A global biogeochemical perturbation across the Silurian–Devonian boundary: Ocean–continent–biosphere feedbacks

Palaeogeography, Palaeoclimatology, Palaeoecology 276 (2009) 244–254

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Palaeogeography, Palaeoclimatology, Palaeoecology j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / p a l a e o

A global biogeochemical perturbation across the Silurian–Devonian boundary: Ocean–continent–biosphere feedbacks Krzysztof Małkowski, Grzegorz Racki ⁎ Institute of Paleobiology, Polish Academy of Sciences, Twarda 51/55, 00-818 Warszawa, Poland

a r t i c l e

i n f o

Article history: Received 15 August 2008 Received in revised form 6 February 2009 Accepted 1 March 2009 Keywords: Silurian–Devonian boundary Carbon isotopes Biogeochemical cycles Terrestrial plants Sea level Climate

a b s t r a c t The large-scale global biogeochemical perturbation across the Silurian–Devonian (S–D) boundary, recorded in the major positive excursion of the δ13C time curve (Klonk Event) of an amplitude of 2.5 to 3.0‰ (max. 4.0‰) in Europe and up to 5.0‰ in North America, reflects a unique combination of palaeogeographic, biogeochemical and evolutionary processes in the late Caledonian geodynamic setting. The steady sulfur isotopic ratios show an overall stability of the S–D oceanic geochemical system as a whole and do not indicate any synchronous changes in anoxic deep oceanic and sediment processes. Therefore this led us to a hypothesis that the crucial changes that contributed to the recorded carbon cycling turnover are related to rapidly evolving ocean–continent–biosphere feedback. Coastal zones of the latest Silurian epicontinental seas accumulated considerable quantities of organic carbon from early vascular vegetation, which explosively expanded to inhabit vast near-coastal shallows and deltas. Large primary production of these early terrestrial plants and rapidly enhanced sedimentary burial of organic carbon were also responsible for CO2 drawdown, which have resulted in reversed-greenhouse effect and a global climatic cooling tendency. This feedback was blocked when the sea level gradually dropped and led to shrinking of the Silurian epicontinental seas and the growing climatic deterioration during the S–D transition limited primary production. Furthermore, continued processes of regressive abrasion and erosion limited the storage of organic carbon, as well as the efficiency of the carbonate factory. During the following Early Devonian greenhouse interval, marine regression and active latest Caledonian tectonism promoted progressive weathering of the sedimentary organic matter. The oxidation of C-rich deposits caused the subsequent growth of CO2 levels in the Early Devonian, culminating in the warming of global climate. © 2009 Elsevier B.V. All rights reserved.

Contents 1. 2. 3.

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Silurian–Devonian boundary isotopic event . . . . . . . . . . . . . Global exogenic system across the Silurian–Devonian transition . . . . . . 3.1. Palaeogeography and geotectonic setting . . . . . . . . . . . . . 3.2. Sea-level changes . . . . . . . . . . . . . . . . . . . . . . . . 3.3. Terrestrial vegetation, continental weathering and climate . . . . . 4. Discussion and implications . . . . . . . . . . . . . . . . . . . . . . . 4.1. Current understanding of the S–D carbon isotopic excursion . . . . 4.2. Proposed extended scenario . . . . . . . . . . . . . . . . . . . 4.3. Time-scale approximation of the S–D carbon cycle perturbation . . 4.4. Comparison with other middle Paleozoic carbon cycle perturbations 5. Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Acknowledgments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

⁎ Corresponding author. Tel.: +48 22 697 88 50; fax: +48 22 620 62 25. E-mail address: [email protected] (G. Racki). 0031-0182/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.palaeo.2009.03.010

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1. Introduction The Barrandian sedimentary succession of Silurian–Devonian (S–D) boundary represents a geological system boundary of international importance, when following the choice of the base of the Monograptus· uniformis graptolite zone it was approved as the first global stratotype section and point (GSSP) at Klonk in the Prague Basin of the Czech Republic (Martinsson, 1977; www.stratigraphy.org/sildev.htm). A review of the significant biotic changes recorded at this reference locality is given by Chlupáč and Hladil (2000), Chlupáč and Vacek (2003) and Brocke et al. (2006). However, worldwide biodiversity shift corresponds to the generally evolutionarily ‘quiet’ S–D boundary (see summary in Walliser, 1996, Copper, 2002, House, 2002, Bambach, 2006). Only a fifthorder global bio-event has been proposed by Kaljo (in Barnes et al., 1996, p. 328), characterized by Walliser (1996, p. 228) as “a minor, relatively gradual but globally traceable event”, exemplified spectacularly by two acme episodes of floating Scyphocrinites that resulted in the common deposition of coarse-grained crinoid limestones (e.g., Vacek, 2007; Lubeseder, 2008). Nevertheless, the Klonk Event, which was paired with a oceanographic turnover negatively influencing many principal fossil groups (such as graptolites, conodonts, chitinozoans, radiolarians, brachiopods and trilobites), was established by Jeppsson (1998, p. 253; see also Afanasieva and Amon, 2006) within the S–D transition. In addition, a coeval major perturbation in carbon cycling has been recently documented (Saltzman, 2002; Buggish and Mann, 2004; Buggisch and Joachimski, 2006; Gill et al., 2007; see also Veizer et al., 1999), being, in terms of the maximal amplitude of its δ13C excursion (>4‰), among the largest well-known biogeochemical turnovers during the Paleozoic, as listed by Holser et al. (1996). When combined with the first occurrence of the cosmopolitan conodont species Icriodus·woschmidti and the graptolite Monograptus·uniformis, the considerable carbon isotope curve peak provides a consistent chemostratigraphic tool for worldwide recognizing the Silurian– Devonian boundary (Williams and Saltzman, 2004). We have recently elaborated new δ13C and δ18O curves, derived from continuous and extended sections in the Dniester Valley, southwestern Ukraine, Podolia (Małkowski et al., in press), as a basis for highresolution stratigraphy of the whole-rock isotopic systematics. The refined chemostratigraphic trends are an all-embracing starting point to unraveling causes of large-scale biogeochemical changes in the Silurian– Devonian exogenic system, which were recently considered by Saltzman

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(2002) and Buggish and Mann (2004). The isotopic perturbation in the global carbon cycle across the S–D boundary, in particular a rapid enhancement of carbon burial, is considered herein in the broad context of profoundly evolving palaeogeographic, geochemical and biospheric processes in the late Caledonian geodynamic setting. 2. The Silurian–Devonian boundary isotopic event Since study of Hladíková et al. (1997, 1999) in the Barrandian basin, including the global stratotype succession, it is well-known that the S–D boundary can be correlated worldwide with a major ‘heavy carbon’ signal (Klonk isotope event of Buggisch and Joachimski, 2006). The positive δ13C excursion was firstly reported from Europe (Buggisch and Joachimski in Schönlaub et al., 1994; Porębska and Sawłowicz, 1997; summarized by Buggish and Mann, 2004) and Australia (Andrew et al., 1994). In particular, δ13C pattern from the·Icriodus hesperius (= I.·woschmidti) conodont Zone in the Jack Limestone of Queensland, Australia demonstrates a distinct and sustained increase from 6.5 to 8‰, but a broader stratigraphic context of this anomalous 13C enrichment in ca. 90 m-thick shallow-water succession is unknown. In addition to these reports from Gondwana, Saltzman (2002) extensively described the highest peak δ13Ccarb values near the S–D boundary from three Laurentian carbonate successions (Central Nevada, Oklahoma and West Virginia; Fig. 1). The maximum δ13C value is within the lowermost Lochkovian in the outer shelf Birch Creek II locality in central Nevada. In the lower part of the section, δ13C values begin near 1‰, and rise gradually in a fluctuating trend to heavier values, finally approaching an extreme height of 5.84‰ at the top of the succession, ca. 20 m above the S–D boundary. The higher Lochkovian record of the region is distinguished by relatively uniform but 13C depleted ratios near 0‰. On the other hand, the carbon isotopic peak (5.75‰) within an extended positive excursion is depicted ca.16 m below the S-D boundary at Strait Creek, West Virginia (Gill et al., 2007, Fig. 4 therein; see also data from Saltzmann, 2002 in Fig. 1) what suggests severe diagenetic modifications of the record over the crucial S–D boundary interval. Buggish and Mann (2004) extensively documented δ13C chemostratigraphic pattern in inorganic and organic carbon of several welldated S–D localities in the Prague Syncline (Barrandian), the Carnic Alps, the Montagne Noire, and the Cantabrian Mountains (Fig. 1). The highest values reach almost 4.0‰ for δ13Ccarb and 26.0‰ for

Fig. 1. The positive δ13C excursion curve at the Silurian–Devonian transition, as recorded in North American (after Saltzman, 2002) and European successions (open symbols include data from literature and from diagenetically altered limestones), compiled by Buggish and Mann (2004, Fig. 11, modified; with kind permission from Springer Science+Business Media).

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δ13Corg. Many deeper-water sections demonstrate rhythmic deposition of dark to black shales and limestones, and δ13Ccarb values frequently altered in these sediments owing to the input of 12C enriched carbon as result of bacterial sulfate reduction during the diagenetic stabilization of the carbonate oozes. However, the positive correlation of δ13Ccarb and δ13Corg confirms that these trend mirror primary carbon isotope ratios. In general terms, the excursion during the·I.·woschmidti·–·I.·postwoschmidti zones reveals an amplitude of 2.5 to 3.0‰ (max. 4.0‰) in Europe, i.e. less prominent than in North America. After the ‘heavy carbon’ signal, a long-lived δ13Ccarb minimum with values between 0 and 1.5‰ is documented higher in the Lochkovian (Fig. 2). The major biogeochemical perturbation in carbon cycling is reliably confirmed in well-exposed carbonate successions of Podolia (Małkowski et al., in press), correlated with brachiopod calcite data of Veizer et al. (1999; see http://www. science.uottawa.ca/geology/isotope_data/). The accelerated ‘heavy carbon’ maximal enrichment corresponds to the S–D boundary (excursion 1 in Fig. 2), although the peak of 4.2‰ in bulk-rock samples (and 4.7‰ in brachiopod calcites) occurs slightly above the graptolite-defined boundary. Comparison with the δ13C values of Buggish and Mann (2004) reveals more negative values for the stratigraphically oldest and, especially, youngest Podolian samples (Fig. 2), confirmed by data of Veizer et al. (1999). However, comparable C13 depletions are reported from some successions (Kaljo et al., 1998, Fig. 7; Kaljo et al., 2003, Fig. 3; Saltzman, 2002, Fig. 3; Buggish and Mann, 2004, Figs. 8 and 9), and manifested also by another global isotopic curve for the Prodoli by Azmy et al. (1998, Fig. 9), based on brachiopod calcites. The differences may reflect partly decoupled, regional carbon isotopic ratios in transiently semi-restricted epeiric basins, and may record such regional factors controlling carbon cycling as input of weathered carbon and the oxidation of organic matter (e.g., Wigforss-Lange, 1999; Saltzman, 2001; Panchuk et al., 2005, 2006; Immenhauser et al., 2008). In addition, in the deep-sea clayey–siliceous succession of

Sudetes (SW Poland), the S–D δ13C excursion measured in organic carbon is as low as 1.5‰ (Porębska et al., 1999). In summary, a major S–D boundary biogeochemical perturbation in carbon cycling, resulting in marine δ13C values as high as 6‰, is now reliably demonstrated in many successions worldwide, and belongs to the largest turnovers in carbon cycling during the Paleozoic (Holser et al., 1996). The similarity of key signature in distant palaeotectonic and palaeogeographic settings argues strongly against the profound δ13C excursion being exclusively an artifact of post-depositional process diversity. Thereby, this large-scale 13C enrichment episode of Klonk Event is reasonably seen by Saltzman (2002) and Buggish and Mann (2004) as a signature of original variations in the carbon isotope ratios of the global oceanic reservoir. 3. Global exogenic system across the Silurian–Devonian transition Understanding of the above described S–D perturbation in the global carbon cycle requires a broad contextual account of interacting palaeogeographical, climatic and biospheric settings. Two significant geochemical proxies should be taken into consideration when evaluating any Silurian–Devonian ecosystem changes, because the exogenic (litho-, hydro-, atmo-, biosphere) system was obviously driven by global tectonic forces on geological time scales, summarized below, via its control of (bio)geochemical cycling (e.g., Prokoph and Veizer, 1999; Veizer et al., 1999; Godderis et al., 2001; Kump et al., 2003): 1. Isotopic ratios of strontium (87Sr/86Sr) in Silurian–Devonian oceanic water co-vary with the composite carbon isotopic time curve, as exemplified by the significant Ludlow–Lochkovian positive 87Sr/86Sr plateau around 0.7087 (Denison et al., 1997; Azmy et al., 1999; Veizer et al., 1999). Ruppel et al. (1996) concluded that the acme of the long-term Silurian–Devonian Sr isotopic cycle occurs near the S–D boundary. In fact, Sr isotopic

Fig. 2. Chemocorrelation of the Podolian δ13C curve documented by Małkowski et al. (in press; Figs. 5–7, with kind permission from Cambridge University Press with the proposed worldwide carbon isotope stratigraphic trends across the Silurian–Devonian transition compiled by Buggish and Mann (2004; with kind permission from Springer Science+Business Media; see Fig.1).

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data, refined by Fryda et al. (2002) at the Klonk stratotype section, suggest the existence of a high-order oscillation at the S–D boundary, with a maximum ratio value of 0.7088. In general, this large-scale excursion is a reflection of enhanced erosion of exposed crystalline crust in uplifting continental regions, particularly the Caledonian Orogen (van Geldern et al., 2006, Fig. 7; see below). 2. Sulfur isotopic evolution of Middle Paleozoic seawater is less refined. The global trend based on the analysis of structurally substituted sulfate in carbonates exhibits an Ordovician–Silurian plateau with δ34S around 26‰ followed by a decrease of the sulfur isotopic ratios during the late Lochkovian and Pragian, below 18.5‰, although a positive shift is possibly recorded near the S–D boundary in sparse evaporite-based data (Kampschulte and Strauss, 2004, Figs. 3 and 5 therein). Conversely, the new refined S isotopic curve from West Virginia (Gill et al., 2007, Fig. 4 therein) exhibits a fluctuating negative trend but markedly below the key S–D transition, where consistently high values (26–29‰) are confirmed. When compared with other geochemical proxies, the sulfur isotopic composition remains surprisingly stable and relatively high, and the positive values of the oceanic sulphate sulfur isotopic signal depended on a continous deposition of isotopically light sulfides in the lower oxygen-deficient oceanic zone. Hence, the profoundly

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evolving oceanographic processes did not influence deeper dysaerobic zones, i.e., an extensive mixing in the postulated stratified World ocean was absent at that time (compare Saltzman, 2005). 3.1. Palaeogeography and geotectonic setting During the Silurian, continents commenced to emerge at the expense of epeiric seas, and the area of the marine shelves diminished from about 40% in the Early Silurian to less than 30% at the beginning of the Devonian (Wilde et al., 1991, Fig. 11; see also Brunton et al., 1998; Kiessling et al., 2003; Miller et al., 2005). The late Ludlow to late Lochkovian interval was dominated by diachronous compression and uplift associated with global-scale Caledonian orogenic movements, exemplified primarily by consolidation of the Arctic–North Atlantic fold belt leading to the vast Old Red continent (Scanian phase of McKerrow et al., 2000; see Khain and Seslavinsky, 1996, pp. 137–152; Ziegler, 1990; Golonka, 2000; Dewey and Strachan, 2003; Ebert and Matteson, 2003; Ford and Golonka, 2003; Sasseville et al., 2008). The geodynamic evolution is clearly recorded in strontium isotopic ratios (a rough proxy for the runoff of continental weathering products delivered from emergent continents into marine basins; see Ruppel et al., 1996; Fig. 4; Veizer et al., 1999, Figs. 5 and 6; Ebert and Matteson, 2003; van Geldern

Fig. 3. Summary diagram of the principal generalized geochemical proxies and other ecosystem indicators used for distinguishing the two principal global ecosystem states (Fig. 4) and for understanding Silurian–Devonian biogeochemical evolution; arrows show subordinate and/or regional only events while line thickness reflects resolution of the data, and differently sized rhombs corresponds to global bioevents (Kaljo et al., 1996; Walliser 1996). Based on generalized curve of Buggish and Mann (2004, Fig. 13) for carbon isotopes (see Fig. 2 for event numbers; for the more detailed Pridoli curve see Fig. 8 in Azmy et al., 1998, Figs. 4, 5 and 7, 8 in Saltzman 2002, and Fig. 3 in Kaljo et al., 2003); the Web Site Ottawa database of Veizer et al. (1999; www.science.uottawa.ca/geology/isotope_data/) for carbon, oxygen and strontium isotopes, slightly modified for timing of the major S–D excursion (see Fig. 5 in Małkowski et al., in press), see also Fryda et al. (2002, Fig. 3) and Jeppsson et al. (2007, Figs. 5 and 6) for strontium isotopic pattern; Edwards et al. (2000, Fig. 1) and Edwards and Wellman (2001, Tables 2.1–2.2) for plant diversity and distribution, respectively (see Raymond et al., 2006 for updated Silurian data); Berner (2003, Fig. 5) for approximated atmospheric composition (RCO2 refers to the ratio of CO2 in the past to that for the pre-industrial present; Fig. 5); and Ross and Ross (1996, Fig. 1) and Walliser (1996, Fig. 3) for sea-level curve (see also Raymond and Metz 1995; Azmy et al., 1998; Berner and Kothavala 2001; Crick et al., 2001; Royer et al., 2004). In general palaeogeographic terms (sensu Heckel, 1972), the epicontinental state A refers to the peneplaned Earth with extensive carbonate shelves in eustatic highstand conditions, while the pericontinental mode B corresponds to the Earth with uplifted continents, narrowed shelves and lowered sea levels. Note that the δ13C excursion over the Silurian–Devonian boundary corresponds mostly to the passage interval between the two states of exogenic system (see Fig. 4) and the regionally disturbed eustatic lowstand setting.

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et al., 2006, Fig. 7), and also in contrasting sea-level histories in particular regions (see below). On the other hand, only some regions are marked by the S–D volcanic activity (e.g., western Avalonia – McKerrow et al., 2000, Fig. 2; Marriott et al., 2009; Baltic basin – Hints et al., 2008; see also summary in Khain and Seslavinsky, 1996, p. 368). Therefore, a new exogenic mode predominated in Early Devonian (Figs. 3 and 4), generally reflecting higher relative elevations of assembled continental masses, especially uplift of the Caledonides and their foreland basins, and extensive drainage systems paired with a development of mainly non-marine facies (e.g., Ziegler, 1990; Dewey and Strachan, 2003). 3.2. Sea-level changes As summarized by Buggish and Mann (2004, p. 537–538) and Brocke et al. (2006, p. 38), the eustatic setting of the S–D passage is ambivalent and is likely obscured by synsedimentary tectonics, as both regression and transgression are alternately indicated for different regions (see summary in Johnson 1996, Fig. 2). For example, the magnetosusceptibility record for the S–D stratotype section indicates a short-lived deepening in the latest Pridoli and an equally brief regressive episode in the earliest Lochkovian, superimposed on an overall transgressive trend (Crick et al., 2001, Fig. 4; compare Vacek, 2007); the deepening tendency is also visible in the Podolian sedimentary record (Skompski et al., 2008; Małkowski et al., in press). However, most compiled sea-level curves show a low-order eustatic lowstand over the Silurian–Devonian boundary (Kaljo et al., 1996; Walliser, 1996; Johnson et al., 1998; House and Gradstein, 2004; Melchin et al., 2004; Miller et al., 2005; Johnson, 2006; Haq and Schutter, 2008; see Fig. 3), with a eustatic drop approximately 50 m

below the shelf margin (Algeo and Seslavinsky, 1995, Fig. 12; Ross and Ross, 1996, Fig. 1a), what caused common hiatuses and erosional gaps (see e.g., Fig. 2 in McKerrow et al., 2000). 3.3. Terrestrial vegetation, continental weathering and climate By the end of the Silurian, a series of evolutionary innovations had led to the initiation of early plant-dominated terrestrial ecosystems (Kenrick and Crane, 1997; Driese and Mora, 2001; Edwards and Wellman, 2001; Raymond et al., 2006; Marriott et al., 2009; see also review of freshwater record in Martín-Closas, 2003), termed as the Phase 1 of land colonization by Ward et al. (2006). This is manifested by still poorly timed advent of niche-partitioning lycopsids, zosterophyllids and trimerophytes, with the first macroscale roots, distributed over all Lochkovian continents (Zosterophyllum Zone of Edwards et al., 2000; see Raymond and Metz, 1995; Driese and Mora, 2001; Edwards and Wellman, 2001; Edwards and Richardson, 2004; Raymond et al., 2006; Steemans et al., 2007; Wang et al., 2007; Hillier et al., 2008; Spina and Vecoli, 2008). As summarized recently by Gensel (2008): “in mid-Late Silurian/Devonian strata, certain benchmark events in early plant evolution, such as acquisition of complex form, larger size, rooting structures, leaves, and seeds, are now known to occur earlier in time than the earlier postulated Middle or Late Devonian“. In particular, a commencement of much higher productivity per plant and possibly greater impact on substrates are linked with the large-sized Late Silurian lycophyte genus Baragwanathia, preserved as 40 cm long fragments (Edwards and Wellman, 2001, p. 23). The oldest coaly lenses or horizons are first reported from Late Silurian epoch as well (e.g., from Ludlow by Strother, 1988).

Fig. 4. Idealized presentation of continental reconstructions for two principal states of the late Caledonian exogenic system across the Silurian–Devonian transition (based on Heckel, 1972, Fig. 1; see also Fig. 3). Latest Silurian timespan is marked by explosive rise of primary production within shrinking epeiric seas due to progressive expansion of early vascular vegetation into extensive near-coastal shallows and deltas (mode A). Intensified end-Caledonian movements and continental uplift promoted gradual palaeogeographic evolution toward predominantly pericontinental mode B during the Lochkovian age. Hence, the biogeochemical equilibrium shifted from effective organic matter burial towards long-term reburial and oxidation due to widespread erosion and weathering processes that altogether reduced highly positive carbon isotope ratios to background values due to plentiful C12 influx.

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Thus, the pre-Devonian role of a terrestrial vegetation in riverine weathering runoff and nutrient delivery to marine habitats may be underestimated (e.g., Retallack, 1997; Boucot and Gray, 2001; Panchuk et al., 2005; Strother, 2006), especially in light of the popular model by Algeo et al. (1995). Significant sedimentary storage of organic carbon, that dynamically lowered atmospheric CO2 concentration, is attributed in the scenario to the later Devonian stratified forests (see also Walliser, 1996; Algeo and Scheckler, 1998; Lenton, 2001; Beerling and Berner, 2005). However, as modeled by Berner and Kothavala (2001) and Berner (2003, 2006), and implied from the proxy CO2·record (palaeosols, stomata; Fig. 3 a in Royer, 2006), the broad S–D transition was an interval of high but falling CO2 levels (see Figs. 3 and 5). This pCO2·drop should result in a reversed-greenhouse climatic effect, interrupting a major warming trend in Middle Paleozoic (see also Fig. 3 in Royer et al., 2004). Likewise, increased oxygen production due to increasingly effective organic carbon (Corg) burial added a significant new non-oceanic O2 source that significantly reversed the slowly dropping trend of oxygen levels in the earliest Paleozoic atmosphere, bringing high values crossing the present level (Berner, 2003, 2006; Ward et al., 2006; Algeo and Ingall, 2006; Fig. 5). Consequently, the early vascular plant communities were affected by wildfires in the oxygen-enriched atmosphere (Scott and Glasspool, 2006), evidenced by the oldest known fossil charcoal from Pridoli coastal mudflat and marsh environments (Glasspool et al., 2004). To recapitulate,·the increasing coupling between the wetland– terrestrial biomass and highly efficient organic matter storage is proposed to have initially been effectively working by the latest Silurian. This may be implied from the middle Paleozoic peak of carbon sequestration (Berner, 2003, 2006; Fig. 5), but this does not mean, however, an amplified Corg biomass storage in the continent interiors. It should be stressed that the large masses of organic carbon on the intensively colonized continental margins were mostly destroyed by uplift in the orogenically active late Caledonian setting. In addition, an increasing marine nutrient budget, due to effective riverine nutrient delivery from plant-mediated and tectonicallyenhanced weathered crystalline massifs during the global mountain building (as guided by the strontium isotopic curve; Fig. 3), is presumed (compare with the Eovariscan scenario of Averbuch et al., 2005). Increased phosphorus fluxes are the most crucial precondition for the large-scale nutrient availability and massive uptake of 12C by marine primary producers (Martin, 1996; Kump and Arthur, 1999; Lenton, 2001; Lenton and Watson, 2004). Major turnovers in the

Fig. 5. Main middle Paleozoic carbon isotopic events against compositional evolution of atmosphere and rate of global organic carbon burial (courtesy of R.A. Berner – compiled from Berner (2003), Fig. 2 therein and Berner, 2006, Figs. 18 and 20; see also Algeo and Ingall, 2006). Note irregularly dropping CO2·levels, mostly due to increasing sedimentary sequestration of primary production and tectonically-enhanced biogeochemical impact of early terrestrial biota that culminated across the broad Silurian– Devonian transition (see Ward et al., 2006).

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global marine fauna are promoted by this progressive terrestralization across the broad S–D transition (Bambach, 1999). 4. Discussion and implications The Silurian Period has previously been considered as an interval of relative ecosystem stability within a greenhouse period (Wilde et al., 1991). However, recent studies of carbon and oxygen isotope stratigraphy instead imply major biotic, climatic, and eustatic changes during this timespan (see summaries in Kaljo et al.,1998, 2003; Munnecke et al., 2003; Calner et al., 2004; Stricanne et al., 2006; Loydell, 2007). Kaljo et al. (2003) have tested the commonly accepted, two-mode (cold/warm) oceanic scenario first developed by Jeppsson (1990, 1998), and reveal many disagreements between the model predictions and observed carbon·isotopic trends. In fact, although several oceanographic models have been proposed, their indicated causal connections among glaciation, oceanographic mode, primary productivity, and the global carbon cycle remain a matter of debate (see Azmy et al., 1998; Calner and Eriksson, 2006; Johnson, 2006; Stricanne et al., 2006; Loydell, 2007; Vecoli et al., 2009). 4.1. Current understanding of the S–D carbon isotopic excursion The Late Silurian was characterized by strong environmental changes, as indicated by several pronounced positive δ13C and δ18O excursions, exemplified by the extreme late Ludlow (Ludfordian) positive carbon isotopic excursion of 8‰ (Azmy et al., 1998) to 11‰ (Wigforss-Lange, 1999), the strongest in the Phanerozoic (Lau Event; Stricanne et al., 2006, Jeppsson et al., 2007). Thus, the well-known abrupt and highamplitude·δ13C peak at the S–D transition (+5.8‰ in Podolian sections) may be seen as a final step of the destabilization of the Late Silurian global ecosystem (Brunton et al., 1998, Fig. 3; Kaljo et al., 2003, p. 63). The hitherto proposed cause of this major biogeochemical perturbation remains conjectural, but the factors certainly triggered increases in carbon pumping and organic matter burial, recorded in organic-rich deposition in some peri-Gondwanan basins (Buggish and Mann, 2004, p. 538). In particular, Saltzman (2002) suggests that the volume of extra carbonate eroded from exposed carbonate shelves during the assumed global sea-level fall may rapidly increased the δ13C index of the riverine weathering input (see also sulfur isotopic data in Gill et al., 2007). This significant regression is seen therefore as a principal cause of enhanced nutrient delivery to the oceans and an effective trigger of an increase in organic carbon burial rates in Middle Palaeozoic settings (see discussion in Buggisch and Joachimski, 2006 and Lubeseder, 2008). The changes are remarkably abrupt and shortlived but non-catastrophic only scenarios are considered for this ‘quiet’ time interval. According to some authors (Bickert et al., 1997; Munnecke et al., 2003), such a large amplitude (above 5‰) cannot be explained easily by fractionation between surface and deep-water isotopic composition due to enhanced marine primary production. On the other hand, the observed 13C acme signal of Klonk Event might be regionally exaggerated by various forcing agents in the epeiric realm (see Immenhauser et al., 2007), especially an elevated level of biological pumping (e.g., Saltzman et al., 2004). In pelagic regimes, the δ13C excursion is extraordinarily negligible (1.5‰; Porębska et al., 1999). This distinctive bathymetric isotopic gradient is repeated in other Silurian positive δ13C excursions (Munnecke et al., 2003; Vecoli et al., 2009; see discussion in Loydell, 2007), confirming – or supposed as such by Herten and Mann (2003) – a main burial loci of organic carbon in shallow-water settings. 4.2. Proposed extended scenario The global circulation of carbon in the Earth system (see review in Kump et al., 2003) is illustrated by well-known models (e.g., Garrels

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and Perry, 1974; Berner and Lasaga, 1989; Kump and Arthur, 1999; Berner and Kothavala, 2001; Berner, 2006). The proposed S–D model evaluates isotopically differentiated qualitative flows between reservoirs, which resulted in the measured isotopic record, while the observed geological data within the study area and other palaeoregions (e.g., Wilde and Berry, 1984; Aldridge et al., 1993), as well as sulfur and strontium isotopic time curves and sea-level changes, are treated as independent test of the developed scenario. It should be stressed once more that the steady sulfur isotopic ratios indicate an overall stability of the S–D deep oceanic geochemical system as a whole (Strauss, 1999; Kampschulte and Strauss, 2004), and therefore the key changes, which contributed to the recorded isotopic anomalies, are related to quickly evolving shallow epicontinental seas–continent–biosphere feedbacks. As statistically shown by Prokoph and Veizer (1999), cyclicities and excursions in C, O and Sr isotope systems all likely reflect plate reorganizations within the Caledonian, Hercynian and Alpine tectonic cycles, even if underlying feedbacks are crudely understood for carbon cycling. The trigger factors relate to the hydrological cycle and continental weathering/riverine input (δ13C = 0‰), as well as volcanic and metamorphic flux of CO2 (δ13C = −5‰; see modeling in Kump and Arthur, 1999). Undoubtedly, geodynamic processes strongly influenced overall rates of organic matter production and storage. In the proposed scenario, two global ecosystem states are crucial for understanding S–D biogeochemical evolution: an epicontinental and an opposite pericontinental states (following Heckel, 1972) for primary production conditions in the palaeoworld costal areas. These reflect fundamental switches in the total balance of the main elements in the Earth exogenic system: oceanic carbonate system, atmospheric CO2·and biospheric CH2O, with buried Corg. The principal modes are linked to (A) favorable conditions for primary production and for the deposition of a considerable amount of Corg in sediments, followed by (B) a gradual reduction of the Corg deposited in sediments during regression by return (reburial) and oxidation of carbon to CO2·during exposure of the sediments to oxic and/or meteorologically-based weathering conditions. Taken as a whole, conditions of geochemical cycling during these two states of the Middle Paleozoic geosystem are greatly opposed (Figs. 3 and 4), and various aspects were documented previously by many workers (e.g., Jeppsson, 1990, 1998; Wenzel and Joachimski, 1996; Bickert et al., 1997; Brunton et al., 1998; Munnecke et al., 2003; Calner et al., 2004; Saltzman, 2005; Johnson, 2006; Cramer and Saltzman, 2007; Immenhauser et al., 2007; Lubeseder, 2008): (1) The epicontinental state A: characterised by generally warm climate, decreasing atmospheric CO2 engaged in primary production and carbon burial, growing amount of atmospheric O2, very active hydrological cycle stimulating strong chemical erosion of peneplanized landmasses, and changeable palaeoceanographic conditions in epicontinental seas (including their salinity), thermal stratification, thermohaline circulation, weak influence of the upwelling zones, dependence of primary production on direct input of land-derived nutrients, extensive carbonate sedimentation, and wide littoral zones. (2) The pericontinental state B: characterized by cooler conditions, extensive weathering driven primarily by progressive uplift of continents marked by narrow littoral zones, increased CO2 consumption, possible production of isotopically light CO2 in volcanic zones activated during orogenic phases and more 12Cenriched CO2 produced by oxidation of Corg, and mixed upper oceanic waters with possibly weaker stagnation in the bottom oceanic zone. Primary oceanic productivity depended on upwelling and land-derived runoff. Even if so many global aspects are implied as responsible for an important dichotomy of conditions in the earth system, the distinction between presence or absence of wide littoral zones should be

strongly·emphasized, paired in general terms with the relative area covered by epeiric seas. The presence of a passive plate–tectonic setting during the Silurian is evidenced by relative oceanic ridge quiescence and extensive epeiric seas, with active reef growth in the near-equatorial belt (see summaries in Wilde et al., 1991; Golonka, 2000; Scotese, 2001; Copper, 2002; Kiessling et al., 2003). Low-relief landmasses displayed negligible erosion rates and suppressed nutrient transfer to marine niches (thus predominant presence of oligotrophic regimes, Martin, 1996; Cramer and Saltzman, 2007; see also Saltzman, 2005). Tectonically driven processes promoted climate change at the end of the Caledonian orogenic activity because they strongly affected global weathering and, therefore, enhanced land-derived nutrient input and lowered CO2·levels (Berner and Kothavala, 2001). These tectonic processes also resulted in totally different sedimentary and trophic conditions in the presence of shrinking epicontinental seas. These factors also played a role in the retreat of carbonate platforms, recorded worldwide in the tropical domain during the Pridoli–Pragian interval (Brunton et al.,1998; Copper, 2002, Fig. 1), when the shelf area covered by carbonate platforms decreased from ~25 to 15% (Kiessling et al., 2003, Fig. 17). Late Silurian–Early Devonian marginal land areas were occupied by escalating primary producers, while no profound changes in the deep-sea carbon sequestration at that time may be inferred from the contemporaneous seawater sulfur isotopic curve (compare Gill et al., 2007). So, following Herten and Mann (2003), a large-scale increase of primary production and especially burial at the sea–land transition and nearshore domains, such as estuaries and river deltas, is hypothesized (see Figs. 3 and 4). Rapidly evolving terrestrial biosphere, paired with orogenically-enhanced continental weathering and organic matter burial (see Averbuch et al., 2005), jointly participated in the developing reversed-greenhouse effect and short-term climatic deterioration (see further discussion of oceanographic consequences in Archer et al., 2004). In fact, Jeppsson (1998, p. 253) considered the Klonk Event as paired with the ending of a warmer and arid (‘secundo’) oceanic phase (see also Kaljo et al., 2003, Fig. 3; Melchin et al., 2004, Fig. 13.3). Noteworthy, the δ18O time curve is also marked by a large δ18O values increase across the S–D boundary in Podolia: by at least 3‰ in micrites (Małkowski et al., 2009, Fig. 5 therein), and by 3.7‰ in more reliable brachiopod calcite samples of Veizer et al. (1999), what is visible also in whole-rock isotopic data from Gondwanan basins. A positive δ18O shift, from −6 to −3.5‰, was described from the Barrandian (Hladíková et al., 1997, 1999), whilst similar chemostratigraphic pattern from the Carnic Alps is interpreted by Buggisch and Joachimski (in Schönlaub et al., 1994) exclusively as a diagenetically controlled signature (see also North American data in Saltzman, 2002). Although there may be indeed a diagenetic influence (see comprehensive discussion of the primary vs. secondary dilemma in Veizer et al., 1999, pp. 74–82, and Li et al., 2006, among others), the main temperature signal is inferred from this intercontinental dataset. Significantly, the climatic cooling pulse, predicted above in our scenario for the Klonk Event, is confirmed by a short-term interregional shallowing of the upwelling system in the Rheic Ocean (Porębska and Sawłowicz, 1997; see also Lubeseder, 2008), as well as recently modeled by Simon et al. (2007). The short-lived lowering of pCO2 to concentrations around 2000 ppmv is shown as a result of effective organic carbon burial because of increasing phosphorus input to the marine reservoirs during the global regression acme. Despite overall sea-level lowstand at the broad S–D transition interval, favorable conditions for primary production and for the deposition of Corg·took place in the dominantly epicontinental-type settings (Fig. 3), possibly controlled by regionally accelerated subsidence in an active late Caledonian geodynamic setting. The orogenically sustained regression in the early Lochkovian, supplemented by high sediment input, continental runoff and possibly a cooling pulse, led to reburial conditions and may destroy considerable

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areas of primary production in pericontinental seas (as recorded in the isotopic curve fall). In concert with the reversed-greenhouse trend, the gradual oxidation of isotopically light carbon caused the subsequent growth of atmospheric CO2 content later in the Early Devonian, followed by the transiently interrupted warming trend of global climate (see e.g., House and Gradstein, 2004, Fig. 14.3), toward the time of Old Red Sandstone deposition in Podolia (see Uchman et al., 2004), but of first reefal build-ups in North Africa (Lubeseder, 2008). 4.3. Time-scale approximation of the S–D carbon cycle perturbation Causes that involve changes in the balance of buried and eroded carbon and/or transformed weathering rates include the carbonate/ silicate rock weathering cycle, with a long time scale on the order of hundreds of thousands to millions of years (Berner, 2003, 2006; Berner and Kothavala, 2001). The stratified ocean model is applied to the global Silurian–Devonian carbon cycle analysis in relation to the described carbon isotopic changes (Fig. 2). The studied successions were undoubtedly deposited at the epeiric sea bottom, i.e., in the upper part or box of the hypothetical ocean. It is assumed that the sampled part of the S–D oceanic carbonate system was not less than 10–20% of the whole and that the system was not much different globally from the modern analog. If so, the upper box of the inorganic carbon reservoir is estimated to have contained between 4000 and 8000 Gt (e.g., Spitzy and Degens, 1985). To this approximation, the biota, dissolved organic carbon and atmospheric CO2 (which are in isotopic equilibrium with the oceanic carbonate system; e.g., Hoefs, 1997; Kump et al., 2003; Archer et al., 2004) may be added (Berner and Kothavala, 2001). The sum of all these values is not less than 5000 Gt for the total S–D carbon reservoir. When the Spitzy and Degens (1985) calculations are taken into consideration for the main δ13C shift from −2 to +4‰, a storage of ~1500 Gt of organic matter was necessary, which had to be stored outside of the main carbonate reservoirs. In the modern world, the productivity of estuaries and river deltas may be ~0.02 g/cm/yr (e.g., Schopf, 1980), and the area of the latest Silurian plant colonization may be globally in order of a 1 million km2, a figure that is reasonable in light of current phytogeographic continental reconstructions (Edwards and Wellman, 2001; Raymond et al., 2006). Hence, if 2% of the total organic production was sequestered annually in sediments, 500,000 yrs would have been sufficient time for the burial of 1500 Gt Corg necessary to cause the 6‰ positive shift of the carbon isotopes curve (from − 2 to + 4‰). This isotopic change prior the S–D boundary was quite abrupt, being equivalent to the sequestration of ~1000 Gt of isotopically light carbon. In the Podolian section, a sharp 4‰ shift occurs through only 6 m of marly limestone deposits and may represent an isotopic change of up to 1‰ per 10,000 yrs. The whole positive trend observed in the same sequence covers ~40 m, a minor part of the >500 m-thick Lochkovian succession that represents ~5 million yrs of deposition (Małkowski et al., in press). Thus, this rapid burial episode may have lasted as little as ~20,000–30,000 yrs, driven by enlarged production/ storage areas or by more efficient production/burial conditions. In contrast, the negative carbon isotopic shift between the S–D boundary and onset of the Old Red facies was much slower (Fig. 2), i.e. with a rate of 1‰ per million yrs. Hence, the observed rate difference is a basis of the concept that the organic matter accumulation depended on fast ecological processes and the opposite reburial tendency depended on geological processes that usually operate more slowly. In fact, a dynamic interplay between different unsteady processes in the photic zone and in the deeper-water carbonate shelf system was probably responsible for the episode of biogeochemical perturbation across the S–D transition. An exchange of isotopically different water masses between both hypothetical boxes of the earliest Devonian ocean remains an alternative scenario for the sluggish organic matter burial-to-reburial process assumed for the

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unsettled Silurian ocean, as proposed by Bickert et al. (1997). Nonetheless, large-scale intra-oceanic mixing and/or upwelling of 13 C-depleted waters were far less probable in the succeeding climatically stable ecosystem, if greenhouse and oceanic stagnation were indeed the dominant Devonian tendencies (Saltzman, 2005). 4.4. Comparison with other middle Paleozoic carbon cycle perturbations Several distinct Silurian episodes of high C- and O-isotope ratio values were recently recognized, representing times of relatively oligotrophic regimes, eustatic lowstands, increased growth of reefs and the formation of extended carbonate platforms over well-aerated shelves (e.g., Jeppsson, 1998; Kaljo et al., 1998, 2003; Saltzman, 2001; Calner and Eriksson, 2006; Johnson, 2006; Stricanne et al., 2006; Cramer and Saltzman, 2007; Loydell, 2007; Vecoli et al., 2009). As stressed by Munnecke et al. (2003), the similar relationships between Silurian and earlier Paleozoic positive isotope excursions, protracted biotic overturns and facies development indicate analogous controlling mechanisms, particularly circulation switches between humid and arid conditions in low latitudes. In this model, carbon isotope ratios are tied to the advection of 13C-rich surface water (arid episodes) or upwelling of 13C-depleted deep water (humid episodes) in a Silurian ocean marked by permanent euxinic deep-water conditions due to the storage of organic carbon in black shales (Bickert et al., 1997; Munnecke et al., 2003; Calner et al., 2004; Stricanne et al., 2006). It is advocated by Cramer and Saltzman (2007), however, that sequestration and burial of organic carbon in deep anoxic water, as a result of altered deep ocean circulation from thermohaline to low-latitude halothermal mode, played a key causal role. However, a convoluted relationship between the magnitude of isotopic and biotic events is evident (Fig. 3; see also Kaljo et al., 1998). When compared with the older Silurian isotopic events, the Klonk Event obviously bears several specific characters determined primarily by the collapse of reefs and probably of the whole carbonate factory, possibly paired with eutrophication, high primary productivity and oxygen depletion in shelf settings (Hladíková et al., 1997; Saltzman, 2002; Herten and Mann, 2003; Buggish and Mann, 2004; Williams and Saltzman, 2004). In particular, the highlighted strongly enhanced bioproduction, evidenced also in the S–D pelagic habitats (Porębska and Sawłowicz, 1997), contrasts with recent explanations of the older isotopic events (Stricanne et al., 2006; Cramer and Saltzman, 2007). Alternatively, despite an extreme difference in biotic responses and environmental perturbation rates, some analogies with the Late Devonian Kellwasser events can be made (see summary in Joachimski et al., 2004; Chen et al., 2005; Buggisch and Joachimski, 2006). As noted already by Andrew et al. (1994), “the different isotopic responses to periods of marked reduction of biomass and biodiversity in the Silurian and the Devonian implies fundamentally different causes and/or responses to the environmental factors leading to the events”. Looking into the diversity of possible causal processes, the contributory role in this difference is ascribed to the increasing biogeochemical impact by terrestrial biota, but essentially modified by evolving climatic, palaeogeographical and tectono-magmatic driving forces (see e.g., atmospheric evolution data in Fig. 5). On the other hand, the terrestrialization factor was neglected even for the Lau Event, merely 3.5 Ma prior the Klonk Event (see summary in Kaljo et al., 1998, p. 309), when the environmental changes connected with the δ13C excursion first affected deeper-water habitats and later the photic zone (Stricanne et al., 2006). In fact, the enormously increasing δ13C values of +6 to +11‰ are suggested to be attributed to photosynthetic activity (Wigforss-Lange, 1999), and despite more heavier strontium isotope ratios during this timespan (Ruppel et al., 1996), the ‘terrestrial signal’, indicated by high abundance of spores during the isotope excursion, is explained by increased aeolian input

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(Stricanne et al., 2006; see also Calner and Eriksson, 2006). Only Vecoli et al. (2009) remarked, in the context of this terrestrial signal during the Lau Event, a continued increase in geographical dispersal, morphological complexity and biodiversity of early vascular plants (see also Porębska et al., 1999). 5. Conclusions 1. The recently refined isotopic systematics provide strong support for the large-scale global biogeochemical perturbation across the Silurian–Devonian boundary (Klonk Event), manifested in almost 6‰ positive shift of the carbon isotopes curve, and offer an improved interpretation of the large changes in the exogenic system (Figs. 3 and 4). The interpretation of the δ13C excursion as recording original seawater chemistry due to an extensive ecosystem disturbance is consistent in the context of a noteworthy global bio-event at that time. 2. The steady sulfur isotopic ratios indicate an overall stability of the S–D oceanic geochemical system as a whole (Kampschulte and Strauss, 2004), thus indicating that the crucial changes that contributed to the recorded carbon cycling perturbation are driven by rapidly changing shallow epicontinental seas–continent– biosphere feedbacks in the non-actualistic stratified Palaeozoic oceanic realm. 3. The primary turnover in carbon cycling is usually seen to commence later in the Devonian with the development of new widespread sinks owing to the spread of large woody plants across mid-Devonian terrestrial settings. A newly proposed key to the storage of considerable quantities of organic carbon during the S–D transition is the rapid expansion of early vascular vegetation into extensive near-coastal shallows and deltas of the suturing lands formed in the course of intermittent late Caledonian orogenic movements (albeit the storage record was mostly destroyed by the synchronous uplift activity). The proposed change in the delivery and cycling of carbon from the deep ocean to inner shelves is a distinctive feature of the isotopic balance perturbation scoped at land–sea transition, when compared with other current models of Silurian isotopic excursions (Munnecke et al., 2003; Calner et al., 2004; Stricanne et al., 2006; Cramer and Saltzman, 2007). 4. This postulated great increase in the primary productivity of primitive terrestrial plants and more efficient organic carbon pumping was responsible for the rapid enhancement of carbon burial (lasting perhaps only a maximum of ~20,000–30,000 yrs), and resulted also in lowering of atmospheric CO2 levels and a predicted cooling during the Klonk Event (Simon et al., 2007). Thus, the isotopic preturbation in the global carbon cycle is seen as the result of a unique combination of palaeogeographic, biogeochemical and evolutionary processes. This process collapsed because the syn-orogenic sea-level fall and cooling (Fig. 3) strongly limited sequestration of organic carbon, thus destroying the efficiency of the carbonate factory. 5. Subsequently, the C-rich deposits were gradually eroded during an Early Devonian sea regression, paired with the geodynamically active end-Caledonian tectonic setting. Despite progress in land plant colonization, the biogeochemical balance shifted toward long-term organic matter reburial and oxidation, and return of the carbon isotope ratios to background values due to a large-scale C12 delivery. 6. The long-term oxidation of sedimentary organic matter resulted finally in the following rise of atmospheric CO2 content throughout the Early Devonian, and caused the warming of global climate. This was an initiation of prolonged greenhouse conditions that culminated catastrophically in the disturbed climatic mode of the Frasnian–Famennian transition (Joachimski et al., 2004; Averbuch et al., 2005).

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