Two distinct sedimentary rock sequences are exposed in the walls of the Grand Canyon (Figure 1). The older, Late Proterozoic sedimentary sequence is comprised of the Grand Canyon Supergroup which consists of the Chuar Group, the Nankoweap Formation, the Unkar Group, and the Sixtymile Formation, and is only found in isolated patches along the main Colorado River corridor and some of its major tributaries (Figure 2). For approximately 500 million years, beginning about 1,250 million years ago (during the Late Proterozoic Era), nearly 13,000 feet of sediments and lava were deposited in coastal and shallow marine environments of an ancient passive continental margin. Then, beginning about 750 million years ago, Basin-and-Range style crustal extension and deformation lifted and tilted these rocks and subsequent erosion removed much of the tilted layers from the Grand Canyon region, leaving only wedge-shaped remnants preserved in large graben structures (Figure 2), mainly observed in the eastern parts of the canyon. A younger sedimentary sequence comprises much of the Paleozoic Era and forms the vast majority of rocks exposed within the canyon and on adjacent plateaus (Figure 1 and Figure 2). These mudstones, sandstones, and limestones are widely distributed in the region, but, because they are essentially undeformed, they total a mere 2,400 and 5,000 feet thick by comparison with Proterozoic rocks. They offer a plethora of evidence interpreted as coastal and marine environments, including several significant marine incursions from the west, developed on a passive continental margin setting between about 550 and 250 million years ago. Rock formations from the Cambrian, Devonian, Mississippian, Pennsylvanian and Permian periods are represented within the sequence.
Mesozoic Era sedimentary rocks do exist regionally, although uplift and erosion during the Early Tertiary removed most of this younger sedimentary sequence from the Grand Canyon area, leaving only small, isolated remnants, particularly southeast of the canyon. Nearby rock outcrops to the north in the Grand Staircase area, and to the east in the Painted Desert and Echo Cliffs, suggest 4,000 to 8,000 feet of Mesozoic sedimentary layers once covered the Grand Canyon region, but were removed by uplift and erosion in the early Tertiary. Cenozoic Era sediments and sedimentary rocks are limited to the western Grand Canyon and to stream terraces and travertine deposits found superimposed on older rocks near the Colorado River itself. Lava flows and associated cinder cones comprise the majority of Cenozoic deposits. Volcanic activity began between about nine and six million years ago and is still ongoing in some areas. One major volcanic field formed on the Shivwits and Uinkaret Plateaus in the northwestern Grand Canyon region, including spectacular lava cascades that poured down the canyon walls like so much candle wax; while similar volcanics formed south of the Grand Canyon in the San Francisco Volcanic Field, an area that may still produce volcanic activity in the future.
The following discussion focuses on the formations comprising the Paleozoic and early Mesozoic sedimentary rocks of the region. Older, Proterozoic age sedimentary rocks of the Grand Canyon Supergroup are covered in another section of this website.
Figure 1. The suite of sedimentary rocks exposed by the downcutting of the Colorado River in Grand Canyon National Park includes an older Proterozoic sequence, and a younger Paleozoic sequence.
Figure 2. A geologic map of the eastern Grand Canyon area indicating the general outcrop locations of Proterozoic crystalline basement and Grand Canyon Super Group sedimentary rocks, sedimentary rocks of the Paleozoic sequence, and geologic structures; note the general juxtaposition of Supergroup rocks against bounding normal faults.
The Paleozoic Sedimentary Rock Sequence
The Tonto Group (by Hannah Slover and Ken Bevis)
The Tonto Group is considered “one of the classic sequences of sedimentary rocks exposed in North America.” Accumulated on a passive continental margin beginning about 525 million years ago, the Tonto Group represents the lowermost sedimentary rocks of the Paleozoic sequence exposed in the Grand Canyon, and is subdivided into three formations, the basal Tapeats Sandstone, middle Bright Angel Shale, and upper Muav Limestone (Figure 1). The rock units are best expressed in the Grand Canyon along a prominent, bench-like surface known as the Tonto Platform formed in the central part of the canyon were the weak Bright Angel Shale has eroded back from the inner gorge (Figure 3). The lower, cliff-forming Tapeats Sandstone is a tan, medium- to coarse-grained quartz-rich sandstone deposited in high-energy coastal plain braided streams, beaches, and intertidal to shallow subtidal wave- and tide-dominated environments (Middleton and Elliot, 2003). The middle Bright Angel Shale Formation is comprised of greenish-gray colored, silty to sandy textured, slope-forming mudrocks (Berthault, 2004; Blakey and Middleton, 2012). These sediments were deposited on shallow near shelf environments and strongly influenced by tidal- and storm wave-generated currents (Middleton and Elliot, 2003). The upper Muav Limestone, tends to transition from olive-green, slope-forming, muddy carbonates, to yellowish-brown, cliff-forming, purer carbonates. It was deposited in subtidal shelf areas dotted with carbonate islands farther offshore (Middleton and Elliot, 2003).
Figure 3. The Grand Canyon’s Tonto Group viewed from the South Kaibab Trail, consists of three formations, upper and lower cliff-formers and a middle slope-former, that crop out along a prominent bench or terrace above the Colorado River’s inner gorge known as the Tonto Platform.
The sedimentary rocks of this group formed on the slowly subsiding, passive margin of western North America that formed after the breakup of the late Proterozoic supercontinent Rodinia (Blakey and Middleton, 2012). Its rock units mark an overall eastward-directed marine transgression of the proto-Pacific Ocean associated with a north-to-south oriented shoreline, but one characterized by multiple transgressive-regressive cycles that resulted in a complex intertonguing of the three formations. This process is nicely described by Blakey and Ranney (2008) as an oft-repeated pattern of “two steps in, one step out, two steps in, again and again” as the sea moved relentlessly eastward over millions of years. Figure 4 displays three paleogeographic maps developed by Ron Blakey that illustrate the depositional settings during accumulation of the Early Middle Cambrian Tapeats Sandstone (Figure 4a), Middle Cambrian Bright Angel Shale (Figure 4b), and Late Middle Cambrian Muav Limestone (Figure 4c), respectively. Overall, Tonto Group deposition reflects the Simple Ideal Model of Fitcher and Poche (2001) (Figure 2 in THE GEOLOGY OF SEDIMENTARY ROCKS). Sediments accumulated during flooding of the passive continental margin of ancient North America during a major marine transgression towards the east, depositing a typical three-fold, onlapping, fining-upward sedimentary rock sequence comprised of shoreline quartz sandstones, near shelf mudstones, and far shelf carbonates (Figure 15 in THE GEOLOGY OF SEDIMENTARY ROCKS).
Figure 4. Cambrian paleogeography of western North American during deposition of the Tonto Group’s Tapeats Sandstone (a), Bright Angel Shale (b), and Muav Limestone (c); original maps are by Ron Blakey and presented in Blakey and Ranney (2008) and Blakey and Middleton (2012).
The Cambrian Tonto Group directly overlies the Great Unconformity, with the 525 million year old Tapeats Sandstone deposited directly on tilted, 1.4 to 1.1 billion year old Grand Canyon Supergroup rocks in the eastern canyon, forming an angular unconformity; and 1.8 to 1.6 billion year old crystalline basement rocks of the Grand Canyon Metamorphic Suite in the west, forming a nonconformity (Middleton and Elliot, 2003; Blakey and Ranney, 2008). This lower contact is a major unconformity developed by erosion during extensive and prolonged subaerial exposure of western North America’s passive margin that probably required substantial uplift brought on by mountain building. In some locations within the canyon, a thick paleosol occurs at the top of Precambrian surface suggestingsignificant chemical weathering of the basement rocks occurred prior to the beginning of Tapeats deposition (Middleton and Elliot, 2003). Wave erosion associated with initial Cambrian transgression may have removed the paleosol in areas where it is absent, and this weathered material provided a ready source of sediment as the sea moved eastward. The surface of the unconformity is very irregular. Lowlands, developed by weathering and erosion of weaker rocks, alternate with bedrock “hills” formed of more resistant material (Middleton and Elliot, 2003; Blakey and Ranney, 2008). The tips of the “hills” are often capped with the Bright Angel Shale rather than the basal Tapeats Sandstone which pinches out against the resistant knobs. Apparently, as sea level rose during the Cambrian transgression, these “hills” formed islands in the advancing sea, but were not inundated until deeper water conditions had ensued later during Bright Angel deposition.
The Tapeats Sandstone forms a conspicuously layered, pervasive, cliff-forming, medium- to coarse-grained sandstone weathered buff to brown in color deposited unconformably above Proterozoic basement rocks; it is most easily recognized where it rests on vertically foliated rocks of the Grand Canyon Metamorphic Suite such as that observed along the Tonto Trail west of Travertine Canyon (Figure 5). The rock unit was named for exposures along Tapeats Creek in the west-central part of the Grand Canyon (Middleton and Elliot, 2003). The sandstone contains abundant quartz and lesser amounts of feldspar (Berthault, 2004), although as the formation progresses upward, the feldspar percentage, the bedding thickness, and the coarseness of the sandstone decrease (Middleton and Elliot, 2003), probably related to less direct input of source material from the crystalline basement below. Overall, the formation ranges from 100 to 325 feet thick, reaching a maximum of 393 feet in Bass Canyon, although its thickness is clearly controlled by the relief of the unconformity it rests on. The base of the formation is most often composed of pebble conglomerates deposited in environments best characterized as “braided streams merging with nearshore marine areas” (Blakey and Middleton, 2012). The composition of the basal zone generally reflects the mineralogy of the underlying Precambrian rocks (Middleton and Elliot, 2003), with highest quartz and feldspar content near areas with abundant Zoroaster Granite.
Figure 5. The Tapeats Sandstone overlies Precambrian basement, in this case crystalline rocks of the Grand Canyon Metamorphic Suite; here viewed along the inner gorge of the Colorado River from the Tonto Trail between Travertine and Boucher Canyons.
The majority of the Tapeats Sandstone forms a distinct cliff characterized by relatively thick bedding (individual beds are up to three feet thick) that exhibits planar and trough cross-stratification and poor horizontal stratification. The upper section tends to be weathered back into a stepped-ledge slope of thin interbeds of fine- to medium-grained lensoidal sandstone and interbeds of mudstone accompanied by trough and ripple cross laminations (Middleton and Elliot, 2003). The transition from Tapeats Sandstone into the Bright Angel Shale forms an intercalated, gradational contact and marks a significant facies transition to lower energy conditions (Middleton and Elliot, 2003 and Berthault, 2004). Body fossils are very rare in the Tapeats Sandstone, but more common in the finer-grained sediments near the top of the unit (Blakey and Middleton, 2012). Fossils have been used to date the Tapeats as late Early Cambrian in the west and early Middle Cambrian in the east, indicating its time-transgressive nature. According to Blakey and Middleton (2012), the Tapeats was deposited along a shoreline characterized by numerous embayments where fluvial deposits derived from braided streams draining westward off the Transcontinental Arch merged with a shallow, high-energy marine environment. It was likely formed in beach, intertidal barrier bar or shoaling environments, or shallow, subtidal zones influenced by tidal currents, although wave processes probably dominated in areas adjacent to resistant, island-forming “islands.”
The Bright Angel Shale is fine-grained and dominated by greenish muds composed mostly of clay (Middleton and Elliot, 2003). The mudstone is accompanied by minor sandy dolomites and silty limestones, and somewhat more common green siltstones and sandstones containing glauconite and dark brown ironstone (Berthault, 2004). The mineral glauconite, common to clastic shelf environments, is responsible for the pervasive green coloration of the rock unit. The formation’s abundance of muds provides a relatively high susceptibility to weathering and erosion and is responsible for its expansive, slope-forming topographic expression. Its name derives from the broad slope of interbedded shale, siltstone, and sandstone exposed just above the Tonto Platform near Bright Angel Creek. The view expressed in Figure 4 shows this slope east of Bright Angel Canyon quit well, although a closer view of the unit is offered by Figure 6, where the formation is exposed along Cottonwood Canyon on the Grandview Loop Trail.
Figure 6. The weak, slope-forming middle unit, and the stronger, cliff-forming upper unit of the Tonto Group, the Bright Angel Shale and Muav Limestone respectively, are nicely exposed in the walls of Cottonwood Canyon as viewed from the Tonto Trail as in begins to round Horseshoe Mesa.
The average thickness of the formation is 320 to 400 feet; however, it thickens to over 450 feet in the western Grand Canyon, thinning to 325 feet at Bright Angel Creek in the central canyon, but reaching only 270 feet thick in the eastern part of the canyon. This complex variation in thickness is due to the time-transgressive nature of the unit; the near shelf conditions that caused its deposition persisted much longer to the west than to the east. Only one member is identified in the Bright Angel Shale. Located in the uppermost portion of the unit in its western canyon exposures, it is comprised of shale, siltstone, and limestone; although as the member progresses to the east, the limestone decreases, and the member becomes entirely shale (consistent with an eastward-directed transgression). According to Middleton and Elliot (2003), the sedimentary structures identified in the formation include “horizontal laminations, small- to large-scale planar tabular and trough cross-stratification, and wavy and lenticular bedding.” They also identified coarsening- and fining-upward sequences in the central Grand Canyon. Trace fossils and body fossils are abundant in the Bright Angel Shale, and Blakey and Middleton (2012) report that the fossils record a “major proliferation of invertebrate fauna.” The fossils place the western edge of the Bright Angel Shale during late Early Cambrian and the eastern side as Middle Cambrian (Middleton and Elliot, 2003), again attesting to its time-transgressive nature. The deposition of the Bright Angel Shale occurred in deeper water, under the lower energy currents of the subtidal shelf environment influenced by both calm weather and storms, and was likely deposited in lagoons or estuaries influenced by tidal action (Blakey and Middleton, 2012).
The Muav Limestone, generally a cliff-forming unit, is the uppermost formation of the Tonto Group (Figure 3 and Figure 6). Blakey and Middleton (2012) report that the formation “records the first widespread accumulation of carbonate sediments ….. in the Paleozoic.” It is composed of yellowish-brown to yellowish-green, silty and sandy, dominantly impure limestone, though the purity improves to the west (Berthault, 2004). Its dominant color provides its namesake. It is mainly calcareous mudstone nearer its base where it tends to form ledgy slopes, and its contact with the Bright Angel Shale is also intercalated and gradational, reflecting intertonguing of adjacent environments related to small-scale, short-duration, repeated transgressions and regressions probably during the highstand of the Cambrian Sea. Its limestones, which become more abundant upward within typical cliff-forming portion of the formation, are chiefly packestone, accompanied by intraformational, flat-pebble (intraclastic) conglomerates, and with minor interbeds of shale, siltstone, fine-grained sandstone, and silty limestone (Middleton and Elliot, 2003), suggesting deposition associated with clear, shallow-marine shoals and platforms (Blakey and Middleton, 2012) occasionally disturbed by storms.
In general, the siliciclastic content of the rock unit increases eastward as the carbonates decrease, in accordance with an eastward-directed transgression. Bedding alternates from thin to thick with small, irregular clay patches, and overall, bedding thickens to the west. The thin beds tend to be composed of micaceous shale and siltstone with a minor amount of sandstone and limestone. Beds are typically structureless or horizontally laminated. Also marking the bedding are small-scale trough, planar tabular, and low angle cross-stratification, as well as fenestral fabrics and desiccation cracks.
Mirroring the bedding is the overall thickness of the formation (Middleton and Elliot, 2003). It is 827 feet thick in the west at the Grand Wash Cliffs, 439 feet thick at the Tuweap Valley in the central canyon, and only 136 feet thick in eastern Grand Canyon. Within the Muav Limestone, seven members can be identified and separated on the basis of fauna, or prominent conglomerate, shale, or limestone layers. The lower four members are only found in the western canyon, but the upper three members spread throughout its entirety. Trace fossils are present in the Muav Limestone, but are not as abundant as they are in the Bright Angel Shale; however, according to Blakey and Middleton (2012), the Muav contains “the first major occurrences of invertebrate fossils”. The western section of the Muav has been assigned to the Middle Cambrian Period and the eastern section to the late Middle Cambrian Period (Middleton and Elliot, 2003), again paralleling the time-transgressive age trend of the two lower units in the Tonto Group.
In the western section of the Grand Canyon, dolostones overlie the Muav Limestone to a thickness of over 400 feet. It is thought that these rocks are Upper Cambrian, though there is no paleontological evidence to support this assumption. Geologists refer to it as the Cambrian Undifferentiated or the Grand Wash Dolomite. Three lithofacies are present within this dolomite sequence. Middleton and Elliot (2003) describe them as “white-to-buff massive dolomite,” “white-to-yellow, very fine-grained, thick-bedded dolomite,” and “gray, fine-grained, thick-bedded dolomite.” Oolitic grainstones and stromatolitic beds lie between these layers. The sedimentary structures include wavy and asymmetric ripple lamination and small-scale cross-stratification, and the dominant trace fossil is horizontal burrows and tracks (Middleton and Elliot, 2003).
Brachiopods and trilobites, though poorly preserved, are the most common invertebrates found in the Tonto Group (Middleton and Elliot, 2003). Brachiopods are most often found in the coarse-grained sandstone of the Bright Angel Shale and the mixed siliciclastics and carbonates of the Muav Limestone; while several dozen species of trilobites have been identified within the Tonto Group, congregating in the coarse-grained sandstone of the Bright Angel Shale and the mudstones of the Bright Angle Shale and the Muav Limestone. Fragmented sponges, primitive mollusks and echinoderms, and algae occur in the Bright Angel Shale and Muav Limestone, but they are not abundant. Trace fossils in the form of tracks, trails, and burrows are evident in all three formations of the Tonto Group, with the majority in the Bright Angel Shale and the least amount in the Muav Limestone (Middleton and Elliot, 2003). They become increasingly abundant in the upper half of the Tapeats Sandstone as it transitions into the Bright Angel Shale, with single and paired vertical tubes and horizontal traces quit common. The “unbranched, straight vertical burrows” created by Skolithos in the fine-to-coarse-grained and cross-bedded sandstone provide evidence of currents capable of active bedload transportation. U-shaped burrows, perpendicular to the bedding in fine- to coarse-grained sandstone of the Bright Angel Shale and Tapeats Sandstone have lead geologists to infer shallow-water deposition. Finally, within the Bright Angel Shale and the Tapeats Sandstone, on interbedded sandstone and mudstones are trilobite crawling tracks, referred to as Cruziana, likely created during a time of fair weather, while the resting tracks, referred to as Rusophycus, are suggestive of storm conditions (Middleton and Elliot, 2003).
The Tapeats Sandstone was deposited on sandy beaches, intertidal sand and mud flats, and shallow, subtidal sand wave complexes (Figure 4a). Sedimentation below the beach zone, beyond the coast, occurred in water depths less than 100 feet. The consistent dip directions of the cross-stratification provide evidence that the islands dotting the coastline had minor influence on sedimentation, simply being a source of sediment themselves. Tidal flat mud deposition ranged from high to low and was intermixed with evidence of sandy tidal channels. At the base of the Tapeats, less mature, coarse-grained sandstone and conglomerates alternating with planar-tabular and trough cross-stratified sandstone fill shallow channels and imply braided stream deposits which grade into marine facies. Large channels occurring near the top of the Tapeats Sandstone exhibit internal structure suggestive of formation by subtidal channel complexes dominated by the ebb tidal phase.
The Bright Angel Shale was deposited on open shelf environments, in water depths between the Tapeats Sandstone and the Muav Limestone, but generally below fair-weather wave base (Figure 4b). A combination of alternating upward-coarsening, upward-fining, cross-bedded sequences, and interbedded sandstone and mudstone indicate subtidal deposition manipulated by tidal currents and storm wave processes. Upward-coarsening sequences, up to 25 feet thick and several tens of kilometers wide, record deposition in a solely subtidal environment though the upper portion of many were in shallow, intertidal waters. The lower, finer sections contain laminated, bioturbated mudstones probably deposited during fair weather. The upper, coarser sections exhibit thick, planar tabular cross-stratification, often with abrupt changes in the dip of foresets that likely formed by migration of sand waves, dunes, and ripples associated with stronger tidal and storm-wave generated currents. Fining-upward sequences typically overlie an erosive base. They begin with pebble conglomerate or sandstone and grade into interbedded fine-grained sandstone and mudstone. The coarse basal layer is representative of high-energy deposition during storm-induced currents capable of transporting coarse material. Symmetric ripples and a high amount of laminated mudstone occur at the top of these sequences. They are marked by a diverse array of trace fossils representative of post-storm resettling of muds and sands.
The Muav Limestone, like the Bright Angel Shale, was also deposited in a subtidal environment (Figure 4c). East of Bass Canyon, intraformational, often imbricated, flat-pebble conglomerates consisting of disc-like clasts of mostly micrite mark an extremely important facies within this rock unit. The intraformational conglomerates occur as both discontinuous lenses and large flat sheets extending for miles, and likely indicate subtidal deposits created during temporary regressions. In addition to a subtidal environment, some limestone and dolostone beds are thought to have been deposited in intertidal and supratidal flats. The laminated dolostones of this formation deposited in the western Grand Canyon area may have been limey sediments bound by tidal flat algal mats. The undifferentiated dolomites found in the western part of the canyon containing thick beds of oolitic grainstones and stromatolites interbedded with fine-grained carbonates are probably representative of shallow subtidal and perhaps intertidal environments, again related to regression.
There were many influences on the sedimentation of the Tonto Group and deposition occurred in a variety of interrelated environments associated with an overall eastward directed marine transgression of a north-south oriented coastline (Figure 4). Complex intertonguing of the rock units was generated in response to multiple regressive phases during dominant transgression onto North America’s passive margin. This passive continental marginal suite records the classic onlapping, fining-upward, sedimentary sequence of the Simple Ideal Model (Figure 2 in THE GEOLOGY OF SEDIMENTARY ROCKS). The Cambrian Tonto Group contains the typical, stacked, beach and nearshore generated sandstone (Tapeats Sandstone), near shelf generated mudstone (Bright Angel Shale), and far shelf generated carbonate facies (Muav Limestone) produced during a marine transgression.
The Tonto Group is geologically unique in several ways. Its rock units are well exposed and readily studied. Not only do they demonstrate a classic marine transgressive sequence, obeying Walther’s Law of Facies and the Simple Ideal Model, but they also describe the time-progressive nature of many formations; that is, the sediments they contain are not the same age everywhere, but change in age as the depositional environment they are associated with migrates. In this case, each formation becomes younger toward the east as the marine transgression gradually inundated the passive continental margin from west to east. And of course, these rocks wouldn’t have been preserved if not for the self-generating nature of passive continental marginal tectonic settings, creating accommodation space for the sediments to accumulate through subsidence (a mere rise in sea level could not account for the thickness of material preserved). Finally, these Middle Cambrian rock units also portray the amazing proliferation of life on planet earth at this time, located in the middle of a very short (geologically speaking), 150 million year time stream in which life went from nothing more advanced than single-celled organisms in the late Proterozoic, to evolve the most complex phyla found on earth today by the Late Ordovician (placing this is context, it took 3,000 million years of earth history to create single-celled organisms).
The Temple Butte Formation (by Hannah Slover and Ken Bevis)
The Temple Butte Formation is a westward thickening carbonate layer in the Grand Canyon deposited during the Late Middle to early Late Devonian Period roughly 385 million years ago (Beus, 2003). The rock unit is named for exposures on Temple Butte in the eastern Grand Canyon. It forms a relatively inconspicuous layer at the eastern end of the canyon where it crops out only as discontinuous lenses of paleo-stream channel fill (Figure 7), and at the western end where it forms a thicker, continuous layer, its relative resistance causes it to merge with the overlying Redwall Limestone cliffs overlying it. Extensive dolomitization of these carbonates makes it difficult to interpret a depositional environment, but a few fossils are present to aid interpretations. In the east, it is likely that the formation was deposited in narrow tidal channels, appearing as discontinuous lenses, while the central and western outcrops were deposited in widespread, shallow, subtidal, marine conditions, resulting in a thicker and more continuous layer. Blakely and Ranney (2008) suggest that the channels were marine estuaries extending finger-like eastward onto the continental margin and coalescing westward into a contiguous shallow marine shelf environment.
Figure 7. The Temple Butte Formation filling the paleochannel of an ancient stream cut into the Muav Limestone in Saddle Canyon (a side tributary of the Colorado River at mile 47 in Marble Canyon).; photo by Garry Hayes from his Geotripper blog, posted August 24th, 2013).
The Temple Butte’s upper and lower contacts are unconformable with the Redwall Limestone and Muav Limestone respectively, although the lower unconformity represents a much more significant time gap. Absent between the Muav Limestone and Temple Butte Formation is the Late Cambrian, all of the Ordovician and Silurian, and most of Early and Middle Devonian sediments, a hiatus of about 120 million years, but the gap between the Temple Butte and Redwall is only upper Devonian, a mere 45 million years at most. Both gaps in the rock record are disconformities; that is, the rock units above and below the erosion surface are both sedimentary and are essentially undeformed. In the west (and also near Lava Canyon in the east) light-gray strata of undifferentiated Cambrian dolomite up to 500 feet thick conformably overlie the Muav Limestone (Beus, 2003). Overall, except for their appearance at Lava Canyon where they have been reported at a thickness of 163 feet, the strata are truncated eastward by the Temple Butte. These unnamed layers do appear to conform with deposition of the Muav Limestone, and Korolev and Rowland (1993) have correlated the undifferentiated dolomite layer in the Grand Canyon to the Banded Mountain Member of the Bonanza King Formation of Nevada, which is thought to be of late Middle Cambrian age. Where the Temple Butte is in direct contact with the Muav Limestone, the unconformable surface is identified by the considerable relief created by channels and depressions up to 100 feet deep. However, there are areas where the channels are absent. The result is a planar contact of gray dolomite from the Temple Butte and the Muav Limestone, making the unconformity difficult to locate even though it represents more than 100 million years of missing time (Beus, 2003).
The southwestern North American plate during the Late Devonian was tectonically quiet, essentially retaining all the features of earlier formed passive continental margin setting (Figure 8); although the Antler volcanic arc was approaching from the west, plowing into slope-rise deposits off the continent’s edge (Blakely, 2014). As the arc encroached eastward, it began to collide with the passive margin of western North America, thrusting these deep water deposits onto the continental shelf and causing regional subsidence that may have aided deposition of the Temple Butte Formation. Overall, the westward thickening sediments of the Devonian Temple Butte are representative of a marine transgression to the east onto a submerged continental shelf (Beus, 2003; Blakey and Middleton, 2012).
Figure 8. Middle Devonian paleogeography of western North America during deposition of the Temple Butte Formation; original map is by Ron Blakey and presented in Blakey and Ranney (2008) and Blakey and Middleton (2012).
The Temple Butte is composed of gray and tan limestone and dolostone with minor traces of siliciclastics, but is often weathered to a purplish color. It occurs as thin, patchy layers, deposited in “deep, channelized surfaces” in the east but is “more continuous” in the west (Blakey and Middleton, 2012). The discontinuous lenses in the east are usually less than 100 feet in thickness, but may reach up to 400 feet in width and are “eroded into the upper surface of the Muav Limestone or the Cambrian undifferentiated dolomite” (Beus, 2003). To the west of Hermit Creek, the formation (composed of continuous dolomite) thickens to more than 450 feet (Beus, 2003).
The basal channel-fill sections of the Temple Butte are usually composed of dolomite or sandy dolomite of a distinct pale, reddish purple color, creating irregular, or even gnarly, beds (Beus, 2003). Basal conglomerates composed of subrounded dolomite pebbles occur, but are very rare. Variations occur between these beds; horizontal in some localities, or conforming to the walls of the channel in others, but when the strata do conform to the curved channel margins, the layers above often truncate them with angular discordance. The upper section of the channel fills is often composed of about twenty percent insoluble residue and an interesting columnar alternation pattern of pale gray and purple dolomite probably caused by weathering. Bedding above these channels is more continuous and can be described as thick and blocky ledges. It is a “fine to medium crystalline dolomite” with a “dark-to-light olive gray” color and often “weathers to a sugary texture” (Beus, 2003). Another, less common, but extensively deposited version of the carbonate in the Temple Butte formation occurs as very fine crystalline dolomite with a grayish-orange-pink color. It is thinly bedded, has evidence of concoidal fracture, often seems porcelaneous, and it weathers to an extremely light gray or yellow. Within this layer of aphanitic dolomite, laminae or thin lenses of rounded, frosted quartz sand occur; although locally, lenses may reach up to 7 feet thick. Most of the dolomite beds form steep, receding ledges with the exception of the central Grand Canyon where they conform to the unbroken, 1550-foot carbonate wall comprised of the Redwall and Muav Limestone. While fossils are abundant in other Devonian formations, few are identifiable in the Temple Butte. Brachiopods, gastropods, corals, fish plates, and trace fossils are all present, but conodont microfossils are the most helpful in determining the age to be of latest Middle Devonian to early Late Devonian (Beus, 2003).
Temple Butte strata truncate older rocks from west to east across the Grand Canyon region. The disconformity at the base of this formation marks a significant stratigraphic break in the Grand Canyon’s Paleozoic sedimentary sequence, representing the latest Cambrian, the entirety of Ordovician and Silurian, and Early through Middle Devonian time. Sediments of the upper Muav Limestone were deposited during eastward-directed marine transgression of the Middle Cambrian period while Temple Butte sediments mark a much later transgression of the Middle Devonian Sea, making the unconformity greater in the east because subaerial exposure would have dominated this region longer. In general, the deposition of the Temple Butte Formation occurred on a passive continental margin. Extensive dolomization and the paucity of recognizable fossils make it difficult to pinpoint exact depositional environments. The Temple Butte of the central and western Grand Canyon seems to have accumulated in shallow, subtidal, open marine conditions which remained more or less static over time despite overall transgression. The aphanitic limestone is more specifically considered to have been deposited in supratidal conditions. Discontinuous lensoidal patches of Temple Butte in the eastern Grand Canyon were formed in tidal channels associated with an intertidal environment. These deposits thicken and coalesce to the west where marine conditions became more contiguous. A comprehensive paleoenvironmental interpretation of Temple Butte deposition is a system of estuary channels in the eastern Grand Canyon area that merged with a more pervasive, gradually subsiding, shallow marine shelf to the west.
The Redwall and Surprise Canyon Formations (by Hannah Slover and Ken Bevis)
The Mississippian Period produced two carbonate dominated formations, the Redwall Limestone and Surprise Canyon Formation. Its basal unit, the Redwall Limestone, forms massive, steep cliffs and is stained red by the iron oxides weathered from the overlying Hermit Formation and Supai Group rocks, and is probably the most easily identified rock unit in the Grand Canyon. Wedged between the Redwall Limestone and the overlying rocks of the Supai Group is found the inconspicuous, discontinuous patches of limestones, siltstones, sandstones, and conglomerates, often high in fossil content, that are now recognized as belonging to the Surprise Canyon Formation (Blakey and Middleton, 2012). Both formations accumulated in shallow-marine passive-marginal environments (Beus, 2003), although of a very different nature and extent (Figure 9). The Redwall Limestone accumulated during the Early to early Late Mississippian associated with two major transgressive-regressive cycles on open-marine continental shelf carbonate platforms (Figure 9a). It is broken down into four member subunits (Beus, 2003). The basal layer is referred to as the Whitmore Wash member and was deposited during the first transgression. This layer is followed by the Thunder Springs member, formed during the subsequent regression. The latter two layers, the Mooney Falls and Horseshoe Mesa members, accumulated during the second transgressive-regressive cycle. The Surprise Canyon Formation was deposited during the Late Mississippian (and Early Pennsylvanian) Periods, in an array of discontinuous karst-generated sinkholes and caves and in the paleovalleys of ancient dendritic drainage systems eroded into the limestone of the Redwall after marine regression and subaerial exposure of much of the carbonate-dominated shelf (Figure 9b). The Surprise Canyon’s limited exposure makes it undeniably the least easily identified Paleozoic rock unit in the Grand Canyon. The variety and distribution of this unit’s deposits suggest a transitional environment from fluvial to estuarine conditions associated with marine transgression onto a karstified coastal plain dissected by stream valleys.
Figure 9. Middle and Late Mississippian paleogeography of western North America during deposition of the Redwall Limestone and Surprise Canyon Formation, respectively; original maps are by Ron Blakey and presented in Blakey and Ranney (2008) and Blakey and Middleton (2012).
The passive continental margin of the Mississippian North American west coast was generally unaffected by tectonism. However, the convergence of an oceanic plate with the North American plate in latest Devonian-Mississippian time resulted in the Antler Orogeny to the west in central Nevada (Beus, 2003). The colliding Antler volcanic arc shoved strata deposited in deep water over the continental shelf to form the Antler highlands (Blakely, 2014). The weight of the thickened crust caused eastern Nevada and western Utah to subside in a foreland basin and thick Upper Devonian and Lower Mississippian strata formed there. Farther east, thinner deposits accumulated on a broad carbonate shelf over much of the Western Interior (Figure 9a). The Grand Canyon sat on the periphery of orogenic activities, and was affected by them only indirectly (Beus, 2003). A broad, but gentle forearc bulge probably enhanced subaerial exposure in the Late Mississippian and helped induce channel incision on the karstified landscape (Figure 9b). Localized uplift also occurred as faults shifted the crust due to the contractions associated with the Antler Orogeny. For example, following the deposition of the Redwall Limestone and prior to the Surprise Canyon Formation, an anticline caused by east-west compression and reactivation of a buried fault was responsible for the truncation of the Horseshoe Mesa member of the Redwall Limestone in the Tanner Trail area of the eastern Grand Canyon.
In general, the Redwall forms soaring, mottled red cliffs spanning five- to eight-hundred feet in height, probably the most recognizable rock unit in the Grand Canyon (Figure 10). It is tallest in the west at roughly 800 feet at Iceberg Ridge, five miles west of the mouth of the Grand Canyon, and just over 400 feet in the easternmost Grand Canyon near the cliffs forming the Palisades of the Desert. The Redwall Limestone is deposited unconformably on Devonian sediments or, where the Temple Butte Formation is absent, on Cambrian sediments of the Muav Formation. Missing between these formations is latest Devonian and earliest Mississippian time. In the western Grand Canyon, the basal beds of the Redwall are of Early Mississippian and its abrupt contact with Devonian sediments is well marked by an erosion surface with up to 10 feet of relief. The more substantial part of the unconformity occurs in the east where the sediments are of the late Early Mississippian Period. The time difference between the basal sediments from west and the east is due to the eastward directed transgression during deposition of the unit. Locally, the unconformity is typically bordered by a conglomerate of angular dolomite or limestone clasts derived from the Temple Butte Formation, but becomes very difficult to locate when the strata on either side of the unconformity are dolomitic. These factors suggest a gentle uplift paired with mild erosion of the strata before deposition of the Mississippian Redwall began (Beus, 2003).
Figure 10. The Redwall Limestone literally forms a massive “red wall” spanning hundreds of feet in height broken only by the occasional stream drainage, one of the most easily recognized Paleozoic rock units in the Grand Canyon; in this view from Yavapai Point looking northeast, in the foreground, the Redwall forms the foundation of Cedar Ridge, and across the canyon in the background, it can readily be observed just above the gray slopes of the Tonto Platform.
Grand Canyon geologist Edwin McKee divided the Redwall into four members (Beus, 2003). He named the basal late Early Mississippian layer “Whitmore Wash” for its type section at that location. Composed almost entirely of pure carbonates, often “pelleted and locally skeletal or oolitic wackestones and packstones,” it is also accompanied by iron oxides and less than 2 percent of gypsum. It is deposited in thick beds, ranging 2 to 4 feet thick, and is dominated by fine-grained limestone in the west and dolomite in central and eastern Grand Canyon. The Whitmore Wash unit spans an area from about five miles beyond the western edge of the Grand Canyon at a thickness of approximately 200 feet to Iceberg Ridge in the east, thinning to around 100 feet. Due to extensive dolomitization, fossils are rare in the Whitmore Wash member. Its upper contact is conformable and easily distinguished from the alternating dark chert and light carbonate beds of the Thunder Springs member. The initial deposition of the Whitmore Wash began as a major transgression resulted in oolitic shoals created by high energy currents in nearshore, shallow subtidal conditions, and were followed by continuous offshore deposition of skeletal grainstones and packstones in quieter water (Beus, 2003).
Following the transgression of the Whitmore Wash was the regression that deposited the Thunder Springs member of the Redwall Limestone in shallower marine conditions. This second unit is composed of alternating thin and dark reddish brown or gray chert beds with light gray carbonate beds. In the west, the carbonate bed is a more pure limestone, while in the east it is more dolomitic. Beus (2003) reports that the chert layers are “silicified blue-green algal mats.” Invertebrate marine fossils are very abundant in the chert beds and include corals, bryozoans, brachiopods, and crinoids. While fossils still appear in the carbonate beds, they are poorly preserved because of the dolomization. The member thickens from 100 feet in the east to about 150 feet in the west. Overlying the Thunder Springs is the Mooney Falls member, along an unconformable contact except in the extreme west. It is a low-angle disconformity suggesting only minor structural activity and erosion (Beus, 2003).
Like the previous two members of the Redwall Limestone, the third and thickest member, Mooney Falls, also thickens from east to west, ranging from around 200 to 400 feet. This member forms the sheer wall plastered in oxidized muds that inspired its name. It is characterized by pure limestone with thin beds or lenses of chert towards the top, but is locally dolomitized. The limestone is thickly bedded, micritic, and accompanied by carbonate grains including oolites, pellets, and skeletal fragments. Invertebrate marine fossils are also abundant in the Mooney Falls member and include solitary and colonial corals, spiriferid brachiopods, and crinoids. The top third of the member in the central and eastern Grand Canyon is also characterized by large-scale, tabular-planar cross-bedding. The second marine transgression was responsible for the deposition of the Mooney Falls in more open marine waters (Beus, 2003). The upper contact of the Mooney Falls member with the overlying Horseshoe Mesa member is conformable and difficult to locate. Geologists identify the contact where the limestone transitions from thickly-bedded medium and coarse limestone to the aphanitic, thin-bedded, ledge-forming limestone of the Horseshoe Mesa (Beus, 2003).
The final and youngest member of the Redwall Limestone is the Horseshoe Mesa member of the early Late Mississippian Period. Also thinnest in the east, it ranges from 45 to 125 feet thick, and wedges out 30 to 40 miles south of the Grand Canyon due to erosion. In most of central Arizona, it is absent from the top of the Mooney Falls member. Composed of thin layers of light gray limestone, it has a texture ranging from mudstone to wackestone, and forms weak, receding ledges. Similar to the underlying Mooney Falls are chert lenses in the lower portion of the Horseshoe Mesa member. Invertebrate fossils are rare, but well preserved, with at least sixteen species of foraminifers identified. As the final regression occurred during the deposition of the Horseshoe Mesa member, the water became increasingly shallow resulting in restricted circulation. The upper contact of the Horseshoe Mesa member with the Watahomigi Formation of the Supai Group forms a major unconformity; locally the Horseshoe Mesa lies in contact with patchy lenses of the Surprise Canyon Formation along a lesser unconformable surface.
The Late Mississippian Surprise Canyon Formation occurs in isolated patches between the underlying Redwall Limestone and the overlying Watahomigi Formation of the Supai Group. It filled the caves and sinkhole depressions of a karstified limestone landscape and the erosional valleys of dendritic river systems with clastic and carbonate sediments (Figure 11a). Sediment-filled sinkholes and caves often had high porosity relative to the surrounding solid limestone and often served as conduits for mineralized fluids. Indurated breccia pipes containing uranium and copper ores were sometimes the result (Figure 11b). These patches occur with V- or gentle U-shaped cross-sections throughout the Grand Canyon and Marble Canyon to the east. Patches in the western Grand Canyon are typically 150 to 200 feet deep and up to half a mile wide, but in the central and eastern Grand Canyon and Marble Canyon, they are shallower and wider, usually reaching a thickness of only 50 feet. The largest paleo-river valley occurs in Quartermaster Canyon, a south rim tributary in western Grand Canyon, with a thickness of 400 feet. The dendritic drainage system flowed east to west, merging to form more continuous patches and fewer, larger channels to the west, where it met with the paleocoastline to form an estuarine environment. The depth of paleovalley incision suggests subaerial exposure associated with gentle uplift or a sea level drop by as much as 100 feet (Beus 2003). Blakey and Middleton (2012) indicate that much of the North American continent was gently uplifted in the Late Mississippian. Streams flowing into the marine trough formed by the still extant Antler foreland basin carved westward-deepening valleys, and these were subsequently flooded to form coastal estuaries and then filled by marine deposits to generate the Surprise Canyon Formation (Figure 9b).
Figure 11. A sinkhole formed in the top of the Redwall Limestone on the south side of Cottonwood Canyon, filled with deposits of the Surprise Canyon Formation (A); because these sinkhole deposits filled with brecciated rock, they had high porosity relative to the surrounding solid limestone and often served as conduits of uranium and copper ore mineralization such as at Horseshoe Mesa (B).
According to Beus (2003), the Surprise Canyon Formation “has the most varied sedimentary lithology of any Paleozoic formation in Grand Canyon.” It is laterally divided into eastern and western units; where in the east, the stratigraphy merges into a single red-brown, slope-forming conglomeratic sandstone or siltstone, while in the western and central Grand Canyon, the Surprise Canyon is divided into three more distinct units and each unit has multiple lithofacies. The lower western unit is a fining-upward conglomeratic sandstone of fluvial origin, forming cliffs and slopes. Beus (2003) describes this basal unit as a “ferruginous pebble-to-pebble and local boulder conglomerate.” Clasts are also present and dominated by chert, with minor amounts of limestone derived from the underlying Redwall, which is enclosed in a quartz matrix (with some hematite). The presence of imbricated cobbles provides evidence for a unidirectional, westward flowing, fluvial current. Conglomerates grade upward into quartz sand or siltstone or carbonaceous shale. The sandstone facies is “yellow to dark reddish brown or purple,” flat-lying, characterized by trough cross-strata or ripple lamination, and contain Lepidodentron log impressions (Beus, 2003). The shale facies is dark and plant fossils have been found in the unit. Beus (2003) reports that this lower unit preserves “a record of continental and fluvial conditions changing to intertidal conditions as the sea transgressed eastward into the estuary and began to trap and rework clastic sediments within the tidal range.”
The middle western unit of the Surprise Canyon Formation is a cliff-forming, skeletal limestone with a coarse grainstone texture and whole or fragmented shells of marine origin. The limestone occurs in thin beds with thicknesses up to 4 inches and alternates with thinner quartz sandstone beds. Due to erosion, this middle unit often truncates the lower unit, and after weathering has a yellowish brown, rusty or purple gray color. While the lower unit has evidence for a unidirectional current, the small-scale trough-cross strata of the middle unit suggest a bimodal current with the dominant direction upstream, or to the east (Grover, 1989). Paired with the abundance of marine invertebrate fossils, the bimodal current is suggestive of an estuary environment dominated by tides. This unit is absent east of the Grand Canyon’s Fossil Bay area (Beus, 2003). The upper western unit is a marine siltstone and silty or sandy limestone, and forms weak slopes to receding ledges. It is dark red-brown to purple in color, and is planar bedded to ripple laminated. Nearly spherical algal stromatolites occur in some portions of the unit, but the relative absence of marine fauna suggest a restricted marine to upper tidal-flat environment as the karstified topography of the Redwall Limestone is finally filled in (Beus, 2003).
The fossil record of the Surprise Canyon is extremely rich and diverse. It contains more than 60 species of marine invertebrates with brachiopods being the most common form. In the middle unit, microfossil invertebrates are “moderately abundant” with 7 species of foraminifiers and 10 forms of conodonts (Beus, 2003). Bone fragments and sharks teeth are common in this unit. Plant fossils are most common in the western part of the lower unit, with multiple types of algal structures and “12 species of plant megafossils” (Beus, 2003).
The contact between the Surprise Canyon and the Watahomigi Formation is often obscured, but where it is exposed, a low-angle unconformity has been observed and the basal unit of the Watahomigi is a thin, widespread, locally discontinuous limestone pebble conglomerate with minor traces of chert clasts (Beus, 2003). If the conglomerate is absent, the contact can be identified by a purplish red calcareous siltstone and mudstone.
Geologists believe that the Redwall Limestone was deposited under shallow, open-marine conditions on a submerged continental shelf across northern Arizona during the Middle Mississippian. The four members of the Redwall were formed during two major transgressive-regressive cycles. Following the final phase of Redwall deposition, seas withdrew from the Colorado Plateau region for a time and the exposed landscape became karstified and incised by west-flowing rivers. A subsequent Late Mississippian to Early Pennsylvanian marine transgression produced the Surprise Canyon Formation. The rapid and significant lateral and vertical facies changes within the Surprise Canyon has lead geologists to infer deposition corresponding to a major east-west estuary system in the Grand Canyon area that emptied onto the paleocoastline near the Nevada-Arizona border. The eastern limit of marine fossils suggests a fluvially-dominated system in the eastern canyon progressing to a more marine-dominated environment in the central to western canyon at this time. The region remained largely unaffected by tectonic activity during the Mississippian, although the Antler Orogeny to the west likely caused sporadic upwarping of the crust, aiding the occasional exposure event and production of unconformable surfaces within the Redwall-Surprise Canyon sedimentary rock sequence.
The Supai Group (by Hannah Slover and Ken Bevis)
Stretching across the Grand Canyon from rim to rim is a thick band of red and tan cliffs and slopes that comprise the Supai Group (Figure 1). To even the untrained eye, the alternating stepped-ledge layers of sandstones, mudstones, and limestones of the Supai Group are easily distinguished from the soft, dark red slopes of the Hermit Shale above and the sheer cliff of the Redwall Limestone below (Figure 12). Helble (2011) suggests that these rocks reflect a diverse array of paleoenvironments that included continental, shoreline, and shallow-marine origins. The Supai Group is divided into four individual formations beginning with the oldest unit: the Watahomigi Formation characterized by red mudstone, siltstone, gray limestone, and dolomite; the Manakacha Formation dominated by quartz sandstone; the Wescogame Formation also of sandstone but intermixed with mudstone and limestone; and finally the Esplanade Sandstone with minor traces of mudstone and gypsum. Ages, determined primarily by brachiopods and fusulinids, indicate a correlation with the Pennsylvanian and Early Permian Periods. Common characteristics among the formations include mud cracks, ripple marks, and worm burrows (Hill, 2009). Vertebrate footprints occur in the Wescogame formation and Esplanade Sandstone (Helble, 2011), probably representing reptile tracks. Deposited prior to the Supai Group during the Mississippian Period is the Redwall Limestone, with the Surprise Canyon Formation sandwiched in discontinuous patches between the two. Following the deposition of the Supai Group is the Hermit Formation, also of the Permian period, and often associated with the Supai Group because of similarities in sediment type and inferred origin even though an unconformity separates the Hermit from the Esplanade Sandstone (Blakey, 2003).
Figure 12. The Supai Group consists of alternating cliff- and slope-forming “redbeds”, layers of resistant sandstones and limestones interbedded with less resistant mudstones that tends to produce a distinctive stepped-ledge appearance in outcrop; the view here is from Shoshone Point.
The oldest rock unit of the Supai Group is the Watahomigi Formation (Figure 1) which accumulated during the Early Pennsylvanian Period. It gradually thickens to about 300 feet in western Grand Canyon and the Grand Wash Cliffs compared to a thickness of 100 feet in the eastern Grand Canyon and in Sycamore Canyon to the South. Overall, it is composed of a mixture of red mudstone, siltstone, gray limestone, dolomite, and some sandstone. Informally, the Watahomigi Formation is divided into three horizontal sections: the lowest, a redbed slope, the middle, a carbonate ledge, and the upper section, another redbed slope. In addition to a horizontal division, the Watahomigi Formation can be divided laterally into an eastern and western facies, eastward being dominated by clastic redbeds and westward by the carbonates. The carbonate sediments in the east are mostly fine-grained limestone, while the limestone is more granular to the west with “abraded fossil fragments, accretal grains, and pellets” (Blakey, 2003). The majority of the mudstone is composed of “non-descript, slope-forming, poorly exposed units with thin, intercalated, bioturbated (disturbed by organisms), bright-orange, limey sandstone” (Blakey, 2003). When the redbeds are exposed, various bed-plane tracks, trails, and burrows along with a basal chert-pebble conglomerate are commonly observed (Blakey, 2003). This combination was deposited as a result of regional subsidence. Sediment from the low-energy, shallow-marine and coastal-plains blanketed the area to create the mudstones, carbonates, and minor sandstones (Blakey, 2012). Where the Watahomigi Formation lies directly on the Redwall Limestone, the contact is typically “sharp and unconformable,” although the abruptness of the contact is less discernible between the Watahomigi and Surprise Canyon Formation. The upper contact with the Manakacha Formation appears to be conformable thought not necessarily of level stratigraphy since the contact joins the slope below with steep slope or cliff above (Blakey, 2003).
Following the deposition of the Watahomigi is the Manakacha Formation (Figure 1). Also formed in the lower Pennsylvanian Period, it marks a change in the depositional facies and their thicknesses that dominated earlier time periods during the Paleozoic era. Previously, carbonates and minor mudstones were the norm, but the Manakacha Formation is largely composed of quartz sand with minor traces of mudstones and limestones, and is thickest in the central Grand Canyon, rather than in the west as are other Supai Group formations. As Blakey (2003) explains “the deposition of the Manakacha reflects a significant [environmental] change across the western interior of the United States”. The formation is typically 300 feet thick in Grand Canyon, while only 150 feet thick throughout the Verde and Chino valleys to the south. It consists of three types of sandstone containing “very fine- to medium-grained quartz grains to ooids, abraded fossils, and peloids” (Blakey, 2003). Units are typically held together with calcite, and the “more limey” sections are accompanied by Jasper (red chert). According to Blakey (2003), the first sandstone is cross-stratified and comprised of trough, planar, and compound bed sets, in thicknesses ranging from 1 to 30 feet. Within these strata are found climbing translatent strata. Supporting an interpretation of eolian deposition for the unit are thin laminae that show reverse grading while each lamina records the migration of a single wind ripple. The geometry of the strata leads geologists to suspect deposition at the forward base and between dunes, on eolian sheets of sand (supported by the horizontal to very low angles of sediments), and due to migration of the eolian dunes (supported by the climbing strata). The second sandstone is described as having “horizontally laminated to very low-angle, cross-stratified, fine-grained, calcareous, and silty units up to 20 feet thick” (Blakey, 2003). Of a deeper red color, this stratum is also characterized by rare ripple lamination and the abundance of local tracks and trails on the bedding planes. This evidence suggests a subaqueous genesis origin, but clues are limited. Finally, the third sandstone found in the Manakacha Formation is very fine-grained sand, ranging from structureless to greatly bioturbated layers. It is thought that both deposits have undergone bioturbation, but the structureless layers were so extensively mixed that all traces of burrows or tracks have been removed. This sandstone’s color ranges from grayish orange to a bright reddish orange. The strata of the Manakacha Formation often display evidence of alteration by diagenesis (growth of calcite and dolomite crystals), and minor amounts of mudstone, fine-grained limestone, and dolomite are common throughout the formation.
As the Manakacha Formation moves toward the northwest, it increases in cross-bedded carbonates, implying deposition by fluid currents (Blakey, 2003). The mixture of river, shoreline, and shallow-marine sediment suggests complex intertonguing of low-energy, intertidal, estuarine, and eolian environments as the sea continued to transgress and regress (Blakey and Middleton, 2012). A regional unconformity occurs at its upper contact with the overlying Wescogame Formation, leaving absent the Middle Pennsylvanian period. The boundary is identified by scouring, channeling, and a coarse lag of conglomerates, although it is vague in the southern Colorado Plateau due to poorer outcrops, and in general is difficult to locate, and was likely associated with prolonged subaerial exposure (Blakey, 2003).
The third formation of the Supai Group is the Wescogame Formation (Figure 1). Deposited during the upper Pennsylvanian period, its lithologic components are exceedingly similar to that of the Manakacha. This sheet-like body of sand expands across the Grand Canyon and the western Mogollon Rim region while maintaining thicknesses from 100 to 200 feet. In the central Grand Canyon and western Chino Valley, sandstone is the dominant facies, while mudstone increases in the east and limestone to the west of Chino Valley. The Wescogame Formation is informally divided into an upper slope unit and a lower cliff unit. In addition, three north-east-trending facies belts have become associated with this formation. The easternmost belt is the redbed facies, located near the Hermit Basin and spanning to the western Verde Valley. The middle belt is the sandstone facies, extending west from the eastern belt through the Shivwits Plateau region. Finally, cross-stratified limestone and limey sandstone composes the limestone facies of the westernmost belt (Blakey, 2003). While the depositional environment is still the same low-energy environments of the Manakacha, a decrease in marine fossils leads geologists to suspect an increase in continental settings (Blakey, 2012). The lower contact is with the Manakacha Formation and, as discussed above, is often difficult to locate because of the variations of the conglomerates and debris that tend to obstruct the view. Most likely marking the disconformity between the Pennsylvanian-Permian boundary is the Wescogame’s upper contact with the Esplanade Sandstone, an unconformable contact that is generally simple to locate by the presence of lag conglomerates and its high relief of up to 50 feet (Blakey, 2003).
The Esplanade Sandstone forms an easily distinguished cliff band in the Grand Canyon directly below the slope-forming Hermit Formation (Figure 13) and is the final formation of the Supai Group (Figure 1). This rock unit is characterized by minor mudstones and bedded gypsum (Blakey, 2012) and has a higher percentage of sandstone than any other in the Supai group (Blakey, 2003). It accumulated during the Lower Permian, and it is correlative with the Pakoon Limestone, a carbonate mud stratum, intertonguing from the west (Blakey and Middleton, 2012). Blakey (2003) describes the formation as a wedge, thickening from around 200 to 250 feet in eastern Grand Canyon and western Mogollon Rim region to 800 feet in the northwest at the western end of the Grand Canyon along the Grand Wash Cliffs. The 800 foot thickness includes the Pakoon Limestone which thickens to about 300 feet along the Grand Wash Cliffs. The Esplanade Sandstone is compositionally very similar to the Manakacha Formation, but it does differ in its dominance of cross-stratification. It also exhibits evidence of eolian deposition due to the abundance of climbing translatent strata. Its color, originally a pale grayish orange to reddish orange, appears dark reddish brown due to staining from iron oxides by the overlying Hermit Formation. The formation contains multiple stacked and interfingering sandstone units, ranging from 5 to 50 feet thick, and divided by “thin, red mudstone; by fine-grained carbonate units; or by prominent, irregular-bedding planes” (Blakey, 2003). It is informally divided into a basal slope unit and the main cliff unit. The Pakoon Limestone formed in the west grades laterally from this basal slope unit, a transition characterized by sandstone grading into dolomite and limestone (Blakey, 2003). The majority of eolian sand in the west is of marine origin while the east is of fluvial origin. The abundance of cross-stratification in the sandstone is supportive of eolian dune and sand sheet environments. The minor mudstones were likely deposited in tidal-estuarine and fluvial floodplains, the gypsum on coastal sabkhas, and the limestone in a shallow-marine environment (Blakey, 2012). Still debated amongst geologists, the Esplanade Sandstone’s upper contact with the Hermit Formation is generally believed to suggest a time of environmental transition and/or a regional unconformity. Different evidence includes 30 to 50 feet of local relief, channeling (originating from the Hermit Formation) into the Esplanade Sandstone, and cross-stratified sandstone 100 feet thick, beveled into the contact. In the west, the lower slope unit transitions into the Pakoon Limestone as indicated, while the upper cliff unit increases in carbonate content. In the Toroweap area, the top of the Esplanade Sandstone experiences bedded gypsum that is “intercalated with cross-stratified sandstone throughout a zone as much as 200 feet thick (Blakey, 2003).
Figure 13. The Esplande Sandstone forms a prominent cliff-band below the Hermit Shale’s equally distinct reddish mudstone slopes; the view here is from the Boucher Trail looking west across upper Hermit Canyon toward Dripping Springs.
The Supai Group is “sparsely and sporadically fossiliferous” (Blakey, 2003). The fossil types consist of trace fossils—including “trackways, burrows, impressions, resting marks, and feeding marks,”— invertebrates, and flora (Blakey, 2003). The abundance of swirls, crinkles, and irregular patterns make bioturbation the most common trace fossil in the Supai Group. Burrowed structures appear most commonly in silty sandstones and carbonates in a variety of types, but cylindrical, smooth burrows are dominant. Trackways occur specifically in the Wescogame Formation and the Esplanade Sandstone, usually created by terrestrial vertebrates. The fossil invertebrates most commonly observed from the Watahomigi Formation are brachiopod fauna, and marine fossils occur most often in deposits of the western Grand Canyon. However, any presence of abraded fossils in fine sandstone or limestone should not lead one to assign a marine deposition to the formation. The sedimentological evidence of eolian deposition suggests that a much more likely factor for the deposition of these particular fossils involves wind deflation of subaerially exposed marine sediments. Finally, flora is sparsely distributed throughout these formations and suggests coastal floodplains created during regressions (Blakey, 2003).
The origin of deposits that comprise the Supai Group is complicated and a subject still under consideration. The initial thought included a compilation of fluvial, deltaic, beach, shallow-marine, or estuarine origins, but after the evidence of wind ripples was identified, an eolian depositional setting seemed much more plausible for much of the material. Controls during the deposition of the Supai Group were very complex, including the continuous eustatic rise and fall in sea level, regional subsidence, and an influx of sand from the north. As Blakey (2003) states, “The heterogeneity of much of the Supai is directly related to this complex intermixing of environments coupled with a broad range of available depositional material.” In general, the Supai Group seems to have been deposited on a broad coastal plain associated with a passive continental margin, with individual settings ranging from shallow marine to continental depending on fluctuating sea level (Figure 14). However, it is thought that the eolian dune deposits occurred in a coastal-plain setting. Quartz sand was transported from the north by wind, but the continuous transgressions and regressions of the Pennsylvanian and Early Permian periods interrupted eolian deposition, preventing full development. While many carbonates were clearly deposited from a shallow-marine origin, minerals such as calcium carbonate can be precipitated within sediments as a result of extreme aridity which causes salty groundwater to rise upward by capillary action and evaporate (Blakey, 2003).
Figure 14. Early Pennsylvanian through Early Permian paleogeography of western North America during deposition of the Watahomigi Formation (A), Manakacha Formation (B), Wescogame Formation (C), and Esplanade Sandstone (D), respectively; original maps are by Ron Blakey and presented in Blakey and Ranney (2008) and Blakey and Middleton (2012).
Though debate continues about the Supai Group’s origins, its formations were likely deposited on a passive continental margin that experienced multiple, transgressive-regressive cycles of short duration. Environments transitioned from shallow marine to continental across a broad coastal plain, but became increasingly characterized by eolian deposition upward through the sequence. While blowing sand associated with an arid climatic setting dictated much of the deposition during the Pennsylvanian and Early Permian in the North American southwest, global tectonics also played a significant role. Assemblage of the supercontinent Pangaea was completed during this period, and compression generated by plate collisions along the southern and southwestern edge of the North American plate penetrated far inland to form the Ancestral Rocky Mountains (Blakey and Ranney, 2008). Erosion of the Uncompaghre Uplift, a western extension of the Ancestral Rockies in western Colorado and eastern Utah, generated copious amounts of sediment that accumulated to vast thicknesses on the eastern Colorado Plateau well into the early Mesozoic. Deposits mainly accumulated in the Paradox Basin on the west side of the Uncompaghre highlands. To the west, in the Grand Canyon region, deposition from these eastern highlands was intermittent and limited to distal, fine clastics deposited on a continental shelf setting. Sandstones and mudstones of the Supai mark the first major influx of clastics into the region since the Middle Cambrian, provided for the most part by erosion of the Ancestral Rockies (Blakey and Middleton, 2012). Repeated glaciations in the Southern Hemisphere, where much of the Pangean landmass was located, probably induced the sea level changes experienced on western North America’s passive margin as multiple, short-term marine transgressions and regressions. During higher sea levels, carbonate platform conditions overwhelmed clastic deposition to prevail in many areas, and during lower sea levels, clastic deposition in the form of eolian sands dominated. Over 150 million years, eolian deposits accumulated during regressions, were planed-off by transgressions or minor fluvial events, and repeated.
Deposition of the lower two-thirds of the Watahomigi Formation likely occurred in a shallow-marine to coastal plain setting along a fluctuating, low-energy shoreline (Figure 14a) (Blakey and Middleton, 2012). The mudstone was probably deposited on along a shoreline tidal flat. It was followed by carbonate deposition as the sea level rose or the Grand Canyon region subsided, with cycles of fine clastic influx followed by erosion marking short-lived regressions. Sediment accumulation was initially restricted to the Antler foreland basin until transgression allowed shallow marine conditions to invade much of the Colorado Plateau region (Blakey and Ranney, 2008). The upper third of the Watahomigi Formation and the entirety of the Manakacha Formation experienced periodic marine transgression into the Grand Canyon region paired with an increase in eolian material from the north during temporary regressions (Figure 14b). Periodic marine flooding of a broad, arid- and eolian-dominated coastal plain prevailed (Blakey and Ranney, 2008; Blakey and Middleton, 2012). After a relatively significant regional unconformity developed during prolonged subaerial exposure in the Middle Pennsylvanian, deposition of the Wescogame Formation in the Late Pennsylvanian marked a substantial change to more pervasive continental conditions paired with eolian deposition across the Colorado Plateau region (Blakey and Ranney, 2008) (Figure 14c). Minor transgressions paired with carbonate platform deposition continued to penetrate inland from the west giving the Wescogame a similar character to the Manakacha below (Blakey and Middleton, 2012). The erosion surface at the upper contact of the Wescogame, representative of the Pennsylvanian-Permian boundary, is thought to have been caused by streams crisscrossing newly exposed terrain resulting from a lowering of sea level. Finally, during the Early Permian, widespread eolian deposition formed the Esplanade Sandstone to the east, while the Pakoon Limestone member of the Esplanade Sandstone formed in shallow-marine, carbonate platform environments of clear water to the west (Blakey and Middleton, 2012) (Figure 14d). Interruptions were again caused by fluctuations in sea level coupled with fluvial activity, eventually transitioning to fluvial conditions as the eolian supply diminished (Blakey, 2003). Overall, the history of depositional settings of the Supai Group is very complex, and as studies continue more definitive, alterations will likely be made to the current interpretations described above.
The Supai Group was deposited on a passive continental margin. Recurring marine transgressions and regressions of the Pennsylvanian and Early Permian periods coupled with the influx of eolian sands caused a battle to ensue for depositional dominance. The sea would regress and eolian dunes would form on the broad coastal flat, only to be interrupted by a brief transgression, limiting eolian development. Carbonate units in the Supai Group are not solely derived from marine origin, but also of eolian (arenaceous limestone), and their interbedding is still not fully understood. To add complications, as the sea transgressed and regressed, debris was washed into the area by ephemeral streams, trapping mud and fine-grained sediment in lagoons, tidal flats, and river channels. Eolian deposition definitely played a significant part in the deposition of the Supai Group, but wind-generated sediments were continuously reworked, altered, or limited in extent by changes in sea level.
The Hermit Formation (by Meagan Redmond and Ken Bevis)
Along the upper walls of the Grand Canyon, the Hermit Formation occupies a distinctly reddish, slope-forming bench above the Esplanade cliff-band and at the base of the contrasting white cliffs of the Coconino Sandstone (Figure 1). The formation overlies the equally reddish cliff bands of the Esplanade Sandstone, the last formation of the Supai Group, and the lowermost Permian rock unit in the canyon (Figure 12 and Figure 13). All five of the sedimentary rock units from the Esplanade Sandstone, upward through the Kaibab Formation that creates the uppermost sedimentary rock layer within the Grand Canyon’s walls, accumulated during the Permian (Blakey and Middleton, 2012). Though many of the Permian age formations are well studied and documented, the Hermit Formation and the Coconino Sandstone are the least explored, with the Hermit Formation in particular being described as “one of the poorest known units in Grand Canyon” due to a seeming lack of interest by geologists in studying the rock unit (Blakey, 2003).
The Hermit Formation was once considered a member of the Supai “Formation”, mostly due to the fact that like the other formations of the later established group, the Hermit has a predominantly red strata and its depositional origin was thought to closely resemble that of the Supai Group’s (Blakey, 2003). In 1922, the Hermit was seperated from the underlying members of the Supai Formation as the Hermit Shale, and in 1975, when geologist Edwin McKee formally established the Supai Group’s four formation (the Watahomigi, Manakacha, and Wesogame Formations and the Esplanade Sandstone), he also renamed the Hermit Shale as the Hermit Formation. More recent publications on the Hermit restrict their discussion to the Hermit’s composition and appearance in outcrop, and its relationship to the Supai Group.
After deposition of the Early Permian eolian sandstones of the Esplande, all previous basins across the Colorado Plateau had been overwhelmed and filled with sediment (Blakely and Ranney, 2008). Lack of accomodation space may then have induced a phase of erosion. The Early Permian Hermit Formation is seperated from the slighly older Esplanade Sandstone by what McKee called a “regional unconformity” (Blakey, 2003). As Blakey (2003) describes, in some areas there is a definable transition from the underlying Esplanade to the overlying Hermit. Yet, in most locations at or near the Grand Canyon, the Esplanade and the Hermit are seperated by “an erosional surface with 30 to 50 feet of local relief.” Other locations also indicate channeling into the Esplanade Sandstone; these channels are filled with Hermit sediments, and thus “originate from within the Hermit and not from the base [of the Esplanade],” and most of these channels are filled by sediments affiliated with the Hermit Formation (Blakey, 2003). This erosional relief, therefore, “may be related to channel cut-and-fill sequences,” meaning that as stream channels cut across the Esplanade Sandstone, the Hermit Formation was also beginning to be deposited, locally filling in some of these channels with sediment.
The main body of the Hermit Formation is comprised of brick-red sandstone and mudstone, with minor pebble conglomerate and eolian loess and sand dune deposits associated with fluvial settings on a broad coastal plain (Blakely and Ranney, 2008; Blakely and Middleton, 2012). The rock unit formed as arid-land river systems spread sediment southwestward from the Uncompaghre uplift; the sandstones and conglomerates compose river channels deposits, while the mudstones were deposited on intervening floodplains (Figure 15). Early Permain Hermit deposition was a time of crustal stability in the Grand Canyon area, although areas to the south and west were beginning to subside in a forearc basin as the complex McCloud volcanic arc initially began to collide with the North American plate; this collision would eventually generate the Sonoma orogeny and reinvigorate the older Antler highlands (Blakely and Ranney, 2008).
Figure 15. Early Permian paleogeography of western North America during deposition of the Hermit Formation; original map is by Ron Blakey and presented in Blakey and Ranney (2008) and Blakey and Middleton (2012).
The Hermit Formation was deposited about 280 million years ago and is named after the Hermit Basin along Hermit Canyon (Wuerthner, 1998) where the unit is well exposed. According to the US Geological Survey, this formation is composed of “red, slope-forming, fine-grained, thin-bedded siltstone and sandstone.” Blakey (2003) also describes the Hermit as consisting mostly of “interbedded, silty sandstone and sandy mudstone.” The dark red-orange slope-forming characteristics of the Hermit help to distinguish the formation from the overlying tan-colored Coconino Sandstone (Timmons, 2003). However, the Hermit Formation’s “fine-grained nature” causes the formation to be more suseptible to erosion, leading to “relatively poor exposures” and making the formation somewhat difficult to study; and thus, to the relative lack of interest by geologists (Blakey, 2003). The Hermit Formation is about “100 feet thick in the eastern portion of the Grand Canyon and near Seligman, AZ to over 900 feet in the Toroweap and Shivwits Plateau areas” and can average about 300 feet thick in the Mongollon Rim region east of Seligman (Blakey, 2003). The rock unit comprises a heterogeneous, chiefly a fine-grained, silisiclastic sequence, with neither vertial or lateral distribution of its components suggesting any distinct trends (Blakely, 2003). Silty sandstone and sandy mudstone are the most volumious and widespread sediments. The units are interbedded rhythmically, although mudstones generally increase upward within the formation. Sandstones range from structureless to ripple cross-laminated and trough cross-stratified; ledge-forming beds average 3 feet in thickness and are structureless. Ripple cross-laminated and trough cross-stratified layers are associated with point bar deposition in meandering streams. Sandy mudstone forms slopes in the Hermit Formation and is commonly featureless.
The Hermit Formation is sparsely fossiliferous, but a broad range of fossils have been identified (Blakely, 2003). Trace fossils, including burrowed structures, possible feeding and resting marks, invertebrate trackways, and bioturbation by plant roots are common. Plant fossils are widespread but few in number within the Hermit, but they do suggest the presence of broad floodplains developed on a semiarid to arid landscape.
Timmons (2003) suggests that because of the continuous sea level fluctuations that had been occuring up to and during this time, “shallow marine, lagoon, and fluvial environments” were possible. He also explains that the Hermit Formation, as well as the other formations of Permian age, were “the first layers deposited in an eolian (wind-dominated) environment in the Grand Canyon region” and that these “layers mark the onset of the coastal sand dune environment that was dominant in the region through the late Jurassic Period” (Timmons, 2003). Blakey and Middleton (2012) indicate that the depositional environment for the Hermit Formation was likely a “broad coastal plain, chiefly in fluvial settings, but also wind-blown dust (loess) and scattered dune deposits of eolian origin.” The eolian deposits that contribute to the formation are also present within the overlying Coconino Sandstone, but in much larger quantities deposited in an “arid, inland dune setting” (Blakey & Middleton, 2012). This suggests that marine regression was approaching its maximum with the deposition of the Hermit Formation, and would later hit its peak with Coconino deposition. During this time, most of Colorado was dominated by the Ancestral Rocky Mountains, and westward draining “ephemeral river systems” were scattered over much of southwestern North America, with a shallow-marine shelf environment restricted to northwestern Utah (Figure 15). Throughout the Grand Canyon area, the contact between the Hermit Formation and overlying Middle Permian Coconino Sandstone is sharp and without gradation everywhere it has been observed (Blakely, 2003). The top of the Hermit is distinguished by descication cracks up the 20 feet deep that are frequently filled with the quartz sand of the Coconino, suggesting a rapid and intense return to aridity in Coconino time.
The Hermit Formation has undergone only limited investigation by geologists mainly because its degraded slopes make observation difficult. The Hermit Formation was deposited in an arid climate, possibly on a broad coastal plain where fluvial settings dominated, but eolian sands continued to be intermittently carried across the Four Corners region. Tectonics did not play a critical role in the Hermit Formation’s deposition; although the Sonoma Orogeny caused areas to the southwest to subside as the McCloud volcanic arc collided with the North American plate. Eustatic sea-level changes associated with Southern Hemispheric glaciation continued to produce cyclical flucuations in sea level, periodically exposing coastal areas to wind erosion and sand transport. The overall marine regression that began at the start of Supai Group deposition, continued erosion of the Ancestral Rocky Mountains, and filling of sedimentary basins across the region, dominantly controlled Hermit Formation’s deposition. Factors such as the presence of mudstone interbedded with sandstone, an overall fining-upward sequence within the formation, and sparse terrestrial plant fossils, indicate that the Hermit was formed in an environment that was somewhat wetter than underlying and overlying eolian sandstones, but also one in which initially higher energy streams where gradually replaced by lower energy streams. It’s dark red, gentle slope-forming appearance sets it apart from the overlying Coconino Sandstone, but an erosional disconformity seperates it from the otherwise similarly colored underlying Supai Group. The paleogeography of the Colorado Plateau region during the time of Hermit Formation deposition was likely a terrestrial-dominated area with a large, shallow body of water lying to the northwest. This trend of water transgressing and regressing from the northwest into and out of forearc basin created by the Antler-Sonoma orogenies would continue into the Triassic, although rock units of that age are almost entirely eroded from the Grand Canyon area.
The Coconino Sandstone (by Meagan Redmond and Ken Bevis)
The massive cliff-band of the Coconino Sandstone, resting directly on the brick-red slopes of the Hermit Formation (Figure 1), may be the most easily recognized and traced formation in the Grand Canyon despite the relative lack of detailed knowledge concerning this poorly studied unit (Middleton, et al., 2003). The formation is Early Permian in age, its deposition dating back to around 275 million years ago. Other formations within the Grand Canyon formed during the Early Permian include the Esplanade Sandstone of the Supai Group and the immediately underlying Hermit Formation. Much of the Early Permian was a time of only minor tectonic activity on North America’s western passive continental margin; although by the latest Early Permian areas to the southwest were actively subsiding in a forearc basin as the Sonoma orogeny progressed associated with collision of the McCloud arc with the passive margin of the North American plate (Blakely and Ranney, 2008). Deposition of the Coconino and prior Permian age formations is considered to have been chiefly in response to “glacial-eustatic sea-level oscillations.” Deposition of these rock units was affected by global-wide glacial activity that caused sea levels to rise and fall (Hopkins and Thompson, 2003). The Coconino Sandstone was deposited during a time when sea level had receded westward, and the Grand Canyon area was dominated by a terrestrial, desert-like environment. Unlike other Permian formations within the Grand Canyon, the Coconino is not subdivided into members.
The Permian west coast of North America underwent several cycles of marine transgressions and regressions while undergoing a major period of overall regression which had begun as far back as Wescogame Formation deposition during the Late Pennsylvanian period, a regression that ultimately led to the deposition of the terrestrial Hermit Formation and the Coconino Sandstone (Blakey and Middleton, 2012). Coconino deposition occurred during a time of extreme aridity, with eolian sands covering two-thirds of Arizona and New Mexico, and yet “local high water tables furnished enough water to allow a diverse reptile fauna to exist” (Blakey and Middleton, 2012). No sedimentary rocks of the latest Early Permian are preserved in the walls of the Grand Canyon, presumably this stable shelf area simply acted as an eolian transport surface for sand dunes because it lacked the capacity to retain sediment. However, the equivalent latest Early Permain rocks of the Schnebly Hill Formation and De Chelly Sandstone were deposited in the Holbrook basin to the southeast where subsidence allowed accumulation of laterally adjacent and intertonging sand dune and tidal flat deposits.
The Coconino Sandstone forms thick, light-colored and mega-crossbedded cliffs directly above the Hermit Formation’s contrasting reddish slopes (Figure 16). Taken in sequence, the rock unit is easily recognized from any North or South Rim overlook as the third layer from the top, where it forms the second riser in a series of cyclopethan stair steps descending into the canyon; the top step is the rim, the top riser is formed by the rim-capping Kaibab Limestone, the second step formed by the Toroweap Formation’s well-vegetated slopes, the second riser the Coconino, and the third step the Hermit Formation’s brick-red slopes. Blakey and Middleton (2012) describe the Coconino Sandstone as pale-yellow in color, consisting “almost exclusively of large-scale cross-bedded sandstone.” This sandstone is composed of “fine-grained, well-sorted” sand with large amounts of quartz and some feldspar, and the formation overall has a tendency to form massive, “high angle” cliffs in the Grand Canyon (Middleton, et al., 2003). McKee (1979) describes the lithology of the Coconino Sandstone as dominantly quartz arenite with a siliceous cement of a white to pink color with sparse amounts of feldspathic sediment present throughout. The grain sizes range from fine to medium and are well sorted, and the grains are well rounded to very well rounded with frosted surfaces that are indicative of wind transport (Blakely, 1990). These characteristics are indicative of mature grains that have undergone substantial transportation, opposed to immature sediments which would contain greater lithics and feldspars, would be less rounded, more poorly sorted, and consist of coarser particle sizes. McKee (1979) indicates that the predominant sedimentary structure found in the formation is steeply dipping, large-scale or mega-crossbedding, although well preserved ripple marks, and slump structures also occur.
Figure 16. The yellowish-tan colored cliffs of the Coconino Sandstone are easily distinguished from the overlying and underlying slopes of the Toroweap and Hermit Formations, forming the second riser (third layer) in the sequence of massive stair steps leading downward into the Grand Canyon; viewed from the Hermit Trail the Dripping Springs tributary of upper Hermit Canyon.
The Coconino Sandstone exhibits significant thickness variations which are generally inconsistent with those of its overlying Permian counterparts, the Toroweap and Kaibab Formations (Blakely, 1990). Bucking the normal trend of rock units in the Grand Canyon, it thickens southeastward and is 65 feet thick near the Grand Wash Cliffs at the western end of the canyon, 96 feet thick at Toroweap Point, and 300 feet thick in the Bright Angel Creek area (Babcock et al., 1974). At the central part of the canyon where the formation is thickest, the Coconino is 600 feet thick near Cottonwood Creek (Middleton, et al., 2003). Overall, the formation thins northward and westward. The directionality in which this “dry” formation thins indicates where marine conditions would most likely have prevailed; in studying sedimentary deposits, geologists know that where sand dunes and other associated terrestrial facies are thickest, the drier the area, suggesting that the Coconino’s depositional environment was more strongly influenced by terrestrial conditions. Where the terrestrial facies begin to thin, “wet” facies become thicker, indicating marine domination in the area. Because the Coconino Sandstone consists exclusively of homogeneous sands that consistently exhibit large scale crossbedding, it is widely accepted that the depositional environment of this rock unit largely resembled wind-blown sand deposits spread over most of Arizona and New Mexico in an enormous desert erg, much like the African Sahara of today (McKee, 1979; Blakely, 1990; and Middleton, et al., 2003) (Figure 17).
Figure 17. Early Permian paleogeography of western North America during deposition of the Coconino Sandstone; original map is by Ron Blakey and presented in Blakey and Ranney (2008) and Blakey and Middleton (2012).
Evidence within the Coconino suggests that there was some surface water present at times. As Middleton et al. (2003) explain, not only does the Coconino Sandstone have trace fossils of reptile footprints, but within the sedimentary structures of the Coconino there are “rain drop impressions.” Much like a modern desert, any rainfall that came during Coconino deposition would have been sparse and short-lived, and the reptiles that thrived in this environment would have been small and hardy, able to live on very little water. Middleton et al. (2003) also describes ripple marks present within the Coconino, and many geologists interpret that the ripple marks were likely the result of active crosswinds (Babcock et al., 1974).
Coconino deposition suggests that the paleogeography of the Colorado Plateau region was terrestrially dominated (Figure 17). The Ancestral Rocky Mountains and Uncompahgre Highlands, formed initially during latest Mississippian-Early Pennsylvanian period, dominated most of Colorado and parts of Utah and New Mexico to the east, while the regrowth of the Sonoma-Antler Highlands to the west was underway related to the ongoing Sonoma Orogeny. A huge desert covered much of Utah, New Mexico, and Arizona in between remnant and growing highlands (Figure 17). Marine and shoreline environments would have been present only at the uppermost corner of Utah and at the southernmost ends of Arizona and New Mexico. These neighboring bodies of water would later transgress inland to inundate the region during Toroweap and Kaibab deposition, and marine depositional environments would transform the Grand Canyon area, ending Coconino deposition (Blakey and Middleton, 2012).
Even though relatively few geologists have studied the Coconino Sandstone, a deciphering of its sedimentology and paleontology indicates that it was deposited in a terrestrial setting dominated by an arid climate where there was relatively little water, except for an occasional brief rain. The Coconino Sandstone exhibits textbook examples of sand dune facies deposited under intense eolian conditions and its lithology, texture, sedimentary structures are all strong evidence of this conclusion. Like a modern desert, wind-blown eolian sand dunes dominated the Four Corners region and only hardy reptiles could live in this harsh environment. Coconino deposition took place at the very end of a significant period of marine regression, and its deposition was eventually terminated by a return of marine conditions to the Four Corners region as a transgression swept across the Grand Canyon area from the west to deposit the Toroweap and Kaibab Formations.
The Kaibab Limestone and Toroweap Formation (by Meagan Redmond and Ken Bevis)
The Kaibab and Toroweap Formations form the uppermost layers exposed in the walls of the Grand Canyon (Figure 1), and are recognized as a distinct pair of cliff and slope bands just below the rim (Figure 16). Both rock units were formed during the Early to Middle Permian period of the Paleozoic era (Turner, 2003), which began about 286 million years ago and lasted until 248 million years ago. Other Permian formations deposited within the Grand Canyon area include the slightly older Early Permian Coconino Sandstone and Hermit Formation; Late Permian rock units are not preserved (Blakey and Ranney, 2008). The overlying Kaibab and underlying Toroweap Formations are so compositionally similar and stratigraphically related that the two were originally paired by geologists into one formation, the Aubrey Limestone, which was later changed to the Kaibab Limestone. It would not be until 1938 that geologist Edwin McKee would ultimately split the Kaibab Limestone into the two formations that we recognize today, the upper Kaibab Formation and the lower Toroweap Formation (Turner, 2003).
Tectonic and sedimentation patterns developed earlier in the Pennsylvanian continued during much of Permian time (Fillmore, 2011). The Colorado Plateau occupied a broad, shallow-marine, continental shelf associated with the passive southwestern margin of the North America continental plate (Blakey and Ranney, 2008). Sediments continued to accumulate as a westward thickening wedge across the region on a gradually subsiding foreland basin comprised of thin crystalline basement and the accreted remains of the Mississippian Antler Orogeny and the ongoing Sonoma Orogeny (Figure 18). Although there is little to no direct evidence of tectonic activity in the Grand Canyon area occurring during Kaibab or Toroweap deposition, lying more or less in the center of a broad basin hemmed in the west by the Sonoma-Antler Highlands, and to the east by the remnants of the Ancestral Rocky Mountains, the Permian period of the Colorado Plateau region as a whole was tectonically active. As Fillmore (2000 and 2011) describes, “The Permian saw the final assembly of the supercontinent known as Pangaea.” The Ancestral Rocky Mountains, the westernmost extension of the Marathon-Ouachita Belt that had formed during the Pennsylvanian in response to North America’s collision with the South American continent, continued to rise isostatically and influence regional drainage patterns and accumulation of sediment (Fillmore, 2000 and 2011). More specifically, the Uncompahgre highlands and adjacent Paradox basin formed both the source and sink for vast quantities of sediment in the Four Corners area, and sediment was contributed distally to most if not all of the Grand Canyon’s Permian formational layers. Throughout Early Permian time, subsidence of the Paradox basin would decrease gradually as it filled with westward-prograding sediments. By the end of the Permian period, the Uncompahgre uplift of the Ancestral Rockies was inactive, although sediment would continue to be generated and shed westward from this decaying highland well into the Triassic (Fillmore, 2011).
Figure 18. Late Early Permian and Middle Permian paleogeography of western North America during deposition of the Toroweap Formation and Kaibab Limestone, respectively; original maps are by Ron Blakey and presented in Blakey and Ranney (2008) and Blakey and Middleton (2012).
Throughout the Permian and into the Triassic, various island arc terranes (generally known as the McCloud Arc) approached North America from the west, eventually colliding along east-dipping subduction zones to initially form a trough-like foreland basin that accommodated sediments from the east. Collision of arc terrains eventually formed the highlands of the Sonoma Orogeny (Figure 18), a probable erosional source on the west side of the basin that now existed between these growing highlands and the remnant Ancestral Rockies to the east (Blakey and Ranney, 2008). The region would experience a number of other mild uplifts during the Permian (Fillmore, 2011), but none would contribute much sediment to the Grand Canyon area’s passive marginal setting. Instead, these uplifts mainly served to limit accommodation space and preservation of sediment in certain areas, especially during the Late Permian. Ensuing regional uplift prevented deposition or induced erosion of previously deposited sedimentary rocks over much of the Colorado Plateau for roughly the next 25 million years (until the Middle Triassic Period of the Mesozoic Era).
The Toroweap Formation underlies the Kaibab Limestone in the Grand Canyon (Figure 16), and the two formations are separated by a regional unconformity (Babcock, et al., 1974). The rock unit is distinguished as a slope-former because its higher mud content makes it relatively less resistant to erosion. The unconformity is an erosional surface, meaning that before the deposition of the Kaibab, the Toroweap Formation was subject to weathering and erosion through subaerial exposure, leading to the conclusion that marine conditions retreated from the Four Corners region for a time. The Toroweap was largely “deposited in open and restricted-marine environments, tidal flats, sabkhas, and eolian dunes fields” (Turner, 2003). Figure 18a depicts the paleogeography of these depositional environments. Turner further explains that “The Toroweap Formation, in its variety of lithofacies, reflects transgressions and regressions of an eastward-advancing and westward-retreating sea.” This is further confirmed by Harris and Tuttle (2004), who describe that the Toroweap Formation’s “nearly 200 feet of sandstone and limestone interbedded with gypsum records marine advances and retreats,” probably associated with an arid coastline. Thus, through lithologic and fossil evidence within the Toroweap Formation, we can see that during its deposition sea level rose and fell multiple times, and the primary source of the water lay to the northwest, advancing southeast into the basin formed between the steadily eroding Uncompaghre and Sonoma-Antler highlands. This is also shown by the formation’s distribution; it thins to the east and south from the Grand Canyon, presumably having less time to accumulate.
The Toroweap Formation has been divided into three members: alpha, beta, and gamma. These members from top to bottom include the Woods Ranch Member (or “alpha”), the Brady Canyon (“beta”), and the lowermost Seligman (“gamma”) (Turner, 2003). The Kaibab Formation has been similarly subdivided and the top of the Woods Ranch Member ends where the “gamma” member of the Kaibab begins. The members of the Toroweap are discussed in ascending order, so that changes in their composition can be associated with fluctuating depositional environments, sediment sources, and tectonic settings.
According to Babcock et al. (1974), the lowermost “gamma” member of both the Toroweap and Kaibab Formations can also be referred to as “the time of transgressing sea” and is a relatively thin unit. The Seligman Member marks the beginning of Toroweap deposition, and begins where the Coconino Sandstone Formation in the Grand Canyon ends. During Coconino deposition, the future Four Corners and Grand Canyon areas were mostly a terrestrial, desert-like environment with “enough water to allow diverse reptile fauna to exist” (Blakey and Middleton, 2012). However, as Babcock et al. (1974) suggest, marine conditions eventually returned to the region, and this transgression was marked by the beginning of Seligman Member deposition.
The Seligman Member largely consists of sandstone and evaporites, and is distinguishable in the Grand Canyon from the Coconino Sandstone because “it is the first non-cross-bedded” and evaporite-containing layer above the Coconino (Turner, 2003). In the Grand Canyon, this unit is less than 45 feet thick. The Seligman contains “thin-bedded dolomite, sandstone, and gypsum,” with the sandstone and evaporites dominating southeastward and the dolomite becoming more apparent northwestward. This change in facies from northwest to southeast can be interpreted as “a transition from coastal sabkha environments on the northwest to a continental sabkha to the southeast” (Langenheim Jr. and Schulmeister, 1987). The evaporites within this member and the ripple marks in the sandstone indicate that the Seligman was likely deposited on a shallow continental shelf environment where “warm restricted-marine waters were evaporated, promoting the precipitation of evaporite minerals” (Turner, 2003). Essentially, the evidence within the Seligman suggests that the water level was deepest to the northwest, where ocean water was coming in and transgressing onto the North American continent. As the water became shallower to the southeast, it became cut off from its original water source and subjected to intense evaporation, allowing for evaporites to form. Evidence within all of the members of both the Toroweap and Kaibab, as well as in other formations within the Grand Canyon, suggest that the main water source was located to the northwest.
The Brady Canyon (or “beta”) Member underlies the uppermost Woods Ranch (“alpha”) Member of the Toroweap, and is described by Babcock et al. (1974) as being “the time of maximum advance.” It is during this time of Toroweap deposition that the marine transgression reached its peak. The Brady Canyon is the thickest member of the Toroweap within the Grand Canyon, and its 280 feet of thickness is almost entirely composed of dolomite (Turner, 2003). Dolomite would usually indicate that deposition occurred in deep water; however, the lithofacies within the Brady Canyon Member have a “mud-supported texture,” which indicates deposition in quiet-water conditions consistent with the shallow water of a carbonate shelf (Turner, 2003). However, as the member moves southeast, the dolomite is “progressively replaced by wackestone, pelletal packstone, and quartzose dolomite,” again showing the shallowing of seawater to the southeast (Langenheim Jr. and Schulmeister, 1987). The Brady Canyon is the only fossiliferous member of the Toroweap; its two major faunal facies are open-marine fauna to the west (brachiopods, crinoids, horn corals) and more restricted, molluscan fauna to the east (bivalves and gastropods). According to Turner (2003), these fossils within the Brady Canyon Member “suggest a late Leondardian age for the Toroweap Formation.”
The contact between the lower Seligman Member and the Brady Canyon Member is described as being “on a regional basis, gradational” meaning that the Seligman gradually fades into the Brady Canyon (Langenheim Jr. and Schulmeister, 1987). However, the top of the Brady Canyon Member is more abrupt and easily distinguishable from the Woods Ranch Member, because at the top of the member there is dolomitic mudstone that has “abundant desiccation cracks” (Turner, 2003). These cracks indicate subaerial exposure, and this exposure indicates that at the end of the Brady Canyon Member’s deposition, sea level dropped and water receded from the area temporarily.
The Woods Ranch (“alpha”) Member is the uppermost member of the Toroweap and underlies the Kaibab Formation. Babcock describes the “alpha” member as being “the time of regressing sea.” Much like the Seligman Member, the Woods Ranch primarily consists of the evaporite gypsum, sandstone, and limestone; yet, the Woods Ranch does contain some dolomite at its base, making it conformable to the underlying Brady Canyon. As Turner (2003) describes, “The Woods Ranch Member typically extends from the top of the aphanitic dolomite that forms the uppermost unit of the Brady Canyon Member to the base of the cliff-forming limestone of the overlying Kaibab Formation.” The Woods Ranch is about 180 feet thick in the Grand Canyon, but unlike the other members of the Toroweap, the Woods Ranch “shows no consistent thickening or thinning trends” (Turner, 2003), meaning that it does not follow the typical pattern of thinning to the southeast or east like other rock units. Most likely, the Woods Ranch Member was deposited in a shallow shelf environment much like the environment that was present during Seligman Member deposition. The evaporites and carbonates within the Woods Ranch indicate periods when water was reintroduced back into the area, and the “eolian sandstones [within] suggest times of subaerial exposure” (Turner, 2003). The repetition of limestone, sandstone, and evaporate within the Woods Ranch Member’s deposition indicates that the member went through a type of “cyclic sedimentation”, possibly influenced not only by sea level but also by seasonal changes in depositional conditions. By the end of Woods Ranch Member, and ultimately Toroweap deposition, water completely receded from the area for a time and erosion took place, leaving behind the unconformity that separates the Toroweap Formation from the Kaibab (Turner, 2003).
The Middle Permian Kaibab Limestone is the uppermost rock unit within the Grand Canyon’s sedimentary sequence (Figure 1) and is discernible from the lower Toroweap Formation by its grey, interlayered slopes and stepped-cliff appearance (Figure 16). It is likely that the Kaibab and not the overlying Triassic Moenkopi Formation is the cap rock of the Grand Canyon area because the Moenkopi had a “less resistant nature” than the Kaibab, and was ultimately eroded away along with the uppermost layers of the Kaibab from most of the Grand Canyon area (Hopkins and Thompson, 2003). The Kaibab Formation was named after the Kaibab Plateau on the North Rim of the Grand Canyon, where it was first discovered and is well exposed. As its name suggests, the Kaibab Limestone is composed primarily of limestone (Wuerthner, 1998), which can reach a thickness of anywhere from 300 feet and, in northwestern Arizona where it is thickest, an excess of 500 feet (Hopkins and Thompson, 2003). The Kaibab overall was likely deposited in a subtidal, shallow-marine environment where minor fluctuations in sea level could abruptly change the depositional environment (Figure 18b), much like the environment present at the time of Toroweap deposition. These changes were not due to tectonic activity, as Hopkins explains: “Considering the quiescent (inactive) tectonic setting of the Grand Canyon region during the Permian, it is most likely that these cycles were caused by glacial-eustatic sea-level oscillations.”
Because the Kaibab and the Toroweap were originally grouped together as one rock unit (McKee, 1938), the Kaibab was also divided into three members: the lowermost gamma, the middle beta, and the uppermost alpha. However, because the rocks within the lower gamma member were confined to the Mogollon Plateau south of the Grand Canyon, “Geologists [now] interpret the gamma member as a facies within the Fossil Mountain (or “beta”) Member” (Hopkins and Thompson, 2003). Even though the Kaibab technically has three members, only two distinct members are now recognized and named: the uppermost Harrisburg Member and the middle/lower Fossil Mountain Member. The Harrisburg Member is typically thinner than the underlying Fossil Mountain Member, but its actual thickness at the time of deposition is impossible to determine accurately because subsequent erosion has stripped away the member’s uppermost portion.
Although the lowermost “gamma” member is no longer formally recognized as its own distinct layer within the Kaibab, the unit is still significant for understanding the depositional environments of the other members of the Kaibab Formation. As with the basal unit of the Toroweap Formation, the gamma layer represents a time of marine transgression. An unconformity separates the Woods Ranch Member of the Toroweap Formation from the gamma member of the Kaibab, indicating that, for a time, water receded from the area and allowed for subaerial exposure and erosion to occur. Then water returned to the area, just as it had after the dry spell during Coconino deposition, when the arrival of water was signified by deposition of the Toroweap’s Seligman Member. Evidence of this occurs in the Kaibab’s gamma member where it consists of “sandstone and dolomite with a rich [marine] fossil assemblage” (Fillmore, 2000). As Fillmore explains, “The gamma member was laid down as the sea advanced eastward across Utah and Arizona….”, and just as with the Toroweap, as marine waters advanced eastward, deeper water conditions would help to deposit the overlying beta member.
The Fossil Mountain Member is described by Fillmore (2000) as being a “massive, cliff-forming cherty limestone” that was deposited when sea level was at its highest. This member typically contains abundant and diverse marine fauna such as brachiopods, bryozoans, crinoids, sponges, and solitary corals. However, eastward within the member, the fauna changes to suggest more restricted conditions and the member becomes more siliciclastic dominated, indicating yet again that the water source lay to the west (Hopkins and Thompson, 2003). As the member thickens westward it can become up to 300 feet thick; and though the Kaibab Formation as a whole is largely known for its limestone composition, the Fossil Mountain Member is “75 percent sandstone or sandy dolostone” with scattered chert, dolomite, and limestone throughout (Hopkins and Thompson, 2003). The Fossil Mountain was likely deposited in a shallow “Kaibab Sea” that regularly experienced sea level rises and falls, extending eastward into the Grand Canyon area during transgression and receding westward toward the continental margin during regression (Fillmore, 2000). The difference between the beta member of the Kaibab Formation and the beta member of the Toroweap Formation is that, during the time of Fossil Mountain deposition, water covered nearly all of Utah, Arizona, and most of southern New Mexico, making this transgression much more extensive than the one that occurred during the Toroweap’s Brady Canyon Member (Blakey and Middleton, 2012).
The uppermost Harrisburg Member of the Kaibab Formation was formed during a time when the Kaibab Sea that had generated Fossil Mountain deposition was receding westward. As Hopkins and Thompson (2003) describe, “units that comprise the Harrisburg sequence reflect deposition within predominantly restricted-marine environments during cyclic westward retreat of the Kaibab Sea.” These repeated “shifts” in depositional environments due to sea level change can be observed throughout the Grand Canyon region with the regular alterations between carbonate, siliciclastic, and evaporite deposits. The Harrisburg Member in particular consists of an array of gypsum, dolostone, sandstone, redbeds, chert, and minor limestone (Hopkins and Thompson, 2003). The large amount of evaporites present within the Harrisburg Member indicates that the climate was arid, the influx of water supply from the ocean was routinely restricted, and water was generally receding from the area (Fillmore, 2000). This restriction of water and associated unstable environment is also reflected in the member’s fauna content. Within the Harrisburg Member, there are fossils of pelecypods, gastropods, some ostracods, and generally fauna that were “hardy individuals tolerant of a greater range in environmental conditions” (Hopkins and Thompson, 2003).
The Toroweap and Kaibab Formations compose the uppermost layers of the Grand Canyon’s walls (Figure 1), forming a lower slope-forming band and upper cliff-forming band at the canyon’s rim (Figure 16). Both rock units were deposited during a time of quiescent tectonic activity. During the time of their deposition, sea level rose and fell in a series of cycles influenced by the waxing and waning of glaciations occurring to the south. As sea level rose during the lowermost “gamma” member of the Toroweap, the Seligman Member, sand dunes that were left behind during the time of Coconino deposition were covered and the landscape began to change from terrestrial- to shallow marine-dominated. By the time of the “beta” Brady Canyon Member, sea level across the Grand Canyon area and greater Four Corners region had reached its peak, covering half of Utah and parts of northwestern Arizona. During deposition of the Woods Ranch Member, sea level began a step-wise fall and marine waters gradually receded to the west. Following Woods Ranch Member deposition, subaerial exposure of the region allowed for erosion to take place, indicated by an unconformity that separates the Toroweap Formation from the Kaibab Formation. The Kaibab’s gamma member marks a new phase of marine inundation of western North America and initiates another cycle of rising and falling sea level. The only significant difference between this new cycle of sediment deposition and that of the cyclic deposition of the Toroweap is that, during deposition of the Kaibab’s beta member, sea level rose enough to cover an even greater area, allowing for most of Utah, Arizona, and the lower half of New Mexico to be almost completely under water. However, depositional environments associated with both formations were unstable and ever fluctuating, indicated by the relative lack of fossil preservation.
Mesozoic Sedimentary Rocks of the Grand Canyon and Grand Staircase
The Chinle and Moenkopi Formations (by Meagan Redmond and Ken Bevis)
The Moenkopi and Chinle Formations comprise two stratigraphically related sedimentary rock units that formed during the Triassic Period of the Mesozoic Era, the time interval from 245 to 208 million years ago (Stokes, 1986). However, “no Mesozoic rocks are exposed within the Grand Canyon” (Colbert, 1974) and because they are not Paleozoic in age, the Moenkopi and the Chinle Formations do not usually appear within the Grand Canyon’s sedimentary sequence. Cedar Mountain (Figure 19) and Red Butte (Figure 20), relatively nondescript hills located near Desert View, AZ in the southeast corner of Grand Canyon National Park and south of Tusayan, AZ respectively, offer the only outcrops of the Moenkopi and Chinle formations within the vicinity of Grand Canyon National Park. Yet Mesozoic rock formations do represent an important part of the Colorado Plateau’s geologic history, and they do crop out in areas near the Grand Canyon, making the formations important to include when discussing the geology of this region. For this reason, I include these early Mesozoic rock units in Figure 1. Alternating, stair-stepped, cliff-slope exposures of these Mesozoic rocks primarily occur to the north in an area aptly called the Grand Staircase, preserved for our viewing pleasure in such parks as the Escalante-Grand Staircase National Monument, Glen Canyon National Recreation Area, and Vermillion Cliffs National Monument. The Moenkopi Formation is believed to have been formed during the Early (and possibly early Middle) Triassic Period, while it is inferred that Chinle deposition occurred during the Late Triassic (Reppening and Cooley, 1969), the two rock units being separated by a significant region-wide unconformity. To understand how these rock units were formed, it is first important to explore what was happening tectonically during the time of their deposition and to understand what the Colorado Plateau region’s landscape, in which the Grand Canyon is now situated, looked like (what were the rock unit’s paleogeographic and paleoenvironmental settings during the time of their formation?).
Figure 19. The only outcrops of the Triassic Moenkopi and Chinle Formations within the boundaries of Grand Canyon National Park occur at Cedar Mountain, in the southeast corner of the park; here viewed from the observation deck at Desert View.
Figure 20. Red Butte lies a few miles south of Tusayan, AZ just east of AZ Hwy 64; capped by an old lava flow originating from the San Francisco Volcanic Field, it preserves a section of Triassic age sedimentary rocks of the Moenkopi and Chinle Formations.
The Triassic period is the first of three time intervals that comprise the Mesozoic era (the second being the Jurassic, which is followed by the Cretaceous). During the Mesozoic Era, which lasted about 183 million years, the world was changing as the supercontinent Pangaea fragmented and the Earth’s distribution of landmasses and ocean basins began to resemble what we see today. The initial shifting of tectonic plates during the Triassic affected “western North America only indirectly because the region was far removed from the rifting margins of the breakup” (Fillmore, 2011). Yet the shifts still had some impact on the Colorado Plateau region as continental masses split apart to form constricted seaways that ultimately grew to form ocean basins, and marine transgressions and regressions came and went across the landscape, leaving their mark as the sedimentary rocks of the Mesozoic (Fillmore, 2011).
Throughout the Triassic period, the Ancestral Rocky Mountains (and the main Ouachita-Marathon mountain belt further to the southeast) “slowly disappeared under the assault of erosion and regional subsidence”, but “endured as low hills….., sporadically submitting small amounts of fine-grained sediment to the west” (Fillmore, 2011). This sediment contributed to the material deposited during the Triassic and therefore to our two Mesozoic formations. Triassic sedimentation can be “characterized by a variety of marine and non-marine deposits” (Stokes, 1986), suggesting the importance of marine transgression and regression and the lateral shifting of transitional marine depositional environments during this period. A more significant and long-lasting shift from a primarily marine to a more non-marine environment likely happened between the Early Triassic Period, when the Moenkopi Formation was deposited, and the Late Triassic, when the Chinle Formation was deposited.
“In a continuation of Late Permian geography, the west margin of Pangaea (proto-western North America) was initially an ocean, although a volcanic island [arc] lay an unknown distance offshore to the west” (Fillmore, 2011), the McCloud Arc of Blakey (2014). The volcanic island arc and the North American plate finally collided in the Triassic to form the Sonoma highlands (or the Sonoma orogenic belt) to the west in central Nevada and a parallel foreland basin in the Four Corners area (Fillmore, 2011). The accreted arc terrane and foreland basin cut off a portion of the west-lying ocean, creating an interior basin occupied by an arm of the ocean opening to the northwest known as the “Moenkopi Sea” (Figure 21) and allowing for the Moenkopi Formation to accumulate (Fillmore, 2011). The climate at the time of Moenkopi deposition was likely still “arid” and prevailed “across the Four Corners region during this time” (Fillmore, 2011).
Figure 21. Early Triassic paleogeography of western North America during early (A) and later (B) depositional members of the Moenkopi Formation; original maps are by Ron Blakey and presented in Blakey and Ranney (2008) and Blakey and Middleton (2012).
The Moenkopi Sea was likely the remnant of a more extensive marine transgression which had flooded the entire passive margin of western North America and covered much of Utah and Arizona during the earlier Permian period, depositing the Kaibab Formation (Blakey and Middleton, 2012). In the Triassic, remnants of the Ancestral Rocky Mountains still lay the east of the interior basin, which since Late Paleozoic time had “endured and continued to shed sediment westward and influence drainage patterns” (Fillmore, 2000). Early Triassic Moenkopi deposition in areas nearer the Grand Canyon was overall the result of the interplay between marine and terrestrial paleoenvironments (Figure 21); the shallow Moenkopi Sea that dominated western Utah and northwestern Arizona and a coastal plain setting dominated by aridland braided rivers occupied eastern Utah and Arizona. Stokes (1986) describes the general paleoenvironment of the Moenkopi Formation in Utah as “mainly marginal marine mud flats”. The Moenkopi Sea likely did not extend into the area that is now the Grand Canyon, but rivers flowing into this sea probably did (from the remnant Uncompaghre highlands to the east and Sonoma highlands to the west), depositing correlative sediment with a more terrestrial signature. This depositional scenario played out over the course of Moenkopi accumulation, with earlier Grand Canyon area environments dominated more by marine conditions (Figure 21a), and later accumulation dominated by terrestrial fluvial conditions as the seas regressed from the region (Figure 21b).
The Sonoma mountain belt and volcanic arc that formed during Moenkopi time ultimately formed the western margin of the Late Triassic Chinle foreland basin in what would become Nevada and Arizona (Fillmore, 2000 and 2011). Trapped between the Sonoma highlands to the west and the ever diminishing Uncompaghre Uplift of the Ancestral Rockies to the east (as well as the Ouachita-Marathon mountains further to the southeast), the interior basin gradually filled with sediment. As millions of years passed, the interior basin was transformed into a terrestrial environment dominated by a sprawling, lowland river drainage system (similar to the Amazon basin of today) flowing generally to the northwest in the direction of the retreating Moenkopi Sea (Figure 22). This shift in paleo-environmental conditions resulted in the Chinle Formation’s deposition, which now overlies the Moenkopi (Blakey and Middleton, 2012). Fillmore (2000 and 2011) describes much of the deposition during the Late Triassic Chinle Formation as being “floodplain” and “river channel deposits”. Fossil plants and pollen preserved in the Chinle’s fluvial sediments record a dramatic shift in paleoclimate toward “wet subtropical” and “warm, strongly seasonal monsoonal” conditions influenced mainly by an equatorial paleolatitude of 5-15° N (Fillmore, 2011), while the volcanic arc surrounding the westernmost side of the Chinle basin would “episodically” erupt to leave evidence of its activity in the form of copious volcanic ash within some of the rock unit’s members.
Figure 22. Late Triassic paleogeography of western North America during early and later depositional members of the Chinle Formation; original maps are by Ron Blakey and presented in Blakey and Ranney (2008) and Blakey and Middleton (2012).
Fillmore (2011) indicates that Chinle Formation deposition gradually shifted from a high energy, braided stream dominated system (Figure 29a) to low energy, “sluggish river and swampy lake deposits” (Figure 29b) as the western Sonoma volcanic arc, southeastern Ouachita-Marathon belt, and eastern Ancestral Rockies continued to supply volcanic material and sediment throughout the Late Triassic. By the end of the Triassic, the Uncompaghre and Ouachita-Marathon highlands had all but disappeared, although Sonoma highland sedimentation continued into the Early Jurassic. Evidence of prevailing paleoenvironmental conditions changes slightly depending on where one examines the Moenkopi and Chinle Formations, but the basic conclusion remains the same. During the Moenkopi Formation’s deposition, the region would have been arid and largely covered by a northwest to southeast transitional setting from shallow sea to coastal plain (Figure 21), while during the Chinle Formation’s deposition that sea regressed and the area became terrestrial, covered with braided to meandering riverine environments and large, shallow, swampy lakes dominated by a wet, subtropical climate (Figure 22).
Although the Moenkopi and the Chinle Formations do not crop out in the Grand Canyon itself, they do appear extensively in areas just north and east of the canyon. If isolated remnants of the Moenkopi and the Chinle occur south and east of the canyon, and both units blanket areas to the north and east, then it is very likely they once accumulated in the same area as the Grand Canyon. According to Colbert (1974), “It seems obvious that rocks representing the three Mesozoic periods once were part of the Grand Canyon sequence…Their present absence from the canyon is merely the result of extensive erosion since Cretaceous time.” The removal of the Moenkopi and Chinle Formations from a broad swath of the Grand Canyon region suggests area-wide uplift in part controlled by the Late Cretaceous-Early Tertiary Laramide Orogeny that built the modern Rocky Mountains to the east and by enhanced downcutting of drainage systems across the region caused by Late Tertiary and Quaternary Basin and Range extension and downdropping to the west. The discussion which follows examines these two associated formations as they are described within the Escalante-Grand Staircase area of Utah, where they are widely distributed and well studied (Sprinkel, et al., 2003). The Moenkopi and the Chinle have paired stratigraphic relationships and it is best to describe their members in ascending order, so that one can better see the sequential facies changes within both formations (from Moenkopi through Chinle) through time, with the Moenkopi being deposited first and the overlying Chinle formation being deposited later.
The Moenkopi Formation throughout its extent on the Colorado Plateau is characterized as being deep red in color, fine grained, exhibiting many thin, horizontal beds, and has a tendency to weather into multiple steep slopes and benches (Fillmore, 2011). The Moenkopi is “dominated by thin alterations of sandstone, mudstone, and shale, which contribute to its horizontally striped appearance” (Fillmore, 2011). These typical characteristics stand out for most of this formation’s members. Figure 23 provides an illustrative example of these features from rock outcrops along the Colorado River in the Glen Canyon National Recreation Area near Lee’s Ferry, AZ.
Figure 23. The Moenkopi Formation exposed in a monoclinal fold at the base of the Vermillion Cliffs near Lee’s Ferry, AZ as viewed from the Spencer Trail; the Chinle Formation’s basal member, the Shinarump Conglomerate, forms the buff-colored cliff resting on the brick-red layers of the Moenkopi.
The Moenkopi Formation can be divisible into 4-6 formal members, depending on where the formation occurs and what park or monument it is being described in (Sprinkel, et al., 2003). The unit is 910-1150 feet thick in the Grand Staircase section of the Grand Staircase-Escalante National Monument in Utah, but only 440-730 feet thick in the Escalante Canyons section to the northwest, indicating that the formation thickens westward (Sprinkel, et al., 2003). Fossils within the Moenkopi represent “a mix of terrestrial and marine taxa,” such as plants, brachiopods, fish, reptiles, amphibians, and reptile footprints (Sprinkel, et al., 2003). This mix of terrestrial and marine fossils within the members of the rock unit indicates that the Moenkopi Sea underwent multiple short-lived transgressions and regressions. The open marine water source of the Moenkopi was located to the northwest and deposition occurred on a “west-sloping” coastal plain where “the very low relief of the coastal plain allowed the smallest sea level changes, or even tidal fluctuations, to shift the shoreline [of the Moenkopi Sea] tens of kilometers…” (Fillmore, 2011). Evidence of these fluctuations is present within the Moenkopi Formation’s members.
In the Grand Staircase section at the monument’s southwestern end, the Moenkopi is divided into six members. In ascending order, the members are the Timpoweap, the Lower Red, the Virgin Limestone, the Middle Red, the Shnabkaib, and the Upper Red. In the Escalante Canyons section to the northeast, however, the formation is divided into only four members which in ascending order include the Black Dragon, the Sinbad, the Torrey, and the Moody Canyon. From northeast to southwest across the monument, most of these members correlate to one another, although they have different names. In ascending order, the correlations follow: the Black Dragon – no Grand Staircase Member equivalent; the Sinbad – correlates with the Timpoweap in the Grand Staircase; the Torrey – correlates with the Lower Red, Virgin Limestone, and Middle Red; and the Moody Canyon – correlates with the Shnabkaib and the Upper Red members in the Grand Staircase (Sprinkel, et al., 2003). The Black Dragon Member, lowermost and oldest member of the Moenkopi Formation, is about 40 feet thick, and consists of laminated to thinly bedded siltstone and silty sandstone. This unit is ripple-marked and intercalated with thin-bedded, fine-grained, micaceous sandstone (Sprinkel, et al., 2003), features that indicate deposition in a transitional, likely nearshore marine environment. The member shows no evidence that any major tectonic activity was occurring during the time of its deposition (Fillmore, 2011). The base of this member, which directly overlies Permian formations, consists of accumulations of chert or quartz pebble conglomerates, and some gypsum. Gypsum is an evaporate mineral, which is “precipitated from mineral-rich waters [within restricted basins] when evaporation rates exceed the influx of water” (Fillmore, 2000). Its presence within the member suggests that the Moenkopi Sea still largely dominated the Utah area, although it was largely cut off from its main source of water (the ocean, which lay to the west). According to Sprinkel, et al. (2003), “it is thought that the [Black Dragon Member] was deposited on a mudflat or tidal flat and in associated lagoons,” on the eastern edge of the Moenkopi Sea (Figure 21a).
Overlying the Black Dragon Member is the Sinbad (or Timpoweap) Member. Fillmore (2011) describes the contact between the Black Dragon and the Sinbad as gradual, from red deltaic deposits upward into fossiliferous limestone. This points to a “progressive rise in sea level” associated with marine transgression, the first of four main changes in sea level that the area would experience in Moenkopi time. The limestone deposits within the Sinbad “mark the highest level that the sea reached during the Triassic period” (Fillmore, 2000); Figure 21a is representative of this member’s paleogeography. In the Grand Staircase-Escalante National Monument, the member is 55 feet thick in the Circle Cliffs area, “fossiliferous,” and is comprised of “limestone, dolomite, and calcareous siltstone” (Sprinkel, et al., 2003). It intertongues with the Black Dragon member below and the Torrey Member above, indicating a close stratigraphic relationship and proximity of adjacent transitional marine environments. The limestone beds of this member weather out as bench-forming ledges and are typically thin (Sprinkel, et al., 2003). The Timpoweap Member, which correlates with the Sinbad and is located at Buckskin Mountain within the monument, resembles its partner in color (light brown to yellow gray), but consists of carbonate rocks, sandstone, chert breccias, and siltstone. Both layers were deposited in a marine environment which “represented a transgression,” but the thin, even beds of the Sinbad in the Circle Cliffs are an indication of “quiet water deposition,” while the chert breccias in the Timpoweap “might indicate deposition in more turbulent waters” (Sprinkel, et al., 2003). The western side of the monument was occupied by deeper water, while the Timpoweap formed in shallower, more wave and/or tidally influenced waters on the eastern side (Figure 21a).
The Lower Red Member rests on the Timpoweap Member in the Grand Staircase area of the monument and correlates with the lower part of the Torrey Member in the Circle Cliffs area (Sprinkel, et al., 2003). It is 140-220 feet thick and thickens westward along with the rest of the Moenkopi, and it consists of “red to chocolate-brown interbedded and thin-bedded siltstone and fine-grained sandstone” (Sprinkel, et al., 2003). This member tends to weather into slopes, but with alternating slight ledges; the ledges forming platy beds that display abundant ripple marks, indicating a transitional marine environment. Sprinkel, et al. (2003) infers that the “Lower Red Member was deposited on a tidal flat traversed by meandering streams.” The presence of meandering streams suggests slight regression of the shoreline or a prograding deltaic environment and it appears that sea level in this area may have shallowed at this time. The Virgin Limestone Member overlies the Lower Red Member in the Grand Staircase area and correlates with the main body of the Torrey Member in the Circle Cliffs area of the monument. It is 10-30 feet thick and thickens to the west. The member is described as being “conspicuous because it is ledge forming” (Sprinkel, et al., 2003) and it consists of yellow-brown sandstone, siltstone, and limestone, and it is not fossiliferous (Sprinkel, et al., 2003).
The Torrey Member overlies the Sinbad Member in the Circle Cliffs area of the monument. It is 240-310 feet thick and like other members in the area, it thickens to the west. “The upper contact with the Moody Canyon Member is intertonguing and sometimes difficult to place” (Sprinkel, et al., 2003). The member is comprised of very fine to fine-grained sandstone and silty sandstone that tend to form thin to medium-bedded ledges, cliffs, and slopes. The “slope-forming constituents” are generally very micaceous, and the “ledge formers” display “ripple marks, load casts, drag structures and animal trails” (Sprinkel, et al., 2003). Features such as ripple marks and drag structures suggest flowing water, which could indicate stream transport and Sprinkel, et al., (2003) have inferred that the “Torrey Member was deposited in a deltaic and shoreline environment.” This transition from marine waters of moderate depth in the Sinbad Member to delta and shoreline environments in the younger Torrey Member indicates continued regression and/or progradation of a terrestrial environment. As Fillmore states, “This change heralds the retreat of the shallow sea and the resurrection of the delta” (Fillmore, 2011). From a paleogeographic standpoint, the Moenkopi Sea at the time of the Torrey Member’s deposition would have appeared shrunken, with some of eastern Utah now subaerially exposed (Figure 21b).
The Middle Red Member of the Moenkopi overlies the Virgin Limestone Member in the Grand Staircase and is described as “the thickest of the Moenkopi members in this section and the least resistant to erosion” (Sprinkel, et al., 2003). It has a slope-forming nature, and it is 280-400 feet thick. The member consists of “medium-brown to chocolate-brown mudstone and siltstone” along with fine-grained silty, ripple-marked sandstone. Gypsum veinlets within this member indicate a correlation with the upper section of the Torrey Member which also contains gypsum. However, the gypsum content in this member “increases upward,” which could indicate an increased evaporation rate and/or a decrease in the influx of water as the Moenkopi Sea regressed. The upper section of this member is believed to correlate with the lower section of the Moody Canyon Member in the Escalante Canyons area of the monument (the Moody Canyon correlates with the Shnabkaib Member in the Circle Cliffs, which sits just above the Middle Red Member). In an indirect way, this makes sense, because the Moody Canyon Member is resting on the Middle Red Member, and one is just transitioning into the other (Sprinkel, et al., 2003).
The Shnabkaib Member overlies the Middle Red Member in the Circle Cliffs area and is 150-250 feet thick. “It is a ledge- and slope-forming unit,” contains silty gypsum and primarily consists of sandstone and “red and green-gray siltstone” (Sprinkel, et al., 2003). In order for gypsum to form, the Shnabkaib Member was “probably deposited in restricted embayments of a sea surrounded by low tidal-flat and mud-flat areas” where a “fresh supply” of salty ions could be deposited regularly (Sprinkel, et al., 2003), indicating that the Shnabkaib Member was deposited in the receding Moenkopi Sea (Figure 21b).
The uppermost correlative members of the Moenkopi within Grand Staircase-Escalante National Monument are the Upper Red Member and the Moody Canyon Member. The Upper Red Member in the Grand Staircase section is 90-180 feet thick, while the Moody Canyon Member in the Escalante Canyons is 200-300 feet thick. Both members are generally red brown in color, and both contain interbedded siltstone, sandstone, or mudstone. Fillmore (2011) reports that the Moody Canyon Member’s, “siltstone units occasionally display ripple marks, and [the] mudstone is mostly laminated.” Both of these members are generally slope-forming, although the Upper Red Member is divided into an upper half weathering into ledges and a lower half into slopes. The inferred depositional environment for both of these members is very similar: “The Upper Red Member is mainly a tidal flat deposit”, while the Moody Canyon Member is “believed to have been deposited in shallow quiet water, such as in ponds or lagoons on a tidal flat” (Sprinkel, et al., 2003).
From the beginning of the Moenkopi Formation onward, we have seen what appears to be a gradual marine regression to the northwest from the Utah area punctuated by one or more minor transgressions, but no significant tectonic activity. However, it is sediments within the Chinle Formation, deposited after the Moenkopi Formation, that show evidence of a substantial transition from a marine-dominated to a terrestrial-dominated environment, as well as the considerable influence of tectonic activity. The Chinle Formation overlies the Moenkopi Formation, however, “the contact between the Lower Triassic Moenkopi and the Upper Triassic Chinle Formation is an unconformity marked by the absence of Middle Triassic strata”, an erosional surface which “represents an omission of up to 10 million years” (Fillmore, 2011). The top of the Moenkopi contains distinctive northwesterly oriented furrows, presumably cut by stream erosion and into which the initial deposition of Chinle material occurred; however, how much of the upper Moenkopi was remove by erosion is not well understood.
The Chinle Formation is described in Colorado Plateau literature and in the Grand Staircase-Escalante National Monument as being lithologically “heterogeneous” and extremely colorful (Fillmore, 2011). Fossils within the Chinle Formation include petrified wood (which is the most abundant), as well as fossils of fish, bivalves, amphibians, dinosaurs, and dinosaur tracks” (Sprinkel, et al., 2003); all indicative of a terrestrial setting. The rock unit’s sediments consist of “varying amounts of fluvial and lacustrine interbedded sandstone, mudstone, claystone, siltstone, limestone, gritstone, and conglomerate” (Sprinkel, et al., 2003). In the Grand Staircase area of Grand Staircase-Escalante National Monument, the rock unit is 500-930 feet thick and 425-750 feet thick in the Circle Cliffs area. It’s members in ascending order include; the Temple Mountain, the Shinarump, the Monitor Butte, the Petrified Forest, the Owl Rock, and the Church Rock; although some of these members are not present in every location within the monument (Sprinkel, et al., 2003). Figure 24 nicely displays exposures of the Chinle Formation at the mouth of the Paria River canyon near Lee’s Ferry, AZ.
Figure 24. Exposures of the Chinle Formation at the mouth of the Paria River canyon near Lee’s Ferry, AZ as viewed from the Spencer Trail; the brightly colored layers in the center of the photograph form the Chinle’s distinctive Petrified Forest Member.
The Temple Mountain Member only occurs within the Circle Cliffs area of the Grand Staircase- Escalante Monument, and on the Colorado Plateau, “the Temple Mountain Member…is mostly confined [well to the north] in the San Rafael Swell” (Fillmore, 2011). It is 50 feet thick, and “it forms slopes and ledges between the Moody Canyon Member of the Moenkopi Formation and the resistant cliff-forming sandstones of the Shinarump Member” (Sprinkel, et al., 2003). The member’s most distinctive characteristic is its “mottled” appearance, with “colors of purple, red, brown, and gray-white” (Fillmore, 2011). The member consists of well-cemented siltstone containing immature quartz, along with some evidence of root casts and some considerable amount of dark minerals, such as tourmaline, which indicates that the source of the siltstone was not erosion of the Moenkopi. It is believed to be “mostly a fluvial deposit that was highly weathered before the Shinarump and later sediments were deposited” (Sprinkel, et al., 2003). These features, as well as the disconformity with the underlying Moenkopi Formation, suggests that at this point in time, this area had been transformed into a terrestrial environment. Fillmore (2011) argues that deposition was controlled by a brief “reinvigoration” of streams flowing through the area during the Middle Triassic in response to “a drop in sea level and probable uplift southeast of the Colorado Plateau”; these rivers carved out “an early incarnation of the Chinle drainage network”, the channels that would eventually receive Shinarump and Temple Mountain deposition.
The Shinarump Member overlies the Temple Mountain, but unlike the Temple Mountain Member which is not widespread, “The Shinarump Member holds the basal position of the Chinle over much of the Colorado Plateau, particularly in northern Arizona and southern Utah” (Fillmore, 2011); its resistant sands and gravels forming a distinctive buff-colored cliff band throughout the region (Figure 23). The unit primarily consists of relatively high energy stream-channel deposits: “Streams meandered northwestward [away from North American continental sourcelands] across an old Moenkopi surface where local thick soils [had] developed” and “locally cut deep channels into this surface” (Sprinkel, et al., 2003). Deposition at this time was dominantly fluvial, filling an internally drained basin (Figure 22a): “During Chinle time, southwest Utah may have been an area of interior drainage, with the Shinarump streams depositing their loads of suspended materials over a broad flat that had developed on the old Moenkopi surface” (Sprinkel, et al., 2003). Fillmore (2011) notes that “Shinarump deposits mark the paths of energetic braided rivers.” Throughout the Utah area, the Shinarump can largely be characterized as having trough cross-bedded sandstone, siltstone, mudstone, and “volcanic-derived pebbles,” indicating that the volcanic arc associated with the ongoing Sonoma Orogeny that initially developed during Moenkopi deposition continued through Shinarump time. (Sprinkel, et al., 2003). As Fillmore (2011) describes, “…numerous tributaries from the remnant Ancestral Rockies and the chain of volcanic mountains to the west and southwest contributed significant amounts of sediment” to the braided rivers of the Shinarump.
The Monitor Butte Member overlies the Shinarump and underlies the Petrified Forest Member, forming gradual intertonguing contacts with its neighbors. It forms stepped-ledge slopes and because of its dominant gray coloration it is commonly included in the Shinarump. It contains gray-green mudstones with interbedded sandstones and siltstones, and it contains petrified wood fragments in some places and shales and coal bed deposits in others, probably related to the prevailing tropical climate and thickly forested conditions of the great paleovalley system in which it was deposited. “Generally, sandstone beds are more common in areas overlying channels of the Shinarump Member, indicating that stream courses established by Shinarump streams were maintained during Monitor Butte time” (Sprinkel, et al., 2003). However, Fillmore (2011) describes that the “The fine sandstone, mudstone, and shale of the Monitor Butte are a striking contrast with coarse underlying strata, reflecting an equally dramatic shift in depositional setting.” The volcanoes that surrounded the Chinle basin to the west and the southwest were becoming more active, releasing “huge quantities of fine ash” into the air that were then carried by wind into the Chinle basin (Figure 22b) and the Moenkopi Sea at this point in time had completely disappeared far to the northwest. As the ash accumulated, rivers became clogged and an overall lower energy system of “lakes, rivers, and wetlands evolved” (Fillmore, 2011). The abundant organic matter that accumulated in these long-lived lakes would later form bogs and eventually the petrified wood and coal deposits that we find within the member (Sprinkel, et al., 2003).
Sprinkel, et al. (2003) reports that “whereas the lower three members of the Chinle generally form cliffs or stepped-ledge slopes, the upper three members generally form slopes, steep slopes, and higher in the section, ledgy slopes.” The Petrified Forest Member, the first of these three upper units, overlies the Monitor Butte. Its most “conspicuous” characteristic is its eye-catchingly bright coloration (Figure 24), which is usually differentiated in bands. An interesting sedimentological feature of this member, along with an abundance of petrified wood, is its quantity of bentonite. “Bentonite, or montmorillonite, is a type of clay produced from the decomposition (devitrification) of volcanic ash. It swells dramatically when wet” (Sprinkel, et al., 2003). “The Petrified Forest Member was deposited on a fluvial plain largely as overbank deposits,” indicating that streams formed in Shinarump time persisted through Petrified Forest time (Figure 22b). The expansive floodplains were “dotted with lakes and ponds” (just as it had been during deposition of the Monitor Butte Member), and at times, as the sediments accumulated, volcanic ash settled on the plain, which was altered to bentonite clay and locally into siliceous brown nodules” (Sprinkel, et al., 2003). Fillmore (2011) indicates that the continued presence of volcanic ash in the Petrified Forest Member was likely caused by the “continued flood of airborne volcanic ash into the basin from the erupting volcanoes to the southwest.” This volcanic activity was sourced in the active volcanic arc of the Sonoma highlands to the southwest, likely the result of continuing subduction along the western margin of the North American continental plate. This member contains abundant fossil wood, and huge logs and coal deposits have been found (Sprinkel, et al., 2003), suggesting that during the time of Petrified Forest deposition, climate in the Chinle basin area still resembled monsoon-like, tropical conditions, allowing for abundant vegetation to grow (Fillmore, 2000).
The Owl Rock Member overlies the Petrified Forest Member to a thickness of 150-250 feet (Sprinkel, et al., 2003). This member is composed of thin lenticular beds of green limestone interbedded with red, brown, and green-gray sandstone and mudstone. The presence of bentonite in this member suggests that the upper part of the Petrified Forest and the lower part of the Owl Rock form a gradational contact, although its decreasing abundance suggests “that volcanic activity in the southwest was waning” (Fillmore, 2011). “Regionally, the Owl Rock Member reflects continued subsidence of the Chinle basin, centered around the Four Corners area, but with a much-reduced sediment supply from surrounding highlands” and its paleogeography depicted by Figure 21b can basically be compared to that of the present-day Everglades, “in which freshwater trickled northwestward to an eventual rendezvous with the sea in central Nevada” (Fillmore, 2011). In the Grand Staircase-Escalante National Monument, it was believed to be deposited “on a fluvial environment on a plain dotted with lakes” (Sprinkel, et al., 2003), while more broadly, on the greater Colorado Plateau, it was considered to be “deposited in a widespread lake and marsh system linked by the occasional wandering river” (Fillmore, 2011).
The uppermost member of the Chinle Formation is the Church Rock member. The member is described as having “orange red sandstone and siltstone” (Fillmore, 2011) and he further notes that the “Church Rock Member has been correlated with the widely scattered Rock Point Member in northeast Arizona and western New Mexico.” The similarity of lithologies and stratigraphic position shows that the member itself is widespread, and suggests a broad geographic distribution of its depositional environment. In the Grand Staircase-Escalante National Monument area, “The sediments were deposited in lakes and by streams on an alluvial plain that [still] sloped away from the Uncompahgre uplift in eastern Utah during Late Triassic time” (Sprinkel, et al., 2003); although on the Colorado Plateau more broadly, “the Church Rock and its equivalents were deposited in a variety of settings, all signaling an increasingly arid climate” (Fillmore, 2011). There are no volcanics within this member, or evidence of tectonic activity, indicating that the area was becoming stable once more.
At the beginning of the Triassic period, Pangaea was undergoing break up, and this tectonic activity had an indirect impact on the depositional environment of the Moenkopi Formation. To the west of the Pangaean supercontinent, a volcanic island arc made its final landfall with North America. This collision and subduction created the Sonoma Orogeny in central Nevada and its accompanying Sonoma highlands, followed by active arc volcanism along the North American southwest coast. Consequently, a foreland basin developed between the Sonoma highlands and the much subdued Ancestral Rocky Mountains, leaving an interior basin occupied by the Moenkopi Sea over most of Utah. Tectonic activity generally remained mild throughout the time of the Moenkopi’s deposition, but by the time of the Chinle Formation’s deposition, the Moenkopi Sea had almost entirely receded to the northwest, leaving behind an immense, northwestward draining paleoriver valley system in the basin between the Sonoma highlands and Ancestral Rockies. Tectonic activity ramped up during Chinle Formation deposition, with evidence of further subduction and growth of a volcanic arc to the southwest represented by copious volcanic ash within some of the Chinle’s later members.
Based on lithological and paleontological evidence, the Moenkopi Formation is inferred to have been deposited in moderately deep to shallow marine and transitional marine environments, and as sea level steadily regressed, they were gradually replaced by terrestrial environments upward within the formation. Between the Moenkopi and Chinle Formations there is an erosional unconformity, indicating that marine conditions completely disappeared, allowing substantial subaerial exposure, weathering, and erosion of the top of the Moenkopi. Dominated by fluvial deposits that contain abundant plant and animal fossils, the Chinle Formation is believed to have formed in a wetter climate than those that preceded or followed its deposition. Starting with the deposition of the Shinarump Member of the Chinle Formation, accumulation of sediment was renewed as vigorous streams and rivers began to flow across the previously deposited Moenkopi Formation. In the Monitor Butte Member, lakes and ponds began to dominate the interior basin as volcanic ash from the west clogged earlier formed rivers, creating abundant vegetation and adding to the water supply in the area. From the Petrified Forest Member through the Owl Rock Member, previously established drainage in this area prevailed, although volcanic activity was significant. Finally, in the Church Rock Member, volcanic activity had waned, and the environment began to transition into drier terrestrial conditions, with streams still present in western areas of the basin.