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Home , Bald Eagle Formation, Hirnantian, Neda Formation, Ordovician, Paleosol

Late Glaciation Initiated by Early Land Evolution andPunctuatedbyGreenhouseMassExtinctions

Gregory J. Retallack*

Department of Geological Sciences, University of Oregon, Eugene, Oregon 97403-1272, USA

ABSTRACT Late Ordovician () glaciation, indicated by periglacial paleosols, tillites, and glacial pavements in Saharan , has been attributed to advances in and carbon sequestration due to evolution of early land , in the same way that - glaciation has been attributed to evolution of forests and glaciations to evolution of grasslands. Two problems for carbon cycle explanations of this glaciation are an Hirnantian CO2 green- house estimated at 4035 ppmv from mass balance models and possibly related Hirnantian mass extinctions. This estimate of high CO2 is a 10-m.yr. average, here evaluated against high-resolution sequences of paleosols that give estimates not only of CO2 but also of paleoclimate and vegetation through the Late Ordovician. paleosols of the eastern United States show increases in depth and in degree of carbon consumption and storage comparable with those preceding glaciation of the Devonian-Permian and Quaternary. Early thalloid land plants with rhizoids may have drawn down atmospheric CO2 to only 166 5 83 ppmv during the Hirnantian, as estimated here from a pedogenic CO2 paleobarometer. Ordovician paleosols also reveal a short-term greenhouse spike to as much as 4670 ppmv immediately before Hirnantian glaciation. This greenhouse spike may be due to thermogenic release from a large igneous province and is comparable with greenhouse spikes during other mass extinctions, such as those of the Late Permian and Late .

Online enhancements: supplementary tables.

Introduction The Late Ordovician (Hirnantian) glaciation is well lack and Feakes 1987; Feakes and Retallack 1988) and documented from periglacial paleosols, tillites, and (Driese and Foreman 1991, 1992) and by glacial pavements in Saharan Africa (Le and comparing paleosols from (Yapp and Poths Craig 2008; Finnegan et al. 2011), but it has been a 1992, 1993, 1994), Nova Scotia (Feakes et al. 1989; paradox for carbon cycle models because it occurred Jutras et al. 2009), and (Retallack 2009a, at a time of a CO2 greenhouse (Yapp and Poths 1992; 2009b). Paleosols have unique advantages for such a Berner 2006) and attendant mass extinctions (Fan study because the same paleosol can furnish estimates et al. 2009; Harper et al. 2014). The Late Ordovician of paleoprecipitation (Retallack 2005), paleotempera- was also a time of evolutionary of early ture (Óskarsson et al. 2009), paleoproductivity, and land plants, with consequences for terrestrial weath- atmospheric CO2 (Breecker and Retallack 2014). In ering that may have ushered in the (Retallack addition, long paleosol sequences are now available 2000, 2003; Lenton et al. 2012). This study attempts from the through Devonian (Retallack 2008, to reconcile apparent contradictions in paleocli- 2009a, 2009b; Retallack and Huang 2011), with tem- matic and biotic history of the Late Ordovician by re- poral resolution equal to marine carbonate curves examining paleosol records from (Retal- of d13C, d18O, and 86Sr/87Sr (Veizer et al. 1999; Pro- koph et al. 2008). Thus, both short-term (!1 m.yr.) perturbations like mass extinctions and long-term (110 m.yr.) trends like paleoclimatic cooling can be Manuscript received May 4, 2015; accepted August 14, 2015; compared both on land and at sea. electronically published November 17, 2015. Hypotheses for glaciations in deep time fall into * E-mail: [email protected]. three general categories: (1) volcanic-tectonic, (2) ocean-

[The Journal of , 2015, volume 123, p. 509–538] q 2015 by The University of Chicago. All rights reserved. 0022-1376/2015/12306-0002$15.00. DOI:10.1086/683663

509 510 G . J . R E T A L L A C K atmosphere circulation, and (3) biotic. Uplift of the Material and Methods Himalaya, for example, has been thought to have The most informative sections of paleosols in the increased Cenozoic chemical weathering and car- in Pennsylvania (fig. 1A)were bon sequestration to draw down a greenhouse atmo- roadcuts 4 km east of Potters Mills (Centre County) sphere (Raymo and Ruddiman 1992). Ocean circula- on US Highway 322 (40.7680047N, 77.3982687W), tion changes, such as initiation of the Antarctic 1 km east of Reedsville (Mifflin County) on Old US Circumpolar Current, have also been implicated as Highway 322 (40.6557317N, 77.596867W),and1km causes of Cenozoic glaciation (Kennett 1982). A biotic east of Loysburg (Bedford County) on State Highway hypothesis for inducing Cenozoic glaciation is the evo- 35 along Yellow Creek (40.1591727N, 78.3707337W). lution of grasslands with , which are moist, These outcrops were fresh in 1983 when first exam- organic-rich of high albedo vegetation (Retal- ined but are now becoming overgrown with crown lack 2003, 2013a). Similarly, Devonian to Permian vetch (Coronilla varia) sown by the state highway de- glaciation can be considered the result of Acadian- partment. Five additional sites in Pennsylvania doc- Hercynian mountain uplift (Raymo 1991), Iapetan umented by Retallack (1985) now show only small ocean enclosure (Crowell 1978), or evolution of for- amounts of and no paleosols. Also reex- ests with deep clayey (Alfisol and ) and peaty amined for this study in 2001 and 2012 is a section () soils (Berner 1997; Retallack and Huang in Tennessee (fig. 1B) studied by Driese and Fore- 2011). Hirnantian glaciation has been attributed to man (1991, 1992) opposite the cafe and viewpoint at Taconic mountain uplift (Kump et al. 1999), amal- gamating and expanded microcontinents (Saltzman and Young 2004), or evolution of early land plants increasing the depth and degree of weathering on land (Retallack 2000, 2003; Lenton et al. 2012). Among the various causes for Late Ordovician mass extinction (Harper et al. 2014), Hirnantian gla- ciation has been a leading suspect (Brenchley et al. 1995), but extinction precedes glaciation when de- tailed stratigraphic records are considered (Lefebvre et al. 2010; Ghienne et al. 2014). Three additional hypotheses for mass extinction will also be consid- ered here: (1) methane and from un- usually large volcanic eruptions, (2) nitric and sulfu- ric acid from asteroid or other extraterrestrial impact, and (3) sulfides from oceanic stagnation. A mecha- nism of thermogenic methane release by feeder dikes of large igneous provinces has been proposed for Late Permian mass extinctions (Retallack and Jahren 2008) and has also been inferred for other mass extinctions (Keller 2011; Bond and Wignall 2014). Black pyritic shales with biomarker evidence for atmospheric re- lease of sulfides and sulfuric acid are known for Late Permian (Grice et al. 2005; Bond and Wignall 2014; Sephton et al. 2015) and Late (van de Schoot- brugge et al. 2013) mass extinctions. Black shales and other evidence of anoxia are also known from the Hirnantian (Armstrong et al. 2009a; Fan et al. 2009; Hammarlund et al. 2012). Asteroid impact and atten- dant interruption of photosynthesis by dust clouds and acid rain has been demonstrated for Late Creta- ceous mass extinction (Alvarez et al. 1995; Retallack 2004), but evidence of impact at other mass extinc- tions is not compelling (Retallack et al.1998; Becker et al. 2004). These various hypotheses for glaciation Figure 1. Ordovician paleosol localities and outcrops of and mass extinction are testable by paleosol evidence Pennsylvania (A) and Tennessee (B) and a geological cross of changes in paleoclimate and atmospheric CO2. section of Pennsylvania (from Thompson 1999; C). Journal of Geology ORDOVICIAN PALEOSOLS 511

Beans Gap (Grainger County) on US Highway 25E of the Tuscarora Formation in Pennsyl- (36.3512617N, 83.3977057W). vania (Schuchert 1916; Thompson 1999) and by the A variety of paleoenvironmentally significant mea- Clinch Sandstone in Virginia, West Virginia, Tennes- sures were taken in the field: Munsell hue, nodule size, see, Georgia, and (Driese and Foreman 1991, depth to calcareous nodules, thickness of paleosol 1992). Although there is no angular discordance, this with nodules, and depth of bioturbation (Retallack and other contacts within the Tuscarora and Clinch 2005). Carbonate in the Juniata Formation is partly Sandstones have been considered disconformable dolomitic (Retallack 1993) and weakly reactive to (Dorsch et al. 1994). acid but is distinctive in nodular form and light color, Geological Age. The precise age of the Late Or- so it was estimated as volume percent in outcrop dovician Juniata Formation has been uncertain with- using a relative abundance card (Terry and Chilingar out age-diagnostic but is constrained by ma- 1955). Separate and carbonate samples were col- rine rocks above and below (fig. 3A). The underlying lected from representative beds for laboratory analy- Formation has a distinctive of large ses: major and trace element geochemical analysis by rhynchonellid (Orthorhynchula linneyi) x-ray fluorescence, ferrous iron by potassium dichro- and bivalves (Ambonychia sp.indet.),ofRichmondian mate titration (by ALS Chemex, Vancouver, British (late Katian) age (Schuchert 1916). At Cumberland Gap, Columbia), and bulk density determined from the Tennessee, 70 m of with late weight suspended in air and then water of a paraffin- Katian marine brachiopods is considered laterally coated clod. Petrographic thin sections were cut from equivalent to the lower Juniata Formation (Thompson the same samples, and 500 points were counted using 1970b). The overlying Tuscarora and Clinch Sandstones a Swift automated point counter to determine grain- have a marine trace assemblage of uncertain size distribution and mineral composition (see tables geological age: Arthrophycus alleghaniensis, S1–S3, available online). verticalis, Daedalus , , , Also a part of this work was assembly of databases , and Monocraterion (Butts 1940; Cotter on late Cambrian to early Devonian paleosol depth 1983). The Tuscarora Formation of Pennsylvania also to carbonate (Retallack 2008, 2009a, 2009b; Retal- contains like Velatitetras with lack and Huang 2011) and recalibration of marine a range of – (Johnson 1985; carbonate isotopic data from old time scales used by Vecoli et al. 2011), thus indicating that the Tusca- Veizer et al. (1999) and Prokoph et al. (2008) to cur- rora Formation began accumulating during the ear- rent understanding of numerical age of geological liest . Another point of correlation for the stages (Gradstein et al. 2012). Clinch Sandstone is the Thorn Hill K-Bentonite Complex 30–480 m above the Juniata Formation at Beans Gap and elsewhere in Tennessee and corre- Geological Setting of Appalachian lated into fossiliferous sections near Cincinnati of Ordovician Paleosols or mid–Early Silurian age (Manzo et al. Stratigraphy. The Juniata Formation is red silt- 2002). Above that again the Castanea Sandstone and stone and sandstone (figs. 1A,1B,2A–2C) widespread , near Reedsville, Pennsylvania, in the eastern Valley and Ridge Province of Pennsyl- have brachiopods, , and of Tely- vania, Virginia, West Virginia, and Tennessee (Thomp- chian or late Early Silurian age (Schuchert 1916). son 1970a,1970b, 1999; Driese and Foreman 1991, Traditionally, the Ordovician-Silurian boundary 1992). To the east, the Juniata Formation conformably has been placed within a hiatus between the Juni- overlies folded gray (fig. 1C), ata and Tuscarora Formations or a hiatus within the which is evidence that it was a clastic wedge from Tuscarora Formation (Dorsch et al. 1994), compa- Taconic mountain uplift (Hatcher et al. 1989). In cen- rable with the Martinsburg-Shawangunk hiatus to tral Pennsylvania, Virginia, and West Virginia, Juniata the east (fig. 1C). However, the Juniata Formation, Formation conformably overlies red and green-gray like the latest Ordovician elsewhere, is entirely nor- sandstones and conglomerates of the Bald Eagle (fig. 1C) mally magnetized (Trench et al. 1991) and fails to or , which in turn conformably show reversals found in the earliest Silurian Red overlie gray shale and sandstone of the Reedsville Mountain Formation, equivalent to the Tuscarora For- Formation (Schuchert 1916). To the west, the Juniata mation (Trench et al. 1991), thus supporting a Late Formation passes laterally into red Queenston Shale Ordovician Juniata Formation and an Early Silurian of New York (Dennison 1976) and red shale and do- Tuscarora Formation. Palynological (Johnson 1985; lostone of the Sequatchie Formation of Virginia, Ten- Vecoli et al. 2011) and tephrostratigraphic (Manzo nessee, Georgia, and Alabama (Thompson 1970b). et al. 2002) dating of the Tuscarora and Clinch Sand- The Juniata Formation is abruptly overlain by quartz stones now indicate that little is missing from the Figure 2. Ordovician paleosols in the field. A, Red clayey paleosols of the Juniata Formation overlain by white Clinch Sandstone in Beans Gap, Tennessee (exposed section is 35 m thick). B, Red clayey paleosols in the lower Juniata For- mation near Reedsville, Pennsylvania (hammer shown for scale). C, Red clayey paleosols in the Bald Eagle Formation near Loysburg, Pennsylvania (hammer shown for scale on Loysburg pedotype). D, Type Tussey sandy paleosol at Loysburg with burrows of Skolithos (hammer shown for scale). E, Bedford pedotype paleosol with prototaxaleans in overlying sandstone at Beans Gap (hammer shown for scale). F, Type Morrison silty clay loam paleosol with fine blocky peds and Scoyenia beerboweri burrow at Loysburg (hammer shown for scale). Journal of Geology ORDOVICIAN PALEOSOLS 513

Figure 3. Stratigraphic levels and correlations of studied paleosols. A, Stratigraphic section near Reedsville, Pennsylvania (after Schuchert 1916; Horowitz 1965; Thompson 1968), and relative stratigraphic position of detailed sections (fig. 4). B, Depth to carbonate nodules (Bk horizon) in paleosols of all Appalachian sections examined.

earliest Silurian. Finally, correlation of global sea level (Thompson 1970a). Alteration of (CAI 4) inferred from cumulative aggradation plots and gamma in Ordovician of central Pennsylvania is ray logs of the Juniata Formation has revealed that evidence of burial temperatures of 1607–2107Cand the upper Juniata Formation is Hirnantian (Diecchio burial depths of 4.6–6.6 km (Epstein et al. 1977). Local and Broderson 1994; Diecchio 1995) and thus cor- sulfide mineralization in the Bald Eagle Formation relative with global marine regression and glaciation at Skytop includes quartz veins with fluid inclu- in Saharan Africa (Le Heron and Craig 2008; Ghienne sion temperatures of 1407–3757C (McWilliams et al. et al. 2014). 2006), but such locally high-temperature alteration Burial Alteration. The Juniata Formation and was not seen at other localities examined for this associated formations have been folded about work. Burial compaction to 70%–75% of the original northeast-southwest strike ridges (fig. 1A,1B) that volume of the paleosols in the Juniata Formation can characterize the Appalachian Valley and Ridge prov- be calculated using a standard compaction formula ince (Hatcher et al. 1989). Trioctahedral chlorite and for calcareous soils (eq. [1] for of Sheldon dioctahedral illite in the Juniata Formation are evi- and Retallack 2001) to obtain the original thickness dence of metamorphism to lower greenschist facies of (Bs [cm]) from the thickness of paleosol (Bp 514 G . J . R E T A L L A C K

[cm]) for a known burial depth (K [km]) of 4.6–6.6 km, as ripple marks, planar bedding, and cross-bedding. as follows: Sedimentary structures in the Juniata Formation have been obscured by a variety of distinctive soil features, including crack networks (peds and cutans; fig. 2F), B B p p . (1) bowl slickensides (stress argillans; Gray and Nickel- s 2 =½ = 0.17K 2 0.62 (0.38 e ) 1 sen 1989; Driese and Foreman 1992), clay skins (illu- viation argillans; fig. 8E), expansion cracks (silans; The red color of the Juniata Formation is from he- fig. 8D), elongate drab reduction spots (gleyans; matite of diagenetic origin (Thompson 1970a), recrys- figs. 2E,7F), and replacive dolomite and nod- tallized during early burial from goethite and other ules (caliche; figs. 7C,8C). Petrographic thin sections iron minerals during soil formation (Retallack 1985). show gradation between laminated argillasepic fab- Gray-green mottles also are ubiquitous in the sur- rics of (fig. 8D) and massive mosepic por- faces of the paleosols (fig. 5) and, like comparable fea- phyroskelic fabrics of soils (fig. 8E). Also comparable tures in Cambrian (Retallack 2008) and Triassic (Re- with soils is micritic carbonate intimately associated tallack 1997) paleosols, may have been created by with burrows and sparry calcite tubes after plant burial gleization of organic matter in original surface remains (fig. 8C). horizons. Early burial gleization is confirmed by mag- Distinctive kinds of Ordovician paleosols have netic susceptibility correlation with iron content due already been named Potters Mills, Faust Flat (Feakes to open alteration in mottles of the Potters and Retallack 1988), Powell Mountain, and Beans Mills clay paleosol (Retallack et al. 2003). Gap pedotypes (Driese and Foreman 1991), and this terminology is here amplified to eight pedotypes in the Juniata and Bald Eagle Formations (figs. 5, 6; Paleosols and Associated Fossils tables 1, 2). Pedotypes are defined on the basis of field Sparsely fossiliferous red beds of the Juniata Forma- differences, such as clayeyness and color. Although tion have been considered nonmarine (Schuchert named for its type locality (figs. 4A, 5, 6), the Potters 1916), and their paleochannel form and sedimentary Mills pedotype represents a specific kind of pale- structures have been compared with those of braided osol, which was also found at Beans Gap (fig. 4B) and streams (Cotter 1978, 1983; Davies and Gibling 2010). Reedsville (fig. 4C). The Beans Gap and Powell Moun- While acknowledging that there are no marine sedi- tain pedotypes, in contrast, were seen only at these mentary structures or body fossils in the Juniata For- locations in Tennessee (Driese and Foreman 1991), mation, Davies et al. (2010, 2011) take the icono- at stratigraphic levels regarded here and by Diecchio clastic view that the Juniata Formation was marine (1995) as part of the basal Silurian marine trans- because it is more clayey and rich in trace fossils gression (fig. 3). than usual for a pre-Devonian fluvial formation. This Each pedotype is based on a type profile, sampled view has been challenged elsewhere (Retallack et al. in detail for petrographic and geochemical study to 2011a), and the following paragraphs detail further test the hypothesis that it is a paleosol (figs.5,6;for evidence for paleosols and terrestrial plants and - comparable data on Beans Gap and Powell Moun- mals in the Juniata Formation. tain pedotypes, see Driese and Foreman 1991). These Field and Petrographic Evidence of Paleosols. Paleo- additional data also allow identification of pedo- sols are recognized by a trinity of (1) root traces, types within classifications of modern soils (table 1), (2) soil horizons, and (3) soil structures (Retallack such as the paleosol classification of Mack et al. 2001b). True roots had not evolved by the Ordovi- (1993), the old (Stace et al. 1968) and new (Isbell 1998) cian (Retallack 2000), but there are beds with ir- Australian classifications, the world map of soils of regular strata-transgressive tubular features that ap- the Food and Agriculture Organization (1974), and pear to be burrows and early land plants (figs. 5–7). the soil taxonomy of the Staff (2010) Many beds in the Juniata Formation have the sharp of the US National Resources Conservation Service. top and gradational lower alteration characteristic Petrographic analysis has confirmed that these Or- of soil horizons (figs. 2A–2C, 5, 6). These truncated dovician paleosols all lack argillic horizons (illuvial tops are strongly ferruginized and clayey and grade clay-enriched horizons of Soil Survey Staff 2010), down into a subsurface zone of carbonate nodules despite their clay skins, bioturbation, and thickness (fig. 2B,2F), like A-Bk-C profiles of paleosols, rather (figs. 2F, 5, 6) comparable with those of Argillisols than bioturbated marine beds. in sed- (Mack et al. 1993), (Stace et al. 1968), Chro- imentary rock is often characterized as massive, mosols (Isbell 1998), Luvisols (Food and Agriculture featureless, and nondescript because it develops at Organization 1974), and Alfisols (Soil Survey Staff the expense of familiar sedimentary structures, such 2010). Figure 4. Detailed measured sections of Ordovician paleosols in Pennsylvania (A, C, D) and Tennessee (B). Columns by graphical log show paleosol development (from Retallack 2001b), volume percent carbonate nodules (estimated from the field card of Terry and Chilingar 1955), and Munsell hue. 516 G . J . R E T A L L A C K

Figure 5. Petrographic data, chemical index of alteration values, and concentration ratios for Ordovician paleosols of the Juniata Formation (Late Silurian Beans Gap and Powell Mountain pedotypes of the Juniata are described in Driese and Foreman 1991). Colors are from a Munsell chart, sand-silt-clay and mineral proportions are from point counting, and molecular weathering ratios were calculated from major-element chemical analysis of the indicated specimens from the Condon Fossil Collection, Museum of Natural and Cultural History, University of Oregon.

Analysis of Chemical Weathering. The most im- different levels within a paleosol can be determined portant weathering reaction on Earth is hydrolysis, by major element chemical analysis. One common which supplies cations to fuel plant and fungal pro- measure of overall hydrolysis is the chemical index duction of acid and reduces both the volume and the of alteration (I [mole fraction]), calculated from mo- mass of soil consumed (Paton et al. 1995). Hydrolysis lar proportions (m) of alumina, lime, potash, and soda is incongruent dissolution of weatherable minerals, according to eq. (2) (Nesbitt and Young 1982). The such as feldspar and biotite, in carbonic acid from lime is noncarbonate lime. dioxide to produce clay and cations in solution (van Breemen et al. 1983; Chadwick and 100(mAl O ) Chorover 2001). The most common of these cations I p 2 3 . (2) 21 21 1 1 1 1 1 are Ca ,Mg ,Na, and K , whose abundance at mAl2O3 mCaO Na2O mK2O Journal of Geology ORDOVICIAN PALEOSOLS 517

Figure 6. Petrographic data and molecular weathering ratios of Ordovician paleosols for the Juniata Formation (Late Silurian Beans Gap and Powell Mountain pedotypes of the Juniata are described in Driese and Foreman 1991). Colors are from a Munsell chart, sand-silt-clay and mineral proportions are from point counting, and molecular weathering ratios were calculated from major-element chemical analysis of the indicated specimens from the Condon Fossil Collection, Museum of Natural and Cultural History, University of Oregon.

The chemical index of alteration is highest near the in temperate climatic source terrains (Passchier and surface of moderately developed Potters Mills and Krissek 2008). Bedford pedotypes but shows less variation within A better indication of individual paleosols in al- weakly developed alluvial paleosols (fig. 5). Chemi- luvium are major oxide concentration ratios of in- cal index of alteration values for Ordovician pale- dividual beds down to likely parent material, fol- osols are similar to those produced by weathering lowing Grandstaff et al. (1986), which reveals the 518 G . J . R E T A L L A C K

Figure 7. Ordovician fossil plantlike remains. A, B, Lepidotruncus sp. indet. from above Bedford pedotype at Beans Gap, Tennessee. C, Radix sp. indet. from Faust Flat pedotype at Potters Mills, Pennsylvania. D, Unidentified im- pression (thalloid liverwort?) dotted with black iron-manganese nodules from type Potters Mills clay paleosol at Potters Mills. E,Unidentified laterally branching calcite tubes (leafy liverwort?) from Faust Flat pedotype at Potters Mills. F, Gray-green mottle surrounded by smaller tubular mottles in red siltstone of the type Potters Mills clay from Potters Mills (D–F). A–C and F are field photographs, but others are in the Condon Fossil Collection, Museum of Natural and Cultural History, University of Oregon (P16295 [D], F35150 [E]).

Potters Mills clay as the most deeply weathered of mation marked in Potters Mills and Bedford pedo- the paleosols and the Faust Flat silty clay as the types compared with other paleosols (fig. 6). least weathered (fig. 5). Molecular weathering ratios Weathering trends within beds are most effec- of alumina over bases are also evidence for hydro- tively explored by tau analysis (Brimhall et al. 1992). lysis in the various paleosols in proportion to ob- The mass transfer of elements in a soil at a given served physical and petrographic degree of soil for- horizon (tw,j [mole fraction]) can be calculated from 23 mation (fig. 6). Other molecular weathering ratios the bulk density of the soil (rw [g cm ]) and parent 23 are designed to evaluate differences between paleo- material (rp [g cm ]) and from the chemical con- sols in gleization (ferrous/ferric iron), salinization centration of the element in soils (Cj,w [weight %])

(soda/potash), calcification (alkaline earths/alumina), and parent material (Cp,w [weight %]). Changes in and clay formation (alumina/silica). Salinization, cal- volume of soil during weathering (strain) are esti- cification, and gleization are subdued and clay for- mated from an immobile element in soil (such as Ti, Journal of Geology ORDOVICIAN PALEOSOLS 519

Figure 8. Bioturbation of Ordovician paleosols in thin section. A, Sand-filled Skolithos burrows from A horizon type Faust Flat silty clay paleosol at Potters Mills, Pennsylvania. B, Clay-lined surface cavity ( or egg chamber) and indistinct subsurface bioturbation from the uppermost A horizon of the type Morrison silty clay loam at Loysburg, Pennsylvania. C, Likely millipede burrow of Scoyenia beerboweri truncating pedogenic micrite and truncated by pedogenic micrite around a tubule with a sparry center. D, Skolithos burrow and relict bedding from the lower Bk horizon of the type Morrison silty clay loam at Loysburg. E, Indistinct bioturbation and birefringence fabric (porphyroskelic mosepic fabric) from the lower A horizon of the type Bedford clay paleosol at Loysburg. Specimen numbers are F36212 (A), R2301 (B), F36211 (C), R2303 (D), and R2298 (E), Condon Fossil Collection, Museum of Natural and Cultural History, University of Oregon.

used here) compared with parent material (εi,w [mole diagenetic processes, such as illitization and albiti- fraction]). This strain is restricted to the centime- zation (Maynard 1992). Juniata Formation beds show ter to meter thicknesses of a single paleosol pro- collapse and loss of weatherable bases and collapse file, whereas burial compaction acts over kilometers and gain of sesquioxides, as in comparable (Sheldon and Retallack 2001). The relevant equations alluvial calcareous paleosols (Retallack 1991). Mio- (3) and (4) (below) are the basis for calculating diver- cene paleosols were chosen for comparison, rather gence from parent material composition (origin in than soils, because they have already been modi- fig. 9). Subscripts in these equations are for immobile fied by burial compaction, which reduces strain and (i) and studied (j) element and for soil (w) and parent mass transfer to a more muted signal than in soils. (p) material. The observed collapse occurs because depletion due to weathering has concentrated weather-resistant r w Cj,w minerals (such as ilmenite) and elements (such as Ti) tj,w p (εi,w 1 1) 2 1, (3) rpCj,p in the upper parts of beds, as in soils, rather than concentrated heavy minerals at the base of the bed, r C ε p p j,p 2 as in graded sedimentary beds (Brimhall et al. 1992). i,w r 1. (4) w Cj,w The loss of some elements occurs because the sur- faces of the beds have been depleted in weatherable Tau analysis of selected beds in the Juniata For- cations, as in soils, rather than cation enriched, as in mation (fig. 9) does not fall into the dilate-and-gain marine clays (Feakes and Retallack 1988). Gain of field of both marine accumulation and alumina is due to clay formation in the surface of the Table 1. Pedotypes of the Juniata and Bald Eagle Formations Type example (level Paleosol classi- US taxonomy FAO map (Food Old Australian New Austra- in m relative to fication (Mack (Soil Survey and Agriculture (Stace et al. lian (Isbell Pedotype Diagnosis Reedsville section) et al. 1993) Staff 2010) Organization 1974) 1968) 1998) Beans Gap Red (5R) clayey (A) horizon Beans Gap clay 723 m Hapludert Chromic Vertisol Brown Clay Brown with drab (5Y) and pyritized at Beans Gap Vertosol Skolithos over slickensided thick (90 cm) clayey (Bw) horizon Bedford Greenish gray (5GY) bedded Bedford clay 58 m at Protosol Fluvent Eutric Alluvial Soil Stratic siltstone (A) with Skolithos Loysburg Rudosol over red (5YR) bedded silt- stone (C) Faust Flat Thin (18 cm) red (2.5YR) silty Faust Flat silty clay Protosol Earthy Sand Earthy Sand Lutic clay (A horizon) with abun- 624 m at Potters Rudosol dant Scoyenia burrows over Mills red (2.5YR) bedded sandstone (C) horizon Loysburg Red (2.5YR) siltstone (A) with Loysburg silt loam Calciustept Calcic Desert Loam Calcenic Scoyenia over red (2.5YR) 57 m at Loysburg Tenosol siltstone with shallow

520 (12–14 cm) scattered small nodules of dolomite and calcite (Bk) Morrison Red (10R) siltstone with Morrison silty clay Calcisol Vertic Calcic Cambisol Desert Loam Hypocalcic Scoyenia over shallow loam 57.5 m at Haplocalcid Calcarosol (18 cm) common nodules of Loysburg dolomite and calcite (Bk) Potters Mills Red (2.5YR) clayey surface Potters Mills clay Calcisol Ustic Calcic Xerosol Calcareous Calcic with drab (5Y) mottles and 625 m at Pottters Haplocalcid Red Earth Calcarosol abundant Scoyenia burrows Mills over shallow (19–36 cm) horizon (Bk) of small dolo- mitic and calcitic nodules Powell Red (5R) slickensided clayey Powell Mountain clay Protosol Sulfaquept Thionic Fluvisol Wiesenboden Supratidal Mountain (A) horizon with drab (10Y) just below Clinch Hydrosol dolomitic mottles over red Sandstone at Powell (10R) bioturbated claystone Mountain and drab (10Y) dolomitic and phosphatic siltstone with Tussey Greenish gray (5GY) bedded Tussey sandy clay Protosol Psammaquent Eutric Alluvial Soil Oxyaquic siltstone (A) with Skolithos loam 39 m at Hydrosol burrows over bedded green- Loysburg ish gray (5GY) sandstone Table 2. Paleoenvironmental Interpretation of the Juniata and Bald Eagle Formations Time of formation Pedotype Paleoclimate Vegetation Paleotopography Parent material (k.yr. ago) Beans Gap Subhumid (mean annual pre- Unknown Worms (Skolithos) and Coastal plain Smectite clay ∼10 cipitation, 720 5 182 mm) brachiopods (Lingula) temperate (mean annual temperature, 11.4754.47C), with marked dry season Bedford Not diagnostic of climate Nematophtyes Alluvial levees Quartzo-feldspathic ∼.1 (Lepidotruncus fortis, silt and sand Radix corrugatus) Faust Flat Not diagnostic of climate Nematophtyes (Radix (Scoyenia Alluvial levees Quartzo-feldspathic ∼.1 corrugatus) beerboweri) and silt and sand worms (Skolithos)

521 Loysburg Subhumid (mean annual pre- Unknown Millipedes (Scoyenia Alluvial terraces Quartzo-feldspathic 3.9 5 1.8 cipitation, 810 5 182 mm) beerboweri) and silt temperate (mean annual worms (Skolithos) temperature, 10.9754.47C) Morrison Subhumid (mean annual pre- Unknown Millipedes (Scoyenia Alluvial terraces Quartzo-feldspathic 5.0 5 1.8 cipitation, 607 5 182 mm) beerboweri) and silt temperate (mean annual worms (Skolithos) temperature, 9.6754.47C) Potters Mills Subhumid (mean annual pre- Thalloid liverwort-like Millipedes (Scoyenia Alluvial terraces Quartzo-feldspathic 6.8 5 1.8 cipitation, 827 5 182 mm) plants beerboweri) silt temperate (mean annual temperature, 11.0754.47C) Powell Mountain Not diagnostic of climate Unknown Worms (Skolithos) and Supratidal flat Smectite clay and ∼.1 brachiopods (Lingula) dolomite Tussey Not diagnostic of climate Unknown Worms (Skolithos) Estuarine sand Quartzo-feldspathic ∼.1 flats silt Figure 9. Mass transport (t) and strain (ε) in three kinds of Ordovician paleosols from Pennsylvania, compared with a variety of late Miocene paleosols from Pakistan, which are also in a deeply buried sequence of calcareous red paleosols (Retallack 1991). Journal of Geology ORDOVICIAN PALEOSOLS 523 paleosols, and gain in potash may reflect burial dia- ation (Melim et al. 2004) or cellulose respiration genesis (Feakes and Retallack 1988; Fedo et al. 1995). (Ehleringer et al. 2000). These correlations and low The few samples in dilate-and-gain quadrants of fig- values are easily distinguished from unaltered ma- ure 9 reflect sedimentary additions to the tops of the rine (fig. 10E), which clusters near 0 for paleosols (cumulic horizons; Retallack 1991, 2000). both d13Candd18O of modern sea shells (Veizer et al. Tau analysis also reveals a different degree of weath- 1999) and is more negative than modern for d18Oin ering: Potters Mill profiles are more weathered than Ordovician (Prokoph et al. 2008) and Cambrian Morrison or Loysburg profiles, and these Ordovician (Surge et al. 1997) marine limestones. The long-term paleosols overlap trends of Miocene paleosols of secular increase in d18O values for sea water may be comparable calcareous red beds (fig. 9). related to changes in ocean crust hydrothermal re- Tau analysis also reveals that these paleosols were gime (Kasting et al. 2006). For these reasons, marine biologically active because depletion in limestones with indistinct correlations between d13C soils depends on ligands released by microbes (Nea- and d18O are poor indicators of paleoclimate, mixing man et al. 2005). The main source mineral for phos- signals of original sea water and later weathering phorus in soils is apatite, which is relatively inert to (Lohman 1988; Melim et al. 2004). In contrast, pedo- carbonic acid attack. Weathering of phosphorus and genic carbonate with tight correlations preserve orig- alkali and alkaline earth bases in Ordovician soils is inal soil signals (Knauth et al. 2003; Huang et al. direct evidence for biotic enhancement of weather- 2005; Ufnar et al. 2008; Retallack et al. 2014). ing inferred by carbon cycle models (Berner 2006). Also distinctive is a pattern of near constant d18O Isotopic Fractionation of Pedogenic Carbonate. A dis- but highly varied d13C in carbonate of marine meth- tinctive diagnostic of carbonate ane cold seeps (fig. 10F; Aiello et al. 2001). Meth- in paleosols (fig. 10A) and soils (fig. 10B) is correla- anogenic carbonate d13C is extremely negative: down tion of d13C and d18O values (Knauth et al. 2003; to –48‰ in marine methane seeps (Peckmann et al. Huang et al. 2005; Ufnar et al. 2008; Retallack et al. 2002) and to –43‰ in wetland paleosols (Ludvigson 2014). Such tight correlations are inherited from the et al. 2013). This “meteoric line” pattern 18 same covariance seen in respired soil CO2 (Ehlerin- arises because of the dominance of d O composition ger and Cook 1998; Ehleringer et al. 2000), and that by water in waterlogged to aquatic habitats. in turn is derived from the same covariance ob- Tight correlations of d13Candd18O are observed served in plant cellulose (Barbour et al. 2002). These only for soils of similar vegetation and climate observations confirm models that show that carbon (fig. 10B) and paleosols of particular geological for- and oxygen isotopic covariance is due to selection mations (fig. 10A) because the isotopic composition of light isotopologs of CO2 during its diffusion into of soil CO2 varies with vegetation and climate (Uf- cells and consumption by photosynthesis (Barbour nar et al. 2008). Moderately significant correlations and Farquhar 2000). Carbon-oxygen isotopic covari- and surprisingly positive values for both d13C and ance is a characteristic ecosystem relationship (Bar- d18O are found on modern carbonate crusts on basalt bour et al. 2002), with strongest correlations in semi- (fig. 10C). arid communities of mixed C3-C4 photosynthetic Fossil Plants. The Juniata Formation predates the pathways and seasonal variations in moisture and oldest known vascular land plant megafossils: Silu- plant productivity (Liu et al. 2014). The slope of the rian (Ludlovian) longifolia with Bo- relationship shows a positive relationship with mean hemograptus bohemicus graptolites (426 Ma) from annual vapor pressure deficit (kPa), or the difference Victoria, Australia (Rickards 2000), and between moisture in air and the amount a plant can pertoni in the Saetograptus leintwardiensis grapto- hold when saturated, a measure of evapotranspira- lite zone (425 Ma) in (Cleal and Thomas 1995). tion (Barbour et al. 2002). Molecular-clock ages of Ordovician have been de- Marine limestone altered by surface weathering termined for liverworts (475 Ma) and (458 Ma), also shows positive correlation (fig. 10D). Such altera- but ages of Silurian have been determined for horn- tion is commonly called burial diagenesis but is due worts (440 Ma) and vascular land plants (422 Ma; to dissolution and replacement from percolation of Magallón et al. 2013). Ordovician plants have been meteoric water, sometimes to exceptionally deep lev- considered , but recent research on els well below current sea level during the last glacial early land plants suggests several extinct prebryo- maximum (Lohman 1988; Melim et al. 2004). Lacus- phytic lineages, and a better term for nonvascular trine marls may also show modest positive correla- Ordovician-Silurian plants is “cryptophytes” (Ed- tions (r2 of 0.4–0.6), attributed to closed basin hydrol- wards et al. 2014). There is also evidence from pal- ogy (Talbot 1990) or climate change (Drummond ynology for glomalean fungi (Redecker et al. 2000) et al. 1995) but perhaps also due to meteoric alter- and cryptophytes (Steemans et al. 2010; Kenrick 524 G . J . R E T A L L A C K

Figure 10. Carbon and oxygen isotopic composition of carbonate in (A) Ordovician paleosols of Pennsylvania (Retallack 1993) and Australia (Retallack 2009b), compared with and Cambrian paleosols of Australia (Retallack et al. 2014) and Silurian paleosols of Pennsylvania (Harrisburg data of King 1998); (B) soil nodules (above Woodhouse lava flow, near Flagstaff, AZ, from Knauth et al. 2003; and in Yuanmou Basin, Yunnan, , from Huang et al. 2005); (C) soil crusts on basalt (Sentinel Volcanic Field, AZ, from Knauth et al. 2003); (D) Quaternary marine limestone altered diagenetically by meteoric water (Key Largo, FL, from Lohman 1988 and Clino Island, Bahamas, from Melim et al. 2004); (E) (open circles) and Ordovician (open squares) unweathered marine limestones (Veizer et al. 1999; Prokoph et al. 2008) and early Cambrian (filled circles), Ajax Limestone, South Australia (Surge et al. 1997); and (F) marine methane cold seep carbonate (Miocene, Santa Cruz Formation, Santa Cruz, CA, from Aiello et al. 2001). Slope of linear regression (m) and coefficients of determination (R2)showthat carbon and oxygen isotopic composition is significantly correlated in soils and paleosols. et al. 2012) during the Katian and Hirnantian. This remains from Ordovician and Silurian rocks of Ap- kind of low, clumped vegetation of well-drained palachia have the distinctive carbon isotopic com- soils has been called a polsterland by Retallack position of land plants (Tomescu and Rothwell 2006; (1992), comparable with modern bryophytic soil Tomescu et al. 2009) but represent wetland marsh crusts of deserts (Belnap et al. 2003). Carbonaceous and microbial mat communities distinct from the Journal of Geology ORDOVICIAN PALEOSOLS 525 dryland vegetation represented by most of the red interbeds or weakly developed paleosols. Comparable calcareous paleosols described here (Feakes and Re- low-diversity assemblages are found else- tallack 1988). where in Cambrian-Ordovician fluvial and lacustrine Fossil impressions and drab mottles in the Juni- rocks (Trewin and McNamara 1994; Mikuláš 1995; ata Formation are suggestive of thallose liverworts Retallack 2008, 2009a, 2009b). The Juniata Forma- (fig. 7D), but some are comparable with leafy liver- tion has not yet yielded marine trace fossils such as worts (fig. 7E). Also locally common, particularly Rusophycus or , common among at least in Tussey and Faust Flat pedotypes, are stout casts 22 ichnogenera in Ordovician (Katian) rocks with un- of irregularly wrinkled trunks engulfed in cross- doubted marine fossils around Cincinnati, Ohio (Os- bedded sandstone, with branching and anastomos- good 1970). ing, rugulose, rootlike structures (fig. 7A,7B). Some Ordovician paleosols of the Juniata Formation have of these strata-transgressive structures are prefer- common burrows of invertebrates with prominent entially encrusted with pedogenic carbonate (fig. 7C). clay-silt backfills (Scoyenia beerboweri Retallack These various trunks are comparable with fossil trunks 2001a). These burrows formed during soil formation described as Lepidotruncus fortis (Fritsch 1908) and because they both cut across and are cut by micritic rootlike structures, Radix corrugatus (Fritsch 1908), carbonate with microtexture (calciasepic) and isoto- from the Katian-Hirnantian (Kříž and Pojeta 1974), pic composition of pedogenic carbonate (fig. 8C). The Kosov Formation, near Chodouň and Vonoklas (re- morphology of backfills, fecal pellets, and burrow- spectively), Czech Republic. Comparable remains (here size distribution are compatible with the creation referred to as Lepidotruncus sp. indet.) are known of S. beerboweri by polyzoniid or similar millipedes from Hirnantian glacial deposits of the El Golea (Retallack 2001b; Retallack et al. 2011). This inter- Member, Tamadjert Formation, near Zmeïlet Barka, pretation was challenged by Davies et al. (2010), who Algeria (Arbey and Koeniguer 1979; Le Heron and considered the oldest millipedes Silurian and were Craig 2008). More informative comparable remains unaware of Late Ordovician (Sandbian) millipede track- (here referred to Radix sp. indet.) are known from ways (Johnson et al. 1994), Cambrian millipede- the early Katian near Ingham like fossils (Hou and Bergström 1998; Budd et al. Mills, New York (Argast 1992). The New York Radix 2001; Retallack et al. 2011), or phylogenomic (Rota- is partly permineralized in calcite and shows a dis- Stabelli et al. 2013; Misof et al. 2014) and biogeo- tinctive of tubular cell walls with a carbon graphic (Shelley and Golovatch 2011) evidence for a 13 isotopic composition (d Corganic vs. Pee Dee belem- Cambrian origin of millipedes. Another remarkable nite [PDB] of 223.5‰ to 225‰) of land plants rather trace fossil found in thin section of the surface of the than marine organic matter (Tomescu et al. 2009). type Morrison clay paleosol (fig. 8B, upper left) is a The tubular histology is comparable with that of clay-lined structure, within an excavated hollow, Devonian logani, the best preserved of containing organic lined spherical structures. This nematophytes, interpreted as a giant basidomycete enigmatic fossil is comparable with egg and molt by Hueber (2001; Boyce et al. 2007), a lichen chambers of millipedes (Romell 1935; Toye 1967). by Retallack and Landing (2014), and a rolled mat Other common trace fossils in the Juniata For- of liverworts by Graham et al. (2010). Rolled mat mation include remains comparable with Skolithos is not plausible for these Ordovician rootlike and (fig. 8A), Circulichnus, Scolicia, Planolites, Palaeo- current resistant structures because they are strata- phycus, and Helminthopsis (Diecchio and Hall 1998; transgressive (fig. 7A–7C). Other terrestrial nemato- Davies et al. 2010), all commonly attributed to worm- phytes are known from Early Ordovician sandstones like or sluglike organisms. (Grindstone Range Sandstone near Wirrealpa, South Also found in the Juniata Formation are rare track- Australia; Retallack 2009a) and Middle Cambrian ways up to 3 cm wide of (Diecchio and phosphorites ( Creek Formation at Mount Mur- Hall 1998; Davies et al. 2010), comparable with ray, Queensland; Fleming and Rigby 1972; Southgate older Ordovician gouldii from the Mid- 1986). dle Ordovician () part of the Tumblagooda Trace Fossils. The diversity of trace fossils in beds Sandstone of Kalbarri, Western Australia (Trewin of a particular pedofabric in the Juniata Formation is and McNamara 1994; Retallack 2009b), and Early low, usually only Skolithus and Scoyenia (table 2). Ordovician () lower Grindstone Range The other ichnogenera of trace fossils known from Sandstone of Wirrealpa, South Australia (Retallack the Juniata Formation (cf. Diplichnites, Circulichnus, 2009a). These Australian Ordovician trackways were Helminthopsis?, Palaeophycus, Planolites, Scolicia?; attributed to large amphibious euthycarcinoids, such Retallack 1985, 2000; Diecchio and Hall 1998; Davies as Kalbarria brimmellae (McNamara and Trewin et al. 2010; Retallack et al. 2011) are in lacustrine 1993). 526 G . J . R E T A L L A C K

Marginal marine trace fossils (Skolithos verticalis) water table or inundated by coastal lagoons with penetrate a firmground atop the Beans Gap clay pa- lingulid and rhynchonellid brachiopods (Schuchert leosol near the top of the Juniata Formation at Beans 1916; Dreise and Foreman 1992). Gap, Tennessee (fig. 4A). In this case, transitional Time for Formation. A useful proxy for the time marine origin is betrayed by local gleization with over which a paleosol formed is the size of carbonate pyritization and phosphatization as well as lingulid nodules in paleosols (Retallack 2005). In modern fragments (Driese and Foreman 1991). soils, the diameter of nodules (S) is related to soil age Ordovician marine bioturbation was thorough and (A [ka]) by equation (5) (R2 p 0.57; SE, 1.8 ka; Re- deep (Sheehan and Schiefelbein 1984), but coeval tallack 2005). physical disruption of soils was far from negligible. A p 3.92S0.34. (5)

New Interpretations of Paleosols Burial compaction has a minor flattening effect on Red, clayey, fractured, and nodularized beds of the carbonate nodules but cannot alter their horizontal Juniata Formation do indeed appear to be paleosols diameter because of lithostatic pressure from the (Retallack and Feakes 1987; Feakes and Retallack 1988; side (Sheldon and Retallack 2001). This proxy is Retallack 1993; Driese and Foreman 1991, 1992), and unlikely to be compromised by lack of vascular land they have much to tell us about Ordovician terres- plants because the formation of soil carbonate is trial environments and life on land (fig. 11). considered largely microbial (Monger et al. 1991). Parent Material. The Juniata and Bald Eagle For- The results of these calculations show soil dura- mations are quartzofeldspathic in mineral compo- tions of 4–7 k.yr. for carbonate in Ordovician pale- sition, derived from of preexisting sedimen- osols. For the 116-m measured section at Potters tary rocks, such as the Martinsburg Formation in Mills (21% of the whole section of Juniata Formation the to the east (Thompson 1999). at Reedsville), the durations for 8 Potters Mills and There is a surprising amount of feldspar in the sand- 11 Faust Flat paleosols add up to a total duration of stones and claystones (figs. 5, 6) compared with Me- 55,500 yr. For the 34-m measured section at Beans sozoic and Cenozoic sediments (Retallack 1991, 1997). Gap (6% of the whole section of Juniata Formation at These differences in other early sandstones Reedsville), the durations add up to a total duration have been attributed to a different regime of weath- of 54,100 yr. If these sections represent the interval ering before the advent of vascular land plants (Dott of low Hirnantian sea level, dated at 445.2–443.8 Ma 2003; Jutras et al. 2009). Tau analysis of paleosols (Gradstein et al. 2012), this 1.4 m.yr. is 25 times the does indicate less mobilization of potash from Or- duration represented by paleosols. Such discrepancies dovician than Miocene paleosols (fig. 9). Parent ma- are normal for geological versus paleosol estimates of terial composition of the soils that supplied sediment geological time of Cenozoic paleosol sequences, which to the Juniata Formation is known only in general are rife with both major and minor disconformities terms, but this article considers mainly soil forma- (Retallack 1998). tion forward from single beds of weathered alluvium Paleoproductivity. Soil productivity can be mea- as a parent material (figs.5,9). sured in a variety of ways, including secondary pro- Paleotopography. The Juniata and Bald Eagle For- ductivity of microbial and root-respired CO2 within mations have long been interpreted as deposits of the soil responsible for hydrolytic supply of nutrients sheet-braided streams draining a coastal plain to the to plants (van Breemen et al. 1983; Chadwick and west of the Taconic Mountains (Cotter 1978). Abun- Chorover 2001). A late-growing-season soil CO2 pa- dant burrows (Scoyenia beerboweri)inmanyofthe leobarometer of Retallack (2009c) from modern soils paleosols are evidence that those soils were well instrumented for CO2 content has been updated by drained and presumably elevated on alluvial terraces Breecker and Retallack (2014) and relates productiv- above paleochannels (Retallack 2001a). Scoyenia- ity (Pr [ppmv CO2]) with depth to carbonate (Ds [cm]) bearing paleosols of the Juniata Formation are red and as equation (6) (R2 p 0.66; SE, 766 ppmv). highly oxidized, but the Bald Eagle Formation and the uppermost Juniata Formation also have paleosols Pr p 35.3Ds 1 588. (6) that are entirely or partially gray-green due to early burial gleization (Thompson 1970a). The drab paleo- The alternative equation for Pr of Cotton and sols also have a different assemblage of trace and body Sheldon (2012) derived from paleoprecipitation esti- fossils, including Skolithos and Lepidotruncus.These mates of paleosols is not suitable for Devonian or paleontological differences can be related to slow older paleosols because even Devonian forests were drainage of gray-green paleosols, closer to alluvial less productive at similar isohyets than modern eco- Journal of Geology ORDOVICIAN PALEOSOLS 527

Figure 11. Reconstructed Ordovician soils and landscapes of the Juniata and Bald Eagle Formations in Pennsylvania. An earlier reconstruction by Feakes and Retallack (1988) showed only inland pedotypes. Beans Gap and Powell Mountain paleosols of the Juniata Formation in Tennessee (Driese and Foreman 1992) are not included because they are in a different region and are also likely to be early Silurian (fig. 3). systems (Retallack and Huang 2011), and this article three separate sections separated by 2000 km in aims to quantify even lower productivity of Ordovi- Western and South Australia (Retallack 2009b). Fur- cian ecosystems. Using equation (6), respired soil thermore, it correlates with global marine isotopic

CO2 can be calculated for paleosols from depth to Boda event and Hirnantian mass extinction (Arm- carbonate nodules once they have been corrected to strong et al. 2009a). Comparable paleoproductivity original soil depths using equation (1). This proxy similar to modern desert soils, including produc- included desert soils in which pedogenic carbonate is tivity spikes, have also been calculated for 383 suc- mainly a microbial precipitate (Monger et al. 1991), cessive Cambrian-Silurian paleosols of the Tumbla- so it is applicable to pre-Devonian paleosols. gooda Sandstone of Western Australia (fig. 12A; The results of these calculations for the Late Retallack 2009b). Rising late Katian terrestrial pro- 13 Ordovician Potters Mills clay paleosol is soil CO2 ductivity was coeval with rising d C in marine car- of 1538 5 766 ppmv, which is low compared with bonate (fig. 12B,12C), which reflects increased ma- 19 other Ordovician paleosols of Pennsylvania and rine productivity (Prokoph et al. 2008).

Tennessee showing increased depth to Bk and thus Soil CO2 levels in Ordovician and Silurian arid- increased late-growing-season soil CO2 until a Hir- land paleosols are significantly less than soil CO2 of nantian decline (fig. 3B). A spike in soil secondary 26,656 ppmv determined by Yapp and Poths (1993, productivity is apparent near the Katian-Hirnantian 1994) for a Late Ordovician paleosol capping the boundary in Pennsylvania and Tennessee. This spike Neda Iron Formation in Wisconsin. This thick (1.5 m) is not just noise in the data because it is also found in lateritic paleosol (Retallack 2003) is very different 528 G . J . R E T A L L A C K

Figure 12. Proxies for terrestrial productivity (A), marine carbon cycle and paleotemperature (B), terrestrial weath- ering (C), and first appearance of land plants and animals (D). Terrestrial productivities were calculated by the method of Breecker and Retallack (2014) using Australian data of Retallack (2009b; circles) and North American data (herein; squares). Isotopic data are from Veizer et al. (1999) and Prokoph et al. (2008), recalibrated to the time scale of Gradstein et al. (2012). Range data are from Retallack (2008, 2009a,2009b), Steemans et al. (2010), Retallack et al. (2011), and Kenrick et al. (2012).

from the calcareous paleosols described here and Sheldon et al. (2002) on the basis of the chemical formed farther north within the tropics of the time, index of alteration without potash, or CIA-K (A p during the latest Katian (McLaughlin et al. 2013). Mod- 100 # mAl2O3/(mAl2O3 1 mCaO 1 mNa2O) [moles]), ern calcareous soils have CO2 concentrations less which increases with mean annual precipitation than 12,400 ppm (Retallack 2009c), but modern lat- (R [mm]) in modern soils (R2 p 0.72; SE, 182 mm), as eritic soils have CO2 concentrations up to 104,000 follows: ppm (Matsumoto et al. 1997). The Ordovician trop- ical lateritic paleosol of the Neda Iron Formation p 0.0197A thus had productivity that was only about 26% of R 221e . (7) modern lateritic soils. The lower productivity of Ordovician compared This formulation is based on the hydrolysis equa- with modern terrestrial ecosystems is more than a sim- tion of weathering, which enriches alumina at the ple pH effect on potassium mobilization inferred by expense of lime, magnesia, potash, and soda. Mag- Jutras et al. (2009). Both calcareous (alkaline) and lat- nesia is ignored because it is not significant for most eritic (mildly acidic) Ordovician paleosols discussed sedimentary rocks, and potash is left out because here show evidence of lower than modern produc- it can be enriched during deep burial alteration of tivity. Nor was the Ordovician a unique time of un- sediments (Maynard 1992). This proxy is not suit- usually alkaline soils between an acidifying primeval able for samples with free carbonate but can be ap-

CO2 greenhouse and acidifying vascular land plants plied to carbonate-free soil above carbonate at depths (Jutras et al. 2009) because potash-rich calcareous pa- comparable with Ordovician paleosols (17–47 cm leosols are known in the Cambrian (Retallack 2008), before burial compaction estimated from eq. [1]). (Driese et al. 1995), and Archean (Wata- This proxy also is applicable to pre-Devonian soils nabe et al. 2004). with clay skins and ped structures comparable with Paleoprecipitation. Paleoprecipitation can be esti- Bw horizons (figs. 5, 6) because carbonic acid from mated from paleosols using the paleohyetometer of microbial is the principal agent of Journal of Geology ORDOVICIAN PALEOSOLS 529

hydrolysis (Paton et al. 1995). Unlike the bed-scale the lowest and the highest nodule in the profile (Ho tau analysis employed here (eq. [2]), this method [cm]), by equation (8) (R2 p 0.58; SE, 22 mm). conflates weathering that contributed to the parent material as well as within the paleosol. An alterna- M p 0.79Ho 1 13.71. (8) tive paleohyetometer uses depth to carbonate nodules in soils, which increases with mean annual precipi- For application to paleosols, the thicknesses of pa- tation (Retallack 2005) but shows a closer relation leosol with carbonate must be corrected for com- to late-growing-season soil CO2 (Breecker and Retal- paction by overburden (eq. [1]). Because soil car- lack 2014). The carbonate paleohyetometer is not bonate is primarily a microbial precipitate (Monger suitable for Devonian or older paleosols because even et al. 1991), this proxy can be applied to pre- Devonian forests were less productive at similar iso- Devonian paleosols with other evidence of life. Es- hyets than modern ecosystems (Retallack and Huang timates of the mean annual range of precipitation 2011). are 46 5 22 mm for the Loysburg silt loam and 42 5 Estimates of paleoprecipitation based on CIA-K 22 mm for the Morrison silty clay loam (both late give subhumid mean annual precipitation: 607 5 Katian) and are 45 5 22 mm for the Potters Mills 182 mm for Loysburg silt loam and 810 5 182 mm clay (Hirnantian). This is moderately seasonal but for Morrison silty clay loam (both late Katian), 827 5 not monsoonal because it is less than a third of the 182 mm for Potters Mills clay (Hirnantian), and 720 5 mean annual range of precipitation of 141–2510 mm 182 mm for Beans Gap clay (early Llandoverian). for modern monsoonal soils (Retallack 1991, 2005). These new estimates are preferable to earlier semi- Like the monsoonal soils and paleosols, the Ordo- arid estimates (Feakes and Retallack 1988) because vician ones have alternately precipitated carbonate the Juniata Formation formed a coastal plain within and hematite (fig. 8C), an indication of alternating 100 km of the sea (fig. 1C) and high marine pro- dissolution-evaporation and reduction-oxidation (Re- ductivity fueled by mineral nutrients is inferred for tallack 1991). Furthermore, monsoonal carbonate high marine carbonate d13C values (Hirnantian Iso- nodules are often unusually abundant and small (2– tope Carbon Excursion [HICE] of Young et al. 2010; 4 mm) rather than scattered and large (2–5 cm; Re- Ghienne et al. 2014) and black shales (Armstrong et al. tallack 2005). 2009a). Paleotemperature. A useful paleotemperature proxy At the 827-mm isohyet in Australia today, trees for paleosols of Óskarsson et al. (2009) is based on are 23 5 2.3 m tall, Bk horizons are 81 5 5 cm deep modern soils under tundra vegetation of Iceland. (Retallack 2012), and late-growing-season produc- This linear regression between mean annual tem- 5 tivity is 3448 766 ppm CO2 (eq. [6]). Ordovician perature (T [7C]) and the chemical index of weather- nematophytes of the Juniata Formation are only 5% ing (I p 100 # mAl2O3/(mAl2O3 1 mCaO 1 mNa2O) of this height, following the maximum likely scaling in molar proportions is given in equation (9) (R2 p presented by Retallack and Landing (2014). Ordovi- 0.81; SE, 0.47C). This proxy is useful for pre-Devonian cian Bk depth was only 33% of modern and second- paleosols with evidence of life because of hydrolytic ary productivity was only 44% of modern for this weathering reactions driven by soil respiration (Paton isohyet: evolutionary advances since the Ordovician et al. 1995). Furthermore, it is based on vegetation have greatly increased terrestrial productivity and with large proportions of nonvascular land plants, their effects on weathering (Berner 1997, 2006). unlike other paleothermometers based on forested Paleoclimatic Seasonality. Driese and Foreman soils (Sheldon et al. 2002; Gallagher and Sheldon (1992) described extensive slickensides in the Beans 2013). Gap and Powell Mountain clay paleosols and com- pared them with the structure of , which are T p 218.5N 1 17.3. (9) swelling clay soils with a marked dry season (Soil Survey Staff 2010). No mukkara structure or gilgai These calculations give temperate paleotempera- microrelief of Vertisols (Paton 1974) was noted in tures of 9.6750.47C for the Loysburg silt loam and these paleosols, which are exceptional for the Juniata 9.0750.47C for the Morrison silty clay loam (both Formation and formed at a stratigraphic level of Si- late Katian), 10.9750.47C for the Potters Mills clay lurian marine transgression (fig. 3). (Hirnantian), and 10.3750.47C for the Beans Gap Seasonality of precipitation in soils, defined as clay (early Llandoverian). These estimates are cool the difference in mean monthly precipitation of the but compatible with equally surprising evidence of driest versus the wettest month (M [mm]), has been Cambrian shoreline ice in Wisconsin (Runkel et al. shown by Retallack (2005) to be related to the thick- 2010), Cambrian tillites in and Ire- ness of soil with carbonate or to the distance between land (Landing and MacGabhann 2010), and Late Or- 530 G . J . R E T A L L A C K dovician tillites and periglacial features of Saharan much of it in soil pores beyond the rhizosphere is Africa (Le Heron and Craig 2008; Finnegan et al. 2011). from microbial respiration (Monger et al. 1991), allow- These paleotemperatures are also cool for paleo- ing use of these proxies for pre-Devonian paleosols. latitudes determined by the Point Tracker program of Scotese (1997): 277S for the Ordovician site in d13 2 d13 2 p Cs 1.0044 Cr 4.4 (10) Tennessee and 257–267S for sites in Pennsylvania. Pa Pr , d13C 2 d13C These are latitudes like those of modern Sandgate, a s d13 1 near Brisbane, Australia, which has a mean annual Cc 1000 (11) d13C p 2 1000, temperature of 19.77C(for1913–1976) and a mean s ½(11.98 2 0.12T)=1000 1 1 sea surface temperature of 237C(for1961–1990; ∂13 p ∂13 1 (12) Australian Bureau of Meteorology 2015). Marine Ca 0.665 Co 6.128. temperatures estimated from d18Oofconodontsin Minnesota and Kentucky are thought to have been Analytical data for this calculation are available

227C (Buggisch et al. 2010), and those from D47 of only for the Potters Mills paleosol, considered co- from Anticosti Island, which was closer to eval with Hirnantian glaciation in Saharan Africa 13 the paleoequator, were 287C (Finnegan et al. 2011). (fig. 3). The isotopic composition of calcite (d Cc) Late Ordovician marine limestones of the Nashville in that paleosol is –5.82‰, and coexisting organic 13 and Cincinnati domes are bryozoan-dominated (rather matter (d Co)is–21.5‰, both versus PDB standard than -dominated) temperate-type carbonates (Hol- (Retallack 1993), which falls within the D13C range land and Patzkowsky 1998). Modeling based on ox- (carbonate-organic) of –14‰ to –17‰ accepted by ygen isotopic composition of marine carbonates Cotton and Sheldon (2012) as well-preserved or- (Armstrong et al. 2009b) shows that tropical waters ganic matter. This isotopic value for organic matter extended south only as far as latitude 227S and that is higher than other Ordovician analyses of Tomescu the intertropical convergence zone extended south et al. (2009), who present no Hirnantian values. Ma- as far as 107S for only a brief interval of the late rine Hirnantian carbonate carbon isotopic values 13 Katian (Boda isotopic event). For most of the Katian (d Cc) are also up to 6‰ higher than before or after- and Hirnantian, tropical waters were no farther south ward (HICE of Young et al. 2010; Ghienne et al. than 107S and the intertropical convergence zone was 2014). Paleotemperature and productivity estimates no farther south than latitude 27N, well north of the for the Potters Mills paleosol are as given in previ- studied sections in Tennessee and Pennsylvania. ous paragraphs. SEs of estimates of atmospheric CO2

Paleoatmospheric CO2. The paleosol paleobarom- using equation (10) were calculated by Gaussian er- eter of Cerling (1991) estimates past atmospheric ror propagation of the five-component SEs (follow-

CO2 (Pa [ppmv]) from the balance within soils at the ing Breecker and Retallack 2014). Atmospheric CO2 depth of carbonate nodules between isotopically light levels calculated using these equations and paleo- 13 soil CO2 (d Cs) compared with isotopically heavy at- temperature of 10.97C (from eq. [9]) for the type Pot- 13 mospheric CO2 (d Ca), and it has been refined for pa- ters Mills clay paleosol is 166 5 80 ppmv, roughly leoproductivity effects (Breecker and Retallack 2014). equal to glacial levels of 180 ppmv (Stocker A variety of terms in the governing equation (10) et al. 2013). below are derived from other analyses and transfer This CO2 estimate of 166 5 80 is much lower than functions. The isotopic composition of soil CO2 4035 ppmv from mass balance modeling in 10-m.yr. 13 13 (d Cs) comes from that of pedogenic carbonate (d Cc) steps (Berner 2006)—too long to incorporate the using temperature-dependent (T [7C]) fractionation known isotopic volatility of the Hirnantian (LaPorte from equation (11) (R2 p 0.93; SE, 0.28‰; Romanek et al. 2009; Ghienne et al. 2014), which lasted only et al. 1992). The isotopic composition of associated 1.4 m.yr. (Gradstein et al. 2012). This estimate is also 13 (d Co) is used as a proxy for re- much lower than 2000 ppm, which is a minimum CO2 13 spired soil CO2 (d Cr) as well as to calculate the concentration associated with the “big five” mass 13 isotopic composition of atmospheric CO2 (d Ca) fol- extinctions (Retallack 2013b). This estimate is also lowing equation (12) (R2 p 0.96; SE, 1.73‰), derived below 3000 ppm (2240–3920 ppm), which Royer from data on modern bryophytes (Fletcher et al. (2006) considers the upper limit of CO2 permitting

2005). Finally, partial pressure of respired CO2 in soil glaciation in a computer-modeled Hirnantian world

(Pr) is derived from its known relationship with depth of lower solar luminosity and no land plants. This lat- to carbonate nodules (Do, corrected for compaction ter assumption is doubtful considering evidence for life in paleosols using eq. [1] in modern soils expressed in Late Ordovician paleosols presented here (fig. 7). in eq. [6]). Some of this respired CO2 in desert soils Estimated paleoatmospheric CO2 from the Hir- studied by Cerling (1991) is from root respiration, but nantian Potters Mills clay paleosol is also lower than Journal of Geology ORDOVICIAN PALEOSOLS 531

a previous estimate of 4670-ppmv atmospheric CO2 (Retallack and Feakes 1988; Driese and Foreman by Yapp and Poths (1992, 1993) for a lateritic paleosol 1991, 1992; Yapp and Poths 1992, 1993, 1994) add pa- in the in Wisconsin. The age of this leogeographic coverage and are similar as far as they paleosol is latest Katian considering marine fauna in go (squares in fig. 12A). underlying rocks (Royer 2006) and carbon isotope Hirnantian Glaciation. The evolution of early land stratigraphy of overlying limestones (McLaughlin plants and increased depth and degree of weathering et al. 2013). The new low atmospheric CO2 estimate and biomass on land is revealed by plantlike fossils for the Potters Mills clay is compatible with Hir- (fig. 7) and thicker Katian paleosols of increased nantian continental glaciation in Saharan Africa (Le productivity (fig. 12A). Furthermore, the pedogenic Heron and Craig 2008), and the greenhouse crisis paleobarometer applied to one particularly well- estimate of Yapp and Poths (1992, 1993) is compat- understood Hirnantian paleosol reveals atmospheric ible with a negative carbon isotope anomaly before CO2 levels of only 166 5 80 ppmv, well below mod- Hirnantian glaciation (Armstrong et al. 2009b; Young eled thresholds for Ordovician glaciation (Royer 2006). et al. 2010). This negative anomaly was also at a time Thus, there is direct evidence from paleosols for of marine mass extinction (Fan et al. 2009; Harper drawdown of atmospheric CO2 by increasingly deep et al. 2014). The conundrum of continental glaciation and profound chemical weathering and biomass ac- at a time of high atmospheric CO2 (LaPorte et al. 2009) cumulation by Late Ordovician nonvascular land is avoided by this new evidence of short-term paleo- plants (Retallack 2000, 2003; Lenton et al. 2012). atmospheric volatility (fig. 12). The Ordovician and Silurian paleosol record also has temporal resolution adequate to show short- term volatility in paleoclimate and paleoproductiv- Conclusions ity on land (fig. 12A). A lateritic late Katian paleosol Paleosol Records. Paleosol records of atmospheric from Wisconsin is direct evidence of 4670-ppmv at- changes with high temporal resolution are used here mospheric CO2 (Yapp and Poths 1992, 1993), which to disentangle conflicting indications of the two is considered a very short-lived greenhouse spike, main issues of this study: Late Ordovician glaciation now correlated (following McLaughlin et al. 2013) and mass extinction. Ordovician paleosols have been with the Boda carbon isotopic shift (Armstrong et al. challenged (Davies et al. 2010), but field and petro- 2009a) and deep-calcic paleosols of the latest Katian graphic features of ancient soils are documented (Retallack 2009b). This is the reason why carbon here: clear pedofabric down from erosional surfaces cycle modeling in 10-m.yr. steps revealed 4035-ppmv

(fig. 2A–2C), including clay skins (fig. 8C,8E), sepic CO2 during the Late Ordovician (Berner 2006), as this plasmic fabric (fig. 8E), and replacive micrite (figs. 7C, greenhouse spike and lower CO2 levels of the Hir- 8C). Ordovician paleosols also have isotopically de- nantian glaciation lasted less than 1.4 m.yr. (Grad- pleted carbonate for both d13Candd13C, with tight stein et al. 2012). correlation between stable isotopic values charac- Hirnantian glaciation has also been linked with teristic of soils and paleosols (fig. 10). Finally, geo- increased silicate weathering from Taconic moun- chemical mass balance reveals that they collapsed tain uplift (Kump et al. 1999) and with amalgamating and lost nutrient cations like soils and Miocene pa- and expanded microcontinents (Saltzman and Young leosols (fig. 9) rather than dilating and gaining nu- 2004). The mountainous source for the Juniata For- trient cations like sediments and sedimentary rocks mation persisted through to the early Devonian as (Brimhall et al. 1992). Thus, evidence presented here source for the Shawangunk and Bloomsburg Forma- supports the interpretation of bioturbated caliche- tions (fig. 1C), which lack any evidence of glaciation bearing beds of the Juniata Formation as paleosols (Retallack 1985; King 1998; Driese and Mora 2001). (Retallack 1985, 1993, 2000; Retallack and Feakes The 10-m.yr. trough in marine carbonate 87Sr/86Sr 1987; Feakes and Retallack 1988; Driese and Fore- (fig. 12C) marks lower riverine input relative to ma- man 1991, 1992; Davies and Gibling 2010) rather rine hydrothermal activity during the Katian and than marine beds (Davies et al. 2010). has been linked to decreased weathering at a time of Long sequences of paleosols in the Tumblagooda active volcanism (Shields et al. 2003). Coming at a Sandstone, Mount Chandler Sandstone, and Indul- time of increased soil thickness and soil respiration kana Shales in aridland outcrops and drill cores of (fig. 12A), this lower fluvial contribution to marine Central and South Australia (Retallack 2009b) give carbonate 87Sr/86Sr may reflect not decreased silicate high-resolution Ordovician and Silurian records of pa- weathering but increased ground cover and water re- leoclimate and paleoproductivity (circles of fig. 12A). tention by bryophytic mats. Patchy outcrops of paleosols in the Juniata and Neda Late Ordovician Mass Extinction. The estimate of

Formations of Pennsylvania, Tennessee, and Wisconsin 4670-ppmv atmospheric CO2 by Yapp and Poths 532 G . J . R E T A L L A C K

(1992, 1993) for a lateritic paleosol in the latest Katian of the Argentine Precordillera (Ortega et al. 2008; Neda Formation of Wisconsin (McLaughlin et al. González-Menéndez et al. 2013). Thus, Late Ordo- 2013) coincides with the main pulse of Late Ordovi- vician mass extinction as a preplay of the Late cian marine mass extinction just below the Nor- Permian extinction by atmospheric pollution of flood malograptus extraordinarius graptolite zone forming basalt eruptions remains a viable hypothesis. the base of the Hirnantian, which was followed by Alternative hypotheses of Late Ordovician mass a second pulse of extinction in the late Hirnantian extinction include Hirnantian cooling (Harper et al. lower Normalograptus persculptus zone (Harper et al. 2014), but the first wave of the extinction occurs be- 2014). These extinction levels are also spikes in fore glacial onset indicated by positive marine carbon depth to paleosol carbonate and inferred productiv- isotopic values and glacial facies correlations (Ghi- ity (fig. 12A) and spikes in weathering inferred from enne et al. 2014), and the second wave of extinction marine carbonate 87Sr/86Sr (fig. 12C). Both pulses of occurs during waning phases of glaciation (Arm- extinction were most severe in high-latitude Gond- strong et al. 2009a; Young et al. 2010). Many mass wanan terranes but were less marked in the tropics extinctions were at times of abrupt global warm- (Harper et al. 2014) and not yet detectable in the rec- ing, not cooling (Keller 2011; Retallack 2013b;Bond ord of early land plant (Steemans et al. 2010; and Wignall 2014). Another plausible culprit for Late Edwards et al. 2014). The extinction levels are both Ordovician mass extinction is oceanic anoxia (Har- at abrupt negative carbon isotopic excursions (pro- per et al. 2014), indicated by black shales with ductivity crises) and positive oxygen isotopic excur- reduced iron, high molybdenum, abundant pyrite sions (warming events) in marine carbonate flanking framboids, and low sulfide sulfur isotopic values (Arm- a ca. 1-m.yr. positive excursion (HICE of Armstrong strong et al. 2009a; Fan et al. 2009). These black shales et al. 2009a; Young et al. 2010). Comparable spike- postdate the mass extinction and coincide with the like perturbations are also known for a succession of positive marine carbon isotopic excursion (HICE of Permian and Triassic extinctions, which rise to the Armstrong et al. 2009a; Young et al. 2010). The de- level of mass extinction when inferred atmospheric layed onset of Hirnantian euxinia is evidence that

CO2 from paleosol and paleobarometers is oceanic anoxia was a consequence of CO2 greenhouse more than 2000 ppmv (Retallack et al. 2011; Retal- spikes rather than a first cause of mass extinctions lack 2013b). Abrupt negative carbon isotopic excur- (Retallack 2013b). There is no clear evidence of Hir- sions have been explained by atmospheric pollution nantian extraterrestrial impact as an extinction with massive amounts of isotopically light CH4, method,eveninsectionsinwhichithasbeensought which within a decade oxidizes to a transient CO2 (Wang et al. 1993). A swarm of is known greenhouse, raising planetary temperature and pre- from the latest Dapingian (Middle Ordovician) lime- cipitation to deepen paleosol carbonate horizons and stones in but is not associated with mass leave isotopic signatures in both organic matter and extinction (Schmitz et al. 2008). carbonate of paleosols (Retallack et al. 2011b; Re- tallack 2013b). The source of such large injections of isotopically light carbon into the atmosphere may ACKNOWLEDGMENTS have been thermogenic cracking of coal or other organic matter around the feeder dikes of flood ba- R. Beerbower first suggested the Potters Mills out- salts (Retallack and Jahren 2008). The Katian nega- crops as potential paleosols, and C. Feakes helped tive carbon isotopic excursion stimulated Lefebvre with laboratory studies to develop their geochemical et al. (2010) to postulate a Late Ordovician flood and petrographic potential. S. Driese helped with basalt, and this prediction may be represented by data and provided a guided tour of the Beans Gap widespread tholeiites in the late Katian Yerba Loca paleosol. I thank D. Breecker, L. Nordt, M. Droser, and Alcaparrosa Formations of the Cuyana Terrane and D. Bottjer for useful discussions.

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