Certain geological features and the formative processes they are derived from are so pervasive in central Oregon, and so fundamental to this section of the website, that they deserve special attention.  We’ll examine fire first and then ice, that is, the processes and products of volcanism and glaciation, because they are so numerous, easily accessed, and readily observed. 

Volcanoes are the most ubiquitous feature of the regional landscape.  There are many varieties of volcanoes; however, all share some processes and features in common as exhibited in Figure 1.  Each volcano is both a magma source, referred to as the vent, and a construct of the magma that is extruded.  Magma typically migrates upward through surrounding country rock from depth to accumulate within a shallow magma chamber.  Fractures formed as the magma pushes through overlying rock serve as one or more conduits, or feeder pipes that allow passage of the magma to the surface, where it erupts from the vent as successive lava flows, often of variable composition, and/or pyroclastic accumulations deposited on the landscape from tephra plumes, pyroclastic flows, lahars, and collapsing lava domes.  The word volcano implies a singular, generally conical feature built on the landscape.  However, volcanoes are invariably a composite structure, formed of interlayered lava and pyroclastic material, but also comprised of successive and varying domal accumulations, piled one upon another, that may include lava domes, cinder cones, composite (strato-) volcanoes and shield volcanoes.  Many volcanoes exhibit an age progression to the extruded materials of which they are made, generally erupting less viscous, more mafic or gaseous magma first and more viscous, more felsic and gaseous magma last, although some volcanoes display more cyclical processes related to subsequent refilling of the magma chamber from new magma sources followed by a new eruptive phase.  

Figure 1.  Common formative processes and physical characteristics of volcanoes.

Generation of Magma

To understand where volcanoes come from, one must first examine the origin of magma; the hot, molten rock that wells up from deep within the earth to extrude from fractures and conduits as lava flows and/or pyroclastic ejecta.  The source for magma is not the earth’s liquid outer core, a common misconception; instead, magma is generated at the relatively shallow depths of 100 to 300 km, through the partial melting of the earth’s crust and upper mantle.  It is most often formed by decompression-melting of the mafic asthenosphere associated with divergent plate boundaries or mantle plumes, or by partial-melting of water-rich felsic crust and/or asthenospheric material in association with subduction at convergent plate boundaries. 

The ingredients necessary for the production of magma involve the interplay between heat, pressure, intragranular fluids (present as gases within very hot rock or magma), and the composition of the material subject to melting.  Heating obviously brings solids closer to their melting points, the more heat, the more likely a solid will melt.  In general, higher pressures prevent melting because the constituent atoms of minerals in rocks are squeezed together and remain solids under high pressure.  Consequently, lowering pressure on already hot rock induces melting.  Intragranular fluids lower the melting point of solids, so the presence of fluids (gases), generally water, allows solid rock to melt at a lower temperature (or heat content) than it otherwise would.  Finally, there are two general trends to explore in relation to rock composition: rock that contains a relative abundance of silica (SiO2) and aluminum (aluminum oxide), generally called felsic rock, will melt at a lower temperature (heat content); while a rock containing a relative abundance of ferromagnesian (Fe, Mg, and Ca) ions, commonly referred to as mafic rock, will melt at higher temperatures (heat content).  Return to our earlier discussion of silicate minerals and Bowen’s Reaction Series to see why this is so (GEOLOGY BASICS – CRUSTAL COMPOSION AND DEFORMATION – Figure 7). 

Loosely speaking, the melting of continental crust generates felsic magma compositionally enriched in silica and aluminum, while melting of mantle rock (asthenosphere) and oceanic crust forms ferromagnesian-rich, mafic magma (Figure 2).  The earth’s crust naturally contains a higher water content (because of its proximity to the hydrosphere) than the mantle, accounting for higher water (and thus gas) content in felsic to intermediate magmas.  As previously discussed, the presence of water in rock causes it to melt at a lower temperature, so water-rich, generally felsic rocks will melt more quickly.  However, the water (now in the felsic magma) also prevents crystallization of silica chains within the magma, so water-rich magma will remain molten longer (at lower temperatures) (Figure 2).  Basically, “wet” magmas remain molten at lower temperatures than “dry” magmas.  The relatively high content of silica in continental crust also correlates with the lower melting temperatures of felsic to intermediate magmas.  As previously discussed, the abundance of silica-rich minerals in generally felsic rock allows melting at lower temperatures, but the resulting felsic-rich magmas would also crystallize sooner (at higher temperatures) (Figure 2), except for the presence of water which retards the crystallization process.  Mafic, ferromagnesium-rich mantle rock melts at greater depth and higher temperatures and pressures, not requiring as much “assistance” from silica and water in the melting process.  Hot, buoyant mantle rock naturally rises toward the surface, and under a lower pressure regime it melts, the resulting mafic magma then remains molten as it rises toward the surface (it remains hot, and the lower pressures reduce its ability to crystallize) (Figure 2).

Figure 2.  Factors that control the composition and viscosity of a magma; factors that in turn play a determining role in the style of volcanic eruption, eruptive products, and the nature of the volcano formed.

The viscosity of magma is its resistance to flow (viscous magma is thick, pasty, and chunky).  Figure 2 indicates that magma viscosity is related to its water content, silica content, and heat content; in general, more water, less silica, and greater heat equates to a less viscous magma.  Once formed, magma’s low density (lower than the surrounding country rock) naturally causes it to rise toward the surface (toward lower pressures).  If the viscosity remains low enough, magma can be extruded at the surface.  Mafic magma reaches the surface most readily because of its high heat content and low silica content.  Felsic to intermediate magmas are much less likely to reach the surface because of their high silica content and low heat content (they have a propensity for crystallization at depth within the crust).  Roughly speaking, about 80% of erupted magma is of mafic composition while intermediate to felsic magmas make up about 10% of erupted magmas each.  Water content plays a crucial role in determining the style of a volcanic eruption, with greater water content equating to more gaseous, explosive eruptions (Figure 2).  As for the eruption products, Figure 2 indicates that when mafic magma is erupted, it tends to form lava flows as flood basalts, shield volcanoes, or cinder cones (with higher gas content).  On the other hand, more felsic magmas tend to form greater pyroclastic materials, composite (strato-) volcanoes, and lava domes.

The composition of magma (and extruded lava) depends on three main factors: 1) the degree of partial melting of the crust or mantle; 2) the degree of assimilation of country rock and/or magma mixing; and 3) magmatic differentiation by fractional crystallization.  Several types of basaltic lavas that easily rise through the crust to erupt at the surface result from a range of 5 to 40% partial melting of mantle and oceanic crust (Figure 3); tholeiitic basalts (requiring 20 to 40% melting) are most commonly produced at subduction zones and mantle plumes.  Emplacement of basaltic magma chambers within continental crust often raises the temperature of the surrounding silica- and water-rich country rock enough to cause significant melting.  The country rock becomes assimilated into the basaltic magma to greater or lesser degree, contaminating it with felsic material to form andesitic magma (Figure 3).  Where tectonic movement and thermal convection in the magma chamber is low, two contrasting magma types develop; lavas erupting at the surface directly from the magma chamber are basaltic, whereas molten country rock erupts as rhyolitic lavas (Figure 3).  If substantial mixing of the magmas occurs, usually requiring significant plate movement and/or magmatic convection, erupted lavas are generally more intermediate in composition (ranging from andesitic to dacitic or rhyodacitic), the relative proportion of basaltic versus rhyolitic volcanism reflecting the initial amount of each magma type (Figure 3).

Figure 3.  The relationships between pooling of magmatic intrusions within continental crust, partial to complete melting of the surrounding country rock, and the generation of magmas with differing compositions.

As magma cools, crystallization ensues.  At specific temperature ranges, certain elemental constituents of the melt are preferentially removed to form specific minerals, a process modelled by Bowen’s Reaction Series (see the section of my website entitled GEOLOGY BASICS – CRUSTAL COMPOSION AND DEFORMATION – Figure 7).  In general, the melt becomes progressively depleted with respect to iron, magnesium, and calcium ions, and enriched in silica through a process known as fractional crystallization, while the magma is said to evolve by magmatic differentiation.  Within the magma chamber, slow cooling and crystallization produces large crystals that sink to the floor of the magma chamber or become plastered to its walls as cumulates.  Volcanic eruptions from a magma chamber can occur at any juncture in the magmatic differentiation process, feeding volcanoes with lavas and pyroclastic material of various compositions.  Inclusions of cumulates torn from the walls of the magma chamber during eruptions can be used to identify stages of fractional crystallization within the source magma chamber.

The magma’s viscosity (its fluidity or tendency to flow) determines how fast it rises toward the surface and whether or not it will see the light of day (Figure 2).  As previously indicated, felsic magmas have higher viscosities because of their lower heat content and enrichment with respect to silica; whereas mafic magmas have lower viscosities because of their greater heat content and lack of silica (they have a greater abundance of iron and magnesium ions).  Felsic magmas are generated by the partial melting of the more siliceous upper portion of water-saturated oceanic crust (more siliceous because of the thick sedimentary cover it carries) where it is subducted at convergent plate boundaries, and by assimilation-melting of siliceous, water-rich, continental crust into the magma derived from partial melting of mafic oceanic crust and asthenosphere as it rises toward the surface.  Figure 4 offers a possible plate tectonic model of the processes and features involved; here, a cross-sectional sketch of a volcanic island arc generated by oceanic-to-oceanic convergence shows the melting of oceanic crustal material to form magma.  This magma would tend to be more mafic because of the melting sources involved, but oceanic-to-continental convergence would form magma of more intermediate compositions.  Sources of felsic to intermediate magma impart a characteristic gaseous nature, and these gases play a significant role in reducing magma viscosity by preventing the crystallization of silicate minerals as the magma rises and cools.  Otherwise, felsic magma would never erupt at the surface, and as it is, it only rarely does so.  Mafic magmas can be generated by decompression-melting of highly mafic asthenosphere and assimilation-melting of mafic oceanic lithosphere and crust in association with oceanic-to-oceanic convergence (Figure 4), but they more commonly form at divergent plate boundaries (Figure 5) and oceanic mantle plumes (Figure 6).  The magma source at these latter locations is naturally low in water content, however, these magmas have a much easier time of it; greater heat and less silica allows it to readily reach the surface as volcanic eruptions (despite its lack of water).

Figure 4.  A cross-sectional model of a volcanic island arc formed by oceanic-to-oceanic collision and its associated arc-sediment deposition.   

Figure 5.  A mature divergent plate boundary exhibiting an ocean basin and well-developed spreading center or rift zone separating oceanic crust and lithosphere of progressively increasing age.

Figure 6.  A mantle plume model, showing the postulated source of magma and associated hot spot volcanism trailing away from the current hot spot in the direction of motion of the overlying plate.

The newly generated magma is lighter than the surrounding solid country rock and it rises buoyantly toward the surface as diapiric plutons (upside-down, tear-drop shaped bodies of melt) to collect in magma chambers at shallow depths beneath the surface (Figure 4).  Pressures build as the often gaseous magma accumulates in the magma chamber, and as it pushes upward, the overlying brittle rock is fractured to form conduits to the surface.  These conduits serve as the volcano’s plumbing system, feeding the volcanic vent on the surface where the magma is extruded as lava and pyroclastics (Figure 1).  Magma generated at divergent plate boundaries is chiefly the result of mantle convection in the asthenosphere which brings hot, plastic, mafic material closer to the surface where decompression induces melting (Figure 5).  This material is rich in ferromagnesian ions, but is low in silica, very hot, and water poor, and so has a low viscosity (Figure 2).  Magma continuously rises through fissures in the spreading plates; some is trapped to form new mafic lithosphere and deeper gabbroic oceanic crust.  Much of the magma rises to shallow depths to form oceanic crust of basaltic sheeted dikes, or pillow-basalts on the seafloor.  Flood-basalt volcanism forms stripes of lava parallel to the divergent boundary’s rift zone.  Over time, the sea floor accumulates a layer of siliceous sediment that naturally thickens toward continental margins, especially tectonically active convergent margins where sedimentation rates are high.  This package of sediment atop basaltic pillows and sheeted-dikes, overlying gabbro is referred to as an ophiolite sequence. 

Magma formed at oceanic hot spots in relation to rising mantle plumes is characterized by a similar process, although eruptions on the sea floor tend to form a long, narrow path of volcanic material as the overlying plate continues to move over the stationary mantle plume, burning a continuous line of hot spots through the plate (Figure 6).  Where volcanism is particularly voluminous, basaltic lavas stack up on the sea floor as seamounts and islands, such as the Emperor Seamount – Hawaiian Island chain in the Pacific Ocean basin.  Magma generated from continental hot spots is more silicic in composition, often resulting in intermediate to felsic volcanism.  The increased silica content is derived from assimilation-melting of continental lithosphere and crust as the original mafic magma burns upward through the overlying continental plate.  The Yellowstone area is thought to be the current position of a long-active mantle plume underlying the western North American plate.

The production of magma at convergent plate boundaries is predominantly related to subduction and partial melting of an oceanic plate under a younger, warmer, less dense oceanic plate (Figure 4), or under a buoyant continental plate.  Partial melting of the upper portion of the subducting plate is in response to its higher proportion of silica (from accumulated sediment on the ocean floor) and greater water content; some melting of asthenospheric material does occur adjacent to the subducting plate as steam is released into it from below.  During oceanic-to-oceanic plate collisions (Figure 4), this magma rises through the overlying oceanic plate and is little changed by assimilation-melting (the original mafic magma simply assimilates more mafic material on its way upward); volcanic eruptions on the sea floor stack up to form arccuate-shaped island chains called island arcs.  Volcanism is initially mafic in composition, but as time progresses and the volcanic arc ages and is subject to erosion (producing sediment that accumulates in the subduction zone), newer magmas become increasingly silicic (closer to intermediate in composition).  During oceanic-to-continental collisions, the generally mafic magma rises through felsic continental lithosphere and crust to build a volcanic arc on the continental margin.  Assimilation-melting of the overlying felsic continental plate produces intermediate volcanism on average, although lesser volumes of more mafic and more felsic volcanism does occur.  Volcanic arcs at convergent plate boundaries are more generally known as magmatic arcs because the magma is more viscous and much of it remains beneath the surface as intrusive bodies forming the roots of mountain ranges.  Magma constituting the volcanic arc of the Pacific Northwest’s Cascade Range has formed in this way. 

Volcanic Eruptions

A volcano is the result of volcanic eruptions, either where the tectonic plate on which the extruded magma accumulates is stable, allowing the material to stack up, or where the eruptions of material are particularly voluminous, outpacing the movement of the plate away from the volcanic source.  Volcanic eruptions occur as two general types (Figure 2): effusive (placid), when the magma erupted is of lower viscosity (and low gas content); and explosive (violent), when the magma erupted is of higher viscosity (or higher gas content) and consequently, it often has a significant component of water.  Although volcanoes can erupt a combination of lava flows and pyroclastic deposits (Figure 1), effusive eruptions tend to produce lava flows (of mafic to intermediate composition) that gently stream from the volcanic vent; whereas explosive eruptions generate copious amounts of pyroclastic materials (of intermediate to felsic composition) and may literally rip their vents apart.

The gas released from magma during volcanic eruptions is mostly water vapor, with a small component of carbon dioxide, and even smaller fractions of methane as well as sulfur and nitrogen compounds.  Deep beneath the ground, gases are under extreme pressure and are dissolved in the magma rather than occurring as bubbles.  As magma rises nearer the surface, internal pressure drops as it equilibrates with atmospheric pressure, the magma becomes supersaturated with respect to water vapor, and the gases are liberated from the magma to form bubbles in a process known as vesiculation (Figure 7).  Note that higher water (gas) content results in earlier vesiculation (at greater depth and pressure) (Figure 7a).  The sudden release of pressure that often begins with collapse of a lava dome plugging the central vent allows an explosion of gas-rich, viscous magma and rock fragments torn from the volcanic neck (Figure 7b).  Envision a bottle of soda; by the uncapping of a bottle of soda, the carbon dioxide dissolved in the liquid soda while under pressure is rapidly liberated to form bubbles which then rise to the surface once the cap is removed, often generating a frothy mix of liquid and gas that spills from the bottle’s mouth. 

Figure 7.  A diagram illustrating the process of vesiculation in a rising magma body; (A) shows the relationship between water content and magma depth, while (B) displays vesiculation within the volcanic system.

Geologists have classified volcanic eruptions into several styles based on the nature of gas release and the pyroclastics that are generated: strombolian, Plinian, and pelean.  Strombolian eruptions, named for those commonly witnessed on the volcanic island of Stromboli near Sicily, occur in association with effusive volcanism.  Gases dissolved in the magma vesiculate in the volcano’s magma chamber, coalescing into larger bubbles at the top of the magma chamber and then rise rapidly toward the vent through the volcano’s feeder pipe (Figure 7b).  Gases expand as they rise because of the difference between the gas bubble’s internal pressure and atmospheric pressure, whereupon the bubbles explode as they exit the vent, throwing fluid lava high into the air.  Intermittent release of huge bubbles leads to occasional impressive upward blasts of lava, whereas more continuous release of smaller bubbles forms lava fountains.  The substantially more violent nature of Plinian eruptions, as depicted in Figure 1, is linked to the rapid removal of a solidified “cap” of volcanic debris obstructing the rise of viscous, gas-rich, intermediate to felsic magma.  Plinian eruptions are named for the Roman scholar Pliny the Younger who reported on the eruption of Mount Vesuvius which buried Pompeii in AD 79.  High viscosity magma prevents vesiculation of the dissolved gases in the ascending magma, leading to an extreme pressure-differential between the gases and atmosphere.  Removing the offending obstruction, often involving oversteepening and mass wasting of a lava dome or flank of the volcanic cone, releases the pressure built up, causing extremely rapid vesiculation and transformation of the lava to highly vesicular pumice; then boom!  The violence of the volcanic explosion shatters the pumice into lapilli and ash which is thrown from the vent in a turbulent column of gases, forming a huge mushroom-shaped tephra plume, often reaching into the stratosphere.  The cloud of gases, condensed water droplets, and volcanic debris is spread downwind, generating a blanket of tephra that often covers vast areas (Figure 1).  Heavier material falling close to the vent, or ejected laterally outward, may roar down the volcano’s flanks as pyroclastic surges and flows.  Surges occur in conjunction with flows, the surge being the lighter and more highly mobile component of hot gases and ash, while the ground-hugging, more viscous flow is comprised of hot, frothy pumice and/or rock fragments. During the eruption of viscous magma, a domal mound of felsic lava may accumulate over the vent.  Rapid growth of the lava dome prevents degassing, and the resultant buildup of pressure causes explosive pelean eruptions that trigger multiple pyroclastic surges and flows down the volcano’s flanks, such as the surge and flow generated by the eruption of Mount Pelee on the Caribbean island of St. Martinique in 1902 for which these eruptions are named.  Plinian eruptions, generating immense volumes of pyroclastic deposits and their associated surges and flows, are common to composite (strato-) volcanoes of volcanic arcs generated at convergent plate boundaries such as the High Cascades of the Cascadia subduction zone. 

Another type of volcanic eruption, known as phreatic or phreatomagmatic eruption, is not related to the magma viscosity so much as it is related to where the eruption occurs.  Particularly violent volcanic eruptions occur when any type of magma comes in contact with groundwater or when lava is extruded into shallow water (depths of 700 meters or less).  Phreatic eruptions occur as violent steam explosions that release little or no lava, usually just ejecting pyroclastic material and the pulverized fragments of country rock from the vent area.  Water vaporizes when it comes in contact with lava, resulting in extremely explosive phreatomagmatic eruptions.  The lava is instantly frozen and fragmented into tiny pieces of volcanic glass and/or fine-grained igneous rock that accumulates as a ring of pyroclastic deposits around the vent.  Evidence of phreatic and phreatomagmatic eruptions is associated with the formation of maars, tuff rings, and tuff cones.

Volcanic Materials and Landforms

Effusive eruptions can occur in conjunction with any magma composition, but generally, because mafic magmas are less viscous, they are more likely to produce lava flows (Figure 2).  Basaltic lava forms two common types of lava flows: pahoehoe and aa, terms derived from the Polynesian language of the Hawaiian Islands.  Each is fundamentally the result of slight differences in silica and gas content, with pahoehoe lava containing less silica and gases.  Pahoehoe lavas are characteristically fluid because they are erupted at very high temperatures between 1100 to 1200 °C and are generally degassed by the time they flow onto the ground.  They may travel considerable distances (tens of kilometers) downslope from their source vent.  When they solidify, their surfaces exhibit smooth, ropy textures, and may initially appear shiny (Figure 8a).  Ropy texture originates when the upper surface of the lava cools to form a viscous plastic that continually refolds on itself as the lava flows forward.  Surfaces may break into platy blocks as warm, plastic lava continues to flow beneath.  Aa lavas are somewhat more viscous because they are erupted at slightly lower temperatures (1000 and 1100 °C) and they contain slightly more silica and gases; however, they travel similar distances.  As the flow moves forward, its front and top solidify and are broken into large chunks that are conveyed along by the hot, mushy interior.  These blocks roll off the steep flow front and are buried, steam-roller-like, by the advancing plastic interior of the lava.  When the flow becomes solidified, its surface (Figure 8b) and base form rubbly fragments between a few inches and a few feet in diameter, while its interior is usually denser and more contiguous.  Andesitic to rhyolitic lavas (Figure 2) are less fluid, erupting at temperatures ranging from 1000 to about 750 °C, and containing increasing concentrations of silica and gas.  They form progressively more jagged, blocky flow tops and bases as more and more of the melt has been crystallized prior to eruption.  A rather unique type of basaltic lava erupted in deep water occurs as pillows.  The pressure of seawater at depths of about 700 meters or more prevents degassing, instead, the lava is extruded gently and accumulates as thick flows composed of a multitude of rounded to oval blobs from about one to ten feet in diameter.  Although basaltic lava rarely erupts violently, it can when the magma is gaseous, and the lava generally forms coarse cinders, pyroclastics comprised of small lava bombs, rock fragments, and scoria that accumulate near the vent.  

Figure 8.  Basaltic lava flows generally consist of two types: pahoehoe lava with smooth, ropy surface textures (A); and aa lava with rough, blocky surface textures (left side of photograph) (B).

Explosive eruptions also occur in conjunction with any magma composition, but generally, because intermediate to felsic magmas are more viscous and gas-rich, they are more likely to produce lava domes and pyroclastic materials (Figure 2).  Pyroclastics form as lava bombs, blocks and rock fragments, scoria, lapilli, and volcanic ash, the finer fraction of which is often referred to as tephra.  The smaller the particles, the further from the source they can be projected (or transported by the wind).  Lava bombs form as fluid or plastic blobs of lava that are ejected into the air and fall to the ground, often molded into spindle-like, projectile, or cowpat-like shapes.  Blocks and smaller fragments are large to small pieces of previously solidified volcanic rock that are ripped from the vent during an eruption and projected through the air.  Scoria cinders form as fragments of gaseous, mushy lava that is solidified upon ejection from a volcano, preserving the plethora of vesicles.  Lapilli are small (5/64 to 2 ½ inches in diameter) glassy and pumiceous rock fragments ejected from a vent, whereas volcanic ash is any volcanic material consisting of smaller particles.  Figure 9a shows an outcrop of pyroclastic material consisting of ash and lapilli (Figure 9b) generated as a two-phased airfall and pyroclastic flow deposit. 

Figure 9.  Typical pyroclastic deposits (A) formed of volcanic ash and pumice lapilli (B); the outcrop (above the head of my former student) records a two-phased accumulation of airfall ash and overlying pyroclastic flow.

As we have seen, volcano formation is linked to magma viscosity, itself determined by magma temperature, composition, and gas content.  The magma’s viscosity controls not only the style of volcanic eruption and the nature of the volcanic materials that are generated by the eruption, but the way in which these materials accumulate on the landscape (Figure 2).  These observations have led geologists to classify volcanoes using a simple, three-fold differentiation based on size and volume: shield volcanoes, composite volcanoes, and cinder cones (Figure 10).  Other volcano forms, including flood basalts, lava domes, and maars, tuff rings and cones, do not fit precisely into this classification scheme.  Flood basalts are so fluid, they really don’t stack up at all; instead, they usually flow across a landscape, covering vast areas.  They are the main volcanic feature formed on the sea floor at divergent plate boundaries, but they can also occur terrestrially (such as the Columbia River Basalts of the Pacific Northwest).  As magma becomes progressively more felsic, the likelihood increases that andesitic to rhyolitic lava will stack up near the vent, or be ejected from the vent as pyroclastic material.  The flow potential of increasingly viscous magmas decreases markedly, instead, the lava oozes out of the vent to form a lava dome or dome complex.  The violent phreatomagmatic eruptions that occur when lava is extruded into shallow water produce pyroclastic material comprised of instantly frozen and fragmented tiny pieces of volcanic glass and/or fine-grained igneous rock that accumulate as deposits of hyaloclastite to form maar volcanoes or tuff rings and cones.

Figure 10.  A simple three-fold classification of volcanoes generally bases on the size and volume of the feature, includes shield volcanoes, composite (strato-) volcanoes, and cinder cones.

To begin a description of the three main volcanic forms, consider that multiple eruptions of fluid, basaltic pahoehoe and aa lava flows often accumulate as broad, gently-sloping, semi-circular volcanic edifices known as shield volcanoes (Figure 11); the name “shield” having been derived from the resemblance of this volcanic landform to an inverted warrior’s shield.  Much of the High Cascade’s platform of the Cascadia volcanic arc is constructed of thin sheets of basaltic lavas forming overlapping shield volcanoes that formed over a central graben during the Pleistocene.  Of the Cascades shield volcanoes, Belknap Crater near McKenzie Pass is a classic example (Figure 12); these shield volcanoes are typically somewhat smaller and steeper than is common because of the eruption of slightly more silicic continental volcanic-arc magmas.  Aa lava flows predominant on their slopes.

Figure 11.  A cross-sectional view of a shield volcano; broad, gently-sloping, semi-circular volcanic edifices comprised of multiple, thin, overlapping and interbedded basaltic pahoehoe and aa lava flows with only a minor component of pyroclastics.

Figure 12.  Belknap Crater in the central Oregon Cascade Range provides an excellent example of a continental volcanic arc shield volcano; dark, basaltic lavas can be seen to spread in all directions from a scoriaceous summit cone.

Mafic magma of slightly higher gas content will erupt more explosively, forming small, steep-sided volcanoes known as cinder cones (Figure 13).  Cinder cones typically host a truncated summit containing a funnel-shaped depression called a crater that marks the position of the vent and feeder system.  Cinder cones are formed of a monogenetic accumulation of basaltic lava bombs, blocks and rock fragments, and scoria usually generated during a singular phase of volcanic activity.  Many of these cones display a flanking breach where denser lava fractured the porous cinders of the cinder cone wall and poured out over the surrounding terrain as flows.  These small volcanoes are common, often occurring as “parasitic” dimples on the flanks of larger volcanic edifices where a smaller vent breached the surface, or as an alignment of cones marking the position of linear fault systems.  Figure 14 presents Yapoah Crater, a typical cinder cone formed along the central Oregon Cascade Crest near McKenzie Pass; lava flows breached the cone’s northwest flank to pour downslope to McKenzie Pass where they turned eastward to follow an old, glaciated stream drainage.

Figure 13.  Cinder cones are small, steep-sided volcanoes similar to shield volcanoes in composition, but the slightly higher gas content of erupted magma causes more pronounced stacking of an assortment of lava bombs, blocks and rock fragments, and scoria.

Figure 14.  Yapoah Crater provides an excellent example of a cinder cone; this one breached on its northwest flank by a lava flow that poured downslope past McKenzie Pass, made a turn to the east and continued flowing down a formerly glaciated stream valley and around the northern slope of Black Crater.

Eruptions dominated by magmas of intermediate composition, which typically produce both effusive and explosive volcanism, result in the formation of composite (strato-) volcanoes (Figure 15).  These volcanoes are volumetrically smaller than shield volcanoes, but larger than cinder cones, and are comprised of more viscous material that tends to accumulate nearer the vent and stack up as steep-sided mountains.  They are formed of a “composite” of alternating strata consisting of lava flows and pyroclastic deposits, hence the name (Figure 16).  The flanks of composite volcanoes often host glaciers and/or permanent snow fields, giving the misleading impression that they are taller/larger than shield volcanoes.  In fact, because many of these volcanoes have formed as the individual components of volcanic arcs perched on the overriding continental plate at a subduction zone, their bases begin at higher elevations and so they only seem to be of more significant size.  Although dominantly constructed of lava flows, these lofty peaks contain substantial amounts of pyroclastic-surge and -flow deposits.  Their steep flanks and the typically moist alpine climates of these mountains combine to produce considerable deposits associated with lahars and debris flow activity.  The Three Sisters (South, Middle, and North), forming a segment of the Cascadia volcanic arc in the central Oregon Cascades offer a classic example of continental volcanic arc composite volcanism (Figure 17).    

Figure 15.  Cross-sectional view of a composite (strato) volcano; note that they are smaller than shield volcanoes, with considerably steeper flanks, and are comprised of interlayered lava flows and pyroclastic material that is generally of intermediate composition.

Figure 16.  The glacially scoured northeast flank of South Sister Volcano in the central Oregon Cascades exposes a fine example of the interlayered lava flows (dark gray) and pyroclastic deposits (oxidized red layers) typical of a composite (strato-) volcano.

Figure 17.  The Cascadia volcanic arc in the central Oregon Cascades displays three exquisitely formed composite volcanoes strung in a row, South, Middle, and North Sister; a series of recently formed lava domes can also be seen to stretch southward from the southeast flank of the South Sister.

The summit areas of some composite volcanoes exhibit volcanic craters similar to their smaller cousins the cinder cones, but others feature caldera, large, often lake-filled depressions, formed by spectacularly violent eruptions that destroy entire mountain tops, such as that of Mount Saint Helens on May 18, 1980.  Calderas are the result of summit collapse caused by the rapid depletion of gaseous magma from the volcano’s subterranean magma chamber during a Plinian eruption and they are triggered by decompression associated with the abrupt removal of overlying material, often pushed or pulled aside by rising magma, ground swell, and mass wasting.  Caldera formation consists of several phases as outlined in Figure 18: (A) the magma chamber fills and begins supplying lava to the volcano, eruptions commence; (B) explosive eruptions begin to exhaust the reservoir of magma, evacuating the upper part of the magma chamber and initiating collapse of the volcano; (C) eruptions cease, magma chamber begins to cool, final collapse into the depleted magma chamber forms a caldera partially filled with volcanic debris, as water percolates into hot volcanic material, stream explosions are released; and (D) subsequent minor eruptions from cooling magma chamber build small volcanoes and accumulate volcanic deposits on the caldera floor as the depression gradually fills with water.  The wide, partially filled depression comprising the summit area of Newberry Volcano forms a spectacular, but relatively little-known caldera in the central Oregon Cascades east of the Cascade Crest (Figure 19).  Today, the caldera forms an immense, semicircular depression, partially filled by subsequent minor cinder cone and lava dome eruptions, and occupied by two lakes.

Figure 18.  Caldera collapse generally consists of four phases of volcanic activity: initially, the magma chamber fills and begins supplying lava to the volcano and eruptions begin (A); explosive eruptions begin emptying the magma chamber, initiating collapse of the volcano (B); volcanic eruptions cease as steam explosions from percolating groundwater continue while the magma chamber cools, contracts, and undergoes final collapse to form a caldera partially filled with volcanic debris (C); and subsequent minor eruptions from the cooling magma chamber build small volcanoes and accumulate volcanic deposits on the caldera floor as the depression gradually fills with water (D).

Figure 19.  The summit of Newberry Volcano suffered a caldera collapse many thousands of years ago during a spectacularly violent eruption, evidence of caldera formation can be observed as an immense, semicircular depression, partially filled by subsequent minor cinder cone and lava dome eruptions, and occupied by two lakes.

When viscous, intermediate to felsic magmas of relatively low volume are erupted, lava with a typically low gas content is squeezed slowly to the surface as an undifferentiated mass to accumulate over the vent as a lava dome or dome complex (Figure 20).  Lava domes are similar in size to cinder cones (Figure 10).  Domes often form in conjunction with the eruption of felsic pyroclastic material; the less dense, gaseous lava that generates pyroclastics is erupted first; the blast crater often serving as the basin in which the more viscous lavas later accumulate.  Lava domes may also develop as a second pulse of less gaseous, more silicic volcanic activity subsequent to caldera formation.  Several silicic domes and their short, stubby lava flows occur in the central Oregon Cascade Range, such as the lineament of silicic lava domes formed along the southeast flank of South Sister Volcano running southward to the Cascade Lakes Hwy that can be observed in Figure 17.

Figure 20.  A lava dome formed by the eruption of relatively small volumes of viscous, intermediate to felsic magmas with atypically low gas content.

The injection of magma upward through groundwater-saturated sediment and rock, or the extrusion of lava up through the floor of a lake creates particularly violent phreatomagmatic eruptions of predominantly pyroclastic material.  These eruptions may generate explosion craters in the original sediment and/or bedrock and the accumulation of tuff rings and cones over the vent area (Figure 21), collectively referred to as maar volcanoes.  Maars often contain small lakes or ponds where the water table intersects the floor of the crater.  Although these landforms are relatively small in volume and steep-sided, much like cinder cones and lava domes, magma composition and the outpouring of lavas play a minor role in the formation of these features; volcanic intrusion combined with ground- and/or lake-water is the key.  Figure 22 shows Table Rock Maar, a tuff ring formed at the northwest fringe of Christmas Valley in Oregon during the Pleistocene.  At the time of its formation, the area was saturated by a shallow water table or even filled with a shallow pluvial lake (Fort Rock Lake).  When a basaltic volcanic eruption came in content with groundwater or lake water, an explosive phreatomagmatic eruption of tuffaceous material ensued, forming the nearly circular tuff ring.  Subsequent erosion by a high pluvial lake level removed much of the outwardly dipping slopes of partially indurated tuff (picture Figure 21 with its outer slopes carved away by wave action) to form the volcanic feature we see today.  A careful examination of the right-hand side of the tuff ring reveals remnants of the outwardly dipping tuffaceous bedding.

Figure 21.  Maar volcanoes typically exhibit explosion craters and an accumulation of pyroclastic material preserved as a tuff ring that results from a violent phreatomagmatic eruption related to the forceful injection of magma upward through water-saturated sediment and rock.

Figure 22.  Table Rock Maar is a tuff ring in Christmas Valley, OR formed during the Pleistocene when a basaltic volcanic eruption encountered saturated sediment or a shallow pluvial lake.