This review presents some of the current knowledge of volcanoes in Hawai'i. It was originally written for a NASA-sponsored workshop about Hawaiian volcanism. We hope that with this review you can gain a better understanding of the processes and landforms that are associated with Hawaiian volcanoes. Many of these processes and features can also be found at other basaltic volcanoes on Earth. Additionally, Kilauea and Mauna Loa and have also become the primary volcanoes used by planetary geologists as analogs for volcanoes on Mars and Venus.
This review presents ideas derived by many volcanologists over the last few decades, the most prominent of whom are George Walker, Dave Clague, Jim Moore, and the late Gordon Macdonald. Scientists at the U.S. Geological Survey's Hawaiian Volcano Observatory, the University of Hawai'i, and elsewhere have made further important contributions to the study of Hawaiian volcanism, and their willingness to share their knowledge is gratefully acknowledged. Hypertext references refer to a bibliography that provides just a taste of the extensive literature available to those interested in studying Hawaiian volcanoes.
Basaltic shield volcanoes comprise a small percentage of Earth's volcanoes (~8%; Simkin et al. 1981). Hawaiian volcanoes are by far the best-studied examples of basalt shields. This means we have really only studied a small sample of a small percentage of Earth volcanoes. This review starts with large-scale structures and works to smaller and smaller features. Keep in mind how they all fit together to form a complex volcano.
Map of the main Hawaiian islands and nearby ocean floor. Note the Hawaiian trough and Hawaiian arch (dotted and dashed lines, respectively). Note also that some of the volcano rift zones extend well offshore.
Adapted from Macdonald et al. (1983); 1 fathom equals 6 feet or 1.8 meters.
The Hawaiian shield volcanoes are the largest volcanoes on earth (e.g. Peterson & Moore 1987) rising some 9 km above the ocean floor (see image), with volumes of 42,500 and 24,800 cubic kilometers (not counting subsidence) for Mauna Loa and Mauna Kea, respectively. Kilauea is a relatively small bump on the flank of Mauna Loa with a volume of 19,400 cubic kilometers. This can be contrasted to an average of ~100 cubic kilometers for strato volcanoes such as Mount Saint Helens (Wood & Keinle 1990). In the other direction, Olympus Mons on Mars rises 24 km above its base and has a volume of almost 4,000,000 cubic kilometers.
Hawaiian volcanoes reach these huge volumes in relatively short periods of time. Mauna Loa is thought to have begun forming on the sea floor some 500,000 years ago, although this is poorly constrained. For Mauna Loa, these numbers yield an average eruption rate of 0.085 cubic kilometers/year or 2.7 cubic meters/second. Interestingly, this is almost exactly the same eruption rate that is seen during low effusion-rate eruptions, and from observation of such eruptions this has been proposed to be the supply rate from the mantle (e.g. Swanson 1972, Dzurisin et al. 1984).
Plate tectonics provides a modern explanation for the presence of the Hawaiian volcanoes and their age progression from young in the southeast to old in the northwest. The lithosphere consists of the crust and uppermost mantle, both of which are rigid, and together can be divided into sections called plates. Beneath the lithosphere is the asthenosphere, a hot plastic layer on which the lithospheric plates can slide. Somewhere beneath the asthenosphere, and possibly as deep as the core-mantle boundary, is a hotspot, and the Hawaiian volcanoes are formed because of it. There are approximately 42 hotspots on earth (Duncan & Richards 1991).
Map of the Hawai'i-Emperor hotspot trace. Sinking and erosion causes the volcanoes to become smaller the farther they are from the hotspot. The last island where any volcanic rock sticks up above sea level is Gardner pinnacles. From Gardner to Kure the islands are atolls - coral reefs built on top of volcanoes that are now submerged. Beyond Kure even the coral has submerged and there are only seamounts (mountains on the ocean floor that don't make it to sea level). The seamounts at the bend have been dated at ~40 million years, and Meiji seamount is about 70 million years old. The ocean crust itself in the vicinity of Hawaii is about 90 million years old. (From Clague & Dalrymple 1987).
The exact nature of hotspots is poorly known, but it is known that they are sources of heat and/or magma that is supplied to the surface. Because they are stationary with respect to the moving lithosphere (as well as with respect to each other), linear chains of volcanoes form on the overlying plates and these volcanoes get older as you look in the direction of plate motion.
At Hawai'i, the Pacific plate is moving at ~9 cm/year towards the northwest. The Hawaiian volcanoes grow progressively older from the submarine volcano Lo'ihi and the island of Hawai'i at the southeast end of the chain through the main islands, through the leeward islands (mostly atolls formed on old submerged volcanoes), and beyond Kure atoll to the Emperor seamounts, the northernmost and oldest of which (Meiji Seamount) is being subducted under Kamchatka. The bend in the Emperor seamount chain reflects a change in plate motion about 40 million years ago.
This is a map of the "geoid" in the vicinity of Hawai'i. If the Earth were all water, or at least if you cut a bunch of canals through the continents so that water could flow freely between the oceans, the geoid would be the surface of that water. This imaginary body of water (the geoid) reacts to the force of gravity, which is not constant everywhere on Earth. Where the force of gravity is stronger more water will be attracted and flow into the area making the geoid height higher. Areas with lower gravity will be the places that give up this water and there the geoid height will be lower. In Hawai'i, however, there is no need to have an imaginary surface of water because we're surrounded by the Pacific Ocean.
Does this mean that the ocean is not "flat" in the vicinity of Hawai'i? Yes! The excess of dense mantle material provided by the upwelling hotspot means that the force of gravity is higher centered over the hotspot. The ocean surface is actually tilted upwards toward the big island of Hawai'i. The units are in meters, meaning that off the north coast of the big island the geoid is about 22 meters higher than the reference level. You couldn't actually see this tilting of the ocean surface with your eyes because it is very gradual and would be masked by waves. (Data courtesy H. Haack).
Uplift caused by the hotspot has bulged the Pacific plate upward over a broad region approximately 400 kilometers wide called the Hawaiian swell. This brings an excess of dense mantle material to near the earth's surface, and this extra mass actually results in gravity being slightly higher around the Hawaiian chain. At the same time, the loading of the volcanoes onto the heat-weakened swell has warped the center downward. This combination of uplift and subsidence (see image) has formed a broad m-shaped profile; the resulting structures are the Hawaiian trough adjacent to the islands, and the Hawaiian arch outboard of that. Recently the arch has been found to be the source of very high-volume lava flows (Holcomb et al., 1988; Clague et al. 1990).
Subsidence is greatest directly over the hotspot, for it is here that the lithosphere is the most thermally weakened and at the same time the greatest amount of lava is being loaded. At Hilo, tide gauge measurements during the past century have recorded sea level going up relative to the land at a rate of 4 mm per year. Worldwide, sea level has been rising ~1.5 mm/year during this time, so the extra 2.5 mm/year change in Hilo must be due to subsidence of the big island (Moore 1987). The southern part of the big island, more directly over the hotspot, must be sinking even faster.
In general, Hawaiian volcanoes during their most-active tholeiite shield stage can be characterized as having gentle slopes extending from the sea floor up to a summit caldera. The submarine slopes of Hawaiian volcanoes are steeper than the subaerial slopes but even they rarely exceed 14º (Mark & Moore 1987). In the alkalic stage the caldera is usually filled in and as noted earlier the slopes become steeper.
The greater proportion of explosive activity provides a good deal of ash causing Hawaiian volcanoes in their post-shield alkalic stage to resemble strato volcanoes to some degree. The most prominent large-scale structures are calderas and rift zones.
Located at the summits of both Kilauea and Mauna Loa are calderas. The areas of the calderas are 15 and 11 square kilometers, respectively, and their depths range from 140 to 170 m. The outlines of the calderas are distinctly non-circular, strong evidence that they result from the coalescence of more than one center of collapse. This is particularly evident in the shape of Moku'aweoweo, the caldera of Mauna Loa. Moku'aweoweo is strongly aligned in a direction that points towards the two Mauna Loa rift zones. The smaller collapse features that have coalesced to form Moku'aweoweo indicate the locations of zones of magma storage. This is evidence that the main magma chamber complex itself is aligned along these directions.
How do calderas actually form on basaltic volcanoes? Basaltic calderas are often shown as if they formed when a big piston-shaped cylinder of rock drops into the magma chamber, but this is not correct. A growing body of evidence is pointing to the idea that the calderas are sag structures that form when support is removed from below the summit. Walker (1988) showed that the cumulative subsidence of the Kilauea caldera has been funnel-shaped. (See below for the collapse and infilling history of Kilauea caldera since the arrival of Westerners. The bottom diagram shows the cumulative collapse to be funnel-shaped, not piston-shaped (adapted from Walker 1988). Additionally, a study of the eroded Ko'olau caldera showed that all the caldera-filling lavas have a centripetal dip (Walker 1988). The reason active Hawaiian calderas have steep walls is that surface rocks are too brittle to sag very well, and they fracture. The flat floors result from the re-surfacing by flows erupted within the calderas.
This resurfacing points to the fact that calderas are very dynamic features that are capable of collapsing and infilling many times. Indeed, ~25% of the surface of Mauna Loa is covered with lavas erupted ~600 years ago during a time when the summit caldera was full and overflowing (Lockwood & Lipman 1987)
This is not to say that distinct collapse events do not occur. Three large pyroclastic units have been mapped at Kilauea, and each can be correlated with a caldera collapse event (but not necessarily a caldera-forming event). The prevailing idea is that when magma drains suddenly from the summit region, support of the caldera floor is removed and collapse occurs. Groundwater is then able to flow inward towards the hot volcanic plumbing, and it flashes to steam, producing phreatic eruptions (as in 1924) or phreatomagmatic eruptions (as in 1790; e.g. McPhie et al. 1990).
Calderas are thus presumed to be have come and gone during the tholeiite stages of all the Hawaiian volcanoes, however, direct evidence is not present at all of them. Mauna Kea and Hualalai, for example, show no evidence of having had calderas in the past. As noted above, the "caldera" of Haleakala, although spectacular, is actually an erosional feature on the East Maui volcano, and evidence of a true volcanic caldera there has likewise not been found.
On some of the other older volcanoes the presence of old calderas manifests itself in sequences of thick flat-lying ponded flows, and areas of preferential erosion. This is particularly the case for East Moloka'i volcano. In these cases the centers of the volcanoes are now occupied by big holes. Caldera-filling lavas are usually easier to erode than flank lava flows because during the active period of the volcano's life the center of the volcano (the caldera) is the zone of greatest thermal and hydrothermal alteration and the rocks are quickly reduced to clays. An alternative idea is that the calderas happened to be in a state of collapse (rather than infilling) at the end of the tholeiite stage.
The image to the right shows the windward side of the old Ko'olau volcano (which comprises the east half of the island of O'ahu). The red line marks the crest of the pali (cliff). Numerous people have suggested that these cliffs are the scarp of the giant landslide that has removed most of the right-hand side of the Ko'olau volcano. If this were so, however, one would not expect there to be any high-standing intra-caldera rocks such as those inside the yellow marker. It seems more likely that the actual scarp is offshore from the present coastline; all the erosion between this scarp and the present pali has been accomplished by non-catastrophic processes since (but probably accentuated by) the giant landslide.
The situation seems to have been reversed on the old Kaua'i volcano. Here the ponded flows were so massive that they present more resistance to erosion compared to the flank lavas, and the outline of the old Kaua'i caldera is today marked by a high. relatively circular plateau. This also assumes that Kaua'i was only a single volcano, an idea recently challenged by Robin Holcomb of the USGS.
To the left is an air view (towards the west) of the caldera of Kaua'i volcano. Here the relatively horizontal plateau (outlined with the dashed light-blue line) consists of thick slightly more erosion-resistant lavas that ponded in the old caldera; it is topographically higher than the deeply eroded flank lavas. The yellow line marks the boundary between flank lavas and caldera lavas. Kalalau Valley and Waimea Canyon are eroding their way into the central plateau. Photo from Macdonald et al. 1983.
Radiating away from the summits of Hawaiian volcanoes are (usually two) linear rift zones. The rift zones conspicuously do not point towards adjacent volcanoes, but instead parallel the volcano-volcano boundaries. Rift zones mark preferred directions of sub-horizontal magma excursions from the magma chamber. Below is a map of the main Hawaiian islands showing rift zones in red lines and volcanic centers as red squares. Note that the rift zones tend to parallel the volcano boundaries, and avoid pointing at each other (from Fiske & Jackson 1972).
At the surface they are characterized by numerous vents, fissures, earth cracks, cinder cones, graben, pit craters, and the sources of lava flows. All of these are indications that magma preferentially intrudes into the rift zones and is also often stored there for periods of time up to a few years.
The vertical air photo on the left shows of a section of the NE rift zone of Mauna Loa. Even without the arrow it is pretty easy to figure out where the axis of the rift zone is. The red numbers give the dates of the flows (from Macdonald & Abbot 1970).
There has been much discussion about the formation and persistence of Hawaiian rift zones (e.g. Fiske & Jackson 1972; Deterich 1988). The general idea is that because Hawaiian volcanoes are close to one another relative to their size, a younger volcano is growing through the flank of an older one. The gravitational stress field caused by the pre-existing volcano tends to yield downslope-directed directions of least compressive stresses. Because dikes orient themselves so that their direction of widening is parallel to this least compressive stress, the dikes end up propagating parallel to the volcano-volcano boundary. Once a preferred direction of dike propagation is established, it is self-perpetuating as long as there is a mechanism for the flanks of a volcano to move outward to accommodate the repeated dike injections.
On the right is a schematic representation of Kilauea (purple) growing on the flank of Mauna Loa (green). Note how Kilauea has been affected by the shape (and hence the
stress orientation) of its huge neighbor, and has adopted the same rift zone orientation (from Fiske & Jackson 1972).
The most popular mechanism for this outward movement is sliding along the volcano-ocean floor interface which consists of easily-deformable sediments (e.g. Nakamura 1982). The focal mechanism for the 1975 M7.2 Kalapana earthquake indicated a slip plane that was nearly horizontal with a slight dip towards at a depth consistent with the base of the volcano (e.g. Lipman et al. 1985). Such an orientation would be expected due to the downward warping of the oceanic lithosphere under the load of the island.
Above is a schematic cross-section through Kilauea and part of Mauna Loa, viewed towards the East. This shows how the seaward flank of Kilauea (and part of Mauna Loa) is pushed southward (to the right) by the intrusion of dikes down the rift zone (away from you into the plane of the diagram). This huge bulk of volcano is probably sliding on ocean sediments that accumulated on the ocean floor during the 90 million or so years between the time that our particular part of Pacific Plate formed and when the Big Island of Hawai'i started to grow.
Rift zones probably become preferred directions of dike propagation due to stress orientations, and they evolve thermally to perpetuate themselves. This means that eruptions are rare elsewhere on the flanks of the shields. Except at the summit, the vents of Kilauea are found exclusively along the rift zones. On Mauna Loa, however, there is a class of vents called "radial vents " (Lockwood & Lipman 1987) that are found on the northern and western flanks. This is the sector on the obtuse side of the angle formed by the two rift zones, and circumferential tension caused by a bending moment set up by the rift zones and the westward push of neighboring may be leading to the formation of these vents (Walker 1990).
To the left is a map of the big island with Mauna Loa in orange. The short white lines are the "radial rifts" that do not fall into either of the rift zones (NERZ and SWRZ). Note that one of these radial rifts erupted through the flank of Mauna Kea, and that another erupted offshore (in 1877). Adapted from Lockwood & Lipman 1987.
Probably the most studied rift zone is the east rift of Kilauea. The northern flank of this rift is stable, probably because it abuts Mauna Loa. The south flank, however, is notably mobile. It has been shown to move seaward during both earthquakes and intrusive events. There is nothing in this direction to buttress the flank so the continued pressure caused by numerous dike intrusions produces this seaward displacement (Swanson et al. 1976; Lipman et al. 1985). This relative displacement between the non-mobile north flank and mobile south flank has caused a wide graben to form along the crest of the rift. Thus even though the rift axis is the locus of most eruptive activity it is in places topographically subdued. Some of the faults bounding this graben are visible near Napau crater.
Vertical air photo of Napau pit crater along the East rift zone of Kilauea. Napau has been almost filled by recent lavas (here making it look smooth relative to the surrounding forest). Note that vents, faults, fissures, and smaller pit craters are all aligned from the lower left (uprift) to upper right (downrift). The actual rift zone is wider than this photo (from Carr & Greeley 1980). Note also that differences in vegetation make flow margins traceable - the dotted white lines outline an old flow that appears to have had a source that is now engulfed into Napau crater.
Continued transport of magma down the rift zone results in the establishment of a thermally efficient conduit probably 2-3 km below the surface. Some evidence for this was provided by the first 10 km of propagation of the dike marking the onset of the Pu'u 'O'o/Kupa'ianaha eruption being aseismic (Klein et al. 1987). This indicates that there was a pre-existing conduit could be utilized by the migrating magma. This distance corresponds rather closely with the distribution of pit craters along the east rift Kilauea. Beyond the first 10 km, earthquakes marked the propagation of the dike.
Pit craters are not explosion craters or vents, but rather they are locations of localized collapse into a void. The above-mentioned conduit is the best candidate for such a void. A pit crater forms from the bottom up by stoping of a cavity.
The schematic diagram on the left shows the formation of a pit crater from the bottom up. This cross-section is cut perpendicular to the line of a rift zone, and "C" represents the main conduit at 2-3 km depth. Note that the crater is not a piston that has dropped and that the top lavas are the last to fall in (from Walker 1988). This process of a void working its way upward through solid rock (like a bubble) is called "stoping", and it commonly occurs in mines. The void and the eventual crater have a volume greater than the conduit, however, the conduit can continually carry material downrift.
Evidence for this is provided by a pit crater called the "Devil's Throat." When first noticed by Westerners, Devils Throat was a hole in the ground a few 10's of m across. A very brave man was lowered through this opening on a winch, and he soon found himself in a huge cavity, much wider than the hole he'd come through. It was evident that he was in a bell-shaped void and that the top layers of lava had not yet collapsed into it. Since then the last layers have fallen in, leaving Devil's Throat with the more typical cylindrical form of a pit crater.
On the right is a photo into Devil's Throat pit crater. Note that numerous pre-existing flows that are exposed in the walls, and the geology student for scale.
Eruptive fissures occasionally cut right across pit craters apparently without noticing the difference in topography. An eruptive fissure can extend from the floor of a pit crater, up the wall, and continue on beyond the rim.
Photo (taken in 1973 by Gordon Macdonald) showing a short curved fissure erupting across the floor of Makaopuhi pit crater. Lava is erupting (but not fountaining) from the near end whereas the far end is emitting only steam. The dotted line marks the base of the far wall of the crater. Makaopuhi is actually a double crater, and the lava cascades mark the boundary between the deeper near half and the shallower far half. The near half was previously twice as deep as the far half; eruptions prior to the one seen here almost made the two levels the same, and this eruption did eventually fill the deeper half to give the crater a single floor level.
In summary, vent, graben, and pit crater distributions yield insight to the preferred directions of magma travel within a Hawaiian volcano. These in turn can yield information about the stress directions within the edifice (e.g. McGuire & Pullen 1989; Rubin 1990). There have been some attempts to tie these stress directions to the stress field within the Pacific plate but it appears that the local stress field caused by neighboring volcanoes is much more important in determining the eventual direction of rift zone formation.
Vents, of course, are the locations from which lava flows and pyroclastic material are erupted. Their forms and orientations can be used to determine many characteristics of the eruption with which they were associated. There are two main endmembers in a spectrum of pyroclastic vents in Hawai'i, spatter vents and cinder cones. Their differences are due mostly to the gas content of the magma that is erupted. Additionally, there are satellitic shields formed during eruptions without fountaining and tuff cones formed during phreatomagmatic eruptions.
As a dike approaches the surface, it generates a zone of tension at the surface. This tension is usually manifested as a pair of cracks with the ground with the area in between often lower than the surrounding elevation (see below). The first phase of a Hawaiian eruption is usually characterized by breaking to the surface of a dike along one of the two fractures resulting in a line of erupting vents commonly called a "curtain of fire" (e.g. Macdonald 1972). After a few hours or few days most parts of the fissure stop erupting and activity is concentrated at one or more separate vents (e.g. Bruce & Huppert 1989). It is these vent locations that usually persist long enough (hours to weeks and sometimes years) to produce significant near-vent constructs. The change from long continuous erupting fissures to one or a few vents must be remembered when mapping eruptive fissures in remote sensing data and relating them to dike dimensions: The near-surface part of the dike is almost certainly longer than any line of near-vent constructs (see discussion in Munro 1992).
Fissures opened in the cinder-covered surface uprift from Pu'u 'O'o in July of 1985. Note that there are two parallel fractures about 50 m apart and forming a small graben. The next morning lava erupted out of the nearest fissure.
The ground surface here is covered by a ~2 m-thick layer of Pu'u 'O'o scoria, and this helped to accentuate the cracks - similar to the way that a small hole dug into sand at the beach will eventually look quite large as sand slumps into the hole.
In the case shown here the actual cracks in the rock under the scoria were only about 10-20 cm wide but so much scoria fell into them that they were wide enough to barely be jumpable.
Spatter refers to blobs of lava thrown a little ways into the air (by expanding gases) that is still molten when it lands.
Spatter ramparts and spatter cones are the vent structures formed by this type of activity. Spatter ramparts are elongate along the trace of an eruptive fissure whereas spatter cones occur as discrete mounds. They range between 1 and 5 m high, are steep-sided, and are composed of agglutinated (stuck-together) spatter. They are steep-sided because the hot spatter blobs are able to stick to each other when they land, and don't flow or roll away.
Small spatter cones forming near Pu'u 'O'o on the east rift zone of Kilauea in July of 1985. Note that the ability of molten spatter to stick together allows the spatter cones to be steep and even vertical. The pahoehoe toes in the foreground are picking up pieces of scoria that cover the ground in this area, and via a "reverse caterpillar motion" are placing these pieces (the dark spots) on top of themselves - stratigraphy is being reversed.
The fountaining associated with the formation of spatter ramparts is usually less than 10 m high.
A small spatter cone on Kilauea erupting in 1992. Note the fluid nature of the individual blobs making up the cone, and their ability to form a steep structure. The glowing orifice from which the spatter is erupting is ~30 cm across. The profile of Pu'u 'O'o can be seen in the background.
At the end of the eruption, lava often drains back into the fissure, forming prominent drainback features. Even if nobody actually witnessed a particular eruption, if you find spatter ramparts or cones associated with it, you can say that the fountaining that formed the cone or rampart was not very high.
Spatter vents from the 1974 eruption of Kilauea along the upper SW rift zone. Note that the last thing that happened was the drain-back of lava into the fissure (red arrows). The higher areas on the left and right are spatter ramparts and are 1-2 m high.
At the high-fountaining end of the spectrum are cinder cones. Cinder cones can be quite large in Hawai'i; those on the summit of Mauna Kea (formed during gas-rich alkalic-stage eruptions) are a few hundred meters high, whereas those on Mauna Loa and Kilauea usually range between 20 and 100 m high. Pu'u 'O'o on the E rift of Kilauea, which formed between 1983 and 1986 is unusual in that it reached a height of 255 m above the surrounding surface (Heliker & Wright 1991).
To the left is a photo of Pu'u 'O'o cinder cone, Kilauea, viewed toward the west. The prevailing right-to-left tradewind direction is obvious from the way that the plume is being blown. During the eruptions that formed Pu'u 'O'o, these same tradewinds built the cone much higher on the downwind side of the vent than the upwind side. Almost all the lava flows therefore came out of the upwind side (i.e. towards where the photo was taken).
As their name suggests, cinder cones consist of cinders, more properly called scoria. Scoria is very vesicular, low density basalt. Lava fountains are driven by expanding gas bubbles; the bubbles are trying to expand in all directions but the only way to relieve the pressure is up out the vent so fountains are usually directed relatively vertically. The Pu'u 'O'o fountains were at times up to 350 m high, and those during the early stages of the Mauna Ulu eruption were up to 500 m high. Because the pyroclasts are thrown so high, they cool before they land and don't stick together. Cinder cones are therefore composed of loose pyroclastics at the angle of repose (~33º).
Right is an image of fountaining at Pu'u 'O'o, Kilauea (July 1985). These particular fountains were ~200 m high, and were sending short fast flows in many directions.
In plan view, cinder cones tend to be roughly circular. They are usually formed later in an eruption when activity has localized to one or more discrete vents. If the precise location of the vent changes during an eruption, the cone loses its simple circular shape, and becomes more complex. Roadcuts through most cinder cones expose very complicated crosscutting relationships relating to the different locations of the fountain centers.
Photos of cinder cones on Mauna Kea (arrows), viewed from the summit of another cinder cone. The reddish color is common to cinder cones and occurs both during and soon after the associated eruption due to the combined efforts of moisture and oxidizing gases. The light blue line marks the Mauna Kea-Mauna Loa boundary. Note that one of the cones (yellow arrow) has been surrounded by (younger) Mauna Loa lavas.
Cinder cones can also be distinctly asymmetric if there was a persistent wind blowing during the eruption and/or they form at the heads of major lava flows. In this second instance they are horseshoe-shaped (see below), with the lava flow issuing out of the open end because during the eruption any pyroclasts that landed on the flow were rafted away.
A Mauna Kea cinder cone viewed from the air. A lava flow field (white outline) has issued from the base of the cone, giving the cone an asymmetric form. The flow spread almost all the way around the cone (white arrows). Magenta lines mark the rims of older cinder cones nearby.
Between the two extremes of spatter ramparts and cinder cones are all gradations. Some pyroclastic constructs consist of alternating layers of agglutinate and cinder, indicating that the vigor of the fountaining varied during the eruption. The early part of an eruption, often called the "curtain of fire", produces mostly spatter ramparts and spatter cones. As the activity becomes localized at one or more points along the fissure, this concentration of activity usually leads to higher fountaining. Cinder cones are built at these points, often at the same time that spatter ramparts are forming at the (less active) ends of the fissure. During the Mauna Loa eruption of 1984, there was a distinct gradation from vigorous fountaining at the main vents, progressing to lower and lower fountaining both up and downrift (Lockwood et al. 1987). At the farthest uprift end of the fissure, only gas was being emitted from a spatter cone that had been active earlier in the eruption.
Approximately 50% of Hawaiian eruptions have no pyroclastic activity associated with them at all. Instead, lava is quietly erupted onto the surface.
This lava flows away in all directions forming a miniature shield volcano. These vents are called "satellitic" or "parasitic" shields, and produce tube-fed flows. Satellitic shields have diameters of 1-2 km, and can be ~100 m high with slopes of only a few degrees.
On the right is an image of Mauna Iki satellitic shield on the SW rift zone of Kilauea. Note the gentle slopes (similar to Kilauea as a whole). The distance from left to right on the horizon is about 5 km
While a satellitic shield eruption is going on, a lava pond usually exists at the summit of the shield. Overflows from the pond build the shield.
Kupa'ianaha lava shield and pond in September 1986 (~2 months after it formed). Notice the shape of the pond, with a large near-circular part and an elongate extension to the left. The main lava tube was fed from the end of the extension. When this photo was taken the pond was also overflowing to the right. A break-out low on the far side of Kupa'ianaha was feeding another surface flow that in this photo had almost reached the contact with 'a'a flows from Pu'u 'O'o (3 km uprift; at left).
There have been 4 major satellitic shields formed on since the arrival of Westerners (Mauna Iki 1919-1920, Mauna Ulu 1969-1974, Kupa'ianaha 1986-1992, and the presently-active vent 1992-who knows?). Including these, 16 satellitic shields have been mapped on Kilauea (Holcomb 1987).
The major eruptive product of Hawaiian volcanoes is lava. Lava flows can form during fountaining eruptions or they can well out of the ground with little or no pyroclastic activity. There are two major types of basaltic lava flow, 'a'a and pahoehoe. These are Hawaiian words that have no meaning other than the type of lava, and they have been adopted by geologists working in other basaltic areas besides Hawai'i. They are different in almost every respect possible except for their chemistry.
Pahoehoe flow (left) flowing over older 'a'a, and in front of an advancing 'a'a flow (background). Note the continuous skin of the pahoehoe being crumpled at the flow front, and compare it to the broken clinker of the active 'a'a. Note also that the pahoehoe flow front is ~30 cm thick whereas the 'a'a flow front behind is almost 2 m high.
Table comparing 'a'a and pahoehoe lava flows.
All their myriad differences can be attributed to different eruption and consequent emplacement mechanisms.
'A'a flows are characterized most obviously by very rough top surfaces, dense interiors, and sometimes rough bottom surfaces. The loose fragments that make up the surface of flows are usually spinose (=clinkers) or more rarely smooth (=blocks). True block lavas are essentially absent from Hawaiian volcanoes (but common on some strato-volcanoes). Clinkers are formed as pasty lava is pulled apart by shearing and twisting during flow. The clinker layer is usually 1-2 m thick, but can be as thin as 10 cm. In general, the spininess of the clinkers is inversely proportional to both the thickness of the clinker layer and of the flow itself.
The dense interior is what actually flows as an 'a'a flow is emplaced, and it carries the clinkers along with it. Clinkers that fall off the front are buried by the advancing flow, generating a bottom clinker layer.
On the left are diagrams of a moving distal-type flow, showing the dense interior (that acts almost like a solid even though it is indeed flowing) with a blocky clinker carapace. The bottom clinker layer forms mostly from material that falls off the front of the flow and is run over (adapted from Macdonald 1972).
'A'a flows in Hawai'i range in thickness from 1-10 meters, and each consists of a few large flow units. The longest post-contact 'a'a flow in Hawai'i is the 1859 Mauna Loa flow (see below), at 51 km in length on land (plus a little more offshore).
Right: Map of the 1859 Mauna Loa "paired" lava flow. The 'a'a flow (orange) was active for 16 days, advanced at an average flow-front velocity of 133 meters/hour, and erupted at a volumetric flow rate of 208 cubic meters/sec. The pahoehoe part (blue) followed immediately after, was active for 285 days, advanced at an average flow-front velocity of 7 m/hour, and was erupted at a volumetric flow rate of 5 cubic meters/sec (from Rowland & Walker 1990).
High discharge-rate eruptions (usually accompanied by vigorous fountaining) lead to high volumetric flow rates, and these form 'a'a flows, which are emplaced at high flow-front velocities. The fastest recorded flow in Hawai'i was the 1950 Ho'okena 'a'a flow of Mauna Loa which advanced down a 5º slope through thick forest at approximately 10 km/hour.
Below: A plot of volumetric flow rate vs. average flow-front velocity. Note the distinct separation of 'a'a and pahoehoe at a volumetric flow rate of about 6 cubis meters per second and a flow-front velocity of about 10 m/hour (adapted from Rowland & Walker 1990).
Hawaiian 'a'a flows can be classified into two main types, proximal-type 'a'a and distal-type 'a'a (Rowland & Walker 1987).
Each can be found at any distance from the vent although the names imply otherwise. Proximal-type 'a'a flows tend to be 1-3 m thick, fast-moving, have thin layers of spiny clinker, little fine material mixed in with the clinker, and their interiors are often vesicular. When moving, the pasty interior of proximal-type 'a'a flows can be observed deforming and flowing, and can be penetrated by a thermocouple or viscometer.
To the right are examples of proximal-type 'a'a flows. In A, the incandescent core could be seen deforming as the flow advanced at a few meters/minute, and if you had been properly protected from the intense radiant heat you could have scooped out a pasty blob with a hammer. In B, note that the top clinker layer is only 10-20 cm thick and that the interior is relatively vesicular.
Distal-type 'a'a flows are often up to 10 m thick, slow moving, have thick top layers of dense smooth-surfaced blocks intermixed with much fine comminuted sand and dust, and their interiors are poorly-vesicular. When moving, the "flow" of the distal-type 'a'a interior is imperceptible, and it isn't possible to penetrate it even though it may be incandescent. There is a complete gradation between these two 'a'a types, and their differences are due to the fact that flows lose so much heat as they flow. This loss of heat increases the viscosity and yield strength of the lava, greatly changing the flow properties.
Distal-type 'a'a flows. In A, which shows the top half of a ~6 m-thick flow margin during the 1984 Mauna Loa eruption, the incandescent lava could not be perceived to flow and the main activity consisted of little particles of incandescent sand-sized particles sliding down the flow edge. Occasionally a large dense block would also tumble down off the top of the flow. Additionally, we could not penetrate the incandescent lava with a stick. In B, a pre-contact Kilauea flow, note the blockiness of the top layer and the massive, non-vesicular quality of the interior. Hammer (arrow) for scale.
The first and most obvious difference is that pahoehoe flows are smooth down to a scale of a few mm. Instead of consisting of only 1-2 large flow units, a pahoehoe flow consists of thousands on thousands of small flow units called toes. Each toe is usually <30 cm thick, 1-2 m long, and 30-50 cm wide.
Close-up photo of an active pahoehoe toe. This toe is about 30 cm wide at its widest. Note how it has erupted out of a crack in a previous toe and is flowing over yet another previous toe (with the ropy texture). Note also that with the sun shining on it, one side of the active toe doesn't look all that different from the surfaces of the older inactive toes; late afternoon and early morning (and night) are the best times for observing lava flows.
Pahoehoe flows are associated with low-effusion rate eruptions and are emplaced at low volumetric flow rates (2-5 cubic meters per second) and slow flow front velocities (1-10 m/hour) [See the A'a page for a velocity comparison chart]. Pahoehoe flows can be just as long as 'a'a flows. The longest post-contact flow was also erupted from Mauna Loa in 1859 (forming the second half of the "paired flow"; Rowland & Walker 1990), and is 47 km long. This strongly contradicts the notion that flow length is directly determined by effusion rate.
The low velocity of pahoehoe flows means that the skin that forms by air-cooling is not disrupted during flow and can maintain its smooth, unbroken, well-insulating surface. Thus the temperature and viscosity of lava do not change very much even tens of kilometers from the vent. The advancing front of a pahoehoe flow consists of hundreds or thousands of active toes. Each stops flowing after a few minutes and becomes inflated (with lava) as the eruption continues. Eventually the cooled skin fractures, often at the seam between two toes, and a new toe forms.
Cross-section of a pahoehoe flow exposed in a sea cliff. Some of the individual flow units have been outlined in white but you can see many others. Two particularly large ones were probably flowing as small lava tubes; the one labeled 'd' drained out at the end of the eruption and the one labeled 'f' solidified full. The dashed pink lines mark the top and bottom of the pahoehoe flow; above and below are 'a'a flows.
Lava has to somehow get from the vent to the front of the flow. There are two main types of "lava transport mechanisms", channels and tubes. It is important to remember that both of these form within the flow as it advances. Only occasionally have lava flows been observed to re-occupy pre-existing tubes and channels.
Channels (especially large ones) are associated with 'a'a flows. Pahoehoe flows sometimes have small channels. Channels form because the margins of active flows experience much more cooling and friction against the ground; they eventually stagnate but meanwhile the central part continues to flow (e.g. Hulme 1974; Sparks et al. 1976).
There are many forms of channels, depending on the degree of distinctness between the levees and the moving centers. The front of an 'a'a flow often moves as a single body of lava and has no channel. The margins of 'a'a flows have a distinctive cauliflower-like shape when viewed from above, which is caused by small surges of lava during flow. As soon as the margins begin to stagnate, however, the velocity difference between the margins and the flow center becomes large enough so that the lava begins to shear to make up the difference. At this point the channel may be evident only by two shear planes parallel to the flow direction. Eventually the levees stop moving altogether and the central part of the flow forms a well-defined, distinct channel. In such a case all the relative motion between the stagnant levees and flowing channel interior is taken up in thin zones at the very edges of the channel, and this makes the channel more efficient. In fact, velocities within channels have been recorded in the range of 60-80 km/hour (remember, this is not the speed of the flow front). Overflows from the channel usually coat the levees and build them up higher than the level of flowing lava. This channel evolution progresses downflow as the flow advances, with the well-defined channel lagging behind the actual flow front (by ~1 km on the 1984 Mauna Loa flow; Lipman & Banks 1987). This means that an 'a'a flow can be quite different depending on where you look at it.
Left: Photograph of a small lava channel a few hundred meters upflow from the 'a'a flow front. This channel was about 3 meters wide.
The lava flowing within a big channel often looks very smooth, and you might want to consider it pahoehoe. However, the lava flowing rapidly in a big channel is almost certainly destined to become 'a'a when it reaches the flow front. This fast-flowing smooth-surface is probably the reason why many people consider pahoehoe to be faster than 'a'a. The important thing to consider though, is what the lava will eventually become and it is pretty clear from observed eruptions that flows with fast-flowing channels feed rapidly-advancing 'a'a flow fronts. Most channels don't have solid crusts because the rapid flow as well as the velocity gradient across a channel means that it is impossible for any crust to remain continuous. A solid crust is a good insulator so because they don't have one, big channels lose a great deal of heat; this combined with the vigorous stirring during rapid flow helps to explain why pahoehoe-looking lava in a fast-moving channel invariably becomes 'a'a by the time it reaches the flow front.
Below: Diagram of a lava channel. The yellow planes are meant to illustrate how the central part of the channel moves faster than the sides. Torques (curved arrows) are set up by this velocity gradient, and continuously disrupt any surface crust that may form. For example, note that points A and B could rotate in different directions and pull the lava between them apart (adapted from Rowland & Walker 1990).
During pahoehoe eruptions, lava continues to flow through earlier-formed toes, particularly those near the center of the flow field. This continued flow causes the walls separating individual toes to become soft and eventually the toes start to coalesce to form a lava tube.
Diagrammatic cross-sections through a pahoehoe flow (same location, different times; grey color is solidified, reds indicate still-fluid lava). Note how the toes on the edges become solidified while continued flow through those in the middle eventually softens and erodes the toe margins until a large master tube forms.
Lava tubes also form when small channels roof over (e.g. Greeley 1971; Peterson & Swanson 1974). Lava tubes are thus self-forming within a flow field, and like channels they develop downflow during an eruption. Lava tubes are very efficient transporters of lava from the vent to the flow front, and Laszlo Keszthelyi has recently calculated that lava flowing in a tube loses only about 1Cº of heat/kilometer. At the flow front, the lava behaves much like a river delta, forming small distributary tubes that continue to branch until they consist of the same type of single flow-unit tubes (toes) that have been forming the flow the whole way downslope.
Above: Diagrammatic cross-sections through a pahoehoe flow (same time, different locations; grey is solid, reds are still fluid). A: Master tube formed within a pile of flow units, B: distributary tube system near flow front, continued flow in some of these will result in their merging to form a master tube, C: flow front consisting of numerous pahoehoe toes (adapted from Rowland & Walker 1990).
The amount of gas in lava at the flow front is somewhat lower than that at the vent because although lava tubes are good thermal insulators, they are not so well sealed that gas can be trapped. Gas escapes through cracks in the tube and also through skylights, places where the roof of a tube has fallen in to expose the flowing lava. Hot escaping fumes are strongly concentrated at skylights and they must be approached VERY carefully.
Right: An active skylight formed on the main lava tube from the Kupa'ianaha vent, Kilauea east rift zone, November 1991. The skylight is about 2 m across.
Whenever the flow front of a flow stagnates on sub-horizontal surfaces (such as on the coastal plain), it starts to grow by inflation as well as by spreading. This growth by inflation has only recently been recognized as an important factor in flow field growth (Walker 1991; Hon & Kauahikaua 1991; Walker 1992). The initial advancing flow front is usually <1 m thick. Subsequent inflation of the flow can increase its thickness up to ~10 m.
Places where inflation is particularly concentrated are called tumuli (singular: tumulus). Tumuli range in size from a few square meters in area and ~1 m high to a few thousand square meters in area and ~10 m high. Tumuli are often ubiquitous on flat-lying portions of a pahoehoe flow field. There are also places that inflate less than the surrounding area. These are called lava-rise pits (Walker 1991).
A large tumulus developed in the 1859 Mauna Loa flow near the coast. Note George Walker (arrow) in the central cleft for scale
Photograph of shelly pahoehoe. Note that the surface skin is only a few cm thick. The lava has not drained away from under the skin but rather within a rapidly quenched outer skin, an initially foam-like mixture of lava and gas bubbles separated into a gas-only top and lava-only bottom. It is easy to see why walking on shelly pahoehoe is not a pleasant pastime.
This bubble-rich flow front pahoehoe has recently been termed s-type pahoehoe, the "s" standing for "spongy" which aptly describes its appearance in cross-section (Walker 1989). While the lava is still molten, vesicles often coalesce just beneath the quenched skin and this forms a mostly gas (and therefore weak) layer under the top 1-2mm of surface glass. This weak surface layer of s-type pahoehoe spalls off, sometimes only within a few days, exposing the texture of the underlying vesicles.
Photograph of a s-type pahoehoe flow unit. The scale indicates 13 cm.
The vesicularity is also evident on the top surface of the lava, which consists of filamentous strands of vesicle walls that have been stretched as the skin deforms.
Close-up of the surface of an s-type pahoehoe toe showing the filamentous structure that forms from stretched an deformed vesicle walls.
The lava that is stored within a flow field before being erupted, on the other hand, is distinctly vesicle-poor. It has been termed p-type pahoehoe (Wilmuth & Walker 1993) because the bottoms of these flow units commonly contain pipe vesicles. Flow units of p-type pahoehoe have an appearance that makes them appear to be less viscous and their thickness is often <10 cm. This low-lying appearance is mainly due to the small number of gas bubbles within the p-type pahoehoe; if you stir it with a rock hammer it definitely "feels" more viscous than the more common gas-rich s-type pahoehoe
Left: S-type and p-type pahoehoe together. Note the thinness of the p-type lava flow units compared to the more billowy and/or ropy forms of the s-type. The blue color is distinctive of fresh p-type pahoehoe.
When first erupted, p-type pahoehoe has a distinct shiny blue color. During weathering it develops a patchy ochre coating.
Above: Mauna Iki p-type pahoehoe erupted in 1920, showing the yellowish secondary mineral that usually develops on this type of lava. Note that the secondary mineral does not form near fractures, perhaps because it dissolves in rainwater that accumulates there.
The top layer of this vesicle-poor lava strongly resembles obsidian and was originally thought to form when heavy rain falls while the lava is flowing (Wentworth & Macdonald 1953). Because you can see p-type pahoehoe flowing during any weather, this idea can easily be shown to be false. The most commonly held idea is that the paucity of vesicles is due to the prolonged storage within the flow field (up to a few days) prior to being forced onto the surface by an influx of new lava. During this storage gas bubbles are able to migrate upward and escape through fractures in the surface carapace. A competing idea (Hon & Kauahikaua 1991) is that p-type pahoehoe is actually richer in gas; the overburden of the solid flow surface has caused the gas to go back into solution.
All lava flows pass through a range of viscosities as they cool and solidify. Therefore you can't say that forms when the viscosity is low and forms when the viscosity is high. If that were the case you would never find any solidified pahoehoe because it has to go through high viscosities on its way to solidifying. The important factor is whether or not the lava is still moving at the time that it has developed the high viscosity (Peterson & Tilling 1980). This helps to explain why a pahoehoe flow sometimes becomes when it flows down a steep slope: If it has cooled to the point that it can no longer deform fluidly, it will instead break or be pulled apart into spiny clinkers. A sustained volumetric flow rate above ~10 cubic meters per second will cause almost any Hawaiian lava to form 'a'a so that if you find an 'a'a flow you can be pretty sure that it was emplaced at a high volumetric flow rate. A pahoehoe flow, on the other hand, tells you that the eruption took place at a low volumetric flow rate.
Volumetric flow rates can in turn usually be associated with discharge rates at the vent, and thus these two flow types can provide a great deal of information about eruptive conditions even if nobody was around to record the eruption.
Below: Map showing most of the large 'a'a and pahoehoe flows on Hawai'i. This information can be used to attempt to understand the plumbing of the rift zones as well as to characterize volcanic hazards (From Rowland & Walker 1990).
For example, long-duration tube-fed pahoehoe eruptions require a mechanically open pathway for magma to travel from the magma chamber. This is because the flow rate through the underground conduits is so low that magma pressure alone is unable to hold open the conduit. Such a mechanically strong conduit can last for long periods of time (years). During these eruptions, summit tilt (which is an indirect indication of magma chamber inflation or deflation) often shows no consistent inflation or deflation. The suggestion is that an open pathway has been established from the magma source (the mantle?) all the way to the vent.
A high-discharge rate eruption, on the other hand, (which produces 'a'a) does have sufficient pressure to erupt where no mechanically strong conduit exists, but only while the magma pressure remains high (Wilson & Head 1988; Hoffmann et al. 1990). This is the reason why high-discharge rate eruptions tend to be short-lived (often hours to days), and are associated with distinct summit deflation events (indicating emptying of the magma chamber, or at least a part of it). One relationship that is commonly found is for one or more short-duration high-discharge rate eruptions to precede a long-duration low discharge-rate eruption. Long-lasting low discharge rate eruptions will occur only if the repeated pumping of magma through the conduits during the earlier eruptions are able to establish an open pathway. Otherwise the eruption ends when the high discharge-rate portion ends (the conduits close off). Thus mapping the distributions of lava flow types can provide information about the nature of volcanic plumbing systems.
Mantle supply rates:
Because the discharge rates of almost all tube-fed pahoehoe eruptions on both and Mauna Loa seem to have been between 2 and 5 cubic meters per second, this has been proposed to be the supply rate to each volcano from the mantle (Swanson 1972; Dzurisin et al. 1984; Rowland & Walker 1990). You may recall from the beginning of this review that dividing the total volume of Mauna Loa by its estimated age yields essentially the same value. This would imply that all magma supplied from the mantle is erupted onto the surface; this is definitely not the case. When the volume of lava erupted onto the surface at Mauna Loa since the arrival of westerners (1778) is divided by the time since 1778, the rate is only 1 cubic meter per second, and a similar calculation for Kilauea yields a value of only 0.2 cubic meter/sec.
These values are 3 and 15 times smaller than the proposed supply rate of about 3 meters per second. These relationships point out the pitfalls of looking only at the surfaces of volcanoes for short periods of their lives, and suggest that the ratio of intruded:erupted magma is high at both Mauna Loa and Kilauea.
The diagrams below show the life stages of a Hawaiian volcano. Note that because of subsidence about half of the volume of the volcano is below the level of the ocean floor. Within the main SE Hawaiian chain, Lo'ihi is in stage 2. Kilauea is a young stage 4 and Mauna Loa is a mature stage 4. Hualalai and Mauna Kea are both in stage 5. Kohala is in stage 6. Moving to Maui, East Maui is in stage 7 whereas West Maui is between stages 7 and 8. East Moloka'i is also between stages 7 and 8. Lana'i (which skipped stage 5), Kaho'olawe, and West Moloka'i are in stage 6, perhaps never to go through stage 7. On O'ahu, Ko'olau is in (still in?) stage 7 (but also skipped stage 5), and Wai'anae is probably between stages 7 and 8. Kaua'i and Ni'ihau are between stages 7 and 8. As you can see, even though a lot of work has gone into figuring out this sequence, the volcanoes themselves have a good deal of variation and haven't perfectly "followed" the sequence. The diagram is adapted from Peterson & Moore (1987).
The Lo'ihi seamount off the southeast coast of Hawai'i was known from bathymetric surveys (see below), and thought to be a large slump. However, dredging in the 1970's recovered fresh lava samples, and the growing HVO seismic network began to record earthquake swarms centered on Lo'ihi. Submersible investigations have confirmed that Lo'ihi is actually the youngest Hawaiian volcano, with its summit some 975 m below sea level (~4000 m above the adjacent sea floor). Geochemical analyses of Lo'ihi samples show many of them to be alkalic basalt. Lo'ihi has a flat top that may be an almost-filled caldera, and as we'll see later this means it probably also has a magma chamber. Many of the pillow lavas observed on Lo'ihi have little or no sediment on them, a good indication of their recent formation.
The map on the right shows bathymetry of Lo'ihi volcano, off the SE coast of the big island of Hawai'i. The colors go from deep blue (~4500 m depth) to red (~1000 m depth). Note the relatively flat summit that is pocked with a couple of pit craters. This summit is probably a filled caldera. A seismic swarm in July of 1996 was accompanied by the formation of a new pit crater near the SW edge of the summit. This diagram was kindly provided by John Smith and Terri Dunnebier of the Hawai'i Mapping Research Group.
Below is a photograph of pillow lavas near the summit of Lo'ihi taken from the research submarine ALVIN, in 1987.
The prominent landmarks around the city of Honolulu (Diamond Head, Punchbowl crater, etc., collectively called the Honolulu volcanic series) started erupting some 1 million years after the last eruption of the tholeiite lavas (Ko'olau apparently never went through a post-shield alkalic stage).
Air photo to the right is of the Koko rift section of O'ahu (the E and of Ko'olau volcano). This prominent line of vents built up after erosion had carved deep valleys into this southeastern end of the Ko'olau volcano. The dashed pink lines mark the rims of phreatomagmatic craters, and the two pink dots mark non-phreatomagmatic "dry" vents. The green line is the crest of the Ko'olau mountains, which in this part of the volcano, roughly marks the top of the giant landslide scarp. Hanauma Bay is obvious with its white sand beach, Koko Crater is the tallest cone, and Manana Island marks the northeast end of the rift.
The other fine example is the East Maui volcano, commonly (but incorrectly) called Haleakala. The "crater" at the summit (which is properly called Haleakala) actually formed from the coalescence of two very large valleys that in the 400,000 years after the cessation of the post-shield alkalic stage, were able to eat out the heart of the volcano. After this long eruptive repose the rejuvenation stage filled these valleys with lava flows and cinder cones, providing the spectacular scenery found today.
The lava erupted in these rejuvenation-stage eruptions is highly alkalic and geochemically indicates that it came from a great depth (Chen & Frey 1983); a few of the deposits contain garnet-bearing xenoliths. These are indicative of a rapid and violent journey from the zone of magma generation. The volume contribution from these rejuvenation-stage volcanics is <<1% of the total for a Hawaiian volcano, and they form monogenetic fields - each vent only erupting once. Like monogenetic fields elsewhere in the world the overall eruption rate during this stage is very small (Walker 1990).
It is not known why this stage of volcanism occurs. There have been numerous explanations put forth, the most popular one today suggests that the lithosphere rebounds upward after having been depressed while directly over the hotspot. This rebounding is because the lithosphere is no longer being thermally weakened and because the overlying volcanoes are eroding. Depressurization due to this uplift would then lead to melting and magma generation. It is not clear that the lithosphere does indeed rebound in that way, however. Perhaps batches of magma attempt to make it to the surface all over under the Pacific plate, and only where the plate has been fractured and weakened by hotspot volcano formation are they able to make it to the surface. On the Ko'olau volcano, many of the rejuvenation stage vents are found to lie along rifts that are perpendicular to the trend of the old Ko'olau volcanic structure. Many of the Honolulu volcanic series vents happened to erupt into shallow seawater, and the eruptions were phreatomagmatic.
It is also during this stage that a magma chamber fully develops to serve as a way-station for ascending magma. A magma chamber migrates upward as the volcano grows, and the magma chambers of Mauna Loa and Kilauea are both 2-3 km below the summits. Although usually depicted as giant balloons, magma chambers are most probably a complex of smaller interconnected voids (more like a magma chamber complex). This idea has been confirmed at by geodetic measurements that show the center of deformation moving around during periods of inflation and deflation (Fiske & Kinoshita 1969).
The end results of all these processes are that lavas erupted during this main shield stage of volcanic life are: 1) hot and fluid because they have an efficient pre-heated plumbing system to get them to the surface; 2) have already lost some of their gas when they eventually erupt because it escaped while the magma was resting in the magma chamber; 3) possibly olivine-rich if the eruption taps the lower part of the magma chamber; and 4) unlikely to include xenoliths because the xenoliths sank to the bottom of the magma chamber.
Another important consequence of the development of a magma chamber is that it can lead to the formation of a caldera. Calderas result from collapse and/or subsidence into the magma chamber, thus a caldera is a sign that an active magma chamber is or once was present, and this in turn implies a high supply to the volcano. Calderas are very dynamic features, however, and at times they can be completely filled in, only to re-form again later.
The lower magma production rate during the post-shield alkalic stage decreases the thermal efficiency of the plumbing system and eventually the main magma chamber solidifies due to the lack of replenishment. This lack of a shallow magma chamber has a great effect on the nature of eruptions during the post-shield alkalic stage. Without the shallow magma chamber to rest in, only large batches of magma are able to make it to the surface, and they have to make the trip from the source region to the surface quickly to avoid solidifying along the way. This means that their xenoliths and large crystals don't have a chance to settle out, nor does gas have a chance to escape.
The consequences of all this are that post-shield alkalic stage eruptions usually consist of large volumes of cooler, gas-rich, xenolith-rich lava, but they are infrequent. The greater gas content means higher fountains and consequently larger cinder cones such as on Mauna Kea (right). Alkalic lava theoretically has a lower viscosity due to its lower silica contents. However, because during this stage the lavas also tend to be cooler, the viscosity increase due to this lower temperature usually outweighs the viscosity decrease due to the lower silica. Post-shield alkalic activity tends to be more concentrated at the summits than during the main shield activity.
This combination of lots of big cinder cones and lots of thicker viscous flows all concentrated near the summits causes Hawaiian volcanoes in this alkalic stage to be noticeably steeper and bumpier than in the tholeiite stage (see below). Mauna Kea and Hualalai are in this stage of development. Hualalai last erupted in 1800 and 1801, and Mauna Kea about 3600 years ago.
The air photo below, aiming towards the SW, shows Mauna Loa, Mauna Kea, and Hualalai. Note that Mauna Loa (ML), which is in its tholeiite shield stage, has a more gradual and smooth profile compared to Mauna Kea (MK) and Hualalai (H), both of which are in their post-shield alkalic stages.
Eventually, the volcano moves so far off the hotspot that magma is unable to be supplied; erosion takes over as the dominant geological process. This erosion is both gradual and catastrophic. Hawaiian lavas are very permeable. Many are vesicular, and lava tubes, clinker layers, and flow boundaries all provide easy pathways for percolating water. For this reason, even in many of the wettest areas of Mauna Loa and Kilauea, erosion is minimal. During the post-shield alkalic stage, however, the greater explosivity of the eruptions deposits many ash and cinder layers. These pyroclastic layers are much less permeable, and they allow streams to form readily.
For instance, the image to the left is Kohala volcano viewed from the north. Note that the original volcano surface (outlined by the dashed white line) is well-preserved in the highlands. Stream capture has allowed Waipi'o and Waimanu valleys to become huge at the expense of their neighbors, and Waipi'o has captured the headwaters of the left-hand branch of Waimanu. The green arrow indicates a "wind gap" between the two large valleys. The dotted yellow line roughly corresponds to the extent of (younger) Mauna Kea Lavas, and the town of Waimea is in the middle distance. Kohala last erupted ~60,000 years ago.
Another consequence of the lower eruption rate is that the coastal plain that formed from lava deltas during the tholeiite stage is not re-surfaced fast enough to avoid being submerged below sea level (the volcano continues to sink even though the eruption rate has decreased). The submerged (once coastal) plains can be identified offshore by bathymetric surveys, and if submerged coral reefs on these can be dated it is possible to determine roughly how long ago that particular volcano finished its tholeiite shield stage (Moore 1987). For example, corals on the shelf off Mauna Kea are about 500,000 years old. The shelf is at a depth of about 1000 meters, yielding a subsidence rate of 2 mm/year.
Bathymetry map (right) showing a set of slope breaks (gradual to steep heading offshore), corresponding to the now-submerged tholeiite-stage coastal shelves of Mauna Kea (magenta arrow), Kohala (yellow arrows), E Maui, W Maui, Kaho'olawe, Lana'i (cyan arrows), E Moloka'i, and W Moloka'i (green arrows). The diagram is adapted from Moore (1987), and the contour interval is 100 fathoms (= 600 feet or 183 meters).
There is another form of volcano degradation that has only recently been recognized as a significant process on Hawaiian volcanoes. Bathymetric surveys in the early 1960's showed two tongues of rough under-sea terrain extending ~180 km offshore from the eastern parts of O'ahu and Moloka'i (Moore 1964).
To the left is a map of the first two mapped giant landslides (north of O'ahu and Moloka'i). These landslides are the windward halves of the Ko'olau and E Moloka'i volcanoes. Note that they traveled across the Hawaiian Deep and climbed up the inner face of the Hawaiian Arch. Counting contours indicates that the slide from O'ahu climbed 2400 feet (~ 750 meters), a good indication that it was moving quickly rather than slowly creeping along (adapted from Macdonald et al. 1983).
Considerable debate raged over whether or not these were landslides. Using GLORIA sidescan sonar, 17 of these giant deposits have since been identified off the 8 main islands, and the general opinion is that they are indeed landslides. Many of them flowed out and sloshed up the island-facing slope of the Hawaiian trough, and must therefore have been moving quickly. Of course, if one of these events were to take place today the results would be devastating.
A deposit of beach cobbles has been identified on Lana'i extending up to an elevation of ~100 m, and it has been attributed to the tsunami generated by the most recent of these giant landslides. (Moore & Moore 1988).
On the right is an image of the deposit of coral and basalt rubble attributed to a large tsunami that washed up the southern flanks of Lana'i approximately 100,000 to 150,000 years ago due to the most recent giant avalanche (probably off the W flank of Mauna Loa). The geologist's right foot is on the contact between underlying lava of Lana'i volcano and the tsunami deposit.
On land, the headwalls of these giant landslides are indicated by steep ocean-facing scarps and slopes. The north coasts of E Moloka'i, and Kohala are prime examples. Below left is an air photo of the north coast of Moloka'i. The 600-1000 m high cliffs are the headwalls of a giant landslide that carried away half of the E Moloka'i volcano. The dotted red line outlines the original shield surface, into which the large valleys of Wailau, Pelekunu, Waikolu, and Wai'ale'ia Valley have cut. Oloku'i is a high point between Wailau and Pelekunu; because of its isolation by steep cliffs, Oloku'i boasts a fine complement of rare native plants.
The western slope of Mauna Loa is very steep, and has been the source of many giant landslides (Normark et al. 1987; Lipman et al. 1988), however, because Mauna Loa is still active, any scarps that may have formed have been mostly mantled by lava flows. The Hilina fault system appears to be a different type of mass-wasting structure. Here, large fault blocks tend to move in small increments rather than in huge catastrophic slides. During the large M7.2 1975 Kalapana earthquake, these blocks subsided up to 8 meters, and a small tsunami was generated (Tilling et al. 1976; Lipman et al. 1985). The Hilina fault scarps are continually resurfaced by lava flows which spread out when they reach the coastal plain.
The giant landslides were at first thought to be problematic because unlike steep strato volcanoes (where landsliding is expected), Hawaiian shields have very gradual slopes and very little ash. When further consideration is made of the structure of the Hawaiian shields, however, the mechanism of catastrophic failure becomes evident. When lava flows into, or is erupted in, shallow seawater, explosions occur. This happens as the volcano first grows through sea level and also when an already-subaerial volcano sends lava flows to the coast. These explosions fragment the lava into sand-sized particles consisting mostly of glass. Additionally, flows break up when tumbling down offshore slopes or being beaten by ocean waves. Lava flows extend the island offshore on top of all this loose material. The result of these processes is that much of the submarine component of all the Hawaiian volcanoes consists of very weak and unconsolidated easily-weathered material. Lava flows on land are mechanically strong but because they are underlain by these deposits of junk, the volcanoes as a whole are weak.
A series of diagrams to the right show how Hawaiian volcanoes have an inherent weakness that can lead to giant landslides (from ideas presented by Dave Clague, USGS). In A, a young volcano erupts pillow lavas on the seafloor; explosive activity is prevented by deep water pressure. In B, the volcano nears the surface; the water pressure no longer prevents explosive mixing of hot lava and seawater, and phreatomagmatic explosions produce a layer of hyaloclastite (yellow). In C, the volcano has grown above sea level so that eruptions no longer encounter seawater; they are not explosive, however, lava flowing into the ocean breaks up and occasionally produces littoral explosions, both of which also generate hyaloclastite. In D, the volcano is continuing to build subaerially; the layers of hyaloclastite are an inherent weakness that may promote giant landslides.
Following is a list of publications referenced in the Hawaiin Volcanism Section of VW.
Bruce PM & Huppert HE (1989). Thermal control of basaltic fissure eruptions. Nature 342: pp. 665-667.
Carr MH & Greeley R (1980). Volcanic Features of Hawaii. NASA pub. SP-403. 211 pp.
Chen C-Y & Frey FA (1983). Origin of Hawaiian tholeiite and alkalic basalt. nature 302: pp. 785-789.
Clague DA (1987). Hawaiian xenolith populations, magma supply rates, and development of magma chambers. Bull. Volcanol. 49: pp. 577-587.
Clague DA, & Dalrymple GB (1987). The Hawaiian-Emperor volcanic chain. U.S. Geol. Surv. Prof. Pap. 1350: pp. 5-54.
Clague DA, Holcomb RT, Sinton JM, Detrick RS, & Torresan ME (1990). Pliocene and Pleistocene alkalic flood basalts on the seafloor north of the Hawaiian islands. Earth Planet. Sci. Lett. 98: pp. 175-191.
Decker RW (1987). Dynamics of Hawaiian volcanoes: an overview. U.S. Geol. Surv. Prof. Pap. 1350:997-1018.
Deterich JH (1988). Growth and persistence of Hawaiian rift zones. J. Geophys. Res. 93. pp. 4258-4270.
Duncan RA & Richards MA (1991). Hotspots, mantle plumes, flood basalts, and true polar wander. Rev. Geophys. 29: pp. 31-50.
Dzurisin D, Koyanagi TY, & English TT (1984). Magma supply and storage at Kilauea Volcano, Hawaii, 1956-1983. J. Volcanol. Geotherm. Res. 21: pp. 177-206.
Fiske RS & Jackson ED (1972). Orientation and growth of Hawaiian volcanic rifts: the effect of regional structure and gravitational stresses. Proc. Roy. Soc. London a-329: pp. 299-326.
Fiske RS & Kinoshita WT (1969). Inflation of Kilauea Volcano prior to its 1967-1968 eruption. Science 165: pp. 341-349.
Greeley R (1971). Observations of actively forming lava tubes and associated structures, Hawaii. Modern Geology 2: pp. 207-223.
Heliker CC & Wright TL (1991). The Puu Oo-Kupaianaha eruption of Kilauea. EOS 72: pp. 521-530.
Hoffmann JP, Ulrich GE, & Garcia MO (1990). Horizontal ground deformation patterns and magma storage during the Puu Oo eruption of Kilauea volcano, Hawaii: episodes 22-42. Bull. Volcanol. 52: pp. 522-531.
Holcomb RT (1987). Eruptive history and long-term behavior of Kilauea Volcano. U.S. Geol. Surv. Prof. Pap. 1350: pp. 261-350.
Holcomb RT, Holmes M, Denlinger RP, Searle RC & Normark WR (1988). Submarine Hawaiian north arch volcanic field. EOS 69: p. 1445.
Hon K & Kauahikaua J (1991). The importance of inflation in formation of pahoehoe sheet flows. EOS 72 #44: p. 557.
Hulme G (1974). The interpretation of lava flow morphology. Geophys. J. Roy. Astro. Soc. 39: pp. 361.
Klein FW, Koyanagi RY, Nakata JS, & Tanigawa WR (1987). The seismicity of Kilauea's magma system. U.S. Geol. Surv. Prof. Pap. 1350: pp. 1019-1185.
Lipman PW, Lockwood JP, Okamura RT, Swanson DA, & Yamashita KM (1985). Ground deformation associated with the 1975 magnitude-7.2 earthquake and resulting changes in activity of Kilauea Volcano, Hawaii. U.S. Geol. Surv. Prof. Pap. 1276: 45 pp.
Lipman PW & Banks NG (1987). Aa flow dynamics, Mauna Loa 1984. U.S. Geol. Surv. Prof. Pap. 1350: pp. 1527-1567.
Lipman PW, Normark WR, Moore JG, Wilson JB, & Gutmacher CE (1988). The giant submarine Alika debris slide, Mauna Loa, Hawaii. J. Geophys. Res. 93: pp. 4279-4299.
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Macdonald GA (1972). Volcanoes. Prentice-Hall Inc., Englewood Cliffs. 510 pp.
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Macdonald GA, Abbot AT, & Peterson FL (1983). Volcanoes in the Sea the Geology of Hawaii. Univ. Hawaii Press, Honolulu. 517 pp.
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McPhie J, Walker GPL, & Christiansen RL (1990). Phreatomagmatic and phreatic fall and surge deposits from explosions at Kilauea volcano, Hawaii, 1790 AD: Keanakakoi ash member. Bull. Volcanol. 52: pp. 334-354.
McGuire WJ & Pullen AD (1989). Location and orientation of eruptive fissures and feeder-dykes at Mount Etna: influence of gravitational and regional tectonic stress regimes. J. Volcanol. Geotherm. Res. 38: pp. 325-344.
Moore JG (1964). Giant submarine landslides on the Hawaiian ridge. U.S. Geol. Surv. Prof. Pap. 501-D: pp. 95-98.
Moore JG (1987). Subsidence of the Hawaiian ridge. U.S. Geol. Surv. Prof. Pap. 1350: pp. 85-100.
Moore GW & Moore JG (1988). Large-scale bedforms in boulder gravel produced by giant waves in Hawaii. Geol. Soc. Am. Special Pap. 299: pp. 101-109.
Moore JG, Clague DA, Holcomb RT, Lipman PW, Normark WR, & Torresan ME (1989). Prodigious submarine landslides on the Hawaiian ridge. J. Geophys. Res. 94: pp. 17,465-17,484.
Munro DC (1992). Applications of remotely-sensed data to studies of volcanism in the Galapagos Islands. PhD Dissertation, University of Hawaii.
Nakamura K (1982). Why do long rift zones develop better in Hawaiian volcanoes--a possible role of thick oceanic sediments. Bull. Volcanol. Soc. Japan 25: pp. 255-267.
Normark WR, Lipman PW, Wilson JB, Jacobs CL, Johnson DB, & Gutmacher CE (1987). preliminary cruise report: Hawaii GLORIA legs 3 and 4, F3-88-HW and F4-88-HW. U.S. Geol. Surv. Open-File Rept. 87-298: 34 pp.
Peterson DW & Swanson DA (1974). Observed formation of lava tubes. Studies in Speleology 2: pp. 209-222.
Peterson DW & Moore RB (1987). Geologic history and evolution of geologic concepts, island of Hawaii. U.S. Geol. Surv. Prof. Pap. 1350: pp. 149-189.
Peterson DW & Tilling RI (1980). Transition of basaltic lava from pahoehoe to a'a, Kilauea Volcano, Hawaii: field observations and key factors. J. Volcanol. Geotherm. Res. 7: pp. 271-293.
Rowland SK & Walker GPL (1987). Toothpaste lava: characteristics and origin of a lava structural type transitional between pahoehoe and a'a. Bull. Volcanol. 49: pp. 631-641.
Rowland SK & Walker GPL (1990). Pahoehoe and a'a in Hawaii: volumetric flow rate controls the lava structure. Bull. Volcanol. 52: pp. 615-628.
Rubin AM (1990). A comparison of rift-zone tectonics in Iceland and Hawaii. Bull. Volcanol. 52: pp. 302-319.
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Swanson DA (1972). Magma supply rate at Kilauea Volcano, 1952-1971. Science 175: pp. 169-170.
Swanson DA, Duffield WA, & Fiske RS (1976). Displacement of the south flank of Kilauea volcano: the result of forceful intrusion of magma into the rift zones. U.S. Geol. Surv. Prof. Pap. 963: 39 pp.
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Tilling RI, Koyanagi RY, Lipman PW, Lockwood JP, Moore JG & Swanson DA (1976). Earthquake and related catastrophic events, island of Hawaii, November 29, 1975: A preliminary report. U.S. Geol. Surv. Circular 740: 33 pp.
Walker GPL (1988). Three Hawaiian calderas: an origin through loading by shallow intrusions? J. Geophys. Res. 93: pp. 14,773-14,784.
Walker GPL (1989). Spongy pahoehoe in Hawaii: a study of vesicle-distribution patterns in basalt and their significance. Bull. Volcanol. 51: pp. 199-209.
Walker GPL (1990). Geology and volcanology of the Hawaiian Islands. Pacific Science 44: pp. 315-347.
Walker GPL (1991). Structure and origin by injection of lava under surface crust, of tumuli, "lava rises", "lava-rise pits", and "lava-inflation clefts" in Hawaii. Bull. Volcanol. 53: pp. 546-558.
Walker GPL (1992). Puu Mahana near South Point in Hawaii is a primary Surtseyan ash ring, not a sandhills-type littoral cone. Pacific Science 46: pp. 1-10.
Wentworth CK & Macdonald GA (1953). Structures and forms of basaltic rocks in Hawaii. U.S. Geol. Surv. Bull. 994: 98 pp.
Wilmuth RA & Walker GPL (1993). P-type and s-type pahoehoe: a study of vesicle distribution patterns in Hawaiian lava flows. J. Volcanol. Geotherm. Res.? in press
Wilson L & Head JW III (1988). Nature of local magma storage zones and geometry of conduit systems below basaltic eruption sites: Puu Oo, Kilauea east rift, Hawaii, example. J. Geophys. Res. 93: pp. 14785-14792.
Wood CA & Kienle J (1990). Volcanoes of North America. Cambridge Univ. Press, Cambridge. 354 pp.