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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
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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
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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;
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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-
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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
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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
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GLOBAL SERIES
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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
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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
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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.
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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).
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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
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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
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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
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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
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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
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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
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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
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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
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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). The authors are grateful
to Dimitri Kaljo (Tallinn, Estonia) and an anonymous referee, and to
Christian Samtleben (Kiel, Germany) and Torsten Bickert (Bremen,
Germany) for many helpful comments on an earlier version of this
manuscript. This paper is a contribution to IGCP 503 (“Ordovician
Palaeogeography and Palaeoclimate”) and to the CNRS-INSU project
SYSTER (“Ordovician Climate”).
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