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This is an Open Access article distributed under the terms of the Creative Commons Attribution 4.0 License (http://creativecommons.org/licenses/by/4.0/).

Icehouses are the less common climate state on Earth, and thus it is notable that the longest-lived (c. 370 to 260 Ma) and possibly most extensive and intense of icehouse periods spanned the Carboniferous Period. Mid- to high-latitude glaciogenic deposits reveal a dynamic glaciation–deglaciation history with ice waxing and waning from multiple ice centres and possible transcontinental ice sheets during the apex of glaciation. New high-precision U–Pb ages confirm a hypothesized west-to-east progression of glaciation through the icehouse, but reveal that its demise occurred as a series of synchronous and widespread deglaciations. The dynamic glaciation history, along with repeated perturbations to Earth System components, are archived in the low-latitude stratigraphic record, revealing similarities to the Cenozoic icehouse. Further assessing the phasing between climate, oceanographic, and biotic changes during the icehouse requires additional chronostratigraphic constraints. Astrochronology permits the deciphering of time, at high resolution, in the late Paleozoic record as has been demonstrated in deep- and quiet-water deposits. Rigorous testing for astronomical forcing in low-latitude cyclothemic successions, which have a direct link to higher-latitude glaciogenic records through inferred glacioeustasy, however, will require a comprehensive approach that integrates new techniques with further optimization and additional independent age constraints given challenges associated with shallow-marine to terrestrial records.

In the six hundred million years since the radiation of multi-cellular life, our planet has experienced several major climate swings between the more typical warm and ice-free or ephemeral ice conditions of a greenhouse climate state and those of a colder glaciated icehouse state. Of these climate turnovers, the transition into the longest-lived and possibly most extensive and intense of the icehouse periods, known as the Late Paleozoic Ice Age (LPIA), occurred primarily during the Carboniferous Period. Substantial evidence indicates that the turnover from the mid-Paleozoic greenhouse into Earth's penultimate icehouse began with relatively short-lived (<1 to 2 myr) and possibly intense glaciation(s) in the late Famennian through Devonian–Carboniferous boundary interval (Streel et al. 2000, 2012; Brezinski et al. 2008, 2010; Caputo et al. 2008; Isaacson et al. 2008; Lakin et al. 2016; Caputo and Santos 2020; Ettensohn et al. 2020).

Glaciogenic deposits are reported for every stage of the Carboniferous and much of the Permian (up through the Wuchiapingian Stage), suggesting the possibility of continued glaciation from the latest Devonian through the Permian Period (Fig. 1a) (Soreghan et al. 2019). Although evidence exists for glaciation in the mid-Early Mississippian (mid-Tournasian), the onset of near sustained glacial conditions and increasing long-term accumulation of continental ice appears to have begun in the late Middle Mississipian (mid- to late Visean) to earliest Serpukhovian (Crowley and Baum 1991; Isbell et al. 2003, 2012; Fielding et al. 2008a, b; Montañez and Poulsen 2013; Griffis et al. 2019, 2021; Soreghan et al. 2019), reaching the peak of glaciation at the close of the Carboniferous and earliest Permian (Fig. 1b) (Fielding et al. 2008b; Isbell et al. 2012, 2013; Montañez and Poulsen 2013; Griffis et al. 2018, 2019, 2021; Soreghan et al. 2019). The onset of the main glacial interval of the LPIA has been previously associated with a biodiversity crisis that saw c. 27% loss of marine genera (Sepkoski 1998; McGhee et al. 2012) followed by low biodiversity and sluggish evolutionary turnover in the oceans through the subsequent 40 million years of the main phase of the LPIA (Stanley and Powell 2003). A broader faunal assessment of marine life at the time, however, reveals a persistent rapid rate of species increase and a near tripling of species from the late Visean through late Asselian, coincident with the main phase of the LPIA (Fan et al. 2020; Shi et al. 2021). This Carboniferous–earliest Permian Biodiversification Event (CPBE) was likely an adaptive radiation made possible by the rapid environmental (i.e. sea-level, ocean circulation and degree of stratification) and climate changes that occurred with the transition into the main glacial phase of the LPIA (Shi et al. 2021).

Fig. 1.

Late Devonian through Permian trends in occurrence and distribution of glaciogenic records, inferred magnitudes of glacioeustasy, and palaeo-CO2 plotted on the Geologic Time Scale 2020 (Gradstein et al. 2020). (a) Number of occurrences (extent), by stage, of glaciogenic deposits in mid- to high-latitude Gondwanan basins compiled by Soreghan et al. (2019). (b) Geographical distribution of glaciogenic deposits and geomorphic features through time; see section ‘Synopsis of ice ages of the late Paleozoic’ for references. Hachured pattern in the ‘Arabian Peninsula and Madagascar’ record represents multiple hiatuses, of unknown age, that are interpreted as glacial advances/cycles. Blue horizontal lines show the temporal position of high-precision CA (ID)-TIMS U–Pb dates for glaciogenic deposits (Gulbranson et al. 2010; Metcalfe et al. 2015; Griffis et al. 2018, 2019, 2021; Phillips et al. 2018; Backhouse and Mory 2020). SIMS (secondary ion mass spectrometry), SHRIMP (sensitive high-resolution ion microprobe), or LA MC-ICPMS zircon ages are not shown given large uncertainties associated with ages. (c) Magnitudes of high-frequency glacioeustasy inferred from different sedimentary archives; modified after Rygel et al. (2008); (d) Palaeo-CO2 estimates and LOESS trendline (30%) with 12 and 86% confidence intervals, courtesy of J. Chen. Palaeosol carbonate- and stomatal-based CO2 estimates of Montañez et al. (2016) and Richey et al. (2020) are shown as solid symbols and pass the screening criteria to be acceptable as initially published (https://palaeo-co2.org). CO2 estimates compiled by Foster et al. (2017) and shown as open symbols are quarantined until further modernization. Vertical grey lines are uncertainties on age estimates of individual CO2 values.

Fig. 1.

Late Devonian through Permian trends in occurrence and distribution of glaciogenic records, inferred magnitudes of glacioeustasy, and palaeo-CO2 plotted on the Geologic Time Scale 2020 (Gradstein et al. 2020). (a) Number of occurrences (extent), by stage, of glaciogenic deposits in mid- to high-latitude Gondwanan basins compiled by Soreghan et al. (2019). (b) Geographical distribution of glaciogenic deposits and geomorphic features through time; see section ‘Synopsis of ice ages of the late Paleozoic’ for references. Hachured pattern in the ‘Arabian Peninsula and Madagascar’ record represents multiple hiatuses, of unknown age, that are interpreted as glacial advances/cycles. Blue horizontal lines show the temporal position of high-precision CA (ID)-TIMS U–Pb dates for glaciogenic deposits (Gulbranson et al. 2010; Metcalfe et al. 2015; Griffis et al. 2018, 2019, 2021; Phillips et al. 2018; Backhouse and Mory 2020). SIMS (secondary ion mass spectrometry), SHRIMP (sensitive high-resolution ion microprobe), or LA MC-ICPMS zircon ages are not shown given large uncertainties associated with ages. (c) Magnitudes of high-frequency glacioeustasy inferred from different sedimentary archives; modified after Rygel et al. (2008); (d) Palaeo-CO2 estimates and LOESS trendline (30%) with 12 and 86% confidence intervals, courtesy of J. Chen. Palaeosol carbonate- and stomatal-based CO2 estimates of Montañez et al. (2016) and Richey et al. (2020) are shown as solid symbols and pass the screening criteria to be acceptable as initially published (https://palaeo-co2.org). CO2 estimates compiled by Foster et al. (2017) and shown as open symbols are quarantined until further modernization. Vertical grey lines are uncertainties on age estimates of individual CO2 values.

The demise of Earth's penultimate icehouse began in the mid- to late Sakmarian Age of the Permian Period and was protracted, extending from c. 290 to c. 255 Ma (Fig. 1b), with post-Sakmarian ice becoming increasingly more restricted after 290 Ma and, in many places, limited to alpine glaciation (Fielding et al. 2008a, b; Isbell et al. 2013; Montañez and Poulsen 2013; Garbelli et al. 2019; Griffis et al. 2019, 2021; Soreghan et al. 2019).

This glaciation history is clearly imprinted on the Carboniferous and Permian stratigraphic record. Glacial deposits and associated geomorphic structures have been documented throughout mid- to high-palaeo-latitude depositional basins of South America, eastern and western Australia and Tasmania, southern and central Africa, southwestern China and Tibet (referred to as the Cimmerian blocks on Fig. 1), India, and the Arabian Peninsula (Fig. 2). In contrast, in the low-latitude tropical and sub-tropical successions of Euramerica (Fig. 2), the LPIA glaciation history is archived in the stratigraphic architectures and stratal stacking patterns of contemporaneous successions that are interpreted to record sea-level fluctuations and climate changes driven by the waxing and waning of Gondwanan glaciers.

Fig. 2.

Palaeogeographical distribution of glaciated basins in Gondwana (light orange shading) and select unglaciated, low-latitude depositional sites (dark orange shading) during the late Devonian, Carboniferous, and Permian periods. Base map and locations of palaeo-ice centres (blue) and ice-flow lines modified after Craddock et al. (2019). Position of Gondwanan glaciated basins from Gulbranson et al. (2010), Wopfner and Jin (2009), Isbell et al. (2012), Bussert (2014), Fan et al. (2015), Griffis et al. (2019), Caputo and Santos (2020), Buso et al. (2020) and Zurli et al. (2021). Dashed lines for Cimmerian blocks (N. Qiangtang–Qamdom, S. Qiangtang–Baoshan, Tethys Himalaya and Lhasa blocks) indicate uncertainty in distribution of glaciomarine diamictites (Fan et al. 2015). Low-latitude basins from Heckel (2013) and Blakey (2008); Midcontinent includes N. Shelf, Arkoma, Anadarko, and Midland basins and eastern shelf of the Midland Basin.

Fig. 2.

Palaeogeographical distribution of glaciated basins in Gondwana (light orange shading) and select unglaciated, low-latitude depositional sites (dark orange shading) during the late Devonian, Carboniferous, and Permian periods. Base map and locations of palaeo-ice centres (blue) and ice-flow lines modified after Craddock et al. (2019). Position of Gondwanan glaciated basins from Gulbranson et al. (2010), Wopfner and Jin (2009), Isbell et al. (2012), Bussert (2014), Fan et al. (2015), Griffis et al. (2019), Caputo and Santos (2020), Buso et al. (2020) and Zurli et al. (2021). Dashed lines for Cimmerian blocks (N. Qiangtang–Qamdom, S. Qiangtang–Baoshan, Tethys Himalaya and Lhasa blocks) indicate uncertainty in distribution of glaciomarine diamictites (Fan et al. 2015). Low-latitude basins from Heckel (2013) and Blakey (2008); Midcontinent includes N. Shelf, Arkoma, Anadarko, and Midland basins and eastern shelf of the Midland Basin.

The following section presents first a mid- to high-latitude perspective of the current understanding of the geographical distribution and timing of continental ice and ensuing deglaciations in the late Paleozoic. Second, it provides a context for discussion of how the glaciation history is archived in the low-latitude stratigraphic record. Third, it provides the foundation for discussion of recent and future efforts to develop a robust astronomical timescale to calibrate stratigraphic, sedimentological, palaeobiological, and geochemical proxy records of Earth's penultimate icehouse and to refine the temporal resolution, accuracy, and precision of the Late Devonian to middle Permian timescales.

The earliest evidence for global cooling and glaciation is in the Fammenian Age of the Late Devonian Period and exists as late Fammenian glaciogenic deposits in the Peru–Bolivian basins, Solimões, Amazonas, Parnaíba and possibly Paraná basins of mid- to high-palaeo-latitude South America (Fig. 2) (Caputo and Crowell 1985; Díaz-Martinez et al. 1999; Isaacson et al. 1999, 2008; Streel et al. 2012; Lakin et al. 2016; Caputo and Santos 2020), suggesting widespread glaciation with multiple ice sheets and/or ice caps (Caputo et al. 2008). Latest Devonian continental ice may have also existed in central and west Africa (Streel and Theron 1999; Streel et al. 2000; Almond et al. 2002; Isaacson et al. 2008; Wicander et al. 2011), the Cimmerian blocks, and in the subtropical palaeo-latitudes of Euramerica (Brezinski et al. 2008, 2010; Ettensohn et al. 2020). Independent evidence for cooling in the Late Devonian comes from an increase in oxygen isotopic compositions (δ18O) of brachiopods and conodonts (Grossman and Joachimski 2020). Substantial shallowing of sea-level (100 to 140 m), recognized globally and interpreted as a eustatic fall, also began in the Late Devonian and culminated immediately prior to the close of the Devonian Period (Isaacson et al. 1999; Kaiser et al. 2011), suggesting maximum ice accumulation for the Devonian at this time (cf. Soreghan et al. 2019). Shorter-term, large-magnitude fluctuations were superimposed on the longer-term sea-level trend and are interpreted to be glacioeustatic (Day et al. 2013; Myrow et al. 2014; Ettensohn et al. 2020). The Hangenberg (Crisis) Event (≤50 kyr), which occurred at the Devonian–Carboniferous boundary, is widely identified by widespread organic-rich black shales (Kaiser et al. 2008) and characterized by a positive carbon isotope (δ13C) excursion in carbonates and organic matter from many stratigraphic successions (Buggisch and Joachimski 2006; Kaiser et al. 2006, 2008; Day et al. 2013). The Hangenberg Event occurred during an interglacial transgression expressed globally with warm-water faunal migrations into the higher latitudes (Kaiser et al. 2006, 2011). This event was the fourth largest biodiversity crisis of the Phanerozoic (McGhee et al. 2012). An astrochronology developed for a low-latitude Devonian–Carboniferous boundary succession (Poland) constrains the duration of the Hangenberg transgression and anoxic event (400 kyr) and supports a hypothesized extremely high eccentricity driver for collapse of small glaciers in western Gondwana (De Vleeschouwer et al. 2013).

Although evidence of glaciogenic deposits exists throughout much of the Mississippian (Fig. 1a) (Soreghan et al. 2019), the distribution of glacial deposits through this interval suggests an initial episode of geographically restricted glaciation in the Early Mississippian (late middle to early late Tournasian) prior to the onset of sustained glaciation that began in the Middle Mississippian (Visean). Reduced ice volume, sea-level rise, and possibly reduced magnitudes of eustatic fluctuations occurred between these two periods (e.g. Rygel et al. 2008; Lakin et al. 2016). Glaciomarine deposits (Fig. 2b) occur in the Parnaiba and Solimões basins in northern Brazil and probable glacial diamictites occur in Bolivian, Argentinian and southern Brazilian basins (Caputo et al. 2008; Rosa et al. 2019; Caputo and Santos 2020) as well as possibly regions of South Africa and the Falkland Islands (Caputo et al. 2008; Lakin et al. 2016). This glacial interval is associated with a major fall in sea-level archived in palaeotropical stratigraphic architecture (Ross and Ross 1988; Kammer and Matchen 2008), and a long-term positive shift in biogenic δ18O values indicating cooling beginning in the Early Mississippian (Fig. 3c) (Bruckschen et al. 1999; Buggisch et al. 2008; Grossman and Joachimski 2020). As with the latest Devonian glaciation, the mid- to late Tournasian interval culminated with a variable magnitude (2 to 7‰), positive carbon isotope excursion (CIE) recorded in stratigraphic successions from Euramerica and South China and named the Kinderhookian–Osagean Boundary Excursion (KOBE; Saltzman 2002) or Tournasian Isotope Carbon Excursion (TICE; L. Yao et al. 2015) (Fig. 3d). The CIE lasted between 6 and 10 myr (Saltzman et al. 2000; Katz et al. 2007; Qie et al. 2011; Cole et al. 2015) and was accompanied by a c. 6‰ positive excursion in organic matter δ15N values (L. Yao et al. 2015), indicating a major change in ocean-circulation during the Tournasian cooling. Multiple shorter-term δ13Ccarb excursions of 105 to 106 year-scale were superimposed on the longer-term trend with some correlating to stratigraphic genetic sequences and inferred substantial sea-level fluctuations (Katz et al. 2007).

Fig. 3.

Late Devonian through Permian trends in occurrence of glaciogenic records, palaeo-CO2, and geochemical proxy records of environmental change plotted on the Geologic Time Scale 2020 (Gradstein et al. 2020). (a) Number of occurrences (extent), by stage, of glaciogenic deposits as in Figure 1a. (b) Palaeo-CO2 estimates and LOESS trendline as in Figure 1d. (c) Trends in brachiopod (black lines) and conodont (blue and grey lines) δ18O values from Grossman and Joachimski (2020). The brachiopod (black) and conodont (grey) δ18O records were compiled using primarily records from palaeo-epicontinental seas; the blue conodont δ18O record is from an open-water slope succession in South China (B. Chen et al. 2016). Decoupling of trends and increased variability in conodont δ18O from epicontinental sea successions have been interpreted as recording variability in surface water salinity and seawater δ18O (Joachimski and Lambert 2015; Montañez et al. 2018). (d) Trend in sedimentary carbonate δ13C compiled by and courtesy of B. Cramer (see Cramer and Jarvis (2020) for details and data sources. KOBE is Kinderhookian–Osagean Boundary Excursion; TICE is the Tournasian Isotope Carbon Excursion. (e) LOESS trend in Carboniferous and Permian seawater 87Sr/86Sr shown with 2.5 and 97.5% confidence intervals; modified after from McArthur et al. (2020); for the Carboniferous: Chen et al. (2021); for the Permian: Wang et al. (2021). Frasnian 87Sr/86Sr trend from McArthur et al. (2020); Famennian not shown given sizeable disparity between datasets.

Fig. 3.

Late Devonian through Permian trends in occurrence of glaciogenic records, palaeo-CO2, and geochemical proxy records of environmental change plotted on the Geologic Time Scale 2020 (Gradstein et al. 2020). (a) Number of occurrences (extent), by stage, of glaciogenic deposits as in Figure 1a. (b) Palaeo-CO2 estimates and LOESS trendline as in Figure 1d. (c) Trends in brachiopod (black lines) and conodont (blue and grey lines) δ18O values from Grossman and Joachimski (2020). The brachiopod (black) and conodont (grey) δ18O records were compiled using primarily records from palaeo-epicontinental seas; the blue conodont δ18O record is from an open-water slope succession in South China (B. Chen et al. 2016). Decoupling of trends and increased variability in conodont δ18O from epicontinental sea successions have been interpreted as recording variability in surface water salinity and seawater δ18O (Joachimski and Lambert 2015; Montañez et al. 2018). (d) Trend in sedimentary carbonate δ13C compiled by and courtesy of B. Cramer (see Cramer and Jarvis (2020) for details and data sources. KOBE is Kinderhookian–Osagean Boundary Excursion; TICE is the Tournasian Isotope Carbon Excursion. (e) LOESS trend in Carboniferous and Permian seawater 87Sr/86Sr shown with 2.5 and 97.5% confidence intervals; modified after from McArthur et al. (2020); for the Carboniferous: Chen et al. (2021); for the Permian: Wang et al. (2021). Frasnian 87Sr/86Sr trend from McArthur et al. (2020); Famennian not shown given sizeable disparity between datasets.

A second and geographically more extensive phase of continental ice accumulation began in the later Visean Age (c. 335 to 330 Ma) of the Middle Mississippian Epoch as indicated by glaciogenic deposits in several peripolar circle regions (Fig. 2; some South American basins, eastern Australia, Falkland Islands, and Malaysia) and evidence of glacially eroded surfaces (González 1990; Caputo et al. 2008; Fielding et al. 2008a, b; Gulbranson et al. 2010; Meor et al. 2014; Alonsa-Muruaga et al. 2018; Rosa et al. 2019). In the lower palaeo-latitudes, this increase in ice accumulation is archived as the first appearance of regional-scale erosion surfaces and karstification in carbonates, with incised valleys stacked stratigraphically, suggesting sea-level fluctuations of a few to several tens of metres (Smith and Read 2000; Wright and Vanstone 2001; Al-Tawil and Read 2003; Bishop et al. 2009; Pointon et al. 2014; Fielding and Frank 2015; Ahern and Fielding 2019, 2021; Chen et al. 2019). This was also the timing of a regional facies change (see section ‘Greenhouse to icehouse stratal architecture and cyclicity’) and a c. 3 to 5°C cooling inferred from a shift in biogenic apatite δ18O values from the UK (Barham et al. 2012) and a more open-water slope setting in South China (B. Chen et al. 2016).

Notably, magnitudes of eustatic fluctuations estimated from stratal patterns and from depths of Middle to Late Mississippian erosion surfaces (i.e. incised valleys) increase rapidly in the late Visean, suggesting a ramping up of ice build-up into the Serpukhovian c. 5 to 10 myr after the onset of repeated occurrences of substantial surface erosion (Fielding and Frank 2015; J. Chen et al. 2016; Ahern and Fielding 2019). Additionally, eustatic magnitudes (Davies 2008; Fielding and Frank 2015) and onlap–offlap reconstructions (Eros et al. 2012a) inferred from Visean and Serpukhovian Euramerican successions indicate short periods of alternating offlap–onlap and higher–lower magnitude fluctuations interpreted as precursor cycles of glaciation and glacioeustasy. Collectively, this evidence suggests a stepwise onset of the main phase of the LPIA (Eros et al. 2012a) analogous to the cycles of expansion and retraction of the Antarctic Ice Sheet several million years prior to the Eocene–Oligocene boundary (34 Ma), heralding the onset of the Cenozoic icehouse (DeConto et al. 2008; Miller et al. 2020).

Although glacial deposits occur in nearly each stage of the Carboniferous and Permian geological record, multiple sedimentological, stratigraphic and geochemical lines of evidence indicate a dynamic glaciation history during the main phase of the LPIA (c. 335 to 290 Ma) that was characterized by a series of discrete several million-year-long glacials (Visser 1997; Isbell et al. 2003, 2012; Fielding et al. 2008a, b; Gulbranson et al. 2010; Montañez and Poulsen 2013; Buso et al. 2020) with ice expanding out of minimally a dozen to up to 30 ice centres (Isbell et al. 2012; Craddock et al. 2019). Notably, periods between discrete intervals of expanded continental ice appear to have had substantially reduced ice volume, in particular for three intervals of the Pennsylvanian Epoch: (1) within the Bashkirian Age (although the timing is less constrained than the other two events), (2) in the late Moscovian Age, and (3) the latest Kasimovian through early–mid-Gzhelian interval contemporaneous with the timing of the Alykaevo Climatic Optimum (Cleal and Thomas 2005; Richey et al. 2020). These inferred warmings and periods of ice retraction led to (1) permanent loss of Carboniferous glacial deposits in west-central Argentina by the latest Bashkirian Age (c. 315 Ma) (Gulbranson et al. 2010; Henry et al. 2010), (2) deglaciation sequences in southern Africa (Stollhofen et al. 2008), and (3) the deposition of widespread marine shales (e.g. the mid- to late Bashkirian Roncador Shale and the Kasimovian Lontras Shale) in the Paraná Basin, Brazil that are interpreted to record transgressions at the culmination of million-year-scale glacial cycles (Buso et al. 2020). These three inferred warmings are further suggested by contemporaneous increases in relative sea-level in Euramerica (Ross and Ross 1987; Eros et al. 2012b) and South China (Huang et al. 2017), and dampened magnitudes of glacioeustasy inferred from time-equivalent low-latitude successions (Rygel et al. 2008; Fischbein et al. 2009; Waters and Condon 2012). Biogenic δ18O values (Fig. 3c) decrease during the latter two intervals, indicating warmer seawater temperatures. For the Moscovian and late Kasimovian events, penecontemporaneous decreases in δ13C values (Fig. 3d) are associated with rises in atmospheric pCO2 (Fig. 3b) (Richey et al. 2020). Whether these inferred warming events and deglaciations were synchronous requires further study and higher temporal resolution and chronostratigraphic control for many Gondwanan and lower-latitude basins.

Between these inferred warmer periods of ice retraction or loss (whether synchronous or not), glacial deposits and geomorphic features record multiple intervals of ice expansion, referred to as ‘glacials’. Sedimentary successions throughout southern Gondwana indicate a geographical expansion of continental ice from the Serpukhovian onward (Fig. 1a, b) with a marked increase in inferred magnitudes of glacioeustasy across the mid-Carboniferous boundary (323.2 ± 0.4 Ma) (Fig. 1c). Mid-Carboniferous glaciation was geographically widespread with palaeo-ice centres in the mid- to high-palaeo-latitudes of eastern Australia (Fielding et al. 2008a, b), northwestern Argentina (Henry et al. 2008; Gulbranson et al. 2010; Buso et al. 2020), southern Brazil (Rocha-Campos et al. 2008; Holz et al. 2010), and possibly in central and southern Africa (Isbell et al. 2008; Stollhofen et al. 2008; Linol et al. 2015). In lower-latitude successions, the ‘mid-Carboniferous’ glaciation is recorded by a basin-wide shift in sedimentation and by the development of widespread erosional and karst surfaces, and, in places, a multimillion-year unconformity (Blake and Beuthin 2008; Rygel et al. 2008; Bishop et al. 2009; Martin et al. 2012; Eros et al. 2012b; J. Chen et al. 2016). Notable cooling and increased ice accumulation at this time is further indicated by a large-scale increase in biogenic δ18O values (Fig. 3c) (Buggisch et al. 2008; B. Chen et al. 2016; Montañez et al. 2018; Grossman and Joachimski 2020). A contemporaneous increase in brachiopod δ13C values (Fig. 3d) is interpreted as a major increase in carbon sequestration and a consequent decrease in palaeo-CO2. Independent proxies of pCO2 indicate that this was also the beginning of a c. 35 myr interval of overall lowest pCO2 for the late Paleozoic (Fig. 3b). Mid-Carboniferous glaciation was associated with substantial biotic turnover and the beginning of strong endemism in marine fauna (McGhee et al. 2012).

A second interval of increased ice buildup may have occurred in the late Bashkirian through first half of the Moscovian (Fig. 1b) as suggested by the re-occurrence of glacial deposits in eastern Australia (Fielding et al. 2008a), the occurrence of the first Pennsylvanian-age deposits in the Paganzo and Rio Blanco basins of NW Argentina (Gulbranson et al. 2010; Henry et al. 2010; Buso et al. 2020), and possibly contemporaneous glaciogenic deposits of a third glaciation cycle in the Parnaiba and Paraná basins, Brazil (Rosa et al. 2019; Buso et al. 2020; Caputo and Santos 2020), although Carboniferous age constraints are limited in the Brazilian successions (Holz et al. 2010; Griffis et al. 2018; Rosa et al. 2019). Glaciogenic deposits of this age also record the onset of glaciation in the Karoo Basin (although weak age constraints make it possible that this could have been earlier; Visser 1997; Stollhofen et al. 2008), Antarctica, and in the eastern Arabian Shield (Martin et al. 2008). A sustained (up to 4 myr) drop in sea-level during this time (Ross and Ross 1987; Eros et al. 2012a, b; Huang et al. 2017) and the return of superimposed high-magnitude short-term glacioeustasy (Bishop et al. 2010; Waters and Condon 2012) further support renewed ice expansion in the first half of the Middle Pennsylvanian Epoch.

A third short-lived (<1 myr) but intense glaciation occurred in the earliest Kasimovian and is recognized primarily by a major drop in sea-level (Eros et al. 2012a), and deep incision and karstification of exposed surfaces, which can be correlated across palaeotropical successions in central Euramerica (Belt et al. 2011; Falcon-Lang et al. 2011). Although this glaciation is contemporaneous with a halving of palaeo-CO2 concentrations (Fig. 1d; from an average of 500 ppm to 200–300 ppm) (Richey et al. 2020), there are no well-dated glaciogenic deposits of this age that indicate a discrete glaciation. Although the highest magnitude glacioeustasy is inferred for this interval (Fig. 1c; Soreghan and Giles 1999), chronostratigraphic constraints preclude establishing this relationship with confidence. The demise of this short-lived ice age is coincident with a marked rise in atmospheric CO2 concentrations (from <300 to up to 850 ppm) over 105 years (Fig. 3b) recorded in a contemporaneous c. 4‰ decrease in marine δ13C values (Fig. 3d). The time-equivalent c. 1‰ decrease in brachiopod δ18O following a short-term rise in values through the early Kasimovian (Fig. 3c) is interpreted as recording abrupt warming at this time of doubling of CO2 concentrations. Notably, the early Kasimovian short-lived and intense glaciation and its subsequent greenhouse gas-forced abrupt deglaciation was characterized by major biotic change with important evolutionary turnover in many marine taxa (Stanley and Powell 2003; Lucas et al. 2021). The loss of arborescent lycopisids and many other wetland groups throughout much of Euramerica occurred during the Kasimovian glacial interval and has been attributed to the crossing of a CO2 viability threshold (Richey et al. 2021). It is not clear as to whether a recently identified significant drop in marine species and genus richness in the Kasimovian (Fan et al. 2020) was in response to the early Kasimovian glacial or the subsequent abrupt CO2-forced warming, or possibly the impact of both. This issue warrants further study.

The apex of late Paleozoic glaciation, however, is broadly considered to have started in the Gzhelian Age (Late Pennsylvanian Epoch) and continued into the Sakmarian Age of the early Permian Period. This inferred peak of glaciation is archived by widespread sub-glacial and glaciomarine deposits throughout South America, western and central Africa, Antarctica, eastern and western Australia, Tasmania, the Arabian Peninsula, India, and the Cimmerian blocks (Fig. 1a, b) (Veevers and Powell 1987; Visser 1997; Wopfner and Casshyap 1997; Fielding et al. 2008a, b; Martin et al. 2008; Mory et al. 2008; Rocha-Campos et al. 2008; Stollhofen et al. 2008; Isbell et al. 2012, 2013; Montañez and Poulsen 2013; Craddock et al. 2019; Soreghan et al. 2019).

The spatial and temporal distribution of glaciogenic deposits in Gondwana (Fig. 1b) indicate broadly a west-to-east progression in glaciation through time (Isbell et al. 2003, 2012; Buso et al. 2020). The extensive glaciomarine deposits of latest Pennsylvanian to early Permian periods (Fielding et al. 2008b) indicate that both glaciers and ice sheets in southern Gondwana reached sea-level (Isbell et al. 2012, 2013). Moreover, U–Pb ages and Hf isotopic compositions of detrital zircons recovered from late Carboniferous through early Permian strata throughout southern Gondwana indicate multiple large ice centres and far-field (1000 to 2000 km) ice transport, suggesting trans-continental ice sheets during the apex of glaciation (Craddock et al. 2019; Griffis et al. 2019). Hypothesized cooling and large ice volume during this time is further suggested by a progressive increase in brachiopod δ18O values (not observed for conodonts) following a previous decline in δ18O values through the latter half of the Pennsylvanian Epoch (Fig. 3c; Grossman and Joachimski 2020).

The apex of glaciation straddling the close of the Carboniferous and earliest Permian (Fig. 1a, b) is archived in mid- to low-latitude successions of Euramerica and South China as a notable basinward shift of facies and relative sea-level fall (West et al. 1997; Olszewski and Patzkowsky 2003; Eros et al. 2012a; Huang et al. 2017; Eriksson et al. 2019), and in carbonate-dominated successions, by substantial exposure and an erosional unconformity (Melvin et al. 2010; Koch and Frank 2011). Notably, the apex of continental ice accumulation was coincident with a 10 myr nadir in palaeo-CO2 when concentrations dropped to well below 300 ppm (Fig. 3a, b) (Richey et al. 2020). The CO2 fall is archived as a ≤6‰ increase in marine δ13C values (Fig. 3d) interpreted to record a major increase in organic carbon burial, possibly involving a shift in the primary locus of carbon burial from land (peats) to the ocean (Chen et al. 2018; White et al. 2020). This supports atmospheric CO2 as the primary driver of cyrosphere evolution during the LPIA (Montañez et al. 2007, 2016; Richey et al. 2020). This apex of late Paleozoic glaciation was also the time of O2/CO2 ratios (700 to 960) nearly double that of present day (550). On land, these atypical atmospheric conditions were coincident with a permanent shift in the composition and architecture of palaeotropical forests (DiMichele et al. 2009; DiMichele 2014) and the acquisition of terrestrial adaptations in crown tetrapods and radiation and eventual dominance of dryland-adapted amniotes, possibly shaping the phylogeny of modern terrestrial vertebrates (Pardo et al. 2019; Lucas et al. 2021).

Over the past decade, many new high-precision, CA(-ID)-TIMS single zircon U–Pb ages have become available and integrated with existing chronostratigraphic constraints for uppermost Carboniferous and lower Permian (Cisuralian Series) depositional basins that hosted palaeo-ice centres in southern Gondwana (blue horizontal lines in Fig. 1b). These U–Pb-calibrated chronostratigraphic frameworks further support a dynamic glaciation history. In the Paganzo and Rio Blanco basins (NW Argentina), U–Pb ages indicate that southwestern Gondwana was mostly terminally deglaciated by 315 Ma (Moscovian Age) (Gulbranson et al. 2010). To the east in South America (Brazil), high-precision U–Pb ages indicate that subglacial deposits and evidence of grounded ice in the Paraná Basin are confined to the Carboniferous System (Fig. 1b; Griffis et al. 2018), with continental ice developing perhaps as early as the middle Visean and possibly even the Tournasian (Rosa et al. 2019). Conversely, glaciogenic deposits persist into Permian successions of eastern and western Australia, Tasmania, and southern Africa (Fig. 1b; Fielding et al. 2008a, b; Mory et al. 2008; Stollhofen et al. 2008; Henry et al. 2012; Isbell et al. 2013; Backhouse and Mory 2020).

It has long been proposed that the eastward progression of permanent loss of ice was associated with diachronous, over tens of millions of years, deglaciation across Gondwana (Caputo and Crowell 1985; Isbell et al. 2003, 2012; Limarino et al. 2014; Moxness et al. 2018). Recent high-precision U–Pb dating of stratigraphic deposits in palaeo-ice centres of west-central Gondwana, however, does not support this hypothesized diachronous loss of ice across this core region of Gondwana. Rather a series of deglaciations in the latest Carboniferous Period through to the late early Permian Period are revealed with each deglaciation synchronous and geographically widespread. The first deglaciation event occurred in the latest Carboniferous Period (300 Ma ± 1 Ma, possibly as far back as 304 Ma) and is expressed as a major transgression across the Paraná (Brazil), Kalahari (Namibia), and Karoo (South Africa) basins of west-central Gondwana (Cagliari et al. 2016; Griffis et al. 2018, 2019, 2021). It is possible, but not demonstrated, that this deglaciation was initiated during the Alekyvo Climactic Optimum of elevated palaeo-CO2. For the Paraná Basin, this deglaciation event marks the terminal demise of glaciers, whereas there is a return to glaciated landscapes in the Kalahari and Karoo basins c. 298 Ma, in the earliest Permian Period (Fig. 1b; Bangert et al. 1999; Werner 2006; Stollhofen et al. 2008; Griffis et al. 2019, 2021; Schemiko et al. 2019; Zieger et al. 2019, 2020). This was synchronous with continental ice accumulation in many peri-polar and polar basins of Gondwana (Fig. 1b; Fielding et al. 2008b; Isbell et al. 2012; Montañez and Poulsen 2013).

A second episode of synchronous ice loss and transgression at 296 Ma (Asselian) across southern Africa is archived in all three west-central Gondwanan basins (Fig. 1b; Visser 1997; Stollhofen et al. 2008; Griffis et al. 2019). This transgression marks the terminal demise of glacial conditions (Fig. 1b) in the Kalahari Basin (Griffis et al. 2021), is the timing of Deglaciation Sequence III in South Africa (Griffis et al. 2019) and is penecontemporaneous with terminal deglaciation (c. 295 Ma) in the Bonaparte Basin of northwestern Australia (296 to 295 Ma; age of P. confluens–P. pseudoreticulata Zone boundary; Phillips et al. 2018; Backhouse and Mory 2020). A return to glacial conditions in west-central Gondwana in the Sakmarian Age is constrained to the Karoo Basin (Fig. 1b). The third and terminal deglaciation for this region occurred at c. 283 to 282 Ma, in the later part of the Cisuralian Epoch (Griffis et al. 2019, 2021) and may have been penecontemporaneous with a deglaciation event in Antarctica. Younger glacial deposits of Guadalupian and Lopingian age, interpreted as restricted alpine glaciers, occur in eastern Australia (Fielding et al. 2008a; Metcalfe et al. 2015; Garbelli et al. 2019) and New Zealand (Waterhouse and Shi 2010), and possibly, based on less well chronostratigraphically constrained deposits, in Antarctica, Angola and India (not shown on Fig. 1b) (Isbell et al. 2008; Wopfner and Jin 2009; Lakin et al. 2016). This suggests that continental ice existed, albeit restricted, in parts of central and eastern Gondwana until c. 255 Ma.

Overall, the geographical distribution of glaciogenic deposits in Gondwanan successions and inferred duration of glaciations decreased after the mid- to late Sakmarian (Isbell et al. 2012, 2013; Fielding et al. 2018b; Garbelli et al. 2019; Soreghan et al. 2019) with ice increasingly restricted to alpine glaciers. In lower-latitude successions, there is a widespread loss of cyclicity in post-mid-Sakmarian successions (e.g. Stemmerick 2008), and stratigraphic indicators of glacioeustasy record diminished magnitudes of sea-level fluctuations (Fig. 1c; Rygel et al. 2008). Together, this suggests a pattern of stepwise demise of the LPIA that mirrors its inferred stepwise onset (see section ‘Onset of the late Paleozoic icehouse’). Reconstructed palaeo-CO2 for the Cisuralian Epoch (Figs 1d & 3b) documents the relatively abrupt onset of rising concentrations at the close of the Sakmarian Age (Richey et al. 2020), with palaeo-CO2 rising above the modelled continental ice-initiation threshold for the early Permian (560 ppm; Lowry et al. 2014). This trend in palaeo-CO2 argues for a greenhouse gas-forcing for the demise of the LPIA (Montañez et al. 2007; Richey et al. 2020). Analogous to the onset of the LPIA, its demise was accompanied by major faunal and floral reorganization and changes in climate and the hydrological cycle (Cleal and Thomas 2005; Powell 2007; Clapham and James 2008; Waterhouse and Shi 2010; McGhee et al. 2013; Montañez and Poulsen 2013; DiMichele 2014; Richey et al. 2020, 2021).

Despite a rich record of glaciogenic deposits and geomorphic features from Gondwana, a major contributor to our understanding of the late Paleozoic glaciation history is the stratigraphic record of the low palaeo-latitudes. Carboniferous and early Permian stratal patterns of Euramerican platform successions (e.g. Midcontinent, Donets, Appalachian, and Illinois basins on Fig. 2), typified by cyclothems, have long been interpreted to record glacioeustatic fluctuations driven by the waxing and waning of high-latitude Gondwanan glaciers (Wanless and Shepard 1936; Crowell and Frakes 1970; Heckel 1977, 1986, 2002, 2008; Ross and Ross 1985; Veevers and Powell 1987; Algeo and Wilkinson 1988; Eros et al. 2012a; Cecil et al. 2014; Fielding et al. 2020). Sedimentary successions that document systematic stacking patterns of lithofacies have been argued to provide a more complete record of the glaciation and deglaciation history than higher-latitude glaciogenic deposits given the potential erosive effects of palaeo-glaciers (e.g. Rygel et al. 2008). Moreover, sedimentary geologists have long recognized that the sedimentary architectures of both carbonate- and siliciclastic-dominated systems differ between successions deposited during greenhouse, transitional, or icehouse climates (e.g. Read 1995, 1998; Lehrmann and Goldhammer 1999; Sømme et al. 2009; Eriksson et al. 2019; Cosgrove et al. 2021). These differences are observed in the stratigraphic architectural elements and stacking patterns, and in the nature of the bounding surfaces between stratal units that collectively define genetic packages referred to as sequences, parasequences, and systems tracts (Fig. 4a–c).

Fig. 4.

Synthetic platform stratigraphy and idealized stratigraphic architectures. Forward-modelled stratigraphies (a–c) represent a 400 kyr duration and were generated using a hierarchy of 400, 100, 40 and 20 kyr sea-level changes (inset); modified after Read (1998). (a) Greenhouse carbonate-dominated stratigraphy for sea-level fluctuations of 400 kyr: 1 m; 100 kyr: 2 m; 40 kyr: 2 m; 20 kyr: 5 m showing layer-cake stratigraphy and stacking of 5th-order parasequences composed of shallow (orange) to deep subtidal (blue) facies capped by tidal-flat carbonates or fluvial to shoreface siliciclastics (yellow). Mud-mounds (within the deeper subtidal facies) form during transgressions. (b) Generalized icehouse stratigraphy showing characteristics of stratigraphic architectures that develop with large-scale changes in accommodation space and restricted progradation of facies: deep-fluvial incision, karstification and palaeosol development, widespread disconformities (red lines), downdip imbrication/shingling of prograding (regressive) parasequences, and complex lateral relationships. (c) Icehouse stratigraphy for 400 kyr-dominated eustasy (inset) and sea-level fluctuations of 400 kyr: 20 m; 100 kyr: 70 m; 41 kyr: 5 m; 19 kyr: 1 m modelled to be analogous to Pennsylvanian cyclothems. Synthetic stratigraphy shows a 4th-order (400 kyr) disconformity bounded sequence within which four disconformity bounded 100 kyr parasequences are nested. Note well-developed ramp-margin (lowstand) wedge and thin backstepping TST with maximum flooding surface (mfs) (darker blue facies) extending far back onto the platform. Shingled parasequences of the highstand and forced regressive systems tracts are composed of shallow subtidal lithofacies and downlap onto the mfs, producing highly variable thickness and composition patterns of cycles by location. Black lines updip are exposure/erosion surfaces. (d) Idealized stratigraphic architecture for a 4th-order Pennsylvanian sequence deposited during the main phase of the LPIA icehouse (modified after Eriksson et al. 2019). Note that coals predate transgressive shales and can occur directly overlying incised valley fills. The diagram is drawn parallel to depositional strike.

Fig. 4.

Synthetic platform stratigraphy and idealized stratigraphic architectures. Forward-modelled stratigraphies (a–c) represent a 400 kyr duration and were generated using a hierarchy of 400, 100, 40 and 20 kyr sea-level changes (inset); modified after Read (1998). (a) Greenhouse carbonate-dominated stratigraphy for sea-level fluctuations of 400 kyr: 1 m; 100 kyr: 2 m; 40 kyr: 2 m; 20 kyr: 5 m showing layer-cake stratigraphy and stacking of 5th-order parasequences composed of shallow (orange) to deep subtidal (blue) facies capped by tidal-flat carbonates or fluvial to shoreface siliciclastics (yellow). Mud-mounds (within the deeper subtidal facies) form during transgressions. (b) Generalized icehouse stratigraphy showing characteristics of stratigraphic architectures that develop with large-scale changes in accommodation space and restricted progradation of facies: deep-fluvial incision, karstification and palaeosol development, widespread disconformities (red lines), downdip imbrication/shingling of prograding (regressive) parasequences, and complex lateral relationships. (c) Icehouse stratigraphy for 400 kyr-dominated eustasy (inset) and sea-level fluctuations of 400 kyr: 20 m; 100 kyr: 70 m; 41 kyr: 5 m; 19 kyr: 1 m modelled to be analogous to Pennsylvanian cyclothems. Synthetic stratigraphy shows a 4th-order (400 kyr) disconformity bounded sequence within which four disconformity bounded 100 kyr parasequences are nested. Note well-developed ramp-margin (lowstand) wedge and thin backstepping TST with maximum flooding surface (mfs) (darker blue facies) extending far back onto the platform. Shingled parasequences of the highstand and forced regressive systems tracts are composed of shallow subtidal lithofacies and downlap onto the mfs, producing highly variable thickness and composition patterns of cycles by location. Black lines updip are exposure/erosion surfaces. (d) Idealized stratigraphic architecture for a 4th-order Pennsylvanian sequence deposited during the main phase of the LPIA icehouse (modified after Eriksson et al. 2019). Note that coals predate transgressive shales and can occur directly overlying incised valley fills. The diagram is drawn parallel to depositional strike.

A stratigraphic sequence, which is independent of time or spatial scales, is defined as a relatively conformable and genetically related stratal package that corresponds to a full stratigraphic cycle and is bounded by basin-wide unconformities and their correlative conformities (Van Wagoner et al. 1990; Catuneanu et al. 2009). Notably, the term ‘sequence’ in sequence stratigraphy is not interchangeable with the use of stratigraphic ‘succession’. Sequences of greenhouse depositional systems can be bounded by flooding surfaces that record minimal change in accommodation space and minimal to no exposure (disconformity) (Fig. 4a). Sequences in transitional or icehouse systems are defined by pronounced bounding unconformities, which pass down depositional dip into correlative conformable contacts (Fig. 4b, c). Parasequences stack into parasequence sets, which in turn define systems tracts (Fig. 4a–c) and are conventionally defined as relatively conformable packages of genetically related strata bounded by marine flooding surfaces. The relevance of the term ‘parasequence’ has been questioned for some depositional settings (Posamentier and Allen 1999; Catuneanu et al. 2009) and has been applied quite variably, although generally referring to cycles nested within larger-scale cycles.

Sequence stratigraphic terminology has been widely applied to late Paleozoic successions given that it provides a framework in which to identify and interpret stratigraphic architectures and link them mechanistically to relative sea-level changes and climate (Weibel 1996; Smith and Read 2000; Al-Tawil and Read 2003; Cecil et al. 2003; Olszewski and Patzkowsky 2003; Feldman et al. 2005; Bishop et al. 2010; Belt et al. 2011; Martin et al. 2012; Eros et al. 2012a; Heckel 2013; Ahern and Fielding 2019, 2021; Eriksson et al. 2019; Fielding et al. 2020, among others).

The following sections present the observed and simulated differences in stratigraphic architectures and stratal stacking patterns for platform/shelf depositional systems that developed under different climate states. The intent of these sections is to illustrate how stratigraphic patterns and relationships have been used to constrain the onset of glaciation in the late Paleozoic, to infer relative changes in ice volume through time, and to estimate magnitudes of glacioeustatic fluctuations from low-palaeo-latitude successions.

Greenhouse depositional systems develop low platform/shelf topography due to small changes in accommodation that are driven by low-magnitude sea-level fluctuations (of a few metres) and for which sedimentation rates can easily keep up and fill the accommodation space (Fig. 4a). The dampened topography permits both siliciclastic and carbonate sediments to prograde across the shelf during low- and highstands of sea-level, creating widespread relatively uniform deposits and ‘layer cake stratigraphy’ with ordered stratal thickness patterns (Lehrmann and Goldhammer 1999). Stratigraphic cyclicity occurs in greenhouse intervals of all ages from which small to moderate magnitude (c. 5 m but up to 20 to 30 m) sea-level fluctuations have been inferred (e.g. Read 1995, 1998; Eriksson et al. 2019; Sømme et al. 2009; Miller et al. 2020). The driving mechanisms, which include draining and filling of continental aquifers and lakes (Jacobs and Sahagian 1993; Wendler and Wendler 2016), however, remain poorly understood. For greenhouse sedimentary systems, the fundamental stratigraphic unit is the 3rd order (c. 0.5 to several myr) sequence. Higher frequency (104-year) parasequences, which are referred to as 5th-order and nested within 3rd-order sequences, are bounded by minimum accommodation flooding surfaces and stack within sequences to define typically thin transgressive and thick highstand systems tracts (Fig. 4a) (Eriksson et al. 2019).

Many carbonate-dominated systems of Early to Middle Mississippian age throughout Euramerica exhibit greenhouse stratal architecture, but with features of cooler climates (relative to the Middle Devonian Epoch). These include widespread deposits of laterally continuous, medium-bedded to massive skeletal-rich (typically bryozoan- and crinoid-rich) muddy (wackestone) to grainy (grainstone) subtidal carbonates (e.g. Katz et al. 2007; Bishop et al. 2009; Manifold et al. 2021). Cool greenhouse carbonates are non-cyclic to cyclic (metre-scale) with little or no evidence of significant subaerial exposure. Non-erosional disconformities typically occur within or at the top of tidal-flat facies or cross-bedded shoreface sandstones, or less commonly, occur as palaeosols. Isolated mud mounds (i.e. build-ups), composed of a crinoid-bryozoan cores, micro-peloidal micrite cements and flanking grainstones, rich in crinoidal debris and binding fenestral bryozoa, developed along Mississippian carbonate-dominated distally steeped ramp margins, foundering margins of platforms, or adjacent to intrashelf basins where sufficient accommodation space existed (Monty et al. 1995, and papers within; Read 1998; Ahr et al. 2003, and papers within).

The first occurrence of icehouse type stratal architecture is observed as early as the mid- to upper Visean (c. 335 to 332 Ma) to as late as earliest Serpukhovian successions (c. 330 Ma) in North America (Al-Tawil and Read 2003; Batt et al. 2008; Bishop et al. 2009; Gastaldo et al. 2009; Fielding and Frank 2015; Ahern and Fielding 2019, 2021; Chen et al. 2019; Smith and Read 1999, 2000), the UK (Wright and Vanstone 2001; Barham et al. 2012; Waters and Condon 2012; Pointon et al. 2014), and China (Wang et al. 2019). Collectively, the observed change in stratal architecture is two-fold. First, for carbonate-dominated platforms, there is a regional-scale facies shift from shallow-marine, carbonate-dominated sequences bounded by non-erosional disconformities to mixed carbonate-siliciclastic or subtidal carbonate sequences bounded by well-developed subaerial erosion and dissolution features (Ramsbottom 1977; Walkden 1977; Somerville and Strank 1984; Horbury 1989; Smith and Read 1999, 2000; Wright and Vanstone 2001; Batt et al. 2008; Waters and Condon 2012; Manifold et al. 2021). In fluvial systems this facies shift is recorded by a major shift in fluvial styles at this time (Allen et al. 2011, 2014). Secondly, the frequency of the predominant cyclicity in carbonate- and siliciclastic-dominated systems shifts from 3rd- (106-year) to 4th-order (105-year) sequences, making the 4th-order sequence the fundamental stratigraphic unit in the uppermost Mississippian Series (Eriksson et al. 2019). This onset of predominant shorter-term cyclicity, however, is not observed in all basins (Manifold et al. 2021), but where observed, it records an increase in the magnitude of the 4th-order sea-level fluctuations that obscures the 3rd-order signal (Fig. 4b, c; see also Fig. 1c) (cf. Lehrmann and Goldhammer 1999). Fourth-order sequences are separated by well-developed subaerial erosion surfaces commonly with significant relief (Eriksson et al. 2019), including development of deeply incised valleys (Fig. 4d) (aforementioned references; summarized in Bishop et al. 2009; Fielding and Frank 2015). Maximum flooding facies (black and grey shales; transgressive limestones) and flooding surfaces are geographically widespread (Fig. 4c) (e.g. Heckel 2008, 2013). This icehouse stratigraphic architecture persists through the remainder of the Carboniferous and into the early Permian Period, but with long-term temporal variability in cycle stacking patterns (discussed below).

The aforementioned regional facies shift is interpreted to archive a dramatic increase in ice accumulation in Gondwana by the close of the Middle Mississippian. The notable unconformities capping sequences reflect markedly increased changes in accommodation due to higher magnitude sea-level fluctuations and the consequent repeated basinward migration of the shoreline by tens of kilometres, promoting a greater degree and duration of surface exposure (Fig. 4b, c; cf. Read 1995, 1998; Sømme et al. 2009; Eriksson et al. 2019). Such rapid changes in accommodation, driven by glacioeustasy, maintain deep platforms, thus limiting the extent of highstand progradation of sediment and leading to unfilled accommodation space and uneven depositional topography (Fig. 4b, c) (Read 1998; Sømme et al. 2009; Alnazghah and Kerans 2018). These changes are recorded in the stratigraphy by vertical juxtaposition of deeper water marine facies over fluvial and estuarine facies and palaeosols (Fig. 4d) and by downdip shingling of prograding (regressive) parasequences, complex lateral relationships, and variable thickness patterns of cycles (Fig. 4b, c) (Read 1998; Lehrmann and Goldhammer 1999; Alnazghah and Kerans 2018).

Thus, the transition from greenhouse to icehouse depositional systems is imprinted in the late Middle Mississippian to earliest Pennsylvanian stratigraphic architecture by the onset of cycles exhibiting well-developed exposure and erosion surfaces (Fig. 4b). The erosion surfaces cap a range of underlying lithofacies, reflecting the limited lateral continuity of some lithofacies (Fig. 4b, c) (Watney et al. 1989; Read 1995, 1998; Lehrmann and Goldhammer 1999; Barnett et al. 2002; Alnazghah and Kerans 2018). The Upper Devonian through lower Permian Arrow Canyon succession of eastern CA, which hosts the mid-Carboniferous GSSP, is one example that records this stepped transition from greenhouse to fully icehouse stratigraphic architectures (Bishop et al. 2009, 2010).

Cyclothems are characteristic of stratigraphic cycles that were deposited on platforms/shelves under icehouse regimes (Watney et al. 1989; Read 1998). The term ‘cyclothem’ was first coined by Wanless and Weller (1932) for the repetition of lithofacies groupings in palaeotropical successions deposited during a single sedimentary cycle and separated below and above by unconformities (Fig. 5). Wanless and Weller (1932) specifically referred to these metre- to tens of metres-thick stratigraphic cycles as the ‘type that prevailed during the Pennsylvanian Period’. By nature of being unconformity bounded stratigraphic units, cyclothems are considered genetic sequences (Weibel 1996; Heckel 2002, 2008; Eros et al. 2012a).

Fig. 5.

Select cyclothems from the Illinois and Donets basins. (a) Stacked Middle Pennsylvanian (Desmoinesian) cyclothems of the Illinois Basin (equivalents of the Midcontinent Pawnee and Altamont ‘major’ cyclothems), modified after Nelson et al. (2011): #1 = deltaic siliciclastics of highstand systems tract (HST) and falling stage systems tract (FSST) into which incises valleys erode; #2 = FSST (regressive) of late highstand to lowstand systems tract (LST); point of maximum rate of sea-level fall denotes beginning of the LST. Incised valley and channel formation with initial fill of fluvial deposits; palaeosol formation on interfluves; #3 = coals of late LST that formed as peats once the rate of subsidence (accommodation space) outpaced the rate of sea-level fall; regional coals (Jamestown, Baker, and Cottage) are not laterally correlatable; #4 = black shales (some phosphatic in lower layers); black shales in some cyclothems (e.g. Energy and Anna shales of Pawnee) are underlain by wedge-shaped deposits of heterolithic grey siltstones and mudstones (estuarine), which exhibit tidal rhythmicity and thicken into channels. Transgressive surfaces (ts) denote onset of transgressive systems tract (TST); truncation of underlying facies and burial of in-situ lycopsids found in living position and rooted in underlying coals, record eustatically driven rapid rise in accommodation space. #5 = open marine muddy to grainy marine limestones; #6 = FSST – same as #1. (b) Pennsylvanian cyclothem types in the Donets Basin succession modified after Eros et al. (2012a). Up-dip and down-dip positions of correlated retrogradational, aggradational, and progradational cyclothems illustrate cross-ramp (hundreds of kilometres) variations in the facies assemblages (FA).

Fig. 5.

Select cyclothems from the Illinois and Donets basins. (a) Stacked Middle Pennsylvanian (Desmoinesian) cyclothems of the Illinois Basin (equivalents of the Midcontinent Pawnee and Altamont ‘major’ cyclothems), modified after Nelson et al. (2011): #1 = deltaic siliciclastics of highstand systems tract (HST) and falling stage systems tract (FSST) into which incises valleys erode; #2 = FSST (regressive) of late highstand to lowstand systems tract (LST); point of maximum rate of sea-level fall denotes beginning of the LST. Incised valley and channel formation with initial fill of fluvial deposits; palaeosol formation on interfluves; #3 = coals of late LST that formed as peats once the rate of subsidence (accommodation space) outpaced the rate of sea-level fall; regional coals (Jamestown, Baker, and Cottage) are not laterally correlatable; #4 = black shales (some phosphatic in lower layers); black shales in some cyclothems (e.g. Energy and Anna shales of Pawnee) are underlain by wedge-shaped deposits of heterolithic grey siltstones and mudstones (estuarine), which exhibit tidal rhythmicity and thicken into channels. Transgressive surfaces (ts) denote onset of transgressive systems tract (TST); truncation of underlying facies and burial of in-situ lycopsids found in living position and rooted in underlying coals, record eustatically driven rapid rise in accommodation space. #5 = open marine muddy to grainy marine limestones; #6 = FSST – same as #1. (b) Pennsylvanian cyclothem types in the Donets Basin succession modified after Eros et al. (2012a). Up-dip and down-dip positions of correlated retrogradational, aggradational, and progradational cyclothems illustrate cross-ramp (hundreds of kilometres) variations in the facies assemblages (FA).

Although ideal vertical lithofacies successions for cyclothems have been proposed (e.g. Wanless 1957; Heckel 1994, 2008), these stratigraphic cycles exhibit variability in their stratal order and range of lithofacies (Fig. 5) (e.g. Wilkinson et al. 2003; Weibel 2004; Eros et al. 2012a; Burgess 2016; Fielding et al. 2020). This stratal variability reflects differences in regional subsidence rates, platform slope, palaeotopography, sediment influx, and climate as well as temporal variability in the magnitudes and rates of sea-level change (Watney et al. 1989; Read 1995, 1998; Eros et al. 2012a). On a larger stratigraphic and spatial scale, however, distinct stratal stacking patterns are recognized (e.g. Fig. 6; Heckel et al. 1998; Olszewski and Patzkowsky 2003; Heckel 2008, 2013; Elrick et al. 2009; Eros et al. 2012a; Ahern and Fielding 2019; Fielding et al. 2020), which, coupled with conodont biostratigraphic events, has led to proposed correlations of cyclothems across Euramerica (Fig. 6) (Cecil et al. 2003; Feldman et al. 2005; Heckel et al. 2007; Heckel 2008, 2013; Belt et al. 2011; Eros et al. 2012a; Schmitz and Davydov 2012; Rosscoe and Barrick 2013; Fielding et al. 2020). Proposed intercontinental correlation of cyclothems (Heckel et al. 2007; Eros et al. 2012a; Schmitz and Davydov 2012), in particular, may offer a chronostratigraphic framework (Fig. 6) of higher resolution for cyclothemic successions than typically afforded by Carboniferous biostratigraphy. But, such correlations, which are based on conodont biostratigraphy, require the assumption that conodonts globally evolved by phyletic gradualism that may be problematic (Lucas 2020) and thus would benefit from further study.

Fig. 6.

Hierarchy of Donets sequences and the proposed equivalence to Midcontinent cyclothems, modified after Eros et al. (2012a). Donets sequences (e.g. Gz-4) shown with predominant limestone marker bed (denoted by red letters and numerals) on left. Composite and longer-term composite sequences shown in blue on the right. U.S. Midcontinent cyclothems shown as ‘major’, ‘intermediate’, and ‘minor’ cyclothems as defined by Heckel (1994); see text for further details. Grey dashed tie-lines indicate Heckel et al.'s (2007) ‘digital correlation’ of Donets limestone marker beds to the Midcontinent cyclothems. Superscript ‘+’ indicates the presence of additional minor cyclothems in the Midcontinent record and ‘*’ denotes less confident inter-continental correlations.

Fig. 6.

Hierarchy of Donets sequences and the proposed equivalence to Midcontinent cyclothems, modified after Eros et al. (2012a). Donets sequences (e.g. Gz-4) shown with predominant limestone marker bed (denoted by red letters and numerals) on left. Composite and longer-term composite sequences shown in blue on the right. U.S. Midcontinent cyclothems shown as ‘major’, ‘intermediate’, and ‘minor’ cyclothems as defined by Heckel (1994); see text for further details. Grey dashed tie-lines indicate Heckel et al.'s (2007) ‘digital correlation’ of Donets limestone marker beds to the Midcontinent cyclothems. Superscript ‘+’ indicates the presence of additional minor cyclothems in the Midcontinent record and ‘*’ denotes less confident inter-continental correlations.

Each genetic sequence (i.e. cyclothem) is inferred to have been deposited within a single cycle of relative sea-level change (Fig. 5a). Cyclothems are bounded by exposure surfaces denoted by mature palaeosols, karstification of carbonate facies (Fig. 4b), and surface erosion expressed as small channel incisions (up to several 100 m wide; <10 m deep) to kilometre-wide incised valleys with up to tens of metres of erosional relief (Fig. 4d; e.g. #2 on Fig. 5a). Exposure surfaces can be quite laterally persistent (Fig. 4b, c) (e.g. Belt et al. 2011; Falcon-Lang et al. 2011), each being attributed to forced regression driven by sea-level fall (Fig. 4b) with the maximum rate of sea-level fall marking the sequence boundary (Feldman et al. 2005; Fischbein et al. 2009; Eros et al. 2012a). A typical stacking pattern of facies associations in a cyclothem (Fig. 5a) includes the following five elements. (1) A basal sandstone made up of single to multi-storey bodies of conglomerate and sandstones (#2 on Fig. 5a), interpretated as fluvial deposits, which overlie incised valley and channel surfaces (sequence boundaries) and, in some cases, by tidally influenced fluvial sandstones. (2) Coals (#3), up to a metre or more thick, contemporaneous with and overlying palaeosols (underclays), which developed on interfluves. (3) Estuarine heterolithic laminated siltstone and muddy sandstone alternations (#4) and/or regionally widespread marine grey to black phosphatic shales (#4). The sharp contact typically overlying coals or other late lowstand facies, is a ravinement surface (i.e. transgressive surface ('ts’ on Fig. 5a) that marks the first significant flooding surface in a sequence (cyclothem). It further denotes the onset of the transgressive systems tract during which time the rate of increase in accommodation space due to sea-level rise rapidly outpaces the rate of sediment supply. (4) The lithofacies (heterolithic siliciclastics, open-water limestones, or black shales) directly overlying the transgressive surface is determined by the rate of and maximum accommodation space provided by the sea-level rise. And the deepest water lithofacies present (e.g. phosphatic nodule-rich black shales; deepest water limestones) marks the maximum flooding surface (‘mfs’ on Fig. 5a). (5) Coarsening-upward and sandstone-dominated fluvio-deltaic deposits of the highstand and forced regressive systems tracts overlie the maximum flooding surface (mfs) and are capped by erosion surfaces or palaeosols (Fig. 4b, c; #1 and 6 on Fig. 5a).

In the US Midcontinent and Donets Basin, a hierarchical order of cyclothems has been noted for decades (Busch and Rollins 1984; Heckel 1994; Izart et al. 2003). In the Midcontinent region, cyclothems that record the largest-scale transgression (i.e. widespread shales that extend across the shelf and into the deeper-water basin) are referred to as major cyclothems, whereas those lacking deepest-water shales and with grey shales or limestones that extend less far across the platform are referred to as intermediate cyclothems (Fig. 6) (Heckel 2008, 2013). Smaller metre-scale cycles, referred to as minor cyclothems are nested within major and intermediate cyclothems. For some successions, the vertical stacking patterns of smaller-scale cycles along with changes in lithofacies composition and stratigraphic architectures have been used to define retrogradational, aggradational, and progradational packages of cyclothems (i.e. systems tracts) (Fig. 5b). In turn, the stratal groupings of retrogradational, aggradational, and progradational cyclothems define unconformity bounded larger-scale composite sequences (Fig. 7). Some smaller-scale cyclothems also exhibit exposure and ravinement surfaces and thus have been referred to as stratigraphic genetic sequences (Olszewski and Patzkowsky 2003, 2008; Fielding et al. 2020).

Fig. 7.

Generalized stratigraphy for the Donets cyclothemic succession, hierarchy of Donets sequences, and inferred relative sea-level (onlap–offlap) curve modified after Eros et al. (2012a, b) and adjusted to the Geologic Time Scale 2020 (Gradstein et al. 2020). Cyclothem types as presented in Figure 5b. Correlation of limestone marker beds (red) relative to sequences and stacking patterns of cyclothem types indicated by the red dashed lines. Position of Donets marker beds and Midcontinent cyclothems on relative sea-level curve indicated by red dashed lines and blue and white alternating bands, respectfully. Offset from the horizontal between the two sets of Donets marker beds is a function of nonlinear rescaling of depth–time assignments for the onlap–offlap curve (and thus the position of the marker beds) to accommodate proposed (Schmitz and Davydov 2012) scaling of cyclothem (Donets and Midcontinent) stratigraphic distribution with high-precision U–Pb ages and stage boundaries. For all columns to the right of the relative sea-level curve, the distribution of Donets cyclothem types, sequences, and limestone marker beds (red) within individual global stages were rescaled linearly to address changes between the 2012 and 2020 Carboniferous timescales. The vertical break between the two parts of the diagram denotes these differences.

Fig. 7.

Generalized stratigraphy for the Donets cyclothemic succession, hierarchy of Donets sequences, and inferred relative sea-level (onlap–offlap) curve modified after Eros et al. (2012a, b) and adjusted to the Geologic Time Scale 2020 (Gradstein et al. 2020). Cyclothem types as presented in Figure 5b. Correlation of limestone marker beds (red) relative to sequences and stacking patterns of cyclothem types indicated by the red dashed lines. Position of Donets marker beds and Midcontinent cyclothems on relative sea-level curve indicated by red dashed lines and blue and white alternating bands, respectfully. Offset from the horizontal between the two sets of Donets marker beds is a function of nonlinear rescaling of depth–time assignments for the onlap–offlap curve (and thus the position of the marker beds) to accommodate proposed (Schmitz and Davydov 2012) scaling of cyclothem (Donets and Midcontinent) stratigraphic distribution with high-precision U–Pb ages and stage boundaries. For all columns to the right of the relative sea-level curve, the distribution of Donets cyclothem types, sequences, and limestone marker beds (red) within individual global stages were rescaled linearly to address changes between the 2012 and 2020 Carboniferous timescales. The vertical break between the two parts of the diagram denotes these differences.

In the Donets Basin cyclothemic succession (Figs 6 & 7) 242 cyclothems of c. 100 kyr duration define (1) composite sequences (c. 400 kyr) that exhibit an internal architecture, which records a long-term cycle of relative sea-level change; (2) and groups of composite sequences that build into ‘longer-term (1.6 ± 0.5 myr) composite sequences’ whose architecture reflects that of the composite sequences that define them (Eros et al. 2012a). An onlap–offlap (relative sea-level) history has been defined based on this hierarchy of cyclicity (Fig. 7). Such hierarchies of cyclicity, recognized in many low-palaeo-latitude basins, have been interpreted by many workers as superimposed scales of sea-level change driven by the interacting influence of astronomical parameters on high-latitude insolation and climate during the LPIA (see following section ‘Cyclostratigraphy as a calibration tool in the Carboniferous’). In addition, sub-metre-scale stratigraphic cycles have been identified in some Carboniferous successions and have been hypothesized to record superimposed decadal- to millennial-scale climatic rhythms (Tucker et al. 2009; Franco et al. 2011; Franco and Hinnov 2012; Kochhann et al. 2020).

As discussed in the previous sections, it is widely accepted that the stratigraphic cycles/sequences of Upper Mississippian to lower Cisuralian (Permian) successions record glacioeustatic fluctuations driven by the waxing and waning of Gondwanan ice sheets. However, robust documentation of a periodic driver of cyclothems is lacking and debate continues over the timescales and magnitudes of inferred sea-level fluctuations. Magnitudes of glacioeustasy have been inferred from Carboniferous and lower Permian deposits using modern facies analogues and depths of incision, with estimates ranging from a few tens of metres to well over 100 m (Fig. 1c) (Heckel 1986; Isbell et al. 2003; Joachimski et al. 2006; Rygel et al. 2008; Bishop et al. 2010; Martin et al. 2012; Montañez and Poulsen 2013; Fielding and Frank 2015; Alnazghah and Kerans 2018; Ahern and Fielding 2019; Fielding et al. 2020). And systematic changes in the stratigraphic patterns of cycles further reveal trends that may indicate shifts between periods of greater and lesser magnitudes of glacioeustasy from the Late Mississippian through early Permian (Fig. 1c) (e.g. West et al. 1997; Smith and Read 2000; Wright and Vanstone 2001; Heckel 2008, 2013; Bishop et al. 2010; Martin et al. 2012; Waters and Condon 2012; Eros et al. 2012a). For example, inferred magnitudes of a couple of tens of metres in the late Visean more than double in the early Serpukhovian (Smith and Read 2000; Ahern and Fielding 2019, 2021) and increase further by the Early Pennsylvanian (Fielding and Frank 2015; Ahern and Fielding 2019), generally mirroring the interpreted changes in continental ice extent (Fig. 1a, b).

Estimates of glacioeustasy through the Pennsylvanian also vary largely (Fig. 1c; Rygel et al. 2008), including posited minimal magnitudes of fluctuation (Dyer and Maloof 2015). Stratal patterns in several Pennsylvanian successions provide evidence for variability in magnitudes of glacioeustasy that appear to be in-step with independently inferred intervals of ice expansion and retraction (Fig. 1a, b). For example, increased amplitudes of glacioeustasy during the first half of the Pennsylvanian have been inferred from Euramerican stratigraphic patterns synchronous with ice expansion. In the Pennine Basin, UK, Waters and Condon (2012) document high frequency of occurrence of unconformities and inferred major marine flooding events with independently reconstructed periods of high-latitude ice expansion. Conversely, they document reduced frequency and magnitude of flooding events and lack of major incision events during intervals of reduced ice volume.

Higher up in the Pennsylvanian succession (upper Moscovian through Gzhelian), larger-scale, myr-scale changes in overall accommodation space and glacioeustatic magnitudes have been inferred from cyclic successions (Feldman et al. 2005; Rygel et al. 2008; Bishop et al. 2010; Eros et al. 2012a, b; Martin et al. 2012; Alnazghah and Kerans 2018; Fielding et al. 2020). These inferred changes are similarly in-step with the independently reconstructed ice buildup to the apex of glaciation and, to some degree, with superimposed shorter-term changes in ice volume (compare Fig. 1a, b). Intriguingly, Alnazghah and Kerans (2018) reconstruct a shift from short- to long-eccentricity cycles coincident with the inferred doubling of eustatic magnitudes in the Kasimovian to Gzhelian Epochs (cf. Bishop et al. 2010) that they attribute to increasing ice volume through this time interval. Estimates of glacioeustatic magnitudes generally decrease through the Cisuralian and Guadalupian epochs of the Permian Period (Rygel et al. 2008), tracking waning ice volume and increasing restriction of ice to alpine glaciers with the demise of the late Paleozoic icehouse.

Late Paleozoic icehouse climate coupled with active tectonics driven by amalgamation of Pangaea promoted geomorphic and climatic diversity, in particular during the Pennsylvanian Subsystem (e.g. Buso et al. 2020). The resulting provincialized flora and fauna makes global biostratigraphic correlations of Pennsylvanian and early Permian successions challenging (Lucas et al. 2021). Global correlations and assigning ages to late Paleozoic strata is further complicated by a relative lack of and uneven geographical distribution of high-precision radioisotope dates. For records that span the LPIA (i.e. Late Devonian through middle Permian), a substantial number of palaeontological and proxy records come from low-latitude successions of Euramerica and the South and North China blocks. There are now a moderate number of high-precision U–Pb ages for low-palaeo-latitude successions (n = 47 Carboniferous; Permian n = 37; Aretz et al. 2020, table 23.2; Henderson and Shen 2020), with many of the dating efforts focused on constraining international stage boundaries. A greater number of high-precision (CA(-ID)-TIMS) U–Pb ages that constrain the timing and location of glacial centres have been published for Carboniferous successions in mid- to high-latitude Gondwanan basins (Gulbranson et al. 2010; Metcalfe et al. 2015; Griffis et al. 2018, 2019, 2021).

Cyclostratigraphy has long been shown to be a powerful tool for correlation (as previously discussed in the section ‘Cyclothems, the fundamental stratigraphic unit of the Pennsylvanian and early Permian’). It has been further demonstrated as a tool for deciphering time in the stratigraphic record if the presence of orbitally forced Milankovitch cycles can be robustly documented in stratigraphic successions (Fischer et al. 1988; Strasser et al. 2006; Hinnov and Ogg 2007). To that end, Euramerican cyclothems, hypothesized to be orbitally forced, have been utilized as an astrochronological tool for developing the age model of the Carboniferous timescale (Davydov et al. 2012; Ogg et al. 2016; Aretz et al. 2020).

Astrochronology, a subdiscipline of cyclostratigraphy, applies the geological record of cyclicity, inferred to archive climate rhythms (i.e. Milankovitch cycles), to astronomical models of palaeoclimate forcing in order to develop a chronometer of orbital (astronomical) resolution (tens to hundreds of kyr increments) (Strasser et al. 2006; Hinnov 2013; Meyers 2015). Astronomical calibration of stratigraphic successions thus permits deciphering the distribution of time in a succession and places constraints on the rates of processes archived in the sedimentary record. If the floating astronomical timescale can be anchored by CA(-ID)-TIMS zircon dates then absolute ages can be assigned to geological and palaeobiological events and the duration and boundaries of biozones and stages within a succession(s). Ultimately, consistent astronomical calibration of contemporaneous successions globally can reveal the temporal relationship of regional and global stages to one another, permit correlation of records globally where alternative chronostratigraphic constraints are associated with large uncertainties or lacking (e.g. continental successions), and refine the geological timescale and vastly improve its temporal resolution (Hinnov 2013). Not only is an astronomical timescale independent of the taxonomic issues that limit biostratigraphy, but it provides a direct link between the astrochronometer and climate change (Meyers 2019).

Statistically robust evidence of astronomical forcing of sedimentary processes has been documented in multiple LPIA successions. A range of spectral and statistical techniques for astrochronological testing and time-calibration (for overviews of methodologies see Weedon 2003; Hinnov 2013, 2018; Meyers 2015, 2019; Sinnesael et al. 2019) have been applied primarily to deeper-water (i.e. open-water) successions, which were deposited in platform margin (i.e. deep shelf or ramp) to slope and/or basinal environments and to a few lacustrine stratigraphic records (Table 1). The focus of astrochronological studies on ‘open water and/or hydrodynamically quiescent’ stratigraphic records reflects that sediments deposited in such environments have a high potential to yield near-continuous depositional archives, to be characterized by relatively low variability in sediment accumulation rates (SARs), and to permit development of high-resolution proxy time series for spectral analysis (e.g. magnetic susceptibility or X-ray fluorescence elemental analysis). These cyclostratigraphy studies (Table 1) have (1) constrained the frequencies of astronomical cycles during the Late Devonian through middle Permian, refining extrapolations of frequencies from astronomical solutions derived for ≤50 myr ago, (2) further confirmed the long-term stability of the long-eccentricity cycle (405 kyr), that in turn can be used to astronomically ‘tune’ sedimentary records, and (3) documented Earth–Moon orbital evolution and provided evidence for chaotic resonance transitions between Earth and Mars orbits during the late Paleozoic (e.g. Fang et al. 2015, 2018a; Wu et al. 2019). Several of these studies have applied the resulting astrochronology to constrain the duration of the Frasnian (De Vleeschouwer et al. 2012, 2013) and Fammenian (Pas et al. 2018) stages and associated boundaries of the Upper Devonian stages, the Serpukhovian Stage of the Mississippian Subsystem (Fang et al. 2018a, b; Wu et al. 2019), all four stages of the Pennsylvanian Subsystem (Wu et al. 2019), and the Roadian and Wordian stages of the Permian Guadalupian Series (Fang et al. 2015). Furthermore, statistical testing for astronomical forcing in open-water LPIA stratigraphic records has documented strong obliquity- and precessional-forcing of climate, seasonal monsoon intensity, oceanic dynamics, and marine primary productivity (Fang et al. 2018a, b; X. Yao et al. 2015; Yao and Hinnov 2019), orbital to millennial-scale climate forcing in the higher latitudes of Gondwana (Kochhann et al. 2020), and long-period astronomical forcing of 3rd-order (106 year) glacioeustasy (e.g. Fang et al. 2015).

Table 1.

Summary of astrochronological studies of late Paleozoic deeper- and quiet-water deposits

Study*MethodsLimitationsOrbital Signal(s)Timescale application
Late DevonianDe Vleeschouwer et al. (2012), (2013) ms-ds; mtmVariable sar; astronomical calibration target (405 kyr); fatsl & S Ecc (c. 100 kyr), lpme, Pr (c. 18 kyr)Constrain duration of Frasnian Stage (6.5 ± 0.4 myr); Giv.–Fras. bdry (383.6 ± 3.0 Ma); Fras.–Fam. bdry (376.7 ± 3.0 Ma)
Pas et al. (2018) ms-ds; mtm, efft esaAstro. calibration target (405 & c. 100 kyr & Ob cycle); high-prec. U–Pb anchored fatsL & S Ecc, lpmo (1.2 myr), ObConstrain Fammenian Stage (13.5 ± 0.5 myr); Fras.–Fam. Bdry (372.4 ± 0.9 Ma)
MississippianFang et al. (2018a, b)ms-ds; mtm, ama & esaAstro. calibration target (405 kyr); fatsL & S Ecc (128 & 95 kyr)., lpmo (1.1 myr), lpme, Ob (33 to 34 kyr); Pr (19 & 15.9 kyr)Constrains duration of Serpukhovian (7.68 ± 0.15 myr)
PennsylvanianWu et al. (2019) ms-ds; mtm, efft amaAstro. calibration target (405 kyr); U–Pb anchored fatsL & S Ecc (136, 122 & 96 kyr)., lpme (2.4, 1.6 & 1.2 myr), lpmo (1.2 myr), lpme, Ob (31 kyr); Pr (22.9 & 19.7 kyr)Constrains duration of Serpukhovian (7.6 myr), Bashkirian (8.1 myr), Moscovian (8.5 myr), Kasimovian (2.87 myr) and Gzhelian (4.83 myr) stages
early PermianKochhann et al. (2020) xrf-ds; mtm, efft & TimeOptRatios of spectral Periodicities tuned to period ratios of astro. calibration targets (132, 124, 98, 96, 35.5, 21.7, 20.6, 17.8, 17.7 kyr);S Ecc, Pr, hemiPr
mid. PermianFang et al. (2015) arm-ds; mtm, esa & amaVariable sar; astro. calibration target (405 kyr); fatsL & S Ecc (c. 95 kyr)., lpmo (c. 1 myr), lpme (c. 2 myr), Ob (c. 44 & 33 kyr), Pr (c. 20 kyr)Constrain duration of Roadian Stage (3.7 ± 0.4 myr); Wordian Stage (2.9 ± 0.4 myr)
X. Yao and Hinnov (2019) (revised Yao et al. 2015)Midpoint-triangle interpolation ds; mtm,Binary lithologic series; variable sar; U–Pb (n = 2) anchored fatsOb (26, 34, 35 & 44 kyr); Pr (21.5, 19.5 & 17.7 kyr); possible L Ecc modulationReveals new Pr and Ob alternations; constrain ages and duration of radiolarian biozones
Huang et al. (2021) ngr-ds; mtm, eha & TimeOpt;Astro. calibration target (405 kyr); fatsL & S Ecc (96.2 to 129.9 kyr)., lpmo (1.2 myr), Ob (40.7, 30 & 25 kyr); Pr (23 & 21–18 kyr)Refines astronomical frequencies for mid. Permian
Study*MethodsLimitationsOrbital Signal(s)Timescale application
Late DevonianDe Vleeschouwer et al. (2012), (2013) ms-ds; mtmVariable sar; astronomical calibration target (405 kyr); fatsl & S Ecc (c. 100 kyr), lpme, Pr (c. 18 kyr)Constrain duration of Frasnian Stage (6.5 ± 0.4 myr); Giv.–Fras. bdry (383.6 ± 3.0 Ma); Fras.–Fam. bdry (376.7 ± 3.0 Ma)
Pas et al. (2018) ms-ds; mtm, efft esaAstro. calibration target (405 & c. 100 kyr & Ob cycle); high-prec. U–Pb anchored fatsL & S Ecc, lpmo (1.2 myr), ObConstrain Fammenian Stage (13.5 ± 0.5 myr); Fras.–Fam. Bdry (372.4 ± 0.9 Ma)
MississippianFang et al. (2018a, b)ms-ds; mtm, ama & esaAstro. calibration target (405 kyr); fatsL & S Ecc (128 & 95 kyr)., lpmo (1.1 myr), lpme, Ob (33 to 34 kyr); Pr (19 & 15.9 kyr)Constrains duration of Serpukhovian (7.68 ± 0.15 myr)
PennsylvanianWu et al. (2019) ms-ds; mtm, efft amaAstro. calibration target (405 kyr); U–Pb anchored fatsL & S Ecc (136, 122 & 96 kyr)., lpme (2.4, 1.6 & 1.2 myr), lpmo (1.2 myr), lpme, Ob (31 kyr); Pr (22.9 & 19.7 kyr)Constrains duration of Serpukhovian (7.6 myr), Bashkirian (8.1 myr), Moscovian (8.5 myr), Kasimovian (2.87 myr) and Gzhelian (4.83 myr) stages
early PermianKochhann et al. (2020) xrf-ds; mtm, efft & TimeOptRatios of spectral Periodicities tuned to period ratios of astro. calibration targets (132, 124, 98, 96, 35.5, 21.7, 20.6, 17.8, 17.7 kyr);S Ecc, Pr, hemiPr
mid. PermianFang et al. (2015) arm-ds; mtm, esa & amaVariable sar; astro. calibration target (405 kyr); fatsL & S Ecc (c. 95 kyr)., lpmo (c. 1 myr), lpme (c. 2 myr), Ob (c. 44 & 33 kyr), Pr (c. 20 kyr)Constrain duration of Roadian Stage (3.7 ± 0.4 myr); Wordian Stage (2.9 ± 0.4 myr)
X. Yao and Hinnov (2019) (revised Yao et al. 2015)Midpoint-triangle interpolation ds; mtm,Binary lithologic series; variable sar; U–Pb (n = 2) anchored fatsOb (26, 34, 35 & 44 kyr); Pr (21.5, 19.5 & 17.7 kyr); possible L Ecc modulationReveals new Pr and Ob alternations; constrain ages and duration of radiolarian biozones
Huang et al. (2021) ngr-ds; mtm, eha & TimeOpt;Astro. calibration target (405 kyr); fatsL & S Ecc (96.2 to 129.9 kyr)., lpmo (1.2 myr), Ob (40.7, 30 & 25 kyr); Pr (23 & 21–18 kyr)Refines astronomical frequencies for mid. Permian

*ms-ds, magnetic susceptibility depth (time) series; xrf-ds, -ray fluorescence depth (time) series; arm-ds, anhysteretic remanent magnetization depth (time) series; mtm, multi-taper method; esa, evolutionary (power spectra) spectral analysis; eha, evolutive harmonic analysis (Meyers 2014); efft, evolutionary fast fourier transform analysis; ama, amplitude modulation envelope analysis (Hinnov 2000); TimeOpt (Meyers 2015).

Sources of uncertainty reflect: sar, sediment accumulation rate; fats, floating astronomical timescale.

L, long (405 kyr); S, short Ecc. (eccentricity); lpmo, long period modulation of obliquity; lpme, long period modulation of eccentricity; Ob, short obliquity; Pr, precession. Durations are provided only for statistically strong signals.

Conversely, the primary stratigraphic archive of glacioeustasy and source of many geological, geochemical, and palaeobiological records of environmental and climate perturbation during the LPIA come from cyclic shallower-water (platform top/shelf) marine (Fig. 4), mixed marine-terrestrial, and fully terrestrial low-latitude successions. Some of the GSSPs and most North American and western European regional stages have been defined in shallow-water and mixed marine-terrestrial successions (Aretz et al. 2020; Henderson and Shen 2020), many of which are cyclic. These features make cyclic shallow-water to terrestrial successions an important link to higher-latitude Gondwanan records of glaciation/deglaciation history.

Eccentricity pacing (short- or long) of low-latitude cyclothems and other contemporaneous stratigraphic cyclicity has long been hypothesized (e.g. Heckel 1986, 2013; Algeo and Wilkinson 1988; Boardman and Heckel 1989; Soreghan and Dickinson 1994; Davydov et al. 2010, 2012; Eros et al. 2012a; Schmitz and Davydov 2012; Montañez et al. 2016; Alnazghah and Kerans 2018; Fielding et al. 2020). High-precision U–Pb ages for several Carboniferous cyclic successions (Rasbury et al. 1998; Gastaldo et al. 2009; Eros et al. 2012a; Waters and Condon 2012; Pointon et al. 2014; Jirásek et al. 2018; Wu et al. 2020) further suggest that an astronomical signal is archived in Carboniferous and early Permian cyclic successions. To date, however, a signal of orbital forcing in these successions has been largely inferred by (1) dividing a duration of time constrained biostratigraphically and/or by radioisotope ages, by the number of observed cycles and/or (2) by applying an inferred astronomical frequency (typically eccentricity) to ‘tune’ a power spectra of stratigraphic cyclicity, but without further statistical analysis that has been applied to deeper-water successions (aforementioned references and Heckel 1986; Izart et al. 2003; Davydov et al. 2010, 2012; Schmitz and Davydov 2012; van den Belt et al. 2015; Alnazghah and Kerans 2018; Fielding et al. 2020). A few studies have more rigorously tested for signals of astronomical forcing in shallow-water and continental cyclic successions with results associated with considerable temporal uncertainty (Weedon and Read 1995; Eros 2011; Gebhardt and Hiete 2013).

The lack of robust validation of astronomical forcing of cyclothems reflects the challenges in signal processing to search for orbital cycles using shallow water to terrestrial stratigraphic records. That is, there are multiple Earth surface processes that can filter the orbitally forced climate signal in the stratal cyclicity and be potential sources of complexity (‘noise’) in testing for astronomical forcing. Potential sources include the imprint of non-orbitally influenced climate and depositional processes, autogenic variability, the influence of highly variable sedimentation rates, differential compaction between lithofacies, the unconstrained distribution and duration of hiatuses, and timescale uncertainties among other factors (Meyers 2019; Sinnesael et al. 2019). Additionally, aspects of the techniques used for astronomical testing can introduce bias during tuning using astronomical calibration targets and signal manipulation. These challenges can lead to substantial error in the derivative orbital timescales including the occurrence of false positives and negatives in the astronomical forcing signal and produce multiple orbital interpretations for a given stratigraphy (for excellent discussions of these issues see Hinnov 2013; Li et al. 2019; Meyers 2019; Sinnesael et al. 2019). New techniques (Table 1) have been developed recently to address many of these issues but have been applied primarily to cyclostratigraphic analysis of deeper-water and lacustrine records of the LPIA (Table 1; see section ‘Advancing a robust astronomical timescale for the LPIA interval’). The focus of the following discussion is on those Earth surface processes that operate in shallow-water marine to terrestrial environments and are sources of potential contamination and distortion of the primary forcing signal in cyclic stratigraphic records. These issues must be fully addressed for further advances in developing an astrochronometer for the LPIA interval.

Shallow-water and terrestrial depositional environments are characterized by highly variable sedimentation rates, which in turn influence the time series power spectrum and complicate the detection of astronomical signal frequencies. Two other sources that impact testing for astronomical signals in shallow water and terrestrial successions include (1) the non-uniqueness of facies–depth relationships, which make stratigraphic depth to time conversions challenging, and (2) the high potential for ‘missing beats’ of sediment deposition and stratigraphic cycles. In subsidence-controlled depositional environments, such as those in which cyclothems and terrestrial sediments are deposited, the interplay of sea-level fluctuations, subsidence, and sediment accumulation controls the thickness of the cycles and the stratigraphic position and duration of hiatuses (Fig. 8a) (Watney et al. 1989; Goldhammer et al. 1991; Read 1998; Purkis et al. 2015). When the rate of sea-level fall outpaces that of the subsidence of the platform/shelf, typically characterized by slow subsidence rates (cm ka−1), the sediment–water interface is exposed with the possibility of no sediment record of multiple beats of orbital band duration (Fig. 8a). The stratigraphic distribution of the hiatuses (whether erosional or non-depositional) within a cyclothemic (cyclic) succession and the absolute time archived within any given hiatus will vary with differences in, and the interaction between, platform slope, subsidence rates, sediment supply and production rates, and rates of sea-level change and thus, are not straightforwardly predictable (Fig. 8a). A substantial percentage (up to c. 85%) of the sea-level (orbital) cycle can be represented in surface erosion, karstification, and palaeosol development (Watney et al. 1989; Goldhammer et al. 1991; Read 1995, 1998; Barnett et al. 2002). It is these issues that underlie the debate over the suitability of the Mississippian–Pennsylvanian GSSP hosted in the shallow-water cyclic succession at Arrow Canyon, NV, USA (Barnett and Wright 2008; Bishop et al. 2009). These issues further explain the temporal and geographical differences in the dominant eccentricity signal (c. 100 v. 405 kyr) inferred from cyclothemic successions deposited in basins and on shelves of differing subsidence rates (Horbury 1989; Fielding et al. 2020; Smith and Read 2000; Wright and Vanstone 2001; Eros et al. 2012a; van den Belt et al. 2015).

Fig. 8.

Forward stratigraphic modelling of four Pennsylvanian cyclothems that formed in response to long-eccentricity modulated sea-level fluctuations (modified after Watney et al. 1989). (a) The stratigraphic succession was modelled using constraints estimated for the low-sloping (<0.1°) Midcontinent shelf: subsidence rate of 0.1 m ka−1, range of appropriate sedimentation rates for carbonates and shales, and sea-level rise and fall rates constrained by those of the Late Pleistocene. The resulting stratigraphic succession is shown horizontally at the top of the diagram. (b) Oxygen isotope record for the Late Pleistocene as a proxy of sea-level change. Comparison of (a) and (b) indicates that deposition of the transgressive and highstand facies in the Pennsylvanian cyclothems (colour vertical bars) would have occurred at points in the sea-level curve analogous to that of the last interglacial sensu stricto (5e) or MIS 5a–e (sensu lato) of the last Pleistocene glacial cycle. The remaining portion of the Pennsylvanian sea-level cycle, including the latter half of the glacial period (MIS 2–4), is primarily captured in the hiatuses (cross-hachured intervals) of the modelled stratigraphy. Horizontal lines indicate the position of sea-level at the peak interglacial (red) and the point at which sea-level falls below the platform (black; ‘sediment–water interface’) and the platform is exposed. Black vertical lines and numbers indicate the amount of sea-level change (m) between the two points. (c) Theoretical model of Carboniferous sea-level fluctuations over a 1.5 kyr period forced by five orbital cycles; modified after Barnett et al. (2002). Synthetic solar insolation curve assumes frequencies of astronomical cycles of 413 (long) and 112 kyr (short) for eccentricity, 34 kyr for obliquity, and 21 and 17 kyr for precession. More recent astrochronology studies reveal eccentricity frequencies more similar to astronomical solutions for the Cenozoic.

Fig. 8.

Forward stratigraphic modelling of four Pennsylvanian cyclothems that formed in response to long-eccentricity modulated sea-level fluctuations (modified after Watney et al. 1989). (a) The stratigraphic succession was modelled using constraints estimated for the low-sloping (<0.1°) Midcontinent shelf: subsidence rate of 0.1 m ka−1, range of appropriate sedimentation rates for carbonates and shales, and sea-level rise and fall rates constrained by those of the Late Pleistocene. The resulting stratigraphic succession is shown horizontally at the top of the diagram. (b) Oxygen isotope record for the Late Pleistocene as a proxy of sea-level change. Comparison of (a) and (b) indicates that deposition of the transgressive and highstand facies in the Pennsylvanian cyclothems (colour vertical bars) would have occurred at points in the sea-level curve analogous to that of the last interglacial sensu stricto (5e) or MIS 5a–e (sensu lato) of the last Pleistocene glacial cycle. The remaining portion of the Pennsylvanian sea-level cycle, including the latter half of the glacial period (MIS 2–4), is primarily captured in the hiatuses (cross-hachured intervals) of the modelled stratigraphy. Horizontal lines indicate the position of sea-level at the peak interglacial (red) and the point at which sea-level falls below the platform (black; ‘sediment–water interface’) and the platform is exposed. Black vertical lines and numbers indicate the amount of sea-level change (m) between the two points. (c) Theoretical model of Carboniferous sea-level fluctuations over a 1.5 kyr period forced by five orbital cycles; modified after Barnett et al. (2002). Synthetic solar insolation curve assumes frequencies of astronomical cycles of 413 (long) and 112 kyr (short) for eccentricity, 34 kyr for obliquity, and 21 and 17 kyr for precession. More recent astrochronology studies reveal eccentricity frequencies more similar to astronomical solutions for the Cenozoic.

Although robust evidence of astronomical forcing of deposition during the LPIA exists in deeper-water marine and lacustrine stratigraphic records, to date, only the potential for identifying astronomical signal frequencies in shallow-water marine, mixed, and terrestrial successions has been shown. The aforementioned sources of signal noise all limit the degree to which a robust astronomical timescale can be developed for cyclothemic successions using typically applied spectral and statistical approaches. This could have implications for the age model for four of the five Pennsylvanian stages in the radioisotopically calibrated composite standard of the Geologic Time Scale (2012: Davydov et al. 2012; propagated in the GTS 2016 (Ogg et al. 2016) and 2020 (Aretz et al. 2020) that was derived using an astronomically tuned (405 kyr) cyclothem-scaled stratigraphy (see Wu et al. 2019, fig. 3). Additional interrogation of Carboniferous and Permian cyclic successions will require a more comprehensive approach and application of novel and integrated approaches.

New objective statistical methodologies have been introduced to address the aforementioned challenges, among others, and to better constrain astrochronological interpretations and their uncertainties (see reviews of Hinnov 2018; Li et al. 2019; Meyers 2019). For example, techniques being used to accurately constraint sedimentation accumulation rates and to estimate their variability include the average spectral misfit (ASM; Meyers and Sageman 2007; Meyers 2008), a Bayesian Monte Carlo approach (Malinverno et al. 2010), TimeOpt (Meyers 2015), evolutive TimeOpt (Meyers 2019), and the use of the correlation coefficient between the power spectra of a data series and that of an associated astronomical forcing series (Acycle of Li et al. 2018; Li et al. 2019). Recently, a community-scale experiment, the Cyclostratigraphy Intercomparison Project (Sinnesael et al. 2019), has defined a general set of guidelines for future studies to work toward reducing uncertainty in astrochronology studies. A key recommendation is to compare different approaches, parameterizations, and signal manipulation and to optimize all information archived in the stratigraphic record.

Of particular relevance to shallow-water marine to terrestrial stratigraphic records, TimeOpt (Meyers 2019) uses sedimentation templates that permit integration of information on the slope of the platform/shelf and sedimentation accumulation rates, along with defined uncertainties. In this way, several of the processes that lead to missing beats and hiatuses in cyclothemic successions (Fig. 8a) and make such sedimentary systems challenging for testing for astronomical forcing can be considered. Further optimization of these approaches is likely needed to better represent the physics of the climate and depositional system and to include process-based null models specific to the environments in which cyclothems and other stratal cyclicity were deposited (Meyers 2019). These and further development of independent age constraints have the potential to rigorously test how the signal of astronomical forcing is imprinted on late Paleozoic shallow-marine to terrestrial stratigraphic record. An astronomical timescale defined for cyclothemic/cyclic successions, in turn, will reveal the orbital pacing of glacioeustatic fluctuations and other environmental and biological changes inferred from these successions and improve regional to global-scale correlations in the low latitudes. Developing astrochronologies for cyclothemic/cyclic successions from different regions will be a robust test of the astronomical timescales developed in deeper water successions and of the suitability of building age models for the Carboniferous timescale based on an astronomically tuned (405 kyr) cyclothem-scaled stratigraphy.

A current synthesis of the Late Paleozoic Ice Age (LPIA) highlights its longevity (Late Devonian Epoch through much of the Permian Period) and its dynamic glaciation history characterized by a series of 106-year-long glacials and intervening warmer periods of ice retraction. Continental ice expanded out of up to 30 ice centres distributed throughout mid- to high-latitude Gondwana with possible trans-continental ice sheets during the latest Carboniferous–earliest Permian apex of glaciation. Evidence for stepwise buildup of continental ice during the Middle Mississippian onset of the main phase of the icehouse is analogous to the cycles of expansion and retraction of the Antarctic Ice Sheet heralding, by millions of years, the onset of the Cenozoic icehouse (34 Ma). The Sakmarian demise of the penultimate icehouse underwent reciprocal stepwise loss of ice synchronous with rising atmospheric CO2 concentrations to > 400 to 500 ppm, with clear implications for our future.

Multiple lines of evidence document repeated environmental and ecosystem perturbations during the LPIA, many in apparent step with fluctuations in palaeo-CO2 within the magnitude of anticipated anthropogenically driven changes (200 to 500 + ppm). During the LPIA, sea-level fluctuated at a hierarchy of timescales with magnitudes of glacioeustatic fluctuations (tens of metres to ≥100 m) broadly varying with continental ice extent. Late Paleozoic biogenic and sedimentary δ13C and δ15N records indicate repeated perturbation of global carbon and nitrogen cycling contemporaneous with changes in ice extent and sea-level. Major faunal and floral reorganizations and changes in biodiversity also occurred at both the onset and the demise of the icehouse, as well as during some of the major climate events within it. Additional high-precision chronological constraints are needed to define the phasing between inferred changes in high-latitude ice volume, warmings and coolings, durations and magnitudes of sea-level fluctuations, perturbations to C, N, and hydrological cycling, and ecosystem disruptions. Currently chronostratigraphic constraints preclude establishing these relationships with confidence. Integration of Earth System models that contextualize proxy records are needed to identify the nature of mechanistic linkages and teleconnections in the late Paleozoic climate system.

The glaciation history is clearly imprinted on the Euramerican Carboniferous and Permian stratigraphic record expressed as temporal variations in stratigraphic architectures, stratal stacking patterns, and the nature of bounding surfaces between stratal units. The transition from greenhouse to icehouse depositional systems occurs in the late Middle Mississippian (c. 335 to 332 Ma) archived as regional-scale facies shifts in carbonated-dominated systems and siliciclastic fluvial styles, and by a shift in the frequency of the predominant cyclicity from 3rd- (106-year) to 4th-order (105-year) sequences. This shift, which persists from the Late Mississippian through early Permian, occurred synchronous with an increase in the magnitude of the 4th-order sea-level fluctuations. Similarly, a change in stratal architectures, including the loss of cyclicity post-mid-Sakmarian captures the diminishing ice volumes and a substantial drop in the magnitudes of glacioeustasy with the onset of the demise of the LPIA.

Euramerican cyclothems in Pennsylvanian and lower Permian successions archive relative sea-level cycles long inferred to be glacioeustatic. In turn, magnitudes of sea-level fluctuations and relative changes in ice volume have been inferred from cyclothemic successions, exhibiting variability in estimated magnitudes of glacioeustasy that appear to be in-step with independently inferred intervals of ice expansion and retraction. Thus, given the strong imprint of climate changes on low-latitude strata, it has been argued that low-latitude successions provide the most complete record of the LPIA glaciation history. Temporal trends in stratal stacking patterns, when integrated with biostratigraphic constraints, have been used to propose intra- and inter-continental correlations and may offer higher resolution chronostratigraphic constraints for cyclothemic successions than typically afforded by other chronostratigraphic tools. Similarly, further age constraints are needed to test all of these aforementioned relationships.

The late Paleozoic is considered a frontier for developing robust astronomical timescales given the Earth system parallels between the LPIA and the orbitally paced Cenozoic icehouse. Astronomically calibrating stratigraphic records permits deciphering the distribution of time in a succession and places constraints on the rates of processes archived in sedimentary records. If radioisotopically anchored, an astronomical timescale permits assignment of ages to geological and biological events and to chronostratigraphic boundaries, and ultimately refinement of the late Paleozoic geological timescale. Statistically robust evidence of astronomical forcing of sedimentary deposition has been documented in several late Paleozoic open- and/or quiet-water successions, reflecting their high potential for yielding near-continuous depositional archives characterized by relatively low variability in sediment accumulation rates. Although eccentricity pacing of low-latitude cyclothems and other contemporaneous cyclicity has long been hypothesized, a signal of orbital forcing in these successions has been largely inferred with no statistically rigorous validation. This reflects the challenges in testing for astronomical forcing in shallow-water to terrestrial successions given the many sources of stratigraphic ‘noise’ that can filter the orbital signal. New techniques have been developed that can address these challenges and better constrain astrochronological interpretations and their uncertainties. Future interrogation of Carboniferous and Permian shallow-water and terrestrial successions for signals of astronomical forcing will need to use an integrated approach that combines new techniques with further optimization and independent age constraints. The emerging field of ‘crypto-tephralogy’, which seeks to define depositional ages from the youngest population of detrital zircons may provide new opportunities.

S. Lucas and J.W. Schneider are thanked for encouraging me to prepare this manuscript and for the editorial handling. The helpful comments of two anonymous reviewers are also appreciated. Bradley Cramer is thanked for graciously sharing a compilation of sedimentary carbonate δ13C presented in Figure 3 and for engaging discussions. Jitao Chen is thanked for his efforts in developing the CO2 compilation presented in Figures 1 and 3. J. Fred Read graciously provided forward models that were used in Figure 4.

IPM: conceptualization (lead), funding acquisition (lead), investigation (lead), methodology (lead), visualization (lead), writing – original draft (lead), writing – review & editing (lead).

Funding for this research came from the Sedimentary Geology and Paleobiology Program, NSF, Earth Science Division (awards EAR1338281 and EAR1729882).

Data sharing is not applicable to this article as no datasets were generated or analysed during the current study.

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Figures & Tables

Fig. 1.

Late Devonian through Permian trends in occurrence and distribution of glaciogenic records, inferred magnitudes of glacioeustasy, and palaeo-CO2 plotted on the Geologic Time Scale 2020 (Gradstein et al. 2020). (a) Number of occurrences (extent), by stage, of glaciogenic deposits in mid- to high-latitude Gondwanan basins compiled by Soreghan et al. (2019). (b) Geographical distribution of glaciogenic deposits and geomorphic features through time; see section ‘Synopsis of ice ages of the late Paleozoic’ for references. Hachured pattern in the ‘Arabian Peninsula and Madagascar’ record represents multiple hiatuses, of unknown age, that are interpreted as glacial advances/cycles. Blue horizontal lines show the temporal position of high-precision CA (ID)-TIMS U–Pb dates for glaciogenic deposits (Gulbranson et al. 2010; Metcalfe et al. 2015; Griffis et al. 2018, 2019, 2021; Phillips et al. 2018; Backhouse and Mory 2020). SIMS (secondary ion mass spectrometry), SHRIMP (sensitive high-resolution ion microprobe), or LA MC-ICPMS zircon ages are not shown given large uncertainties associated with ages. (c) Magnitudes of high-frequency glacioeustasy inferred from different sedimentary archives; modified after Rygel et al. (2008); (d) Palaeo-CO2 estimates and LOESS trendline (30%) with 12 and 86% confidence intervals, courtesy of J. Chen. Palaeosol carbonate- and stomatal-based CO2 estimates of Montañez et al. (2016) and Richey et al. (2020) are shown as solid symbols and pass the screening criteria to be acceptable as initially published (https://palaeo-co2.org). CO2 estimates compiled by Foster et al. (2017) and shown as open symbols are quarantined until further modernization. Vertical grey lines are uncertainties on age estimates of individual CO2 values.

Fig. 1.

Late Devonian through Permian trends in occurrence and distribution of glaciogenic records, inferred magnitudes of glacioeustasy, and palaeo-CO2 plotted on the Geologic Time Scale 2020 (Gradstein et al. 2020). (a) Number of occurrences (extent), by stage, of glaciogenic deposits in mid- to high-latitude Gondwanan basins compiled by Soreghan et al. (2019). (b) Geographical distribution of glaciogenic deposits and geomorphic features through time; see section ‘Synopsis of ice ages of the late Paleozoic’ for references. Hachured pattern in the ‘Arabian Peninsula and Madagascar’ record represents multiple hiatuses, of unknown age, that are interpreted as glacial advances/cycles. Blue horizontal lines show the temporal position of high-precision CA (ID)-TIMS U–Pb dates for glaciogenic deposits (Gulbranson et al. 2010; Metcalfe et al. 2015; Griffis et al. 2018, 2019, 2021; Phillips et al. 2018; Backhouse and Mory 2020). SIMS (secondary ion mass spectrometry), SHRIMP (sensitive high-resolution ion microprobe), or LA MC-ICPMS zircon ages are not shown given large uncertainties associated with ages. (c) Magnitudes of high-frequency glacioeustasy inferred from different sedimentary archives; modified after Rygel et al. (2008); (d) Palaeo-CO2 estimates and LOESS trendline (30%) with 12 and 86% confidence intervals, courtesy of J. Chen. Palaeosol carbonate- and stomatal-based CO2 estimates of Montañez et al. (2016) and Richey et al. (2020) are shown as solid symbols and pass the screening criteria to be acceptable as initially published (https://palaeo-co2.org). CO2 estimates compiled by Foster et al. (2017) and shown as open symbols are quarantined until further modernization. Vertical grey lines are uncertainties on age estimates of individual CO2 values.

Fig. 2.

Palaeogeographical distribution of glaciated basins in Gondwana (light orange shading) and select unglaciated, low-latitude depositional sites (dark orange shading) during the late Devonian, Carboniferous, and Permian periods. Base map and locations of palaeo-ice centres (blue) and ice-flow lines modified after Craddock et al. (2019). Position of Gondwanan glaciated basins from Gulbranson et al. (2010), Wopfner and Jin (2009), Isbell et al. (2012), Bussert (2014), Fan et al. (2015), Griffis et al. (2019), Caputo and Santos (2020), Buso et al. (2020) and Zurli et al. (2021). Dashed lines for Cimmerian blocks (N. Qiangtang–Qamdom, S. Qiangtang–Baoshan, Tethys Himalaya and Lhasa blocks) indicate uncertainty in distribution of glaciomarine diamictites (Fan et al. 2015). Low-latitude basins from Heckel (2013) and Blakey (2008); Midcontinent includes N. Shelf, Arkoma, Anadarko, and Midland basins and eastern shelf of the Midland Basin.

Fig. 2.

Palaeogeographical distribution of glaciated basins in Gondwana (light orange shading) and select unglaciated, low-latitude depositional sites (dark orange shading) during the late Devonian, Carboniferous, and Permian periods. Base map and locations of palaeo-ice centres (blue) and ice-flow lines modified after Craddock et al. (2019). Position of Gondwanan glaciated basins from Gulbranson et al. (2010), Wopfner and Jin (2009), Isbell et al. (2012), Bussert (2014), Fan et al. (2015), Griffis et al. (2019), Caputo and Santos (2020), Buso et al. (2020) and Zurli et al. (2021). Dashed lines for Cimmerian blocks (N. Qiangtang–Qamdom, S. Qiangtang–Baoshan, Tethys Himalaya and Lhasa blocks) indicate uncertainty in distribution of glaciomarine diamictites (Fan et al. 2015). Low-latitude basins from Heckel (2013) and Blakey (2008); Midcontinent includes N. Shelf, Arkoma, Anadarko, and Midland basins and eastern shelf of the Midland Basin.

Fig. 3.

Late Devonian through Permian trends in occurrence of glaciogenic records, palaeo-CO2, and geochemical proxy records of environmental change plotted on the Geologic Time Scale 2020 (Gradstein et al. 2020). (a) Number of occurrences (extent), by stage, of glaciogenic deposits as in Figure 1a. (b) Palaeo-CO2 estimates and LOESS trendline as in Figure 1d. (c) Trends in brachiopod (black lines) and conodont (blue and grey lines) δ18O values from Grossman and Joachimski (2020). The brachiopod (black) and conodont (grey) δ18O records were compiled using primarily records from palaeo-epicontinental seas; the blue conodont δ18O record is from an open-water slope succession in South China (B. Chen et al. 2016). Decoupling of trends and increased variability in conodont δ18O from epicontinental sea successions have been interpreted as recording variability in surface water salinity and seawater δ18O (Joachimski and Lambert 2015; Montañez et al. 2018). (d) Trend in sedimentary carbonate δ13C compiled by and courtesy of B. Cramer (see Cramer and Jarvis (2020) for details and data sources. KOBE is Kinderhookian–Osagean Boundary Excursion; TICE is the Tournasian Isotope Carbon Excursion. (e) LOESS trend in Carboniferous and Permian seawater 87Sr/86Sr shown with 2.5 and 97.5% confidence intervals; modified after from McArthur et al. (2020); for the Carboniferous: Chen et al. (2021); for the Permian: Wang et al. (2021). Frasnian 87Sr/86Sr trend from McArthur et al. (2020); Famennian not shown given sizeable disparity between datasets.

Fig. 3.

Late Devonian through Permian trends in occurrence of glaciogenic records, palaeo-CO2, and geochemical proxy records of environmental change plotted on the Geologic Time Scale 2020 (Gradstein et al. 2020). (a) Number of occurrences (extent), by stage, of glaciogenic deposits as in Figure 1a. (b) Palaeo-CO2 estimates and LOESS trendline as in Figure 1d. (c) Trends in brachiopod (black lines) and conodont (blue and grey lines) δ18O values from Grossman and Joachimski (2020). The brachiopod (black) and conodont (grey) δ18O records were compiled using primarily records from palaeo-epicontinental seas; the blue conodont δ18O record is from an open-water slope succession in South China (B. Chen et al. 2016). Decoupling of trends and increased variability in conodont δ18O from epicontinental sea successions have been interpreted as recording variability in surface water salinity and seawater δ18O (Joachimski and Lambert 2015; Montañez et al. 2018). (d) Trend in sedimentary carbonate δ13C compiled by and courtesy of B. Cramer (see Cramer and Jarvis (2020) for details and data sources. KOBE is Kinderhookian–Osagean Boundary Excursion; TICE is the Tournasian Isotope Carbon Excursion. (e) LOESS trend in Carboniferous and Permian seawater 87Sr/86Sr shown with 2.5 and 97.5% confidence intervals; modified after from McArthur et al. (2020); for the Carboniferous: Chen et al. (2021); for the Permian: Wang et al. (2021). Frasnian 87Sr/86Sr trend from McArthur et al. (2020); Famennian not shown given sizeable disparity between datasets.

Fig. 4.

Synthetic platform stratigraphy and idealized stratigraphic architectures. Forward-modelled stratigraphies (a–c) represent a 400 kyr duration and were generated using a hierarchy of 400, 100, 40 and 20 kyr sea-level changes (inset); modified after Read (1998). (a) Greenhouse carbonate-dominated stratigraphy for sea-level fluctuations of 400 kyr: 1 m; 100 kyr: 2 m; 40 kyr: 2 m; 20 kyr: 5 m showing layer-cake stratigraphy and stacking of 5th-order parasequences composed of shallow (orange) to deep subtidal (blue) facies capped by tidal-flat carbonates or fluvial to shoreface siliciclastics (yellow). Mud-mounds (within the deeper subtidal facies) form during transgressions. (b) Generalized icehouse stratigraphy showing characteristics of stratigraphic architectures that develop with large-scale changes in accommodation space and restricted progradation of facies: deep-fluvial incision, karstification and palaeosol development, widespread disconformities (red lines), downdip imbrication/shingling of prograding (regressive) parasequences, and complex lateral relationships. (c) Icehouse stratigraphy for 400 kyr-dominated eustasy (inset) and sea-level fluctuations of 400 kyr: 20 m; 100 kyr: 70 m; 41 kyr: 5 m; 19 kyr: 1 m modelled to be analogous to Pennsylvanian cyclothems. Synthetic stratigraphy shows a 4th-order (400 kyr) disconformity bounded sequence within which four disconformity bounded 100 kyr parasequences are nested. Note well-developed ramp-margin (lowstand) wedge and thin backstepping TST with maximum flooding surface (mfs) (darker blue facies) extending far back onto the platform. Shingled parasequences of the highstand and forced regressive systems tracts are composed of shallow subtidal lithofacies and downlap onto the mfs, producing highly variable thickness and composition patterns of cycles by location. Black lines updip are exposure/erosion surfaces. (d) Idealized stratigraphic architecture for a 4th-order Pennsylvanian sequence deposited during the main phase of the LPIA icehouse (modified after Eriksson et al. 2019). Note that coals predate transgressive shales and can occur directly overlying incised valley fills. The diagram is drawn parallel to depositional strike.

Fig. 4.

Synthetic platform stratigraphy and idealized stratigraphic architectures. Forward-modelled stratigraphies (a–c) represent a 400 kyr duration and were generated using a hierarchy of 400, 100, 40 and 20 kyr sea-level changes (inset); modified after Read (1998). (a) Greenhouse carbonate-dominated stratigraphy for sea-level fluctuations of 400 kyr: 1 m; 100 kyr: 2 m; 40 kyr: 2 m; 20 kyr: 5 m showing layer-cake stratigraphy and stacking of 5th-order parasequences composed of shallow (orange) to deep subtidal (blue) facies capped by tidal-flat carbonates or fluvial to shoreface siliciclastics (yellow). Mud-mounds (within the deeper subtidal facies) form during transgressions. (b) Generalized icehouse stratigraphy showing characteristics of stratigraphic architectures that develop with large-scale changes in accommodation space and restricted progradation of facies: deep-fluvial incision, karstification and palaeosol development, widespread disconformities (red lines), downdip imbrication/shingling of prograding (regressive) parasequences, and complex lateral relationships. (c) Icehouse stratigraphy for 400 kyr-dominated eustasy (inset) and sea-level fluctuations of 400 kyr: 20 m; 100 kyr: 70 m; 41 kyr: 5 m; 19 kyr: 1 m modelled to be analogous to Pennsylvanian cyclothems. Synthetic stratigraphy shows a 4th-order (400 kyr) disconformity bounded sequence within which four disconformity bounded 100 kyr parasequences are nested. Note well-developed ramp-margin (lowstand) wedge and thin backstepping TST with maximum flooding surface (mfs) (darker blue facies) extending far back onto the platform. Shingled parasequences of the highstand and forced regressive systems tracts are composed of shallow subtidal lithofacies and downlap onto the mfs, producing highly variable thickness and composition patterns of cycles by location. Black lines updip are exposure/erosion surfaces. (d) Idealized stratigraphic architecture for a 4th-order Pennsylvanian sequence deposited during the main phase of the LPIA icehouse (modified after Eriksson et al. 2019). Note that coals predate transgressive shales and can occur directly overlying incised valley fills. The diagram is drawn parallel to depositional strike.

Fig. 5.

Select cyclothems from the Illinois and Donets basins. (a) Stacked Middle Pennsylvanian (Desmoinesian) cyclothems of the Illinois Basin (equivalents of the Midcontinent Pawnee and Altamont ‘major’ cyclothems), modified after Nelson et al. (2011): #1 = deltaic siliciclastics of highstand systems tract (HST) and falling stage systems tract (FSST) into which incises valleys erode; #2 = FSST (regressive) of late highstand to lowstand systems tract (LST); point of maximum rate of sea-level fall denotes beginning of the LST. Incised valley and channel formation with initial fill of fluvial deposits; palaeosol formation on interfluves; #3 = coals of late LST that formed as peats once the rate of subsidence (accommodation space) outpaced the rate of sea-level fall; regional coals (Jamestown, Baker, and Cottage) are not laterally correlatable; #4 = black shales (some phosphatic in lower layers); black shales in some cyclothems (e.g. Energy and Anna shales of Pawnee) are underlain by wedge-shaped deposits of heterolithic grey siltstones and mudstones (estuarine), which exhibit tidal rhythmicity and thicken into channels. Transgressive surfaces (ts) denote onset of transgressive systems tract (TST); truncation of underlying facies and burial of in-situ lycopsids found in living position and rooted in underlying coals, record eustatically driven rapid rise in accommodation space. #5 = open marine muddy to grainy marine limestones; #6 = FSST – same as #1. (b) Pennsylvanian cyclothem types in the Donets Basin succession modified after Eros et al. (2012a). Up-dip and down-dip positions of correlated retrogradational, aggradational, and progradational cyclothems illustrate cross-ramp (hundreds of kilometres) variations in the facies assemblages (FA).

Fig. 5.

Select cyclothems from the Illinois and Donets basins. (a) Stacked Middle Pennsylvanian (Desmoinesian) cyclothems of the Illinois Basin (equivalents of the Midcontinent Pawnee and Altamont ‘major’ cyclothems), modified after Nelson et al. (2011): #1 = deltaic siliciclastics of highstand systems tract (HST) and falling stage systems tract (FSST) into which incises valleys erode; #2 = FSST (regressive) of late highstand to lowstand systems tract (LST); point of maximum rate of sea-level fall denotes beginning of the LST. Incised valley and channel formation with initial fill of fluvial deposits; palaeosol formation on interfluves; #3 = coals of late LST that formed as peats once the rate of subsidence (accommodation space) outpaced the rate of sea-level fall; regional coals (Jamestown, Baker, and Cottage) are not laterally correlatable; #4 = black shales (some phosphatic in lower layers); black shales in some cyclothems (e.g. Energy and Anna shales of Pawnee) are underlain by wedge-shaped deposits of heterolithic grey siltstones and mudstones (estuarine), which exhibit tidal rhythmicity and thicken into channels. Transgressive surfaces (ts) denote onset of transgressive systems tract (TST); truncation of underlying facies and burial of in-situ lycopsids found in living position and rooted in underlying coals, record eustatically driven rapid rise in accommodation space. #5 = open marine muddy to grainy marine limestones; #6 = FSST – same as #1. (b) Pennsylvanian cyclothem types in the Donets Basin succession modified after Eros et al. (2012a). Up-dip and down-dip positions of correlated retrogradational, aggradational, and progradational cyclothems illustrate cross-ramp (hundreds of kilometres) variations in the facies assemblages (FA).

Fig. 6.

Hierarchy of Donets sequences and the proposed equivalence to Midcontinent cyclothems, modified after Eros et al. (2012a). Donets sequences (e.g. Gz-4) shown with predominant limestone marker bed (denoted by red letters and numerals) on left. Composite and longer-term composite sequences shown in blue on the right. U.S. Midcontinent cyclothems shown as ‘major’, ‘intermediate’, and ‘minor’ cyclothems as defined by Heckel (1994); see text for further details. Grey dashed tie-lines indicate Heckel et al.'s (2007) ‘digital correlation’ of Donets limestone marker beds to the Midcontinent cyclothems. Superscript ‘+’ indicates the presence of additional minor cyclothems in the Midcontinent record and ‘*’ denotes less confident inter-continental correlations.

Fig. 6.

Hierarchy of Donets sequences and the proposed equivalence to Midcontinent cyclothems, modified after Eros et al. (2012a). Donets sequences (e.g. Gz-4) shown with predominant limestone marker bed (denoted by red letters and numerals) on left. Composite and longer-term composite sequences shown in blue on the right. U.S. Midcontinent cyclothems shown as ‘major’, ‘intermediate’, and ‘minor’ cyclothems as defined by Heckel (1994); see text for further details. Grey dashed tie-lines indicate Heckel et al.'s (2007) ‘digital correlation’ of Donets limestone marker beds to the Midcontinent cyclothems. Superscript ‘+’ indicates the presence of additional minor cyclothems in the Midcontinent record and ‘*’ denotes less confident inter-continental correlations.

Fig. 7.

Generalized stratigraphy for the Donets cyclothemic succession, hierarchy of Donets sequences, and inferred relative sea-level (onlap–offlap) curve modified after Eros et al. (2012a, b) and adjusted to the Geologic Time Scale 2020 (Gradstein et al. 2020). Cyclothem types as presented in Figure 5b. Correlation of limestone marker beds (red) relative to sequences and stacking patterns of cyclothem types indicated by the red dashed lines. Position of Donets marker beds and Midcontinent cyclothems on relative sea-level curve indicated by red dashed lines and blue and white alternating bands, respectfully. Offset from the horizontal between the two sets of Donets marker beds is a function of nonlinear rescaling of depth–time assignments for the onlap–offlap curve (and thus the position of the marker beds) to accommodate proposed (Schmitz and Davydov 2012) scaling of cyclothem (Donets and Midcontinent) stratigraphic distribution with high-precision U–Pb ages and stage boundaries. For all columns to the right of the relative sea-level curve, the distribution of Donets cyclothem types, sequences, and limestone marker beds (red) within individual global stages were rescaled linearly to address changes between the 2012 and 2020 Carboniferous timescales. The vertical break between the two parts of the diagram denotes these differences.

Fig. 7.

Generalized stratigraphy for the Donets cyclothemic succession, hierarchy of Donets sequences, and inferred relative sea-level (onlap–offlap) curve modified after Eros et al. (2012a, b) and adjusted to the Geologic Time Scale 2020 (Gradstein et al. 2020). Cyclothem types as presented in Figure 5b. Correlation of limestone marker beds (red) relative to sequences and stacking patterns of cyclothem types indicated by the red dashed lines. Position of Donets marker beds and Midcontinent cyclothems on relative sea-level curve indicated by red dashed lines and blue and white alternating bands, respectfully. Offset from the horizontal between the two sets of Donets marker beds is a function of nonlinear rescaling of depth–time assignments for the onlap–offlap curve (and thus the position of the marker beds) to accommodate proposed (Schmitz and Davydov 2012) scaling of cyclothem (Donets and Midcontinent) stratigraphic distribution with high-precision U–Pb ages and stage boundaries. For all columns to the right of the relative sea-level curve, the distribution of Donets cyclothem types, sequences, and limestone marker beds (red) within individual global stages were rescaled linearly to address changes between the 2012 and 2020 Carboniferous timescales. The vertical break between the two parts of the diagram denotes these differences.

Fig. 8.

Forward stratigraphic modelling of four Pennsylvanian cyclothems that formed in response to long-eccentricity modulated sea-level fluctuations (modified after Watney et al. 1989). (a) The stratigraphic succession was modelled using constraints estimated for the low-sloping (<0.1°) Midcontinent shelf: subsidence rate of 0.1 m ka−1, range of appropriate sedimentation rates for carbonates and shales, and sea-level rise and fall rates constrained by those of the Late Pleistocene. The resulting stratigraphic succession is shown horizontally at the top of the diagram. (b) Oxygen isotope record for the Late Pleistocene as a proxy of sea-level change. Comparison of (a) and (b) indicates that deposition of the transgressive and highstand facies in the Pennsylvanian cyclothems (colour vertical bars) would have occurred at points in the sea-level curve analogous to that of the last interglacial sensu stricto (5e) or MIS 5a–e (sensu lato) of the last Pleistocene glacial cycle. The remaining portion of the Pennsylvanian sea-level cycle, including the latter half of the glacial period (MIS 2–4), is primarily captured in the hiatuses (cross-hachured intervals) of the modelled stratigraphy. Horizontal lines indicate the position of sea-level at the peak interglacial (red) and the point at which sea-level falls below the platform (black; ‘sediment–water interface’) and the platform is exposed. Black vertical lines and numbers indicate the amount of sea-level change (m) between the two points. (c) Theoretical model of Carboniferous sea-level fluctuations over a 1.5 kyr period forced by five orbital cycles; modified after Barnett et al. (2002). Synthetic solar insolation curve assumes frequencies of astronomical cycles of 413 (long) and 112 kyr (short) for eccentricity, 34 kyr for obliquity, and 21 and 17 kyr for precession. More recent astrochronology studies reveal eccentricity frequencies more similar to astronomical solutions for the Cenozoic.

Fig. 8.

Forward stratigraphic modelling of four Pennsylvanian cyclothems that formed in response to long-eccentricity modulated sea-level fluctuations (modified after Watney et al. 1989). (a) The stratigraphic succession was modelled using constraints estimated for the low-sloping (<0.1°) Midcontinent shelf: subsidence rate of 0.1 m ka−1, range of appropriate sedimentation rates for carbonates and shales, and sea-level rise and fall rates constrained by those of the Late Pleistocene. The resulting stratigraphic succession is shown horizontally at the top of the diagram. (b) Oxygen isotope record for the Late Pleistocene as a proxy of sea-level change. Comparison of (a) and (b) indicates that deposition of the transgressive and highstand facies in the Pennsylvanian cyclothems (colour vertical bars) would have occurred at points in the sea-level curve analogous to that of the last interglacial sensu stricto (5e) or MIS 5a–e (sensu lato) of the last Pleistocene glacial cycle. The remaining portion of the Pennsylvanian sea-level cycle, including the latter half of the glacial period (MIS 2–4), is primarily captured in the hiatuses (cross-hachured intervals) of the modelled stratigraphy. Horizontal lines indicate the position of sea-level at the peak interglacial (red) and the point at which sea-level falls below the platform (black; ‘sediment–water interface’) and the platform is exposed. Black vertical lines and numbers indicate the amount of sea-level change (m) between the two points. (c) Theoretical model of Carboniferous sea-level fluctuations over a 1.5 kyr period forced by five orbital cycles; modified after Barnett et al. (2002). Synthetic solar insolation curve assumes frequencies of astronomical cycles of 413 (long) and 112 kyr (short) for eccentricity, 34 kyr for obliquity, and 21 and 17 kyr for precession. More recent astrochronology studies reveal eccentricity frequencies more similar to astronomical solutions for the Cenozoic.

Table 1.

Summary of astrochronological studies of late Paleozoic deeper- and quiet-water deposits

Study*MethodsLimitationsOrbital Signal(s)Timescale application
Late DevonianDe Vleeschouwer et al. (2012), (2013) ms-ds; mtmVariable sar; astronomical calibration target (405 kyr); fatsl & S Ecc (c. 100 kyr), lpme, Pr (c. 18 kyr)Constrain duration of Frasnian Stage (6.5 ± 0.4 myr); Giv.–Fras. bdry (383.6 ± 3.0 Ma); Fras.–Fam. bdry (376.7 ± 3.0 Ma)
Pas et al. (2018) ms-ds; mtm, efft esaAstro. calibration target (405 & c. 100 kyr & Ob cycle); high-prec. U–Pb anchored fatsL & S Ecc, lpmo (1.2 myr), ObConstrain Fammenian Stage (13.5 ± 0.5 myr); Fras.–Fam. Bdry (372.4 ± 0.9 Ma)
MississippianFang et al. (2018a, b)ms-ds; mtm, ama & esaAstro. calibration target (405 kyr); fatsL & S Ecc (128 & 95 kyr)., lpmo (1.1 myr), lpme, Ob (33 to 34 kyr); Pr (19 & 15.9 kyr)Constrains duration of Serpukhovian (7.68 ± 0.15 myr)
PennsylvanianWu et al. (2019) ms-ds; mtm, efft amaAstro. calibration target (405 kyr); U–Pb anchored fatsL & S Ecc (136, 122 & 96 kyr)., lpme (2.4, 1.6 & 1.2 myr), lpmo (1.2 myr), lpme, Ob (31 kyr); Pr (22.9 & 19.7 kyr)Constrains duration of Serpukhovian (7.6 myr), Bashkirian (8.1 myr), Moscovian (8.5 myr), Kasimovian (2.87 myr) and Gzhelian (4.83 myr) stages
early PermianKochhann et al. (2020) xrf-ds; mtm, efft & TimeOptRatios of spectral Periodicities tuned to period ratios of astro. calibration targets (132, 124, 98, 96, 35.5, 21.7, 20.6, 17.8, 17.7 kyr);S Ecc, Pr, hemiPr
mid. PermianFang et al. (2015) arm-ds; mtm, esa & amaVariable sar; astro. calibration target (405 kyr); fatsL & S Ecc (c. 95 kyr)., lpmo (c. 1 myr), lpme (c. 2 myr), Ob (c. 44 & 33 kyr), Pr (c. 20 kyr)Constrain duration of Roadian Stage (3.7 ± 0.4 myr); Wordian Stage (2.9 ± 0.4 myr)
X. Yao and Hinnov (2019) (revised Yao et al. 2015)Midpoint-triangle interpolation ds; mtm,Binary lithologic series; variable sar; U–Pb (n = 2) anchored fatsOb (26, 34, 35 & 44 kyr); Pr (21.5, 19.5 & 17.7 kyr); possible L Ecc modulationReveals new Pr and Ob alternations; constrain ages and duration of radiolarian biozones
Huang et al. (2021) ngr-ds; mtm, eha & TimeOpt;Astro. calibration target (405 kyr); fatsL & S Ecc (96.2 to 129.9 kyr)., lpmo (1.2 myr), Ob (40.7, 30 & 25 kyr); Pr (23 & 21–18 kyr)Refines astronomical frequencies for mid. Permian
Study*MethodsLimitationsOrbital Signal(s)Timescale application
Late DevonianDe Vleeschouwer et al. (2012), (2013) ms-ds; mtmVariable sar; astronomical calibration target (405 kyr); fatsl & S Ecc (c. 100 kyr), lpme, Pr (c. 18 kyr)Constrain duration of Frasnian Stage (6.5 ± 0.4 myr); Giv.–Fras. bdry (383.6 ± 3.0 Ma); Fras.–Fam. bdry (376.7 ± 3.0 Ma)
Pas et al. (2018) ms-ds; mtm, efft esaAstro. calibration target (405 & c. 100 kyr & Ob cycle); high-prec. U–Pb anchored fatsL & S Ecc, lpmo (1.2 myr), ObConstrain Fammenian Stage (13.5 ± 0.5 myr); Fras.–Fam. Bdry (372.4 ± 0.9 Ma)
MississippianFang et al. (2018a, b)ms-ds; mtm, ama & esaAstro. calibration target (405 kyr); fatsL & S Ecc (128 & 95 kyr)., lpmo (1.1 myr), lpme, Ob (33 to 34 kyr); Pr (19 & 15.9 kyr)Constrains duration of Serpukhovian (7.68 ± 0.15 myr)
PennsylvanianWu et al. (2019) ms-ds; mtm, efft amaAstro. calibration target (405 kyr); U–Pb anchored fatsL & S Ecc (136, 122 & 96 kyr)., lpme (2.4, 1.6 & 1.2 myr), lpmo (1.2 myr), lpme, Ob (31 kyr); Pr (22.9 & 19.7 kyr)Constrains duration of Serpukhovian (7.6 myr), Bashkirian (8.1 myr), Moscovian (8.5 myr), Kasimovian (2.87 myr) and Gzhelian (4.83 myr) stages
early PermianKochhann et al. (2020) xrf-ds; mtm, efft & TimeOptRatios of spectral Periodicities tuned to period ratios of astro. calibration targets (132, 124, 98, 96, 35.5, 21.7, 20.6, 17.8, 17.7 kyr);S Ecc, Pr, hemiPr
mid. PermianFang et al. (2015) arm-ds; mtm, esa & amaVariable sar; astro. calibration target (405 kyr); fatsL & S Ecc (c. 95 kyr)., lpmo (c. 1 myr), lpme (c. 2 myr), Ob (c. 44 & 33 kyr), Pr (c. 20 kyr)Constrain duration of Roadian Stage (3.7 ± 0.4 myr); Wordian Stage (2.9 ± 0.4 myr)
X. Yao and Hinnov (2019) (revised Yao et al. 2015)Midpoint-triangle interpolation ds; mtm,Binary lithologic series; variable sar; U–Pb (n = 2) anchored fatsOb (26, 34, 35 & 44 kyr); Pr (21.5, 19.5 & 17.7 kyr); possible L Ecc modulationReveals new Pr and Ob alternations; constrain ages and duration of radiolarian biozones
Huang et al. (2021) ngr-ds; mtm, eha & TimeOpt;Astro. calibration target (405 kyr); fatsL & S Ecc (96.2 to 129.9 kyr)., lpmo (1.2 myr), Ob (40.7, 30 & 25 kyr); Pr (23 & 21–18 kyr)Refines astronomical frequencies for mid. Permian

*ms-ds, magnetic susceptibility depth (time) series; xrf-ds, -ray fluorescence depth (time) series; arm-ds, anhysteretic remanent magnetization depth (time) series; mtm, multi-taper method; esa, evolutionary (power spectra) spectral analysis; eha, evolutive harmonic analysis (Meyers 2014); efft, evolutionary fast fourier transform analysis; ama, amplitude modulation envelope analysis (Hinnov 2000); TimeOpt (Meyers 2015).

Sources of uncertainty reflect: sar, sediment accumulation rate; fats, floating astronomical timescale.

L, long (405 kyr); S, short Ecc. (eccentricity); lpmo, long period modulation of obliquity; lpme, long period modulation of eccentricity; Ob, short obliquity; Pr, precession. Durations are provided only for statistically strong signals.

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