Palaeogeography, Palaeoclimatology, Palaeoecology 531 (2019) 109259
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Review article
Global events of the Late Paleozoic (Early Devonian to Middle Permian): A review
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Wenkun Qiea, , Thomas J. Algeob,c,d, , Genming Luob, Achim Herrmanne ⁎
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a
CAS Key Laboratory of Economic Stratigraphy and Palaeogeography, Nanjing Institute of Geology and Palaeontology and Center for Excellence in Life and Paleoenvironment, Chinese Academy of Sciences, Nanjing 210008, China b State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences, Wuhan 430074, China c State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Wuhan 430074, China d Department of Geology, University of Cincinnati, Cincinnati, OH 45221, USA e Department of Geology and Geophysics, Coastal Studies Institute, Louisiana State University, Baton Rouge, LA 70803, USA
ARTICLE INFO
ABSTRACT
Keywords: Late Paleozoic Ice Age LPIA Land plants Paleoclimate Paleoceanography Mass extinction
The Late Paleozoic (Early Devonian to Middle Permian) was an interval of profound changes in Earth-surface systems, reflected in dynamic interplay among the biosphere, hydrosphere, atmosphere, and geosphere. Major events transpired, including the colonization of landmasses by vascular plants, the assembly of the supercontinent Pangea, two first-order mass extinctions (the Frasnian-Famennian and Devonian-Carboniferous boundary events), and the most severe icehouse climate mode of the Phanerozoic (the Late Paleozoic Ice Age, LPIA). The goals of the present review are (1) to summarize major global developments in climate, oceanography, and paleobiology during the Late Paleozoic, (2) to examine the roles of land plant evolution, global tectonics, and large igneous province magmatism in driving these developments, and (3) to serve as an introduction to the 23 contributions to this special issue of Palaeogeography, Palaeoclimatology, Palaeoecology, indicating how they advance our knowledge of various scientific issues related to these developments.
1. Introduction During the Early Devonian to Middle Permian (~410 to 260 Ma), the colonization of landmasses by vascular plants and the assembly of the supercontinent Pangea had profound effects on the Earth, leading to dynamic interactions among its climatic, environmental, and biotic systems (Algeo and Scheckler, 1998; Montañez and Poulsen, 2013; Wilson et al., 2017; Chen et al., 2018a). The Devonian and earliest Carboniferous were characterized by rapid expansion of terrestrial floras, frequent marine anoxic episodes and biotic crises, and long-term climatic cooling (Algeo et al., 1995; Joachimski et al., 2009; Fig. 1). These trends culminated in two first-order mass extinctions, the Kellwasser Event at the Frasnian/Famennian boundary (FFB) and the Hangenberg Event at the Devonian/Carboniferous boundary (DCB), which were marked by pronounced climate instability, glacio-eustatic lowstands, and perturbations of marine biogeochemical cycles (Kaiser et al., 2016; Huang et al., 2018a, 2018b, 2018c, 2018d). Atmospheric CO2 concentrations declined through most of the Devonian (Royer et al., 2004; Berner, 2006), falling to near-modern levels (280 ppm) for
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the first time in Earth history by the Early Carboniferous (Foster et al., 2017; Fig. 1). Reduced greenhouse gas concentrations were a major factor in initiation of the Late Paleozoic Ice Age (LPIA, ~340–285 Ma). Southern Hemisphere (Gondwanan) glaciation resulted in large glacioeustatic fluctuations, reorganization of ocean circulation, and bio-evolutionary events such as the mid-Carboniferous and Sakmarian-Artinskian crises (Saltzman, 2003; McGhee et al., 2012; Wang et al., 2013). Significantly, the Early Devonian to Middle Permian interval experienced climate-associated changes in environments and biotas at a scale commensurate with impending present-day global changes and, thus, is of relevance for anticipating the consequences of present and future climate change (Karl and Trenberth, 2003; Ceballos et al., 2015). This special issue of Palaeogeography, Palaeoclimatology, Palaeoecology is thematically dedicated to ‘Global Events of the Devonian to mid-Permian: Prelude and Progression of the Late Paleozoic Ice Age’. These studies cover a wide range of sites in North America, Europe and China providing a near-global perspective on Late Paleozoic events (Fig. 2). Given the large number of contributions (23) to this special issue, we have divided them into two volumes organized
Corresponding author. Correspondence to: T.J. Algeo, State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences, Wuhan 430074, China. E-mail addresses:
[email protected] (W. Qie),
[email protected] (T.J. Algeo).
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https://doi.org/10.1016/j.palaeo.2019.109259
Available online 05 July 2019 0031-0182/ © 2019 Published by Elsevier B.V.
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Fig. 1. Late Paleozoic global records. (A) Orogenic events (gray), global climate conditions (red-purple-blue), and occurrence of continental icesheets (thin blue) (Montañez and Poulsen, 2013). (B) Conodont δ18O record and bioevents (Buggisch et al., 2008; Joachimski et al., 2009; Chen et al., 2016). (C) Carbonate δ13C record (Saltzman and Thomas, 2012). (D) Seawater strontium isotope record (blue = McArthur et al., 2012; red = Chen et al., 2018a). (E) Terrestrial organic matter burial flux (North America only) and atmospheric CO2 concentrations (Nelsen et al., 2016; Foster et al., 2017). Abbreviations: EPME = end-Permian mass extinction, ESA = eastern South America, WSA = western South America, SA = southern Africa, EA = eastern Australia, WA = western Australia, AR = Arabian Peninsula, TA = Tasmania, ANT = Antarctica, IN = India, WE. = wetland, TR. = transitional, S.D. = seasonally dry. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)
around broad time intervals. The first volume contains 11 studies focused on the Devonian to early Mississippian, and the second volume contains a further 12 studies examining the mid-Mississippian to Middle Permian (Fig. 2). In this introductory review paper, we (1) undertake an up-to-date review of research on Late Paleozoic global events, examining the patterns, processes, causes, and consequences of major developments in tectonics, geography, oceanography, climate, and the biosphere, and (2) put the contributions of this special issue into the context of contemporaneous events, examining how each study makes an addition to our expanding knowledge of co-evolution of the Earth system during the Late Paleozoic. The 23 studies contained in the two volumes of this special issue provide broad coverage of topics related to Late Paleozoic global events. Grouped topically, these studies investigate the following broad themes:
2) 3) 4) 5) 6) 7)
1) the timing and style of glaciations in west-central Gondwana (Buso
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et al., 2017; Fallgatter and Paim, 2017; Fedorchuk et al., 2018; Griffis et al., 2018), sedimentary and eustatic responses to the LPIA (Sardar Abadi et al., 2018; Chen et al., 2018b), ocean circulation patterns and paleobiogeographic changes (Davydov and Cózar, 2017; Herrmann et al., 2018; Huang et al., 2018c; Shen et al., 2018), biogeochemical cycling in Late Paleozoic oceans (Liu et al., 2018b; Tuite et al., 2019; Turner et al., 2018; Zhang et al., 2018), event analysis of the Frasnian-Famennian boundary (Chang et al., 2017; Uveges et al., 2018) and Devonian-Carboniferous boundary (Kalvoda et al., 2019), refinements to Devonian biostratigraphy using conodonts (Zhang et al., 2019) and palynomorphs (Liu et al., 2018a), taxonomy and habitats of Late Paleozoic land plants (Guo et al., 2017; Tewari et al., 2018; Wan et al., 2019; Xu et al., 2018).
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Fig. 2. Paleogeographic and stratigraphic distribution of studies in this special issue. The time scale is after Gradstein et al. (2012) and Ogg et al. (2016), and paleogeographic maps are from Ron Blakey (www2.nau.edu/rcb7). Dashed stage boundaries have not been formally defined. Numbered bars (yellow = near-field studies; blue = far-field studies) are keyed to individual papers: 1–Wan et al. (2019), 2–Xu et al. (2018), 3–Shen et al. (2018), 4–Zhang et al. (2019), 5–Chang et al. (2017), 6–Uveges et al. (2018), 7–Tuite et al. (2019), 8–Guo et al. (2017), 9–Kalvoda et al. (2019), 10–Liu et al. (2018a), 11–Liu et al. (2018b), 12–Griffis et al. (2018), 13–Fallgatter and Paim (2017), 14–Buso et al. (2017), 15–Fedorchuk et al. (2018), 16–Chen et al. (2018b), 17–Herrmann et al. (2018), 18–Turner et al. (2018), 19–Davydov and Cózar (2017), 20–Huang et al. (2018c), 21–Sardar Abadi et al. (2018), 22–Zhang et al. (2018), 23–Tewari et al. (2018). Papers 1–11 are in volume 531 Part A, and papers 12–23 are in volume 531 Part B. Abbreviations: SC = South China, AP = Appalachian Basin, MC = Mid-Continent, BEL = Belgium, RTG = Rheic-Tethys gateway, KZ = Kazakhstan, IR = Iran, IN = India. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)
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2. Global developments of the Devonian and Early Mississippian: prelude to an ice age
between 30°S and 45°S paleo-latitude, reflecting a major increase of alpine glaciation. The development of widespread glaciation in the latest Devonian is consistent with a major decline in atmospheric CO2 at that time (Royer et al., 2004; Berner, 2006). These climatic changes had consequential effects on marine systems, including major perturbations of the C-N-S cycles (Simon et al., 2007; Qie et al., 2015; Liu et al., 2016), expansion of phosphorite and radiolarite deposition (Racki, 1999), contraction of the latitudinal range of metazoan reefs (Copper and Scotese, 2003) and other biota such as calcareous foraminifera (Kalvoda, 1986), and ultimately the end-Devonian mass extinction (Kaiser et al., 2008, 2016). Conodont faunas provide evidence of continuous cooling through the Famennian Stage (Girard et al., 2018). The earliest Mississippian represents a transition from warmer Devonian conditions to the glacial world of the LPIA. Marine C and N isotope records exhibit large perturbations during the mid-Tournaisian (the “Tournaisian carbon-isotope excursion”, or TICE) and mid-Visean (the “Visean carbon-isotope excursion”, or VICE) that correlate with cooling events (Buggisch et al., 2008) and that are suggestive of major changes in ocean circulation and marine biogeochemical cycles (Yao et al., 2015; Liu et al., 2018b–this volume; Fig. 2).
2.1. Long-term trends The Devonian was a period of major atmospheric, climatic, and environmental changes, many of which can be linked to the spread of vascular land plants (Algeo et al., 2001; see Section 5.1) and to the incipient assembly of the supercontinent Pangea (Kidder and Worsley, 2004; see Section 5.2). Atmospheric CO2 levels declined from > 10 PAL (present atmospheric level) to ~1 PAL (Royer et al., 2004; Berner, 2006), falling to close to (pre-industrial) modern levels (280 ppm) for the first time in Earth history by the Early Carboniferous (Foster et al., 2017; Fig. 1). Modeling estimates of pCO2 decline have been independently confirmed by soil carbonate C-isotopes (Mora et al., 1996; Ekart et al., 1999) and land-plant stomatal densities (Chaloner and McElwain, 1997; Willis and McElwain, 2014). The long-term pCO2 decline during the Devonian was due largely to enhanced burial of marine organic matter and weathering-related consumption of atmospheric CO2 (Algeo et al., 1995, 2001; Berner, 1997), with enhanced burial of terrestrial organic carbon in coal swamps becoming more important during the Carboniferous (Berner and Raiswell, 1983). Weathering in upland areas contributed more to CO2 consumption than in lowland areas, and the combination of widespread orogenic activity and colonization of upland areas by arborescent vegetation may have been integral to the steep decline of atmospheric CO2 leading to the LPIA (Falcon-Lang, 2004). A concurrent and causally interrelated rise in atmospheric pO2 occurred during the Devonian, with an especially rapid increase during the Famennian to Early Carboniferous (Algeo and Ingall, 2007; Glasspool and Scott, 2010). This rise has now been documented by multiple proxies, including Mo isotopes (Dahl et al., 2010), Ce anomalies (Wallace et al., 2017), and carbonate I/Ca ratios (Lu et al., 2018), with further support from modeling studies (Krause et al., 2018). Rising atmospheric pO2 led to more frequent wildfires, as indicated by sedimentary charcoal records (Falcon-Lang, 2000; Scott and Glasspool, 2006; Glasspool et al., 2015), and to a progressive increase in average marine animal sizes (Heim et al., 2015). Declining atmospheric pCO2 during the Devonian was a major cause of global climatic cooling. Average temperatures cooled almost continuously from the Early Devonian through the Late Mississippian, with the exception of a short-term warming trend in the late Givetian to early Frasnian (Buggisch et al., 2008; Joachimski et al., 2009; Chen et al., 2016; Fig. 1). Prior to the FFB, global climate was characterized by a low pole-to-equator temperature gradient, reef growth to paleolatitudes of 40–50°, and relatively humid conditions at mid to high latitudes (Copper and Scotese, 2003). A major cooling event occurred at the Frasnian-Famennian boundary (FFB), as shown by positive shifts in marine δ18O records of up to 3–4‰ (Joachimski and Buggisch, 2002) (Fig. 1). In the eastern Paleotethys Ocean, peri-equatorial sea surface temperatures (SSTs) are estimated to have declined by ~6 °C during the FFB event (Huang et al., 2018b). The FFB event was accompanied by a ~3‰ increase of δ13Ccarb and widespread black shale deposition, pointing to links between carbon burial and climatic cooling (Buggisch and Joachimski, 2006). Although near-field sedimentological evidence of Gondwanan glaciation at this time is lacking, a pronounced eustatic fall at the FFB, recorded inter alia by widespread subaerial exposure of continental shelves and local evaporite deposition, is thought to have been due to contemporaneous glaciation (Devleeschouwer et al., 2002; Isaacson et al., 2008; Song et al., 2017). Preceding the LPIA, the earliest known Gondwanan glacial deposits are of latest Famennian age, representing short-lived (< 300 kyr), small-scale alpine glaciation in Bolivia, Peru, Brazil, and South Africa (Caputo, 1985; Caputo and Crowell, 1985; Díaz-Martínez and Isaacson, 1994; Caputo et al., 2008; Isaacson et al., 2008; Fig. 1). In the Appalachian Basin, Brezinski et al. (2008) reported glacigenic diamictites from Pennsylvania, Maryland, and West Virginia that were deposited
2.2. Short-term events The Devonian Period was characterized by a series of short-term events that were broadly similar, marked by concurrent shifts toward greater marine anoxia, deposition of black shales, positive excursions in C and S isotope records, and climatic cooling (Algeo et al., 1995; Goddéris and Joachimski, 2004; Buggisch and Joachimski, 2006; Van Geldern et al., 2006; Becker et al., 2012; Fig. 1). The two largest and best-studied of these events are the ~372-Ma FFB Kellwasser Event (Riquier et al., 2006; Girard and Renaud, 2007) and the ~359-Ma DCB Hangenberg Event (Kaiser et al., 2008, 2011, 2016), but the series began with the ~400-Ma (Emsian) Daleje Event (Becker and Aboussalam, 2011; Tonarová et al., 2017) and included at least 5–6 additional, widely recognized events (Algeo et al., 1995; DeSantis and Brett, 2011; Becker et al., 2012). The association of these events with positive δ13Ccarb excursions is well-established (Joachimski et al., 2001, 2002; Song et al., 2017; Fig. 1), suggesting that each event was driven by episodes of enhanced burial of organic carbon, although the nature of the trigger for carbon burial remains in debate (see Section 5). The Frasnian-Famennian boundary (FFB) crisis consists of two events, recorded by the Lower and Upper Kellwasser Horizons (LKH and UKH) of the late rhenana and linguiformis conodont zones, respectively (Racki et al., 2002; Fig. 1). These horizons are characterized by local organic carbon enrichment, shifts toward more reducing watermass conditions, and positive δ13C and δ18O excursions (Joachimski et al., 2001, 2002). The FFB crisis was once viewed as driven by eustatic transgression, leading to black shale deposition (Buggisch, 1991), but subsequent work has shown that enhanced organic carbon burial was the driver, leading to global cooling, continental ice formation, and subsequent eustatic regression (Algeo et al., 1995; Joachimski et al., 2001, 2002). The FFB mass extinction eliminated ~20% of families and ~50% of genera among marine invertebrates (Sepkoski, 1996; Alroy et al., 2008). The mass extinction particularly affected warm tropical marine faunas, and it decimated the Middle Paleozoic reef community consisting of tabulate corals and stromatoporoids (Copper, 2002). The mass extinction is thought to have been caused by some combination of oceanic anoxia (Algeo et al., 1995; Bond et al., 2013) and climatic cooling (Copper, 1986; Joachimski and Buggisch, 2002), both of which were most severe in tropical marine areas. In the present volume, the timing of the FFB event relative to carbon-cycle perturbations is explored through high-resolution correlation of conodont biostratigraphic data and carbon-isotope records (Chang et al., 2017–this volume; Fig. 2). The Devonian-Carboniferous boundary (DCB) crisis is represented 4
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by the uppermost Famennian Hangenberg beds, recording a sequence of events (Kaiser et al., 2016; Fig. 1). The lower crisis interval is marked by a minor eustatic fall, the onset of widespread black shale deposition (Hangenberg Black Shale in Europe; Cleveland Shale in Appalachian Basin; Lower Bakken Shale in Williston Basin; Algeo et al., 2007), and a positive carbon isotope excursion. It is the main extinction level for marine invertebrates such as ammonoids, trilobites, conodonts, stromatoporoids, and corals, with ~16% of marine families and ~21% of marine genera becoming extinct (Sepkoski, 1996; Alroy et al., 2008). This event is marked by concurrent changes in marine microfaunal and phytoplankton assemblages (Kalvoda, 1986; Riding, 2009). The middle crisis interval was characterized by a major eustatic fall (probably > 100 m) that led to progradation of shallow-water siliciclastics (Pashin and Ettensohn, 1995) as well as widespread unconformities (Over et al., 2019). The glacio-eustatic origin of this regression is demonstrated by miospore correlation with widespread diamictites of South America and Africa (Streel et al., 2000) and by evidence of tropical mountain glaciers in eastern North America (Brezinski et al., 2008). The upper crisis interval is characterized by initial post-glacial transgression, a second global carbon isotope spike, and opportunistic faunal blooms (Kaiser et al., 2016). Both the FFB and DCB were characterized by major eustatic falls. At the FFB, the regressive part of the IId(2) T-R cycle terminated within the Lower triangularis Zone, and a major transgression began at the base of the IIe T-R cycle in the Middle triangularis Zone (Johnson et al., 1985; Brett et al., 1990; Smith and Jacobi, 2001; Over, 2002). The DCB is characterized by an even larger eustatic fall, recorded by prograding lowstand siliciclastics of the Hangenberg Sandstone in Germany (Kaiser et al., 2016), the Bedford-Berea succession in the Appalachian Basin (Bjerstedt and Kammer, 1988; Pashin and Ettensohn, 1995; Cecil et al., 2004), and the middle siltstone (or sandstone) members of the Bakken and Exshaw formations in the Williston and Elk Point basins of western North America (Playford and McGregor, 1993; Hartwell, 1998; Caplan and Bustin, 2001). The ensuing earliest Carboniferous transgression resulted in deposition of an extensive black shale unit across the North American craton, comprising the Sunbury Shale and upper Gassaway Member of the Chattanooga Shale in the Appalachian Basin, the upper Clegg Creek Member of the New Albany Shale in the Illinois Basin, the Upper Black Shale members of the Bakken and Exshaw formations in the Williston and Elk Point basins, and stratigraphic equivalents elsewhere. At present, placement of the DCB is inconsistent at a global scale with respect to these eustatic markers: it is regarded as coincident with the peak eustatic regression in North America (Sandberg et al., 2002; Algeo et al., 2007) but with the post-glacial transgression in Europe (Kaiser et al., 2008, 2016; Becker et al., 2012; see review in Corradini et al., 2017). Two studies in the present volume provide evidence of major changes in the marine nitrogen cycle in conjunction with black shale deposition during the Late Devonian. Uveges et al. (2018–this volume; Fig. 2) document minimum δ15N values of −2 to +1‰ in Appalachian Basin black shales during the FFB event, which is consistent with widespread cyanobacterial (diazotrophic) nitrogen fixation (n.b., a similar negative δ15N shift was reported from near the DCB; Martinez et al., 2019). This pattern is a hallmark of nitrogen limitation of marine primary productivity, which can be triggered by enhanced phosphorus availability. Increased phosphorus availability at the FFB was the product of some combination of enhanced organic P recycling under euxinic conditions (Ingall et al., 1993; Algeo and Ingall, 2007), leading to increased P remineralization fluxes (Murphy et al., 2000a, 2000b; Sageman et al., 2003), and subaerial weathering and riverine delivery of new P to the oceans (Algeo et al., 1995; Lenton, 2001). Tuite et al. (2019–this volume; Fig. 2) study C-N-P-S-Fe systematics in two Upper Devonian black shale intervals from the Appalachian Basin: (1) the Lower and Middle Huron members of the Ohio Shale, which are of early to mid-Famennian age, and (2) the Cleveland Shale Member of the same formation, which is of latest Famennian age. Critically, these two
intervals are interpreted to represent contrasting paleoclimatic conditions, with the former associated with a greenhouse regime and the latter with an (incipient) icehouse regime. The greenhouse interval was characterized by low O2 and pH, strong recycling of N and P, and limited oxidation of ammonia (i.e., nitrification), resulting in an ammonium-dominated water column and low sediment δ15N values. In contrast, the icehouse interval featured higher O2 and pH, more limited recycling of N and P, and strong water-column nitrification, resulting in a nitrate-dominated system and higher sediment δ15N values. These findings highlight the influence of the marine N cycle on organic carbon fluxes and biogeochemical cycling, and its potential role in modulation of long-term climate change (cf. Algeo et al., 2014). 3. The Late Paleozoic Ice Age (LPIA) 3.1. Timing and extent of continental glaciation The end-Devonian glaciation was just the first event in a series of glaciations that extended throughout the Carboniferous and Early Permian, now known as the Late Paleozoic Ice Age (Davydov et al., 2012). The earliest Carboniferous glaciation may have been a transient event around the Kinderhookian-Osagean boundary, i.e., mid-Tournaisian (Saltzman, 2002; Richardson and Ausich, 2004; Fig. 1). Following gradual cooling and possible short-lived glacial events during the late Tournaisian and early Visean (Buggisch et al., 2008; Liu et al., 2018b–this volume; Fig. 2), the main phase of the LPIA, marked by sustained growth of large continental ice sheets in Gondwana, began in the middle to late Visean (Smith and Read, 2000; Fielding et al., 2008a, 2008b; Bishop et al., 2009; Limarino et al., 2014). This event coincided with the onset of sea-level oscillations with amplitudes of 20 to 100 m at orbital timescales, widespread deposition of low-latitude cyclothems, and increased continental weathering (Waters and Condon, 2012; Chen et al., 2018b; Fig. 1). Conodont δ18O records show a series of positive shifts in the late Tournaisian that reached a relatively invariant plateau in the Visean (Buggisch et al., 2008; Chen et al., 2016; Fig. 1), documenting global cooling and extensive icesheet growth, with development of a new quasi-stable climate mode during the Visean. During the LPIA, multiple small- to moderate-size icesheets, ice caps, and alpine glaciers existed on the Gondwanan supercontinent (Isbell et al., 2003; Fielding et al., 2008a, 2008b; Montañez and Poulsen, 2013; Fig. 1). The histories of growth and decay of these icesheets, while not completely known, is thought to reflect significant diachroneity, with a shift in the main locus of glaciation from west to east across Gondwana as it traversed the South Pole during the Carboniferous. This scenario has gained acceptance over older models depicting a single, extensive ice sheet centered over Gondwana (e.g., Crowley and Baum, 1991). Refined biostratigraphic frameworks and new radioisotopic age data have assisted in revising the timing and extent of growth of individual icesheets (Isbell et al., 2003; Fielding et al., 2008a, 2008b; Lakin et al., 2016; Griffis et al., 2017), facilitating an understanding of the glacial dynamics and climatic forcings and feedbacks that operated during the LPIA (Horton et al., 2007, 2010, 2012; Montañez and Poulsen, 2013; Lowry et al., 2014). For example, although the relationship of glacial-interglacial cycles to atmospheric CO2 changes during the LPIA was long uncertain (Horton and Poulsen, 2009), recent work has begun to quantify atmospheric pCO2 variation at this timescale (Montañez et al., 2016). The LPIA varied in intensity over its ~60-Myr duration (Fielding et al., 2008a, 2008b; Montañez and Poulsen, 2013). Some glacial phases may have been quite cold, as suggested by possible evidence of tropical glaciers (Soreghan et al., 2014), although the existence of lowland glaciation during the LPIA was not supported by GCM modeling (Heavens et al., 2015). Based on the geographic extent of nearfield glacial deposits, continental ice mass is thought to have reached an initial peak in the Middle Pennsylvanian (Desmoinesian; Moscovian), contracted to a minimum in the Late Pennsylvanian (Missourian5
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Virgilian; Kasimovian-Gzhelian), and expanded to a second maximum in the Early Permian (early Wolfcampian; Asselian-Sakmarian) (Isbell et al., 2003, 2012). However, González-Bonorino and Eyles (1995) proposed an inverse relation between ice mass and glacial deposit extent for the LPIA, arguing that icesheets reached their maximum size in the Early Pennsylvanian (Morrowan-Atokan; Bashkirian) but left a meager record owing to destruction by later glacial advances. This hypothesis has received support from the recent high-resolution conodont O-isotope study of Chen et al. (2016), who determined that peak δ18O values for the LPIA interval were reached in the Bashkirian. Although marine fossil δ18O records do not distinguish between shifts due to climatic cooling (temperature effect) and continental ice mass (seawater δ18O effect), these influences generally covary positively (e.g., Shackleton, 1987), supporting interpretation of the Bashkirian δ18O peak as evidence of maximum continental ice mass during the Early Pennsylvanian. However, most Paleozoic oxygen-isotope records come from epeiric seas, which are subject to large salinity fluctuations (e.g., Rosenau et al., 2014; Joachimski and Lambert, 2015; Montañez et al., 2018), introducing a complication to the interpretation of δ18O data as a record of glacial intensity. Shifting ice centers resulted in a diachronous termination of the LPIA. In southwestern Gondwana (Argentina), there is no evidence of glaciation after the Early Pennsylvanian (Bashkirian) (Gulbranson et al., 2010), whereas in central-western Gondwana (Brazil, southern Africa), major glaciation persisted into the late Early Permian (Cisuralian), with strong deglaciation accompanied by a sharp rise in atmospheric pCO2 in the mid-Artinskian (Montañez et al., 2007; Frank et al., 2008). In eastern Gondwana, continental glaciation may have continued longer (Fielding et al., 2008a, 2008b; Fang et al., 2017; Garbelli et al., 2019), although these deposits require improved age control (Metcalfe et al., 2015). The Middle/Late Permian boundary coincided with a second-order mass extinction event (the Guadalupian-Lopingian boundary crisis; Clapham et al., 2009; Xie et al., 2017; Huang et al., 2018a, b, c, d; Fig. 1) that has been linked to the eruption of the Emeishan Large Igneous Province (Zhou et al., 2002; Shellnutt, 2014). The record of this event in South China was probed Zhang et al. (2018–this volume; Fig. 2), who generated a high-resolution, multiproxy marine productivity and redox history across the mid-Capitanian mass extinction horizon. They demonstrated gradual increases in productivity and the intensity of oceanic euxinia during the GLB crisis, followed by a shift to oxic shallow-water environments with low productivity in its aftermath, consistent with control by Emeishan LIP magmatism.
record more distal sediment sources, linked to development of complex surface fluvial and glacio-fluvial drainage systems. Icesheets in southern and western Gondwana emanated from at least four large ice centers, the Windhoek Highlands in southern Africa, the Congo and Tanzania in eastern Africa, an Uruguayan ice center, and the Ellsworth Mountains in Antarctica. Buso et al. (2017–this volume; Fig. 2) investigated glacial deposits in the Upper Itararé Group of Santa Catarina State in southern Brazil, correlating three major glacial cycles to subsurface records in the Paraná Basin. They identified five successive glacial subcycles and their component glacial advance and deglacial systems tracts, providing important clues to the durations of both long-term and short-term glacial cycles of the LPIA. Fallgatter and Paim (2017–this volume; Fig. 2) studied outcrops of the Itararé Group in the Alfredo Wagner area, along the eastern margin of the Paraná Basin in southern Brazil. They assembled an 8-km-long, strike-oriented panel correlation showing vertical stacking of six facies associations as well as the morphology of the basal nonconformity. They infer that the deglacial deposits of the Itararé Group were influenced by topographically controlled ice streams, instead of non-confined, widespread ice lobes as previously suggested, and that the ice streams flowed outward via topographically controlled corridors from an ice cap centered in southern Africa. Fedorchuk et al. (2018–this volume; Fig. 2) examine the sedimentology, stratigraphy, and provenance of sediments infilling the Mariana Pimentel and Leão paleovalleys on the Rio Grande do Sul Shield of southernmost Brazil, investigating regional sediment sources and reactivation of Neoproterozoic basement structures during the Carboniferous and Permian. Their detrital UePb zircon age analysis yielded a single population of Neoproterozoic ages (ca. 800–550 Ma) that is unrelated to eastern African basins. On this basis, they reject the hypothesis of a single, massive icesheet flowing westward out of Africa and, instead, infer that glacial sediments on the southern margin of the Paraná Basin were the product of a separate icesheet lobe that advanced northward from Uruguay (cf. Fedorchuk et al., 2019). 4. Consequences of Late Paleozoic Ice Age 4.1. Oceanic-climatic effects The dynamic nature of the LPIA triggered near-continuous climatic and oceanic changes. The waxing and waning of Gondwana icesheets led to glacio-eustatic fluctuations, temperature fluctuations, and perturbations of ocean circulation and marine biogeochemical cycles (Grossman et al., 2008; Chen et al., 2016). The scale of sea-level fluctuations driven by late Paleozoic glaciations has been a matter of debate. Glacio-eustatic fluctuations of at least 100 m have long been inferred to account for observations of widespread water-column stratification, deep-water euxinia, and site-specific sequence stratigraphic relationships in low-paleolatitude marine deposits (Heckel, 1977; Soreghan and Giles, 1999; Algeo and Heckel, 2008; Joachimski et al., 2006; Rygel et al., 2008; Elrick and Scott, 2010). Substantial eustatic fluctuations are also implied by repeated cycles of bauxite formation in South China and elsewhere (Weng et al., 2018; Yu et al., 2019). However, studies based on near-field glacial deposits have argued that sea-level fluctuations greater than ~50 m were not possible given evidence for numerous, small icesheets rather than a single large one, as well as the lack of contemporaneous Northern Hemisphere continental icesheets (Isbell et al., 2003, 2012, 2016; Montañez and Poulsen, 2013). Further research will be needed to resolve this issue. The links between LPIA glaciation, as represented by near-field glacial sedimentary records, and contemporaneous changes in oceanic environments and ecosystems, as represented by far-field marine stratigraphic records, were complex and remain incompletely understood (e.g., Frank et al., 2008). In this volume, such links are explored in six contributions, including studies of Mississippian platform carbonates in
3.2. Field studies of glaciation in western Gondwana With movement of the Gondwanan Plate across the South Pole, continental glaciation began in the latest Famennian in central South America and Africa (Fielding et al., 2008a, 2008b). Growth of icesheets in western Argentina commenced in the Visean (mid-Mississippian) and may have subsequently shifted to central-western Gondwana (Brazil and southern Africa) (Gulbranson et al., 2010, 2015; Montañez and Poulsen, 2013). However, the timing, extent, and style of continental glaciation in Gondwana remain incompletely known. Four studies in the present volume (Griffis et al., 2018–this volume; Buso et al., 2017–this volume; Fallgatter and Paim, 2017–this volume; Fedorchuk et al., 2018–this volume) provide new information concerning the Late Paleozoic glacial history of South America. Griffis et al. (2018–this volume; Fig. 2) report new UePb ages and Hf isotope compositions of detrital zircons recovered from diamictites in two key areas of glaciation, the Paraná Basin of southern Brazil and the Tepuel Basin of southern Argentina. Differences in UePb ages are interpreted as due to changes in the spatial distribution and volume of icesheets through the glacial history of this region. Older (Middle Mississippian to Middle Pennsylvanian) glacial deposits were iceproximal, whereas deposits of the final deglaciation in the Paraná Basin 6
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South China (Chen et al., 2018b–this volume) and coral biostromes in central Asia (Huang et al., 2018a, b, c, d–this volume), MississippianPennsylvanian boundary foraminiferal assemblages (Davydov and Cózar, 2017–this volume), Upper Pennsylvanian cyclothemic units in North America (Herrmann et al., 2018–this volume; Turner et al. (2018–this volume), and Pennsylvanian-Permian shallow-water carbonates in Arabia (Sardar Abadi et al., 2018–this volume). Whereas Pennsylvanian and Lower Permian cyclothems are wellstudied, cyclic sedimentary successions of Mississippian age produced through glacio-eustatic forcing have received comparatively less attention. Chen et al. (2018b–this volume; Fig. 2) present a detailed analysis of sedimentary facies and depositional cycles from MiddleUpper Mississippian shallow-water platform and carbonate slope successions in the tropical Paleo-Tethys Ocean, in order to assess far-field sedimentary responses to early stages of LPIA glaciation. Four upper Visean cycles and four Serpukhovian cycles were identified in the carbonate platform Yashui section, which correlate well with their counterparts in deep-water slope successions of South China and with shallow-water cyclothems in the Donets Basin, Ukraine (Eros et al., 2012), providing evidence of a glacio-eustatic driver. The final closure of the peri-equatorial Rheic Ocean between Laurussia and Gondwana had major effects on global climate, ocean circulation, and paleobiogeographic evolution, but the timing of this event is not well-constrained. During the Mississippian-Pennsylvanian transition, carbon isotope ratios show a major positive shift, with an increase of up to 3.0‰ in the Paleo-Tethyan region and divergence of δ13C values between North America and Europe (Bruckschen et al., 1999; Mii et al., 2001). This pattern is thought to reflect enhanced interoceanic variability due to closure of oceanic gateways during the assembly of Pangea (Saltzman, 2003). Davydov and Cózar (2017–this volume; Fig. 2) compiled warm-water benthic foraminifera (WWBF) data along the Rheic-Tethyan passage from 206 literature sources as well as their own collections, providing new constraints on the timing of oceanic gateway closure. Their paleobiogeographic analysis indicates that the Alleghenian Isthmus between North America and the Western Tethys appeared during the Mississippian-Pennsylvanian transition, resulting in highly endemic foraminiferal assemblages in North America and Europe by the early Bashkirian (earliest Pennsylvanian). They conclude that closure of the Rheic-Tethyan gateway was a major factor in intensification of LPIA glaciation during the Bashkirian. Shifting of continents and closure of oceanic gateways markedly altered oceanic circulation patterns during the Carboniferous, leading to changes in heat transport and regional climatic conditions. Huang et al. (2018a, b, c, d–this volume; Fig. 2) document upper Visean (Middle Mississippian) coral biostromes in the Yamansu Formation of the eastern Tianshan area (Xinjiang Province, western China) that flourished in an island arc setting on the margin of the Kazakhstan Plate at mid-Northern Hemisphere paleolatitudes. Twenty-two coral genera were identified, although a single fasciculate coral species (Siphonodendron irregular) dominates the 0.2- to 5-m-thick biostromes. Compared to widespread Siphonodendron biostromes in the tropical Paleotethys region, formation of the Yamansu Formation biostrome was controlled not only by intrinsic factors but also by paleoceanic changes―specifically, northward redirection of warm ocean currents that allowed Siphonodendron and other rugose corals to thrive at higher latitudes. These changes in oceanic circulation were presumably driven by progressive narrowing of peri-equatorial oceanic gateways prior to their final closure around the Mississippian-Pennsylvanian boundary. In the North American Midcontinent, Middle to Upper Pennsylvanian successions are composed of cyclic alternations of thin transgressive limestones, offshore gray to black phosphatic shales, thick regressive limestones, and nearshore to terrestrial shales with paleosols and coal beds (Heckel, 1977, 1994; Herrmann et al., 2015; Algeo and Herrmann, 2018). These “Kansas-type cyclothems” were deposited in the broad, semi-restricted epeiric Midcontinent Sea, within which largescale estuarine-type (“superestuarine”) circulation operated (Algeo and
Heckel, 2008; Algeo et al., 2008a, 2008b). Turner et al. (2018–this volume; Fig. 2) refine the superestuarine circulation model through geochemical and mineralogical analysis of the Upper Pennsylvanian Heebner Shale and its stratigraphic equivalents at seven sites across the Midcontinent and Illinois basins. Their dataset documents spatial patterns of variation in sediment composition that provide insights into specific aspects of internal circulation within the Midcontinent Sea. Clay and silt fraction proxies and major-element ratios suggest counterclockwise gyral circulation, which was probably caused by a combination of upwelling on the southern Midcontinent Shelf margin and discharge of large rivers into the NAMS. Physical separation of the deep watermasses of the Midcontinent and Illinois basins by the submarine Mississippi River Arch (Algeo and Herrmann, 2018) controlled spatial variation of geochemical features of the Heebner Shale related to primary productivity, redox conditions, and water-column denitrification rates, as reflected in TOC, TN, δ13Corg, δ15N, and trace-metal records. Conodont biofacies models for Carboniferous cyclothems of the North American Midcontinent region have generally regarded glacioeustatically controlled water depths as the main influence on conodont distributions (Heckel, 1977; Boardman et al., 1995). Herrmann et al. (2019–this volume; Fig. 2) demonstrate that controls on conodont biofacies were, in fact, more complex based on a study of three geographically widely separated outcrop sections of the Middle Pennsylvanian Excello Shale (i.e., the core shale of the Lower Fort Scott cyclothem) representing a ~500-km transect across the Midcontinent Shelf. They conclude that, in addition to water depth, nutrient levels and watermass salinities were major influences on conodont biofacies. Although Gondolella was regarded in the past as a deep-water taxon similar to Idioprioniodus, it is found in the Excello Shale only within a narrow interval identified as the “regressive condensation surface” (RCS) of the Lower Fort Scott cyclothem. The RCS marks the onset of glacio-eustatic fall, pycnocline breakdown, and enhanced upward mixing of nutrient-rich deep waters, stimulating marine productivity (Algeo et al., 2004). A pronounced epibole of Gondolella in association with the RCS indicates that this taxon was an opportunistic form that bloomed under eutrophic conditions. With regard to salinity influences, Hindeodus is found exclusively above the RCS at more proximal (landward) sites, in which conodont δ18O compositions are ~1‰ heavier than in contemporaneous offshore sections. This distribution suggests that this taxon favored warm nearshore areas subject to limited freshwater runoff, with elevated salinities developing as a consequence of climate aridification accompanying the onset of renewed Gondwanan glaciation. Savoy and Mountjoy (1995–this volume; Fig. 2) document warmwater, shallow-marine carbonates within Upper Pennsylvanian-lowermost Permian strata from the Paleo-Tethys (Arabian margin of northeastern Gondwana), finding that their sedimentologic and paleobiologic properties differ greatly from coeval cool-water carbonates deposited at a similar paleolatitude (30° S) in the Panthalassic Ocean (Bolivia). They use a climate model (Community Climate System Model Version 3, CCSM3) to analyze controls on warm-water transport to higher southern paleolatitudes in northeastern Gondwana. These simulations suggest that, during the Late Paleozoic glacial maximum, spatial variations in sea-surface temperatures exerted a strong forcing on oceanic circulation patterns, thereby generating regional differences between marine ecosystems of the Paleo-Tethys and Panthalassic oceans with regard to faunal diversity and abundance. 4.2. Biotic effects The LPIA had a major impact on marine ecosystems, resulting in biodiversity losses, increased faunal provincialism, and restructuring of marine communities (McGhee et al., 2012; Wang et al., 2013). About 24% of marine invertebrate genera were lost in the Serpukhovian crisis, which is ranked seventh largest among Phanerozoic biodiversity crises (Sepkoski, 1996). A global analysis of Mississippian brachiopod 7
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occurrences, consisting of 2123 species belonging to 344 genera at 1156 localities, revealed an increase in faunal provincialism at the onset of the LPIA, probably due to closure of the Rheic Ocean and a steeper paleolatitude thermal gradient (Qiao and Shen, 2014). Marine ecosystems developed a new bioevolutionary mode, characterized by extremely low origination and extinction rates that persisted from the Late Mississippian to the Early Permian (Stanley and Powell, 2003; Clapham and James, 2008, 2012; Fig. 1). This mode was attributed to the effects of climatic cooling and accentuated seasonality, which are thought to have favored shallow-water species having large geographic ranges, broad ecological niches, and large and relatively stable populations with sluggish turnover rates. The termination of the LPIA around the Sakmarian-Artinskian boundary triggered major changes in marine ecosystems, including deep losses among rugose corals, brachiopods, fusulinaceans, and conodonts, in South China (Wang et al., 2013) and elsewhere. The Carboniferous Period was marked by major changes in terrestrial ecosystems as well. Coastal coal swamps inhabited by the fern-like Rhacophyton had appeared by the Late Devonian, but a shift to dominance of arborescent lycophytes occurred during the Early Carboniferous (Scheckler, 1986a). Cyclic changes in coastal vegetation occurred in response to glacio-eustatic cycles―in Nova Scotia, for example, transgressions led to retrograding coastal mires dominated by Lepidodendron and Lepidophloios, highstands produced prograding distributary wetlands dominated by lycopsid-pteridosperm-sphenopsid communities, and regressions yielded well-drained alluvial plains dominated by fire-prone cordaite and/or Sigillaria communities (Falcon-Lang, 2003, 2004). During much of the Pennsylvanian, tropical wetland floras were relatively static (Cleal et al., 2012), although by the Late Pennsylvanian epeirogenic uplift and climate change had triggered a long-term retreat of coastal swamps in North America and elsewhere (Cleal et al., 2011). By the Early Permian, a gradually warming climate and large-scale drying led to a seed plant-dominated world, beginning first at high latitudes and proceeding toward the tropics (DiMichele et al., 2001; Montañez et al., 2007). Concurrent changes occurred among terrestrial vertebrate (Ahlberg and Milner, 1994; Sahney et al., 2010) and insect faunas (Harrison et al., 2010; Engel et al., 2013). Of particular note is the “mid-Carboniferous diversification event”, in which animal colonization of coastal terrestrial environments was stimulated by large-scale eustatic transgressions and the widespread establishment of aquatic environments across the fresh-brackish spectrum (Falcon-Lang et al., 2015; Minter et al., 2017).
Bashforth, 2004; van Hoof et al., 2013) triggered major changes in geological processes related to silicate weathering and elemental cycling (Algeo et al., 2001; Le Hir et al., 2011). Late Devonian paleosols show increases in depth, horizonation, and compositional maturity, and some volumetrically important modern soil orders appeared at that time, e.g., alfisols and ultisols, characteristic of base-depleted forest ecosystems (Retallack, 1986, 1990, 1997). In the short term, rates of physical and chemical weathering increased due to the transition from largely unvegetated to vegetated upland areas, during which rhizosphere expansion accelerated the mechanical breakup of substrates (Stallard, 1985; Johnsson, 1993), as evidenced by seawater 87Sr/86Sr values (Fig. 1), sediment-mass anomalies, clay mineral assemblages, and paleomagnetic susceptibility data (Algeo et al., 1995; Crick et al., 1997, 2002; Hladil, 2002). Increased nutrient fluxes to epicratonic seas are suggested by evidence of eutrophication, including more frequent phytoplankton blooms and widespread microbial erosion of carbonate platforms (Caplan et al., 1996; Martin, 1996; Murphy et al., 2000a, 2000b; Joachimski et al., 2001, 2002; Peterhänsel and Pratt, 2001; Whalen et al., 2002; Goddéris and Joachimski, 2004). However, these weathering rate changes were transient owing to the long-term stabilizing effects of denser, thicker root mats as well as buffering of atmospheric pCO2 via the negative weathering-temperature climate feedback (Berner, 1992, 1994; Gwiazda and Broecker, 1994). Ultimately, land plants served to stabilize land surfaces including hillslopes and river levees, and to reduce physical weathering and sediment yields, contributing to a shift from braided channels to meandering channels in lowland river systems (Algeo et al., 2001; Davies and Gibling, 2011; Gibling and Davies, 2012; Gibling et al., 2014). The simultaneous development of widespread marine anoxia and biotic crises suggest a link to weathering processes and riverine nutrient fluxes, reflecting strong terrestrial-marine teleconnections (Algeo et al., 1995; Algeo and Scheckler, 1998). Recent work has resulted in new Devonian plant fossil finds in China (Guo et al., 2017–this volume; Xu et al., 2018–this volume; Xue et al., 2018; Fig. 2) and elsewhere. These studies are beginning to shed light on patterns of turnover among early plant clades, revealing a major transition during the Middle Givetian as dominant early clades (trimerophytes, zosterophyllophytes) went into decline and euphyllophytes, especially lignophytes, became more important. Palynomorph studies have assisted in refining regional miospore zonation schemes and dating Devonian-Carboniferous successions in remote areas such as South Tibet (Liu et al., 2017–this volume; Fig. 2). The nature of terrestrial habitats occupied by early land plants is often difficult to determine owing to a paucity of in situ fossil finds (most plant remains are transported). Wan et al. (2019–this volume; Fig. 2) show that the stable carbon isotopic composition of fossil land plants can be used to gain new insights into early terrestrial ecosystems. In this study, 309 measurements of δ13C in 12 plant genera (Archaeopteris, Drepanophycus, Genselia, Haskinsia, Leclercqia, Pertica, Psilophyton, Rhacophyton, Rhodeopteridium, Sawdonia, Tetraxylopteris, and Wattieza) yielded values ranging from −20.3‰ to −30.5‰ with a mean of −25.5‰, similar to the distribution for modern C3 land plants. This δ13C dataset revealed both intra- and intergeneric variation. Intrageneric variation is expressed as a small (mean 0.5‰) 13C-enrichment of leaves relative to stems that may reflect variation in compoundspecific compositions. Intergeneric variation is expressed as a much larger (to ~5‰) spread in the mean δ13C values of coeval plant genera that was probably controlled by taxon-specific habitat preferences and local environmental humidity (cf. Diefendorf et al., 2010). Among Early Devonian taxa, Sawdonia yielded the most 13C-depleted values (−27.1 ± 1.7‰; 1σ), reflecting lower water-use efficiency that was probably related to growth in wetter habitats, and Leclercqia, Haskinsia, and Psilophyton yielded the most 13C-enriched values (−23.0 ± 1.6‰, −22.3 ± 1.3‰, and − 24.8 ± 1.6‰, respectively), reflecting higher water-use efficiency probably related to growth in drier habitats. Following the onset of the LPIA, tropical forests of Laurasia
5. Controls on Late Paleozoic climate change and global events 5.1. Land plant evolution Terrestrial ecosystems underwent major changes during the Devonian. Early Devonian vascular plants were small (< 0.5 m height), had simple thread-like root systems lacking regrowth potential, and were confined to moist, low-lying habitats such as floodplains (Gensel and Andrews, 1984, 1987; Edwards and Berry, 1991; Raven and Edwards, 2001; Gensel et al., 2001). By the end of the Givetian (late Middle Devonian), plants had developed secondary supporting tissues and large, extractive root systems, allowing arborescent (tree-sized) growth (Chaloner and Sheerin, 1979; Mosbrugger, 1990; MeyerBerthaud et al., 2010; Xue et al., 2015; Xue et al., 2018; Fig. 1). The first forests appeared in the Givetian, signaling a major change in terrestrial ecosystems with potential long-term effects on global climate (Stein et al., 2012). By the Frasnian (early Late Devonian), aneurophyte and archaeopterid progymnosperm forests were common in floodplain habitats (Scheckler, 2001). The appearance of seed plants in the mid to late Famennian led to “greening” of the upland areas in continental interiors by the Early Carboniferous (Beck, 1981; Scheckler, 1986b; Marshall and Hemsley, 2003). The spread of vascular land plants to upland areas (Falcon-Lang and 8
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experienced repeated ecosystem restructuring involving climate-driven shifts between wetlands and seasonally dry biomes (cordaitaleans and conifers) at timescales of 105 to 107 yr (DiMichele et al., 2009; FalconLang, 2003; Falcon-Lang and DiMichele, 2010; Montañez, 2016). A permanent shift to seasonally dry woodland floras occurred during latest Carboniferous time (Sahney et al., 2010; Falcon-Lang et al., 2011, 2017), coincident with widespread low-latitude aridification and a rapid decline in the seawater 87Sr/86Sr record, indicating complex linkages between floral evolution, tectonism, and weathering. The end of the LPIA brought further major changes in the types and geographic distribution of terrestrial floras (Poulsen et al., 2007). The Permo-Carboniferous was the main period of coal formation in Earth history, in large part due to the expansion of tropical wetland forests and a vast supply of biodegradably resistant organic matter, which led to low atmospheric pCO2 and high pO2 (Berner, 2006; Krause et al., 2018; Fig. 1). The development of lignin biosynthesis by plants during the Middle Devonian (Weng and Chapple, 2010) may have preceded the evolution of lignin-degrading fungi in the early Permian by ∼100 Myr (Floudas et al., 2012), although this idea has been challenged (Nelsen et al., 2016). The resulting increase in burial of terrestrial organic carbon (Fig. 1) is likely to have been a contributing factor to low atmospheric CO2 levels during the Carboniferous and, thus, a primary trigger of the LPIA (Robinson, 1990; Montañez, 2016). Indeed, it has been argued that massive coal burial nearly triggered a Snowball Earth event during the LPIA (Feulner, 2017). Organic geochemical studies of late Paleozoic plant fossils are rare to date but can provide insights regarding their biogeochemistry and ecology. Tewari et al. (2018–this volume; Fig. 2) examine the molecular signatures, including lipid-derived aliphatic and aromatic biomarkers and tetramethylammonium hydroxide thermochemolysis, of solventextractable and non-extractable organic matter of the taxon Glossopteris collected from Permian successions of eastern India. The Glossopteris plant was able to biosynthesize abietic acid and related plant terpenoids as indicated by the presence of aromatic diterpanes in the leaf extract. Thermochemolysis data reveal preserved lignin structures in the fossilized Glossopteris leaf and stem samples, which may have contributed to the proliferation and dominance of Glossopteris in Gondwana during the Permian.
1989; Savoy and Mountjoy, 1995; Day et al., 1996; Witzke and Bunker, 1996; Sandberg et al., 2002; Algeo et al., 2007), including in South China (Xu et al., 2018–this volume; Fig. 2). This long-term rise led to deposition of the first-order Kaskaskia Sequence in North America (Sloss, 1963) and was probably driven by accelerated mid-ocean ridge spreading at a global scale (Gaffin, 1987). The seawater 87Sr/86Sr record has long been used to constrain the timing and magnitude of tectonic events, continental weathering, and paleoclimate changes (Fig. 1). Recently, Chen et al. (2018a) presented a ~38-Myr-long, high-resolution (100-kyr) 87Sr/86Sr record for conodont apatite recovered from the Naqing section, South China. From the middle Visean to early Bashkirian, 87Sr/86Sr values increased rapidly as a result of enhanced continental weathering triggered by the Variscan Orogeny, influenced by expansion of rainforests in low-latitude regions (Chen et al., 2018a). Throughout much of the Pennsylvanian, sustained high weathering rates due to westward propagation of the Variscan Orogeny from Europe to North America led to a ~15 Myr plateau of high 87Sr/86Sr values and low atmospheric CO2 levels (Fig. 1). New generations of plate-tectonic reconstructions (e.g., Domeier and Torsvik, 2014; Matthews et al., 2016; Young et al., 2018) are providing more detailed information about plate vectors and rates of motion, oceanic plate locations, and spreading center distributions. A major rise in eustatic elevations during the Middle to Late Devonian (Johnson et al., 1985; Sandberg et al., 2002) was associated with a substantial increase in rates of plate motion (Domeier and Torsvik, 2014; Matthews et al., 2016; Young et al., 2018), implying larger MOR volumes (Gaffin, 1987). Improved plate reconstructions provide the basis for testing models of supercontinent formation and dispersal. The formation of Pangea during the Carboniferous and Permian followed the ‘introversion model’ of Murphy and Nance (2013), which calls for formation of subduction zones on previously passive margins followed by slab avalanche events. New paleobiogeographic data are also helping to refine these models. For example, recent fossil plant finds from the Yellow Sea shelf serve to demarcate the boundary between the Yangtze (South China) Plate and the Sino-Korean (North China) Plate (Guo et al., 2017–this volume; Fig. 2).
5.2. Tectonic activity
Large igneous provinces can have significant global climatic effects and trigger shifts between icehouse, greenhouse, and hothouse states (Ernst and Youbi, 2017). Climatic warming can result from large-scale outgassing of greenhouse gases such as CO2 and SO2, as is well-documented for the Permian-Triassic boundary Siberian Traps (Sun et al., 2012) and Triassic-Jurassic boundary CAMP eruptions (Ruhl et al., 2011). On the other hand, subsequent cooling or even global glaciations can occur due to rapid atmospheric CO2 drawdown via weathering of LIP-related basalts (Swanson-Hysell and Macdonald, 2017). Contemporaneous LIP volcanism may have played an important role in Late Devonian climatic perturbations. This was a time of voluminous flood basalt magmatism on the Siberian Platform, where the eruption of the Viluy Traps began around the Frasnian-Famennian boundary and continued into the Early Carboniferous (Courtillot et al., 2010). On the Russian Platform, the Pripyat–Dniepr–Donets (PDD) LIP was emplaced between the late Frasnian and the late Famennian, with peak magmatic activity during the late Famennian (Kusznir et al., 1996; Wilson and Lyashkevich, 1996). The PDD covered a minimum area of ~1 × 106 km2 with a total magma volume > 1.5 × 106 km3 (Kravchinsky, 2012). More importantly, the PDD was situated in a periequatorial location (Abrajevitch et al., 2007; Scotese, 2013), which may have led to rapid weathering of the flood basalts under humid tropical conditions (Goddéris and Joachimski, 2004) and, thus, strong drawdown of atmospheric CO2 levels. In recent years, Hg has emerged as a proxy for volcanic inputs to stratigraphic successions (Sanei et al., 2012), providing evidence of the relationship of LIPs to key geologic event intervals (Grasby et al., 2013;
5.3. Large igneous province (LIP) magmatism
All of the paleofloral changes discussed above occurred against a backdrop of widespread tectonism. In North America, the Acadian Orogeny was caused by a collision between several Acadian microcontinents and the eastern margin of the Laurentian Craton (Murphy and Keppie, 2005; Fig. 1). The Acadian Orogen was active from the latest Early Devonian (Emsian) through the end of the Devonian, marked by a progressive shift in peak intensity from the Canadian Maritime provinces southward through the New England region to the southern Appalachians (Ettensohn, 1992; Faill, 1997; Bradley and Tucker, 2002; Dunning et al., 2002; Thompson and Hermes, 2003). Roughly contemporaneous orogenies were underway in northern Laurentia (the Ellesmerian; Stoakes, 1980; Embry, 1991; Stevenson et al., 2000) and in Baltica (the Eovariscan; Tait et al., 1997; Echarfaoui et al., 2002). During the Late Devonian, the western margin of North America became tectonically active after prolonged passive margin subsidence (Bond and Kominz, 1991). Tectonic and eustatic influences on Devonian sedimentation have been extensively documented in regional (chemo-)stratigraphic studies of the Appalachian Basin (Murphy et al., 2000a, 2000b; Werne et al., 2002; Sageman et al., 2003), the Midcontinent (Day et al., 1996; Witzke and Bunker, 1996), the Western Canadian Basin (Whalen et al., 2000), and other locales globally. Eustatic elevations were generally rising through the Middle and Late Devonian, resulting in widespread inundation of cratons and relatively deep-water conditions in many cratonic-interior basins (Johnson et al., 1985; Johnson and Sandberg, 9
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Sial et al., 2016; Percival et al., 2017; Shen et al., 2019a). However, recent work has also shown that Hg can be concentrated through a variety of mechanisms, and that spikes in Hg levels are not always due to volcanic inputs, mandating caution in application of this proxy (Shen et al., 2019b; Them et al., 2019). Kalvoda et al. (2019–this volume; Fig. 2) investigate the mercury (Hg) chemostratigraphy of two Devonian-Carboniferous boundary sections (Lesní lom, Czech Republic, and Duli, South China). They observe Hg peaks at the DCB in both sections, although covariant relationships indicate an association of Hg with pyrite at Lesní lom and with clay minerals at Duli. The latter pattern may indicate transport of Hg from land areas to the marine realm during the Devonian-Carboniferous transition. The authors infer a possible relationship to the multi-phase Viluy Traps LIP, although its younger phase (364.4 ± 1.7 Ma) slightly preceded the documented Hg anomalies. Hg enrichments at the DCB were reported in two other recent studies, Racki et al. (2018) and Paschall et al. (2019), demonstrating that the end-Devonian Hg anomalies–regardless of their cause–are a global signal. There is evidence for LIP magmatism in the Early Permian, suggesting that the associated CO2 emissions may have been a contributing factor to the termination of the LPIA. The Skagerrak-Centered LIP (SCLIP) in northwestern Europe covers a large area (~0.5 × 106 km2) and records a flare-up of igneous activity at ~297 Ma (Torsvik et al., 2008). In NW China, spatially and temporally associated mafic–ultramafic and syenitic intrusions and volcanic rocks form the ~275 Ma Tarim LIP, covering an area > 250,000 km2 in the western and central parts of the Tarim Basin (Zhou et al., 2009; Li et al., 2011). These episodes of flood basalt activity may have played an important role in Early to Middle Permian global climate change.
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6. Summary The Early Devonian to Middle Permian (~410 to 260 Ma) was an interval of major tectonic, climatic, oceanic, and biotic changes. The collision of Laurasia and Gondwana led to global-scale tectonic activity and the formation of the supercontinent Pangea, and the spread of vascular land plants resulted in a huge increase in organic carbon burial and atmospheric CO2 drawdown. These events were the main drivers of the global climatic transition from a Middle Paleozoic greenhouse to a Late Paleozoic icehouse. The Late Paleozoic Ice Age (LPIA) commenced with short-lived glaciations at the Frasnian-Famennian boundary (FFB) and Devonian-Carboniferous boundary (DCB) that coincided with two first-order mass extinctions, although many minor biocrises also occurred during this interval. More sustained growth of continental icesheets began in the mid-Mississippian (Visean), and a major intensification of LPIA glaciation occurred around the MississippianPennsylvanian boundary, although it remains in debate whether continental ice volume during the LPIA reached a peak in the Early or Middle Pennsylvanian or the Early Permian. The developments of the Earth system during the Early Devonian to Middle Permian interval record the dynamic interplay of its geo-, hydro-, bio-, and atmospheric systems. The 23 contributions to this special issue of Palaeogeography, Palaeoclimatology, Palaeoecology advance our knowledge of various scientific issues related to these developments. Acknowledgments We thank Howard Falcon-Lang and Isabel Montanez for critical readings of the draft manuscript. References Abrajevitch, A., Van der Voo, R., Levashova, N.M., Bazhenov, M.L., 2007. Paleomagnetic constraints on the paleogeography and oroclinal bending of the Devonian volcanic arc in Kazakhstan. Tectonophysics 441 (1–4), 67–84. Ahlberg, P.E., Milner, A.R., 1994. The origin and early diversification of tetrapods. Nature
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