International Glaciospeleological Survey
N A T U R A L
WASHINGTON GEOLOGY
VOL. 27, NO. 2/3/4
DECEMBER 1999
2 Washington Geology, vol. 27, no. 2/3/4, December 1999
Observations of
Glacial, Geomorphic, Biologic,and Mineralogic Developments
in the Crater of Mount St. Helens, Washington
Charles H. Anderson, Jr., and Mark R. Vining
International Glaciospeleological Survey
547 SW 304 St.; Federal Way, WA 98023
INTRODUCTION
Mount St. Helens is an active andesite-dacite volcano that is currently in a semi-dormant state after a catastrophic explosive eruption in May 1980 and subsequent eruptions through 1986. During these eruptions, a dacite dome called Lava Dome has grown over the volcanic vent in the crater. Since the winter of 1982–83, the crater floor has been progressively covered by a layer of snow, firn, and ice mixed with rock debris.
This paper describes firn caves and recent geomorphic, biologic, and mineralogic developments in the crater of Mount St. Helens. The caves are a system of melt passages that have formed in the crater ice body since the mid-1980s. Glaciologists have described geothermal firn and ice caves in other volcanic craters (Kiver and Mumma, 1975; Kiver and Steele,1975; Le Guern and others, 1999) and have sometimes discussed their origin. No one, however, has yet provided detailed observations of the evolution of such a system.
On Mount St. Helens, we have had a unique opportunity to study the interaction of geothermal energy with the accumulation of alpine snowpack from its inception after a major eruption.
The International Glaciospeleological Survey (IGS) began investigative work in the crater in 1981 (Anderson and others, 1998). IGS is made up of fewer than 100 people, who are amateur to seasoned mountaineers and cave explorers. Professionally, they include a mix of scientists, engineers, and non-technical people. Yearly surveys began in 1982 with sketch mapping, description, and photography of cave passages, snow, firn, and ice. This investigation involved reconnaissance mapping and sampling from 1981 through 1998 by members of the IGS with the permission of the U.S. Forest Service and Mount St. Helens National Volcanic Monument.
Figure 1. View into the crater from the north in the summer of 1999. Firn and glacier ice reach around both sides of the Lava Dome (center). One patch of ice is clearly visible on the right. Avalanche debris falls from the far crater walls onto the ice body and becomes incorporated into it. Debris flows have formed in the loose, unstable crater floor (center foreground). The August 1997 debris flow and its semicircular scarp are slightly to the left of center. This is the locality of the calcium-rich ‘calcite’ streams referred to in following figures. Loowit canyon, with its falls and landslide, is on the left side of the photo.
CRATER SNOW, FIRN, AND ICE
A growing body of firn and ice mixed with rock debris, which we call the ‘crater ice body’, has accumulated in the crater of Mount St. Helens since 1982 (Figs. 1, 2, and 3). The shade of the steep crater walls to the east, south, and west protects this accumulation. The crater headwall rises to 2550 m (8365 ft) on the south (Fig. 4). The contiguous crater floor ice body extends from a maximum elevation of 2000 m (6560 ft) south of the Lava Dome, downward to the northeast and north around both sides of the dome. The crater floor north of the dome (1800 m or 5900 ft in elevation) hosts only seasonal snow accumulations.
The crater ice body is an incipient glacier that continues to grow. It is not readily apparent from a distance that glacier ice is present in the crater, because it is hidden by snow, firn, and rock debris. Snows stacking higher each year have locally compressed the lower layers (visible in the caves) into dense, crystalline ice. Glacier development is suggested by crevasse formation and the banded texture of alternating higher- and lower-density ice caused by recrystallization under stress (Sharp, 1960). Small areas of ice visible on the south crater wall behind the Lava Dome also exhibit crevasses and flow texture, indicating that a new glacier is forming (Fig. 4).
The crater ice body shows signs of flow (crevasses) around both sides of the Lava Dome toward the north side of the crater. At least two large radial (relative to the crater center) crevasses are present in the ice body, adjacent to the Lava Dome on the east and west. Both crevasses penetrate to the lowest layers of the ice body. We first noticed the crevasse on the northwest side in September 1994, after the roof of an ice cave collapsed.
The crater ice body has been expanding since the winter of 1982–83 (D. A. Swanson, Hawaiian Volcano Observatory, oral commun., 1999). Its volume increased from approximately 28 million m3 (37 million yd3) of uncompacted snow and firn in 1988 (Mills and Keating, 1992) to more than 53 million m3 (69 million yd3) of snow and compacted firn and ice by 1995 (our estimate). As of late 1998, Anderson and Greninger estimated that the crater contained over 71 million m3 (92 million yd3) of snow, firn, and ice‘. The thickness in places along the crater walls had reached as much as 140 m (460 ft).
Because of the limited quantity of bulk density data we collected for the crater ice body, the mean bulk density, and therefore the total mass of ice, can only be approximated. We measured the bulk density of ice at the base of a crevasse (Fig. 2) as 0.85 g/cc in September of 1994. We measured the bulk density of ice in the lowest cave passage as 0.86 g/cc in September of 1996. We obtained bulk density measurements by cutting samples with a cylindrical saw and weighing and measuring them in the field.
Our estimates and maps of the crater ice body are based on visual observations and local surface surveys. The areal distribution of snow and firn varies throughout each year. It is greatest in the spring when the winter snowfall first starts to melt. It decreases through the summer to a low in fall as winter snowfall returns. Our maps for 1997 and 1998 are not correlated to the same point in the season. Figure 3 is based on photos from September; Figure 4 is based on early season photos. The cave surveys date from September, but the snowpack areal limits do not. This doesn’t make much difference near the dome, but is substantial on the crater wall.
Figure 2. View from a helicopter looking south and down to the 1997 debris flow. The calcite covered streambed is visible as a white streak leading down from the nose of one debris-flow tongue in the lower center right. Bacteria and calcite coatings occur in streams running alongside and extending below the debris-flow tongue. A calcite-covered streambed was buried by the east lobe of the debris flow. Warm springs rise near the end of the flow. (Reproduced from videotape.)
CRATER FIRN CAVES
Bodies of firn and ice exposed to conditions above freezing tend to develop internal systems of water drainage. Flow of warm air subsequently expands these conduits, forming interconnected cave networks. The well-known ice caves of Mount Rainier occur in stagnant ice bodies such as Paradise Glacier (Anderson and others, 1994; Schmoe, 1926), the summit crater ice body (Kiver and Steele, 1975), and active glaciers such as the Carbon River Glacier (Halliday and Anderson, 1970).
The firn caves on Mount St. Helens are in the crater ice body next to the Lava Dome (Fig. 5). Cave passages form above fumaroles and fractures in and adjacent to the dome. The passages form a circumferential pattern around the dome, with their entrances on the dome flanks. Subglacial fumaroles and relatively warm air currents form and maintain the passages.
To date, we have found more than 2415 m (7925 ft) of cave passages in the crater ice body.
1 Modified from Mills and Keating (1992). We used new thickness data collected by IGS members. Our computations also included the volume of rock debris derived from the crater walls, which was estimated from changes in our topographic maps (1988 base map of Mills and Keating) and more recent photographs. We also used altimeter readings at several
locations around the Lava Dome and crater walls. We have not tried to keep a numerically defensible account of ice volumes as Mills and Keating did. They analyzed topographic maps that captured crater wall changes based on a series of aerial surveys. We have only been watching the changes in general snowpack level through the years and seasons near the dome perimeter, mainly because it directly affects our cave surveys. We used control points on the dome, placed by USGS personnel.
Our figures for snowpack volume are subject to the same correlation problem mentioned above. There are two reasons for this: (1) volume decreases, even if the mass remains the same, because of the steady metamorphism of snow to firn to glacier ice, increasing average density, and (2) melting reduces the mass (and therefore the volume even more) through a season. We used the same topographic basemap for each year. Therefore, our maps are not suitable for volumetric computations. Our maps are more records of our interpretation of the extent of the snowpack and the location of the caves.
Figure 3. Sketch map on a simplified topographic base showing the cone of Mount St. Helens, the crater ice body (delineating firn and ice areas), the August 1997 debris flow, and other features in and around the crater. Note the Loowit landslide in Loowit canyon, which is located inside TheBreach to the east of the debris flow (Fig. 1).
The cave system is dynamic, responding to ice body growth and decay processes. Ablation, caused by outside air circulation, gradually enlarges cave passages. Basal melting of the whole ice body tends to diminish the caves. Increases in geothermal activity in the crater are expressed by the rapid enlargement of ‘steam cups’, dome-shaped melt pockets localized near fumaroles (Kiver and Steele, 1975). Air circulation converts these into the typical scalloped ceiling and wall forms seen in ice caves (Anderson and others, 1994) (Fig. 6). We believe the Mount St. Helens caves to be approximately in balance with the present geothermal heat release, because they have reached an overall stable morphology. Individual passages were observed to change over time, but the system as a whole remains much the same. Changes in the geothermal activity or climate would be expected to affect the dimensions and location of these caves, as well as ceiling, wall, and ablation features.
Cave Description
We mapped the Mount St. Helens caves by compass and steel tape survey. All gear was carried on foot. We recorded our observations on the surface and inside the caves with videotape and still camera. We visually estimated the physical dimensions of rooms and cave features.
Figure 4. Sketch map on a topographic base showing the location of firn cave entrances, crevasses, rockfalls, and the surface extent of snow, firn, and glacier ice in the crater ice body on Mount St. Helens in September of 1998. The queried dotted lines between glacier ice and firn indicate that an ice front is probably concealed under the firn.
We found entrances to and mapped 15 firn caves around the perimeter of the Lava Dome from 1996 through 1998 (Fig. 5). Some have spectacular large rooms. Most have small rooms and crawlways. Cave features include scalloped ceilings and walls (Fig. 6), moulins in the ceiling, multiple domes connected by crawlways, and skylights. In winter, short-lived ice stalactites, stalagmites, and helictites form inside the caves from water dripping from protrusions on the cave ceiling (Fig. 7). Cave floors are formed by the crater floor and, in places, the dome flanks. Room sizes range from 4.6 by 4.6 by 2.4 m (15 by 15 by 8 ft) high to 12 by 24 by 6 m (40 by 80 by 20 ft) high. Most caves occur in the presence of fumaroles. Other caves form adjacent to the dome where melt water undermines the ice body.
Figure 5. The Mount St. Helens crater firn cave system as mapped in 1997. This figure is based on tape and compass survey, which is inherently prone to distortion over long distances without an opportunity to close loops (as is the case with caves distributed around the Lava Dome). Passage shapes and relative sizes are fairly accurate here, but positions relative to the dome are not—we approximated the cave positions around the Lava Dome. The ice and rock debris avalanche shown is now covered by snow and ice. Diamonds indicate locations where photos and the ice-density sample were taken.
Six main entrances and numerous smaller ones lead down the 40-degree slope of the dome flank (Figs. 8 and 9). Passages paralleling the slope contours are surprisingly horizontal. Without geothermal control, passage patterns would be dendritic and follow the crater slope.
Descending passages have vertical sides and ceilings that are convex upward. Passages paralleling the slope contours are often shaped like right triangles with the 90-degree angle located at the junction of the downslope ice wall and the ice ceiling. Floors are composed of mud with up to boulder-size volcanic rubble and slope about 30 degrees. Against the Lava Dome flanks, the slope may exceed 40 degrees.
Ridge-like accumulations of rock debris from the Lava Dome form in many places on the floor of cave passages. They are composed of unsorted, unstratified mud and rock debris derived from the upslope portion of the cave floor. In some places, these ridges are in contact with the downslope ice wall and, in others, they occur toward the middle of the passage. The ridges probably started out as rock debris caught against the passage wall. Passage walls appear to retreat in response to the production of warm geothermal gas emanations. As the walls retreat, the ridges are stranded closer to the middle of the passage.
Progressive Recrystallization of Crater Ice
Generally in ice caves, older firn is distinguished from recrystallized recent snow by textural differences and stratigraphic relationships. Winter snowpacks from multiple years persist and provide the pressure increase necessary to convert snowfall into a permanent ice body. As recrystallization continues, individual ice crystals in the deepest layers grow together to form a rigid fabric with limited permeability (glacier ice).
From 1986 to the present, we observed the gradual change from snow to firn to glacier ice in cave passages (ice bulk densities were not measured systematically). An abrupt decrease in percolating water occurred in the final stage of the transition.
An incipient glacier has developed and grown on the Mount St. Helens crater floor. Through the heavy winter snowfalls and mild summers of the 1980s and 1990s, a continued sequence of yearly net snow accumulation enabled the ice body to persist.
Geothermal Activity in the Caves
The Mount St. Helens Lava Dome is the locus of the active volcanic vent and a source of volcanic gas emanations. The caves are primarily a result of the concentration of heat. They are localized at active fumaroles and form as conduits of venting for the heated gases. They are further enhanced by the drainage of heated surface water from the dome directly into the ice body.
Hundreds of small fumaroles emit considerable quantities of steam that frequently impair visibility in the firn caves and make mapping, photography, and other observations difficult. Some of these fumaroles make audible hissing and gurgling noises. Although the rising heat and steam cause the ice walls and ceilings to drip constantly, we have not observed appreciable quantities of standing or flowing water in the caves, perhaps because the permeability of the crater floor allows seepage. Changes in passage dimensions and location (from periodic observations and resurveys of the caves) indicate changes in heat-flow and the location of volcanic emanations.
Sulfurous fumes occur locally in the caves. Gases from the numerous fumaroles and circulating surface air mix throughout the cave passages. The presence of breathable air in the known cave system indicates that volcanic gases are rapidly mixed with fresh air and removed from the caves. Earlier workers occasionally observed minor carbon dioxide accumulations (D.A. Swanson, Hawaiian Volcano Observatory, oral commun.,1999). Although we have not come across any passages with bad air, we carry portable hydrogen sulfide and carbon monoxide detectors as a routine safety precaution.
Cave Ablation
Within the caves, evaporation, sublimation, and heat conduction are the major ablative processes (Anderson and others, 1994). Since the caves are sheltered from sunlight, radiation from the sun has no direct influence on cave ablation, but energy from heated ground and fumaroles has an appreciable effect. The main control of cave ablation is the amount of air flow against the cave walls. In cave networks possessing substantial vertical relief, trunk passages tend to form as major meltwater conduits and remain dominant because air circulation is enhanced by convection.
As cave ablation and surface ablation continue through a summer season, it is normal for the cave ceiling to approach and intersect the ice surface progressively over time. If the ice is fractured, or perhaps after winter snow adds weight to the ceiling, a cave passage may experience ceiling failure. In either case, the cave system suddenly gains a vent to outside air. The effect of venting in summer is to allow cold cave air out and warm outside air in. The effect in winter is reversed.
The importance of ablation vents is exaggerated when there is any superimposed restriction in the system, such as winter snow or a rockfall blocking other entrances. In this case, the vent entrance becomes the major means of communication with outside air. When all vents to the surface are closed, the ordinary glacier cave becomes dormant. In a cave that has internal heat sources, the ablation process can continue by convection, even when all external openings are blocked. This type of system is therefore less seasonally dependent and may evolve faster than an ordinary glacier cave.
Figure 7. Ice stalagmite in the lowest cave passage (Fig. 5).
Figure 6. A typical cave passage in Mount St. Helens crater firn adjacent to the flanks of the Lava Dome (Fig. 5). Dacite boulder debris forms talus at the angle of repose, about 30 degrees. The scalloped ceiling and walls continually drip cold water during the summer, but ice stalactites form at these points during the winter. Bill Greninger, IGS team member, is looking up at scallops on the cave walls. Photo taken July 28,1997.
Figure 8. A typical cave entrance adjacent to the Lava Dome, looking in. IGS team members are climbing down the flank of the Lava Dome to the entrance to the lowest passage that is parallel to the slope contours. Photo taken Oct. 5, 1997.
Figure 9. The northwesternmost cave entrance as seen from inside the cave (Fig. 5). An IGS
member climbs with the Lava Dome in the background. Photo taken September 1997.
FAUNA OF THE CRATER AND CRATER CAVES
There is little direct evidence of animals inhabiting the crater floor, with one exception—mice were reported on the crater floor north of the dome in 1982 (D. A. Swanson, Hawaiian Volcano Observatory, written commun., 1999). Deer have visited the lower part of the crater on occasion, leaving only tracks for the careful observer to notice. We have seen insects, including honeybees, ladybird beetles, and carpenter ants, in the crater environs, presumably blown in by winds. We also found a mountain beaver skull, probably left by a predatory bird. Fauna observed during ice cave exploration include insects and ice worms that are presently inhabiting the cave and snowfield environment. Similar species are known from ice caves at Mount Rainier (Anderson and Halliday, 1969; Anderson and others, 1994).
Biologists have long sought the primitive, cold-adapted beetles of genus Grylloblatta in the glaciers and craters of Mount Rainier, Mount Baker, Mount Hood, and Mount St. Helens. We observed an unidentified species of Grylloblattain September of 1997 on the ice surface on the northwest side of the Lava Dome. Grylloblattids are also known from the Paradise and Stevens glacier caves of Mount Rainier (Halliday and Anderson, 1970).
Mountain climbers have observed ice worms (Oligochaeta: Plesiopora Enchytraeidae) of the species Mesenchytraeus solifugus rainierensis in snowfields of several Cascade mountains, especially Mount Rainier (Rod Crawford, Burke Museum, oral commun., 1998). In August of 1996, we collected a living specimen from approximately 1 cm (0.4 in.) beneath the surface of an ice wall in the largest of the Mount St. Helens firn caves. These worms are thought to migrate through the ice in a diurnal cycle, taking advantage of pore spaces between ice crystals to move about.
We collected nymph and adult stoneflies (Plecoptera: Perlodidae) of the species Rickera sorpta on the surface of the ice body and in cave interiors (Rod Crawford, Burke Museum, oral commun., 1998), which are also found in the Paradise and Stevens Glacier caves of Mount Rainier (Anderson and others, 1994). Stonefly nymphs are aquatic. The near-mature state of specimens collected at Mount St. Helens indicates that they had crawled out of water for the molt to adulthood. The dark coloration of nymphs makes them almost invisible against the dark bottom of a cave pool. Nymphs are extremely sensitive to warmth—one collected specimen expired after approximately fifteen seconds of exposure to human body heat.
GEOMORPHIC CONDITIONS IN THE CRATER Crater Floor Environment
The present crater floor is underlain by loose, porous, and permeable debris from the landslide caused by the collapse of the upper third of the volcano during the 1980 eruption. The bulk of the debris avalanche flowed downward and to the north, filling in parts of the Spirit Lake basin and upper valley of the North Fork Toutle River. Subsequent eruptions, including the later part of the May 1980 eruption, covered the landslide surface with juvenile pumice and tephra deposits, smoothing the landslide topography and creating what is known today as the Pumice Plain.
The first lava domes formed at the top of the volcanic conduit were wholly or partially destroyed by explosions (Holcomb and Colony, 1995). After the October 1980 eruption, dome growth gradually covered the fringe areas of crater-filling rockfall talus cones (Mills, 1992). These cones are intercalated with accumulating snow. The whole body was insulated and compacted by its own mass. Later tephra eruptions have added only minor amounts to the sediment pile. The most volumetrically significant addition to the post-1986 crater floor environment, therefore, is accumulated ice and rock debris. Through 1988, the rock debris fraction of post-1980 crater fill gradually dropped from 100 percent to about 65 percent of the total (Mills, 1992).
The most active surface processes taking place in the crater are (1) continued landslides from the steep crater walls, (2) fluvial down cutting in the stream courses that have established themselves across the crater floor, and (3) debris flows developing from slope failure on the north crater floor. Perhaps the most significant subsurface process acting on the crater floor contents is percolation of meteoric water and consequent alteration and leaching of the volcanic minerals. Several small surface streams flow intermittently from the crater ice body. Snowmelt and rain percolating through fractures in the Lava Dome and through the permeable crater fill, rise in geothermal springs that feed the crater streams.
Degradation of the Crater Floor
Nearly two decades of precipitation and runoff have eroded and leached material from the thick, unconsolidated mass of volcanic debris on the crater floor. Streams draining the crater have cut through this material and formed steep-walled canyons with unstable slopes (for example, Loowit canyon on the northeast flank of the crater, Fig. 1; Shevenell and Goff, 1995). These canyons are too dangerous to be used as conduits for crater access (Anderson and others, 1998). Workers in the crater have observed repeated slope failures and small slides.
In the spring of 1997, an ice and rock debris avalanche from the crater walls formed a tongue about 25 m (83 ft) in height, 150 m (500 ft) in length, and 15 m (50 ft) in width on the ice surface near the southwest side of the dome (Fig. 5). We estimated that the deposit was about 40 percent rock debris. The tongue froze and lasted through the summer of 1997.
Figure 10. (top) Deposits from the August 1997 debris flow in The Breach (Fig. 3). This view is looking downslope (north) along the path of the debris flow, the reverse of that shown in Figure 1.
Dark colored rock debris around the tongue speeded surface ablation of the ice with heat collected from solar radiation.
In August of 1997, a debris flow was triggered by the failure of a mass of saturated crater-floor material at The Breach (Figs. 1, 2, 3, and 10). The semicircular, steep-walled scarp was originally 20 m (65 ft) deep and about 150 m (500 ft) wide. The deposit extended about 700 m (2300 ft) downslope, from the 1700 m (5580 ft) elevation at the scarp brink to 1550 m (5085 ft) at the lowest point. The scarp cuts across the bed of a geothermal stream that now rises from the scarp floor and feeds clear heated pools that appear to be free of living matter. Two streams exit the scarp mouth and flow through the debris flow deposit in recently excavated gullies. One tongue of the debris flow followed the original stream and filled that stream course. Post-debris flow seepage was diverted around and through the deposit, producing additional springs and seeps throughout its length. We measured water temperatures as high as 80oC (175oF) in pools in the scarp and temperatures of 50oC (120oF) or greater downstream of the debris flow.
Another slide occurred in September 1997 in the east part of The Breach, passing down the Loowit drainage (Figs. 1 and 3). Water-saturated loose volcanic material collapsed to form a lahar that roared out of the crater and reached past Loowit Trail below on the Pumice Plain. The trail was temporarily closed for rebuilding after the slide, and only recently reopened. Similar slides must be expected in the future from the over steepened canyon walls of The Breach area. (Note steep, unstable walls in Fig. 1).
Figure 11. (middle) Samples of calcite deposits formed in a calciumrich geothermal stream below the debris flow tongues on the crater floor of Mount St. Helens. These samples were taken from stalactites that form on rock projections in the stream. The water temperature of the stream at the sample site was 49oC (120oF).
Figure 12. (bottom) SEM photomicrograph of calcite encrusting bacterial strands in a sample taken from a geothermal streambed. The rounded objects are sulfur bacteria. Note rhombohedral crystal terminations.
Photo courtesy of Robert Folk, University of Texas at Austin.
Calcite and Bacterial Growth in Geothermal Streams
Calcite (CaCO3) is actively precipitating from solution in the stream water that rises from the scarp floor mentioned above. It has formed deposits of travertine and tufa as flowstone, dripstone, helictites (cored by bacterial filament aggregates), and cave pearls.These coatings have formed on the streambed and hang from steps and waterfalls. Samples of the calcite coating (Fig. 11) exhibit compact, fan-shaped aggregates of acicular (needlelike) to bladed crystals as much as 1 mm in cross section. These appear to be pseudomorphs after aragonite bundles. Figure 12 is a scanning electron microscope (SEM) photomicrograph of a flowstone surface from a waterfall overhang. We previously (1996 and 1997) observed and filmed calcite growth in thermal streambeds now covered by the debris flow.
Figure 13A. Travertine dripstone growths at a streambed overhang. These calcite growths continue to expand forming travertine stalactites. (Reproduced from videotape.)
Figure 13B. Calcite coatings growing on and engulfing red sulfur bacteria strands. (Reproduced from videotape.)
Water percolating through freshly exposed loose material in the debris flow supplies nutrients and mineral components to the streams. Red (sulfur), orange (iron), and minor green (chlorophyllic) bacterial slime coats the streambed and accumulates in streambed pockets (Folk, 1993). We observed (summer 1998) flourishing bacterial growths in the presence of abundant water seeping from gully walls. Downstream of the debris flow for about 0.5 km, heavy coatings of calcite had grown on streambed rocks and encapsulated bacterial growths. These encrustations actively grow in flowing water and in the splash zone along the stream banks. Helictites grow as thin calcite coatings on strands of red bacteria that hang from rocks in the streambed. Calcite coatings continue to grow on and engulf the bacterial colonies (Fig. 13A,B). Remains of the bacterial growths can be found inside hollow flowstone crusts. SEM microscopy indicates the presence of bacteria and nannobacteria, similar to those described by Folk (1993), in the growths. Only incipient, very thin calcite coatings grew in the scarp pools and streams leading out of the scarp mouth.
Figure 14. Pre-August-1997 geothermal ‘calcite’ stream issuing from Mount St. Helens crater. The white calcite coating highlights the streambed. This view is upstream of the white streambed visible in Fig. 2. This section of the stream was buried by the August 1997 debris flow.
Figure 15.The geothermal stream is coated by a thick mantle of travertine up to 15 cm in thickness. This entire growth of this sample occurred during a single summer season. The tape is graduated in inches.
Calcite deposition in streams of The Breach area has been rapid and continuous (Fig. 14). From September 1997 through August 1998, at least two episodes of calcite deposition took place in a gully cut into the 1997 debris flow deposits. Older calcite-coated terraces are preserved on the walls of the gully 1 to 2 m above the present calcite-coated streambed, indicating that the newest coatings developed after the most recent gully-deepening erosion. Within a one-year period, calcite stalactites and stalagmites (Fig. 15) grew to a maximum size of 27 cm (11 in.) in diameter and 30 cm (12 in.) in length, and calcite cave pearls grew to 3.6 cm (1.4 in.) in diameter.
We believe the supply of calcium to thermal streams derives from the leaching of fresh, porous dacite in the crater by percolating meteoric water. The chief process affecting the chemistry of crater runoff has evolved from degassing of newly injected magma (waning to insignificance about 1985) to passage of meteoric water through the crater floor deposits in a manner too fast to attain equilibrium (Shevenell and Goff, 1995). Such undersaturated ground-water conditions could leach mobile components from a large volume of crater deposits. High rainfall produces a high flux of water through the dome area and out the crater mouth. Heated ground water resurges where the local unconfined water table intersects the crater floor. Farther downstream, calcite precipitates in the rapidly cooling surface streams. Nutrients derived from decomposition of volcanic material appear to support the bacterial population of crater streams.
The presence of red sulfur bacteria indicates that sulfur is an active component in the aqueous chemistry of the crater environment and a prominent source of acidity in the water that acts to digest crater rocks. Elemental sulfur from magmatic emanations interacts with oxygen-rich meteoric water to produce an acidic ground-water system in the dome area. At the hot springs (where nothing is growing and no calcite is present), the pH is about 6.5 and the temperature is 55oC (130oF). At the sampling locations where red bacterial colonies are in contact with actively growing calcite deposits, the water is somewhat alkaline (pH about 8) and the temperature is about 35o to 40oC (95–105oF).
CONCLUSIONS
A growing body of firn and ice mixed with rock debris, which we call the ‘crater ice body’, has accumulated in the crater of Mount St. Helens since 1982. Its mean bulk density is increasing with each passing year, and the transition from snow to firn to glacier ice (with active crevasses) is presently taking place. Net ice mass budget balances have been positive in the crater
since 1986, when the snowpack was first recognized to be growing.
Ice caves form above fumaroles that are located along fractures in the Lava Dome and the surrounding crater floor. Cave passages are gradually enlarged by ablation caused by geothermal sources beneath the ice and by outside air circulation. Passages grow laterally and vertically toward the surface, leading to ceiling collapse. The network of fumaroles has produced a ring of relatively horizontal passages that are connected to the surface by a number of ascending entrance passages.
Changes in geothermal activity in the crater of Mount St. Helens have become noticeable through cave passage observation and remapping. Calcite precipitated from geothermal streams on the crater floor produces coatings as thick as 15 cm (6 in.) thick in a single year. Chlorophyllic and later sulfur and iron bacteria are associated with these streams. In the summer of 1997, a small debris flow developed in the crater north of the Lava Dome, and later the same year, another flow occurred in Loowit canyon.
Increased thermal activity could mobilize crater ice to produce debris flows that could affect the discharge and sediment load in Toutle River. Our mapping and investigations of the crater environment could furnish additional indicators of geothermal activity and incipient geomorphic changes that could augment information provided by remote surveys.
ACCESS TO THE CRATER
The crater of Mount St. Helens can be a dangerous place, particularly because of snow and rock avalanches. Other hazards include invisible snow caves and unstable slopes. The potential also exists for pockets of ‘dead’ (oxygen-depleted) air and unexpected explosions and discharges of volcanic ash. The U.S. Forest Service strictly regulates access to the crater of Mount St. Helens. The area is part of the Mount St. Helens National Volcanic Monument, and special permits are required for any activities other than visitation of public facilities. The authors have a crater access permit for the purpose of scientific study. No one should attempt to approach Mount St. Helens by foot or by air without written clearance from the Forest Service.
ACKNOWLEDGMENTS
The authors are grateful to International Glaciospeleological Survey members for assistance with mapping and to staff of the U.S. Forest Service at Mount St. Helens National Volcanic Monument for logistical assistance and advice in conjunction with permits and crater entry. From the Washington Division of Geology and Earth Resources, we thank Wendy Gerstel and Patrick Pringle for critical review and Jari Roloff for editorial and graphic assistance in preparing this paper. Robert Folk of the University of Texas at Austin, Texas, provided SEM photos and identified bacterial components. Rod Crawford of Burke Museum, University of Washington, identified our insect and worm specimens.
crawlway – a cave passage that can be navigated only by crawling.
crevasse – a deep, nearly vertical fissure formed in ice, firn, or snow caused by movement over an uneven surface. A crevasse suggests that movement is taking place (Sharp, 1960).
cave pearl – an unattached, subspherical to spherical calcite concretion formed in splashing or dripping water, usually deposited on a sand particle or rock fragment nucleus.
dripstone, flowstone – mineral coatings (usually calcite, but may be other minerals or ice) deposited by precipitation from water flowing over an exposed surface, usually found in caves. The distinction indicates the nature of water flow during growth: dripstone forms free-hanging or free-standing deposits; flowstone forms as a wall or floor coating.
firn – a material that is transitional between snow and ice, being older and denser than snow but not yet transformed into glacier ice. Snow becomes firn after existing through one summer melt season; firn becomes glacier ice when its permeability to liquid water drops to zero.
glacier ice – a naturally accumulated ice that has reached a bulk density in excess of 0.82 g/cc. It possesses an intergrown crystalline matrix and flows plastically under its own weight.
helictite – a curved, angular, or dendritic twig-like growth from a flowstone or dripstone surface.
ice, ice body – an accumulated body of firn and ice in the Mount St. Helens crater, regardless of its density, texture, or fraction of non-ice content (air and rock debris).
moulin – a circular, nearly vertical hole or shaft in the ice of a glacier, formed by percolating surface water and enhanced by air circulation.
pseudomorph – a mineral whose outward crystal form is that of another mineral species from which it has been changed by alteration, substitution, or some other process.
rock debris – rock fragments that have fallen from the crater walls after the eruption.
skylight – an opening to outside light in the ceiling of a cave.
stalactite – a cylindrical or conical dripstone deposit that hangs from the ceiling of a cave.
stalagmite – a cylindrical or conical dripstone deposit that rises from the floor of a cave.
travertine, tufa – a dense, finely crystalline massive or concretionary limestone of white, tan, or cream color, commonly having a fibrous or concentric structure and splintery fracture; formed by rapid chemical precipitation of calcium carbonate from solution in surface or ground water, as by agitation of stream water or by evaporation. The spongy or less compact variety is called tufa.
REFERENCES CITED
Anderson, C. H., Jr.; Behrens, C. J.; Floyd, G. A.; Vining, M. R., 1998, Crater firn caves of Mount St. Helens, Washington: Journal of Cave and Karst Studies, v. 60, no. 1, p. 44-50.
Anderson, C. H.; Halliday, W. R., 1969, The Paradise ice caves, Washington—An extensive glacier cave system: National Speleological Society Bulletin, v. 31, no. 3, p. 55-72.
Anderson, C. H., Jr.; Vining, M. R.; Nichols, C. M., 1994, Evolution of the Paradise/Stevens Glacier ice caves: Journal of Cave and Karst Studies, v. 56, p. 70-81.
Folk, R. L., 1993, SEM imaging of bacteria and nannobacteria in carbonate sediments and rocks: Journal of Sedimentary Petrology, v. 63, no. 5, p. 990-999.
Halliday, W. R.; Anderson, C. H., Jr., 1970, Glacier caves—A new field of speleology: Studies in Speleology, v. 2, pt. 2, p. 53-59.
Holcomb, R. T.; Colony, W. E., 1995, Maps showing growth of the Lava Dome at Mount St. Helens, Washington, 1980–1986: U.S. Geological Survey Miscellaneous Investigations Series Map I- 2359, 1 sheet, scale 1:5,000.
Kiver, E. P.; Mumma, M. D., 1975, Mount Baker firn caves, Washington: The Explorers Journal, p. 84-87.
Kiver, E. P.; Steele, W. K., 1975, Firn caves in the volcanic craters of Mount Rainier, Washington: National Speleological Society Bulletin, v. 37, no. 3, p. 45-55.
Le Guern, Francois; Ponzevera, E.; Lokey, W.; Schroedel, R. D., 1999, Mt. Rainier summit caves volcanic activity [abstract]. In Northwest Scientific Association, A century of resource stewardship and beyond—Mount Rainier National Park 100th Anniversary Symposium: Northwest Scientific Association, p. 40. Mills, H. H., 1992, Post-eruption erosion and deposition in the 1980 crater of Mount St. Helens, Washington, determined from digital maps: Earth Surface Processes and Landforms, v. 17, no. 8, p. 739-754.
Mills, H. H.; Keating, G. N., 1992, Maps showing posteruption erosion, deposition, and dome growth in Mount St. Helens crater, Washington, determined by a geographic information system: U.S. Geological Survey Miscellaneous Investigations Series Map I-2297, 4 sheets, scale 1:10,500.
Schmoe, F. W., 1926, Ice caverns of Paradise: Nature Magazine, v. 7, no. 6, p. 347-348.
Sharp, R. P., 1960, Glaciers: Oregon State System of Higher Education Condon Lectures, 78 p.
Shevenell, Lisa; Goff, F. E., 1995, Evolution of hydrothermal waters at Mount St. Helens, Washington, USA: Journal of Volcanology and Geothermal Research, v. 69, no. 1-2, p. 73-94.
The latest information on Mount St. Helens is reported at the U.S. Geological Survey’s Cascades Volcano Observatory website at http://vulcan.wr.usgs.gov.
10 Washington Geology, vol. 27, no. 2/3/4, December 1999
