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Authors requiring further information regarding Elsevier’s archiving and manuscript policies are encouraged to visit: http://www.elsevier.com/copyright Author's personal copy Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 Contents lists available at ScienceDirect 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 Ordovician and Silurian sea–water chemistry, sea level, and climate: A synopsis Axel Munnecke a,⁎, Mikael Calner b, David A.T. Harper c, Thomas Servais a,d a GeoZentrum Nordbayern, Fachgruppe Paläoumwelt, Loewenichstraße 28, D-91054 Erlangen, Germany Department of Earth and Ecosystem Sciences, Lund University, Sölvegatan 12, SE-223 62 Lund, Sweden c Statens Naturhistoriske Museum (Geologisk Museum), Øster Voldgade 5-7, DK-1350 København K, Denmark d FRE 3298 du CNRS Géosystèmes, Université de Lille1, SN5, F-59655 Villeneuve-d'Ascq Cedex, France b a r t i c l e i n f o Article history: Received 15 April 2010 Received in revised form 3 August 2010 Accepted 6 August 2010 Available online 14 August 2010 Keywords: Ordovician Silurian Sea level Stable isotopes Climate a b s t r a c t Following the Cambrian Explosion and the appearance in the fossil record of most animal phyla associated with a range of new body plans, the Ordovician and Silurian periods witnessed three subsequent major biotic events: the Great Ordovician Biodiversification Event, the end-Ordovician extinction (the first animal extinction and second largest of the five mass extinctions of the Phanerozoic), and the Early Silurian postextinction recovery. There are currently no simple explanations for these three major events. Combined extrinsic (geological) and intrinsic (biological) factors probably drove the biodiversifications and radiations, and the appearance and disappearance of marine habitats have to be analysed in the frame of changing palaeogeography, palaeoclimate and sea-water chemistry. The present paper reviews the relationships of the three biotic events to chemical and physical processes occurring in the ocean and atmosphere during the Ordovician and Silurian, including sea-level changes, geochemical proxies (δ13C, δ18O, 87Sr/86Sr) of the ocean waters, and the evolution of the atmosphere (oxygen and carbon dioxide content). © 2010 Elsevier B.V. All rights reserved. 1. Introduction 1.1. The palaeobiological context During the Ordovician and Silurian, profound changes occurred in the planet's ecosystems. Marine life was characterised by a major diversification, the Great Ordovician Biodiversification Event (GOBE), a major extinction, the end-Ordovician event and a subsequent recovery during the Early Silurian. These events are part of a continuum from the evolution of the first metazoans at least by the Ediacaran, the skeletalization of animals during the late Neoproterozoic, the explosion of body plans during the Early to Mid Cambrian, and the massive diversification of benthic marine life during the Ordovician, consequently with demersal and nektonic organisms radiating during the Devonian (Klug et al., in press). Pivotal to this process was the GOBE, but there is currently no single explanation (Servais et al., 2009, 2010; Zhang et al., 2010b). Perhaps a coincidence of biological and geological factors combined to help drive and encourage the biodiversification. Irrespective of its causes, the diversification changed the oceans forever and set a new agenda for marine life (Harper, 2006). The cascading increase in biodiversity at species, genus and family hierarchies was apparent at global levels with the high provincialism of Early to Mid Ordovician faunas, at regional levels with the development of new community types, particularly in deeper water and in and around reefs, and thirdly ⁎ Corresponding author. E-mail address: [email protected] (A. Munnecke). 0031-0182/$ – see front matter © 2010 Elsevier B.V. All rights reserved. doi:10.1016/j.palaeo.2010.08.001 at local levels where more animals were squeezed into existing communities. The oceans were no longer sterile expanses of water, being filled now by phyto- and zooplankton, punctuated by blooms, and including larvae and animals such as the graptolites. Community structures were better organised and more densely packed with the expansion of the number of so called ecological guilds, signalling a range of new feeding strategies and life modes. Tiering structures developed both above and within the substrates while the bioerosion and encrustation of hard surfaces offered a new range of ecological opportunities. The Palaeozoic evolutionary fauna was relatively stable, surviving the end-Ordovician and late Devonian extinctions for some 200 million years. The end-Permian extinction event virtually destroyed its suspension feeding networks and a new ecosystem, based on the detritus feeding ecosystem of the Modern evolutionary fauna and a more explicit arms race, the escalated interactions between predators and prey, diversified and intensified during the Triassic. Additionally during the Ordovician and Silurian, widespread biogenic carbonate factories were established through the generation of heavily skeletalised organisms together with metazoan reefs with consequences for the longer term function of the carbon cycle and the planet's climate. The biological signals for these events have become well established during the last two decades, however, their relationships to chemical and physical processes occurring in the world's ocean and atmosphere are far from clear. Nevertheless through biological, physical and geochemical proxies, the roles of sea–water chemistry and sea level on the planet's climate and evolution are now being more accurately investigated, not least through a range of new techniques and carefully collected field data. In particular ocean Author's personal copy 390 A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 geochemistry and sea-level change is making a huge impact on our understanding of the palaeoclimate and palaeogeography of the Ordovician and Silurian together with the evolution of their biotas. These advances are in part a tribute to the networks and results of two major international projects, the most recent, IGCP 503 with a specific focus on the climate and geography of the two periods. 1.2. IGCP 503 The International Geoscience Programme (IGCP) 503 ‘Ordovician Palaeogeography and Palaeoclimate’ commenced in 2004 and was completed in 2009. It was based on the previous, highly successful IGCP project focused on the Ordovician, IGCP 410 ‘The Great Ordovician Biodiversification Event’ that extended from 1997 to 2002 (see Webby et al., 2004). The main objectives of IGCP 410 were to construct diversity curves for all marine invertebrates during the Ordovician biodiversification, but also to establish a new stratigraphical standard that would permit intercontinental correlation, not only of strata, but also of palaeodiversity patterns and trends of all fossil groups during the radiation. Following project 410, the new programme IGCP 503 focused specifically on the search for the biological and geological triggers of the Ordovician biodiversification. The main goals were, as indicated in the title of the project, to understand the influence of changing geography and climate on the Ordovician radiation. However, because such changes in the Ordovician could only be understood within a broader frame, many Cambrian and Silurian workers also participated in the programme. Project 503 was therefore not limited to the Ordovician Period, and most meetings and field trips covered the entire Lower Palaeozoic. The scientific output of the project, comprising several hundred published papers is, of course, difficult to summarise. Servais et al. (2009, 2010) have reviewed many results of the project, indicating that a continuous sea-level rise between the Early Cambrian and the early Late Ordovician broadly matches the diversification of the marine invertebrates during these periods. Palaeogeography can also be linked to the Early Palaeozoic radiation, as it coincides with the breakup of the supercontinent Rodinia in the late Precambrian. The formation of numerous smaller continents triggered the biodiversification, with seafloor spreading and continental dispersal at their maxima during the Ordovician, together with the greatest extension of tropical shelves of the entire Phanerozoic. Several special issues have published results of project 503, with many of them including some of the major advances presented at the main annual meetings (Munnecke and Servais, 2007; Owen, 2008; Servais and Owen, 2010). The present special issue highlights climate and sea-level changes during the Early Palaeozoic and the range of investigative techniques currently available for their study. The present paper reviews our knowledge of these parameters and their relationships to the Great Ordovician Biodiversification Event, the endOrdovician extinction, and the subsequent Silurian radiation. 1.3. The Ordovician and Silurian world The Ordovician and Silurian world witnessed three major biotic events, the Great Ordovician Biodiversification Event (Webby et al., 2004; Harper, 2006), the end-Ordovician extinction (Barnes, 1986; Rong and Chen, 1986; Rong and Harper, 1988; Barnes et al., 1995; Sheehan, 2001a) and the Early Silurian recovery (Rong and Harper, 1999). These three events helped develop the complexity of the Palaeozoic evolutionary fauna and established the pattern of marine life in the Ordovician and Silurian world (Sheehan, 2001b). In general terms Palaeozoic oceans were characterised by short trophic chains dominated by suspension-feeding organisms, evolved during persistent intervals of greenhouse climate. This ecosystem contrasted with that of the subsequent Mesozoic and Cenozoic eras, dominated by deposit-feeding communities linked to more bioturbated substrates, complex community structures driven by a more pervasive arms race. Early Palaeozoic biodiversifications in most marine groups were spectacular and sustained, setting the agenda for subsequent marine life on the planet. The majority of metazoan groups appeared first at the base of the Palaeozoic (Budd, 2008), increasing in diversity during the Cambrian Explosion and Ordovician Radiation to establish an ecosystem that survived some 250 million years of Earth history. Nevertheless Cambrian ecosystems were probably quite different from those to follow during the Ordovician and Silurian periods, characterised by relatively few megaguilds, poorly-structured communities and a relatively sterile water column. The transition from the Cambrian to Ordovician worlds was a major turning point in Earth history. Much evidence now suggests that the late Cambrian was characterised by warm oceans with widespread anoxia and dysoxia and probably low saturation states for calcite and aragonite (Pruss et al., 2010). Despite the appearance of calcified skeletons in both solitary and colonial organisms in the late Neoproterozoic (Wood et al., 2002), the Cambrian carbonate factory was dominated by physical and microbial processes rather than by biogenic material. Carbonate build-ups and reefs were rare following the virtual extinction of the archaeocyathans in the mid Cambrian. The Ordovician Period experienced a truly massive rise in marine biodiversity (Sepkoski, 1981) accompanied by an increase in the biocomplexity of marine life (Droser and Sheehan, 1997) marking ‘The Great Ordovician Biodiversification’ as one of the two most significant radiation events in the history of marine life. The unique environmental conditions through the Ordovician Period have been emphasised in a number of publications (e.g., Jaanusson, 1984). Extensive, epicontinental seas developed during sea-level high stands (Algeo and Seslavinski, 1995; Pratt and Holmden, 2008), driven by an extended greenhouse climate, were associated with virtually flat seafloors and restricted land areas, many probably represented only by occasional, emergent archipelagos. Sea levels were most probably the highest of the Palaeozoic and possibly the highest of the entire Phanerozoic (Hallam, 1992; Miller et al., 2005; Haq and Schutter, 2008), and there are no modern analogues to the epicontinental seas of the Ordovician Period. Magmatic and tectonic activity was intense and persistent with rapid plate movements and widespread volcanic activity. Possibly even mantle plumes were associated with climatic and faunal changes (Barnes, 2004; Lefebvre et al., 2010-this volume). Island arcs and mountain belts provided sources for clastic sediment in competition with the carbonate belts associated with the platforms on most of the continents. The continents were widely dispersed (Cocks, 2001) driving provincialism. Such biogeographical differentiation was extreme, affecting plankton, nekton and benthos, and climatic zonation, particularly in the southern hemisphere, was pronounced. Provincial differentiation amongst the benthos was also marked with biogeographic differences persisting until near the end of the period (Williams, 1973), when these were disrupted by the endOrdovician glaciation (Rong and Harper, 1988; Owen et al., 1991). Together these conditions were, nevertheless, clearly ideal for allopatric (geographic) speciation processes together with opportunities for canalization of ecological niches (Harper, 2006). Climate and environmental proxies for this interval, that are advancing our understanding of the background to the diversification, extinction and recovery of biotas, are in a rapid stage of development. Isotope shifts in carbon, oxygen and strontium are providing vital clues to the cycling of carbon, temperature fluctuations and the input of terrigenous material associated with orogenic activity. Moreover, sea-water chemistry is proving essential to our understanding of skeletal secretion. The link between sea-level change (e.g., McKerrow, 1979) and climate change together with tectonic activity has been known for some time but refined regional sea-level curves (see Section 2; Figs. 1, 2) are providing a more accurate and precise assessment of these global processes and their influence on biotic evolution at taxonomic, community and ecosystem levels. Today, it is more and more clear that the Great Ordovician Biodiversification Event was an accumulation of biodiversification Author's personal copy A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 events (taking place at different times on different continents within different phyla) that covered the entire Ordovician and were part of a wider Cambrian–Ordovician radiation (Servais et al., 2010). This longterm radiation was interrupted by the first of the ‘Big Five’ mass extinctions: the end-Ordovician extinction, considered as being the second most important extinction of Phanerozoic marine life (the Permian–Triassic extinction being the most severe, e.g., Sepkoski, 1981). This end-Ordovician extinction was considered for many years to be abrupt and directly correlated with the Hirnantian glaciation, with the disappearance of about 85 % of marine species (e.g., Sheehan, 2001a). Two pulses of extinction have been recorded, and discussed in the frame of the glacial intervals in the Late Ordovician. The first pulse was related to the beginning of the glaciation with an important sealevel fall. The second pulse of the extinction was related to the end of the glaciation when sea level started to rise again and oceanic circulation stagnated, marking the end of a long interval of ecologic stasis (Ecologic-Evolutionary Unit) (Sheehan, 2001b; Brenchley et al., 2003). However, in the last few years, several authors noted that the global cooling at the end of the Ordovician was not as abrupt as previously thought (e.g., Saltzman and Young, 2005). Temperatures (and sea level) were decreasing since the middle part of the Late Ordovician, accompanied by a decrease of biodiversity in many fossil groups that apparently began much earlier than the Hirnantian glaciation (e.g., Servais et al., 2008). The Hirnantian glaciation probably only marked the final Ordovician phase of a long interval of overall global cooling. By contrast the Silurian was considered a relatively short but calm period most noted for the beginning of the greening of the land and the radiation of the gnathostome and theledont fishes (see numerous articles in Holland and Bassett, 1989; Bassett et al., 1991; Landing and Johnson, 1998, 2003). On a broader scale, the Silurian is sandwiched in between the Late Ordovician ice-house climate and Devonian extreme greenhouse conditions. Similar to the Ordovician, it is characterised by an archipelagic distribution of several continents in low latitudes (Laurentia, Baltica, Siberia, Kazakhstania), a vast north polar ocean, and the supercontinent Gondwana extending from equatorial latitudes to the South Pole. The sea level was high, large shallow epicontinental seas were widely distributed, and the continents had a low relief. Terrestrial plants were quantitatively insignificant and thus had little influence on the global carbon cycle. During the Silurian, the Iapetus Ocean, that separated Laurentia and Baltica, was closed leading to the Caledonian orogeny. Major extinction events comparable to those of the Ordovician or Devonian periods were unknown (Kaljo et al., 1995), and, except for the Malvinokaffric realm (the southern temperate zone typically represented by the low-diversity Clarkeia (brachiopod) fauna from Gondwanan Africa and South America), reefs were widely distributed. Reefs are reported more or less throughout the entire Silurian, but their distribution through time seems to be clustered. The earliest Llandovery is characterised by the near absence of reefs. The first Silurian reefs appeared in the midAeronian (Li and Kershaw, 2003). Intervals with higher abundances of reefs are the mid to late Aeronian, latest Telychian to early Sheinwoodian, late Homerian, late Gorstian to early Ludfordian, and mid-late Ludfordian (Brunton et al., 1998; Copper, 2002). However, since Silurian biostratigraphy is mainly based on graptolites, which normally are very rare in reefal limestones, the precise stratigraphic correlation of many reef deposits with respect to the graptolite biostratigraphy is still somewhat uncertain. In the past two decades our picture of the apparently ‘calm’ Silurian has changed dramatically (see review in Calner, 2008). Investigations of stable carbon and oxygen isotopes suggest a highly volatile ocean–atmosphere system (e.g., Samtleben et al., 1996; Saltzman, 2001; Kaljo et al., 2003). The presence of four major positive stable carbon isotope excursions in the Silurian (early Wenlock, late Wenlock, late Ludlow, Silurian–Devonian boundary) suggest that fundamental changes in the global carbon cycle were 391 much more frequent in the comparatively short Silurian Period than in any other system of the Phanerozoic (see Section 3). The amplitudes of the Silurian stable isotope excursions are extremely high compared to Mesozoic and Cenozoic excursions, and there is no general agreement on the palaeoenvironmental changes responsible for these excursions (see Section 3). The carbon isotope excursions are also characterised by elevated oxygen isotope values, and are closely correlated with extinction events and with lithological changes (summarised in Munnecke et al., 2003). At the very beginning or even prior to the increase of C- and O-isotope values, many groups of organisms were affected. Especially conodonts, graptolites and trilobites, but also acritarchs, chitinozoans, ostracods, brachiopods, and corals show extinctions, sometimes of a step-wise nature; organisms living in hemipelagic environments were more strongly affected than organisms occupying shallow-water settings. Munnecke et al. (2003) postulated similar but unknown controlling mechanisms for the major Silurian isotope excursions based on their lithological, palaeontological, and geochemical similarities. 2. Sea-level development 2.1. Introduction Sea level exerts a first-order control on the three-dimensional facies architecture of marine sediments and is intimately related to sea-floor and substrate evolution, benthic ecology, and biodiversity in shallow-water cratonic seas. The interaction of sea level and biodiversity is important on long-term as well as short-term time scales. The close correlation between the earliest Palaeozoic firstorder transgression and increased marine biodiversity (Webby et al., 2004; Servais et al., 2009) is, for example, an excellent example of long-term changes. Similarly, the majority of the most profound extinctions in Earth history are in one way or another related to sealevel changes (Hallam and Wignall, 1999). These changes are often abrupt and possibly related to loss of ecospace (regressions) or anoxia (transgression). The temporal change in biofacies upwards through an individual parasequence is an example of this interaction during shorter time-scales. For the reasons above, accurate knowledge of sea level is central to most studies dealing with Palaeozoic marine environments. The Ordovician and Silurian periods as a whole have long been regarded to equate with an extended greenhouse interval only interrupted by a short-lived Hirnantian glaciation (Brenchley et al., 1994, 2003). Indeed, the two periods record the highest sea levels during the Palaeozoic Era, reaching peak levels of about 200 m above present-day sea level in the Sandbian and Katian (Haq and Schutter, 2008). A growing body of evidence, however, now indicates that a ~30-myr-long cool interval interrupted this long greenhouse phase, the ‘Early Palaeozoic Icehouse’ (EPI) of Page et al. (2007). This actually echoes the view previously proposed by Frakes et al. (1992). The Early Palaeozoic Icehouse includes seven glacial maxima of which four occurred in the Late Ordovician and three in the Llandovery. The implications of the EPI are immense for Early Palaeozic sea-level history since it implies that the latitudinal climate gradient must have been much steeper within this time interval than in the Early and Mid Ordovician and in the post-Llandovery Silurian, which itself would have affected the amplitudes of sea-level change. Numerous Ordovician and Silurian sea-level curves have been published in the last few decades, from various palaeogeographic domains and based on different techniques. The curves are useful for explaining regional faunal developments or even continent-wide diversification but are often difficult to apply outside their type areas. As discussed below, the proposed ‘global’ sea-level curves for the Ordovician and Silurian show many similarities at a stage level (probably reflecting 1st or 2nd order trends) but less agreement at time slice (Webby et al., 2004) or stage slice (Bergström et al., 2009b; Author's personal copy 392 A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 Cramer et al., in press) levels. At this resolution the curves often show conflicting patterns, and sometimes are even in mutual opposition (mirrored curves). The reasons for this have been discussed in the literature and can probably be explained by a varying and often inadequate degree of biostratigraphic resolution (e.g., due to facies controls and faunal provincialism), or the poor integration of various stratigraphic schemes. Apart from this, sampling strategy and the choice of investigative methods may affect the result. There are many pitfalls and numerous difficulties in interpreting sea-level change. Sea-level change occurs at several time-scales simultaneously extending from 1st order changes (global tectonics) to 5th order changes (Milankovitch-cyclicity) and, to quote Catuneanu (2006), the interpreter of stratigraphic successions must “wear the right glasses” and identify the level at which correlation is to be achieved (commonly 3 rd or 4th order). Another problem is to isolate the tectonic component of sea-level change, e.g., by backstripping techniques (e.g. Loi et al., 2010-this volume). This is especially necessary on active continental margins and within areas with icecover because of isostatic adjustment of the lithosphere. The problem of mirrored sea-level curves was elegantly illustrated by Zhang et al. (2006), who detected opposite sea-level trends in the Late Ordovician through earliest Silurian along the Laurentian continent. In a recent landmark paper on global Palaeozoic sea-level change, Haq and Schutter (2008) tackled similar problems by defining ‘reference districts’ in intracratonic basins, a concept that was first used by Johnson (1996). Although affected by mantle convection currents (dynamic topography), these types of basins are relatively stable, tectonically, and thus better preserve eustatic signals. This raises another problem, however, that of stratigraphic completeness. Many of the Lower Palaeozoic intracratonic basins constitute only a few hundred metres of thin sedimentary rocks with frequent and often substantial hiatuses. In these types of settings most strata represent transgressions and highstands of sea-level, simply because hinterland relief is too low to generate clastic material during regressions; in addition lowstand weathering of carbonate rocks does not produce any sediment. Thus, parts of the Ordovician with its monotonous temperate carbonates, is a typical example, yielding little clastic sediment (see also Kanygin et al., 2010-this volume). The relative expansion/condensation of sedimentary successions can easily result in erroneous correlations if the sampling frequency is too low. A rule of thumb is, of course, to increase the number of samples with increased condensation. But this can also affect the cross-sectional symmetry of constructed sea-level curves, that should not be compared only on the basis of shape. Another important part of any sea-level curve is the horizontal scale, showing the relative magnitude of sea-level change. Proxies for deducing this include absolute depths for benthic assemblages (Hancock et al., 1974; Brett et al., 1993), the interpreted depth of the fair- and storm-weather wave bases (Harris et al., 2004; Immenhauser, 2009; Loi et al., 2010-this volume), degree of coastal onlap as a measured distance in seismic diagrams, and the preserved relief along unconformities (Johnson et al., 1998). In summary, the possibilities of generating vast numbers of potential errors in any sea-level curve are clearly immense and present a considerable challenge. Is it even possible to create reliable global sealevel curves for the Early Palaeozoic? For example, the global sea-level curve produced by Haq and Schutter (2008) implies that the Palaeozoic Era records as many as 172 eustatic events, ranging in amplitude from a few tens of metres to about 125 m and with sea levels more than 200 m above that of the present-day. This is an interesting concept since ‘a few tens of metres’ is a significant sealevel rise for the widespread, shallow cratonic seas of the Ordovician and Silurian, and undoubtedly produced clear facies shifts. This study should also be contrasted with known causes for eustatic sea-level change, which only include changes in the volume of the ocean water or changes in the volume of the ocean basins (e.g., Miller et al., 2005). Most of these changes are slow and the only mechanism that can lower sea level substantially at subzonal time-scales are continental glaciations (see, e.g., Miller et al., 2005). 2.2. Techniques for measuring sea-level change Over the last few decades, a wide array of techniques has been developed for interpreting sea-level change. These can be categorised as based on a) sedimentary proxies b) physical proxies, c) biological proxies, or d) geochemical proxies (cf. Johnson, 2006). 2.2.1. Sedimentary proxies These include traditional facies analysis, outcrop-based or through geophysical methods in the subsurface. Process-based sedimentology has made great advances in the last few decades, and a wealth of detailed facies models for analysis of shallow marine deposits and their proximality trends have been published. Also the dynamics of various depositional systems are now well understood. Marine sedimentary facies and facies associations can therefore easily be related to energy boundaries such as the fair- or storm-weather wave base or the shoreline, and thereby readily provide a relative sea-level curve, especially when combined with information about the bio- or ichnofacies (Immenhauser, 2009). The effect of sea-level change on environments below the storm wave-base is relatively minor and successions of graptolite-yielding shales are therefore of limited use in describing sea-level change. The same facies principles are applicable to the allochthonous sediments of carbonate-dominated shelves and flat-topped platforms although the depositional profiles and terminology are different. Carbonate rocks have traditionally signalled regressions; commonly it is in fact the opposite. Carbonate platforms produce and accumulate most of their sediments during sea-level rise and highstand, whereas the same situation leads to drowning of source areas, trapping of sediments in inshore areas, and the general starvation of outer shelf areas in clastic depositional systems. The latter produce most of their sediments when sea level is lowered due to the increased depositional gradients and hinterland area exposed to weathering and erosion. A drop in sea level has severe effects on carbonate platforms, which are then exposed, and production of sediments is substantially reduced. For this reason, the bulk of the strata in low-relief intracratonic carbonate successions are arguably transgressive to highstand deposits. In studies of carbonate rocks, the classification scheme by Dunham (1962) is by far the most useful since it describes the depositional texture and not only the grain size of the rock. For example, these textures were used by Harris et al. (2004) in their study of Late Ordovician sea-level changes in Baltoscandia. Sedimentary proxies form the basis for the interpretation of stratigraphical cyclicity and sequence stratigraphic analysis, starting with the recognition of parasequences and their vertical stacking pattern. Early published sea-level curves were heavily influenced by the Global Cycle Chart produced by the Exxon group and the necessity of a eustatic explanation for observed sea-level variations led to a shoehorning of results. Sequence stratigraphy matured as a technique when outcrop-based analyses became common, when a fourth systems tract was included in the original model, with the introduction of soft terms such as ‘relative sea-level change’, and not least with the increasing understanding that carbonate and siliciclastic depositional systems respond to sea-level change in virtually the opposite way (Kendall and Schlager, 1981; Schlager, 1991; Strasser et al., 1999). It is significant that most of these conceptual changes occurred as late as the 1990s. Sequence stratigraphy is today the primary method for conducting basin analysis, virtually independent of scale and purpose, and for the establishment of relative sea-level curves. However, importantly, variations in sediment-transport rate and carbonate productivity can result in parasequence stacking patterns similar to those produced by sea-level oscillations (Burgess, 2001). There are two ways in which sea-level change can be measured; relative to basement (‘relative sea- Author's personal copy A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 level change’) or relative to the centre of the Earth (‘eustatic sea-level change’). A third, commonly-used, term is ‘depositional depth’ (local bathymetry) which can change rapidly through autocyclic processes such as delta progradation or even vertical reef platform growth. Independent of cause and scale, a full sea-level cycle as described in sequence stratigraphy evolves through four datums: onset of baselevel fall, end of base-level fall, end of regression, and end of transgression (Catuneanu, 2006). The terminology associated with these four datums and their associated stratal surfaces, however, currently lacks consensus. Apart from ‘Transgressive–Regressive sequences’ and ‘Genetic sequences’, at least three other types of depositional sequences are currently in use (Catuneanu, 2006). There is still no unique terminology or approach that fits the various depositional settings and clearly a measure of variation in the terminology is necessary to describe all eventualities. This lack of consensus, however, has not only led to confusion in the literature but also hindered a formalisation of sequence stratigraphical terminology. Such formalisation is now underway and is being currently discussed and debated by the International Subcommission on Stratigraphic Classification (ISSC) and the International Working Group on Sequence Stratigraphy (IWGSS). 2.2.2. Physical proxies Preserved relief along regional unconformities has provided an independent method for calibrating and measuring the magnitude of proposed eustatic sea-level changes in the Silurian (Johnson et al., 1998). The method provides direct evidence for the minimum magnitude of sea-level change, measured as the distance between the base of the erosional surface and the top of the buried strata. The obvious drawback is that regional unconformities with some relief are often difficult to identify and measure properly. Preserved topography in Ordovician and Silurian rocks has for example been documented by Johnson et al. (1998), Calner and Säll (1999), Desrochers (2006), and Loi et al. (2010-this volume). Recognition of geomorphology related to glaciation (valleys, striations) or karst weathering processes (epior endokarst) are other ways to detect sea-level lowstands. Both types have a continental origin, and the low preservation potential substantially limits their abundance in the rock record. 2.2.3. Biological and/or ecological proxies Biological evidence for relative sea-level change is based on the depth-ranges of fossilised macro- or microbiota (or their combination) in the rocks (e.g. Brenchley and Harper, 1998). The classic study by Ziegler (1965), based on shelly benthos in the Silurian of the Welsh Borderland, laid the foundations for the use of fossil biotas as a tool for the reconstruction of depositional depth. This study, and a subsequent refinement by Ziegler et al. (1968), developed the concept of depthspecific benthic communities, based on the pioneering work by the Danish zoologist Johannes Petersen on living marine communities (e.g., Petersen, 1918). Five communities inhabiting broad zones parallel to, and with increasing distance from the shoreline were defined; the Lingula, Eocoelia, Pentamerus, Stricklandia and Clorinda communities. The Clorinda community was transitional to graptolitefacies in distal shelf environments. The benthic community concept was further expanded by Boucot (1975) to include a wider array of shelly benthos and he therefore re-named them ‘benthic assemblage zones’ (BAs). He defined BA1-BA6 which were later assigned absolute depths by Brett et al. (1993). The addition of BA6 includes graptolitic black shale formed in offshore environments. Benthic assemblage zones have been widely used for more than four decades, without major conceptual changes, to infer transgressive–regressive cycles, primarily in the Silurian (e.g., Johnson et al., 1981; Landing and Johnson, 2003), but also in the Ordovician (McKerrow, 1979). The parallel improvement of biostratigraphy has increased the validity of such curves. The method has been subject to some criticism although (e.g., Jeppsson, 1990), and a recent evaluation of the method through 393 comparison with δ18O data of the component species (Azmy et al., 2006) has shown that care is required, especially in outer shelf environments where overlap between the pre-defined assemblages may occur. In addition, also factors such as shelf morphology (clastic shelf versus carbonate platform), turbidity of sea water, oxygenation, water–energy, or nutrient conditions affect the distribution of benthic communities and thereby the output of the method (the original model was based on a clastic shelf setting). Several other groups are also valuable for analysis of bathymetry. In benchmark studies for the Ordovician (Fortey, 1975) and Silurian (Thomas, 1979), trilobite communities were related to depth gradients; these models have been subsequently modified for a wide range of Early Palaeozoic settings. Pelagic graptolites show an increase in diversity from nearshore to offshore marine environments (see e.g., Berry and Boucot, 1972 and the ‘graptolite assemblage scheme’ of Chen, 1990), and can thus form a basis for the establishment of trends in sea level. This approach was used by Egenhoff and Maletz (2007) to recognise a series of maximum flooding surfaces in the Lower Ordovician shelf succession of Baltoscandia. Based on the comparison between graptolite and brachiopods Boucot and Chen (2009) demonstrated that fossil plankton, too, can be used as depth indicators for Palaeozoic strata. An alternative conceptual method for carbonate-dominated shelves, although rarely used, is conodont community analysis (Zhang and Barnes, 2002; Zhang et al., 2006). Biological data are highly useful since they can be statistically analysed and assigned to pre-defined depth zones. Nevertheless large and representative samples are required to ensure that conclusions are valid. Moreover it is virtually impossible to prove if water depth or a function of water depth, such as ‘energy levels’ is the factor controlling the distribution of marine benthos. 2.2.4. Geochemical proxies In the last decade geochemical proxies have also played an increasingly important role in tracking sea-level change since the chemical fractionation of 16O during the formation of continental ice sheets leave a surplus of 18O in the oceans, which then is incorporated into the structure of calcitic shells or apatite of tooth elements (conodonts) (see Section 3). 2.3. Ordovician The Ordovician Period records the highest sea levels of the entire Palaeozoic (Haq and Schutter, 2008) and was therefore characterised by extensive continental submergence. In this sense the period is unique since many of the settings we study are fundamentally different from ‘the typical’ shelf seas that most sea-level models are based on. This is well exemplified in Baltoscandia were the Floian– Dapingian Ortoceratite Limestone formed in an extensive basin of very low relief, inherited from the underlying sub-Cambrian peneplain that forms the basement. The very slow sedimentation rates, fold structures, and the exceptional continuity of distinct hardgrounds in the Orthoceratite Limestone are a few of the reasons that depositional depths in the range of several hundreds of metres have been inferred (Lindström, 1963; see also Nielsen, 2004), a figure that is exceptionally high in cratonic interiors. Although a few studies have assigned sequence stratigraphic terminology to this type of deposits they still represent a huge questionmark in terms of sea-level change. In comparison to the Silurian, relatively few sea-level curves exist for the Ordovician. More recent sea-level curves covering all or parts of the Ordovician have been published for Avalonia (Woodcock, 1990), Laurentia (Ross and Ross, 1992, 1995), Western Gondwana (Heredia and Beresi, 1995), Baltoscandia (Dronov and Holmer, 1999; Nielsen, 2004; Dronov, 2005), the Yangtze Platform (Su, 2007), and Siberia (Kanygin, et al., 2010-this volume). More detailed intercontinental correlations have been proposed by Nielsen (2004) and by Su (2007), but a standard global curve has not been produced. The Author's personal copy Age (million years) A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 GLOBAL SERIES & STAGES 394 Laurentia Baltoscandia Baltoscandia Avalonia Western Gondwana Yangtze Platform Siberia (Ross and Ross 1995) (Nielsen 2004) (Dronov 2005) (Woodcock 1990) (Heredia and Beresi 1995) (Su 2007) (Kanygin et al. 2010this volume) High Low High Low High Low High Low High Low High Low High Low 480 KATIAN SANDBIAN DARRIWILIAN DAPING. FLOIAN LOWER ORDOVICIAN 470 MIDDLE ORDOVICIAN 460 TREMADOCIAN 450 UPPER ORDOVICIAN HIR. SIL. Fig. 1. Compilation of different sea-level reconstructions for the Ordovician. most detailed Ordovician curve is based on Baltoscandia (Nielsen, 2004) and, thus, does not necessarily reflect the global pattern. The curve is based on facies and palaeontological observations and interprets the classical carbonate mud mounds of Baltoscandia as reflecting periods of sea-level lowstand (see opposite interpretation by Calner et al., 2010). Nielsen (2004) assumed a magnitude-range of 250 m between extreme lowstand and highstand situations and subdivided the Ordovician into three extended lowstand intervals and three extended highstand intervals, each superimposed by several shorter-term regressions and transgressions. In terms of general trends, there is good degree of agreement in the published sea-level curves (Fig. 1) from Laurentia (Ross and Ross, 1995), Baltoscandia (Nielsen, 2004), the Yangtze platform (Su, 2007), and the very recent curve from Siberia (Kanygin, et al., 2010-this volume). On a broader scale, global sea level rose through the Early Ordovician to peak for the first time in the Floian. This interval was followed by a fall and relatively low sea level through the Dapingian and Darriwilian stages. An extended interval of global lowstand occurred in the mid to latest Darriwilian (Ross and Ross, 1995; Nielsen, 2004; Su, 2007; Kanygin et al., 2010-this volume) or a little later, in the Sandbian (Woodcock, 1990; Dronov, 2005). The Mid Ordovician lowstand was followed by a further highstand in the Katian before the global drop in the Hirnantian. The sea-level curves from Avalonia (Woodcock, 1990) and Western Gondwana (Heredia and Beresi, 1995) deviate most from this general pattern (Fig. 1), requiring additional explanations. 2.3.1. Early Ordovician Sea-level data from all investigated palaeocontinents indicate an overall transgression in the Early Ordovician, confirming the global trend. This transgression has been observed in Baltica, Siberia, Laurentia, and South China, and seems also be present on other parts of Gondwana and Avalonia. The contrasting curves from Baltica are based on different methods of investigation and focused on different parts of the Baltoscandian basin. However, both Dronov (2005), for the Baltic region (St. Petersburg area), and Nielsen (2004), focusing on the Scandinavian part, documented a transgression in the earliest Ordovician Pakerort Stage, followed by a regression at the base of the Varangu Stage. It appears that transgression across Baltica continued, interrupted by smaller regressions, up to the latest Floian, when highest sea levels of the Early Ordovician were reached. In South China, this overall transgression on the Yangtze Platform is very well documented. It includes the entire Early Ordovician and continued up to the base of the Darriwilian (e.g., Su, 2007). This transgression is also observed on Laurentia, where Ross and Ross (1992, 1995) illustrated an overall transgressive trend from the earliest Ordovician to the middle Floian. Prior to the Early–Middle Ordovician boundary regression occurred, thus during the middle Floian the highest sea levels were recorded, similar to Baltica. A new sea-level curve from Siberia (Kanygin et al., 2010-this volume) can partly be correlated with the trends demonstrated from Baltica and Laurentia, located both at similar latitudes. Moreover, a general transgression has been observed in the Lower Ordovician of Siberia, with highest sea levels in the middle Tremadocian and the middle Floian. The curves from Gondwana and the peri-Gondwanan ‘terranes’ must be considered preliminary, and additional information is needed before accurate curves can be drawn. 2.3.2. Middle Ordovician At a global level (Miller et al., 2005; Haq and Schutter, 2008), continuous sea-level rise during the Early Ordovician ceased during the Middle Ordovician. Haq and Schutter (2008), for example, documented that Middle Ordovician sea levels were similar to the high levels of the Floian, and that a further overall transgression Author's personal copy Age (million years) GLOBAL SERIES & STAGES A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 395 Baltica Gondwana (NSW) Global curve Global curve Global curve Global curve American Midwest (Johnson 1996) (Johnson 1996) (Loydell 1998) (Azmy et al. 1998) (Haq and Schutter 2008) (Johnson 2010-this volume) (Spengler and Read 2010) High Low High Low High Low High Low High Low High Low High Low LUD. SHEIN. HOMERIAN GOR. WENLOCK 420 LUDLOW PRIDOLI DEV. 440 AERONIAN RHUDDANIAN LLANDOVERY TELYCHIAN 430 Fig. 2. Compilation of different sea-level reconstructions for the Silurian. Data from Johnson (1996, 2010-this volume), Loydell (1998), Azmy et al. (1998), Haq and Schutter (2008) and Spengler and Read (2010). characterised the latest Middle Ordovician (latest Darriwilian). At a regional (continental) level, some authors noted even regressive events, with sea levels falling after the Floian maximum. Ross and Ross (1992, 1995), for example, illustrated a strong regression on Laurentia, with sea levels reaching their lowest levels during the middle part of the Darriwilian. Nielsen (2004) observed a similar trend on Baltica, while Dronov (2005) observed remaining high levels, comparable to the global curve (Haq and Schutter, 2008). Both Siberia and South China also document regressions in the Darriwilian, but sea level rose again significantly before the Middle-Upper Ordovician boundary. Although far from being complete, the sea-level curves for the different palaeocontinents and also the global curve indicate a rapid trangression below the Middle-Upper Ordovician boundary. transgression starting in the Darriwilian up into the late (but not latest) Katian (Ross and Ross, 1992, 1995). The Hirnantian, especially, is particularly well investigated, more recently by, e.g., Dahlqvist and Calner (2004), Armstrong et al. (2005, 2006, 2009b), Rey et al. (2005), Brenchley et al. (2006), Le Heron (2007), Le Heron and Craig (2008), and Loi et al. (2010-this volume). Of particular interest have been the sea-level changes during the glaciation on Gondwana, but also on other continents (e.g., Desrochers et al., 2010-this volume). Despite this interest, the intraHirnantian sea-level changes are not fully understood, although two regressive phases and an intermittent minor transgression are now inferred in some areas (Nielsen, 2004; Brenchley et al., 2006). 2.4. Silurian 2.3.3. Late Ordovician The highest sea levels of the Palaeozoic have been recorded in the lower part of the Upper Ordovician (e.g., Miller et al., 2005; Haq and Schutter, 2008). Following the beginning of an overall transgression in the upper part of the Middle Ordovician, it appears that sea levels continued to rise up to the middle part of the Katian. The subsequent sea-level fall was not, however, abrupt in the uppermost part of the Ordovician, related to the rapid glaciation on Gondwana during the Hirnantian, but probably commenced much earlier. This sea level trend is recorded from most palaeocontinents. For Baltica, both Nielsen (2004) and Dronov (2005) noted the highest sea levels in the middle part of the Katian. The Laurentian trend indicates a A substantial number of Silurian sea-level curves have been published in the last two decades; for example, Johnson (1987), Johnson et al. (1991, 1997), Ross and Ross (1996), Johnson (1996), Tesakov et al. (1998), Loydell (1998), Artyushkov and Chekhovich (2001, 2003), Lazauskiene et al. (2003), Landing and Johnson (2003, includes tens of relative sea-level curves from various palaeocontinents), Johnson (2006), Antoshkina (2007), Brett et al. (2007, 2009), Haq and Schutter (2008), and Johnson (2010-this volume). In addition to these curves, which cover all or most of the Silurian, there have been numerous curves dealing with parts of the system, which will not be discussed in detail here. The perhaps most cited global curves over the last years are those Author's personal copy 396 A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 of Ross and Ross (1996), Johnson (1996, updated 2006), and Loydell (1998). These three curves are based on different techniques and partly show a strongly conflicting pattern. The recent curve by Haq and Schutter (2008) is herein compared to one of these curves (see Johnson, 2010-this volume). On a broad scale, global sea level rose through the Early Silurian and thereafter declined through the Late Silurian, the latter possibly related to a retardation of the long-term tectono–eustatic cycle (Johnson, 1996) (Fig. 4). Peak levels were reached in the late Telychian (Ross and Ross, 1992; Johnson, 1996; Loydell, 1998) or possibly in the early Homerian (Haq and Schutter, 2008). Based on the identification of depositional sequences, Ross and Ross (1996) concluded that intra-Silurian sea level fluctuated by no more than 60 m. Their curve was based exclusively on sections within Laurentia, including a comparison with Baltica. Since these areas were part of the same continent (Laurussia) it is questionable if their curve actually depicts a global pattern. Markes E. Johnson and his collaborators have played a key role in the research on Silurian sea-level change and provided a huge set of sea-level curves from several continental blocks (e.g., Johnson and McKerrow, 1991; Johnson, 1996; Johnson et al., 1997; Tesakov et al., 1998; Baarli et al., 2003). These curves have the advantage of being tied to the same generalised graptolite zonation (Koren et al., 1996) and on the same principal method (Benthic Assemblage zones). The resulting global Silurian sea-level curve (Johnson, 1996; updated by Johnson, 2006 and refined further by Johnson, 2010-this volume) records ten highstands in the Silurian of which those in the Llandovery have been associated with interglacials in Gondwana (Johnson, 1996). Several of the highstands were calibrated against buried coastal topography by Johnson et al. (1998) showing that transgressions in the Silurian ranged from a magnitude of several tens of metres to more than 70 m. The Early Silurian sea-level curve by Loydell (1998) was based on recurrent incursion of graptolitic dark grey laminated shale and mudstone into shelf sequences. Due to the high abundance of graptolites in this facies the corresponding transgressions could be dated with exceptional precision, even down to subzonal level (corresponding to a few 100 kyr). As discussed, though, by Loydell (1998), the regressions were more difficult to date because of the rarity of graptolites and domination by long-ranging taxa. Silurian shallow-water sediments are best dated by conodonts and chitinozoans; in the late 1990s, however, these corresponding biostratigraphic schemes were poorly integrated with graptolite biostratigraphy. This integration has improved considerably in the last decade and not least chemostratigraphy has helped drive the integration of the shelly and graptolite biofacies (Cramer et al., 2010b). The recent sea-level curve produced by Haq and Schutter (2008) shows fifteen highstands in the Silurian and implies fluctuations in the order of 140 m (see discussion by Johnson, 2010-this volume). 2.4.1. Llandovery sea-level changes Reworked glaciogenic sediments and ice-produced structures of Early Silurian age imply that climate continued to exert a strong control on sea-level development after the Hirnantian glaciation. Evidence for continental ice-sheets include data from southern Libya, and the Amazonas Basin, the Parnaiba Basin, and the Peru-Bolivia basins of South America (e.g., Grahn and Caputo, 1992; Caputo, 1998; Díaz-Martínez and Grahn, 2007). Based on the available biostratigraphical data, the glacial maxima occurred in the early Aeronian (gregarius-?magnus Zone), late Aeronian (sedgwickii Zone), and late Telychian (lapworthi-insectus zones) (see summary by Page et al., 2007). The first maximum is thus very close to or even overlaps, temporally, with the first highstand of the Silurian, indicated at or near the boundary between the Rhuddanian and Aeronian stages (Johnson, 2006) or slightly later, in the triangulates to (?lower) magnus biozones (Loydell, 1998). A second highstand occurred near the convolutus-sedgwickii zonal boundary (Johnson, 1996; Loydell, 1998), and thus just predates the second glacial maximum of Page et al. (2007). This is in contrast to a rigorous conodont community analysis from Anticosti (cf., Zhang and Barnes, 2002), which suggested a highstand in the lower parts of the convolutus Zone, followed by a major regression peaking in the upper convolutus Zone, which in turn is followed by a transgression associated with the first appearance datum of sedgwickii. Moreover, Ross and Ross (1996) identified a lowstand in the upper convolutus Zone based on analysis of depositional sequences. The discrepancies between the Loydell (1998) and Johnson (2006) curves and those of Zhang and Barnes (2002) and Ross and Ross (1996) are thus related to the stratigraphic position of the lowstand, and might be explained by the use of four investigative different investigative techniques. There is little agreement also between the Telychian sea-level curves, which in part are in very marked contrast. For example, the third highstand of Johnson (2006), at the guerichi-turriculatus zonal boundary, correlates with a well-defined lowstand on the curve of Loydell (1998; his Stimulograptus utilis Subzone sea-level fall). The middle Telychian is characterised by a clear sea-level drop with a maximum lowstand close to the griestoniensis-crenulata zonal boundary (Loydell, 1998; Johnson, 2006). The late Telychian constitutes the third glacial maximum of Page et al. (2007). This is consistent with the sea-level curve of Loydell (1998) which shows a major lowstand in the Cyrtograptus lapworthi graptolite Zone (basal amorphognathoides conodont Zone) with relatively low sea levels continuing into the earliest Wenlock. Johnson (2006) argued that his fourth highstand peaks near the boundary between the celloni and amorphognathoides zones, just before Loydell's lowstand. Hence, the discrepancy between these two curves is necessarily not as large as it may seem and might be a question of biostratigraphic correlation or the influence of tectonics. The curve of Ross and Ross (1996) shows a clear lowstand, somewhat later, in the upper amorphognathoides Zone. Geochemical data are not consistent with the correlation of the third glacial maximum with the late Telychian. Oxygen isotopes from brachiopods (Brand et al., 2006) and from conodont apatite (Lehnert et al., 2010this volume) suggest warm conditions in the late Telychian and the onset of cool glacial conditions in the earliest Wenlock. 2.4.2. Wenlock sea-level changes Published post-Llandovery curves show less variation and less magnitude in sea-level changes than those for the Llandovery, but this interval is also less well known. Apart from possible glacial deposits of earliest Wenlock age (Loydell, 2007, p. 543) sea-level changes cannot presently be tied to any geological evidence for glaciation such as tillites. The most cited sea-level curves continue to show contrasting patterns for the earliest Wenlock. Loydell (1998) suggested that rising sea level through the latest Telychian reached its highest point for the entire Sheinwoodian in the murchisoni graptolite Zone. This highstand, however, correlates with the lowstand proposed by Johnson (2006) for the time-equivalent procerus conodont Zone. The next major highstand occurred in the riccartonensis Zone and may (separated by a brief phase of minor regression) have continued into the early Homerian (Johnson, 2006). This relatively extended highstand thus overlaps with the middle Sheinwoodian rigidus Zone highstand identified in many areas of the world (Johnson, 1996; Loydell, 1998). The middle Homerian is particularly well studied because of the international interest in the positive δ13C excursion and the faunal extinctions associated with the Mulde Event (Cramer et al., 2006a; see review by Calner, 2008; Barrick et al., 2009). This coupled isotopic–biotic event is associated with major facies changes, suggesting a distinct regression in the uppermost Cyrtograptus lundgreni graptolite Zone, on several palaeocontinents including Laurentia, Baltica and peri-Gondwana, and independent of basin type (Calner, 2008). There is some evidence that this sea-level drop was particularly large. On Gotland, even the most distally preserved parts of the carbonate platform were exposed, subaerially (Calner, 2002), and a minimum of 16 m of palaeorelief is preserved at a coeval level in the more proximal parts of the platform (Calner and Säll, 1999). The lowstand was followed by a marked transgression (the Author's personal copy A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 largest in the Silurian according to Haq and Schutter, 2008) in the Pristiograptus parvus and Gothograptus nassa zones (on Gotland the LAD of flemingii occurs just below the karst surface, and the FAD of nassa is close above the karst surface; see Calner et al., 2006). This transgression resulted in the widespread deposition of fossiliferous mudstones in Laurussia; the Mulde Brick-Clay Member on Gotland (Calner and Jeppsson, 2003), and the Waldron Shale in the American Midwest (see correlation by Cramer et al., 2006a). A very similar unit has recently been documented from the northern Midland Platform of central England (Ray et al., 2010). The Wenlock Epoch ended with a regression, which is locally represented by a major unconformity (Eriksson, 2004). 2.4.3. Ludlow sea-level changes The Ludlow commenced with a major transgression that peaked near the top of the Neodiversograptus nilssoni graptolite Zone (Johnson, 1996; Loydell, 1998). According to Johnson (2006) the next highstand occurred within the Ludfordian and peaked in the snajdri Zone. This is in sharp contrast to recent data from Gotland (Calner and Eriksson, 2006; Eriksson and Calner, 2008), the Urals (Mishutina, 2007), and Poland (Kozłowski and Munnecke, in press), which suggest a profound sea-level drop at this time. The sea-level fall started close to the LAD of P. siluricus and low sea levels continued through the Icriodontid Zone (sensu Jeppsson, 2005), in some areas with a minor intervening transgression (Eriksson and Calner, 2008). The main post-lowstand transgression occurred in the lowermost O. snajdri conodont Zone. 2.4.4. Pridoli sea-level changes The Pridoli constitutes only ca 2.7 million years of the Silurian Period and sea-level trends are poorly known. Johnson (2006) placed his youngest Silurian highstand in the Pridoli, although a more precise correlation has not been possible. 2.5. Summary of Ordovician and Silurian sea-level development The Ordovician and Silurian sea-level curves have been based on a variety of different techniques, and a few studies have attempted to integrate a number of these different techniques (e.g., Azmy et al., 2006). The most cited curves are in general agreement but show contrasting patterns for several of the major sea-level changes (Figs. 1, 2). The reason for this can possibly be found in the investigative methods selected, inadequate biostratigraphic control in less-well studied areas, the lack of integration of conodont and graptolite biostratigraphical schemes, the erroneous usage of terminology, or even the way the symmetry of sea-level curves are drawn (see, e.g., discussion in Ray et al., 2010). The many discrepancies between these curves reinforce the need for a common language and a common stratigraphic framework. Koren et al. (1996) published a generalised graptolite zonation for the Silurian, which primarily was designed to form the basis for the construction of palaeogeographic maps for the 1996 James Hall Meeting. This scheme was extensively used in the compilation of Silurian sea-level curves by Landing and Johnson (2003). This has the advantage that zones are easily identified even in areas with little biostratigraphic data, but the biostratigraphic resolution is significantly reduced, since some of these generalised zones combined two or more regional biostratigraphic biozones. The recent compilation of Silurian stratigraphy by Cramer et al. (in press) may form a more robust stratigraphic basis for future analyses. 3. Geochemical proxies (δ13C, δ18O, 87 Sr/86Sr) 3.1. Introduction In the past few decades geochemical proxies have become powerful tools for both palaeoenvironmental reconstructions and stratigraphic correlations. Many environmental changes are reflected 397 by changes in the geochemical composition of the ocean and atmosphere, and because of the (geologically) short mixing time within the ocean-atmosphere system, which is in the order of thousands of years, the geochemistry of sedimentary rocks and fossils can be used as proxies to reconstruct these changes. Among the different methods available, the stable isotope geochemistry is the most powerful, especially regarding chemostratigraphy. Palaeozoic rocks, however, very often have experienced significant diagenetic alteration which can modify the original chemical signature in the rock. Nevertheless, careful investigations have shown that (a) some proxies are rather resistant to diagenetic changes, and (b) the degree of alteration can be assessed by a range of different methods, e.g., trace element analysis, SEM, and cathodoluminescence (e.g., Holser, 1997; Samtleben et al., 2001; Brand, 2004). In the following section, a brief summary of the possibilities and pitfalls of some of the most widely used geochemical proxies used in the Early Palaeozoic is presented (δ13Ccarb, δ13Corg, δ18O, 87Sr/86Sr). For more detailed reviews the reader is referred to, e.g., Berger and Vincent (1986), Kump and Arthur (1999), Goddéris et al. (2001), Bickert (2006), Weissert et al. (2008), and Immenhauser et al. (2008). 3.1.1. δ13Ccarb values In nature, the stable carbon isotope 12C is more abundant (98.89%) than 13C (1.11%; Craig, 1953). During photosynthesis 12C is preferentially incorporated into organic material, and marine organic matter therefore is strongly depleted in 13C (ca.−25‰ δ13C) (Hayes et al., 1999). Hence, ocean surface water dissolved inorganic carbon (DIC) is usually enriched in 13C because phytoplankton preferentially remove 12 C, whereas deeper water is depleted in 13C because nearly all of the organic matter produced at the surface is remineralised by bacteria. In anoxic oceans, like the modern Black Sea, the fractionation between surface water and deep water is considerably greater because a large portion of organic matter is not remineralised and is deposited as sapropels (Fry et al., 1991). Enhanced δ13Ccarb values therefore are often explained in terms of either increased productivity (e.g., Wenzel and Joachimski, 1996) or enhanced burial of (isotopically light) organic matter (e.g., Cramer and Saltzman, 2005). In contrast to δ18O values (see below) the δ13Ccarb values from carbonate rocks are usually less affected by diagenetic changes because in many cases the system is more or less closed for carbon (rock-buffered) and the pore fluids contain very little carbon which potentially could change the isotopic composition of the rock. Even dolomites can preserve the original carbon isotopic composition (Ling et al., 2007; Kaminskas et al., 2010). However, care has to be taken especially when the rocks have been exposed subaerially to the potential influence of soil-derived fluids (Joachimski, 1994), and in rocks with a high content of organic matter because the latter has very low δ 13C values which might alter (lower) the rock values (Immenhauser et al., 2008). In addition calcite-cemented sandstones should not be used because of the uncertain origin and age of the carbonatic cement. The most reliable δ13C-curves have been reconstructed by analysing well-preserved brachiopod shells, because brachiopods are composed of diagenetically rather stable lowmagnesium calcite (e.g., Samtleben et al., 1996, 2000; Wenzel and Joachimski, 1996; Azmy et al., 1998; Shields et al., 2003; van Geldern et al., 2006). Analysing rock samples, however, enables much higher resolution and continuous sampling, which is nearly impossible if brachiopods are used. In most cases the curves reconstructed from rocks broadly follow the brachiopod values (Kaljo et al., 1998, 2001, 2004, 2007a,b; Cramer et al., 2010b). Cramer et al. (2010a) have demonstrated the value of high-resolution δ13C stratigraphy for global correlation in the Palaeozoic. Based on chemostratigraphic and sequence stratigraphic correlation, the authors show that the first occurrence of the Silurian conodont Kockelella walliseri within the stratigraphic sequences of Laurentia is essentially one full stratigraphic sequence lower than in Baltica. Author's personal copy 398 A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 In the Palaeozoic several δ13Ccarb excursions with amplitudes of more than 5‰ have been reported (see below). Most of these excursions are recorded from different palaeocontinents, thus clearly indicating that they are not the result of a diagenetic alteration but signal global environmental changes. The absolute values and the amplitudes, however, may vary from section to section, which may be the result of differences in δ13C of the DIC of shallow epeiric seas and the open ocean (Holmden et al., 1998; Panchuk et al., 2005; Melchin and Holmden, 2006b). It is interesting to note that many of the Palaeozoic excursions are significantly greater than the excursions observed in the Mesozoic and Cenozoic. 3.1.2. δ13Corg values As outlined above 12C is preferentially incorporated in organic material during photosynthesis, and marine organic matter therefore is strongly depleted in 13C. Most of the modern warm-water plankton, for example, exhibits δ13Corg values between − 17 and − 22‰ (Bickert, 2006). Both the δ13C composition of organic matter and that of inorganic carbon depend on the isotopic composition of the dissolved inorganic carbon (DIC) in sea water. Therefore, any change in the isotopic composition of the DIC should result in parallel changes in δ13Ccarb and δ13Corg. The amplitudes, however, may be different because the sedimentary record of δ13Corg of the total organic carbon is controlled by several factors, e.g., light intensity, organism cell geometry, heterotrophic reworking, species composition of the phytoplankton community, and input of terrestrial organic carbon (Hayes et al., 1999). Paired analyses of δ13Ccarb and δ13Corg (Δδ13C) are used as a proxy for the pCO2 of the atmosphere because the Δδ13C values are to a large degree controlled by photosynthetic fractionation (εp), which is in part dependent on the concentration of dissolved CO2 in sea water (e.g., Popp et al., 1989; Freeman and Hayes, 1992; Hayes et al., 1999; Kuhn, 2007). There is, however, an ongoing discussion on the applicability of this proxy (see review in Bickert, 2006, p. 319). Differences in peak magnitudes between δ13Ccarb and δ13Corg may also reflect changes in the composition and abundance of certain isotopically distinct sources of organic matter and in the bulk sedimentary organic carbon (Fanton and Holmden, 2007). Investigations of the isotopic composition of organic matter are mostly carried out, either because the rocks do simply not contain enough carbonate for δ13Ccarb measurements (e.g., Underwood et al., 1997), or because information is required with respect to (palaeo-) productivity and/or pCO2 (e.g., Cramer and Saltzman, 2007; Gouldey et al., 2010-this volume). Due to the fact that δ13Corg values are potentially affected by a large number of different primary (e.g., heterogeneity of organic matter) and secondary processes (including thermal alteration and migration of hydrocarbons) their values are often highly variable and thus their stratigraphic use is limited compared to δ13Ccarb values (see discussion in Delabroye and Vecoli, 2010). 3.1.3. δ18O values In nature, oxygen occurs mainly in form of the 16O isotope (99.8%), the 18O isotope is much less abundant. The ratio of 18O/16O (given in delta notation) in naturally formed calcium carbonates depends mainly on the isotopic composition of the surrounding seawater and on the temperature during precipitation. Also the mineralogy of the precipitates (calcite vs. aragonite) is of importance (Grossman and Ku, 1986). A one per mil change in δ18O values corresponds roughly to a change in temperature of 4 °C (Shackleton, 1987). Rain and snow, and therefore also ice caps, contain water with low δ18O values down to − 50‰ (due to an enrichment in 16 O along with a Raleigh fractionation). Consequently, the global ocean water is enriched in 18 O during intervals of glaciation. For the late Pleistocene, an increase in δ18O values of 1‰ indicates a global, glacially-induced sea-level drop of ca. 100 m. (Shackleton, 1987). In addition, the δ18O values of seawater are also affected by fractionation effects due to evaporation and precipitation at the sea surface, admixture of water masses containing other 18O/16O ratios, e.g., melt water and alluvial and meteoric runoff, and the global isotope content of the oceans (Craig and Gordon, 1965). Because changes in these processes also affect the salinity of the ambient seawater, the δ18O values in modern oceans show a correlation with salinity, varying between 0.1 for tropical and 1.5 for polar surface water masses, with a global mean of 0.49 (Craig and Gordon, 1965). Changes in δ18O values are therefore not easy to interpret since they indicate usually a combination of changes in temperature and the hydrological cycle. Modelling the correlation of δ18O values with palaeo-salinity is difficult especially for the Palaeozoic because of the totally different plate tectonic configuration and oceanic circulation patterns of today's oceans (Cocks, 2001), and the – in most cases – unknown salinity of the ocean water. Hay et al. (2006), for example, calculated the Ordovician and Silurian ocean salinity to be roughly 10‰ higher than that in modern oceans. In general, the δ18O values of ancient (especially Palaeozoic) sea water are a very controversial issue (see review in Wallmann, 2001) because the measured values are significantly lower than those of modern oceans (Veizer et al., 1999). Several authors argue that the value of seawater is buffered by both seawater/rock interactions at the midocean ridges and by continental weathering and recycling of subducted water, and therefore the values should be rather stable throughout Earth History (Muehlenbachs, 1986). Others argue that, e.g., changes in the proportion of high-temperature processes at mid-ocean ridges can result in a long-term secular trend in the oxygen isotopic composition of seawater (Veizer et al., 1999; Wallmann, 2001). A further complication is the fact that δ18O values are often diagenetically altered and therefore do not reflect the original signal. Because pore fluids always contain large amounts of oxygen (in their water molecules; high oxygen fluid/rock ratios) the oxygen isotopic composition is much more prone to diagenetic alteration compared to that of carbon isotopes. Because both micritic and sparitic lithified limestones consist largely of diagenetically precipitated calcium carbonate cement (Bathurst, 1975) their oxygen isotope values should not be used for palaeoenvironmental reconstructions (Weisert et al., 2008). There is an ongoing discussion regarding which fossils or precipitates are the best carriers of marine oxygen isotopic compositions in Palaeozoic rocks. A large number of authors agree that brachiopods represent the most reliable fossils (e.g., Samtleben et al., 1996, 2000, 2001; Wenzel and Joachimski, 1996; Azmy et al., 1998, 2006; Veizer et al., 1999; Goddéris et al., 2001; Shields et al., 2003; Ernst and Munnecke, 2009) whereas others argue that biogenic phosphate, e.g., conodont elements, is more reliable (Wenzel et al., 2000; Lehnert et al., 2007b,c, 2010-this volume; Zigaite et al., 2010). Phosphates were thought to be rather resistant to diagenetic alteration, however, there is increasing evidence that early diagenesis associated with microbial activity can modify the isotopic composition of phosphatic components (Sharp et al., 2000). It is, however, interesting that despite differences in amplitudes and absolute values, the δ18O values measured from brachiopods and conodonts show broadly parallel trends, at least in the Ordovician and Silurian (see below, Figs. 3, 4). 3.1.4. 87Sr/86Sr values Strontium has four stable, naturally occurring isotopes (84Sr, 86Sr, 87 Sr, 88Sr). Among these isotopes only 87Sr is radiogenic. It is produced by the decay of rubidium-87. Strontium has a long residence time in the ocean of about 2.4 × 106 years (Faure, 1986). In contrast to the carbon isotopes, the Phanerozoic 87Sr/86Sr curve is characterised by long-term fluctuations, and the values are widely used as a chemostratigraphic proxy (Qing et al., 1998; McArthur and Howarth, 2004). In order to minimise possible diagenetic effects on 87Sr/86Sr data most authors use the lowermost values for a given time slice because of the tendency of diagenetic alteration to increase the values (Veizer and Compston, 1974; Shields et al., 2003). Author's personal copy A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 399 Fig. 3. Synoptic presentation of different Ordovician geochemical proxies. The 87Sr/86Sr isotopic composition of seawater is primarily controlled by the subaerial weathering of continental crust (releasing more radiogenic Sr into the ocean) or young basaltic rocks (providing more nonradiogenic Sr), by halmyrolytic alteration of young basaltic crust, and by precipitation and weathering of marine carbonates (Faure, 1986). Therefore, changes in the 87Sr/86Sr isotopic composition of the ocean through time can be used as a proxy indicator of global tectonic evolution (e.g., Veizer et al., 1999). 3.2. Ordovician δ13Ccarb development In the past 15 years, numerous papers have been published on stable carbon isotopes from Ordovician carbonate rocks. Most studies were carried out on sections in Baltica (Ainsaar et al., 1999; Kaljo et al., 2001, 2003, 2004; Brenchley et al., 2003; Schmitz and Bergström, 2007; Hints et al., 2010; Bergström et al., 2010a; Bergström et al., 2010b-this volume) and Laurentia (Long, 1993; Buggisch et al., 2003; Ludvigson et al., 2004; Young et al., 2005; Melchin and Holmden, 2006a; Fanton and Holmden, 2007; Buggisch, 2008; LaPorte et al., 2009; Ainsaar et al., 1999, 2010; Young et al., 2010-this volume), but also in Gondwana (Marshall et al., 1997), India (Suttner et al., 2007), Korea (Hong et al., in press), and China (Wang and Yang, 1994; Yang and Wang, 1994; Jiang et al., 2001; Young et al., 2008; Bergström et al., 2009b; Wang et al., 2009). A generalised δ13Ccarb curve was published by Bergström et al. (2009a) based on data from Argentina (Tremadocian to Dapingian), Estonia (Darriwilian to Sandbian), and North America (Katian to Hirnantian) (Fig. 3). This curve shows varying, but overall decreasing values in the Tremadocian. In the Floian, Dapingian, Darriwilian and Sandbian the values remain relatively constant (mostly between 0 and + 1‰) with the exception of a small positive excursion in the mid-Darriwilian. The Katian shows more variable values with up to five small excursions. The largest positive excursion is reported from the Hirnantian, with peak values of N6‰ in Baltica (Kaljo et al., 2001; Brenchley et al., 2003; Ainsaar et al., 2010) and Laurentia (Schmitz and Bergström, 2007; Ernst and Munnecke, 2009). Several authors have described a gradient in δ13C values with higher values in proximal settings, and lower values in more distal settings (Kaljo et al., 2004; Melchin and Holmden, 2006a; LaPorte et al., 2009). Melchin and Holmden (2006a) and LaPorte et al. (2009) have explained the differences in peak magnitudes between proximal and distal settings by local carbon cycling in tropical and subtropical epeiric seas. On a regional scale, δ13Ccarb curves represent excellent parastratigraphic tools and are extremely useful for stratigraphic correlations, e.g., in North America and in Baltica (e.g., Ludvigson et al., 2004; Bergström et al., 2010a; Bergström et al., 2010b-this volume). Up to now, at least five of the Ordovician positive δ13Ccarb excursions have been reported from more than one palaeocontinent which may be used for intercontinental correlation as an alternative to classical biostratigraphy (see summaries in Bergström et al., 2010a; Bergström et al., 2010b-this volume, and Ainsaar et al., 2010): the midDarriwilian excursion, the early Katian Guttenberg (GICE) and Kope excursions, the mid-Katian Waynesville excursion, and the Hirnantian excursion (HICE). It seems possible that the other excursions in the Author's personal copy 400 A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 Fig. 4. Synoptic presentation of different Silurian geochemical proxies. Late Ordovician also represent global perturbations of the carbon cycle. For a reliable correlation, however, more precise biostratigraphic data are required. Even for the Hirnantian, which includes one of the largest short-term δ13Ccarb excursions (HICE) of the Phanerozoic (δ13Ccarb up to 8‰; Schmitz and Bergström, 2007), there is no general agreement on the precise correlation of δ13Ccarb values with biostratigraphy (Melchin and Holmden, 2006a; Hints et al., 2010). It is not the topic of the present paper to discuss these complex problems; for a further discussion the reader is referred to the extensive reviews by Kaljo et al. (2008), Melchin (2008), Delabroye and Vecoli (2010), and Hints et al. (2010). Most authors agree, however, that the peak values of the HICE correlate with the N. persculptus graptolite Biozone. The causes of the Ordovician δ13Ccarb development remain unclear. Vascular plants had not yet evolved, and the contribution of the bryophytes to the global carbon cycle was slight. Several authors have speculated that sea-level changes were the major driving force (e.g., Kump et al., 1999; Buggisch et al., 2003; Kaljo et al., 2003; Fanton and Holmden, 2007). However, Bergström et al. (2010a) show that some excursions occur in transgressive intervals whereas others are related to regressive strata; a simple connection between sea level and δ13Ccarb excursions is not obvious (compare Figs. 1 and 3). Buggisch et al. (2010) have shown that the GICE δ13Ccarb excursion clearly postdates a δ18O excursion measured from conodonts (see below), and therefore a global cooling event cannot be the primary factor, at least for this carbon isotope excursion. This has some support from Sweden where facies support that carbonate mud mounds grew during a transgression contemporaneously with the development of GICE, and were terminated by a distinct regression first when GICE peak values were reached (Calner et al., 2010). Additionally, there is – on a broad scale – no sedimentological and/or palaeontological evidence, neither for enhanced planktonic productivity nor for increased deposition of organic-rich deposits during the δ13C excursions (Melchin and Holmden, 2006a). Most papers that have analysed the Ordovician δ13Ccarb development focus on the Hirnantian excursion which correlates with the secondlargest mass extinction in Earth history (Marshall et al., 1997; Brenchley et al., 2003). The Hirnantian is associated with an extensive glaciation on Gondwana and a globally lowered sea level (see above). Kump et al. (1999) provided a ‘weathering hypothesis’ as explanation for the Hirnantian glaciation and δ13C excursion (see also Villas et al., 2002). These authors argued that the Taconic orogeny caused a long-term decline in CO2 through the increased weathering of silicate rocks, resulting in global cooling and eventually the growth of ice sheets. Increased weathering of carbonate platforms in response to the glacially-induced sea-level drop was regarded as the driving force of the δ13C excursion. Due to the problems of correlating the δ13Ccarb curves from lowlatitude carbonate sections both with glacial sediments from (peri-) Gondwana and with δ13Corg curves (e.g., from the Global Standard Section and Point, GSSP) the precise relationship with the sea-level Author's personal copy A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 change is difficult to assess (see above, and discussion in Delabroye and Vecoli, 2010). A problem is that the maximum drop in sea-level predates the maximum increase in the isotope values, both for δ13C and δ18O. For example, on Anticosti (Canada) Desrochers et al. (2010this volume) reconstruct two major transgression/regression cycles in the Ellis Bay Formation. A first interval corresponding to ‘moderate’ ice sheets on Gondwana is represented by the upper Grindstone and Veleda members on western Anticosti, a second interval corresponding to ‘large’ ice sheets is represented by the upper Lousy Cove Member and the Laframboise Member. The stable carbon isotopes, however, show a minor peak immediately before the first glacial maximum (and slightly decreasing values during the glacial maximum), and a second, major maximum only in the Laframboise Member, during a major sea level lowstand interrupted by a brief transgressive episode. The values in the preceeding upper Lousy Cove Member, however, remain low. Young et al. (2010-this volume) present a paired analysis of Hirnantian δ13Ccarb and δ13Corg from Laurentia and Baltica indicating enhanced pCO2 in the atmosphere resulting from a reduction in silicate weathering. They argue that the elevated pCO2 values eventually led to deglaciation. 3.3. Silurian δ13Ccarb development The first δ13Ccarb curves, which covered nearly half of the Silurian (latest Llandovery to latest Ludlow), were published coincidentally by two different groups, both working on brachiopods from the Silurian of Gotland (Samtleben et al., 1996; Wenzel and Joachimski, 1996). In the 1990s it was not clear whether or not these δ13C curves represented regional or global trends, but initial results from other palaeocontinents (Talent et al., 1993) indicated the global nature of these trends. Three major positive δ13C excursions have been documented in the Wenlock-Ludlow interval (in the Sheinwoodian, Homerian and Ludfordian; Fig. 4). These seminal results initiated a series of studies dealing with Silurian carbon isotopes, most of them from the Baltic States and Gotland (Kaljo et al., 1997, 2003; Azmy et al., 1998; Heath et al., 1998; Wigforss-Lange, 1999; Samtleben et al., 2000; Munnecke et al., 2003; Martma et al., 2005; Kaljo and Martma, 2006) and from North America (Azmy et al., 1998; Saltzman, 2001; Noble et al., 2005; Brand et al., 2006; Cramer et al., 2006a,b,c; Melchin and Holmden, 2006b; Munnecke and Männik, 2009), but also from Podolia (Azmy et al., 1998; Kaljo et al., 2007a; Małkowski et al., 2009), Russia (Wenzel, 1997), the UK (Azmy et al., 1998), Norway (Wenzel, 1997), Poland (Kozłowski and Munnecke, in press), the Carnic Alps (Wenzel, 1997; Buggisch and Mann, 2004), and Australia (Jeppsson et al., 2007). Although differences exist in both the base line values and the amplitudes (Cramer et al., 2010b) these studies have confirmed the global extent of the Silurian δ13Ccarb curve. Even in non-tropical palaeolatitudes, parts of the curve have been confirmed (Hladíkóva et al., 1997; Buggisch and Mann, 2004; Lehnert et al., 2007a), highlighting the stratigraphic value of the Silurian stable carbon isotope curve. In analogy to the work by Bergström et al. (2009a) for the Ordovician, Cramer et al. (in press) have compiled a general Silurian δ13C-curve (Fig. 4) and have subdivided the Silurian into distinct Stage Slices. Since the late 1990s several minor δ13C excursions have been added to the general picture, e.g., in the early and late Aeronian (Wenzel, 1997; Kaljo et al., 2003; Põldvere, 2003), the early Telychian (Kaljo and Martma, 2000; Munnecke and Männik, 2009), and in the Gorstian/Ludfordian boundary interval (Samtleben et al., 2000). To date, among these small excursions only the early Aeronian and early Telychian excursions have been reported from more than one palaeocontinent. Most of the Silurian isotope excursions coincide with distinct lithological and biotic changes. In low latitudes, intervals of high carbon (and oxygen; see below) isotope values are in many cases characterised by the growth of reefs and the formation of extended 401 carbonate platforms (Munnecke et al., 2003), although in appropriate settings reefs also occur during times of low δ13C values (Loydell, 2008), and argillaceous sediments can also be deposited during phases of elevated isotope values, e.g., the Mulde Brick Clay Member on Gotland or the Waldron Member in N America (Samtleben et al., 2000; Calner et al., 2006; Cramer et al., 2006a). Often, the onset of the excursion is associated with an unconformity indicating a significant drop in sea-level (see above; Calner, 1999; Cramer et al., 2006b,c; Eriksson and Calner, 2008; Kozłowski and Munnecke, in press). The sediments deposited during these excursions contain depauperate or impoverished fossil assemblages, especially with respect to conodonts, graptolites, and trilobites. Although each of these events has its own characteristics, their conspicuous similarities indicate similar driving mechanisms (Munnecke et al., 2003). A major challenge for future research is to establish if and how these intervals of C (and O, see below) isotope excursions relate to sea-level changes. As outlined above, the scenario is complex (compare Figs. 2 and 4), and at least some of the excursions contain several sea-level cycles. The amplitudes of the Silurian stable isotope excursions are extremely large compared to Mesozoic and Cenozoic excursions. Classical interpretations such as productivity changes cannot explain these extreme amplitudes (Bickert et al., 1997). The identification of the strongest δ13C excursion of the entire Phanerozoic in this interval (Ludfordian) with δ13C maximum values of up to 12‰ appears especially surprising given the fact that the Silurian previously had been considered a time of relatively stable environmental conditions (Bassett et al., 1991). At least since the early Telychian, the δ13C excursions are associated with biotic extinction events. At the very beginning or even prior to the increase of the isotope values, many groups of organisms are affected. Especially conodonts, graptolites and trilobites, but also acritarchs, chitinozoans, ostracods, brachiopods, corals, and even vertebrates show extinctions, sometimes in a step-wise manner, and organisms living in hemipelagic environments were more strongly affected than organisms occupying shallow-water settings (see review in Munnecke et al., 2003; Stricanne et al., 2006; Eriksson et al., 2009). Another striking feature of the Silurian δ13C excursions is the coincidence of high δ13C values and the formation of microbial carbonates (stromatolites, oncolites) and oolites as reported from various palaeocontinents (Munnecke et al., 2003). Microbial carbonates, which are formed mostly by calcifying cyanobacteria, seem to be more abundant in times of high δ13C values. For example, in the early Sheinwoodian microbial carbonates are reported from Gotland (Samtleben et al., 1996; Nose et al., 2006; Munnecke, 2007) and from the Gaspé Peninsula (Bourque, 2007, Desrochers, pers. commun. 2010), in the Homerian from the Much Wenlock Limestone Formation (Ratcliffe, 1988; Kershaw and Li, 2007) and from Gotland (Calner and Säll, 1999; Samtleben et al., 2000; Calner and Jeppsson, 2003), and in the Late Ludlow from Gotland (Samtleben et al., 2000; Calner, 2005a, b; Nose et al., 2006; Jeppsson et al., 2007; Munnecke, 2007), Scania (S Sweden; Wigforss-Lange, 1999), Poland (Kozłowski and Munnecke, in press), and Australia (Jeppsson et al., 2007). It is interesting to note that also the maximum Hirnantian isotope excursion (HICE) is characterised by microbial carbonates (oncolites) in low-latitude shallow-water settings, e.g., on Anticosti (Lespérance, 1981a,b). Because the onsets of the δ13C excursions are characterised by extinction events (see reviews in Munnecke et al., 2003 and Calner, 2008) whereas the microbial carbonates occur later (close to the maximum values), Calner (2005a) interpreted the Late Ludlow mass occurrence of microbial carbonates as a resurgence of an “anachronistic facies”. Alternatively, a rapid change in the pH of the sea water might have been responsible (the Palaeozoic sea water was less buffered than modern ocean water; Ridgwell, 2005). In contrast to most calcifying organisms, which secrete calcium carbonate inside their cells, calcifying cyanobacteria secrete calcium carbonate outside Author's personal copy 402 A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 their cells, i.e., more or less in contact with the ambient sea water. Therefore, changing pH values of the sea water would have a much stronger effect on the calcification rate of cyanobacteria than on other organisms. This hypothesis, however, has to be tested by future studies. It has been argued that the sudden appearance of microbial carbonates in the Silurian is due to elevated carbonate saturation states in the oceans. But here is an important contradiction: the time periods of elevated saturation state, as published by Riding and Liang (2005a,b), stretch over several millions of years whereas the Silurian microbial resurgences are very short-lived (in the range of a biozone). Hence, there is today no proven correlation between short-term anomalies in saturation state and the development of microbial carbonates, and additional controls need to be considered to explain the Silurian microbial resurgences (Calner, 2005c). Different hypothesis have been proposed in order to explain the Silurian carbon isotope excursions, and to date there is no general agreement on the steering mechanisms. Several authors attribute the excursions to glacial events (Azmy et al., 1998; Kaljo et al., 2003; Brand et al., 2006), with weathering of carbonate platforms during lowered sea level as driving mechanism for the carbon isotope excursion (Noble et al., 2005; Melchin and Holmden, 2006b). Others argue that changes in oceanic circulation driven by climatic changes (e.g., between humid and arid climate) are responsible (Bickert et al., 1997; Munnecke et al., 2003; Martma et al., 2005), or latitudinal changes in the formation of deep water and associated changes in deep-ocean circulation (Jeppsson and Aldridge, 2000, 2001; Cramer et al., 2006c). According to Wigforss-Lange (1999) an increase in photosynthetic activity of cyanobacteria and algae was responsible for the increase of the δ13C values in the Late Ludlow. Deposition of organic-rich black shales is known to result in 12C-depleted ocean waters and thus lead to positive δ13Ccarb excursions. However, there is no sedimentological evidence of an overall increase in black shale deposition during the isotope excursions (Munnecke et al., 2003). In contrast, it seems that these phases are characterised by a decrease of organic-rich sediments on the shelf. This apparent problem can be resolved if an anoxic deep ocean is assumed, where the sequestration and burial of 12C took place (Cramer and Saltzman, 2005). Despite these differences in opinions and the multiplicity of hypotheses, all researchers probably agree that the Silurian carbon isotope excursions (directly or indirectly) result from climatic shifts. The causes for these climatic cycles, however, remain unclear. 3.4. Ordovician δ13Corg development Compared with δ13Ccarb relatively little is known about the Ordovician δ13Corg. Data for the isotopic composition of organic material have been published virtually exclusively from the Late Ordovician, especially from N. America (Melchin and Holmden, 2006a; Fanton and Holmden, 2007; Young et al., 2008, 2010; LaPorte et al., 2009), the UK (Underwood et al., 1997; Challands et al., 2009), Baltica (Young et al., 2010), and China (Wang et al., 1997; Fan et al., 2009; Zhang et al., 2010a). A compilation for the entire Ordovician is therefore currently not possible. Data from the Early Ordovician were recently presented by Zhang et al. (2010a) from the Yangtze Platform in South China, showing a large, positive (8‰) excursion in the middle Floian and more or less continuously decreasing values towards the Late Ordovician. Fanton and Holmden (2007) presented data from the Mohawkian of Iowa, correlated with the late Sandbian and early Katian by Bergström et al. (2009b). They identified six small (mostly below 1.5‰) positive δ13Ccarb excursions, and four of them correspond to small δ13Corg excursions, which, however, show slightly larger amplitudes compared to the δ13Ccarb data. The values presented by Challands et al. (2009) for the late Katian are measured from rocks that have experienced anchizone metamorphism, and it is not clear if they represent primary values. In China, the Guttenberg (GICE) δ13Ccarb excursion (early Katian) is accompanied by an excursion in δ13Corg but in sections from N. America no clear correlation is observed (Young et al., 2008). A strong excursion for δ13Corg in the Ordovician is observed in the Hirnantian, e.g., at Dob’s Linn in Scotland (Underwood et al., 1997), on Anticosti, Canada, and in Estonia (Young et al., 2010-this volume). The data from Anticosti and Estonia demonstrate that the large Hirnantian δ13C excursion (HICE; Fig. 3) can be recognised in both δ13Ccarb and δ13Corg but the latter shows significantly lower amplitudes (Young et al., 2010-this volume). The precise correlation of the Hirnantian δ13Ccarb and δ13Corg excursions, however, was recently questioned by Melchin and Holmden (2006a) and Delabroye and Vecoli (2010). Fan et al. (2009) presented δ13Corg values from the Wangjiawan Riverside section (close to the Hirnantian GSSP) showing a double-peaked excursion with a maximum amplitude of ca. 2‰ in the Hirnantian. The maximum values are measured from the Kuanyinchiao Bed, mainly a carbonate which unfortunately does not contain stratigraphically useful graptolites. It is, however, correlated indirectly with the lower N. persculptus biozone by graphic correlation (see review in Delabroye and Vecoli, 2010). According to Fan et al. (2009), the δ13Corg values start to increase in the latest Katian (pacificus Zone), show an initial maximum in the extraordinarius Zone and a second, larger maximum in the lower persculptus Zone. The values retreat to baseline values in the upper persculptus Zone. A similar double-peaked δ13Corg excursion was presented by LaPorte et al. (2009) from the Vinini Creek Section in Nevada, and from three different sections in Arctic Canada by Melchin and Holmden (2006a). Little is known about the environmental mechanisms leading to the development of Ordovician δ13Corg. The influence of local fluctuations in nutrient cycling and phytoplankton growth rates was recently highlighted by Young et al. (2008) and LaPorte et al. (2009). The large (8‰) Floian δ13Corg excursion is interpreted as a result of enhanced burial of organic matter (in an unknown area) (Zhang et al., 2010a). This resulted in a lowering of atmospheric CO2, and may have contributed to the Early/Mid Ordovician cooling proposed by Trotter et al. (2008) (Zhang et al., 2010a). Fanton and Holmden (2007) correlated the small Sandbian/Katian positive excursions of Iowa with sea-level highstands. During high-stands the nutrient-rich upwelling zone shifted landward, which stimulated local primary production in epeiric sea settings, and, consequently, enhanced burial of organic matter. Fan et al. (2009) correlated the two δ13Corg peaks in the Hirnantian with short-lived glacial intervals on Gondwana, i.e. the maximum sea-level low stands. Differences in amplitudes between sections from each palaeocontinent mainly resulted from the position of a particular section relative to the margin of the basin. These authors favour the weathering of carbonates as the cause of the (comparatively small) isotope excursion in South China. Similar conclusions were drawn by Melchin and Holmden (2006a). Based on paired analysis of δ13Ccarb and δ13Corg from Anticosti and Estonia, Young et al. (2010-this volume) argue that the Hirnantian atmospheric pCO2 was elevated. Expanding ice sheets reduced the fraction of continental silicates available for weathering. As a result, pCO2 began to rise, which eventually led to the deglaciation. 3.5. Silurian δ13Corg development To date, only few papers present Silurian δ13Corg data. Melchin and Holmden (2006b) discussed data from the Llandovery from Arctic Canada. The values scatter around −29.5‰ in the Rhuddanian and Aeronian, and around −30.5‰ in the early Telychian. In the latest Rhuddanian/early Aeronian a small (ca. 1‰), probably double-peaked excursion is observed, and a ca. 3‰-excursion is documented in the late Aeronian (Fig. 4). Gouldey et al. (2010-this volume) present data from the Llandovery of the Ikla core (Estonia), showing decreasing values in the Rhuddanian, more or less constant values in the Aeronian, and a shift back to heavier values in the Telychian. Loydell and Frýda (2007) have Author's personal copy A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 analysed a section in Wales, spanning the Llandovery–Wenlock boundary. The values increase from ca. −29‰ (Cyrtograptus insectus Zone) to ca. 27.5‰ (Monograptus riccartonensis Zone), with a small internal maximum at the Cyrtograptus murchisoni / Monograptus firmus zonal boundary. Increasing values in the early Sheinwoodian are also reported from Nashville (Tennessee) by Cramer and Saltzman (2007). The values, however, show a large scatter. A pronounced early Sheinwoodian positive excursion is demonstrated in the Twilight Creek Section (Bathurst Island, Arctic Canada), starting at −29.5‰ in the late Telychian, reaching peak values of −27.0‰ in the early Sheinwoodian, and finally decreasing rapidly to values of ca. −31.5‰ in the late Sheinwoodian (Noble et al., 2005). In the same region, the Homerian exhibits a strong, double-peaked excursion with an amplitude of more than 3‰ (Noble et al., 2005; Lenz et al., 2006; Fig. 4). Buggisch (2008) presents δ13Corg data from Ellesmere Island (Canada), with a positive excursion in the Llandovery, and a pronounced excursion in the Ludlow/Pridoli. Due to poor biostratigraphic control, however, it is yet not possible to precisely correlate these data. The greatest δ13Corg excursion is observed in the Late Ludlow of Gotland, with peak values of −22‰ (Fig. 4, unpublished data, A. Munnecke). A positive δ13Corg excursion correlating with the S/D boundary δ13Ccarb excursion is reported from the Czech Republic and from the Carnic Alps by Buggisch and Mann (2004). Although the Silurian database for δ13Corg is still relatively small, it seems that at least the major δ13Ccarb excursions (early Aeronian, early Sheinwoodian, Homerian, late Ludlow) are also evident in the organic carbon data (Fig. 4; see also Buggisch and Mann, 2004), however, the amplitudes of the excursions vary. The excursions are attributed to enhanced burial of organic matter (Cramer and Saltzman, 2007; Loydell and Frýda, 2007), or to the weathering of exposed carbonate platforms (Noble et al., 2005; Melchin and Holmden, 2006b). Cramer and Saltzman (2007) considered that changing pCO2 values in the atmosphere were responsible for the differences in the amplitudes between δ13Ccarb and δ13Corg in the early Sheinwoodian. 3.6. Ordovician δ18O development For the Ordovician, stable oxygen isotope data were mostly measured either from well-preserved brachiopod shells (Brenchley et al., 1994, 2003; Marshall et al., 1997; Veizer et al., 1999; Shields et al, 2003; Hints et al., 2010) or conodonts (Lehnert et al., 2007b; Trotter et al., 2008; Buggisch et al., 2010). Long (1993) and Armstrong et al. (2009a) presented and interpreted data for the Late Ordovician measured from rocks. However, as each limestone is inevitably diagenetically altered (otherwise it would still be a soft sediment), and because the δ18O values are prone to diagenetic alteration, values measured from rocks should not be used for palaeoenvironmental reconstructions unless the primary nature of the values can be verified (see Section 3.1.3). To date, only two studies cover, more or less, the entire Ordovician: Shields et al. (2003) presented data from 182 brachiopods from Laurentia, Baltica, South China and Australia, and Trotter et al. (2008) have analysed 179 conodonts from Gondwana (Australia) and Laurentia (Canada) (Fig. 3). Although these two studies differ somewhat in detail (which is probably at least partly the result of different palaeoenvironments and palaeolatitudes) and in the absolute values, they both show an overall trend towards heavier values during the Ordovician, with the maximum values observed in the Hirnantian (Fig. 3), which is interpreted as cooling, with a glacial maximum in the Hirnantian. Except for the HICE, there seems to be no correlation with the δ13C trend (Fig. 3). In a recent study Buggisch et al. (2010) presented δ18O data from conodonts from Minnesota and Kentucky displaying a short-lived ~1‰-excursion just above the prominent Deicke K-bentonite, clearly preceding the GICE. The authors argued that (a) the large volcanic eruptions led to global cooling, and (b) the GICE δ13C excursion 403 postdates this cooling event. Unfortunately, no data are presented from strata beneath the Deicke K-bentonite. A temperature change during the late Katian Boda Event, which is either assumed to be a time of comparatively warm climate preceding the Hirnantian glaciation (Fortey and Cocks, 2005), or a time of global cooling (Cherns and Wheeley, 2007), has not yet been confirmed by δ18O data. A clear correlation of δ18O and δ13C data is, however, documented for the HICE, for example, an approximate 4‰ excursion in δ18O parallel to a δ13C excursion is reported from brachiopod shells in the Ruhnu core in Estonia (Brenchley et al., 2003) and from the Stirnas18 core in Latvia (Hints et al., 2010). The resolution, however, of reliable δ18O data is much lower than that for δ13C. Nevertheless, most authors agree that the Hirnantian δ18O excursion reflects the glacial maximum. Two problems remain: (a) the 4‰ shift measured from brachiopods requires a sea-level fall of N100 m and a drop of 10 °C in tropical surface water temperatures (Brenchley et al., 1994, p. 297), which seems unrealistically high, and (b) the proposed development of the Gondwana ice sheets does not precisely match the track of the δ18O values. Peak values are reported from the Laframboise Member on Anticosti Island, Canada, deposited during a major low-order sea level lowstand. The underlying regressive upper Lousy Cove Member, however, does not show elevated δ18O values (Brenchley et al., 1994) although it is interpreted as being deposited during the onset of the second, major glacial pulse (Desrochers et al., 2010-this volume). 3.7. Silurian δ18O development Compared to the Ordovician, the Silurian δ18O curve shows a higher resolution, which is based mainly on brachiopod data presented by Azmy et al. (1998) for the Llandovery from different regions, by Heath et al. (1998) for the Llandovery and early Wenlock of Estonia, and by Samtleben et al. (1996, 2000, 2001) for the Wenlock and Ludlow of Gotland. Additional brachiopod data were published by Wenzel and Joachimski (1996) and Munnecke et al. (2003) from the Silurian of Gotland, and by Munnecke and Männik (2009) from the Llandovery (Telychian) of Anticosti Island (Canada). Silurian δ18O data from phosphatic organisms was presented by Wenzel et al. (2000) from the Silurian of Gotland, by Lehnert et al. (2010-this volume) from the Telychian and Sheinwoodian of Estonia, by Lehnert et al. (2007c) from the Ludlow of the Czech Republic, and by Zigaite et al. (2010) from the Pridoli of Lithuania. The brachiopod data show a decreasing trend from the early Rhuddanian (−3‰) to the late Telychian (−5.5‰), interrupted by two small positive excursions in the early and late Aeronian (Azmy et al., 1998). In the Wenlock and Ludlow, the δ18O curve closely parallels the δ13Ccarb curve (Fig. 4). Pronounced positive excursions are reported in the early Sheinwoodian, the Homerian, and in the late Ludlow. Additionally the δ18O data from phosphatic organisms show a clear correlation with the δ13Ccarb data (Wenzel et al., 2000) (Fig. 4), although the recent study by Lehnert et al. (2010-this volume) indicated that the increase in δ18O values during the Ireviken Event (close to the Llandovery/Wenlock boundary) apparently occurs somewhat later (in the lower part of the Lower K. ranuliformis Zone) than the brachiopod values which start to increase in the Upper P. procerus Zone. This offset, however, is based on the isotope analysis of only two conodonts and therefore awaits confirmation by further studies. Also the decrease of the δ18O values seems to be later compared to brachiopods (Lower versus Upper K. walliseri Zone) (Lehnert et al., 2010-this volume). The environmental interpretation of the Silurian δ18O data is currently a matter of debate (e.g., Bickert et al., 1997; Heath et al., 1998; Brand et al., 2006; Loydell, 2007, 2008; Cramer and Munnecke, 2008). By comparing δ18O values from brachiopods of different contemporaneous facies from the Silurian of Gotland Samtleben et al. (1996, 2000) have shown that oxygen isotope values are more strongly affected by local environmental conditions than carbon Author's personal copy 404 A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 isotopes. Interpretation of the brachiopod oxygen isotope record of the Silurian of Gotland in terms of only palaeotemperature variations and assuming a marine isotopic composition close to the modern global mean (δ18Oseawater = 0‰), require temperatures ranging from 33 to 46 °C, which is unrealistically high (Bickert et al., 1997). These authors therefore attribute the fluctuations between low and high values to mainly salinity changes resulting from oscillations between more humid and arid climatic conditions in low latitudes, respectively. It may be possible, however, that the Silurian sea-surface temperature was considerably higher than modern temperatures, as suggested by Came et al. (2007). Those authors used the ‘clumped isotope thermometer’ which examines the ordering of 13C and 18O into bonds with each other in the calcite (or aragonite) lattice. Based on investigations of Telychian brachiopods from Anticosti Island Came et al. (2007) reconstructed SSTs (sea surface temperatures) around 35 °C, and attribute these high values to the high CO2 values inferred for this time slice (see Section 4.2). A correlation between δ18O and δ13Ccarb values is observed for both the major Silurian excursions and the Hirnantian excursion (see above; Munnecke et al., 2003), and, consequently, most authors attribute the Silurian δ18O excursions to intervals of lower sea-surface temperatures and expanded ice sheets on Gondwana (e.g., Azmy et al., 1998; Kaljo et al., 2003; Brand et al., 2006; Eriksson and Calner, 2008; Lehnert et al., 2010-this volume; Zigaite et al., 2010). An opposite conclusion, however, was suggested by Cramer and Saltzman (2007), who attributed the excursions to intervals of high sea level. These authors associated the positive δ18O excursions with increased carbonate production and burial in epeiric seas, decreasing the [CO2− 3 ]. Heath et al. (1998) measured brachiopods from the Ruhnu core in Estonia, and highlighted the fact that their Llandovery δ18O data did not show evidence for major glacio-eustatic sea-level fluctuations (Grahn and Caputo, 1992). In general there seems to be a clear mismatch between δ18O data and sea-level reconstructions based on sequence stratigraphy. For example, Calner et al. (2004) argued as follows: “In contrast to the siliciclastic depositional system, carbonate platforms produce and deposit most of their sediments during highstand situations. This is primarily due to the increased areal extent of platform flooding and the associated increase in space available for skeletal carbonate production. This is well illustrated on Gotland. Here, the expansion and thickening of reef complexes across distal platform marls imply that reef barriers formed during relative highstand of sea-level. Such substantial progradation of reef complexes onto argillaceous limestone and marl deposited in deeper, distal settings can be seen e.g. in the Lower Wenlock north of Visby and in the Late Wenlock of the Klintehamn area.” This means, in terms of sequence stratigraphy, that the Upper Visby and Högklint formations on Gotland represent highstand deposits (Calner et al., 2004), whereas the δ18O data indicates glacial conditions at least for the upper part of the Upper Visby Formation and the Högklint Formation (Lehnert et al., 2010-this volume, Section 5.1). A sea-level drop is recorded for the Late Ludlow excursion, both on Gotland (Eriksson and Calner, 2008) and in the Holy Cross Mountains in Poland (Kozłowski and Munnecke, in press), which are positioned on the opposing sides of the same foreland basin. Although the sequence stratigraphic interpretations in both studies differ somewhat, in both papers the maximum isotope values, however, are reported from sediments indicating rising sea level (middle and upper Eke Formation on Gotland, and upper Bełcz Member in Poland; in the latter only δ13C values are available). By contrast, in the Vidukle core from Lithuania, which also belongs to the same foreland basin, the Late Ludlow δ13C excursion is reported from strata which are, according to data from benthic assemblage zones, interpreted as encompassing an interval of falling sea level (Martma et al., 2005). In summary, the connection between sea level and δ18O data is far from being understood; more information is required. The greatest Silurian δ18O fluctuations are reported from post-Llandovery strata (Fig. 4), whereas glacial deposits are of Llandovery to earliest Wenlock age (Hambrey, 1985; Grahn and Caputo, 1992; Díaz-Martínez and Grahn, 2007; Loydell, 2007). The Llandovery with its proven glacial deposits and therefore also proven glacially-induced sea-level fluctuations should show the greatest δ18O fluctuations but it does not. In this respect it is significant that the late Ludlow positive δ18O excursion has the same magnitude as the Hirnantian excursion (compare Figs. 3 and 4), but up to now no glacial deposits have been identified in this time slice. 3.8. Ordovician 87 Sr/86Sr development The Ordovician is characterised by a large drop in 87Sr/86Sr values, from ca. 0.7090 to 0.7079 (Fig. 3; Yang and Wang, 1994; Qing et al., 1998; Shields et al., 2003), and this general trend is explained in terms of the reduction in rates of tectonic uplift generated by the waning of PanAfrican mountain-building (Qing et al., 1998; Shields et al., 2003). However, a major drop in seawater 87Sr/86Sr is observed at the Darriwilian–Sandbian transition (Qing et al., 1998; Veizer et al., 1999; Shields and Veizer, 2004). This drop is one of the most rapid changes in 87 Sr/86Sr recorded for the entire Phanerozoic, and because of the lack of correlation between the strontium ratio and other parameters (Fig. 3) a straight-forward explanation is difficult. Servais et al. (2010) highlighted the fact that this major drop in 87Sr/86Sr coincides with a major turnover in the composition of reefs from microbial-dominated reefs in the Early and Middle Ordovician to metazoan-dominated reefs in the Late Ordovician, and speculated that this biological turnover was at least in part the result of a lowering of the carbonate saturation state in the ocean. Qing et al. (1998) and Shields et al. (2003) assumed that the large drop in 87 Sr/86Sr values was the result of lower continental erosion rates and increased submarine hydrothermal exchange rates. These authors assumed that the large transgression postulated for the Darriwilian– Sandbian transition (Haq and Schutter, 2008), drowned the extensive cratonic areas which consequently reduced the source of radiogenic strontium. The transgression itself might have been the result of increased sea-floor spreading, which increased the input of nonradiogenic strontium. The hypothesis proposed by Shields et al. (2003) has been recently confirmed by numerical modelling (Young et al., 2009). 3.9. Silurian 87 Sr/86Sr development The most detailed Silurian 87Sr/86Sr curves based on brachiopod data were provided by Ruppel et al. (1996) and Azmy et al. (1999). Additional data were published by Qing et al. (1998) and Veizer et al. (1999). Data from rock samples are provided by Denison et al. (1997) and Gouldey et al. (2010-this volume). Although the slopes of the curves presented by Ruppel et al. (1996) and Azmy et al. (1999) are not consistent, due to the fact that different timescales have been used, the general trend is more or less the same in both curves. The values increase from ca. 0.7080 in the earliest Llandovery to 0.7087 in the Pridoli. There seems to be two intervals when the shifts were more rapid (in the Sheinwoodian and Gorstian; Fig. 4). As the radiometric age control on the Silurian lacks precision, it is not possible to prove that these features represent ‘real’ events. Ruppel et al. (1996) depicted several high-frequency cycles (≤1 conodont zone) superimposed on the long-term rise of 87Sr/86Sr values, which contradicts the wide-spread opinion that due to the long residence time of strontium in the ocean (Section 3.1.4), short-term changes should be impossible to record. However, because many of these short-term changes are supported by more than one data point and sometimes also by data from different localities they are apparently real, which, according to Ruppel et al. (1996) “requires rethinking of models of strontium isotope flux in marine basins.” The general increase in Silurian 87Sr/86Sr values is interpreted by Azmy et al. (1999) as a result of an increased alluvial flux of radiogenic Author's personal copy A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 strontium due to progressive warming (increased weathering of sialic rocks), and as result of increased continental input in response to the early phases of the Acadian Orogeny (Qing et al., 1998). 4. Evolution of the atmosphere In this section, a brief summary of our current knowledge of the oxygen and carbon dioxide composition of the atmosphere is presented. In contrast to the previous section, which focused on the presentation of measured data, there is as yet no geochemical proxy, which can be directly and unequivocally related to atmospheric gas composition; and thus the concentrations reported must be critically assessed. We are not aware of a single paper published dealing exclusively with either the Ordovician or the Silurian oxygen or carbon dioxide content of the atmosphere. Instead, in most papers the entire Phanerozoic is considered, and therefore detailed curves for atmospheric oxygen during the Ordovician and Silurian are still lacking (e.g., Berner, 1999, 2001, 2006a,b; Berner et al., 2000, 2003, 2007; Royer et al., 2001, 2004; Rothman, 2002; Royer, 2006; Algeo and Ingall, 2007). 4.1. Oxygen content of the atmosphere Berner (2006) Bergman et al. (2004) 4.2. Carbon dioxide content of the atmosphere Carbon dioxide is a greenhouse gas occurring at an average concentration of about 383 parts per million by volume in the modern GEOCARB III (Berner, 2001) incl. estimate of errors modern value Algeo and Ingall (2007) tary rocks is required. Especially for Palaeozoic rocks, however, these parameters are often not known and have to be approximated, or even estimated. Pelagic deep-sea sediments, for example, have been completely subducted, and the preserved rock record has usually experienced strong alteration, together with erosion and weathering. Generally, the oxygen content of the Ordovician and Silurian atmosphere is reconstructed with values well below the modern level (which is about 21%). An exception is the Silurian atmosphere, reconstructed by Berner (2006b) which shows an increase from about 19% to approximately 23% (Fig. 5). The oxygen content of the reconstructed Ordovician and Silurian atmosphere, however, differs significantly in the literature, with values between about 6% (Bergman et al., 2004) and N20% (Berner, 1999). Based on the Corg:P ratio Algeo and Ingall (2007) inferred generally low values in the Ordovician and Silurian (O2 content of ~ 16%) but with a rise to nearly ‘modern’ values peaking in the Ordovician (~18%) (Fig. 5). The low O2 content reconstructed for the Early and Middle Palaeozoic should have encouraged poorly ventilated oceans, and this is in good accordance with the widespread occurrence of black shales (Berry and Wilde, 1978). Low levels of atmospheric oxygen would result in a shallow boundary between oxic surface water and anoxic or dysoxic deeper water masses. Consequently, the bacterial remineralisation of organic matter in the water column is reduced, prompting the increased deposition and burial of 12C-enriched organic matter, and, consequently, driving strong δ13C fractionation between surface water and deeper water masses (Bickert et al., 1997). smoothed record of proxy data (Royer et al. 2006) Cambr. Ordovician Silur. modern value Devonian Carboniferous Permian The atmospheric oxygen content depends on a variety of different biological and geological factors, e.g., rate of photosynthesis, the marine phosphorous cycle, weathering of organic matter and pyrite in sedimentary rocks, reduction and removal of DIC (dissolved inorganic carbon) from sea water, and geochemical reactions of mid-ocean-ridge basalts (for a detailed review see, e.g., Berner, 2001, 2006b; Algeo and Ingall, 2007). For reliable reconstructions of atmospheric gases analyses of the amount and composition (calcium carbonate, calcium sulphate, pyrite, organic matter, phosphorous) of deposited sedimen- 405 10 20 atmospheric O2 (%) 30 0 2000 4000 6000 atmospheric CO2 (ppm) Fig. 5. Atmospheric oxygen and carbon dioxide in the Palaeozoic. 8000 Author's personal copy 406 A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 atmosphere. The CO2 content of the atmosphere is in equilibrium with the ocean waters; however, the amount of carbon dissolved in the 2− modern oceans (as CO2, H2CO3, HCO− 3 , or CO3 ) is about 50 times larger than in the atmosphere. The CO2 content of the ocean/atmosphere system depends on a variety of factors, e.g., rate of photosynthesis, weathering of silicates and carbonates, volcanism, nutrient cycling, burial/oxidation of organic matter, and deposition of carbonate rocks (Goddéris et al., 2001; Royer et al., 2001; Berner, 2003). At present only four widely used palaeo-pCO2 proxies for preQuaternary rocks exist (see extensive review in Royer et al., 2001): δ13C of pedogenic carbonates, paired δ13Ccarb and δ13Corg analyses of marine sediment, density of stomata on the leaves of C3 plants, and δ11B of planktic foraminiferans. Among these, only paired analyses of δ13Ccarb and δ13Corg can be used to reconstruct the palaeo-pCO2 of the Early Palaeozoic because vascular plants and planktic foraminiferans had not yet evolved (e.g., Hayes et al., 1999; Young et al., 2010-this volume). The validity of this method, however, is currently debated (see review in Bickert, 2006), and “pre-Devonian estimates should be treated cautiously” (Royer et al., 2001). Alternatively, CO2 levels can be modelled. For this approach, factors such as global oceanic circulation, mountain building, solar radiation, submarine and continental weathering, and vegetation, have to be estimated (Berner, 2001, 2003, 2006a; Berner and Kothavala, 2001). Nevertheless, where available, modelled data show a good correlation with that reconstructed from proxy data (Fig. 5; Royer, 2006). There is, however, a debate on ancient CO2 levels, especially for pre-Palaeozoic rocks associated with Precambrian world of the early faint Sun (Rosing et al., 2010). The CO2 levels reconstructed for the Palaeozoic by Berner (2001) and Royer (2006) are summarised in Fig. 5. Despite the large number of possible errors the results show very high values above 2000 ppm, probably around 4000 ppm, which is more than ten times higher compared to those of the modern atmosphere. A complicating factor for any atmospheric reconstruction is that carbonate saturation and pH of sea water in the Palaeozoic was probably totally different from those of the modern ocean (Ridgwell, 2005). The Ordovician and Silurian are attributed to a time of ‘calcite seas’ which means the marine precipitates (ooids, marine cements) are mostly low-magnesium calcite instead of aragonite and highmagnesium calcite (Sandberg, 1983). The calcite-dominated mineralogy probably resulted from low Mg/Ca ratios in the ocean water caused by high sea-floor spreading, which is in accordance to the overall high sea level during this time (Stanley and Hardie, 1999). The high CO2 content of the atmosphere was probably not the reason for the ‘calcite seas’ (Stanley and Hardie, 1999). Today, even the comparatively small increase of anthropogenic CO2 in the atmosphere by burning fossil fuels has a strong influence on sea-water pH and thus on the precipitation of calcium carbonate, which eventually will result in a destruction of modern shallow- and deep-water coral reef ecosystems within the next few hundred years (Kleypas et al., 1999; Guinotte et al., 2006). If the modern atmosphere had CO2 levels even approaching those reconstructed for the Early Palaeozoic all carbonates would dissolve. Probably, however, carbonate saturation in the Palaeozoic ocean was much higher because of a lack of the ‘pelagic carbonate sink, which characterises the modern ocean, i.e. pelagic oozes composed of planktic foraminiferans and coccoliths were absent with pelagic calcareous plankton playing only a very minor role in the Palaeozoic (Ridgwell, 2005; Munnecke and Servais, 2008). Because of the numerous sources and sinks, the greenhouse gas CO2 has the potential to influence the global climate over a wide range of timescales (tens of years to millions of years; Royer et al., 2004). And because the Phanerozoic δ18O trend, derived from calcitic and aragonitic shells, does not correlate with the CO2 data, Veizer et al. (2000) questioned the importance of CO2 for global climate change (or, alternatively, the validity of the CO2 reconstructions). This apparent flaw, however, can be resolved if changes in the ocean pH are integrated with δ18O temperature reconstructions (Royer et al., 2004). Somewhat enigmatic, though, is the occurrence of the large Late Ordovician (to early Silurian) glaciation during a time characterised by high CO2 contents in the atmosphere (Figs. 1 and 5) (Gibbs et al., 1997; Kump et al., 1999; Poussart et al., 1999). Although some evidence suggests a longer interval of glaciation in the Late Ordovician and Early Silurian, the dominant glacial phase is consistently placed within the Hirnantian (e.g., Pope and Steffen, 2003; Saltzman and Young, 2005; Díaz-Martínez and Grahn, 2007; Vandenbroucke et al., 2009, 2010; Desrochers et al., 2010-this volume; Loi et al., 2010-this volume; Videt et al., 2010-this volume). For present-day conditions, the threshold for the initiation of widespread glaciations on continental areas is about 500 ppm (Royer, 2006). Royer (2006, p. 5669), however, calculated that this threshold was probably much higher (~ 3000 ppm) in the Early Palaeozoic due to an approximately 4% lower solar luminosity at that time. Herrmann et al. (2004a,b) modelled the Late Ordovician climate and concluded that even a comparatively low CO2 content of 8xPAL would not be sufficient to initiate the growth of ice sheets, other factors must have been involved. Alternatively, it might be possible that a strong but shortlived drop in CO2 is simply not detected yet because of poor proxy coverage and/or the low temporal resolution (10 m.y.) of the GEOCARB model of Berner (2001). Based on coupled δ13Ccarb and δ13Corg analyses, however, Young et al. (2010-this volume) argued for the opposite — a CO2 increase at least close to the glacial maximum. 5. Climatic implications 5.1. Ordovician climate Critical to our understanding of Ordovician and Silurian ecosystems and environments is the development of an accurate model for climate change and climatic conditions during the two periods. Some authors have linked climate change directly to evolution (see Harper, 2009). There are a number of challenges not least that surface water temperatures across the globe have a range of about 30 °C and much of the planet is subjected to significant seasonal variations in temperature. Traditionally a range of biotic and sedimentological indicators, such as latitudinally-controlled faunal and floral provinces together with climatically-sensitive sediments, for example evaporites, calcretes, tillites, coals, kaolins and bauxites, have been widely used (e.g. Boucot, 2009). More recently a number of geochemical proxies (see this paper) based on isotope data are providing a much more accurate picture of temperature change whereas sea-level curves, linked to eustasy, are providing a powerful means of linking transgressions and regressions to the waxing and waning of major ice sheets. Conventional climatic curves for the Ordovician (e.g., Frakes et al., 1992) indicate generally warm conditions throughout much of the Ordovician but with an icehouse interval through the Ordovician– Silurian boundary. Frakes et al. (op. cit.) show a cooling trend during the period with the warmer conditions of the Early Ordovician interrupted by a short-lived cool interval. Some authors have suggested that the widespread Early Ordovician ‘Ceratopyge Limestone regression’ represented the evidence of the first Ordovician ice age. It is only relatively recently that geochemical proxies are confirming an overall decrease in Ordovician temperatures (Trotter et al., 2008). Cooler waters may have provided more hospitable conditions for marine life to diversify (Trotter et al., 2008) or increased calcium carbonate saturation probably aided the precipitation of the heavier skeletons of the Palaeozoic benthos (Pruss et al., 2010). 5.2. Silurian climate The Silurian has traditionally been regarded a greenhouse period with little or no ice at the poles and weak latitudinal climate gradients. As is clearly indicated herein, this view has recently been abandoned Author's personal copy A. Munnecke et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 389–413 due to the immense amount of geochemical data that imply unstable oceanic-atmospheric conditions. An Early Silurian glacial record is also fairly well established, and it is clear that at least the Llandovery belongs within an icehouse period that started in the early Katian–the “Early Palaeozoic Icehouse” of Page et al. (2007). However, up until now there is no consensus on Silurian climate development, especially for the post-Llandovery. An early attempt to introduce climatic cycles into Silurian stratigraphy, was promoted by the oceanic model developed by Jeppsson (1990) (see also Aldridge et al., 1993; Jeppsson et al., 1995). It was based on observed temporal changes in lithology and conodont faunas in the carbonate platform rocks of Gotland. Jeppsson (1990) identified alternating Primo and Secundo episodes. Secundo episodes are assumed to be characterised by a more arid climate at low latitudes favouring the expansion of reefs and associated sediments throughout the tropics in shallow-water settings whereas the more humid climate during Primo episodes resulted in the increased transport of terrigenous material into the sea, favouring argillaceous limestone deposition. This often-cited, seminal model, which was – in more or less modified versions – invoked also to explain the Silurian stable carbon and oxygen isotope signal (Samtleben et al., 1996; Bickert et al., 1997; Cramer and Saltzman, 2005; Cramer et al., 2006c), has received much criticism (e.g., Loydell, 1998, 2001; Kaljo et al., 2003; Johnson, 2006), but it has without any doubt greatly stimulated scientific discussion, and has focussed attention, quite rightly, on other factors besides only sealevel changes. As discussed in Section 2, sea level has demonstrably changed significantly during the Silurian, but the interactions of sealevel change and stable isotope geochemistry are still poorly understood. Especially the fact that the isotope excursions are connected with extinction events is important in this respect because most of the extinctions occur at the onset of the excursions, or even slightly earlier (see review in Munnecke et al., 2003; Jeppsson and Calner, 2003). It is self-evident that these extinctions cannot be the immediate result of the environmental changes producing the isotope excursion — instead both extinctions and stable isotopic events are most likely the result of the same yet still enigmatic processes that occurred shortly before the isotope excursions and extinctions (Cramer and Munnecke, 2008). 6. Conclusion Studies on the Ordovician and Silurian systems have, during the last two decades, advanced dramatically through the acquisition of new data from many new and previously poorly-studied sections throughout the world. More importantly, however, a range of new geochemical techniques and sea-level proxies applied to existing and new sections are rapidly developing our knowledge of ocean chemistry and sea level change. The results of IGCP 410 and 503 have sharpened our focus on the significance of climatic and environmental change for the evolution of Early Palaeozoic biotas at a critical time in Earth history. These projects have identified the need for the continued careful sampling of sections against a wellconstrained biostratigraphy across a wide range of palaeolatitudes. They have also identified the importance of developing and applying new proxies, e.g., Osmium isotopes, 13C–18O ‘clumped’ isotopes, or Ca isotopes (Farkaš et al., 2007; Finlay et al., 2010; Tripati et al., in press), and innovative techniques while building global databases from precisely assembled local and regional datasets. Fluctuations in climate and environment in deep time provide a dynamic to ancient marine ecosystems, moreover the organisms themselves present important feedback processes. Essential to our understanding of Early Palaeozoic earth systems is the continued search for links and relationships between environments, ecosystems and evolution. This can only be achieved by a multidisciplinary approach advocated by the recent IGCP 410 and 503 projects. 407 Acknowledgments This paper summarises many of the activities of IGCP 503 (“Ordovician Palaeogeography and Palaeoclimate”) that extended from 2004 to 2009. We acknowledge all the participants of the project who discussed many aspects of the present manuscript with us, and in particular our three co-leaders Alan Owen (Glasgow, Scotland), Li Jun (Nanjing, China) and Peter Sheehan (Milwaukee, Wisconsin, USA). AM acknowledges funding from the German Research Foundation (DFG Mu 2352/1), and DATH financial support from the Danish Council for Independent Research (FNU). MC acknowledges financial support from the Swedish Research Council (VR) and from the Crafoord Foundation. TS is grateful to the Alexander von Humboldt-Foundation for a Nachkontakt-Programm at the University Erlangen-Nuremberg (Germany). 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