighter plagioclase floated upward to form an anorthosite crust, and the heavier minerals olivine and pyroxene sank toward the bottom of the magma ocean to form the moon’s upper mantle. The greatest volume of the highland crust was generated between 4.5 and 3.9 billion years ago (bya), but it continued to develop up to ~ 3.2 bya. Episodic melting of the mantle rocks beneath the highland crust generated the younger mare basalts. Heavy bombardment of the lunar crust in its early stages of formation (4.5-3.2 bya) provided fractures for rapid ascent of these mantle melts and depressions for them to pond in. The greatest volume of the mare basalts erupted from 3.9 to 3.2 bya, but reduced bombardment and basaltic volcanism continued until ~2.0 bya. Since ~2.0 bya, there has been only very minor eruptions, local degassing, and occasional impacts.

High-resolution photographs of Mare Imbrium reveal spectacular lava flow fronts, extending as much as 1200 km from their source vents. However, these features appear to be unique to Mare Imbrium. Most other mares have smooth surfaces upon which discrete flows cannot be recognized. These are probably composed of compound flows composed of numerous thin flow units.
The mare lava flows are associated with a number of volcanic vents, in the form of domes and scoria cones. The domes are typically found in clusters. They are circular to elliptical, typically a few kilometers across and several hundred meters high. The basaltic domes on the moon appear to be the lunar equivalents of shield volcanoes found on earth. Construction of the relatively steep slopes of the lunar “shields” is poorly understood. Their steep slopes are probably not related to the extrusion of relatively viscous basalt, but rather to the style of extrusion, which may involve eruptions of short duration mixed with minor episodes of ash eruption. The lunar scoria cones are similar in size to terrestrial scoria cones, and they appear to be associated with older fissure systems.
A common and intriguing volcanic landform found in the lunar maria are narrow, winding valleys called sinuous rilles. These range in width from tens of meters to 3 km, and in length from a few kilometers to 300 km. Many originate in irregular craters. One of the largest of these rilles is Hadley Rille in the Imbrium basin, which was visited by Apollo 15 astronauts. Although most workers believe that the sinuous rilles are a combination of lava channels and collapsed lava tubes. However, the mode of lava channel formation remains controversial. Some believe that the channels are “constructional” due to the build up of lava levees along the sides of the channel. However, the majority of the sinuous rilles are thought to the product of thermal erosion of the mare surface by exceptionally hot, fluid lava, similar in composition to the ancient eruptions of komatiite on the Earth’s surface billions of years ago.

Dr. Eugene Shoemaker was among the world’s leaders in the study lunar and planetary geology. He established the Branch of Astrogeology with the U.S. Geological Survey, he helped train the Apollo Astronauts, and he was head of the scientific team to map the moon’s surface during Project Clementine in 1994. He helped establish the Sky Watch program for tracking earth-crossing asteroids and comets. The program culminated in the discovery (with his wife Carolyn and David Levy) of the Shoemaker-Levy comet that collided spectacularly with Jupiter in 1994. In 1992, he received the highest scientific honor in the U.S., the National Medal of Science. Tragically, Dr. Shoemaker died in a car accident in 1997. He had always said that his greatest unfulfilled dream was to go to the moon. As a tribute to this great man, his ashes were carried to the moon during the Lunar Prospector Mission in 1998, thus making Eugene Shoemaker the first human to have his remains buried on another planetary body.

Mars, “the red planet,” derives its name from the Roman god of war. With a diameter of 6,794 km, it is about half the size of the earth. Like earth, it contains a thin crust overlain by a rocky, partially molten mantle. The crustal rocks are composed of basaltic lavas. Although Mars may have been tectonically active in its early history, there is no evidence of geologically recent horizontal motion similar to plate tectonic motion on earth. Volcanism on Mars may be partly derived from the melting of hot mantle plumes. The lack of lateral motion means that these mantle hotspots would remain fixed relative to the Martian surface. This mechanism may account for regional uplift and volcanism at several sites on the Tharsis bulge, the extensive light-colored region on the west side of the image shown here. The Tharsis bulge is associated with a number of exceptionally large volcanoes, which show as dark red spots on the western side of the image. The enormous Valles Marineris canyon system is visible in the center of the image, east of the Tharsis bulge. This huge feature contains numerous canyons, with depths from 2 to 7 km, exposed over a distance of 4500 km (for a close-up view, double-click Valles Marineris). The Marineris canyons were not eroded by running water, but rather by stretching and cracking of the crust during mantle-plume uplift of the Tharsis bulge. However, there is clear evidence of ancient water erosion on Mars, which may date back to nearly 4 billion years ago.
There have been 23 scientific missions to Mars, which include Martian flybys, orbiters, and unmanned landings. The first spacecraft to land on Mars was Mariner 4 in 1965, followed by Viking 1 in 1976 and Mars Pathfinder in 1997.

The volcanic features on Mars are very similar in shape, but not in scale, to those found on earth, and they probably formed by similar processes. In contrast to the volcanic features found on the moon, those on Mars were generated on terrains of variable ages. Thus, volcanism has been an important process throughout Martian history. Numerous volcanic landforms can be found in the older cratered highlands and in the younger volcanic plains surrounding them. However, the most impressive volcanic landforms are associated with the extensive, hotspot-related uplifts of Tharsis and Elysium plateaus. The volcanic features described below include the giant central volcanoes, peterae, tholi, and rootless cones.

The most spectacular volcanic features on Mars are the isolated, giant basaltic shield volcanoes called Montes (singular Mons), Latin for mountain. The largest of these are four giant shield volcanoes associated with Tharsis uplifted region (for a close-up view of two of these, double-click Tharsis Shields). All four dwarf the Mauna Loa shield volcano, the largest mountain on earth. The largest of the four is Olympus Mons, shown in the image here. Olympus Mons is the largest volcano in the solar system, with a base diameter of ~600 km and 25 km of relief from the summit to the plains surrounding it abrupt basalt scarp (see image).
These large basaltic shield volcanoes contain large central calderas and gentle slopes that extend outward at angles usually less than five degrees. The volcano flanks have radial surface textures formed from lava channels and long narrow flows.
The large scale of the Tharsis shield volcanoes suggests that they formed from massive eruptions of fluid basalt over prolonged periods of time. Similar eruptions on earth are associated with flood basalt provinces and mantle hotspots. However, on earth the source region for hotspot volcanism moves laterally as lithospheric plates travel across the stationary mantle plumes beneath them. Without this mechanism of lateral movement, the Martian surface remains above the plume source so that huge volumes of lava will erupt from a single central vent over many millions of years of activity, thus generating a single shield volcano of enormous volume. With this in mind, it is interesting to note that the volume of Olympus Mons is roughly equivalent to the total volume of basalt in the Hawaiian-Emperor seamount chain.

Compared with the Mons volcanoes, Tholi and Peterae are much smaller volcanic vents. A Tholus volcano is an isolated domal mountain with a central crater, having slopes that are usually steeper than those associated with both Mons and Peterae volcanoes. A Patera volcano is dominated by an irregular or complex caldera with scalloped edges, surrounded by very gentle slopes of usually less than one degree. The lower flanks of both Tholi and Paterae are often buried by more recent volcanic plains. This suggests that some of these volcanoes may be the tops of buried shield volcanoes. Several Tholi and Paterae volcanoes contain sinuous channels carved into their flanks, probably by hot effusive lava flows. Some researchers, however, suggest that the channels were eroded by pyroclastic flows. The scoured flanks of many Tholi and Paterae suggest that they may indeed be associated with a component of pyroclastic activity, in contrast to the massive effusive activity associated with Mons-type eruptions.

High-resolution images from the Mars Global Surveyor have revealed numerous, small cone-shaped structures (< 250 m in diameter) superimposed on the surface of fresh lava flows. However, none of the cones appears to be the source of erupted lava. Similar volcanic features on earth are said to be “rootless,” in that they are not positioned above a magma source. Some researchers suggest that these small structures were produced by the interaction of molten lava with a water-rich source. This idea has stimulated a significant amount of interest because, if true, these Martian psuedocraters may mark the locations of shallow water or ice at the time the lava was emplaced.

Volcanism on Mars
Mars has the largest shield volcanoes in the solar system. It also has a wide range of other volcanic features. These include large volcanic cones, unusual patera structures, mare-like volcanic plains, and a number of other smaller features. However, volcanic features are not common. There are less than 20 named volcanoes on Mars, and only 5 of these are giant shields. Also, volcanism occurs mostly within three regions. Even the mare-like plains cluster near these regions. The main cluster of volcanoes and lavas is in Tharsis. A much smaller cluster of three volcanoes lies in Elysium. Lastly, a few paterae are near the Hellas impact basin.
Differences from Moon
Like the Moon, volcanism on Mars is very old. The mare-like plains on Mars are the same age as the lunar mare, roughly 3 to 3.5 billion years old. However, volcanism lasted much longer on Mars than on the Moon. It also seems to have changed over time. Volcanism in the highland paterae and mare-like plains on Mars stopped 3 billion years ago, but some of the smaller shields and cones erupted only 2 billion years ago. The giant shield volcanoes are even younger. These volcanoes formed between 1 and 2 billion years ago. The youngest lava flows on Olympus Mons are only 20 to 200 million years old. These flows are very small, however, and they probably represent the last gasp of martian volcanism. Thus, the odds of finding an active volcano on Mars today are very small.
Like the Moon, Mars shows no sign of plate tectonics. It has no long mountain chains, and there is no clear global pattern to the volcanism. Over half of Mars is heavily cratered like the lunar farside. Unlike the Moon, however, most martian volcanism lies outside large impact basins. Instead, the mare-like plains are mostly near the largest volcanoes. These plains also are not limited to the lowest elevations. Indeed, some lava plains are much higher than the cratered uplands. Lava plains may lie at lower elevations as well. However, thick layers of dust and sediment cover both the Northern Lowlands and the large basin floors. These layers reflect a long history of winds, glaciers and flood events. They also hide any volcanism that may have occurred in the low areas on Mars.

The concentration and duration of volcanism into these two regions are attributed to the evolution of a long-lived mantle hotspot.

Although Venus is similar to Earth in size, mass, and density, its surface environment is hostile, with mean surface temperatures of 475 degrees Centigrade, a lack of water, and a thick atmosphere dominated by CO2 gas. More than 98% of the Venusian surface has been mapped at a resolution of about 100 meters by NASA’s Magellan mission in 1990 through 1994. Regional variations in surface height are shown in the image left, where highlands are depicted in white to tannish colors, uplands in greenish colors, and lowlands in bluish colors. The entire surface of Venus is younger than expected due to a massive resurfacing event associated with flood basalt volcanism. These vast lava plains are best preserved in the lowland regions. The relative ages of impact craters on the lava plains suggest that this voluminous outpouring of basalt could be as young as 300 million years ago, or as old as 1600 million years ago. All of the well-developed volcanic features discussed below appear to be younger than the global resurfacing event.

The lack of water erosion on Venus adds to the preservation of volcanic edifices on the planet’s surface. Thus far over 1700 volcanic centers have been identified on Venus. These vary greatly in size, shape, and in the nature of eruptive activity. Most of these landforms can be subdivided into one of the following catagories: central-vent volcanoes, domes, ticks, calderas, coronae, and arachnoids.

Central-vent volcanoes are characterized by the radially dispersed lava flows extruded from near the summit of a well-defined, generally circular, volcanic edifice. These volcanoes are highly variable in size, although most have diameters <100 km. Most, but not all, of the Venusian volcanoes resemble shield volcanoes on earth. Two of the largest of these are Sif Mons (2 km high and 200 km in diameter) and Sapas Mons (1.5 km high and 400 km in diameter).

Over 152 volcanic domes have been recognized on Venus. Most are circular in form and steep sided, with one or more summit craters and radial fractures. Their upper surfaces, however, vary from convex to exceptionally flat. Venusian domes are much larger than their terrestrial counterparts, with diameters ranging from 20 to 30 km and heights averaging about 300 meters. The flat-topped domes on Venus are sometimes referred to as pancake domes.
The composition of the domes on Venus is unknown. Although their smooth surfaces are consistent with basaltic compositions, their steep sides suggest that they are composed of viscous lava, consistent with high-silica, rhyolitic compositions. The concentric and radial fracture pattern shown by many domes (see above) is consistent with the stretching of dome surfaces during their formation as lava wells up from interior vents and spreads laterally on the surface.

A group of enigmatic volcanic edifices, referred to as ticks, are similar in some respects to volcanic domes in that they are circular, steep-sided, and of similar diameter. They differ, however, in their association with lava flows that breach one or more sides of the structure. The west side of the image shown here, for example, contains dark flows that emanate from a shallow summit crater and traveled west along a lava channel. The origin of these structures is not well understood. Ticks are generally associated with rift zones. One theory suggests that they are generated by dome formation within a rift, followed by localized extrusion of lava along one of the rift fractures. Thus far, about 50 such ticks have been mapped on the surface of Venus.

The term caldera is used here to describe circular to oblong depressions formed by surface collapse, similar to those found on earth. Like their terrestrial counterparts, they are associated with numerous, closely spaced ring fractures along the topographic rim of the structure. A total of 97 calderas have thus far been identified on Venus. Most range from 60-80 km in diameter, with depths up to several kilometers. In many cases, the caldera floors are covered with post-collapse lava flows.

Coronae are immense circular to oblong structures, generally between 200 and 250 km in diameter, although some can be as large as 1100 km across. These features appear to be unique to Venus. They have positive topographic expressions, with a surrounding moat or trough, and typical heights of more than 2 km above the surrounding plains. Their interiors, however, can be either topographically positive (dome-shaped) or negative (bowl-shaped). They are further characterized by closely spaced ring-shaped ridges or partially closed fractures. Radial fractures may also be present. The interiors of most coronae contain numerous, smaller volcanic features, such as lava flows, calderas, shield volcanoes, domes, and lava channels.
The genesis of coronae appears to be associated with the arrival large mantle plumes beneath the Venusian lithosphere. The progressive evolution of these features involves both uplift and collapse, either concurrently, in sequence, or in cycles. Ring fractures develop as the surface bulges outward above the rising mantle plume, but subsequent relaxation of the uplift often results in collapse of the corona’s interior along the concentric fractures. The abundant volcanic features found in the interior reflect melting of the plume head and subsequent extrusion of these melts at the surface. Thus far, 208 separate coronae have been identified.
Arachnoids are smaller than most coronae, but have a similar genesis. They features differ from coronae in that they contain a network of both radial and concentric structures that gives the appearance of a spider sitting on interconnected webs of bright lines, hence the name arachnoid, which is Greek for spider. The radial structures appear to be dike intrusions that stand out as ridges. Although it is thought that both coronae and arachnoids form in association with rising plumes of mantle, the magmatism associated with arachnoids may be largely intrusive, whereas magmatism associated with coronae is largely extrusive. There have been 265 arachnoids thus far identified on Venus.

Volcanoes on Venus
Venus has more volcanoes than any other planet in the solar system. Over 1600 major volcanoes or volcanic features are known (see map), and there are many, many more smaller volcanoes. (No one has yet counted them all, but the total number may be over 100,000 or even over 1,000,000). These volcanoes come in a variety of forms. Most are either Large Shields or Smaller Shield volcanoes, but there are also many Complex Features, several Unusual Constructs, and a few Large Flow Features. None is known to be active at present, but our data is very limited. Thus, while most of these volcanoes are probably long dead, a few may still be active.
Venus is like the Earth in many ways. It is nearly the same size and it has a similar bulk composition. Of all the planets, its orbit around the sun is the closest to EarthUs orbit. It has both clouds and a thick atmosphere. Like the Earth, it even has a fairly young surface age (~500 million years). Despite these basic similarities, however, Venus differs greatly from the Earth in detail.
First, since the atmosphere is mostly CO2, Venus has an extreme Greenhouse Effect. In fact, the surface temperature on Venus is about 470!C (about 880!F). Further, the surface air pressure on Venus is about 90 times greater than that at sealevel on Earth. This is roughly equivalent to the WATER pressure on Earth one kilometer beneath the oceanUs surface. These surface conditions have two effects. (1) There is no water on the surface of Venus. Indeed, there is almost no water in the air, either. The clouds are mostly made of sulfuric acid and they are much, much higher than most clouds on the Earth. (2) Due to the high atmospheric pressure, the winds on Venus are also relatively slow. Thus, neither wind nor rain can really affect the surface on Venus. As a result, volcanic features will look freshly formed for a long time.
Second, Venus shows no evidence for plate tectonics. There are no long, linear volcano chains. There are no clear subduction zones. Although rifts are common, none look like the mid-ocean ridges on Earth. Also, continent-like regions are rare, and show none of the jigsaw fits seen on Earth. Thus, where volcanism on Earth mostly marks plate boundaries and plate movements, volcanism on Venus is much more regional and much less organized.
Third, volcanism on Venus shows fewer eruptive styles than on the Earth. Almost all volcanism on Venus seems to involve fluid lava flows. There is no sign of explosive, ash-forming eruptions on Venus, and little evidence for the eruption of sludgy, viscous lavas. This may reflect a combination of several effects. First, due to the high air pressure, venusian lavas need much higher gas contents than Earth lavas to erupt explosively. Second, the main gas driving lava explosions on Earth is water, which is in very short supply on Venus. Lastly, many viscous lavas and explosive eruptions on Earth occur near plate subduction zones. Thus, the lack of subduction zones should also reduce the likelihood of such eruptions on Venus.

Large Shield Volcanoes
Venus has over 150 large shield volcanoes. These shields are mostly between 100 and 600 km across, with heights between about 0.3 and 5.0 km. The largest shields, however, are over 700 km in diameter and up to 5.5 km in height. For reference, Mauna Loa is ~120 km across at its base, and it has a total height of ~8 km (from the sea floor). Thus, the large Venusian shields are broader, but much flatter then the largest shield volcanoes on Earth. Indeed, the largest shield volcanoes on Venus cover nearly the same area as Olympus Mons, which has a basal diameter of ~800 km. (Note, Olympus Mons is still much bigger than the Venusian shields due to its immense height.)
These large shields all look much like shield volcanoes on Earth. They are mostly covered by long, radial lava flows. They all have very gentle slopes. And most also have some form of central vent or summit caldera. Thus, we think that these shields forme d from basalts, much like the shield volcanoes in Hawaii. The venusian shields, however, show a map pattern which is quite different from that seen on Earth (see reference map). Namely, the shields on Venus are widely scattered, and they show no linear vo lcano chains like those on Earth. This suggests that Venus does not have active plate tectonics, and also that most volcanism on Venus is related to mantle hotspots.

Very few large volcanoes are in the lowland plains or in the few highland regions. Rather, most of the large shields lie in the green to light yellow regions. Also note how the volcanoes are clustered into left and right sides of the map (dotted lines). T hese regions contain 2 to 4 times more volcanoes (for their size) than the planet as a whole. The reason for this grouping is not fully known, but it may reflect a cluster of closely spaced hotspots in a region of marked crustal failure. This region conta ins a large network of crisscrossing rifts and major fault zones.

Smaller Shields on Venus
Small shield volcanoes are very common on Venus. In fact, there may be over 100,000 of such shields less than 20 km across. These small shields usually occur in clusters, called shield fields. Over 550 shiled fields have been mapped (see Reference Map), and most of these are between 100 and 200 km across (nearly the same size as the state of Rhode Island). The shield fields are widely scattered on Venus, but they mostly occur in the lowland plains and lower upland regions. Very few are found at higher elevations or in the chaotic tessera.Many have also been partly buried by later lava plains. Thus, the shield fields seem to be fairly old, and they may have formed during the earliest stages of plains volcanism.
At slightly larger sizes, true shield volcanoes are rare. Of the 270 volcanoes between 20 and 100 km in size, only about 70 seem to be real shields. The rest are unusual features that may reflect different lava types. Most of the shields are also fairly small, less than 30-40 km across, and 25 of them show an unusual, petal-like set of bright lava flows (see images). Thus, there seem to be two distinct scales of volcanism on Venus: large and very small. These two scales may reflect differences in the nature or volume of mantle melting beneath the volcanoes.
Unusual Volcanoes on Venus
Most volcanoes on Venus are shields, but a few are not. These volcanoes all formed from very thick, viscous lavas, and they fall into three types. First are the so-called “pancake” domes. Second are a related class of “tick-like” structures. And third are a few volcanoes with thick, fan-shaped or banded flows. Since most basalt lavas are very fluid and fairly thin, these volcanoes do not seem to be made from basalt. Rather, they may mark quartz-rich or granitic lavas. Still, there are theories that gas-rich, “frothy” basalts could produce such viscous lavas. Also, few of these features lie in or near the Venusian highlands. Thus, these three volcano types may differ greatly from the volcanoes seen on Earth’s continents.
Most of the non-shield volcanoes on Venus are pancake domes, shown here in yellow. Like the shield volcanoes, these domes are widely scattered. They also often form small groups or clusters. Where most venusian shield fields have well over 20 small shields, however, few dome clusters have more than 5 or 6 domes. Further, the pancake domes are rarely found near shield volcanoes on Venus. Instead, many domes lie near corona structures in the lowland plains.
The ticks, shown here in red, show a pattern similar to the pancake domes. Like the pancake domes, most ticks occur in the lowland plains far from the largest shields on Venus. Many are found near the lowland coronas, and some are quite close to one or more pancake domes. The ticks also look something like the pancake domes. Thus, it is thought that many ticks are modified dome features.
Finally, fan-shaped and banded flows are quite rare. They are shown on this map in purple, and most are located in one part of the southern lowland plains. Neither their setting nor their distribution shows any clear trends, so little is known about these flows. Still, they look a lot like the most viscous lava flows on Earth. Thus, they may not have formed from basalt lavas. Instead, they may mark a rare set of granitic or rhyolitic lavas on Venus.

The spectacular planetary body Io is the innermost moon of Jupiter, and the most volcanically active body in the solar system. It is the only body outside the earth to exhibit active volcanism on a massive scale. The first images to arrive from the Voyager 1 spacecraft as it passed by Io in 1979 astounded the scientific community by revealing active volcanic plumes rising up to 300 km above Io’s surface. As additional images arrived from the Hubble Space Telescope and the Galileo spacecraft (1996-2000), we have come to realize that this remarkable body is dotted by hundreds of volcanic centers, about 70 of which are active. As shown in the image here, Io’s surface is extraordinarily colorful, with yellows, oranges, reds, and blacks reflecting compositionally diverse eruptions, from great outpourings of basaltic lava to massive deposits of sulfur. The lack of impact craters on Io’s surface is consistent with a young period of vigorous volcanism, which has buried older structures with lava, sulfur, and pyroclastic materials forming the dimpled, variegated crust we see today.

The first images obtained from Voyager 1 in 1979 showed a strange umbrella-shaped feature that appeared to be protruding from Io’s surface. Soon it was realized that the feature was an eruption plume, largely composed of SO2 gas, rising 260 km above the surface. Additional Voyager images revealed additional volcanic plumes. A total of 15 active plumes have now been identified from the images obtained through the Voyager and Galileo Missions, two of which are shown here.
Spacraft instruments have provided us with important data on the nature of these explosive eruptions. Voyager’s Infra-red Interferometer Spectrometer (IRIS) indicates that most of the active plumes correspond with anomalously hot areas on Io’s surface. Data from Galileo’s Near-Infrared Mapping Spectormeter (NIMS) appears to suggest that many of these high-temperature explosive eruptions are driven by SO2 gas emission, whereas others appear to be driven by volatilized sulfur emission. In contrast, explosive eruptions on earth are driven largely by H2O and CO2 gas emission.

Unlike the surface of most other planetary bodies, the extensive volcanic activity on Io can generate dramatic changes to it’s surface over very short timescales. The colored surface is largely covered by volcanic plains composed of overlapping lava flows and volcanic plume deposits dominated by pyroclastic particles composed largely of condensed SO2 frost, sulfur, and in places, subordinate silicate tephra.