Phanerozoic Evolution of the Sedimentary Cover of the North American Craton

Chapter 2

Phanerozoic Evolution of the Sedimentary Cover of the North American Craton Peter M. Burgess School of Environmental Sciences, University of Liverpool, Liverpool, United Kingdom

Chapter Outline Introduction Definition of a Craton Tectonic Elements of the North American Craton The Canadian Shield The Cratonic Platform Intracratonic Basins Cratonic Margins Controls on Evolution of the Cratonic Cover Eustasy Extension and Thermal Reequilibration Intraplate Stress Dynamic Topography Related to Subducting Slabs Dynamic Topography Related to Supercontinent Cycles, Mantle Convection Cells, and Plate Amalgamation and Dispersal Magmatic Controls Phanerozoic Evolution of the Cratonic Platform Cover The Sauk Sequence (Late Precambrian to Early Ordovician)

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The Tippecanoe Sequence (Middle Ordovician to Early Devonian) Kaskaskia Sequence (Mid-Early Devonian to Late Mississippian) Absaroka Sequence (Late Mississippian to Early Jurassic) Zuni Sequence (Middle Jurassic to Early Paleocene) Tejas Sequence (Late Paleocene to Present) The North American Intracratonic Basins The Michigan Basin The Illinois Basin The Williston Basin The Hudson Bay Basin Summary Acknowledgments References

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INTRODUCTION The North American craton extends north to south for >5500 km from midcontinental United States to the Arctic Islands, and for >4700 km west to east, from western Canada to eastern Greenland. Given the geographical extent of this geological feature, the diversity of strata preserved on it, and the volume of data produced by >100 years of study, it is perhaps surprising that the Phanerozoic history of the craton is often viewed as, by definition, “tectonically stable.” The history recorded in the cratonic Phanerozoic strata suggests a rather more complex evolution. The craton was sometimes tectonically quiescent, sometimes flooded almost entirely to produce extensive shallow epeiric seas like nothing existing on Earth today; sometimes it was peneplained, sometimes being uplifted, warped, and tilted with material being stripped off by erosion at geologically rapid rates. This chapter describes some of the diverse structure and strata preserved on the craton and provides an interpretation of its history. The chapter gives a regional overview of cratonic platform strata across North America, and also focuses on the four largest intracratonic basins to provide a history of cratonic deposition and deformation. The geological history summarized here lends support to a view (e.g., Sloss and Speed, 1974) of the North American craton as a dynamic geological entity with a varied and sometimes spectacular evolution that can enhance our understanding of both Earth’s surface and mantle processes.

The Sedimentary Basins of the United States and Canada. https://doi.org/10.1016/B978-0-444-63895-3.00002-4 © 2019 Elsevier B.V. All rights reserved.

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40  The Sedimentary Basins of the United States and Canada

DEFINITION OF A CRATON In simple terms a craton is a stable, strong, unyielding area of lithosphere. However, the term is difficult to define in more detail (Leighton, 1990) because cratonic areas usually have a history of intense Precambrian deformation, followed by relative tectonic quiescence during Phanerozoic time. This history reflects the ongoing tectonic process of continental accretion; parts of the craton now safely ensconced within the continent were, in Proterozoic time, areas of tectonic activity on the craton margin. Likewise, currently active continental margins may in the future form part of the stable cratonic area when the next phase of orogenesis forms a new accreting, active margin outboard of the current margin. Conversely, areas currently within the craton may become zones of renewed plate-margin tectonic activity due to rifting and oceanic spreading, which in the most extreme case could split the current craton into two or more separate cratonic blocks. The term craton was first used in the mid-20th century, during development of ideas about geosynclines, to contrast the relatively young, active continental margins with the more ancient, apparently less active continental interiors. Problems arose when geologists attempted to identify features marking the boundary between craton and geosynclines and realized that in terms of stratal thickness and characteristics, the difference is difficult to define (Sloss, 1988a). Progress in understanding plate tectonics resolved this problem by allowing a distinction based on the underlying subsidence mechanism. Areas formerly termed geosynclines were redefined as rifted margins undergoing thermal subsidence, or active margin basins formed by a variety of subsidence mechanisms including rifting and lithospheric loading. The craton was then assumed to be the area of old, strong lithosphere beyond the reach of these marginal tectonic mechanisms (Sloss, 1988a). For the purposes of this chapter, and bearing in mind that the preceding is a time-limited definition, the craton in North America is considered to be the area underlain by Precambrian basement that has not been subject to plate-margin processes during the Phanerozoic (Fig. 1). This does not mean that the craton has not been subjected to significant tectonic uplift, subsidence, and deformation, but simply that the tectonic processes affecting cratonic interiors are epeirogenic, and so do not result directly as a consequence of active plate margin processes.

TECTONIC ELEMENTS OF THE NORTH AMERICAN CRATON Subdividing the craton into different elements serves to illustrate the different tectonic mechanisms operating within the craton, and to illustrate progressive cratonic evolution.

The Canadian Shield A shield is part of a craton and is defined by the American Geological Institute Glossary of Geology as a large region of exposed basement rocks, commonly with a very gently convex surface, surrounded by sediment-covered platform. In North America the shield covers a large part of eastern and central Canada and also covers most of Greenland, although currently buried beneath the Greenland ice cap (Fig. 1). Shield rocks consist of predominantly granitic gneisses and greenstone belts composed of metasedimentary and metavolcanic rocks, and can be subdivided into seven provinces with different geological histories, based on spatial distribution, cross-cutting relationships, and isotopic ages (Hoffman, 1989). Did the shield persist in its present state, subaerially exposed and without any sedimentary cover, throughout the Phanerozoic? This would imply absence of tectonic uplift or subsidence in a stable, quiescent setting. However, there is evidence that much of the shield has been covered with platform strata at certain times during the Phanerozoic and has undergone uplift and erosion at other times. Sloss (1963, 1988b) pointed out that Middle and Upper Ordovician distal marine strata are in many places terminated by erosional truncation beneath overstepping Devonian strata, strongly suggesting that the Ordovician and Devonian strata covered a significantly greater portion of the shield before being removed by erosion. Other examples of similar stratal terminations are recorded by Bunker et al. (1988) and Collinson et al. (1988). Limestone xenoliths found in kimberlites on the Canadian shield are Middle Devonian in age and appear to record deposition of >750 m of Devonian and younger strata, preserved on the craton until Middle Jurassic time, and removed by erosion prior to the Late Cretaceous epoch (Cookenboo et al., 1998; Patchett et al., 2004). Nd isotope data show that the shield was a dominant source of cratonic sediment during early Paleozoic time (Patchett et al., 1999), concordant with observations from Sloss (1988b), but much of this sediment probably originated from further north than the present southern limit of the shield (Collinson et al., 1988). Patchett et al. (2004) also used Nd isotope data from strata from the Sverdrup basin of the Canadian Arctic Islands to show that much of the shield was probably covered by Ordovician to Middle Devonian carbonate units, that the northerly derived Upper Devonian siliciclastic sedimentary rocks probably covered about one-half of the shield in its western and northern portions, and that this cover was progressively removed through Mesozoic time. Late Ordovician and early Silurian conodont assemblages preserved in carbonate xenoliths from Late Jurassic to Early



Phanerozoic Evolution of the Cratonic Cover Chapter | 2  41

FIG. 1  Map of North American cratonic tectonic elements. (Modified from Bally (1980). Tectonic element names taken from Stott and Aitken (1993) and Sloss (1988b).)

Cretaceous kimberlites on southeastern Baffin Island indicates 270–305 m of Lower Paleozoic strata was removed prior to the late Jurassic (Zhang and Pell, 2014). Collectively, this evidence clearly indicates epeirogenic activity on the shield throughout the Phanerozoic leading to cycles of flooding, reemergence, and tilting.

The Cratonic Platform The term platform is defined in the American Geosciences Institute’s Glossary of Geology as a part of a continent covered by flat-lying or gently tilted sedimentary rocks, underlain by a complex of rocks that were consolidated during earlier deformations. Platform areas can be considered sedimentary basins in the sense that they are areas of sediment deposition and burial, so net accumulation, of thickness over intervals of geological time (Aitken, 1993). However, they are not

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FIG. 2  Map of stratal thickness on the North American craton. (Modified from Sanford (1987).)

basins in the sense of an area undergoing differential subsidence relative to the surrounding area of stable basement. This latter definition applies to intracratonic basins, but not to cratonic platforms. Most cratonic basins within the platform show significantly greater thicknesses of preserved strata than do the surrounding platform areas, e.g., >4.5 km of strata in the Michigan basin, compared to ~1 km of strata on the surrounding platform area (Fig. 2). Based on this there is some distinction between platform and basin in terms of thickness of preserved strata, and our understanding of the various subsidence mechanisms that may be involved is improving (see “Controls on Evolution of the Cratonic Cover” section). A significant feature of the cratonic platform is the presence of epeirogenic arches. An epeirogenic arch is an intraplatform high that subsides less rapidly than surrounding platform areas, leading to formation of relatively thin strata, or is uplifted, leading to erosion and local unconformity if then reburied. Arches are significant because they separate platform areas and appear to have acted as important elements in the paleogeography of the craton. They are also important because they provide a clear indication of epeirogenic tectonic activity within the craton throughout Phanerozoic time. Although arches may subdivide the platform at particular points in time, and areas enclosed by arches are sometimes



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referred to as intracratonic basins, such areas are still best considered as platform rather than intracratonic basin if there is a distinct difference in total preserved sediment thickness (Fig. 2).

Intracratonic Basins Intracratonic basins are areas on the craton, at some distance from the craton margin, undergoing differential subsidence relative to the surrounding area of cratonic basement. They are thus distinguished from platform areas by significantly greater thicknesses of preserved strata, e.g., >4.5 km of strata in the Michigan basin, compared to ~1 km of strata on the surrounding platform area. Four North American intracratonic basins are identified on this basis, namely the Michigan, Illinois, Williston, and Hudson Bay basins (Figs. 1 and 2). Evolution of the Michigan, Illinois, and Williston basins as discrete basins began in Late Cambrian time, suggesting perhaps that they are in some way associated with the tension along plate-scale stresses that led to breakup of a Late Precambrian supercontinent (Armitage and Allen, 2010). Initiation of the Hudson Bay basin occurred in Ordovician time, so a link with supercontinent breakup is more difficult to define. All four basins were receiving sediment throughout much of early Paleozoic time and were influenced to varying degrees by marginal and craton-wide tectonic events.

Cratonic Margins Tectonic mechanisms and events affecting the cratonic margins are varied in their duration and spatial extent, so there is commonly no clear distinction between marginal and epeirogenic processes. For example, the Colorado Plateau is a largely intact cratonic block surrounded by an area of deformation related to Sevier-Laramide compressional and late Cenozoic basin and Range extensional tectonics (Dumitru et al., 1994). Despite its apparent cratonic nature, the plateau area has been subject to active-margin processes at least since the Jurassic period (Allen et al., 2000) (see Chapters 6 and 8). Similarly, cratonic areas in eastern North America were subjected to margin processes such as flexural loading and slab-related effects from Ordovician time onwards during the Taconic, Acadian, and Alleghanian orogenies (Fig. 3) (Beaumont et al., 1987; Coakley and Gurnis, 1995). However, there is an identifiable difference in the magnitude at which these various tectonic processes operated. Stratal thicknesses increase into the cratonic margins (Fig. 4) as a result of higher rates of subsidence on the margins. The various cratonic margin basins are described throughout the rest of this book; in this chapter the influence of certain marginal tectonic episodes on cratonic strata is described.

FIG. 3  Chronostratigraphic (Wheeler) diagram showing the six cratonic depositional sequences defined by Sloss (1963). Absolute ages are those used in Sloss (1988b). (Modified from Catacosinos et al. (1990).)

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FIG. 4  Diagrammatic cross-sections of Sauk, Tippecanoe, and Kaskaskia strata illustrating thinning of the lower Paleozoic passive margin strata onto the craton, development of arch and basin geometries on the craton, and the often composite nature of the megasequence-bounding unconformities. Note that the datum in the section is the base of the Absaroka sequence. (Redrawn from Bally (1989).)



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CONTROLS ON EVOLUTION OF THE CRATONIC COVER Although studied for over 100 years, the mechanisms responsible for generating the North American cratonic sequences, both on and across the platform, and within the four intracratonic basins, remain somewhat enigmatic. This was especially true for the cratonic platform, where it has been difficult even to explain the basic observation of >1 km of Phanerozoic, predominantly Paleozoic, strata that cover the platform (Burgess and Gurnis, 1995). However, new understanding of tectonic processes has enhanced understanding significantly. In the broadest sense, there are two important candidate processes: eustatic changes in sea level, and tectonic uplift and subsidence. However, this could be considered a rather simplistic view because long-term eustatic changes are themselves driven by tectonic processes, namely changes in oceanic spreading and subduction rates, and cratonic tectonic uplift and subsidence may occur by a number of different processes. Also, both eustatic and tectonic processes can themselves trigger a cascade of depositional, denudation, and isostatic processes, the products of which can be difficult to derange and decipher. A range of these tectonic processes is described here, some of which provide plausible evidence-based explanations for development of cratonic basins and platform megasequences.

Eustasy Definition of the cratonic sequences of North America by Sloss (1963) provided the foundation for much of the sequence stratigraphic model then copied by other workers (e.g., Vail et al., 1977). Vail and coworkers chose to emphasize eustasy as the primary mechanism responsible for the relative sea-level changes necessary to explain cratonic sequence development. Indeed, much of the Paleozoic eustatic curve in Vail et al. (1977) is derived from North American cratonic stratal patterns, and the idea of cratonic platforms and their passive margins as relatively stable locations to record global sea-level history remains important even in more recent eustatic curves (e.g., Miller et al., 2005). However, there is a very basic observation that demonstrates that eustasy was not the only contributor to the relative sea-level changes recorded by cratonic sequences. Long-term eustatic oscillations certainly must have contributed to development of the transgressive and regressive sequence elements, but long-wavelength postdepositional tilting of the cratonic strata and angular sequence-bounding unconformities, both ubiquitous features of North American cratonic strata (e.g., Figs. 4 and 5), obviously require a tectonic mechanism, and cannot be explained by eustatic change alone. Applying such simple reasoning to North American cratonic stratal patterns often provides a reasonable indication of the degree of tectonic and eustatic influence on relative sea-level fluctuations. Moucha et al. (2008) address this point from a different direction. They presented results from numerical modeling of mantle convection that show how even cratonic regions undergo vertical displacements on a time scale of tens of millions of years, making it unlikely that they are sufficiently tectonically stable to permit reliable reconstruction of eustatic history. Note that in this case, a tectonic component in the relative sea-level change usually refers to local or regional tectonics causing surface uplift and subsidence. On a larger scale, ridge spreading rate influences rate of subduction, which, through dynamic topography, influences ocean volume and hence eustasy. Therefore, on large spatial scales, and at time scales of more than a few million years, distinctions between tectonic and eustatic forcing may be misleading—they are different aspects of the same process (Gurnis, 1990).

Extension and Thermal Reequilibration Stretching of the lithosphere by extensional stress creates extensional sedimentary basins, formed by a combination of initial subsidence due to active normal faulting, followed by subsidence due to postrift thermal reequilibration (McKenzie, 1978). This model has been widely applied and used to explain various intracratonic basins, including many of the North American examples (Haxby et al., 1976; Sleep et al., 1980). The Illinois basin is underlain by a Precambrian rift system (Braille et al., 1982), so there is a clear mechanism for subsequent early Paleozoic thermal subsidence. Later subsidence in the Illinois basin was more difficult to explain by extensional stress because it seemed to occur beyond the time required for thermal reequilibration of the lithosphere-asthenosphere boundary (Quinlan, 1987). However, Armitage and Allen (2010) show how protracted low-strain-rate stretching of thick continental lithosphere creates subsidence curves with a near-constant slope. These subsidence histories are similar to those of North American cratonic basins when other tectonic effects are removed (Armitage and Allen, 2010) (Fig. 6). This suggests that, as Armitage and Allen (2010) state, “Cratonic basins are therefore part of the rift–drift suite, occupying a portion of the existence field at low stretch factors and low extensional strain rate.” Forward modeling of subsidence from the Williston Basin suggest beta = 1.12 and a strain rate of 0.73 × 10−16 s−1. For the West Virginia Basin beta = 1.27 and the calculated strain rate was 1.53 × 10−16 s−1 (Fig. 6).

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FIG. 5  Cross-section through the southern midcontinent region of the United States, showing the development of the Sloss sequences in the area, the influence of the Nemaha Uplift on stratal patterns, and the thickening-to-the-west wedge of Cretaceous strata developed over slab-related dynamic topography. (Modified from Bunker et al. (1988).)



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FIG. 6  Backstripped and forward-modeled cratonic subsidence curves from Armitage and Allen (2010) showing how background subsidence in cratonic locations can be explained by lithospheric stretching and cooling. (A) Backstripped water-loaded subsidence curves from the Williston and West Virginia basins, with intervals of accelerated subsidence and uplift identified ready to remove. (B) Remnant background subsidence, with extensional model subsidence curves showing best-fit values of stretching factor beta for these curves.

As Allen and Armitage (2011) pointed out, “Cratonic basins generally lack well-developed initial rift phases, marked by arrays of extensional faults and associated graben and half-graben.” This lack of direct evidence for rifting in cratonic basins has tended to complicate interpretations of their formative mechanisms, though as Allen and Armitage (2011) also state, “this may be due in part to the poor seismic imaging of the base of cratonic basins preserved on land.” The Michigan basin is underlain by a Neoproterozoic rift system (Nunn et al., 1984) and with prolonged, perhaps episodic, low-strainrate extension, this could explain the Paleozoic subsidence history (Nunn et al., 1984). There was only indirect evidence for existence of a rift system beneath the Williston basin (e.g., Kent, 1987; Nelson et al., 1993) and the same was thought to be true of the Hudson Bay basin (Quinlan, 1987; Roksandic, 1987). More recently Lavoie et al. (2015) have argued that “Normal (or transtensional) faults imaged on seismic reflection profiles provide clear evidence for crustal extension during the deposition of the older sedimentary package,” but it seems more likely the faulting is transtensional in nature and postdates deposition of the oldest basin-fill strata (see “The Hudson Bay Basin” section for further discussion). In the Williston Basin, faults are interpreted in seismic data, but image quality is poor, and strata appear to have little or no offset across the faults (e.g., Redly, 1998, Fig. 6.5b). So both the Williston Basin and the Hudson Bay basin currently lack strong evidence for a significant episode of synrift faulting. This may change with better seismic imaging, but if not, it may be that the lack of large extensional faults imaged on seismic indicates that deformation does not localize on large discrete structures at very low strain rates (Armitage and Allen, 2010), but this perhaps requires experimental evidence. Despite the inconsistent evidence for synrift faulting, it is now the case that North American cratonic basin subsidence histories are explicable by the low strain rate model, suggesting that extensional tectonics is a dominant process in forming intracratonic basins.

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However, a few puzzles remain that extensional models cannot explain so well. Most fundamentally, what controls the distribution of extensional stress across the craton, and thus what controls the distribution of cratonic basins versus platform areas? In the latter, the same low strain rate low extension subsidence may operate, but why do these areas have even lower beta values? Inherited basement structure may play a part, but how exactly this works for basins and platforms remains unclear. Related to this, why are several North American basins approximately circular in plan view? And finally, why do the basin fill strata include phases of tilting and unconformity development that cannot be explained through simple ongoing thermal subsidence?

Intraplate Stress Stress due to various tectonic events and applied at plate margins can be transmitted laterally through the plates because the lithosphere has some rigidity and can behave as an elastic material. Such stress is known as in-plane intraplate stress (Cloetingh, 1986). Cloetingh (1986) pointed out that a change in stress of >1 kbar can produce 50 m of vertical displacement of the lithosphere. This occurs by amplifying preexisting curvature of the lithosphere, transferring in-plane stress to vertical movements. As well as causing regional uplift and subsidence, and probably more important that that process, stress transmitted through the lithosphere may also cause reactivation of previously existing structures (e.g., Braun and Shaw, 2001; Marshak et al., 2003). Reactivation of faults may regenerate old basement highs, and partly invert basins or cause new periods of extension and subsidence. Consequently, intraplate stress variations may be responsible for onset or cessation of deposition in the cratonic interior, contributing to formation of cratonic sequences (Braun and Shaw, 2001; Burke et al., 2003). Changes in stress of this magnitude are likely to occur during periods of major plate reorganization, when the location, orientation, and nature of plate boundaries change significantly, causing uplift and subsidence on cratonic margins and in cratonic interiors via intraplate stress and reactivation of old structures (Quinlan, 1987; Ziegler, 1988; Marshak and Paulsen, 1996; Marshak et al., 2003). This would suggest a genetic link between tectonic activity at the cratonic margin and cratonic sequence development, as suggested by various authors (e.g., Sloss and Speed, 1974; Burke et al., 2003). One example of this process is the formation of the Michigan and Williston basins as discrete entities in Late Silurian time, during closure of the Iapetus Ocean (Leighton and Kolata, 1990). Another example is the cratonic tectonism associated with the sub-Absaroka unconformity (see “Absaroka Sequence (Late Mississippian to Early Jurassic)” section), generated by collisional plate convergence in the Marathon and Ouachita orogenesis (Sloss, 1988b). Large-scale tectonic features of the craton such as the Transcontinental Arch and the Nemaha Uplift (Figs. 1, 4, and 5) are also evidence of intraplate stress, particularly since the arch had a complex history of uplift and subsidence throughout the Paleozoic, probably linked to inherited basement structure, but requiring reactivation by variations in intraplate stress (Marshak et al., 2003). Intraplate stress as a mechanism for unconformity generation has been criticized because synchronous activity across the entire craton may be difficult to explain, since it would require very large stress fields to develop simultaneously across the whole craton. However, synchronous in the context of cratonic sequences may only mean within a few million years, so this may be less of a weakness in the intraplate stress model than it first appears. Also, examples of this kind of synchronous effect have been documented. On the African plate, obduction of ophiolites onto the northeastern margin of Afro-Arabia appears to have generated an unconformity related to intraplate deformation, during a period of 2 MY in the Senonian, across >20 million km2, from Oman to the Atlas, and from Kenya to Equitorial Guinea (Guiraud and Bosworth, 1997; Burke et al., 2003). Given this, application of the intraplate stress model as a mechanism for unconformity generation in the Phanerozoic for North America does not seem unreasonable.

Dynamic Topography Related to Subducting Slabs Subducting slabs are cold and therefore dense relative to the surrounding mantle, forming a positive mass anomaly within the mantle. This positive mass hangs in the mantle and exerts a force on the overlying lithosphere. The force is transferred via viscous mantle flow and acts on the base of the lithosphere, dragging it down to cause a dynamic topographic low. Such topography is considered to be dynamic because it is generated by mantle buoyancy forces that are constantly changing, due to convection within the mantle, but slowly enough that the dynamic topography can be considered in equilibrium with the evolving forces at any instant in time. A subducting slab may therefore produce a region of depressed lithosphere, extending as much as 2000 km into the craton, from a maximum near the active continental margin (e.g., Mitrovica et al., 1989; Gurnis, 1993; Husson, 2006; Flament et al., 2012) (Fig. 7). Such a dynamic topographic low will develop by subsidence as the subducting slab penetrates the mantle. The amplitude of the topography will depend on the degree and angle of slab penetration into the mantle and on the temperature of the slab; older oceanic crust will be colder and denser than



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FIG. 7  Mechanisms generating dynamic topography driving vertical motions of the cratonic interior.

younger crust, representing a greater mass anomaly in the mantle and generating a higher amplitude of dynamic topography. Consequently, dynamic topography will peak at the point of maximum slab penetration and thermal age (Burgess and Moresi, 1999). The dynamic topography will then gradually be reduced, causing uplift, as the slab’s thermal age decreases during final stages of ocean subduction, and/or the slab is detached and descends into the mantle (Gurnis, 1993; Burgess and Moresi, 1999). Since North America had several episodes of subduction at its margins during Phanerozoic time, a contribution to the tectonic evolution of the craton from slab-related dynamic topography is to be expected. Burgess et al. (1997) studied this effect, with particular emphasis on two episodes: subduction of Iapetus oceanic lithosphere during early Paleozoic time, and subduction of Pacific oceanic lithosphere during the Mesozoic and Cenozoic Cordilleran orogeny. We found that the early Paleozoic subduction episode probably contributed to development of Lower Paleozoic strata in eastern North America, as also described by Coakley and Gurnis (1995), but a more detailed subduction history is necessary to elaborate further. In contrast, Mesozoic and Cenozoic slab evolution is better constrained (Heine et al., 2008). In this case subsidence analysis and combined mantle and stratigraphic modeling highlight the likely role of slab-related dynamic topography in the development of Cordilleran and cratonic stratal patterns in western North America (e.g., Cross and Pilger, 1978; Cross, 1986; Pang and Nummedal, 1995; Burgess et al., 1997; Burgess and Moresi, 1999; Liu and Nummedal, 2004; Liu et al., 2008, 2011; Liu and Gurnis, 2010; Heller and Liu, 2016). For example, subsidence and uplift that led to development of the Cretaceous Interior Seaway, with Cretaceous strata distributed across a large area of the craton, and various other regional geological features of the southwest United States, are increasingly explicable as the consequence of penetration of the flat Farallon slab beneath southwestern North America in Late Cretaceous and Paleogene time. A key observation is that thick Upper Cretaceous strata are deposited from the Sevier foreland basin region in Utah, Wyoming, and so on, all the way east to Kansas and beyond, which is not possible to explain with eustasy alone (Bond, 1976) but is also difficult to explain just as a foreland basin. Burgess et al. (1997), following Cross and Pilger (1978), Cross (1986), and Mitrovica et al. (1989) modeled these spatially extensive strata as the products of a dynamic topography low related to penetration of the Farallon slab beneath North America, with a phase of low-angle to flat penetration in the Late Cretaceous, and uplift as the slab descended into the mantle. More recently Liu et al. (2008) used flexural backstripping to demonstrate that dynamic topography most likely created 800–1800 m of anomalous subsidence from eastern to western Wyoming. Liu et al. (2011) and Heller and Liu (2016) followed this up with more detailed analysis showing that the loci of anomalous subsidence moved eastward from c. 98 to 74 Ma, most likely related to the west-to-east passage of the Farallon slab. They further suggest that regional variations in subsidence rates, and a wave of uplift from 95 to 60 Ma, support the hypothesis that the thickened Farallon slab here represents a subducted oceanic plateau. Other recent analyses of spatial

50  The Sedimentary Basins of the United States and Canada

d­ istribution of thickness of Cretaceous strata, for example in western Wyoming (Leary et al., 2015) and over a broader region of Wyoming, Utah, Colorado, and South Dakota (Heller et al., 2003), also support this hypothesis. Accumulation of Cretaceous and Paleogene strata in this dynamic topographic low was followed by surface uplift during reversal of the topography as the slab thermal age decreased (Burgess et al., 1997), the slab dip increased, and finally the slab detached. This is a possible explanation for why many of these near-horizontal strata are now elevated significantly above sea level and being eroded (Burgess et al., 1997). More detailed modeling (Liu and Gurnis, 2010) supports this hypothesis. They predict and document a west-to-east wave of dynamic subsidence across the Colorado Plateau, that peaked c. 86 Ma. This was then followed by ~1.2 km of uplift, a first phase in the latest Cretaceous, and a second in the Eocene with a cumulative uplift of ~1.2 km by the Oligocene. The final effect of Farallon slab subduction was a tilt downward to the northeast before the Oligocene, then reversed to the southwest during the Miocene in response to buoyant mantle upwellings, possibly recorded in Miocene Ogallala Formation strata on the western Great Plains (Heller et al., 2003). A weakness with the slab-related dynamic topography model in the past was the lack of obvious dynamic topography on modern convergent margins, such as in Southeast Asia. In this area various mantle flow models predict 1–2 km of dynamic topography yet the maximum amplitude of dynamic topography in Southeast Asia has been calculated as no >300 m (Wheeler and White, 2000). However, Husson (2006) demonstrates that a significant fraction of the topographic variations observed above the Scotia, Mariana, and Hellenic subduction systems is consistent with simple theoretical models of dynamic stresses induced by subduction. From this, Husson (2006) suggests that slab-related dynamic topography is difficult to detect in modern tectonic systems only because isostatic processes tend to mask it. So, of course, interpretations of slabrelated dynamic topography remain open to debate. However, the evidence is mounting that this is a key control on cratonic subsidence and uplift and hence sequence formation in the cratonic interior of North America. It is also most likely a key control globally across many other tectonic regions.

Dynamic Topography Related to Supercontinent Cycles, Mantle Convection Cells, and Plate Amalgamation and Dispersal Heat energy is mostly lost from the mantle via convection and consequent volcanism at oceanic spreading centers. Formation of supercontinents that persisted for tens of millions of years, covering a wide region of mantle, prevent this method of heat loss, trapping mantle heat produced by radioactive decay in the core. Such mantle insulation may produce a rise in temperature of ~20 K throughout the mantle (Gurnis and Torsvik, 1994). Thermal expansion due to this temperature increase produces stress on the base of the lithosphere, creating dynamic topography (Anderson, 1982) with an amplitude of ~150 m (Fig. 7). Similar effects can also be produced by large descending plumes interacting with internal mantle viscosity boundaries (Pysklywec and Mitrovica, 1998). Dynamic topography is also created by large-scale mantle up- and downdwellings, part of the convective circulation process, so that plate motion can also drive long-wavelength subsidence and uplift of cratonic interiors as they move over dynamic topography highs and lows (Gurnis et al., 1998; Gurnis, 2001; Heine et al., 2008; Moucha et al., 2008; Spasojevic et al., 2008, 2009). As a consequence of this dynamic topography mechanism, continents may well experience uplift during supercontinent formation and persistence, followed by subsidence as the supercontinent breaks up and the fragments drift off hot mantle onto adjacent, relatively cool mantle (Anderson, 1982; Gurnis, 1988). Following on from these considerations of large-scale plate motions and dynamic topography, there is also an unresolved question whether mantle convection drives plate tectonics, or vice versa (Anderson, 2001), suggesting that our understanding of these processes remains incomplete. Despite the complexity of the processes involved, there is some evidence that this signature of large-scale convection and supercontinent-related dynamic topography is consistent with patterns of Phanerozoic North American cratonic sequence development. The two longest duration lacunae in North America are the base-Sauk and the base-Zuni unconformities (Fig. 3). Both formed during periods when North America was part of a supercontinent, initially the late Neoproterozoic, and then Pangea in the late Paleozoic and early Mesozoic. Cratonic erosion and nondeposition would have been accentuated by dynamic topographic highs creating an emergent craton (Sloss and Speed, 1974). Conversely, the three Paleozoic unconformities are of shorter duration (Fig. 3) and formed during a period when North America was one of several dispersed continents, overlying cool mantle, and thus having relatively low or submerged (Sloss and Speed, 1974) elevation. Subsidence analysis identifies an anomalously large subsidence event in Late Devonian to Mississippian time, explained by Kominz and Bond (1991) as due to the final stages of assembly of Pangea over a dynamic topographic low (Figs. 8 and 9). A further striking feature of the base-Zuni unconformity is its east-west asymmetry, reflected by the general absence of Mesozoic cratonic strata in eastern North America compared with extensive Jurassic and Cretaceous deposition in the west. This may be in part due to slab-related dynamic topography (see “Dynamic Topography Related to Subducting Slabs” section), but numerical modeling (Burgess et al., 1997) suggests that it may also be in part due to continent-scale tilting up-to-the-east.



Phanerozoic Evolution of the Cratonic Cover Chapter | 2  51

FIG. 8  Reduced subsidence curve (Kominz and Bond, 1991) from a stratigraphic section in north-central Iowa. Minimum and maximum curves are derived from appropriate compaction corrections. The curves represent components of eustasy, changing continental freeboard, and tectonic subsidence. Note the large subsidence event during the Devonian. (From Kominz and Bond (1991).)

Tilting would have resulted from increasing amounts of uplift eastwards across North America toward the hottest mantle situated beneath the center of Pangea (Fig. 7). The extended duration of the sub-Tejas unconformity in eastern North America, and the current elevated topography of North America and the resultant predominantly erosional regime, suggest that the dynamic topographic high is somehow persistent, or that elevation is being maintained by some other mechanism, at least locally (e.g., erosional processes on the Blue Ridge escarpment, Spotila et al., 2004). Note that Burke et al. (2003) also alludes to this general mechanism of thermal insulation beneath Pangea (see their Fig. 2) to explain the elevation of Africa and the generation of an apparently atypical number of deep-seated plume events (but see also Coltice et al. (2007) who suggest that thermal insulation alone can generate melting, without any requirement for plumes). Hartley and Allen (1994) observed that African intracratonic basins have a distinctive gravity signature, requiring either a cold, dense region in the underlying upper mantle, or a downward-acting dynamic force on the base of the lithosphere. From this they suggested that basins such as the Congo Basin may be situated above convective mantle downwellings, which would generate both the downward force and the elevated density. Such “cold-spot basins” would be expected to have sedimentary fills developed over a prolonged period of geological time, to lack well-developed rift precursors, to have low to moderate heat-flow histories, and to have an approximately circular platform. Based on this and descriptions

52  The Sedimentary Basins of the United States and Canada

Tippecanoe Ordovician Sil.

Sauk Cambrian

Peam. Ka-G2

Mosc Bash

Serp

Vis

Fr-Fa

Tour

350

Penn.

300

250 Ma

1. Williston Basin Formation top

Unconformity

1500 2000

EI-Gv

400

450

500

Gd-Em

Prid Lud Wen Lidy

Ashg

Card

Lld

Llv Areg

Trem

1000

LC

500

550

MC

EC 0

Abasaroka

Kaskaskia Devonian Miss.

2. Michigan Basin

Iowa baseline Probable minimum water depth correction

3.Lllinois Basin

7 8

4

5 6

lac

3

hia

2

pa

Iowa platform

ns

1 10 9

Ap

4. Southern Oklahoma Aulacogen

Cordillera

Depth in Metres

Location map

N 500 km

5. W. Virginia

6. Allegany CO., MD.

7. Hector Lake

8. Castle Mountain

9. Schell Creek

10. Bear River Range

FIG. 9  Reduced subsidence curves (Kominz and Bond, 1991) from locations in North America, numbered by location on the map. Effects of compaction and sediment loading have been removed. The curves represent components of eustasy, changing continental freeboard, and tectonic subsidence. The latter dominates the signal in terms of amplitude. Note the large subsidence event in the Devonian, developed to varying degrees in all 10 curves. (From Kominz and Bond (1991).)



Phanerozoic Evolution of the Cratonic Cover Chapter | 2  53

from Hartley and Allen (1994), a downwelling “cold spot” mechanism seems to fit several of the observed properties of the North American cratonic basins, perhaps making it a plausible contender to explain at least some aspects of the formation of the basins, as well as some of the uplift and subsidence history on the intervening platforms that presumably overlie relatively hot mantle. Downwelling of this type, persistent for tens of millions of years during the Paleozoic, could contribute to long-term subsidence history of cratonic basins. However, given the duration of cratonic basins, long-term downwelling subsidence would require a relatively stationary position for North America relative to underlying shallow mantle (Burke et al., 2003). Perhaps more realistic is a hypothesis of variations in vertical motion by this mechanism. For example, Petersen et al. (2010) used a combination of mantle convection and stratigraphic forward models to show how similar small-scale convection could generate stratigraphic sequences. More work is required to investigate these potentially important processes.

Magmatic Controls Magmatic underplating is the process of adding igneous melt material, less dense than the asthenosphere, to the base of the lithosphere, causing isostatic adjustment and uplift as a consequence (Brodie and White, 1995). For example, addition of a thickness of 5 km of basaltic rock would cause an initial uplift of 625 m, which combined with isostatic response to denudation would lead to a total exhumation of ~2.5 km, assuming typical densities for asthenosphere, crust, underplated material, and sedimentary rock (Brodie and White, 1995). Although it is a well-understood process, clearly it requires igneous activity to generate significant volumes of melt, some of which seems likely to be extruded at the surface, particularly if there are crustal-scale faults to serve as conduits for rising magma. Therefore, if there is no evidence of surface volcanics associated with uplift events, it is perhaps unlikely that underplating was the cause of the uplift. Another possible magmatic mechanism is intrusion of anorogenic granites that postdate orogenic intrusive events by several million years. Klein (1995) has postulated that intracratonic basin formation was controlled by intrusion of such granites. However, much of the proposed mechanism as described by Klein (1995) appears to have little to do with granite intrusion and more to do with extensional and compressional intraplate stress variations causing rift subsidence and block uplift, respectively. Klein (1995) proposes that the granites acted to focus intraplate stress, generating intracratonic basins, but it seems more likely that intraplate stress variations simply reactivated preexisting tectonic structures created during earlier phases of rifting, orogenesis, or general terrane accretion.

PHANEROZOIC EVOLUTION OF THE CRATONIC PLATFORM COVER According to Sloss (1963) the Phanerozoic strata covering the North American craton can be subdivided into a number of stratigraphic sequences. Sequences are defined as “rock-stratigraphic units of higher rank than Group or Supergroup, traceable over major areas of continent and bounded by unconformities of interregional scope.” Students and practitioners of sequence stratigraphy will recognize this definition, because it triggered development of many of the later sequence stratigraphic models and methods. Sloss synthesized a large amount of outcrop and subsurface data from the North American continental interior and, through this synthesis, identified a number of interregional unconformity surfaces that could be traced and correlated across the continent. Interestingly, the interregional unconformities are commonly not obviously distinguishable in their local development from other less areally extensive unconformity surfaces (Sloss, 1963). Based on identification of interregional unconformities, Sloss (1963) defined six craton-wide sequences in the Phanerozoic strata of North America, namely the Sauk, Tippecanoe, Kaskaskia, Absaroka, Zuni, and Tejas sequences (Fig. 3). In general terms, these sequences thin toward and onlap onto cratonic platform areas and thicken into intracratonic and marginal basins (Figs. 4 and 5). The interregional unconformities are of greatest duration in the cratonic interior of the continent and pass laterally into conformable successions in the marginal basins. Within this broad pattern there is considerable variability in unconformity development. For example, the sequence bounding unconformities may be less well developed in relatively rapidly subsiding intracratonic basin centers than they are on adjacent uplifted arches (Figs. 4 and 5). However, in all cases the strata beneath the unconformity are older than, and the strata overlying the unconformity younger than, the point of maximum offlap in adjacent marginal basins. Burgess et al. (1997) linked the large-scale pattern of cratonic unconformity development to the effects of dynamic topography related to supercontinent formation and breakup (see “Dynamic Topography Related to Supercontinent Cycles, Mantle Convection Cells, and Plate Amalgamation and Dispersal” section), an interpretation supported by subsequent work (e.g., Heine et al., 2008). Using an improving subsurface data set, the original six Sloss sequences have been further subdivided into subsequences (e.g., Sloss, 1988b). Each subsequence is also bounded by an interregional unconformity, but these may not be present across the entire craton. Subsequences are approximately equivalent to the megacycles of Haq et al. (1987).

54  The Sedimentary Basins of the United States and Canada

The Sauk Sequence (Late Precambrian to Early Ordovician) Sauk strata overlie an interregional, in many places angular, unconformity on late Precambrian sediments and older metamorphic rocks. Abundant stratigraphical evidence suggests that this basal-Sauk unconformity represents a buried land surface showing relief ranging from a few tens of meters that was fully transgressed and covered, to much larger features, such as the transcontinental arch that remained exposed throughout Cambrian time (Sloss, 1988b; Frazier and Schwimmer, 1987). Basal Sauk strata are remarkable for their uniformity, being composed of compositionally mature pure quartz sandstones, with only a few local exceptions. Flooding of the basal-Sauk land surface probably occurred due to cratonic subsidence, linked to rifting, continental drift, and development of dynamic topography during breakup of a late Neoproterozoic supercontinent (Sloss, 1988b; Bally, 1989; Hoffman, 1989; Burgess et al., 1997). Distribution of Sauk strata is shown in Fig. 10. Progressive cratonic transgression was interrupted by periods of minor regression, so that the Sauk sequence can be subdivided into three subsequences (Sauk I, II and III) that progressively overstep and onlap onto the cratonic platform. However, Runkel et al. (1998) point out that these sequence bounding unconformities can be hard to locate in outcrop or core, due to lack of mineralogical and textural changes, easily eroded substrate, and lack of vegetation and weathering to produce soil profiles. On the cratonic margins, Late Proterozoic and early Cambrian (Sauk I) strata are basal sandstones passing upwards to interbedded sandstones and shales, capped by thick carbonate strata. On the western and northwestern margin of the craton, Upper Proterozoic Sauk I strata of the Windermere Supergroup are relatively thick (Fig. 10) and represent mostly deep-water deposits, often felspathic and coarse-grained, formed in marginal basins, possibly with episodes of active rifting and likely some glacial influence (Hein et al., 1994, see also Chapter 5). A remarkable feature of the Sauk megasequence is the lateral extent of sandstone units that can be up to several hundred meters thick yet extend for hundreds of kilometers (Sloss, 1988b). Runkel et al. (1998, 2007) constructed a regional-scale depositional model for these sandstone units across six states in the

2

13

11

9 10

3 1 9 6 3 1 9 3 8 7

4

12

m/MY 5

5–10

40–50

10–20

50–60

20–30

60–80

30–40

80–100 >100

FIG. 10  Sauk sequence (latest Proterozoic to early Ordovician) net subsidence rates. Especially significant labeled features are (1) Cordilleran shelf basin; (3) Appalachian shelf; (4) Ouachita margin; (5) Marathon margin; (8) Mississippi River-Reelfoot rift system; (9) Transcontinental arch; (10) Michigan basin; (13) Williston basin. Note that areas of accumulation with rates <5 m/MY are not distinguished from locations with no preserved Sauk strata. (From Sloss (1988b).)

Phanerozoic Evolution of the Cratonic Cover Chapter | 2  55



Upper Mississippi Valley region and suggested that the units were deposited through four trilobite biozones during three distinct conditions of relative sea level, separated by subtle sequence-bounding unconformities. They suggest that the uniform sandstone lithology was due to abundant sand supply on a stable cratonic platform with relatively slowly, uniform subsidence, but significant relative sea-level changes, that led to low preservation of finer-grained fair-weather deposits. On a larger scale, these sand sheets were part of the “inner detrital belt,” surrounded outboard by carbonate platforms, at least later on in Sauk deposition, and beyond that by an “outer detrital belt” of more siliciclastic but generally finer-grained deeper-water strata forming the continental slope in margin basins (Frazier and Schwimmer, 1987; Myrow et al., 2012; their Fig. 1; see also Chapter 5). Younger Middle Cambrian (Sauk II) strata on the platform are predominantly dolomitized carbonates, with sandstone tongues common near the margin of the Canadian shield. Middle to Late Cambrian (Sauk III) strata, dominated by carbonates, overstep earlier Sauk deposits and onlap Proterozoic and Archaen basement on the Transcontinental Arch and the cratonic shield, representing ongoing Sauk transgression. Topographic features generated by these resistant Precambrian lithologies were emergent during much of Cambrian time but were subsiding and being gradually buried. Final burial of most of the Transcontinental Arch (Figs. 1 and 5) did not occur, however, until latest Mississippian time. The North American intracratonic basins first developed during deposition of the Sauk sequence, although in some cases Proterozoic precursor rift basins underlie them. Initiation of the Illinois and Michigan basins occurred during Middle Cambrian (Sauk II) time, and the Williston basin became established during Middle-Late Cambrian (Sauk III) deposition. Sauk sequence deposition ended because of relative sea-level fall toward the end of early Ordovician time. This relative sea-level fall was probably triggered by a change in stress fields due to plate convergence and creation of the Laurentian active margin during the Taconic orogeny, but possibly also had a component of glacio-eustastic fall related to continental ice sheet growth (Leighton and Kolata, 1990), though evidence for this glacioeustasy remains sparse.

The Tippecanoe Sequence (Middle Ordovician to Early Devonian) Distribution of Tippecanoe strata is shown in Fig. 11. Erosion of Sauk sequence strata beneath the sub-Tippecanoe unconformity is variable across the craton. In some places, commonly on arches, Tippecanoe strata overstep Sauk strata and lie

13

15

6

10 11

5

9 12 15

7

5

6 4 14

15

8 1

3

2

m/MY 5–10

30–40

10–20

40–50

20–30

>50

FIG. 11  Tippecanoe sequence (mid-Ordovician to early Devonian) net subsidence rates. Significant labeled features are (3) Mississippi Valley graben; (4) Illinois Basin; (5) Cordilleran shelf basin; (6) Appalachian shelf; (7) Cincinnati arch; (10) Michigan basin; (11) Wisconsin arch; (12) Kankakee arch; (13) Williston basin; (15) Transcontinental arch.

56  The Sedimentary Basins of the United States and Canada

directly on Precambrian basement (Figs. 4 and 5). Locally, dissolution weathering of Sauk carbonates created karst topography, later buried by Tippecanoe strata. The Tippecanoe sequence is divided into two subsequences by a disconformity, most likely related to the long-term glacio-eustatic fall caused by the latest-Ordovician earliest-Silurian Gondwanan glaciation (Sloss, 1988b; Delabroye and Vecoli, 2010). There is no evidence of angular discordance on the disconformity, apart from in the Hudson Bay basin, and the tectonic setting appears similar in both subsequences. Basal Tippecanoe strata exhibit onlap onto the craton, ranging in age from Middle Ordovician (Tippecanoe I) strata in marginal basins, to Silurian (Tippecanoe II) strata in places on the margin of the Canadian Shield and the Transcontinental Arch. This relationship indicates that the Tippecanoe transgression took tens of millions of years to complete and that the shield and cratonic arches were areas of positive relief relative to much of the rest of the cratonic platform, as they had been during Sauk deposition. Even in the marginal basins, the sub-Tippecanoe unconformity surface has a relief of up to 50 m. Details of the sub-Tippecanoe unconformity in eastern Ontario suggest possible subtle tectonic activity (Dix and Molgat, 1998) and the sequence developed under the influence of compressional collisional tectonics on the eastern cratonic margin, probably due to subduction and collision during the Taconic orogeny. The eastern passive margin became a foreland basin at this time, formed by flexural loading (see Chapters 3 and 4). Tectonic differentiation of the craton was initially low, similar to conditions prevailing during Sauk deposition, but increased during Tippecanoe deposition, so that various arches and basins became increasingly well defined, evolving toward a pattern that was then typical of mid-Paleozoic time (Sloss, 1988b). Siliciclastic rocks are present throughout the US cratonic area at the base of the Tippecanoe sequence, forming a sheet of pure compositionally mature quartz arenite, probably sourced from weathering of crystalline shield areas. Volumetrically, however, the sequence is dominated on the craton by younger carbonate strata that include numerous reef systems. Thick Silurian evaporite strata are also common, particularly in the intracratonic basins (see “The Michigan Basin,” “The Illinois Basin,” “The Williston Basin,” and “The Hudson Bay Basin” sections). Tippecanoe strata covered many arches, much of the western craton, and much of the Canadian shield. Evidence for this includes patterns of onlap and overstep, lack of marginal facies at the limit of present Tippecanoe outcrop, various outliers, and presence of Ordovician and Silurian carbonates preserved as diatreme xenoliths within igneous intrusions in shield areas (Sloss, 1988b; Cecile and Norford, 1993). The extent of Tippecanoe strata is significant because it suggests that the Tippecanoe marine transgression culminated in marine flooding of much of the craton, both platform and shield areas, indicating a period of exceptionally elevated relative sea level. This may have been due, in part at least, to dynamic topography, because at this time the continents were dispersing, moving from dynamic topographic highs to dynamic topographic lows, and therefore undergoing relative sea-level rise (Burgess et al., 1997) (see “Dynamic Topography Related to Supercontinent Cycles, Mantle Convection Cells, and Plate Amalgamation and Dispersal” section). This is also consistent with the relatively short duration unconformities separating the Sauk, Tippecanoe, Kaskaskia, and Absaroka sequences (Fig. 3).

Kaskaskia Sequence (Mid-Early Devonian to Late Mississippian) The basal Kaskaskia unconformity truncates strata of all older ages, from Sauk through Tippecanoe, demonstrating that there had been significant amounts of differential uplift and erosion on the unconformity surface (e.g., Bunker et al., 1988) (Figs. 4 and 5). Various arches were reactivated, presumably under a compressional Acadian orogeny stress regime (Fig. 3), related to final closure of Iapetus, and the landscape was dissected by fluvial channels (Leighton and Kolata, 1990). As with older sequences, the basal Kaskaskia transgression took several million years to reflood the entire craton, although initial subsidence, flooding, and relative sea-level rise were rapid (Figs. 8 and 9) (Kominz and Bond, 1991). The latter authors interpreted the rapid subsidence event, and the overall dominance of compressive stress, to indicate the initial assembly of continents over a dynamic topographic low produced by cold mantle (see “Dynamic Topography Related to Supercontinent Cycles, Mantle Convection Cells, and Plate Amalgamation and Dispersal” section), an interpretation supported by Burgess et al. (1997). Distribution of Kaskaskia strata is shown in Fig. 12. Kaskaskia sequence strata are subdivided into two subsequences, Kaskaskia I ranging from mid-Early Devonian to latest Devonian, and Kaskaskia II, from latest Devonian to Late Mississippian. Kaskaskia I strata progressively onlap basement structures, but in some cases are absent from basement highs (e.g., the Transcontinental Arch, Figs. 4 and 5) that were subaerially exposed throughout the Devonian period and only reflooded in Mississippian time when they were draped with Kaskaskia II strata. Basal Kaskaskia I sandstones occur locally and are either reworked Sauk and Tippecanoe material or derived from the exposed Canadian Shield. Much of the rest of the subsequence on the craton is dominated by thick bioclastic and biohermal carbonates that pass laterally into black shales and cherts in the marginal basins. Evaporite units within the carbonates represent periods of sabkha formation during periodic regressions, and carbonate deposition was also interrupted occasionally by periods of anoxia, generating black shales.

Phanerozoic Evolution of the Cratonic Cover Chapter | 2  57



Kaskaskia I 1 9

10 5 12

11

2

6

4

3

7

8

m/MY

(A)

5–10

40–50

10–20

50–60

20–30

60–70

30–40

>70

7

1

6 4

5 9 4 3

2

8

m/MY

(B)

5–10

30–40

10–20

40–60

20–30

>60

Kaskaskia II

FIG. 12  (A) Kaskaskia I subsequence (mid early Devonian to late Devonian) net subsidence rates. Significant labeled features are (2) Cincinnati arch; (6) Illinois basin; (10) Michigan basin; (12) Catskill siliciclastic wedge. (B) Kaskaskia II subsequence (latest Devonian to Late Mississippian) net subsidence rate. Significant labeled features are (1) Williston basin; (3) Reelfoot-Illinois basin; (4) Cordilleran foreland basin; (6) Michigan basin; (8) Black Warrior basin; (9) Appalachian foreland basin.

58  The Sedimentary Basins of the United States and Canada

The unconformity separating Kaskaskia I from Kaskaskia II strata was marked by uplift and erosion on arches and domes, but basinal areas suffered only a minor hiatus. However, Kaskaskia II strata record onset of significant siliciclastic influx from the Appalachian, Ouachita, Marathon, and Antler orogenic belts forming on the continental margins in the east, south, and west (Fig. 3), and from the partly reemergent Canadian Shield. Initial Kaskaskia II deposition shows evidence of restricted anoxic conditions, but this changed as transgression progressed, circulation improved, and carbonate deposition was established (Sloss, 1988b). Previously emergent basement features were transgressed and buried. Orogenic activity increased during the Late Mississippian, previously dominant carbonate production across the craton was finally smothered, and subsequent deposition was dominated by siliciclastic material derived from orogenic margins. Consequently, in many areas the youngest Kaskaskia strata preserved are sandstones and shales, in places forming cyclothems capped by coal.

Absaroka Sequence (Late Mississippian to Early Jurassic) The sub-Absaroka unconformity records extensive denudation that removed up to thousands of meters of older strata, particularly from old and new areas of positive relief (e.g., the Transcontinental Arch). In places, denudation removed all older Phanerozoic strata, so that basal Absaroka units rest directly on Precambrian basement. Much of this cratonic tectonic activity can be related to orogenic events on the western, southern, and eastern cratonic margins creating the Antler, OuachitaMarathon, and Alleghenian orogens, respectively (Sloss, 1988b; Leighton and Kolata, 1990; see also Chapters  7 and 8). Compressive stress from orogenic events transmitted into the cratonic interior reactivated numerous basement structures, causing uplift and erosion. Local uplifts on the margins of the intracratonic basins left them isolated from the rest of the platform. Leighton and Kolata (1990) postulated that the widespread uplift due to compressive stress fields may have been related to collision of Gondwana and Laurussia along the Eurasian-African margin. Given the icehouse setting, the degree to which a glacio-eustatic component also influenced development of the unconformity is debatable (Beuthin, 1994; Ettensohn, 1994). Distribution of Absaroka strata is shown in Fig. 13. As in older sequences, the basal Absaroka transgression was prolonged, generating upper Mississippian to Triassic transgressive units (Sloss, 1963). Once again, the oldest Absaroka strata are found in the marginal basins, and the youngest strata drape cratonic arches. In contrast to the older sequences, Absaroka cratonic strata are dominantly siliciclastic. Much of this material was derived from within the craton, either from the shield or from emergent arch areas, but marginal orogenic topography contributed sediment to the edges of the craton. The Absaroka sequence is split into three subsequences, spanning latest Mississippian to Early Permian (Absaroka I), mid to Late Permian (Absaroka II), and Triassic to Early Jurassic (Absaroka III) (Sloss, 1988b). The subsequences are generally defined on the basis of relatively minor disconformities; generally, the Absaroka sequence records a steady shallowing-upward trend. However, Bally (1989) noted that Absaroka I strata were influenced by a stress regime related to the Ouachita orogeny (Fig. 3), while Absaroka II deposition was dominated by the Alleghenian orogeny (Fig. 3). Both these orogenies occurred during assembly of the Pangea supercontinent. In contrast, Absaroka III deposition marked the onset of the subsequent breakup of Pangea. Absaroka I strata in southern, western, and midcontinent areas drape a complex mosaic of uplifted and eroded fault blocks and arches (Fig. 13), with proximal coarse siliciclastic material and evaporitic strata generally forming more distally (see also Chapter  7). Delta systems prograded southwestward on the eastern part of the craton, sourced from the Appalachian and eastern Canadian Shield areas. Maximum cratonic transgression occurred in Pennsylvanian time, establishing predominantly marine conditions in the southwest, and passing into more terrestrially dominated deltaic strata in the northeast, perhaps influenced by glacio-eustatic and local tectonic factors to create cyclothems. Absaroka II strata indicate considerably less tectonic differentiation (Fig. 13), with deposition of alternating marine and continental strata across a broad, apparently stable, cratonic area (Sloss, 1988b). Siliciclastic sediment influx from the northeast continued, but the southern and western cratonic areas were dominated by carbonate deposition. The trend of southwestward retreat of the sea was continued during deposition of Absaroka III strata. This subsequence records a final episode of cratonic deposition in mostly terrestrial conditions and limited to the western margin of the craton (Fig. 13). Absaroka deposition was finally terminated by relative sea-level fall and development of a craton-wide erosional unconformity surface, cutting into Pennsylvanian strata in the east, and Permian to Lower Jurassic strata in the west. The prolonged regression during later stages of Absaroka deposition and the consequent long midcraton hiatus before onset of Zuni deposition are consistent with gradual elevation and up-to-the-east tilting of North America over a dynamic topographic high generated by supercontinent insulation (Burgess et al., 1997).

Zuni Sequence (Middle Jurassic to Early Paleocene) Deposition of Zuni strata, ranging in age from mid-Jurassic to early Paleocene, was initially limited to the Cordilleran margin in western North America (Fig. 14). The sequence then records progressive overstep onto the craton into Late Jurassic and

FIG. 13  (A) Absaroka I subsequence (latest Mississippian to Early Permian) net subsidence rates. Significant labeled features are (1) Ouachita margin; (2) Marathon margin; (3) Uncompahgre uplift; (10) Paradox basin; (11) Eagle basin; (12) Denver basin; (13) Anadarko basin; (14) Fort Worth basin; (16) Delaware basin; (18) Nemaha uplift; (21) Michigan basin; (22) La Salle anticline; (23) Cincinnati arch; (26) Central Kansas uplift; (22) Illinois basin; (30) Appalachian basin. (B) Absaroka II subsequence (mid and Late Permian) net subsidence rates. Significant labeled features are (1) Delaware basin; (2) Midland basin; (3) Central Platform basin; (4) Anadarko basin; (5) Alliance basin; (6) Williston basin; (7) Colorado Plateau area. (C) Absaroka III subsequence (Triassic to Early Jurassic) net subsidence rates. Significant labeled features are (1) Extensional rift basins; (2) Permian basin; (3) Williston basin; (4) Western limit aeolian sandstones.

60  The Sedimentary Basins of the United States and Canada

FIG. 14  (A) Zuni I subsequence (mid-Jurassic to earliest Cretaceous) net subsidence rate. (B) Zuni II subsequence (mid-early Cretaceous to earliest late Cretaceous) net subsidence rates. (C) Zuni III subsequence (late Cretaceous to early Paleocene) net subsidence rates. Significant labeled features are (1) Wind River uplift; (2) Front Range uplift; (3) Kaibab uplift; (4) Red Desert-Hanna basin; (5) Piceance-Washakie basin; (6) Denver basin; (7) Pedregosa basin.



Phanerozoic Evolution of the Cratonic Cover Chapter | 2  61

Cretaceous time, and onset of deposition in the Gulf and Atlantic rift-margin basins. Overstep peaked in the Late Cretaceous, with deposition of marine and terrestrial strata over much of the cratonic platform and the Canadian Shield, indicated by significant preserved thickness, or by remnant outliers (Bunker et al., 1988; Sloss, 1988b). This pattern is similar to that developed during Sauk deposition in Cambro-Ordovician time (Sloss, 1988b) and probably represents somewhat similar tectonic conditions, with deposition occurring during cratonic subsidence driven by evolving dynamic topography related to ongoing supercontinent breakup (Burgess et al., 1997). The vast extent of Cretaceous deposition is an indication of the extent of the Cretaceous Interior Seaway that flooded much of North America (see Chapter 9). This is, in part at least, due to the low-angle penetration of the Farallon slab beneath western North America and the resulting dynamic topographic low, best developed on the Cordilleran margin, but extending across much of the craton (Heller and Liu, 2016; and see “Dynamic Topography Related to Subducting Slabs” section). High global sea level no doubt also played a part in generation of the Cretaceous seaway, but thickness trends and long-wavelength tilting of the strata indicate the influence of slab-related dynamic topography (Burgess et al., 1997, 1997; Burgess and Moresi, 1999, 1999; Cross, 1986; Cross and Pilger, 1978; Pang and Nummedal, 1995; Liu and Nummedal, 2004; Liu et al., 2008, 2011; Liu and Gurnis, 2010; Heller and Liu, 2016). Division of the Zuni into three subsequences (Sloss, 1988b) indicates that the pattern of gradual overlap during transgression of the Interior Seaway was interrupted twice, once in the early Cretaceous epoch, separating Zuni I and Zuni II subsequences, and again at the beginning of the Late Cretaceous, separating Zuni II and Zuni III subsequences. In both cases, relative sea-level falls caused terrestrial deposition, subaerial erosion, and hiatus. The Zuni sequence is thus subdivided into three subsequences. The hiatus at the base of the Zuni II subsequence was particularly significant, representing around 10 MY of nondeposition, and having an equivalent unconformity developed throughout Europe (Leighton and Kolata, 1990). Consequently this subsequence boundary is interpreted as formed by a prolonged eustatic low. During early Zuni I deposition the cratonic interior was the main source of siliciclastic detritus, but orogenic loads developed in the west became the dominant sediment source from Late Jurassic time onwards (Sloss, 1988b). Throughout Zuni deposition, large volumes of siliciclastic sediment were shed from the developing mountains, extending 300 km or more eastwards. Further east, thick marine shales and carbonates accumulated across much of the cratonic platform. Subsidence rates were eventually outpaced by sediment supply rates, and perhaps by eustatic fall, and finally terminated by a change to tectonic uplift (Leighton and Kolata, 1990). Consequently Paleocene platform deposition was mostly terrestrial, and Zuni deposition was terminated by development of a major unconformity, often showing angular discordance, indicative of significant tectonic tilting.

Tejas Sequence (Late Paleocene to Present) Development of the sub-Tejas unconformity in the west can largely be related to deformation of crustal blocks during the Laramide orogeny (Sloss, 1988b). Elsewhere on the craton a combination of low eustatic sea level and long-wavelength uplift can be invoked (Burgess et al., 1997; Burgess and Moresi, 1999). Deposition of Tejas strata was areally more limited than in previous sequences. Laramide crustal blocks and associated small basins filled with late Paleocene to middle Eocene Tejas I fluvial and lacustrine deposits. Tejas II and Tejas III strata consist of large volumes of fluvial strata deposited in the west, shed from and burying the Laramide Mountains. Although cratonic Tejas volumes are negligible compared to strata from previous sequences, noncratonic Tejas strata form thick successions in the Gulf Coast and Atlantic passive margin basins (see Chapters 13 and 14). The area of western North America previously occupied by Zuni and Tejas depocenters was uplifted during Neogene times to elevations of 1.5 km or more, and the Zuni and Tejas strata tilted, exposed, and was subjected to erosion. This process can be explained reasonably well by the decay of a dynamic topographic low associated with the Farallon slab. The dynamic topographic low was responsible in part for accommodation creation in the Zuni and Tejas depocenters. During the final stages of subduction and slab detachment, the dynamic topography was reversed and rock and surface uplift occurred (Mitrovica et al., 1989; Burgess et al., 1997). The current elevated hypsometry of the rest of the North American craton is more difficult to explain but, given similar hypsometries on other continents, may be caused by some long-wavelength mantle effect, perhaps elevated temperatures persisting from Pangean mantle insulation. Relatively low oceanic water volumes due to significant water storage in polar ice caps also contributes to the observed elevated hypsometry.

THE NORTH AMERICAN INTRACRATONIC BASINS The large-scale geometry and stratal architectures of the four intracratonic basins discussed here are summarized in Fig. 15. Their increased stratal thickness relative to surrounding platform areas and their spatial distinction from marginal basins are shown in Fig. 2.

62  The Sedimentary Basins of the United States and Canada

W Big Horn Mtn. 1

E Powder River Basin

Cedar Creek Anticline

Sea level

Sea level

Upper cretaceous Lower cret

1 (km)

Nesson Anticline Paleocene

Jurassic Cambrian

2

Permian

OR

D

Mississippian Devonian

3

Silurian

Williston Basin 4

Salt Zuni Absaroka Kaskaskia Tippecanoe Sauk Precambrian sediments Crystalline precambrian

Narwhal Sea level 1 (km)

2

Ord

ovic i

an

Sea level

Devo

nian

Siluria

n

1 (km) 2

Hudson Bay

Pennsylvanian Sea level

Sea level

Mississippian Devonian

1 2

2 Ord

3

4

5

(km)

5

Pennsylvanian Mississippina n n nia uria vo l i e S D

n

vicia Ordo

Sea level

2 Cambrian

Hudson Bay

Williston

1 (km)

2 3

(km)

3

Michigan Basin

1

200 km

Cambrian

4

Sea level

100

1

Sil

(km)

0

3

Michigan 1200 km

Illinois

Illinois Basin FIG. 15  Schematic cross-sections of the four intracratonic basins of North America. The sequences of Sloss (1963) are coded by different stipple patterns. (From Bally (1989) and Leighton and Kolata (1990).)



Phanerozoic Evolution of the Cratonic Cover Chapter | 2  63

The Michigan Basin The Michigan basin is an oval-shaped intracratonic depocenter covering most of the state of Michigan. It contains around 4500 m of strata composed mostly of carbonate and evaporite rocks with only subordinate siliciclastic sedimentary rocks (Fisher et al., 1988) (Fig. 15). The basin is situated atop a Precambrian failed rift system (Nunn et al., 1984). Phanerozoic strata within the basin range in age from Cambrian to Jurassic and represent the Sauk to Zuni sequences of Sloss (1963) (Fig. 16).

Phanerozoic History of the Basin Deposition in the basin began in Late Cambrian time, filling a Precambrian structural low with up to 2000 m of sandstone, shale, and sandy dolomites (Fisher et al., 1988) forming a series of deepening-upward cycles (Catacosinos et al., 1990) similar to other Sauk strata elsewhere (Myrow et al., 2012). These initial deposits form the Sauk II sequence, and unconformably overlie Precambrian rocks. Sauk II strata pass conformably up into overlying strata of the Sauk III sequence.

FIG. 16  Chronostratigraphic diagram of strata in the Michigan basin. (From Catacosinos et al. (1990).)

64  The Sedimentary Basins of the United States and Canada

In this case the Sauk II–III subdivision is based on biostratigraphic correlation of lithostratigraphic units with other areas of Cambrian strata in North America. Sauk III strata are dominated by shallow-water and peritidal dolomites (Catacosinos et al., 1990) with subordinate sandstones and lateral transitions into shale toward the basin center. Deposition of the Sauk III sequence was terminated by a developing unconformity ranging in age across the basin from Late Cambrian to Middle Ordovician. Erratically distributed basal units of the overlying Tippecanoe I sequence suggest deposition over a highly irregular surface, perhaps a karst terrain developed on underlying Sauk III carbonates. The sequence shows a general fining-upwards trend, from basal nearshore marine sandstone units, into overlying shale and shallow-marine platform carbonate strata. During deposition of this sequence, the basin departed from its previous bulls-eye pattern and tilted downwards to the east for a period of 10–15 MY. This tilting event has been ascribed to subsidence due to dynamic topography forming above a slab of oceanic lithosphere being subducted westwards beneath the craton from the Iapetus margin to the east (Coakley and Gurnis, 1995). Tippecanoe I deposition was terminated at the end of the Ordovician period by development of a basin-wide unconformity. Deposition recommenced in Silurian time, forming the Tippecanoe II sequence. Strata in this sequence are predominantly carbonates and evaporites, and include abundant patch reefs, as well as a barrier reef carbonate, up to 210 m thick, that encircles the basin (Fisher et al., 1988; Catacosinos et al., 1990). Strata younger than Tippecanoe II are absent on the northern edge of the basin, and in the rest of the basin Tippecanoe II strata are overlain unconformably by the Kaskaskia I sequence. Kaskaskia I strata are dominated by shelf carbonates, shale, and sabkha evaporites. The upper parts of the sequence include black shales interfingering with less organic-rich gray green shales. The top of the sequence is marked by presence of shale interpreted as prodelta deposits, formed by a river flowing south from Canada. Kaskaskia II strata overlie Kaskaskia I conformably, but occupy a significantly smaller area, due largely to a combination of nondeposition and post-Mississippian erosion over structural highs. The sequence is dominated by terrigenous input. Deltaic deposition continued in the east, with more proximal deposits forming on a delta top. In other parts of the basin marine shales and sandstones dominate, with only occasional thin limestones. In total, the Kaskaskia I and II subsequences reach a maximum 530 m thick. During latest Mississippian, the Michigan basin was an area of low positive relief, bypassed by large rivers transporting sediment into the Illinois basin to the south (see “The Illinois Basin” section). Pennsylvanian Absaroka I strata are limited to the basin center, lie disconformably atop the Kaskaskia II strata, and are no >230 m thick. Interbedded shales, coals, and limestones indicate alternating marine and nonmarine deposition typical of Pennsylvanian strata throughout North America and Europe. Strata forming the Jurassic Zuni I sequence are even more geographically restricted and <120 m thick. They have been interpreted to represent fluvial paleovalley fills, either of Kimmeridgian or Bajocian-Bathonian age. Decreasing depositional area in the basin, either due to postdepositional erosion (in the case of the Absaroka strata), or limited deposition (in the case of Zuni strata), can be interpreted in the wider context of supercontinent formation and uplift related to mantle insulation (see “Dynamic Topography Related to Supercontinent Cycles, Mantle Convection Cells, and Plate Amalgamation and Dispersal” section).

The Illinois Basin The Illinois basin is a saucer-shaped, oval depression filled with up to 7000 m of Paleozoic strata ranging in age from Upper Cambrian to Permian (Fig. 17), with only minor thicknesses of Mesozoic and Cenozoic strata. Approximately equal volumes of siliciclastic and carbonate lithologies occur. Lower Cambrian strata fill a failed rift beneath the younger basin fill (Bushbach and Kolata, 1990). Basement rocks are Precambrian granite and rhyolite dated as between 1420 and 1500 Ma. The Sauk, Tippecanoe, Kaskaskia, and Absaroka sequences are well developed in the basin, but Zuni and Tejas strata are poorly represented, apart from a Quaternary cover over much of the basin (Fig. 17).

Phanerozoic History of the Basin The oldest strata in the basin are Lower and Middle Cambrian sediments filling the New Madrid Rift complex (Braille et al., 1982; Collinson et al., 1988) in the southern part of the basin. The rift probably formed during the break-up of the Rodinia supercontinent (Braille et al., 1982) and rifting had probably ended by Late Cambrian time. Postrift subsidence was more widespread, so that the rest of the Sauk sequence was deposited in late Cambrian and early Ordovician time over an area extending north away from the rift complex (Fig. 11). Potsdam Supergroup strata are predominantly sandstones and form the lower part of the Sauk sequence, present over much of the basin and sourced predominantly from the shield areas to the north. Knox Supergroup carbonates overlie the sandstones in the south, and interfinger in the north and east. During Sauk



Phanerozoic Evolution of the Cratonic Cover Chapter | 2  65

FIG. 17  Chronostratigraphic diagram of strata in the Illinois basin. (From Bushbach and Kolata (1990).)

66  The Sedimentary Basins of the United States and Canada

deposition the basin was open to the south, forming a large embayment on the edge of the craton. Separation of the Illinois and Michigan basin areas by arch uplift began during Sauk deposition. The whole Sauk sequence is interpreted as products of a shallow, subtidal marine system recording a gradual transgression and deepening, terminated finally by development of the sub-Tippecanoe unconformity surface. The unconformity represents a 10–15 MY hiatus, and in places up to 31 m of relief is preserved on the unconformity surface. Deposition of the Tippecanoe sequence began in the south and spread northwards. The thickest part of the sequence is in southern Illinois and western Kentucky. Tippecanoe strata are dominated by carbonate rocks, with only limited volumes of siliciclastic deposition recorded. The Tippecanoe I and II subsequences are defined by an unconformity separating Ordovician and Silurian strata, marked by slight erosion (Collinson et al., 1988). Several other hiatuses and minor erosion surfaces indicate more complex relative sea-level history than during Sauk deposition, in part due to local and regional uplift, and in part due to glacioeustasy (Collinson et al., 1988). Tippecanoe I strata represent predominantly peritidal and shallow-subtidal deposition, with deeper subtidal deposition in the west at the end of the Ordovician. Similar conditions prevailed during Tippecanoe II deposition, but with local reef buildups, and more siliciclastic input. Most siliciclastic strata were derived from the Ozark uplift, the Transcontinental Arch, and from the Canadian Shield, but some of the Upper Ordovician fine-grained siliciclastic strata originated from an Appalachian source. Erosion on the sub-Kaskaskia unconformity mirrors the distribution of Tippecanoe thicknesses. Least erosion occurred in the deepest parts of the basin in southern Illinois and western Kentucky (Fig. 11), and in places deposition was uninterrupted. The greatest amounts of erosion, cutting down to Ordovician or pre-Tippecanoe, occurred over the Ozark Dome and the Northeast Missouri Arch (Kolata and Olive, 1990). Various tectonic elements within the basin were uplifted and eroded during unconformity formation (Collinson et al., 1988), and in places solution fissures up to 22 m deep, and channels up to 16 m deep, were formed (Devera and Hasenmueller, 1990). Kaskaskia strata are again carbonate dominated, but with an increased clastic input relative to earlier sequences. The sequence is divided into subsequences at the Devonian-Mississippian boundary, but the boundary is largely conformable, identified only from biostratigraphic data (Collinson et al., 1988). Further transgression and deposition of carbonate strata across the basin followed initial deposition of sandstone above the basal unconformity in the basin center. These strata were in turn overlain by laminated black shales (New Albany Group), deposited across large portions of the craton (see “Phanerozoic Evolution of the Cratonic Platform Cover” section for an overview of Sloss sequences), and indicating extensive anoxic marine conditions. Carbonate deposition in the basin resumed, forming the Mammoth Cave Supergroup (Fig.  17). Mixed siliciclastic and carbonate deposition occurred in the latest Mississippian when delta systems associated with a large river system prograded into the basin, sourced either from the Canadian Shield or from the adjacent Appalachian mountains (Collinson et al., 1988; see also Chapter 4). Onset of clastic deposition is probably linked to development of the Alleghenian and Ouachita orogenies. Kaskaskia deposition terminated as the basin became subaerially exposed, karst topographies formed, and fluvial systems became deeply incised. Maintaining the pattern of earlier unconformities, the sub-Absaroka unconformity decreases in duration southwards to the deepest parts of the basin. The entire basin area was subaerially exposed, with contemporaneous tectonic activity causing surface uplift and accentuating erosion so that fluvial valleys up to 130 m deep and 32 km wide were formed (Collinson et al., 1988). Early Absaroka deposition was marked by continued tectonic activity and intrabasinal highs either being eroded or accumulating relatively shallow-water deposits. Deposition commenced in the south where over 700 m of sediment accumulated, and transgression deposited strata progressively northwards. Absaroka strata are dominantly siliciclastic, formed in fluvial, deltaic, and shallow-marine environments. The basin opened to the south and had a low gradient, so that small fluctuations in relative sea level caused large lateral shifts in depositional environment, reflected in numerous cycles of delta progradation and subsequent marine flooding. Deposition was continuous through into Permian times in parts of the basin, and coal maturity and shale compaction indicate that certain areas of Pennsylvanian strata were covered by significant thicknesses of Permian strata, since eroded. There is no evidence of deposition younger than Early Permian until Cretaceous strata of the Zuni sequence were deposited. Thus the base-Zuni unconformity represents a hiatus of approximately 200 MY, probably associated with whole-continent tilting up-to-the-east by dynamic topography related to supercontinent insulation (Burgess et al., 1997; see “Dynamic Topography Related to Supercontinent Cycles, Mantle Convection Cells, and Plate Amalgamation and Dispersal” section). Various faults and arches were active during the sub-Zuni hiatus, and igneous rocks of Permian age occur in southern Illinois (Collinson et al., 1988). Active tectonic features were eroded and reduced to base level during the hiatus, so that the first Zuni strata in the Late Cretaceous epoch are nonmarine, overlap eroded Ordovician to Mississippian strata, and in some areas overlie a basal soil (Kolata and Olive, 1990). Zuni strata reach a maximum thickness of around 600 m in the southern part of the basin, and are poorly preserved elsewhere, with only scattered outliers in western Illinois preserving a maximum thickness of 30 m. These outliers are important, however, because they represent some of the easternmost e­ xamples of Upper Cretaceous strata and were deposited in a nearshore marine environment. They were therefore probably deposited on the



Phanerozoic Evolution of the Cratonic Cover Chapter | 2  67

eastern margin of the Cretaceous Interior Seaway (Kolata and Olive, 1990), in accommodation generated by the easternmost limits of slab-related dynamic topography (see “Dynamic Topography Related to Subducting Slabs” section). Petrological data suggest a source in the metamorphic terranes of the Appalachian area to the east. Unconformities within the Zuni strata suggest numerous oscillations in relative sea level; depositional environments range from terrestrial to shallow marine. Marine Paleocene strata overlie Cretaceous strata across an unconformity and mark the final deposition of the Zuni sequence. Tejas strata in the Illinois basin are Pliocene in age and consist of terrestrial gravel. The basin area thus remained emergent throughout Cenozoic time after an initial period of Paleocene marine deposition. Latest Tejas deposition consists of Quaternary glacial deposits, with notably low preservation potential.

The Williston Basin The Williston basin is another elliptical intracratonic depocenter, filled with up to 4900 m of strata spanning the whole of Phanerozoic time (Fig. 18). Upper Cambrian strata initiated the basin fill and it was initially a cratonic-margin basin, only becoming intracratonic during the Cordilleran orogeny as material was accreted to the western continental margin. Data on basement rocks are rather sparse, but Archean Canadian shield rocks extend beneath the basin, and there is also evidence for island arc and oceanic crust basement to the west of the shield rocks (Green et al., 1985; Gerhard et al., 1990; Kent and Christopher, 1994). There is abundant evidence for influence of faults and arches active throughout Phanerozoic time and influencing stratal architectures.

Phanerozoic History of the Basin Deposition of Upper Cambrian Sauk strata began when the basin area was still only an embayment on the western cratonic margin. Thinning of Sauk strata across structural features suggests that there was significant relief on the base-Sauk unconformity surface. Sauk strata are dominantly quartzites and siliciclastic conglomerates, passing upwards into conglomerates with carbonate clasts. Sauk deposition terminated as the Taconic orogeny began on the eastern cratonic margin, perhaps suggesting some genetic link caused by intraplate stress variations. Although the first ~120 m of Tippecanoe strata are marine siliciclastic rocks, formed during initial transgression of the sub-Tippecanoe unconformity surface, the rest of the sequence (up to ~650 m thick) is dominated by carbonate and evaporite strata. The carbonates are mostly shallow-marine limestones and dolomites, interfingered with evaporite strata, also formed subaqueously during intervals of reduced basin circulation (El Taki and Pratt, 2012). Various structural elements within the basin became established during or at the end of Tippecanoe deposition, some with significant relief influencing depositional patterns. The Tippecanoe sequence is terminated by a basin-wide unconformity surface with some evidence for partial karst development on underlying carbonate strata (Gerhard et al., 1990). As in other North American cratonic basins, Kaskaskia strata can be subdivided into Devonian (Kaskaskia I), and Upper Devonian to Mississippian (Kaskaskia II) successions, based on a regional, though not basin-wide, unconformity (Fig. 18). The unconformity is particularly well developed atop structural features, indicating differential uplift, and is absent in the basin center (Gerhard et al., 1990). Both subsequences are carbonate and evaporite dominated, with only minor, finegrained siliciclastic strata present, in contrast to the pattern seen in other intracratonic basins in the east. Initial transgressive strata in the Kaskaskia II sequence are the Bakken Shale, which is an important petroleum source rock in the basin (Gerhard et al., 1990). The Williston basin underwent renewed subsidence and the marine connection to the Alberta basin in the west was cut. Both events were probably related to development of subduction and a foreland basin system, the Elk Point basin, to the west during the Ellesmerian orogeny (Sloss, 1988b; Leighton and Kolata, 1990). Kaskaskia II strata show abundant evidence of fluctuating relative sea level, with rapid transgression followed by gradual progradation (Gerhard et al., 1990). Basin margin strata were deposited in shallow-marine and peritidal settings and subjected to extensive subaerial exposure prior to final burial. Basin center strata comprise a series of shallowing-upward carbonate cycles, capped by crinoidal bioclastic banks and sabkha facies. Following the pattern of increasing structural influence, the post-Kaskaskia unconformity is present across the basin but best developed on various uplifted arches. Tectonic influence on the unconformity development is indicated by contemporaneous uplift in nearby areas such as the Canadian Shield (Aitken, 1993) and the Sioux Arch in South Dakota. This also manifests itself in a change of depositional style in the Absaroka sequence when siliciclastic strata derived from these uplifted extrabasinal sources became dominant, with the main sediment supply from the south. Depositional environments consisted of terrestrial siliciclastic systems interfingering with basinal marginal-marine siliciclastic and evaporite environments. These marginal marine systems prograded during deposition of the Absaroka strata, and the basin became increasingly restricted and saline, with more basin-center evaporites being deposited. The final phase of Absaroka deposition was marked by a decrease in salinity and deposition of carbonates followed by a final basin fill of fine- to medium-grained siliciclastic strata. Deposition was finally terminated by erosion, presumably triggered by a new phase of relative sea-level fall. By this time, accretion on the active margin to the west had left the basin in a truly intracratonic setting.

68  The Sedimentary Basins of the United States and Canada

FIG. 18  Chronostratigraphic diagram of strata in the Williston basin. (From Gerhard et al. (1990).)



Phanerozoic Evolution of the Cratonic Cover Chapter | 2  69

Zuni strata are Jurassic to Paleocene in age, and consist of mixed carbonate and siliciclastic rocks. Deposition was not continuous; Jurassic and Cretaceous strata are separated by an unconformity representing a 26 MY hiatus (Gerhard and Anderson, 1988). Jurassic deposition was more carbonate dominated, but siliciclastic strata became more prevalent into the Cretaceous when the area was part of the Cretaceous Interior Seaway, accumulating a series of transgressive and regressive cycles characteristic of the succession at this time (see Chapter 8). The rest of the Zuni sequence, and the subsequent Tejas sequence follow the patterns observed in the Interior Seaway, summarized in Chapter 8.

The Hudson Bay Basin The Hudson Platform is comprised of a series of basins with intervening arches (Fig. 19). The largest of these basins is the Hudson Bay basin. The basin was initially a relatively simple saucer-shaped depression, but evolved a more complex geometry later in the Paleozoic when preexisting fault systems and arches were reactivated by development of regional stress fields. Lavoie et al. (2015) argue that seismic images provide evidence for early Paleozoic synrift faulting in the basin. Seismic images in Lavoie et al. (2015) show what look like high-angle faults with vertical displacement in the order of 200–300 ms, and in some cases a branching-upward geometry. The structural style depicted in these seismic images is more consistent with a transtensional origin for the faults. For example, there are no half-graben or tilted fault block structures visible, no associated stratigraphic thickness changes and the faults seem to be forming negative flower structures in at least one location. This is therefore only very weak evidence for synrift extension, and therefore previous interpretations of later Paleozoic faulting remain most consistent with available evidence. Basement rocks consist of ensialic Archean protocontinent in the west and Lower Proterozoic fold-belts, including ophiolites, in the east (Roksandic, 1987). The preserved basin fill is mostly limited to Ordovician to Devonian strata, though there are minor areas of Cambrian and Cretaceous strata present (Fig. 19). Total composite thickness does not exceed 2500 m (Figs. 15 and 19).

Phanerozoic History of the Basin During the Cambrian period the Canadian Shield area was still topographically high and therefore remained emergent. The transgressive episode recorded in Sauk strata across much of the US portion of the North American craton had little recorded effect in the Hudson Bay area, with Sauk strata present only in small areas in the south-east. These Sauk strata consist of c. 60 m of unfossiliferous orthoquartzite sandstones and conglomerates, overlain by sandy and stromatolitic dolostones (Sanford, 1987). From Middle Ordovician time, marine transgression was more pronounced, flooding the basin and onlapping strata onto the Transcontinental Arch (Fig. 19). Approximately 180 m of strata accumulated, directly overlying Precambrian basement. A middle Ordovician regression ended with renewed transgression, and deposition of between 50 and 200 m of mixed carbonate and siliciclastic strata, once again onlapping and thinning onto the Transcontinental Arch to the west. Limited evidence suggests that the transgression may have peaked with a marine connection through to the St. Lawrence Platform (Fig. 1). The trend of increasing transgression and encroachment of the Hudson Platform, Canadian Shield and the Transcontinental Arch continued during deposition of Ordovician to Silurian Tippecanoe strata, representing the continued influence of a developing dynamic topographic low formed as the continents dispersed (see “Dynamic Topography Related to Supercontinent Cycles, Mantle Convection Cells, and Plate Amalgamation and Dispersal” section). Marine connections were established across the whole platform and as far as the Arctic and St. Lawrence platforms. During this interval the Hudson Bay basin first became an identifiable depocenter, distinct from the surrounding platform area, suggesting operation of a relatively focused subsidence mechanism in the area of the basin. A hiatus separates Ordovician from Silurian strata, defining the Tippecanoe I and Tippecanoe II subsequences. Ordovician and Silurian lithologies include black bituminous oil shales, and carbonates that formed thick basin-margin successions, including extensive barrier reef units that ringed the basin and fringed adjacent tectonic arches (Sanford, 1987). Conditions for development of these Middle Silurian reef systems may have been aided in part by reactivation of the various structural arches underlying the central portion of the basin. Reactivation caused uplift, initially producing shoaling in the central basin area, and finally generating an emergent arch (the Hudson Bay Central High, Pinet (2015)). This arch facilitated reef growth and divided the basin into eastern and western subbasins (Roksandic, 1987; Pinet, 2015). Uplift of surrounding cratonic areas left the Hudson Platform as an inland sea with only narrow connections to the Arctic Platform and to the Hudson Straits in the northeast. This episode of fault reactivation and uplift may have been a consequence of far-field stress transmitted from the continental margin during the Silurian-earliest Devonian Salinian orogenic events (Pinet, 2015). Kaskaskia strata are separated from underlying Tippecanoe strata by an early Devonian unconformity surface (Roksandic, 1987). However, Kaskaskia strata continue many of the trends and characteristics of Tippecanoe strata, although Kaskaskia strata are not present in the northeast of the basin. The Hudson Bay Central High was progressively buried during Middle and Late Devonian time (Fig. 19). Some initial Kaskaskia deposits developed locally are terrigenous redbed units that interfinger with marine carbonates. Most of the

70  The Sedimentary Basins of the United States and Canada

FIG. 19  East-west restored sections: Canadian craton. CO1, early Cambrian to Canadian; O1, Whiterockian to Blackriverian; O2, Trentonian to Edenian; OS1, Maysvillian to Wenlock; SD1, Ludlow to Gedinnian; D1, Seigenian to Eifelian; D2, Givetian to Famennian. (From Sanford (1987).)



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FIG. 20  Chronostratigraphic diagram of the Hudson Platform area. (From Roksandic (1987).)

sequence is composed of mixed siliciclastic and carbonate strata, with Upper Devonian black bituminous shales developed in the southeast of the basin, possibly sourced from the nearby Appalachian orogen (Sanford, 1987). Kaskaskia deposition was terminated by marine regression, and the Hudson Platform was subjected to extensive erosion throughout the rest of Phanerozoic time (Fig. 20), although it is not clear how much Absaroka strata was deposited during Mississippian and Pennsylvanian time and subsequently eroded. Continental deposits are only preserved locally on the platform (Sanford, 1987). If the uplift in the basin occurred in Mesozoic times, it may have been due to uplift of eastern North America over a dynamic topographic high developed by Pangean mantle insulation (Burgess et al., 1997).

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SUMMARY ●

● ●

●

●

The North American craton is not simply an unchanging, stable platform accumulating strata and influenced only by changes in global sea level. Rather, viewed on a time scale of tens to hundreds of millions of years, it is a dynamic tectonic environment influenced by various plate tectonic and mantle processes. The Sloss cratonic sequences record the history of this dynamic tectonic environment. Variations in dynamic topography generated by subducting lithospheric slabs, thermal insulation of mantle beneath supercontinents, and plate motion over convective upwellings and downwellings can explain much of the large-scale architecture of the Sloss sequences, but more detailed plate tectonic reconstructions and associated mantle convection models are necessary to further test and develop these explanations. Intraplate stress clearly played a large role in generating the cratonic sequences by reactivating preexisting structures and driving subsidence and uplift. Variations in intraplate stress through time can be related to some degree to orogenic and other plate tectonic events occurring on the cratonic margins. Given present available evidence and theory, the North American cratonic basins seem most likely to be due to a combination of long-term low strain rate lithospheric extension and focused intraplate stress variations, in some cases with an element of long-wavelength tilting due to subduction-induced dynamic topography. This is by no means a definitive conclusion, however. Much work remains to be done to test and confirm or refute these ideas.

ACKNOWLEDGMENTS Thanks to Mike Gurnis who nurtured my interest in the North American cratonic sequences and the influence of mantle dynamics, and a posthumous acknowledgment to Larry Sloss for his monumental and groundbreaking work in describing, defining, and understanding North American cratonic sequences. Thanks to Andrew Barnett, Wes Gibbons, Octavian Catuneanu, Brian Pratt, and Andrew Miall for reading and improving previous drafts of this work.

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