The origin of the name “Turtle Mountain” has never been definitely explained.
Between 1810 and 1870, Métis hunters from the Red River area followed trails north and south of the feature, to reach the buffalo herds. When viewed from the south, the upland appeared to the Métis as a turtle on the horizon with the head pointing westward and the tail to the east. Another account says that the feature was named for an Ojibwa Indian, “Makinak,” (turtle) who walked (ran?) its entire length in one day. The Ojibwa often took their names from things in nature, and the turtle was an important figure in their religious tradition. Other names that have referred to Turtle Mountain include Makinak Wudjiw, LaMontagne Torchue (French for ‘Turtle Mountain’), Turtle Hill, Beckoning Hills, and the Blue Jewel of the Plain.
Still another possible origin for the name might be the painted turtles, which are plentiful in the area today. The only “semi-official” information I could find that referred to the origin of the name was included in the early accounts of government cartographers, who noted that, from a distance, the profile of the plateau resembles the back of a turtle. Patrick Gourneau, in his book History of the Turtle Mountain Chippewa explained why “Turtle Mountain” is not “Turtle Mountains.” He states the following: “The naming of Turtle Mountain goes back a long time, versions from white men and Indians. To mention only three Chippeway versions, it indicates that it was the early Chippeway migrants from the woodlands of the east who named it Turtle Mountain. None of the three versions carry the name Turtle Mountains. As far back as my memory goes, I have not ever heard a full blood term the hills as Turtle Mountains, and same applies to the “Mechifs.” The Chippeway name is “Mekinauk Wudjiw” (Turtle Mountain). If it was Turtle Mountains it would be “Mekinauk Wudjiw wum” (plural). The “Mechifs” referred to the hills as “La Montagne Torchue.” “La Montagne Torchue” is French meaning Turtle Mountain. — quoted in Trail of Misgivings by Daniel F. Jerome, 2006, 280 p.
The name “Turtle Mountain” has been misused for so long, that it has become common practice to use the plural or the singular interchangeably to designate the area. I will use the singular form in this article.
For many people, mention of Turtle Mountain brings to mind the International Peace Garden, which straddles an area of three and a half square miles on the U.S.- Canada (North Dakota – Manitoba) border. North Dakota, after all, is known as the Peace Garden State. The Peace Garden was established in 1932 as a symbol of the peaceful relationship between the two nations. The North Dakota portion of the Peace Garden is in Rolette County on the west side of U.S. Highway 281.
Turtle Mountain rises 600 to 800 feet above its surroundings, high enough to receive significantly more precipitation than the surrounding grassland. As a result of the heavier precipitation, Turtle Mountain is forested. The hills cover an area of about a thousand square miles, half in North Dakota, and half in Manitoba. Along with river bottom land and the forested Pembina Hills to the east, Turtle Mountain is one of the few extensive wooded areas in the region. The predominant covering of aspen is interspersed with black poplar, ash, birch, box elder, elm, and bur oak. A large part of the vegetation consists of shrubs like hazel, chokecherry, saskatoon, nanny berry, dogwood, highbush cranberry (pembina), and pincherry. Fire played an important role in the development of present-day vegetation. Prior to settlement, Turtle Mountain was periodically swept by fire caused by lightning and by human activity. Plains Indians recognized that a heavy growth of new plants appeared in burned areas. They knew too that forests did not attract bison, an important food source, so they routinely set fire to the wooded areas. Prairie winds then carried the fires for many miles. This practice may represent one of the earlier attempts by humans to attract animals by manipulating the environment.
Turtle Mountain is basically an erosional feature, a broad area, resulting when younger sediments were left standing when the surrounding older materials were eroded away. Unlike the Killdeers, though, Turtle Mountain was then glaciated and the resulting glacial landforms greatly changed the area. Had the area not been glaciated, Turtle Mountain might be more like the Killdeer Mountains, although much broader and probably not so prominent a feature. The area of Turtle Mountain is underlain by rocks of the Cretaceous Fox Hills and Hell Creek formations and the Paleocene Cannonball Formation, all covered by a thick layer of glacial sediment. In early Pliocene or earliest Miocene time, five or six million years ago, the area that is now Turtle Mountain was part of a broad, northeast-sloping plain. Rivers and streams flowed over the plain from the west and southwest, making their way to Hudson Bay. Then, in Pliocene time, maybe four million years ago, erosion increased markedly and large amounts of material were removed as deep valleys dissected the plain. I am unsure why this cycle of erosion began. Perhaps the area was uplifted by geologic forces so that streams began to cut down and into the sediments over which they had been flowing or (more likely) the climate may have changed.
The erosion removed sediment and shaped new hills and valleys. Gradually, as streams carried the sediments surrounding Turtle Mountain away to Hudson Bay, a large mesa, or perhaps a range of buttes, remained where the hills that comprise Turtle Mountain stand today. The reason the outlier developed where it did is not clear. The uppermost bedrock unit (beneath the covering of glacial sediment) of Turtle Mountain is the Tertiary Cannonball Formation, which is not notably resistant to erosion. It is possible that some kind of resistant layer was present throughout much of the erosion cycle, perhaps a part of the lower Bullion Creek Formation. Additional drilling in the area may eventually penetrate a remnant of some resistant material that has not yet been found. If any resistant layer exists, it is everywhere buried beneath glacial sediments.
About three million years ago, the climate turned colder and, as snow built up to great depths near Hudson Bay, glaciers formed and the ice flowed southward, out of Canada into North Dakota. As the climate fluctuated during the Ice Age, glaciers advanced and receded, flowing over and around Turtle Mountain several times. About 25,000 years ago, the Late Wisconsinan glacier flowed southward over Turtle Mountain for the last time. During the most recent major glaciation, Turtle Mountain was continuously buried beneath the actively moving glacial ice for about 10,000 years.
The movement of the glacial ice over the obstruction formed by the Turtle Mountain upland caused the ice to become compressed, resulting in shearing within the glacier, especially on the west and north sides of the area. The shearing of the ice at the edge of Turtle Mountain caused large amounts of rock and sediment to be incorporated into the ice. As the climate moderated between 15,000 and 13,000 years ago, the glacier became thinner and its margin receded northward. Because Turtle Mountain rises 600 to 800 feet above the surrounding area, and because ice 200 or 300 feet thick can flow under its own weight, the flow of glacial ice on the lowland adjacent to Turtle Mountain continued for a while. At the same time, the glacier on top of Turtle Mountain stagnated, leaving several hundred feet of debris-covered ice covering the surface on the upland.
In areas surrounding Turtle Mountain, where shearing of material into the glacier had not been as intense, the ice was cleaner and it simply melted away, leaving only a small amount of sediment. In contrast, as the debris-covered stagnant ice over Turtle Mountain melted, the debris it contained gradually became concentrated at the surface of the ice, resulting in an increasingly thick insulating layer that greatly retarded the rate of melting. Thus, even though the glacier had stopped flowing, and had stagnated over the Turtle Mountain upland by 13,000 years ago, the layer of insulation that built up on top of the stagnant glacier kept it from completely melting for another 3,000 years. It was not until about 10,000 years ago that the last glacial ice on Turtle Mountain melted.
The glacial sediment on the stagnant glacier covering the Turtle Mountain upland was irregularly distributed and, for this reason, the ice there melted unevenly. This uneven melting caused the upper surface of the stagnant ice to become hilly and pitted with irregular depressions. The glacial sediment on and within the ice was saturated with water from the melting ice and it was highly fluid. It slid down the ice slopes in the form of mud flows and filled the depressions. Thick accumulations of debris in depressions on the stagnant glacier insulated the ice beneath, keeping it from melting quickly. Newly exposed ice, from which the insulating debris cover had recently slid, melted rapidly. The result was a continual reshaping of the surface of the stagnant, sediment-covered glacier.
The environment over Turtle Mountain gradually stabilized and the lakes flooding the sediment-lined depressions on the stagnant glacier became more temperate. Most of the water in the lakes came from local precipitation, rather than from melting ice. Precipitation at the time was greater than it is today; probably more than 50 inches a year, and the mean annual temperature was a few degrees cooler than it is today. Eventually, all the stagnant ice over Turtle Mountain melted, and all of the material that had been on top of and within the glacier was distributed in its current position, forming the hilly “collapse” topography found in the area today. These landforms are referred to by geologists as “hummocky collapsed glacial topography,” or “dead-ice moraine.” The modern landscape on Turtle Mountain is characterized by hundreds of lakes and ponds, by hummocky topography, and also by some broad, flat areas that stand well above the surrounding rougher land, along with some flat, lowland areas. Many of the higher flat areas are old lake plains, underlain by silt and clay that were once surrounded by glacial ice. These areas are referred to as “elevated lake plains.” Some of the lower flat areas are covered by stream deposits of gravel and sand. No streams flow for any great distance throughout the area.
Several years ago, billboards were posted around North Dakota in an effort to entertain and catch motorists’ attention. One of them, outside of Mandan, read “North Dakota Mountain Removal Project Completed.” The billboard referred to the image many people have of North Dakota as a flat and featureless land, but the sign ignored the fact that, within our borders are at least half-a-dozen features bearing the name “mountain.” Was the removal project a failure?
One of our mountainous areas is the Killdeer Mountains in western North Dakota, about 40 miles due north of Dickinson. Another is Turtle Mountain, home of the International Peace Garden on the North Dakota-Manitoba border. (Turtle Mountain is singular, not plural; I’ll explain later). It may seem odd that the Killdeer Mountains, Turtle Mountain, and several other features in North Dakota are called “mountains.” The idea may be related somewhat to scale. When viewed by a person who has recently traveled over eastern North Dakota, the features may be impressive, but I wonder what they might have been named if more of our settlers had come via Wyoming or Montana.
We have several more places in North Dakota that bear the name “mountain.” The town of Mountain in Pembina County was settled by Icelanders in 1873. Mountain is situated on the former shoreline of glacial Lake Agassiz, and the view from there, over the Red River Valley, is impressive. Just north of Mountain, the hilly area along the Pembina River Valley in northeastern North Dakota is sometimes referred to as “Pembina Mountain,” but the term “Pembina Hills” is commonly used as well. The steep escarpment is also referred to as the “Pembina Escarpment” or “Manitoba Escarpment.” Pioneer geologist David Dale Owen, when he traveled through the Red River Valley in 1848, commented on Pembina Mountain thus: it is “in fact no mountain at all, nor yet a hill. It is the terrace of table land – the ancient shore of a great body of water that once filled the Red River Valley.” People have been critical of the kind of mountains we have in North Dakota ever since! (Owen was from Indiana).
Other named “mountains” in North Dakota include Devils Lake Mountain in southeastern Ramsey County, Blue Mountain in western Nelson county, Lookout Mountain in northeastern Eddy County, and the Prophets Mountains in western Sheridan County. All of these features are ice-thrust hills or complexes of ice-thrust topography that stand as high as a few hundred feet above the surrounding areas. Near Medora, in Billings County, we have Tracy Mountain, but we don’t have many “mountains” in southwestern North Dakota – in that area the term “butte” is used more often. Several hundred formally named features called “hills” or “buttes” are found in North Dakota, as well as a few “points” and “ridges.” Many of these are at least as impressive as some of our mountains. We also have many features that fit the formal definition of a mesa, but very few of them have been referred to as mesas. I won’t dwell any longer on the vagaries of naming topographic features. The names don’t necessarily make much sense. We do manage to communicate, at least if we stay close to home. This article will deal mainly with the Killdeer Mountains and I will follow it with an article on Turtle Mountain. Both features are of considerable scenic beauty no matter what you want to call them and both have interesting stories to tell. The Killdeer Mountains
The Killdeer Mountains consist of two large, flat-topped buttes in Dunn County. They cover an area of 115 square miles and rise from 700 to 1,000 feet above the surrounding plains. The entire elevated Killdeer Mountain region is about nine miles long and six miles wide. The highest elevation in the area is 3,314 feet, which is 192 feet lower than the highest point in the state (White Butte). The term “Killdeer” is presumably a translation of a Sioux phrase: “Tah-kah-p-kuty” (the place where they kill the deer).
The caprock on the Killdeer Mountains consists of a 300-foot-thick sequence of siltstone, sandstone and carbonate beds that belong to the Miocene-age Arikaree Formation. One of the most conspicuous, ledge-forming units within the Arikaree Formation is found about 150 feet below the caprock. Known as the “burrowed marker unit” or “wormy marker bed,” it is a sequence of hard, erosion-resistant interbedded siltstone and sandstone with some carbonate lenses (the burrows in the bed were dug by clams living in the sediment before it hardened; the organisms that did the digging were similar to modern “shipworms”). The Arikaree Formation lies on top of the Eocene-age Chadron Formation, a sequence of yellow to green sandy mudstone, clayey sandstone, and pebbly sandstone. The tree-covered, slopes around the flanks of the Killdeer Mountains are mainly landslide topography, consisting of materials that have fallen or slid from higher up. A little farther away, the grassy or farmed, less-hilly areas are underlain mostly by the Golden Valley Formation, a Paleocene to Eocene-age rock unit. The Paleocene Sentinel Butte Formation, which underlies the Golden Valley Formation beneath the Killdeer Mountains, occurs at the surface in a broad area surrounding the Killdeer Mountain upland.
The two main buttes that make up the Killdeer Mountains coincide with areas that were once lakes in which sandy and limy sediments, along with some stream deposits, accumulated during Miocene time. Repeated volcanic eruptions in the Rocky Mountains to the west produced large amounts of ash, which blew eastward, fell to the ground, and washed into the lakes, forming tuffaceous (meaning they contain volcanic ash) sandstones.
A new erosion cycle began about five million years ago, long after the lakes had filled with sediment, and dried up. The relatively hard tuffs and freshwater limestone and sandstone beds that had been deposited in the Miocene lakes were much more resistant to erosion than were the surrounding sediments. Because of their resistance to erosion, these hard materials remained standing above the surrounding area as the softer Golden Valley and Sentinel Butte sediments were eroded away by streams and rivers. The Killdeer Mountains, with their resistant caprock, are the result of that erosion; they are the modern manifestation of ancient lake beds. The topography has undergone a complete reversal; areas that were once low are now high due to their resistance to erosion.
Two sites in the Killdeer Mountains are of particular interest. The Killdeer Mountain Battle State Historic Site is located on the southeast edge of the Killdeer Mountains, seven miles northwest of the town of Killdeer (Section 34, T. 146 N., R. 96 W.). The Battle of the Killdeer Mountains took place on July 28, 1864 when General Sully and 2,200 troops used artillery on 6,000 Teton and Yanktonai Sioux in revenge for the uprising of Santee Sioux in southern Minnesota. Sully decimated the Sioux, killing many of them and destroying their camp and equipment. Less than a year earlier, on September 3, 1863, Sully had accomplished a similar feat at the Battle of Whitestone Hill, where his troops killed, captured or wounded 300 to 400 Sioux Indians. Medicine Hole is, indeed, a hole in the ground, but it is not a cave in the traditional sense because it did not form as most caves do. No solution of carbonates was involved, and there are no stalactites or stalagmites. It is, rather, a crack in the ground, where a large block of material has begun to fall away from the main body of the southern butte of the Killdeer Mountain. Medicine Hole is located on private land and, as I write this, in 2015, the area is not open to public access. Please respect the wishes of the land owner. The Killdeer Mountains support the largest deciduous forest in southwestern North Dakota, except for the forests on the floodplains bordering the major rivers. The Killdeer Mountain forest consists largely of aspen and oak, with some ash, elm, birch, and juniper, along with shrubs such as chokecherry, willow, plum and buffaloberry. The forest is interesting in that it contains species that tend to be found in more boreal settings, areas that may be 200 miles or more to the northeast.
In summary, the Killdeer Mountains are an erosional feature, preserved because of their resistant caprock of tuffaceous sandstone and limestone. Erosion of the area began in late Miocene time, and continued into Pleistocene time, resulting in gravel-covered, flat, sloping surfaces (pediments) around the flanks of the Killdeer Mountain uplands. These gravel deposits, up to ten feet thick, were derived from the sandstone and limestone beds higher up in the Killdeers. The gravels are themselves resistant to further erosion and they help to retard the rate of the ongoing, modern erosion cycle. The current erosion cycle began when the nearby Little Missouri River was diverted by a glacier from its northerly route so that it flowed (flows) eastward to its modern confluence with the Missouri River. As a result of the diversion, and the resulting steeper gradient over which it flows, the Little Missouri River began to erode vigorously, carving the badlands through which the modern river flows today. Although the Killdeer Mountains show no evidence of ever having been glaciated, their modern topography dates largely to the Pleistocene. Old ice wedges can be seen in the pediment gravels in places, testimony to the time when the area was subjected to tundra conditions during one or more of the glacial epochs. There were no glaciers over the Killdeers, but continuous frigid conditions provided a tundra ecosystem. *Frost Wedges: Frost wedges are common, but many areas of patterned ground that have been interpreted to be frost polygons are really dessication cracks developed in silcrete. These are much older than frost wedges, such as this one, which formed in loose materials.
Governor William L. Guy’s secretary thought she felt a sonic boom on Monday, July 8, 1968. The State Capitol Building shook a bit, but most people did not feel the shaking or, if they did, they did not recognize it for what it was, a 4.4-magnitude earthquake. The earthquake was centered just southwest of Huff and it was felt over a 3,000-square-mile area. In Huff, and on nearby farms, the quake rattled dishes, window and wood-frame houses creaked, but no damage was reported anywhere in the state.
At least 12 additional earthquakes have been felt in North Dakota. The most widely felt earthquake in North Dakota occurred at about 9 p.m on May 15, 1909. It was a shock that rocked the northern Great Plains. The epicenter of this tremor was near Avonlea, Saskatchewan, near the North Dakota-Montana- Saskatchewan border. The Avonelea earthquake was felt throughout North Dakota and western Montana as well as in the adjacent Canadian Provinces. It broke windows and dishes and cracked plaster and masonry.
Some of the largest earthquakes in U.S. history, in the early part of the 19th century, were likely felt in North Dakota. They are known as the New Madrid quakes, after the town of New Madrid in southeastern-most Missouri. A series of four strong quakes (and hundreds of smaller ones) occurred – two on December 16, 1811, and one each on January 23, 1812 and February 7, 1812. The largest of the quakes was felt from the Gulf of Mexico to Canada and from the Rocky Mountains to the Atlantic coast. The potential remains for more devastating earthquakes in the New Madrid area, and if that happens, we will likely feel it in North Dakota.
Other earthquakes that have been centered and felt in North Dakota include one in the southeastern part of the state in 1872; one near Pembina in 1900; three in the Williston area in 1915, 1946, and 1982; and one each in the Hebron area in 1927; near Havana in 1934; and the Selfridge area in 1947. Earthquakes centered near Morris, Minnesota were felt in southeastern North Dakota in 1975 and 1993.
Almost all earthquakes are caused by sudden slippage along faults in the upper few hundred miles of the Earth’s outermost shell (the crust). Most of them occur at the boundaries between the several large plates, which fit together to form the crust. These plates move, in some places pulling away from one another, as along the spreading parts of the mid-Atlantic Ridge, sliding past one another, such as along at the San Andreas Fault in California, or colliding into one another, such as in the Pacific Northwest (the Cascadia Fault in British Columbia, Oregon, and Washington) where the Pacific Plate is pushing under the North American Plate. In all of these, and other comparable areas, the potential is higher for a severe earthquake.
Movement along these plates is slow but steady, most of the time approximating one to ten millimeters per year (about the rate at which your fingernails grow), but the continual slow movement causes stress to build. When the stress finally exceeds the strength of the rocks, they break and snap violently into a new position (the last time that happened in the Pacific Northwest was in 1700).The point of rupture, which may be many miles beneath the Earth’s surface, is known as the focus or hypocenter of an earthquake; the epicenter is the point directly above the hypocenter. The process of breaking (known as faulting) creates vibrations called seismic waves. We feel these waves as earthquakes. Earthquakes can occur anywhere enough elastic strain builds up to drive fracture propagation along a fault plain.
The sides of a fault may move past each other smoothly, without causing an earthquake if there are no irregularities along the fault surfaces. However, nearly all faults do have irregularities, which tend to cause frictional resistance. If a fault becomes locked–“stuck” in place, continued relative motion between the plates leads to increasing stress, and therefore, stored strain along the fault surface. This continues until the stress has risen sufficiently to break through the friction, allowing sudden sliding over the locked portion of the fault, releasing the stored energy.
Large earthquakes are among the most devastating natural events that can occur. If an earthquake occurs near the edge of a continent, it may generate a tsunami, which can result in a massive flood when it comes ashore. Tsunamis are long-wavelength, long-period sea waves produced by the sudden or abrupt movement of large volumes of water. In the open ocean, the distance between the wave crests can be more than 60 miles, and the wave periods can vary from five minutes to an hour. A tsunami wave can travel from 375 to 500 miles per hour, faster in deep water, slower if the water is shallow. Tsunamis can travel thousands of miles across open ocean, inundating far shores several hours after the actual earthquakes that generated them. One of the most devastating tsunamis in recent years occurred on December 26, 2004 as a result of a massive earthquake (known as the Sumatra-Andaman Earthquake), with a Richter Scale rating of 9.3 off the west coast of Sumatra, Indonesia. The earthquake that triggered the tsunami was the third largest earthquake ever recorded and it also had an unusually long duration of shaking, about ten minutes. The resulting 100-foot-high tsunami wave killed more than 230,000 people when it came ashore in Sumatra and Thailand.
Other recent, large earthquakes include one in 1960 in Chile and one in 1964 in Alaska (the Good Friday Earthquake). The most devastating (in terms of casualties) earthquakes in recorded history were the 1556 Shaanxi earthquake in China, which killed 830,000 people, and the 1976 Tangshan earthquake, also in China, which killed 655,000 people. The fatalities in the Chinese earthquakes were due to direct earthquake damage, not tsunamis, which more often account for most of the fatalities.
Seismic waves (from the Greek seismos, meaning “caused by an earthquake”), actually consist of two kinds of waves: surface waves and body waves. Body waves travel deep into the Earth’s mantle, and even through its core, before reaching the surface, whereas surface waves travel along the Earth’s surface.
The seismic waves of large earthquakes can induce natural oscillations in the Earth and cause the entire planet to ring like a bell for hours, or even days. The reverberating tone is much too low for us to hear, but seismographs can record the low-frequency oscillations. The recorded sound can be played back at, say, 10,000 times faster than the original. We can then, as it were, listen to the Earth. It is a strange experience that sounds like being in a forest on a windy day, with occasional brief falling tones and longer, rather melodic tones, similar to an orchestra tuning up. Every now and again we hear sharp noises that sound like a branch breaking. Sometimes we hear sounds like a herd of animals stampeding through a forest, smashing off branches and breaking them underfoot. I have listened to several such recordings. You can access samples of earthquake sounds on a variety of internet sites.
Most earthquakes that originate in North Dakota are probably related to deeply buried structures in the Precambrian basement. These structures contain numerous faults, but because they are so deeply buried, their extent and locations are poorly known. Movement on any of the faults could produce small to moderate earthquakes. Small earthquakes can also occur when layers of sedimentary rock collapse into voids left by the dissolution of underlying salt beds. Northwestern North Dakota is underlain by thick layers of salt at depths ranging from 4,000 to 12,000 feet. Salt is a geologically unstable mineral, readily dissolved in water and, when burdened under the tremendous mass of overlying sediments, it can flow and deform. As the salt moves, the support for overlying layers may be removed. The overlying layers can settle downward gradually, or they may collapse suddenly, creating a comparatively shallow, small earthquake.
Seismographs around the world record earthquakes with a magnitude of about 4.5 or greater; seismic waves of smaller tremors dissipate before being recorded by distant instruments. For an earthquake with an epicenter in North Dakota to be recorded, it would have to have a magnitude of 3.3 or greater. The 1968 Huff earthquake, which I mentioned earlier, is one of only about a half-dozen that have been instrumentally verified to have epicenters in North Dakota, although it is likely that other small reported earthquakes have had epicenters within the state. Tremors of Richter magnitude 3.0 or less are often felt by persons favorably situated, so more small tremors could have occurred in the state than instrumentally-verified records suggest. The only permanent seismic monitoring station in North Dakota is located near Maddock, southwest of Devils Lake.
In 1978, along with Alan Kehew, Erling Brostuen, and Ken Harris, I investigated reports by farmers of loud, “banging” sounds in Emmons and Dickey counties in south-central North Dakota. It was a drought year and cracks formed in fields of alfalfa. We determined that the alfalfa, which has a particularly deep root system, was de-watering the soil, causing it to shrink. The cracks – and the noises – were due to shrinkage of the soil, which caused deep cracks to form. As the cracks formed, material at their edges fell into them, many of which were 10 to 15 feet deep. The banging noises may have been due either to material falling into the cracks or, perhaps, the soil may have pulled apart with force as it shrank, causing the sounds. The occurrences I just described are not earthquakes, as they do not involve tectonic forces. They might be characterized as examples of “pseudo-earthquakes,” or perhaps “soilquakes.” Similar events – landslides, etc. – might also be mistaken for earthquakes. We published the results of our “alfalfa study” in a scientific journal: (Bluemle, J. P., Kehew, A. E., Brostuen, E. A.,. and Harris, K. L., 1978, Alfalfa and the occurrence of fissures on the North Dakota prairies, The Prairie Naturalist, Vol. 10, pages 53 – 59).
As a sort of afterthought, because I am so often asked about it, I’ll add a note on current concerns about possible earthquakes being triggered by hydraulic fracturing (“fracking”) activity in our oil-producing areas. At the depths at which they are performed in North Dakota, hydraulic fracturing procedures are unlikely to cause earthquakes. It is possible, though, that injection of waste fluids into certain geologic formations could trigger small earthquakes, as has been reported in some places (Texas, Oklahoma, Ohio, etc.). The likelihood of a damaging earthquake due to this activity in North Dakota is remote, although pollution of groundwater is possible.
The closest I ever came to directly experiencing a significant earthquake was on September 26, 1997, when my wife, Mary, and I were in Assisi, Italy. We spent several hours sightseeing in Assisi, some of it in the Basilica of St. Francis of Assisi. That evening, while we were staying in Chianciano Terme, a nearby village, an earthquake caused extensive damage in Assisi, killing several people in the church we had been in a few hours earlier.
The United States Geological Survey lists North Dakota among ten states that are least likely to suffer earthquake damage. Some other nearby “earthquake-poor” states include Iowa, Minnesota and Wisconsin. Infrequent, small earthquakes may occur near to and within North Dakota, but unless one occurs in a remarkably unfortunate location, it is unlikely that any serious damage will occur. It ought to be safe to visit churches here.
Geologists tend to concoct unusual names for the things they study. “Dead-ice moraine” may sound odd to some of you. It’s a name for a kind of landform found in parts of North Dakota. Dead-ice moraine sounds odd enough, but can you believe it is found along with things called “doughnuts” and “puckered lips”? First of all, the word “moraine” is an 18th century French word. It was coined by Horace de Saussure to refer to “a heap of earth or stony debris” (de Saussure did not initially realize he was referring to glacial deposits). I’ll explain the “dead” part of dead-ice moraine later.
Dead-ice moraine is also referred to as “hummocky collapsed glacial topography” or “stagnation moraine.” It has irregular topography, formed as the last glaciers were melting at the end of the Ice Age, between about 12,000 and 9,000 years ago. The most extensive area of dead-ice moraine in North Dakota is found on the Missouri Coteau, which extends from the northwest corner to the south-central part of the state (coteau is French for “little hill”) Other extensive areas of dead-ice moraine are Turtle Mountain in north-central North Dakota and the Prairie Coteau in the southeast corner of the state near Lidgerwood. All three areas are uplands that stand above the nearby lower land. The landforms on Turtle Mountain are identical to those on the Missouri Coteau and Prairie Coteau, but Turtle Mountain has a woodland cover, the result of several inches more annual precipitation than the other areas.
North Dakota’s areas of dead-ice moraine generally make for poor farmland as they are rough and bouldery. They do, however, include a lot of excellent rangeland and thousands of depressions, which may contain lakes, ponds, and sloughs known as prairie potholes . The dead-ice moraine of the Missouri Coteau is known as the prairie-pothole region (the so-called North Dakota “duck factory”). The dead-ice moraine is essentially undrained, except locally. No rivers or streams flow for any appreciable distance in any of the three dead-ice regions – Turtle Mountain, the Missouri Coteau, or the Prairie Coteau. Dead-ice moraine formed when glaciers advanced against and over steep escarpments as they flowed onto the three upland areas. The land rises as much as 650 feet in little more than a mile along parts of the Missouri Escarpment, which marks the eastern and northeastern edge of the Missouri Coteau. Similar prominent escarpments border the Prairie Coteau and Turtle Mountain in North Dakota and Manitoba, especially the west side of Turtle Mountain near Carbury. When the glaciers advanced over these escarpments, the internal stress resulted in shearing in the ice. The shearing brought large amounts of rock and sediment from beneath the glacier into the ice and to its surface.
Eventually, as the Ice Age climate moderated, the glaciers became thinner and could no longer flow over higher land, although they kept flowing through lower areas. When the ice on the uplands became detached from still-flowing ice, the glaciers on the uplands stopped advancing and stagnated (or “died”). As the stagnant glacial ice melted, large amounts of sediment that had been dispersed through the ice gradually accumulated on top of the ice, which was several hundred feet thick. The thick covering of sediment on the stagnant glacier helped to insulate the underlying ice, helping to preserve it and prolonging the time it took to melt. As a result, it took several thousand years for the ice to melt. Geologists have determined that insulated, stagnant glacial ice continued to exist on the Turtle Mountain and Missouri Coteau uplands until about 9,000 years ago, nearly 3,000 years after actively moving glaciers had disappeared from North Dakota.
In places where the debris on top of the ice was thickest, the glacier was slowest to melt. If little or no insulating debris covered the glacial ice, melting was quicker and the ice had entirely melted away by 12,000 years ago.
As the stagnant ice on the uplands slowly melted, the glacier surface became more and more irregular. The soupy debris on top of the ice continually slumped and slid, flowing into lower areas, eventually shaping the hummocky, collapsed glacial topography – dead-ice moraine – found today over the uplands. As the stagnant glacial ice melted, and debris slid from higher to lower places, a variety of unusual features resulted. Long ridges formed when sediment slid into cracks in the ice. Such ridges may be straight or irregular, depending on the shape of the cracks. Often, cracks that formed in a rectilinear pattern when the glacial ice was disintegrating, became partly filled with debris that slid into them. Today, we see nearly straight, intersecting ridges, where the ice cracks had been. These ridges are called “disintegration ridges.” Mounds of material collected in holes and depressions in the ice. If the mounds were cored by ice, when the ice cores melted, the centers of the mounds collapsed, forming circular-shaped ridges – “doughnuts.” Some of the doughnuts are breached on two sides because the debris cover on a mound of ice slid off two sides of the mound. Some geologists have referred to such features as “puckered lips.” Wherever part of the covering of debris slid off an area of ice to a lower place, the newly exposed ice then melted more quickly transforming what had been a hill into a hole or depression. Such reversals of topography continued until all the ice had eventually melted.
The insulating blanket of debris on top of a stagnant glacier was so thick in places that the cold temperatures of the ice had little or no effect on the surface of the ground. Trees, grasses, and animals lived on the land surface overlying the stagnant glacial ice. As conditions gradually stabilized, water collected in lakes in depressions on the debris-covered glacial ice. Most of the water in the lakes was probably the result of local precipitation rather than from melting ice. Precipitation at the time was greater than it is today, probably 50 or 75 or more inches of rainfall a year. The mean annual temperature was only a few degrees cooler than it is today.
Surrounding the ponds and lakes, the debris on top of a stagnant glacier was forested by spruce, tamarack, birch, and poplar, as well as aquatic mosses and other vegetation, much like parts of northern Minnesota today. This stagnant-ice environment in North Dakota, 10,000 years ago, was in many ways similar to stagnant, sediment-covered glaciers in parts of Alaska today. Fish, clams, and other animals and plants thrived in the numerous lakes. Wooly mammoths, bear, caribou, wapiti, and other large game roamed the broad areas of forested, debris-covered ice.
During the years I was mapping North Dakota geology, I occasionally came across Ice Age fossils in North Dakota’s dead-ice moraine: caribou bones, mammoth teeth, fossil fish (mainly perch), and various kinds of snails, but paleontologists studying the Ice Age fauna and flora in detail have found many more kinds of Ice Age fossils than I noticed.
Prehistoric people probably lived on the insulated glaciers in North Dakota 10,000 years ago without realizing the ice lay only a few feet below. Or, if they did realize it, they likely accepted it as a normal situation (and I suppose it was normal for that time). Eventually, all the buried ice melted, and all the materials on top of the glacier were lowered to their present position, resulting in the hilly areas of dead-ice moraine we see today.
Every summer, even during the coldest part of the Ice Age, some melting took place on a glacier’s surface and along its margin. Melting occurred during each summer season – even more when the climate warmed for periods of several years at a time – 20 or 30 years – periods of time comparable to the kinds of swings we see in North Dakota’s climate today. The farther south the glacier advanced, into more temperate zones, the more the amount of melting challenged the health of the glacier in those areas until a balance was finally struck between 1) the rate at which the glacial was ice advancing, 2) warmer climates to the south, and 3) overall climate warming due to the approaching end of the Ice Age. Gradually, the balance among these three factors shifted farther north (and east) and ice began to disappear in those parts of North Dakota that were glaciated.
The position of the edge of an ice sheet at any given time was determined by the balance between melting and the rate at which the glacier was flowing. While the climate remained cold (at average annual temperatures below freezing), a continental ice mass became thicker and the edge of a glacier advanced. When it warmed a little, perhaps with average temperatures a bit cooler than those we have now, the glacier margin melted back about as fast as new ice could be supplied. Even though the glacier was moving, its edge neither advanced nor receded (but melting was taking place on the glacier surface). Given still warmer conditions, the surface of a glacier melted more rapidly; the ice thinned, and the glacier’s edge melted back faster than new ice was being supplied. Areas the ice had covered gradually became deglaciated.
As the margin of a glacier melted, debris that had been frozen into the ice many miles to the north was freed and deposited on the ground. This “glacial sediment” consisted of a blended sampling of the various kinds of rock and sediment over which the glacier had flowed. Glaciers advancing into North Dakota from the northeast deposited mainly sandy, granite-rich materials they had picked up as they flowed over the Precambrian rocks of northwest Ontario. Ice coming from the northwest brought a mixture of sandy and clayey materials it had accumulated as it flowed over broad expanses of Cretaceous shale and Tertiary sands of southern Alberta and Saskatchewan. Ice advancing straight southward from Manitoba, up the Red River Valley, deposited carbonate-rich sediment it had picked up north of Winnipeg where Paleozoic limestone and dolostone are exposed today. If you travel north of Winnipeg to Stony Mountain or Stonewall, Manitoba, be sure to notice the quarries, now producing some of the same materials that glaciers brought to North Dakota, perhaps 17,000 years ago. Studying the composition of the glacial sediment is one way that geologists can determine the direction from which the ice came, and the kind of land it flowed over.
Eventually, as each glacier melted (North Dakota was probably glaciated between 10 and 20 times during the past three million years), the land gradually became free of ice. No reversal of ice flow is involved when the glacier recedes; I emphasize that “retreat” of a glacier refers to the melting of the ice. Three different kinds of ice wasting occurred, at different times and places in North Dakota. The first occurred when the glacier margin may have been far to the south, in Iowa and South Dakota. The result of melting, perhaps over a period of relative warmth of hundreds or thousands of years, resulted in the loss of much of the ice mass off the glacier’s surface. The “view from above” in North Dakota would not have changed much – everything was still all ice – but the thickness of glacial ice covering the land was diminished in thickness by hundreds or thousands of feet, even before the glacier margin had receded into North Dakota.
The second way a glacier wasted (at least from our North Dakota perspective) was when the ice margin was nearby. When that happened, wasting involved frequent change in the position of the edge of the active glacier. As a glacier melted, and after it had become thinner, its active margin gradually receded because the volume of ice arriving was insufficient to replace the ice lost at the edge due to melting. Shrinkage of this kind caused the ice margin to melt back, sometimes in a step-like fashion, the flow of ice pausing long enough at times for the forward movement of the glacier to deliver piles of sediment (moraines) to the receding ice margin. A year of glacial activity might involve the margin moving forward a short distance during a winter; then, during the following summer, the margin receding a slightly greater distance. During this phase, one of the most important things taking place, at least during summer seasons, was the deposition of large amounts of gravel and sand being deposited in front of the glacier by water flowing from the melting, sediment-laden ice. The net effect of this second phase of glacial melting was deglaciation; land that had been covered by ice saw the light of day again, after about 20,000 years.
The third way a glacier wasted involved large-scale stoppage of ice movement, leaving large parts of the glacier stranded, sometimes over broad, mainly upland areas, detached from the main body of still-actively-flowing ice on surrounding lowlands (the plains surrounding Turtle Mountain, for example). In North Dakota, this was important over upland places like the Missouri Coteau and Turtle Mountain. Areas of “stagnant,” or “dead” ice on the uplands then continued to melt slowly. Landforms resulting from the melting of such stagnant ice are distinctive and much different from those that were constructed during the step-wise retreat of active glacial ice I described earlier.
Much of North Dakota’s modern landscape reflects its latest encounter with glaciers during the Ice Age. While glaciers flowed into and over the state, carrying the pulverized rock and soil debris they had picked up along their routes, they sheared off old bedrock landforms, smeared on new layers of sediment, and built new landforms. They filled old river valleys with sediment at the same time rivers of meltwater were flowing from the glaciers carving new valleys. In some places, the glacial ice forced existing rivers to follow different routes; in other places it completely obliterated and concealed what had been rivers and valleys. Cold winds blowing over sand that had earlier been deposited on floodplains and in lakes built dunes and spread a veneer of silt (loess) over much of the state.
Most of the sediment associated with the action of glaciers of the most-recent glaciation is soft. It is “unconsolidated,” and does not hold together well (you can dig it with a shovel). An exception: earlier glaciers also deposited sediment. Nearly all of this earlier sediment has eroded away, but in those places where we have found it exposed, or drilled into it, it may be cemented. A jackhammer may be more appropriate than a shovel for digging in such cemented deposits. However, the softer, looser materials that form most North Dakota glacial deposits are much more common. Sediments related to glaciation in North Dakota can be grouped into three main types: till, lake sediment, and outwash.
1. Till was deposited directly from the ice, mostly in the form of mud flows, which slumped or flowed into their current position as the ice melted. Till consists of silty, sandy, pebbly clay, as well as cobbles, or even large boulders.
2. Lake sediment is layered material that was deposited in lakes, which formed on and near the glacier. Such sediment consists mainly of layers of fine-grained silt and clay, deposited on lake floors, along with some sand and gravel, which collected as beaches along the shores of lakes , many of which were dammed by glaciers.
3. Outwash consists of material deposited by running water. Some outwash may be cemented into a kind of stony concrete, but most of it is loose sand and gravel that was washed out of the melting glacier (hence the name “outwash”). Outwash was deposited by streams and rivers flowing through meltwater valleys or as broad, often nearly level sheets of sand ahead of a melting glacier.
Where they are present, sediments deposited directly by glaciers, and by wind and water associated with glaciation, form a thick covering on top of much of the preglacial (bedrock) surface. In central North Dakota, in Sheridan County, the glacial sediment is over 700 feet thick in places. Ten miles northwest of Tolley in northwestern Renville County, it is at least 800 feet thick, the thickest I can document in the state. The amount and thickness of glacial sediment can vary considerably over short distances so it is likely that even thicker deposits than I mentioned exist in places. Over much of the glaciated part of the state, the glacial materials average 150 feet thick.
At any given location, the glacial deposits may consist of two or more layers of till, interbedded with lake beds, alluvial sediments or other materials. In some places, soils, which had developed on the surface of an earlier glacial, river, lake, or wind-blown deposit, were buried when a new layer of glacial material was deposited. These old, buried soils (paleosols) were formed during long intervals of weathering and exposure, like the one we are enjoying now. Paleosols are among the best indicators in the geological record for multiple episodes of glaciation during the Ice Age. The characteristics of a paleosol also help us understand the climatic conditions (forest or grassland, wet or dry, cool or warm, etc.) at the time it formed.
The best places to see several multiple layers of glacial deposits in North Dakota are near Riverdale, at the Wolf Creek inlet to Lake Sakakawea in McLean County and in Beulah Bay, about 15 miles north of the town of Beulah in Mercer County. In both locations, two and, in some places, three discrete till units, separated by cemented gravel layers or paleosols, are being eroded by waves along the lake.
Drill-hole data in eastern and southeastern North Dakota provide evidence that at least a half dozen glacial advances have occurred there since the Ice Age began. Parts of southwestern North Dakota were glaciated during some of the earlier glaciations, but (apart from some rare exceptions) glacial landforms are not found there today because they were eroded away long ago. Glacial lake sediments and river gravels containing glacially derived materials can be found as far southwest as Dickinson and near the Killdeer Mountains, and I have tentatively identified patches of hard, cemented till and glacial river gravels near Bowman and Rhame, places usually considered never to have been glaciated.
Modern soils are an important link to our geologic past. Fresh glacial deposits consist of a mixture of materials, and, because their sources are so varied, they provide the combination of nutrients necessary for fertile soil. In the glaciated part of the state, North Dakota’s soils consist of the weathered exterior of materials left by glacial action. In the thousands of years that have elapsed since the ice sheets disappeared, constantly changing climate, physical and chemical weathering, accumulation of prairie and woodland plant litter, development of root systems, and burrowing activity by organisms have all contributed to the transformation of glacial deposits into the rich soils that form the basis for much of our agricultural wealth.
Glaciers in North Dakota: Part One
Glaciers are giant bodies of ice, formed from snow that survives from year to year. Accumulations of snowfall from past years compact into a substance called firn, a recrystallized residue of snow left over from past seasons. During the summers, when temperatures are warm enough for rain instead of snow, the rainfall adds to the mass of a glacier, eventually freezing and becoming part of the glacier. With time and additional snow cover, the whole mass gradually solidifies into hard ice: a glacier.
The color of pure glacial ice, if it is clear and lacking the various rocks and sediments often found in glaciers, is ice-blue. Drop a piece of glacial ice into a glass of warm water and it may literally “explode.” Any air trapped in the ice, thousands of years ago, and pressurized by the great overlying weight of the glacier, escapes with force from the piece of ice as it melts in the glass. By analyzing these trapped pockets of air, scientists can learn what our atmosphere consisted of in the past.
When an ice mass becomes thick enough and heavy enough to flow, it “officially” becomes a true glacier. A glacier flows slowly away from the place where it is thickest. It may flow at a few feet a year although, in some circumstances, the flow rate may be much faster. In the northern hemisphere, glaciers expand mainly southward, away from polar regions because the temperatures to the south are warmer than to the north. A glacier flows most easily when it is warmer and less brittle. It will move much faster at 30 degrees F. than at minus 30 degrees F.
A glacier doesn’t glide placidly over the land. It scratches and grinds the underlying bedrock surface, picking up pieces of rock and soil, dragging them along and using them as tools to scour the ground beneath the ice.
Glaciers don’t normally flow uphill, but they do fill lowlands and overtop them, much like flood water.
Glaciers in mountainous areas, unlike broad ice sheets in places like Greenland and Antarctica, come with a sense of “scale.” The mountain peaks in the distance and the valley walls that hold the glacier help you to orient yourself. But suppose you are standing on a snow-covered, continental-size glacier (dressed in a heavy parka). You see nothing but frozen wasteland, nothing but whiteness. No buildings, no fences, no trees, no landmarks. Only emptiness. No sound but the wind. On a cloudy day, sky and ice blend; making it nearly impossible to distinguish the horizon marking their boundary.
The most-recent major glacial episode in North Dakota is referred to as the Wisconsinan glaciation. It began approximately 90,000 years ago and ended 11,500 years ago, but glacial conditions were not continuous during that entire time. An initial pulse of glaciation (the Early Wisconsinan, 90,000 to 70,000 years ago), was followed by withdrawal of the ice, which was probably complete by 65,000 years ago. Between 65,000 and 35,000 years ago, North Dakota’s climate alternated between combinations of warmer, wetter, cooler, and drier periods, much as it does today. Then, about 30,000 years ago, a second major pulse of glaciation, the Late Wisconsinan, began. The Late Wisconsinan glacier reached its maximum extent between 18,000 and 16,000 years ago when it covered all but the southwestern part of the state. The position of the Missouri River approximates the maximum extent of the Late Wisconsinan glacier. Active glaciers melted completely from the state by 11,500 years ago.
By the time each glacier advanced, the land ahead of it may have become deeply frozen permafrost and, as the ice moved over the frozen rock and soil, it picked up chunks of these materials, incorporating them into the glacier itself. In the areas where it formed, west of Hudson Bay, the materials beneath the thickening ice were mainly crystalline igneous and metamorphic rocks such as granite and gneiss. Farther south, the glaciers flowed over layers of sedimentary rock. Whatever the ice flowed over, it picked up some of it and carried it along.
A moving glacier may be likened to a huge excavation and grading machine that does its job of eroding by plucking and abrasion. Plucking, the more important of the two, is based on a freeze-thaw cycle. The cycle begins when downward pressure melts the ice at the base of a glacier. Water seeps into cracks in the rock beneath the glacier. When the water freezes, the expanding ice plucks rock fragments and incorporates them into the debris near the base of the glacial ice.
After a glacier covers an area for a while, and a considerable thickness of ice lies on the land, the materials beneath a glacier gradually thaw, a combined result of the pressure of the overlying ice and the natural upward flow of heat from the Earth’s interior. The ground surface beneath the North Dakota glaciers was not frozen, and the base of a glacier may have been a muddy mass. This facilitated even more sediment being incorporated into the base of the moving ice.
In some places ground water in the saturated sediments beneath the heavy weight of the glacier built up great pressures due to the weight of the overlying ice.
Besides transforming materials beneath the ice into a mud-like mixture, the water, because it was pressurized, tended to force — squeeze–the sub-glacial sediments upward, into the base of the moving glacier. Sediments beneath the ice were smeared out as they were carried along with the advancing mass of ice.
A glacier seldom behaves like a bulldozer, pushing debris ahead of it. It does, however, incorporate debris as it moves by freezing it onto its base. In whatever way the boulders, gravel, sand, silt, and clay beneath a glacier became part of the moving glacier mass, both ice and sediment flowed forward and the farther the glacier traveled, the more material it accumulates. Glaciers advanced over North Dakota several times, and each time, when they melted, they dropped their entire load of rock and sediment, material gathered from places previously overridden. Some of the material carried by the glacier ended up far from where it had originated. We find rocks in North Dakota that came from northern Saskatchewan, Manitoba, and Ontario. We also find chunks of shale that came from only a few dozen feet away. The sediment a glacier was carrying finally came to rest when the last ice melted.
During the active life of a glacier, every crystal of ice, every boulder, sand grain and fragment of rock within the ice, is moving, slowly making its way away from the center of snow and ice accumulation.
In North Dakota, the movement was generally southward, away from the Keewatin center of ice accumulation west of Hudson Bay. Apart from its overall southward progress, variations in the topography over which the glacier advanced locally affected the direction of flow.
When the glaciers that covered much of North Dakota eventually melted, all of the material they had been carrying was laid down on the land surface where the glacier had been. This included everything from large boulders (erratics) to fine-grained material: sand, silt, and clay. This “glacial sediment,” deposited directly from the melting ice, is known as “till.” Till was deposited as a kind of stony mud that eventually dried out after the glacier melted away. Usually, the till amounted to a few tens of feet of material, but after several glaciations, it might have accrued to several hundred feet: the materials from several glaciations, stacked one on top of another.
Agreement on the origin of the name of the “Missouri” River is difficult because too many contradictory explanations exist. The name apparently comes from a Siouan Indian word, “ouemessourita” or “emissourita,” translated by early French explorers as “those who have wooden dugout canoes,” or “river of the large canoes,” or “town of the large canoes,” or any of several other possibilities. One source says the term was the name the Illinois Indians used for the native people who lived in the Mississippi River Valley, probably mainly on the eastern (Illinois) side of the river; another source says they lived in what is now the State of Missouri.
The Missouri River originates near Three Forks, Montana, where the Gallatin, Jefferson, and Madison rivers come together. It flows 2,341 miles to St. Louis, where it joins the Mississippi River. This makes the combined Missouri-Mississippi River, at 3,709 miles, the fourth longest river in the world, after the Nile, Amazon, and Yangtze. The entire length was once riverine environment but, due to the dams that have been built along its route, approximately a third of the length is now reservoirs – lake environment rather than river. Listed from upstream to downstream, the dams are: Fort Peck in Montana, Garrison in North Dakota, Oahe, Big Bend, and Fort Randall in South Dakota, and Gavins Point on the South Dakota-Nebraska border.
Along with its valley, the Missouri River is largely a product of glaciation. Before North America was glaciated, all the drainage in North and South Dakota, eastern Montana, and northern Minnesota was north or northeastward into Canada. There was no “Missouri River” carrying drainage from the northern mid-continent region to the Gulf of Mexico. The way I define the Missouri River requires that its water ultimately reach the Gulf of Mexico as it does today, and that it carry water draining from the Rocky Mountains and northern Great Plains. Prior to glaciation, no such river existed. Why is the situation today so different than it was before the Ice Age?
The modern Missouri River Valley in North Dakota consists of several discrete valley segments that differ markedly from one another. Some of the segments are broad: six to twelve miles wide from edge to edge, with gentle slopes from the adjacent upland to the valley floor. Others are narrow: less than two miles wide, with rugged valley sides – even badlands slopes in places. Most of the wide segments trend from west to east whereas the narrow segments are mainly north-south. The Bismarck-Mandan area is one of only a few exceptions, and I’ll explain why shortly.
The west-east segments of the Missouri River Valley are wide because they coincide with much older valleys that existed long before the area was glaciated. Old, mature river valleys, which formed over long periods of time (hundreds of thousands or millions of years), tend to be broad with gentle slopes. Younger valleys formed more quickly (tens of thousands of years), and are usually narrower with steeper sides. An example of a wide segment is the forty-mile-long, west-east segment of the Missouri River Valley upstream from Garrison Dam. This part of the valley, now flooded by Lake Sakakawea, was once the route of a river that flowed east to Riverdale. However, the river didn’t turn south at Riverdale, as it does today. Rather, it continued eastward past Riverdale, and on past Turtle Lake and Mercer, flowing into northeastern North Dakota. For convenience, I’ll refer to this ancient river as the “McLean River.”
East of U. S. Highway 83, the route of the old McLean River valley is a broad, low area, partly buried beneath tens to hundreds of feet of glacial sediment. Lake Audubon, Turtle Lake, Lake Brekken, Lake Holmes, Lake Williams, Lake Peterson, Pelican Lake, Blue Lake, Brush Lake, and other smaller lakes mark the former route of the McLean River through eastern McLean County. However, continuing farther east, the McLean River valley becomes so deeply buried beneath glacial deposits that it would be nearly impossible to know its route from a study of the surface topography. Fortunately, hundreds of test holes were drilled during studies of the ground water resources of the glacial deposits so we have a good idea of the route the river followed into northeastern North Dakota.
Another wide, west-east trending segment of the modern Missouri River Valley, between Stanton and Washburn, is an eastward continuation of the modern Knife River. Prior to glaciation, the Knife River flowed east in its modern valley to Stanton, but it continued eastward from there, past Washburn. A few miles east of Washburn it turned slightly northeastward. The ancient Knife River joined the McLean River near the town of Mercer and the combined Knife-McLean River continued northeastward to the Devils Lake area. It then flowed north along the east side of Turtle Mountain area into Canada.
Still another wide segment of the Missouri River Valley in northwestern North Dakota extends from near the modern Missouri River /Yellowstone River confluence, northeastward to Williston.
This six-to-eight-mile-wide section of the valley coincides with the pre-glacial route of the Yellowstone River through that area. Prior to glaciation, the Yellowstone River continued to the north, past Williston, following a route that is now mainly buried. The pre-glacial route coincides with the modern route of the Little Muddy River as far as Zahl, about 30 miles north of Williston. North of Zahl, the old Yellowstone River valley into Canada is so deeply buried that its route is known only through drill-hole data. The river entered Saskatchewan about six miles north of Crosby.
The Missouri River Valley between Williston and New Town, now flooded by Lake Sakakawea, follows the same route as did an east-flowing, mid-Ice Age — but probably not pre-glacial – river. This part of the Missouri River Valley is somewhat narrower than most other east-west segments of the valley in North Dakota, and it is also younger than most of them. It is a continuation of a mid-Ice Age river that flowed east from Montana. In Montana, the route of this river coincides with the modern route of the Missouri River past Wolf Point, Poplar, and Culbertson. The Montana segment of the mid-Ice Age river joined the north-flowing Yellowstone River near Buford.
At Bismarck-Mandan, the Missouri River Valley is about two miles wide at the Interstate Highway 94 crossing, but on the south side of Bismarck the valley broadens to six miles wide. The widening southward seems contrary to my earlier comment that north-south segments of the valley tend to be narrow. There is a reason for this exception though. The valley widens at Bismarck-Mandan because, prior to glaciation, the Heart and Little Heart rivers, which today flow into the Missouri River, joined a few miles east of Bismarck. The combined (preglacial) Heart/Little Heart River continued flowing eastward, joining the Cannonball River in southern Burleigh County, near Moffit. The old, combined Heart/Little Heart valley still exists as a broad lowland south and southeast of Bismarck. It is now a wide spot in the Missouri River Valley.
The Heart/Little Heart river system was probably dammed several times by glacial ice advancing westward. Each time a glacier advanced, a lake formed ahead of – west of – it in the Heart/Little Heart valley. The lake (or lakes) are referred to as glacial Lake McKenzie. At least once, and possibly several times, glacial Lake McKenzie overflowed, carving what is now the Missouri River valley south of the Bismarck-Mandan area.
When the (preglacial?) Heart River flowed eastward, through the Bismarck area, it deposited a thick gravel deposit which now lies buried about 100 feet beneath the Missouri River. Bismarck’s new (2013) water-intake structure withdraws ground water from this old Heart River gravel deposit.
When the McLean River valley was blocked by a glacier in the Riverdale area midway through the Ice Age, a large proglacial lake formed ahead (to the west) of the ice in the valley. This lake might be considered to be the “original” Lake Sakakawea: an early ice-dammed lake that predated the Corps of Engineers version of Lake Sakakawea by thousands of years. When the lake overflowed, near where Garrison Dam is today, the resulting flood quickly carved a narrow spillway trench south to the Stanton area.
Similarly, the Knife River, which flowed past Stanton and on to the Washburn area, was dammed by glacial ice just east of Washburn and the valley was flooded upstream beyond Washburn. The resulting lake overflowed and spilled southward into the Burnt Creek-Square Butte Creek drainage, carving a narrow trench from a few miles east of Washburn to the Bismarck-Mandan area. The modern Missouri River flows in that trench today.
And, as I noted, when the Heart/Little Heart River was dammed by a glacier, which probably advanced as far west as Sterling, glacial Lake McKenzie formed. The lake overflowed southward, forming a new valley, now flooded by the northernmost part of Lake Oahe.
The youngest and narrowest segment of the Missouri River Valley in North Dakota is at New Town, between Four Bears Bridge and Van Hook Bay. As recently as 13,000 years ago, a glacier blocked the Missouri River from its route around the north and east side of New Town. The old river route (prior to 13,000 years ago) is now a broad valley, known as the “Van Hook Arm,” flooded by Lake Sakakawea. The glacier dammed the valley, causing a lake to form upstream (to the west) of the point of blockage. Thick layers of lake sediment, known as the “Crow-Flies-High silt,” were deposited in the ice-dammed “Crow-Flies-High Lake.” Crow-Flies-High Lake extended westward from the New Town area to near Williston. In many places between these two cities, exposures of the bedded lake silt deposits occur at elevations as high as 70 feet above the modern, maximum reservoir level (1850 feet) of Lake Sakakawea. The lake rose until it overflowed southward, cutting the channel now spanned by the Four Bears Bridge.
Other “Missouri” River Routes
Up to now, I’ve tried to explain the origin of the modern route of the Missouri River. That’s not the end of the story though. The modern route of the Missouri River is only the most recent of many routes that earlier “Missouri” rivers followed through North Dakota at various times during the Ice Age. These rivers also carried runoff water from as far away as the Rocky Mountains, through North Dakota, on its way to the Gulf of Mexico. However, most of these routes, mainly in northern and eastern North Dakota, are now buried beneath thick accumulations of glacial sediment. Whatever routes these rivers followed, they had to have flowed generally eastward and southward because their original, northerly and northeasterly routes into Canada were blocked by ice each time glaciers advanced into the state. Test drilling, done to study ground water resources, has helped us identify least least some parts of the old “Missouri” River routes. There are dozens of them.
One of several early routes of the Missouri River, determined by test-hole drilling, took the river southward past Cooperstown and Valley City to the southeastern corner of the state. Another route took the river southeastward past Jamestown. In the northern part of the state, rivers like the Yellowstone were diverted from their northerly routes to easterly and southeasterly routes, past places like Columbus, Kenmare, and Minot. These buried valleys can be considered to be early “Missouri” River routes. The array of buried river valleys is really amazing – and so complicated – and such a great number of possible routes exist, that it is impossible to work them all out. All of them are now buried beneath hundreds of feet of glacier sediment, and most of them have no surface evidence whatsoever.
However, not all of the early “Missouri River” routes through North Dakota are deeply buried. In the western part of the state, a version of a Missouri River formed when an early glacier advanced at least as far southwest as the Hebron area. The margin of that glacier coincided with what is now a prominent, broad valley, known as the Killdeer-Shields channel. The channel extends southeastward from the Killdeer Mountains, past Hebron and Glen Ullin, to the Fort Yates area, crossing the modern Missouri River Valley, and continuing through southwestern Emmons County into South Dakota. No river flows through the Killdeer-Shields channel today, but an early Missouri River flowed in it, perhaps for a longer period of time than the current Missouri River has flowed in its modern route. Interstate Highway 94 crosses the valley about half way between Dickinson and Mandan. Good views of the Killdeer-Shields channel can be seen just north of Richardton and between Hebron and Glen Ullin. Old U.S. Highway 10 and the Burlington Northern Santa Fe Railroad follow the old channel from Hebron to Glen Ullin.
I realize that my description of the changes in the routes the various “Missouri” Rivers have followed since the Ice Age began is complicated. Even so, it doesn’t begin to account for the evolution of all of the changes in the vast array of routes that rivers followed during the Ice Age in North Dakota.
Most of the narrow, north-south segments of the modern Missouri River Valley correspond to places where glaciers diverted then-existing rivers southward. Glaciers in the central part of the state diverted northeast-flowing rivers, like the Knife, McLean, and Heart-Little Heart, and Cannonball, forcing them to flow southward from the points of diversion, forming the north-south segments of the modern Missouri River. Glaciers advancing into northwestern North Dakota diverted mainly north-flowing rivers, like the Yellowstone and Little Missouri, away from their routes into Canada, forcing them to flow to the east and south.
The modern Missouri River Valley is a “composite” feature, consisting of older, wide pre-glacial segments, formed over long periods of time prior to the Ice Age, along with younger, narrow segments that were cut relatively quickly at various times during the Ice Age. The parts of the Missouri River Valley that extend mainly from west to east are wider and much older than are the narrower segments that extend from north to south. Many of the early “Missouri” River routes followed for varying periods of time during the Ice Age in northern and eastern North Dakota were later buried beneath thick deposits of glacial sediment.
The current route of the modern Missouri River Valley is only the latest in a continuing series. After the next glacier has come and gone, a new version of the Missouri River will likely follow a different route than does the river today.
Rain and melting snow, wind, frost, and other forces of erosion have carved our badlands into intricate shapes. Since the Little Missouri River began to form the badlands, it has removed an enormous amount of sediment from the area. In the southern part of the badlands, near the river’s headwaters and close to Devils Tower in northeastern Wyoming and adjacent Montana, the river has cut down about 80 feet below the level at which it had been flowing before it was diverted by a glacier farther north. Near Medora, the valley floor is 250 feet lower than the pre-diversion level. Still farther downstream, in the North Unit of Theodore Roosevelt National Park and near the confluence of the Missouri and Little Missouri rivers, and nearer to where the glacier diverted it, the east-trending portion of the Little Missouri River flows at a level that is 650 feet deeper than when it was diverted.
The average rates of erosion in the badlands, assuming they started to form about 640,000 years ago, can be calculated as follows:
Headwaters area in Wyoming: 0.15-inch/100 years;
Medora area: 0.5-inch/100 years;
Confluence area near Mandaree – Missouri and Little Missouri rivers: 1.25 inch/100 years.
These rates may seem tiny but, over time, erosion has removed a huge amount of sediment. Approximately 40 cubic miles of sediment have been eroded and carried away by the Little Missouri River from the area that is now the badlands. Most of that sediment now lies beneath the water of the Gulf of Mexico.
The rates of erosion I’ve noted are long-term averages, but erosion goes on at highly irregular rates. Locally, considering only the past few hundred years, the badlands have undergone four separate periods of erosion and three periods of deposition. Since about 1936, new gullies have been cut to their present depths. It may seem a paradox that, although running water is the main agent of erosion, badlands formation tends to be most intense when water is in short supply. Why? Because erosion tends to be more vigorous during times of drought when the vegetative cover is too sparse to protect the soil from the occasional rain storm or spring snow melt. When precipitation is sufficient for the growth of heavy vegetation, the soil is better protected from severe erosion.
Streams and rivers carry sediment away from the area of the badlands, but most of the actual “on-the-spot” erosion is a result of slopewash. In places where vegetation is sparse, the soil and rock materials are easily weathered, forming loose surfaces that slide downslope easily, slumping and sliding during showers or when the snow cover melts.
The Badlands Landscape
The shapes, sizes, and configurations of the hills, buttes, valleys, and other landforms in the badlands are not entirely happenstance. Differences in hardness of the materials result in differences in resistance to erosion. Nodules and concretions help to shape a landscape ranging from beautiful, to desolate – even grotesque. Hard beds of sandstone or clinker cap many of the small buttes. Variations in permeability (permeability is a measure of the ease with which water can move through porous rock) have similar effects; rain and melted snow soak into the more open and permeable sands, resulting in only minimal erosion. When water flows over the surface of tighter, less permeable sediment, such as clay, it abrades and erodes the material, carrying some of it away. The presence or absence and the character of the vegetation also play important roles in governing the rate of erosion. Grass usually helps to control erosion more effectively than does forest vegetation.
The irregular placement of hard nodules and concretions may result in the development of rock-capped pillars, known as “hoodoos,” mushroom-like shapes perched on stalks of clay. In places, slopes are covered by nodules of siderite (iron carbonate). As they weather out of the surrounding materials, becoming concentrated on the surface, the copper-colored nodules form an erosion-resistant armor, which temporarily slows the rate of erosion. Clinker beds are also much more resistant to erosion than are the softer surrounding beds. We commonly see buttes capped by red clinker beds.
Erosional “pipes” sometimes form in gullies and ravines where surface runoff is focused. “Piping” results where runoff can flow downward into small cracks and joints. Pipes are common in places where surface runoff erodes cavities vertically downward through the soft rock. With time, the initial pathways may widen at depth into caves the size of small rooms. The average depth of vertical pipes is about 10 to 15 feet, but some are much deeper. The tops of pipes may be partially concealed making hiking treacherous. I have seen the bones of animals, such as rabbits and deer, at the bottoms of pipes (so far I haven’t seen any human bones). The animals fell into the holes and could not get out.
The geology is only part of the badlands story. The weather and climate, vegetation, animals, birds, insects, sounds and aromas–all of these, along with the human history and the ranching heritage, work together to complete the story of the badlands.
I think the North Dakota badlands are particularly beautiful because of their parklands; wooded areas that occur in draws and on north-facing slopes. Heavy vegetation in the badlands in places like Little Missouri State Park adds to the scenery. Evergreens, such as the Rocky Mountain juniper, ponderosa, and creeping juniper are interspersed with quaking aspen, cottonwood, and poplar. Limber pines are found in the badlands in the southwest corner of the state, near Marmarth.
I’ve hiked and camped in the badlands many times. Evening summer showers accentuate the colors and the clinker beds assume intense shades of red and orange. The fresh, pungent aroma of wet sage and cedar enhance the experience. At night, the stark, intricately eroded pinnacles can seem unreal. In the moonlight or in a night lightning storm, it is easy to imagine the strange shapes as ruins of a magical city, rather than structures of mere sand and clay. Blend in the sound of coyotes conversing and the badlands environment is complete.
If asked what he or she knows about North Dakota’s geology, an average resident will likely mention the badlands first. That’s true too of visitors, many of whom come to the state to see our best-known natural feature, the scenic badlands along the Little Missouri River.
The badlands landscape is a rugged and hilly one, best viewed from above, looking down on the hills, not up at them, as we usually view buttes. From the rim of the “breaks,” the point where we descend into the badlands, an intricately eroded landscape of sparsely wooded ridges, bluffs, buttes, and pinnacles lies before us. Black veins of lignite coal may be seen eroding out of the steep badlands slopes. Reddish bands of clinker add vivid colors to the area. Pieces of petrified wood, as well as fossil stumps and logs, litter the surface. Behind us stretch rolling plains, interrupted only by occasional buttes.
The American Indians, who inhabited the area when the European settlers arrived, referred to badlands as “mako sica,” (“land bad”). Early French explorers translated and added to this, referring to “les mauvais terrers a’ traverser” (“bad land to travel across”).
General Alfred Sully, preparing to cross the badlands in August of 1864, described them as “hell with the fires burned out.” Theodore Roosevelt, who lived for a while in the Little Missouri Badlands in the 1880s, described them as “fantastically beautiful.” I prefer TR’s description.
Age of the Badlands Materials
Badlands topography is found in several places on the plains of the U.S. and Canada. The best-known badlands in the United States are the extensive “Big Badlands,” along the White River in western South Dakota. Near Dickinson we have the “South Heart Badlands” (known also as the “Little Badlands”) where we find layers of sedimentary rock, equivalent (same materials, same geologic age) in part to those in South Dakota’s Big Badlands. The South Heart Badlands are an erosional remnant of what was once a large butte or group of buttes. The South Heart Badlands are carved mainly from strata of Eocene and Oligocene age, ranging between 55 and 25 million years old. The youngest beds belong to the Miocene Arikaree Formation sandstone (22 million years old), which caps some badlands buttes.
North Dakota’s Little Missouri Badlands extend from near the Little Missouri River’s headwaters in Wyoming near Devils Tower to the point where the Little Missouri River joins the Missouri River in western North Dakota. The materials being eroded in these, our most extensive area of badlands, are much older than those in the South Heart Badlands.
The oldest materials in the badlands are in the southwest corner of the state, near Marmarth, where Cretaceous-age Hell Creek Formation beds (about 65 million years old) have been carved into badlands. The dark and somber, gray and purple beds of the Hell Creek Formation contain dinosaur fossils. Small patches of badlands, carved from the Hell Creek formation can also be seen along State Highway 1806 between Huff and Fort Rice in Morton County.
However, the main area of the Little Missouri Badlands is that which has been carved largely from the Paleocene Bullion Creek and Sentinel Butte formations, which were deposited between 58 and 56 million years ago. The beds that have been eroded into these badlands are too young for dinosaur fossils; the dinosaurs were already extinct when they were deposited.
Between 70 and 40 million years ago, a major mountain-building event known as the Laramide Orogeny (orogeny = “mountain forming”) formed the Rocky Mountains in Montana and Wyoming. As the mountains rose, they were attacked by intense erosion, providing sediment to eastward-flowing rivers and streams. The rivers delivered the eroded sediment to western North Dakota’s coastal plain, an area that could be likened to today’s Mississippi River Delta (central North Dakota was an inland sea at that time). Sediment from the eroding mountains accumulated into thick layers of soft, poorly lithified siltstone, claystone, and sandstone: materials that were deposited on river floodplains and in swamps in what is now western North Dakota. These are the sediments we see exposed today in the Little Missouri Badlands.
In addition to the stream-transported sediments, clouds of volcanic ash, blown eastward from the rising Rocky Mountains during the Laramide Orogeny, collected in layers that were later weathered to clays ( “bentonite”). When wet, the clay absorbs water and swells, and it can become slippery when wet so don’t try walking or driving on it. When the beds dry, they assume a surface texture, similar in appearance and consistency to popcorn, with colors ranging from white to bluish-gray or black.
Why the Badlands Formed
Even though the layers of sedimentary rock exposed in North Dakota’s Little Missouri Badlands range from Cretaceous through Eocene in age (65 to 50 million years old), the badlands themselves–the hills and valleys we see today–are not nearly that old. Before a glacier diverted it, the Little Missouri River flowed northward through a broad, smooth valley, joining the early Yellowstone River in northern Williams County. The Little Missouri and Yellowstone rivers came together near Alamo (about 30 miles north of Williston) in a place now buried beneath 400 feet of glacial deposits. From there, the combined Yellowstone-Little Missouri River flowed northeastward into Canada.
The diversion of the Little Missouri River, away from its route to the north, probably happened sometime prior to the deposition of a volcanic ash bed on the glacial sediment blocking the channel (the ash was deposited as a result of a volcanic eruption in the area of Yellowstone Park 640,000 years ago). It is possible, though, that an earlier glacier might have diverted the river – the 640,000-year figure is a minimum date; erosion of the badlands may have begun as early as 3.5 million years ago.
Since it was diverted by glacial ice, the Little Missouri River has flowed over a shorter and steeper route than it did prior to its diversion. That part of the river’s route today, from the point where it makes its sharp turn toward the east in the area of the North Unit of Theodore Roosevelt National Park, is east rather than north as it had been before a glacier diverted it. When the river assumed its new, shorter route toward the Gulf of Mexico, it began a vigorous erosion cycle, cutting down more rapidly and deeply and sculpting badlands topography. The badlands then, are an indirect result of glacial activity, even though the only conspicuous direct evidence of glaciation remaining in the area is an occasional glacial erratic on the upland in northern McKenzie County.
As you travel through western North Dakota, notice the multicolored layers and brick- or glass-like masses of baked and fused clay, shale, and sandstone. These baked materials, known as clinker, but often referred to locally as “scoria,” formed in areas where seams of lignite coal burned, baking the nearby sediments to a natural brick. Clinker beds range in thickness from a few feet to more than 50 feet in North Dakota, with even thicker beds in Wyoming and Montana.
The first recorded reference to clinker that I know of was by William Clark, who made the following entry in his journal while wintering at Fort Mandan (March 21, 1805):
Saw an emence quantity of Pumice Stone on the sides & feet of the hills and emence beds of Pumice Stone near the Tops of them, with evident marks of the hills having once been on fire. I Collecte Somne of the different sorts i.e. Stone Pumice & a hard earth, and put them into a funace, the hard earth melted and glazed the others two and the hard Clay became a pumice Stone glazed.
When Lewis and Clark arrived at Beulah Bay, about 20 miles west of present-day Riverdale, on April 10, 1805, Lewis noticed a seam of lignite burning along the face of an outcrop. He commented:
“the bluff is now on fire and throws out considerable quantities of smoke which has a strong sulphurious smell.”
On April 16, 1805, Meriwether Lewis wrote the following:
I believe it to be the strata of coal seen in those hills which causes the fire and birnt appearances frequently met with in this quarter. where those birnt appearances are to be seen in the face of the river bluffs, the coal is seldom seen, and when you meet with it in the neaghbourhood of the stratas of birnt earth, the coal appears to be presisely at the same hight, and is nearly of the same thickness, togeter with the sand and a sulphurious substance which usually accompanys it.
Following Lewis and Clark, numerous explorers mentioned seeing clinker as they traveled through the region. They included Larocque (1805), Maximilian (1833), Nicollet and Fremont (1839), and Audubon (1843). Some of these explorers believed the clinker beds had a volcanic origin, but Lewis and Clark were correct in their appraisal that clinker formed as clay and sand were heated by the underlying lignite when it caught fire due to natural causes, such as lightning or prairie fires.
Several early explorers reported seeing coal fires in the northern Great Plains. Over the years, range fires have ignited lignite beds many times. At Buck Hill, in the South Unit of Theodore Roosevelt National Park, a lignite seam burned from 1951 until 1977. During early October, 1976, prairie fires burned over large areas in the southwestern part of the state, igniting underground lignite seams in at least 30 locations over a 7,000-acre area near Amidon. Most of the fires were extinguished before the following spring, but some of them burned for several months. Again, in July, 1988, several lignite seams were ignited by widespread fires in the badlands. Juniper tree roots burning downward from the surface, into the coal, ignited some of the fires.
Burning lignite is limited to depths where adequate air circulates from the surface. The level of the water table may control this depth (burning can’t take place in water-saturated materials). While veins of coal are burning, fumes from the smoldering coal can alter the growth habits of nearby vegetation, causing it to grow in unusual shapes. After the fires go out, the vegetation reverts to its normal shape, common elsewhere in the badlands. Near Amidon, a stand of junipers grew as columnar-shaped trees for many years while a nearby lignite seam burned, producing ethylene gas, which altered the growth habit of the trees. Since the fire went out, the trees have resumed their normal, more bush-like shapes.
Heat from burning lignite beds hardens, melts, or sinters the overlying and surrounding rocks into brick or glass. Sintering is a process that fuses material into a hard mass, without melting it, much like bricks are baked in a kiln. When lignite burns, it may be transformed to an ash bed that takes up only a fraction of the space the lignite did before it burned. Thin layers of white ash, mostly potash, lime, and other inorganic, non-combustible minerals, can sometimes be found at the base of clinker beds.
The baking process oxidizes iron-rich minerals, mainly to red shades, but black, gray, purple, yellow, and other hues are common. The hue and intensity of the colors depends upon the mineral composition, the grain size of the material that was baked, and how hot a temperature was reached during the baking process. The brick-red color, which is most common, is due primarily to the presence of the mineral hematite (iron oxide: the same as common rust). Following a rain shower, wet clinker beds are much brighter in color.
By the time the materials overlying a burning lignite bed cool and collapse, they are hard, and usually partially fused by baking. As they slump, falling into the burned-out space, the baked, melted, and sintered materials may hold together, resulting in a mass that can be as much as 75 percent air space. After the clinker cools, the empty spaces provide convenient living places for small animals, such as rattlesnakes.
Several prominent clinker zones are found throughout the Little Missouri Badlands. The clinker forms a cap on many hills and ridges over extensive areas. Clinker resists erosion because it is harder than unbaked rocks and also because heating and subsidence during the baking process produce fractures that allow water to infiltrate, minimizing surface runoff. Erosion often leaves clinker as a cap on red-topped knobs, ridges, and buttes standing above the more subdued nearby topography developed on less-resistant, unbaked materials. Some widespread areas of clinker are particularly scenic; good examples can be seen along the Red Hills Road south of Sentinel Butte, along the Bennie Pier road in McKenzie County, and on parts of the Scenic Loop Drive in the South Unit of Theodore Roosevelt National Park.