EPSC 425: Sediments to Sequences
Winter Semester, 2014 Contents
1 Introduction to Stratigraphy 1 1.1 Sedimentary Geology ...... 1 1.2 What is Stratigraphy? ...... 1 1.3 Strata as Temporal Archives ...... 2 1.4 Lithostratigraphy ...... 5 1.5 Depositional Environments ...... 8 1.6 Lithofacies and Associations ...... 14 1.7 Correlation ...... 16
2 Geophysical Data 18 2.1 Well Logs ...... 18 2.2 Seismic Exploration ...... 24
3 Sequence Stratigraphy 27 3.1 Introduction ...... 27 3.2 Fundamental concepts ...... 29 3.3 Shoreline trajectories and sediment stacking patterns ...... 30 3.4 Base-level changes and the development of systems tracts ...... 32 3.5 Sequence boundaries ...... 36 3.6 Stratal terminations ...... 38 3.7 Types of Sequences ...... 38 3.8 Sequence Stratigraphy of Carbonate Platforms ...... 42
4 Learning to Tell Geological Time 44 4.1 Introduction ...... 44 4.2 The Geological Time Scale ...... 45 4.3 Telling Time in the Stratigraphic Record ...... 47
5 Radiometric Techniques 51 5.1 Introduction ...... 51 5.2 U-Th-Pb System ...... 52 5.3 K-Ar System ...... 57 5.4 Re-Os System ...... 59 5.5 Radiocarbon Dating ...... 59
6 Magnetic Stratigraphy 62
7 Biochronology 65 7.1 Biostratigraphy ...... 65 7.2 Biochronology ...... 65
8 Sedimentary Cycles 67 8.1 Introduction ...... 67 8.2 Stratigraphic cycles ...... 67 8.3 Climate-independent cyclic sedimentary patterns ...... 68 8.4 Sedimentary parameters linked to climate change ...... 68 8.5 Astronomical Forcing ...... 71
9 Chemical Stratigraphy 76 9.1 Introduction ...... 76 9.2 Underlying Principles ...... 76 9.3 Development and Application of the Method ...... 78 9.4 Stable Isotope Systems ...... 79 9.5 Radiogenic isotope systems ...... 89 Chapter 1
Introduction to Stratigraphy
1.1 Sedimentary Geology
Sedimentary Geology encompasses any ﬁeld of science dealing with sediments or sedimen- tary rocks – that is rocks formed by Earth surface processes that begin with weathering on land, followed by erosion, and ﬁnally the physical settling of grains or chemical precip- itation of minerals from air, water, or ice. Sediments (unlithiﬁed) and sedimentary rocks (lithiﬁed) are archives of information about environmental, tectonic, and biological condi- tions that prevailed at the time they were laid down; even their very existence is revealing to the interested geologist. Sediments are largely derived from older sedimentary, igneous and metamorphic rocks, and themselves may be metamorphosed or melted to form other rock types. Hence, they represent one arc in the continuous and long cycle of the shaping and reshaping of Earth’s landscape, as envisioned by James Hutton.
1.2 What is Stratigraphy?
Stratigraphy is the study of geological strata. It may not be immediately obvious to a non- stratigrapher why it is important and it may come across as rather boring to many. But stratigraphy is a process for applying order to the world of stratiﬁed rocks. Stratigraphy involves documenting sedimentary successions, interpreting them, correlating sedimentary units across basins, and where possible, dating them, either directly or indirectly. Stratig- raphy is the temporal and environmental framework within which much of Earth’s history has been studied, debated, and deduced. Stratigraphy is also indispensable in petroleum geology, for all petroleum source rocks are sedimentary, as is virtually every reservoir. Similarly, stratigraphy is important both in exploration of stratabound ores and in hy- drogeology, because most major aquifers are in porous sediments, and their lithology and spatial distribution control the extent of the reservoir and ﬂuid ﬂow within it.
Stratigraphy encapsulates both sedimentary and volcanic rocks. However, it is funda- mentally intertwined more closely with sedimentary geology than igneous petrology, and volcanic rocks are studied only insofar as they may comprise a component of a stratigraphic succession. That said, volcanic rocks can be extremely useful in stratigraphy as markers for correlation and for the potential for radiometric dating.
Geology has only been a scientiﬁc discipline since the early 19th century. However, curi-
1 Chapter 1: Introduction to Stratigraphy 2 ous intellectuals contemplated the deeper meaning of sedimentary rocks long before that. Aristotle, Leonardo Da Vinci, and Nicolas Steno all famously endeavoured to understand sedimentary rocks, with varying degrees of success. Steno (1638–1686), a Danish anatomist, was certainly the most successful geologists of these early dilettantes; for his eﬀorts, he is often regarded as the proto-father of geology. This early curiosity about stratigraphy was certainly not a coincidence: sediments and sedimentary rocks cover much of Earth’s sur- face. They also contain fossils, which are captivating to many.
Steno formulated three basic principles (Steno’s Laws) about sedimentary rocks that con- tinue to govern stratigraphic approaches today:
• Law of Original Horizontality and Lateral Continuity
• Law of Cross-cutting Relationships
We can add to this list another basic principle: the law of inclusion, which states that if one rock is incorporated into rock (say a pebble, or even a fossil in most cases), it the former must be older than the latter. These laws are all fairly obvious to any geologist, but they are powerful nonetheless, because they enable us to order geological events chronologically. Time is of course indispensable into trying to reconstruct Earth history, and most often, we cannot rely upon radiometric dating techniques to provide ages for rocks.
1.3 Strata as Temporal Archives
The stratigraphic record is the closest thing we have to an almanac of past geological events on the surface of the earth. Based on Steno’s Laws, we can assume that strata that have not been horribly tectonized or otherwise perturbed should provide some sort of temporal record, with the lowest rocks being the oldest and the youngest rocks of partic- ular succession being the youngest. As such, they are an Archive of Earth history. This is why Earth historians turn to the sedimentary record to reconstruct and interpret past geological events, as varied as the origin of animals to meteorite impact events. However, the stratigraphic record is deeply ﬂawed. For one, sedimentary basins do not cover the globe and many sedimentary basins that have existed over the history of the Earth have been partially or wholly destroyed. Another problem is that the record is not a truly continuous record of past events–that is, it is no tape recorder. Imagine a shallow marine sedimentary environment. Many of the sediments deposited and ultimately preserved are most likely deposited in a series of geologically instantaneous or short-lived events, such as storms, earthquakes, and other unusual conditions. During the intervening time, little or no deposition may take place, or what sediments are deposited may be wiped out by a subsequent event. Hence, the shallow marine sedimentary record only captures snapshots of conditions in any one basin. This irregularity and discontinuity in sedimentation makes it diﬃcult to transport stratigraphic height or depth into time.
Because most of the sedimentary rocks found on the continents are from the continental shelf environment, stratigraphers face a severe uphill challenge to reconstruct past geologi- cal events from the extant rock record. They can improve their odds by studying an entire Chapter 1: Introduction to Stratigraphy 3 sedimentary basin–or at least, more than a single drill core a stratigraphic section. By in- tegrating data from across a basin, on can ﬁll in more details of past geological events and at the same time, better interpret the depositional history of a basin. Hence, such basin- wide studies are important both in Earth history and in exploration for petroleum or other strata-bound mineral resources, whose distribution is controlled by the temporal-spatial evolution of the basin ﬁll and its subsequent burial and diagenesis. As will be discussed throughout the remainder of this course, many diﬀerent tools are available to stratigraphers to reconstruct basin-wide stratigraphic geometries and basin evolution. Biostratigraphy is a tried and true method of establishing time equivalence, hence correlation of sedimen- tary strata. However, a variety of other sedimentologicaly, geochemical, and geophysical techniques are also applied. Indeed, the state of stratigraphy is such that many geologists become specialists in just one or a small handful of these techniques.
The issues of incompleteness and disturbance to the sedimentary record are greatly reduced in the study of deep sea sediment cores. A consortium of scientists has been drilling the seaﬂoor for over four decades now. It started as the Deep Sea Drilling Project (DSDP) in 1968 by the United States, was internationalized as the Ocean Drilling Program (ODP) in 1985, and became the Integrated Ocean Drilling Program (IODP) in 2004. The objective of the program is to sample sub-seaﬂoor environments, which are of course overwhelmingly sedimentary (although they do drill into both oceanic and continental crust as well). This program has produced countless drill cores from all of the ocean basins and is far and away the greatest source of information on Earth history over the past 200 million years.
1.3.1 Stratigraphic subdivisions Stratigraphers naturally want to subdivide the sedimentary record into discrete units for the purpose of correlation, evaluating spatial variability of time equivalent units and fossils, and many other purposes. If all sedimentary rocks could be easily dated, this would be a rather trivial aﬀair. However, it turns out that it is often extremely diﬃcult to date rocks directly, and because the stratigraphic record is so often disturbed by erosion, tectonism, and other indignities, discerning ages of rocks is often very diﬃcult. Even applying relative ages often involves correlation. Consequently, the subdivision of the stratigraphic record is not straightforward.
In the early days of geology, in the 1800’s, two diﬀerent approaches emerged to apply order to the stratigraphic record: lithostratigraphy and biostratigraphy. Lithostratigraphy, de- scribed in more detail below, entailed the physical description of the rocks and correlation between separated sections based on lithological divisions and contrasts. This gave rise to a hierarchical division of the actual stratigraphic record, as shown in Table (1.1), and the practice of correlation based on these subdivisions. At the same time, early palaeon- tologists, working from the principal of faunal succession formulated by William Smith, began to break down the stratigraphic record basic on fossil assemblages. Biostratigra- phy emerged as a powerful tool in establishing relative ages, and hence in correlating, particularly in rocks of homogenous lithology. These sometimes competing, but also com- plementary approaches are deeply intertwined; for example, sedimentary structures and constituents are commonly biogenic and sedimentary environments commonly control the type and preservation of fossils. Lithostratigraphy and biostratigraphy combined were the basis for the early construction of the geological time scale, hence chronostratigraphic sub- Chapter 1: Introduction to Stratigraphy 4 division of the sedimentary record. It was not until radiometric techniques were developed in the middle of the 20th century that actual dates came to be applied to the geological time scale. Since this time, the geological time scale has been continually revised as improved dating techniques are developed and more key rock boundaries are dated and correlated.
Despite the advent of radiometric dating, establishing the ages of strata remains a chal- lenge. Fortunately, other stratigraphic techniques have been developed that have provided powerful new means of correlation. Both magnetostratigraphy and chemostratigraphy are widely applied today and oﬀer highly complementary methods of correlating rock pack- ages of diverse lithologies. Other techniques for ordering and understanding the deposition of sedimentary rocks that will be discussed in detail in subsequent chapters, including sequence stratigraphy and cyclostratigraphy.
Table 1.1: Subdivisions of rock units/geological time for lithology, biostratigraphy, chronos- tratigraphy, and geochronology. Modiﬁed from Nagy and Bjorlykke (2010).
Lithostratigraphic Biostratigraphic Geochronological Chronostratigraphic Supergroup Assemblage zone Eon Eonothem Group Range zone Era Erathem Subgroup Acme zone Period System Formation Interval zone Epoch Series Member Biozone Age Stage Bed Chron Chronozone
1.3.2 Stratigraphic contacts Arguably as important as strata themselves are the contacts that separate strata. The nature of these contacts is important because they indicate whether sedimentation was eﬀectively continuous (in geological time) or episodically interrupted by intervals of non- deposition or erosion. The following are the main types of sedimentary contacts:
• angular unconformity
You may come across the term diastem, which is the geologically fancy way of implying an interval of non-deposition or local erosion, without any major change in depositional environment. In this sense, it is something like a minor disconformity. Chapter 1: Introduction to Stratigraphy 5
Traditional stratigraphy is focused largely on the physical character of stratiﬁed rocks. That is, it entails the description of the grain size, colour, sedimentary structures, sed- imentary textures, nature of bedding and bed boundaries. Lithostratigraphic units are deﬁned as the “bodies of rocks, bedded or unbedded, that are deﬁned and characterized on the basis of their lithologic properties and their stratigraphic relations.”1. The typical way of representing these data is on a stratigraphic log, where height (for a measured section in outcrop) or depth (for a drill) core is plotted on the y-axis. Hence, individual logged units are repented by boxes whose thickness corresponds to its actual thickness or height. The nature of the x-axis is variable, but the most traditional method is to depict grain size, where narrow boxes represent ﬁner grained sediments and wider boxes coarser grain sedi- ments. The boxes themselves usually have some kind of colour or symbology that reﬂect the lithology of the unit, and additional characteristics such as style of bedding, secondary features, or key sedimentary structures are shown in symbols or text to the right (Fig. 1.1). However, the x-axis may also be used to reﬂect relative water depth (where narrow is deeper and wider is shallower). Other times, no variation along the x-axis is shown (most common in drill cores). Rather than simply lithology, lithofacies (see below) are shown instead. In short, there are many ways to depict sediments and sedimentary rocks in logs.
The style and quality of stratigraphic logs is highly variable. This variability reﬂects dif- ferences in the scale of the logged section or core, the purpose of the log, and of course, individual preferences in plotting the stratigraphic logs. The result is that it can often be diﬃcult for one stratigrapher used to a particular logging scheme to appreciate and comprehend another stratigrapher’s log. For this reason, it makes sense to ﬁnd the right comprise between universality and practicality with regards to a speciﬁc goal or project.
The fundamental lithostratigraphic unit is the formation, which should be deﬁned such that its upper and lower boundaries are easily recognizable in the ﬁeld (hence based on clear lithologic properties) and the unit as a whole is mappable at roughly 1:50K scale. Formation thickness can be highly variable and depend on when an area was mapped and the unit deﬁned, the nature of the exposure (well exposed units are commonly further subdivided), and of course, the nature of the sediment. Most formations range from about 50–300 m-thick (Nagy and Bjorlykke, 2010), but exceptions may be as little as a few meters thick (for example some highly distinctive and mappable cap dolostones to Neoproterozoic snowball glaciations) and up to many kilometres thick. Formations should be named after a geographic feature in the area in which it occurs (e.g. mountains, towns, or water bodies) and should be formally deﬁned in the literature.
A protocol has been established for the formal deﬁnition of formations and other lithos- tratigraphic units1 Formal deﬁnition requires reference to a stratotype, or type section, the precise location of which is given (so other geologists can easily investigate it). Ideally, the stratotype is placed where the formation is best represented and where the boundaries are easily visible. Because these boundaries should be recognizable in the ﬁeld (i.e., map- pable), they are typically deﬁned on the basis of lithology rather than other stratigraphic zonation boundaries or techniques. Where possible, the same name should be used for a
1From the International Stratigraphic Guide, available from www.stratigraphy.org 1. Chapter 1: Introduction to Stratigraphy 6
SILICICLASTICS conglomerate 150 sandstone siltstone shale
CARBONATES Lithology grainstone 100 stromatolite ribbonite
SYMBOLS ooding surface cross-lamination 50 hummocky cross- stratication
0 (m) m si ms cg
Figure 1.1: A sample stratigraphic log of the Chandindu Formation, Yukon, modiﬁed from (Kunzmann et al., 2014).
formation wherever it is easily recognizable, and where multiple names have been applied in the literature, the ﬁrst formally deﬁned name has the priority. As you might imagine, signiﬁcant problems arise where many diﬀerent workers mapping in one area establish dif- ferent nomenclature. It is not uncommon for one formation to have multiple names, which of course leads to confusion and ineﬃciency. By convention, formations that have not been formally deﬁned are noted by lowercase ‘f’ in ‘formation.’
Similar concepts apply for other lithostratigraphic subdivisions (Table 1.1). Formations may be subdivided into members, in eﬀectively the same way that the formations them- selves are established. Members may also be mappable and can be formally deﬁned, al- though more often than not, they remain informal. Although the next smaller subdivision Chapter 1: Introduction to Stratigraphy 7
Foreland Mulden Group Basin
Maieberg Fm Tsumeb
Subgroup Ghuab Fm
Ombaatjie Fm Gruis Fm Passive Rasthof Fm Margin Abeneb Subgroup Chuos Fm Otavi Group Devede Fm
Okakuyu Fm DAMARA SUPERGROUP DAMARA
Subgroup Beesvlakte Fm Ombombo Rift Nosib Group Basin
Figure 1.2: The lithostratigraphic subdivision of the Neoproterozoic Damara Supergroup in northwestern Namibia. Notice that each group corresponds to a diﬀerent tectonic phase in the Wilson Cycle.
of a member is a bed, beds are not typically named unless they are particularly distinct or useful in regional correlation. Laminations are by deﬁnition thinner (<1cm) than beds, but are not a subdivision of a bed. Rather they are a thinner equivalent of a bed. Both laminations and beds are considered to have been deposited during a single depositional event. The contact or boundary between two beds is known as the bedding surface, and the nature of this surface (e.g. straight versus wavy, transitional versus sharp, scoured, mud-rich) oﬀers important information about the type of deposition and the depositional environment.
Formations are grouped together into subgroups or groups, which typically represent a series of diﬀerent units formed in a single tectonic environment, but perhaps separated by unconformities or major lithological changes. Groups may then be lumped into super- Chapter 1: Introduction to Stratigraphy 8 groups, which usually comprise one tectonic cycle (for example, a wilson cycle comprising rift, drift, and collision successions; Fig. 1.2.)
1.5 Depositional Environments
Sediments accumulate in sedimentary basins, which require some long-term source of ac- commodation space for the sediments to accumulate and be preserved. Sedimentary basins can be somewhat arbitrarily divided into a suite of depositional environments, each of which can be characterized based on their sediment types and sources, hydrological con- ditions, tectonic controls, preservation potential, biology, chemical conditions, and other parameters. Although certain lithologies and sedimentary structures may occur in multiple diﬀerent environments, a trained sedimentologist can often deduce a depositional setting of a package of sediments based on the ensemble of characteristics2.
Depositional environments are commonly classiﬁed into multiple, somewhat arbitrarily chosen categories. Here I have decided to use 1) continental, 2) siliciclastic coastal (sili- ciclastic), deep marine, carbonate platform, and glacial. Depositional environments are discussed in detail in Sedimentary Geology and are only brieﬂy reviewed here.
1.5.1 Continental Environments Continental environments are obviously those that occur on land. Although there are many places where sediments are deposited and stored temporarily on land, they are not places where they are ultimately preserved as sedimentary rocks. Those that manage to lithify remain vulnerable to erosion. Nevertheless, rocks deposited in terrestrial environments are important components of the sedimentary record, and of course, the main source of information about the ﬂora and fauna that inhabited the continents in the past.
Alluvial fans Alluvial fans are lenticular, convex (in transverse section) accumulations of usually coats- grained sediments that form where rivers exit mountain gorges. The streams drop their sediment bed loads, which may contain considerable and large sediments due to strong stream ﬂow and conﬁned channels in the mountains. In longitudinal cross-section, alluvial fans and concave up, which reﬂects the rapid facies changes away from the point at which they leave mountains; conglomerates and sands deposited by sheet ﬂoods and debris ﬂows give way to channelized sands and silts towards the toe of the fan. Fans will commonly coalesce laterally (to form bajadas) and interﬁnger with lacustrine or ﬂuvial deposits down- stream.
Whereas fans do occur in humid environments, they are best known and most important in dry environments which are prone to ﬂash ﬂoods and have less vegetation to stabilize sediments. 2It should also be noted that sedimentologists often do not agree on the depositional setting and it is not uncommon for two successive sedimentologists to interpret one package of rocks as representing contrasting depositional environments Chapter 1: Introduction to Stratigraphy 9
Rivers Rivers are broadly subdivided into braided and meandering. Braided rivers form braid plains, which are dominated by coarse (sand and gravel) sediments deposited as bars in rivers that experience considerable ﬂuctuation in discharge (say, due to spring melting in adjacent mountains). Most ﬁne-grained sediment is washed down river, although minor ﬂoodplain muds may be deposited.
Meandering rivers produce a distinct geometry of deposits with characteristic sedimentary elements. The channels of meandering rivers migrate by cutting in the outer bank (cutbank) of curved channels, while a point bar on the opposite side accretes laterally in the same direction. The result is that the channel ﬁlls with a ﬁning upward sequence of coarse- grained sediments. The broad ﬂood plain adjacent to the channels, along with oxbows (abandoned meanders), ﬁll predominantly with muds delivered during ﬂoods. The ﬂood plains may accumulate signiﬁcant organic matter, in the form of organic-rich muds, peat, or coal. During ﬂoods, channels may breach their banks, forming distinct ﬁning-upward, lens- shaped deposits within otherwise muddy sediments. Meandering river systems produce a sedimentary package that is characterized by shoestrings of sand (the point bars) encased in ﬂoodplain muds.
Lakes Large lakes of the sort that are prone to being preserved in the stratigraphic record form in structural depressions. Water-dominated lakes are open systems, and as such, do not have a lot of chemical sediments. Close-system lakes, such as Salt Lake Utah, tend to accumulate chemical precipitates, such as carbonate, salts, and other evaporites. Lakes may have many similar features to marine systems, such as deltas, beaches, stromatolites, and even turbidites and other gravity ﬂow deposits. However, they lack tidal inﬂuence and tend not to have smaller waves due to the shorter fetch compared to open marine settings. Where there are abundant fossils that can be used to indicate a fresh water or otherwise non-marine system, it can be diﬃcult to distinguish between lacustrine and marine environments.
Deserts Although many deserts are not sandy, sand dunes are the characteristic sedimentary feature of arid environments. Sand dunes form distinctive deposits with large, tangential cross-beds that are commonly amalgamated. Whole dunes are rarely preserved, and it is the toe of the lee side of the dune that is most often preserved owing to the stabilization of the sands by vegetation, groundwater, or ephemeral (playa) ponds and lakes. The sands in ancient dune deposits are typically reasonably well sorted, often with a bi-modal distribution, and moderately to well rounded grains consisting mainly of quartz. Iron oxide coatings and frosted or pitted grain surfaces are common.
1.5.2 Clastic Coastal Environments Most sedimentary rocks of marine origin preserved on the continents were deposited in the coastal environment. This is where sediments are delivered, either by coastal erosion or rivers, and given the rapid diminution of energy away from the coast, most coarser-grained Chapter 1: Introduction to Stratigraphy 10 sediments are trapped here. The main inﬂuences on the style of deposition on clastic coastal environments are the eﬀects of rivers (deltas), tides, and waves. While rivers act to deliver sediments, tides and waves redistribute those sediments, producing distinct bedforms and other depositional elements in the process. It common to subdivide the clastic coastal environment into three sub environments: river-dominated, tide-dominated, and storm- dominated.
River-dominated coastlines As you would expect, river-dominated coastlines are heavily inﬂuenced by deltas. Deltas can be subdivided into three broad settings: the delta-plain, delta-platform, delta-slope, and pro-delta.
The upper delta plain is the subaerial (above mean high tide) portion of the delta, and as such in inﬂuenced mainly by ﬂuvial processes. Swamps commonly occur between river channels.
The lower delta plain sits between low tide and high tide, and as such, is largely inﬂuenced by tidal processes. Here the river branches into a series of sandy, distributary channels, which are separated by muddy, brackish, distributary bays.
The delta front is that part of the delta lying between low tide and about 10 m – that depth eﬀective by fair-weather waves. It consists mainly of sands forming bars where distributary channels enter the ocean and ridges of sand arranged parallel to the coast-line by break waves.
The pro-delta consists of sloping beds of sand and silt that form the characteristic foresets of a delta system. Depostion is through a combination of settling from jets and plumes of sediment and gravity ﬂow processes.
Sediments seawater of the pro-delta are shelf muds. When a delta progrades, it gives rise to a coarsening upward (shelf muds through delta front sands), then ﬁning upwards (into interdistributary muds) sequence. Deltas commonly migrate laterally (or radially) as individual channels clog with sediments and avulse, givinh rise to an apparent (but usually false) cyclicity.
Tide-dominated coastlines Tide-dominated coast lines are distinguished by broad tidal ﬂats, estuaries, and subaque- ous, linear elements arranged perpendicular to the shoreline. Sands tend to be heavily cross- bedded, and in places my preserve bidirectional cross-sets. Prograding, tide-dominated systems produce a ﬁning-upward sequence, from subaqueous, tidal sands to mud-cracked muds and silts with ﬂaser, wavy, and/or lenticular bedding.
Wave-dominated coastlines Storm-, or wave-dominated coastlines are those in which the retribution sand is largely the result of waves, which forms long beach and barrier island systems resulting in elongate, shore-parallel sand systems. Barriers are sandy, shore-parallel islands or peninsulas that Chapter 1: Introduction to Stratigraphy 11 are separated from the mainland by lagoons or marshes. Lagoons are quiet environments, often anoxic due to protection from waves and storms. Storms and tides generate washover deposits in the for of beds and tidal deltas.
The beach environment is typically fairly narrow, but can be volumetrically important in prograding systems. The beach environment is subdivided into three domains:
• backshore occurs above mean high tide and consists mainly of of dune sands along with minor storm deposits.
• foreshore is the zone between mean high tide and mean low tide. It experiences steady swashing of waves (swash zone), giving rise to parallel and heavy mineral laminations common. Low-angle cross-sets also occur due to a change in slope of the foreshore between seasons.
• shore face is the zone between faire weather wave base and mean lot tide. It includes the breaker zone, where waves begin to touch the seabed and water become turbulent, and the surf zone, where breaking waves become translational waves, setting up a strong coastward current. Trough cross-bedding, symmetric ripples (some quite large), and asymmetric ripples are al common. Beach systems give rise to mixed sands and muds oﬀshore.
Storm-dominated coastlines Storm-dominated coastlines are those in which strong storms tend to transport sediments away from the coast-line towards the middle and outer shelf. During storms, coastal set up drives a current away from the beaches, which entrains sand and carries it out to the open shelf, where it is simultaneously reworked by chaotic, large, storm waves. Character- istic bedforms in this environment include swaley cross-stratiﬁcation (SCS) and hummocky cross-stratiﬁcation (HCS). Both bedforms are the result of the interaction between uni- directional and multi-direction (storm wave) currents (Dumas and Arnott, 2006) and are distinguished by other forms of cross-bedding by
• curved, fanning laminations
• low-angle lamina intersections
• sharp, often erosional, lower bounding surface
• concave up cross-sets that grade upward into convex up cross-sets and vice, versa
SCS is occurs closer to the fair-weather wave base, is coarser grained (medium sand), and more of an erosional feature (it ﬁlls scours). HCS occurs closer to storm-weather wave base and, comprises silt and ﬁne sand, and as its name implies, forms hummocks. Storm inﬂuence beds that do not contain SCS or HCS are distinguished pinch and swell structures, gutter casts, or wavy heterolithic beds, often containing starved ripples. Storms also tend to kick up a lot of ﬁne-grained sediments that is also transported oceanward, often in nepheloid currents, which are gravity ﬂows driven by buoyancy. That is, the waters close to the seabed that entrain the ﬁne sediment are denser than the surrounding seawater and so ﬂow downslope, towards the edge of the continental shelf. Chapter 1: Introduction to Stratigraphy 12
Figure 1.3: Depiction of typical sedimentary facies (see below for a description of ‘facies’) deposited during fair-weather and storm-weather conditions on a clastic continental shelf. From Dumas and Arnott (2006).
1.5.3 The Deep Marine Environment The deep marine environment consists of the outer shelf (neritic), slope, rise, and and abyssal plain. The outer shelf shelf environment is dominated by ﬁne-grained pelagic sed- imentation, along with silts and ﬁnd sands transported by storms. Coarser sands may be transported to the outer shelf in submarine channels connected to river systems (and often a vestige of lower sea level).
On the slope, sediment deposition is the result of both gradual, semi-continuous pelagic sedimentation and gravity ﬂow processes, which include slumps (breccias, syn-sedimentary folds, detachment surfaces), debris ﬂows (diamictites), and turbidity currents (turbidites, e.g., the Bouma sequence). Grain ﬂows are less common, but also occur. On the slope, these gravity ﬂow deposits tend to be channelized, and the channels have similar structures to ﬂuvial channels, such as levees and overbank deposits.
Along the toe of the slope (the rise), gravity ﬂow deposits often accumulate in lobes, similar in form to alluvial fans.
In the open, deep ocean, sedimentation is dominated by ﬁne-grained sediment. This pelagic sediment is in the form of either wind-blown dust or the ﬁne tests of silica- or carbonate- secreting microorganisms. The main culprits of this type of sediment are the diatoms Chapter 1: Introduction to Stratigraphy 13
(siliceous photoautotrophs), radiolarians (siliceous heterotrophs), coccolithophora (calcare- ous photoautotrophs), and foraminefera (largely calcareous heterotrophs). Calcareous sed- iments are subject to the dissolution in deep waters: they begin to dissolve at the lysocline and are completely absent at the carbonate compensation depth. The result is that deepest marine sediments lack carbonate sediments. In shallower, but still deep water sediments, the rapid disappearance of carbonate sediments in a drill core may reﬂect shoaling of the CCD due to, for example, global warming and acidiﬁcation of the oceans.
1.5.4 Carbonate Platforms Most carbonate systems occur on stable, shallow, tropical shelves that receive little clas- tic input. Much of the carbonate is biogenic, derived from corals, bryozoans, sea grasses, coralline algae, and many other carbonate-secreting organisms. Most require sunlight and warm waters to precipitate carbonate. Where such conditions occur, carbonate production can be eﬃcient and rapid, such that sediment accumulation easily keeps up with any local rise in sea level (say, due to ongoing subsidence).
Rimmed carbonate platforms are protected from the open ocean by a reef system, which usually contains a diverse mixture of reef-buliding organisms and reworked carbonate grains (including ooids). Seawater of the barrier, carbonate slumps onto the talus slope and is re- worked by gravity ﬂow deposits, much the same way as on the continental slope. Similarly, carbonates may form turbidites, diamictites, and grain ﬂows. Basinal facies consist of ﬁnely laminated carbonates, perhaps with laminations or partings of pelagic mud. Shoreward of the barrier island is mostly shallow water, comprising sand shoals, lagoonal deposits, and tidal ﬂats.
Carbonate platforms may also be storm dominated, in which case they closely resemble clastic, storm-dominated shelves. The fundamental diﬀerence is the source of sediments: the are produced within the basin on a carbonate shelf, and are epiclastic, or extrabasinal on a clastic shelf.
Carbonate platforms are highly susceptible to environmental change. An inﬂux of clastic sediments or abrupt deepening event can shut oﬀ carbonate deposition, which may only return with a return to formerly stable conditions.
1.5.5 The Glacial Environment Glacial environments can be quite diverse, because they may occur both in the marine and continental realm and may entail large ice sheets or small valley glaciers. In the terres- trial environment, glacial sediments include tills, which are ice contact deposits, namely moraines and lodgement tills (that form at the base of glaciers where the pressure of the ice pushes or lodges sediment into open spaces). Glacial tills are often diamictites, meaning they are mud-supported. The clasts may be far ranging (or exotic) in tills deposited by large ice sheets, faceted, and/or striated. Periglacial environments include glacial wash-out fans, which occur downstream of a glacier and typically experience high sedimentation rates. Other features that might help distinguish ancient terrestrial glacial environments include
• Eskers: deposits from subglacial streams Chapter 1: Introduction to Stratigraphy 14
• Drumlins: whale-shaped waveforms that are sculpted sub glacially into unconsoli- dated sediments • Roche moutonn´ee: an erosional, tear-shaped feature sculpted into bedrock and facing downstream
The glacial-marine environment is often characterized either by very high sedimentation rates, where glaciers are subglacial streams are depositing a lot sediment, or extremely low sedimentation rates, where ice shelf ice buttresses glaciers and impedes pelagic sedimenta- tion.
Sedimentation from the front of a calving glacier results from meltwater ﬂow and plumes, density currents, gravity ﬂows, and ice rafting. Generally, coarser and more poorly stratiﬁed environments occur proximal to the ice grounding line and better stratiﬁed and over all ﬁner-grained sediments occur more distally (Fig. 1.4.)
Figure 1.4: cartoon depicting depositional processes and sediment types in the glacial- marine environment, from Edwards (1986).
1.6 Lithofacies and Associations
The term facies is one of those annoying words that is hard to deﬁne, but is widely used in the Earth sciences. The problem is partly that is used in many diﬀerent ways. In terms of sedimentology, the term facies refers to the suite of of lithological and biogenic features of sedimentary rock, which together, point to deposition in a speciﬁc environment. These features include the texture, composition, and sedimentary structures. The term lithofacies eﬀectively means the same thing, but is more speciﬁc so it is not confused with other uses of the word ‘facies.’
The facies concept is extremely useful for a variety of reasons. First, it forces a sedimen- tologist/stratigrapher not just to describe the rocks, but also interpret their depositional signiﬁcance. This is particularly important if we consider the fact that many lithological features may occur in diverse environments. Consider, for example, ripple cross-laminated sands. This may occur in aeolian, ﬂuvial, coastal marine, or even deep marine environ- ments, so just identifying a ripple cross-laminated sandstone alone is not suﬃcient for Chapter 1: Introduction to Stratigraphy 15 parsing the depositional setting of that sandstone. Other features, such as the maturity of the sands, the nature of the bedding planes, or associated lithologies (facies) will come into play in determining what those sediments represent.
Another practical advantage of identifying facies is that it makes it much easier to log and correlate sections. So, for example, say you are working on a series of sedimentary succes- sions through some Eocene continental deposits in the Bighorn Basin of Wyoming/Montana. Rather than logging sections where you break out every single bed (which might be inordi- nately tedious and arguably not useful), you might instead identify a series of facies, which you deﬁne and interpret. Once this has been done, you might simply log facies. Some of these facies might be ﬂood plain muds (muds with bioturbation, root casts, maybe thin coal stringers, occasional graded sand lenses), crevasse-splay sands (grades beds with rip- ple cross-laminations, commonly overlain by ﬂood-plain muds), and channel gravels(lag conglomerate and trough cross-bedded, coarse sand. Pebble imbrication and tool marks common), and point bar sands (moderately sorted sands, trough-cross bedded sands, grad- ing upward into ripple cross-laminated sands).
By plotting facies in a stratigraphic log, you impose your interpretation on that log. Some purists may argue this is not appropriate, objective approach. However, consider the alter- native. If you simply logged lithologies and structures rather than facies and did not get around to publishing your results, then those stratigraphic logs are virtually useless down the road. It would in most cases be diﬃcult for you and nearly impossible for another stratigrapher to interpret the depositional setting based on those logs alone.
1.6.1 Walther’s Law Retrograding and prograding shorelines give rise to stacked sedimentary facies. When we see a succession of facies in sedimentary section (uninterrupted by unconformities), Walther’s Law tells us that this reﬂects sedimentary environments that occurred adjacent to one another. Hence, we can use Walther’s Law to reconstruct depositional environments and whether rocks were deposited during a transgression or regression. And even though it should now be evident that the inherit instability of many depositional environments in time makes correlating sedimentary strata based on lithology alone is prone to error (assuming the goal is to correlate time-equivalent strata), these same ﬂuctuations yield patterns that are often correctable (for example, sharp transgressive or regressive surfaces).
1.6.2 Facies Associations You can consider ‘facies’ as one rung in a hierarchy of sedimentary features, where in- dividual laminae or beds are the fundamental unit, and these are grouped in a facies. Whereas many facies might be speciﬁc to a certain, broadly-deﬁned depositional environ- ment, many other facies might be non-unique. However, when interpreted in the context of their neighbouring facies, they should (if properly described and interpreted), point un- ambiguously to a speciﬁc depositional environment. A grouping of depositional facies is a facies association. Chapter 1: Introduction to Stratigraphy 16
It is rarely suﬃcient in the Earth sciences to study only a single drill core or stratigraphic section. In order to gain more information about a sedimentary unit, the depositional envi- ronment, the tectonic origin and evolution of a basin, the possible distribution of econom- ically important units, etc., correlation is required. Correlation is a way of tying together the envisioned original spatial and temporal connections between diﬀerent locations. In basin analyses, basin-scale correlations are necessary to reconstruct tectonic setting and the evolution of the basins.
Figure 1.5: Stratigraphic restoration of the Neoproterozoic Otavi platform (see also Fig- ure 1.2) on the southwestern promontory of the Congo craton, northern Namibia, from (Hoﬀman, 2011). This stratigraphic cross-section is based on dozens of measured sections, correlated through a combination of sequence stratigraphy, lithostratigraphy, chemostratig- raphy, and geochronology. Only through correlating all of these sections can the geome- try and evolution of the basin be deduced. Such information is essential in mineral and petroleum exploration of basins and tectonic studies. It is, of course, also very important in interpreting records of environmental evolution.
There is no one single way to make correlations, and as you might guess, correlations are subject to debate, argument, confusion, and change. Classically, correlations were based largely on either lithological features or biostratigraphic features. So, for example, if a formation boundary is marked by a sharp shift from shale to sandstone, then this boundary might be readily correctable over a region. However, that boundary may well be diachronous, corresponding to the gradual progradation of the sandy unit into the basin. Chapter 1: Introduction to Stratigraphy 17
Biostratigraphy can overcome this diachroneity problem in that biostratigraphically deﬁned zone boundaries mark turnover in fossils, and so should not be diachronous. However, a shortfall of biostratigraphy is that it is environment-dependent in that many organisms only occur in certain environments, and environments themselves govern the nature of fos- sil preservation. So the disappearance of certain fossils up section could indicate a change in environments to one not prone to that type of organism rather than a key biostrati- graphic boundary.
For many purposes, temporal correlations are the ideal in stratigraphy. This is because geologists and palaeontologists will commonly want to compare environmental conditions across a basin within temporal time slices. So in order to understand the spatial dimension of a mass extinction and subsequent recovery, it is important to correlate temporally so that the broader structure of the environment can be deduced in each of those time slices. Sometimes, temporal correlations are not so diﬃcult. As we’ll see from sequence stratig- raphy, certain boundaries the recur within stratigraphic packages are precisely or nearly time-equivalent across a basin, and so can be used for temporal boundaries. Volcanic tuﬀ are the ultimate correlation tools, both because they can be traced across multiple environ- ments and because they are amenable to precise radiometric dating. However, more often than not, temporal correlations are often not easy. Much of this course will, in fact, be dedicated to various stratigraphic techniques that serve to enhance temporal correlations. Chapter 2
2.1 Well Logs
There is exist much more data from drill holes and drill cores than from outcrop sections, although much of this data is proprietary and largely unavailable to the public. Nonethe- less, drill holes are extremely important in petroleum exploration, basin analysis, sequence stratigraphy and ocean drilling. When a hole is drilled and a core is recovered, the core is logged at the drilling platform when it is ﬁrst brought to the surface, but usually logged again in more detail in a core library or perhaps some temporary core repository. Logs of drill cores are similar to stratigraphic logs measured in the ﬁeld, with a few important diﬀerences. First, drill cores provide a high percentage of recovery, compared to outcrop sections, which are usually partially obscured by non-exposure or vegetation. Second, the rock is fresh compared to outcrop sections that are weathered. For visual logging purposes, this can be both an advantage, and a disadvantage, particularly for carbonates, where weathering highlights sedimentological features. Third, they are one-dimensional–that is, one cannot look laterally for the occurrence or importance of facies changes, faults, or other 2–3D phenomena which might signiﬁcantly inﬂuence interpretation of a core. Nevertheless, drill cores provide a unique opportunity to established detailed lithological descriptions of a section of rocks. They are also much better suited for chemical and geophysical analy- sis due to their continuity and relatively low degree of alteration due to surﬁcial weathering.
Drill core is extremely expensive to produce, particularly in deep drill holes. Hence, holes are cored only when there is speciﬁc need for the stratigraphic information the core will provide. In other cases, the cuttings removed from drilling might be retained for logging and geochemical purposes, but these obviously do not provide for high resolution study and sampling. On the other hand, a variety of geophysical logs are obtained during drilling that provide continuous data that are important in establishing lithology, porosity, saturation with water, oil, or gas, and other characteristics of the penetrated stratigraphic succes- sions. Classically, these data were collected in increments of the drill hole after drilling and prior to casing (if the hole was to be cased) using a sonde, which contained the various gadgets necessary to make the measurements and was attached to a wire connected the surface (hence the common term wireline logs). A calliper that measured the radius of the drill hole, was also passed through the hole. These days, logging tools are commonly incor- porated inside the drill bit such that the measurements are made simultaneous with drilling.
18 Chapter 2: Geophysical Data 19
Log readings may be strongly inﬂuenced by the drilling mud which is used in the drilling process to balance the water or oil/gas pressure in drill hole and transfer cuttings back to the surface (Bjorlykke, 2010). Some of this mud remains, lining the hole (Fig. ??). This mudcake may penetrate into the sediments, throwing of certain geophysical measurements.
Figure 2.1: Cartoon depicting the invasion of drilling ﬂuids into formations. Note the mud cake is the particulate component of the mud that lines the bore hole, and it is the ﬂuids from that mud that invade the more permeable part of the formation. Because electric measurements are dependent on electrical diﬀerence between the drilling mud and the formation ﬂuids, this invasion of drilling ﬂuids can eﬀect the quality of measurements. Figure from Bjorlykke (2010).
Below are descriptions of the most important type of well logs, along with a few other imaging and analytical techniques that may be applied to drill holes or cores.
2.1.1 Electric logs Electrical logs were the ﬁrst to be employed in drilling because they are relatively easy to acquire. The two classic electric logs, spontaneous potential and resistivity, are useful in concert for determining the location of oil or gas saturated permeable rock. These readings must be made before a hole is cased.
Spontaneous-Potential The spontaneous potential or self-potential (S.P.) log was the ﬁrst geophysical log to be used on drill holes (Selley, 1998). This measurement, in millivolts (mV), records the electric potential between an electrode located in a sonde in the drill hole and another electrode at the Earth’s surface. The electric charge that is measured is caused by the ﬂow of ions (mainly Na+ and Cl−) between formation waters and drilling mud. As such, it only works in mud-ﬁlled and un-cased drill holes where the drilling muds have a diﬀerent salinity than Chapter 2: Geophysical Data 20 the formation waters. The ﬂow of ions is driven by diﬀusion. Any deﬂection of the of the measurement from an arbitrarily chosen shale baseline reﬂects diﬀusion. Consequently, the measurement is controlled both by the diﬀusion gradient and the permeability. Given that most drilling muds are fresh, if the drill core is through marine sediments, those sediments should have marine formation waters, hence relatively high ionic concentrations. In this case, the ﬂow of ions is from sands to the muds, and a negative reﬂection is recorded. A similar but opposite scenario will occur where drilling muds are salty and formation waters are fresh (but the deﬂection from the shale baseline will still give away where the porous sands are). If there is no strong contrast in salinity between the drilling muds and the sediments, then the S.P. log will show very little deviation.
Resistivity Resistivity, measured in ohms, is the resistance of a rock or sediment to the ﬂow of an imposed electrical current. Electrical resistivity is measured on electrodes in contact with the rocks in the well wall. Most minerals are insulators, and hence are highly resistive to electrical ﬂow. The exceptions are the clay minerals and Cl salts, which are conductive. In other lithologies, most conduction takes place within the pore ﬂuids and the extent of that conduction is a direct function of the quantity of dissolved salts and the permeability of the rock. Consequently, sediments with marine (salty) pore ﬂuids are less resistant to elec- trical currents than fresh waters. Resistivity measurements are sensitive to the inﬁltration of drilling ﬂuids into permeable formations, so some sort of correction has to be made to account for this inﬁltration. As a consequence, the electric properties of the drilling mud must be well known. If the drilling mud is oil-based, S.P. measurements cannot be made.
Resistivities in rocks can vary greatly. Solid, well-cemented rock is highly resistive. Porous rocks that are saturated in oil, gas, or freshwater are also resistive. Salty pore waters are not. Hence, resistivity can be used to establish the contacts between oil and water or gas and water in a reservoir. Resistivity logs can be used in combination with SP logs to make qualitative assessments of the distribution of permeable units saturated with oil, gas, or fresh water (Fig. 2.2). Because the reading reﬂects porosity, the relative intensity of the S.P. can be an indicator of how pure a sand is, hence the degree of sorting and abundance of mud or cement occluding pore spaces.
2.1.2 Radioactivity logs The two most important radioactivity logs are the gamma ray and neutron logs. Both can be measured on a drill hole after it has been cased.
Gamma ray Gamma ray logs, is measured in API (American Petroleum Institute) units and records the natural radioactivity emitted by a rock (or sediment). The main sources of radioactivity in typical sedimentary rocks are from the elements K, Th, and U. Shales tend to be relatively rich in all of these elements compared to sandstone, such that gamma ray logs are used most often as lithological indicators. Of course, there are exceptions. Sandstones with abundant feldspar and mica or signiﬁcant concentrations of heavy minerals will have higher gamma ray intensity than clean, quartz sandstones. Limestone, halite, and gypsum/anhydrite also have low U, Th, and K concentrations and so give low gamma ray readings. However, Chapter 2: Geophysical Data 21
Figure 2.2: Cartoon showing how coupled spontaneous potential (S.P.) and resistivity (R) logs can be used to discriminate between impermeable and permeable units and saturation with a non-conductive (fresh water, oil, gas) versus a conductive (salt water) ﬂuid. A) continuous shale; B) A tight sandstone or carbonate bound by shale; C) A permeable unit saturated with salty pore ﬂuids; D) A permeable unit saturated with fresh water, oil, or gas. Modiﬁed from (Selley, 1998).
K-rich salts, as you might expect, show up very clearly in gamma ray logs. Organic matter concentrates U from oxidized seawater. As a result, black shales, which are common source rocks, tend to have the highest gamma ray readings (hot shales, reaching upwards of 200 API.
Whereas older gamma ray scintillation counters did not distinguish between radioactivity emitted by K, Th, and U, newer instruments are able to based on the diﬀerent energy levels of radiation emitted by diﬀerent isotopes. As a consequence, K, Th, and U concentrations can be roughly measured in drill holes and provide more precise measurements of shale content since Th content tends to be less variable in shales than K and U.
Neutron logs Neutron logs are based on the bombardment of the rock with neutrons from a radioactive source. These neutrons trigger the emission of gamma rays from hydrogen in the rocks, which is then measured. Hydrogen is concentrated, of course, in water and petroleum, Chapter 2: Geophysical Data 22 such that high neutron activity reﬂects porous and saturated rocks (note that it is not sensitive to the connectivity of the pore spaces, and so is not a measure of direct measure of permeability). The neutron log response is sensitive to the width of the bore hole, which can be variable due to collapse in the hole ore other factors, so it has to be calibrated using the calliper log, which is simply a measurement of the eﬀective diameter of the hole.
Shales tend to contain more hydrogen then other lithologies, and so yield a response in the neutron log that is not related to porosity. In a similar way, impure sandstones with a signiﬁcant clay matrix record a false high porosity. The neutron log can also be used to distinguish between oil and gas because gas has a higher concentration of hydrogen atoms than oil (or water), and so yields a higher radioactivity (Fig. 2.3). Overall, the neutron log is the most reliable measure of porosity.
Figure 2.3: Cartoon illustrating the variable log response to lithology, porosity, perme- ability, and saturation. From Bjorlykke (2010). Note that by convention, logs that reﬂect lithology tend to be shown on the left, while those that indicate porosity occur on the right
Density logs Density logs measure the density of rocks and their pore ﬂuids by emitting gamma radiation and then recording the amount of gamma radiation returning from the rock. The return radiation is a function of the electronic density in the rock, in turn determined by the bulk density of the rock. Corrections must be made for the borehole diameter and thickness of mudcake on the bore hole wall.
2.1.3 Acoustic log Acoustic, or sonic logs measure the velocity of sound waves through rocks, and as such, provide an indication of porosity. As such, it is particularly sensitive to cementation. Accoustic logs may also be used to distinguish between what phase occupies that pore Chapter 2: Geophysical Data 23 space. Accoustic data is useful not only as log, but also because it provides travel times that can be used in geophysical techniques, namely seismic reﬂection data.
Figure 2.4: Depictions of typical well log proﬁles (Gamma Ray and Resistivity) through sediments deposited in various diﬀerent depositional environments. Note how GR in par- ticular can pick upward coarsening successions. Within pure sands, GR is fairly insensitive to grain size, but resistivity picks out upward ﬁning successions through their eﬀect on porosity.
2.1.4 Other Drill Hole and Drill Core Techniques Temperature logs Temperature logs measure the temperature of the borehole and case be used to determine the geothermal gradient. The geothermal gradient is useful for subsidence modelling of a sedimentary basin to determine its subsidence history, hence the timing when oil-prone source rocks entered the oil and gas windows. Where other rock properties are known or measured, heat ﬂow models can be established. Again, these are useful for establishing when source rocks would have reached maturation or over maturation. Such numbers and models are also essential in geothermal exploration.
Image logs The walls of a well can be directly imaged, revealing layering, sedimentary structures, fractures, and other ﬁne resolution physical characteristics of the drill hole.
Nuclear Magnetic Resonance Nuclear Magnetic Resonance, or NMR is a technique used to image bodies in medicine (MRI’s) and is increasingly employed in petroleum exploration. In short, this method imposes a magnetic ﬁeld on the formation, to which hydrogen nuclei respond. The response Chapter 2: Geophysical Data 24 of the hydrogen nuclei to the imposition and removal of this magnetic ﬁeld reﬂect the porosity of the formation. In combination with other logs, it can reveal information about composition as well.
2.2 Seismic Exploration
Seismic surveying yields 2D to 3D images of the subsurface through the estimation of physical properties based on the travel times of waves that are propagated through the Earth (Mondol, 2010). Because this techniques enables evaluating physical properties over large spatial scales, while still high enough resolution to pick up ﬁner scale structure in the subsurface, it is immensely important in petroleum exploration. Seismic data can be acquired on land or at sea; in both cases, it require a source for the acoustic waves. On land, the source is usually a vibrating truck, and less commonly, dynamite. At sea, air guns are used. Both require an array of phones to collect the returning seismic data, which then must be processed and stacked. Fully processed seismic data yield seismic sections that reveal the internal structure of the subsurface. Reﬂectors highlight contrasts in seismic velocity, and as such, pick out formation boundaries, faults, diapers, water and gas satu- rated formations, and other important physical data. The depth axis on a seismic section is in two-way travel times rather then actual depth. Whereas this information alone is highly useful in seismic interpretation and petroleum exploration, it can also be converted to actual depth with an appropriate velocity model for the subsurface. Such a velocity model is most easily developed through drilling and acoustic log data.
The quality of seismic data is highly sensitive to the techniques employed in the emissions of sound waves and the collection and processing of data. Vastly improved techniques over the past 30 years has revolutionized seismic exploration and resulted in amazingly high resolution and informative data. In addition, software for imaging data, coupled with more powerful computers, has enabled much more sophisticated ways of imaging this data. 3D seismic is now routine scientists can manipulate this data in many ways.
2.2.1 Basic Principles of Seismic Surveying Seismic waves come in four varieties: compressional (P-), shear (S-), Rayleigh and Love waves. All seismic waves travel at distinct velocities in diﬀerent Earth materials. The Rayleigh and Love waves are surface waves, which means they propagate along surfaces; they are of limited use in seismic exploration. P- and S-waves are body waves, which propagate radially outward from the source and are especially useful in exploration. In P-waves, particle motion occurs along the same axis as wave propagation. In S-waves, par- ticle motion is transverse to the direction of propagation, which entails displacement, or shear of the material as the wave passes. The result is that S-waves travel slower in solids than P-waves and do not propagate at all in ﬂuids, which do not have shearing capacity.
In seismic surveying, seismic waves are generated by pulses at the source, and the ampli- tudes and travel times of waves returning to the surface are recorded. When the waves encounter boundaries between layers of rock with diﬀerent seismic velocities, the waves will refract according to Snell’s law: Chapter 2: Geophysical Data 25
Figure 2.5: A comparison of 2D seismic sections acquired at sea showing how much better the imaging is using a bottom versus surface streamer of phones to collect the seismic data. But even the lower quality image above easily picks out the immense angular unconformity and normal fault in the subsurface. From Mondol (2010).
sinθ1/sinθ2 = V2/V1 (2.1) where V1 and V2 are the velocities in layers 1 and 2, respectively, and θ1 and θ2 are the angles of the incident and refracted waves (measured relative to normal incidence), respectively. If the seismic waves are incident at the velocity boundary at an angle, they will also generate reﬂected waves. The extent of the seismic energy that is reﬂected is dependent upon the impedance of the the two layers: the density times the seismic velocity: The reﬂection coeﬃcient, R, is deﬁned as
ρ V − ρ V R = 2 2 1 1 (2.2) ρ2V2 + ρ1V1 where ρ1 and ρ2 are the densities of the two layers. This equation shows that large contrasts in impedance give rise to high reﬂection coeﬃcients, and hence, greater reﬂective energy. The diﬀerent factors that govern impedance are lithology, porosity, type of pore ﬂuids, and extent of pore saturation. Shale and sandstone have signiﬁcantly diﬀerent impedances, and Chapter 2: Geophysical Data 26 so contacts between these to lithologies produce signiﬁcant reﬂected energy. Limestone has both high density and high velocity, and results in even higher impedance contrasts, unless it is very porous (Mondol, 2010).
The energy reﬂected from boundaries of impedance contrast (reﬂectors) is collected in a series of receivers (phones), which covert the energy (usual from P-waves) into an elec- tric signal. The data are presented as a vertical series of transverse waves which form a time series, where the time component is the two-way travel time (the time that a wave required to travel to a reﬂector and back to the surface) and amplitude is proportional to the strength of the reﬂector (Gluyas and Swarbrick, 2004). Because the angle at which the measured waves strikes the reﬂectors (assuming they are horizontal) is high, the reﬂective energy at the receives is low. To enhance the signal-to-noise ratio of the data, multiple receivers measure the arrival of waves reﬂected at a series of angles. The seismic traces from each of the receives is then stacked to provide a single trace for the region directly below the source. Stacking requires developing a model for the depth to the reﬂector and the seismic velocity through that layer.
The subsurface is generally not a single, ﬂat surface. Oﬀsets in reﬂectors (faults), and dipping and curved reﬂectors give rise to diﬀractive noise. Other disruptions and compli- cations to the seismic traces include multiples, which are from internal reﬂections within a layer with large impedance contrasts, signals reﬂecting back to the plane of seismic section from outside that section, and background seismic noise. Migration is the process dealing with these complications and attempting to position reﬂectors in the correct spatial loca- tion.
In order to visualize reﬂectors, the positive (or negative) portions of the seismic traces (the wavelets) are commonly shaded or coloured. When a complete, processed, seismic proﬁle is produced, it shows the spatial distribution of the reﬂectors and their terminations. At this point, a geologist can begin to interpret the data. However, for accurate interpretations and predictions of petroleum potential or other features, it is necessary to convert the two- way travel time (depth axis) to actual depth. This requires a combination of a velocity model and a geological interpretation. That is, velocity must be assigned to every layer, and these layers must be correlated across the seismic proﬁle.
2.2.2 Seismic Stratigraphy Seismic proﬁles provide 2D or 3D images of the sedimentary ﬁll of sedimentary basins. Such proﬁles are impossible to obtain on surface exposures and permit rather straight- forward correlation of major unconformities, which are the boundaries between tectonos- tratigraphic packages (megasequences). The sedimentary ﬁll within a megasequence can than be interpreted based on the reﬂectors, and operating on the assumption that these reﬂectors reﬂect time correlative or fault boundaries. In short, seismic stratigraphy is the stratigraphic interpretation of seismic data. Seismic stratigraphy was the predecessor to sequence stratigraphy, and so will be discussed in more detail in the following chapter. Chapter 3
It has long been recognized that sedimentary successions are separated into cycles or pack- ages separated by distinct boundaries, such as subaerial unconformities. In the 1950’s and early 1960’s, Harry Wheeler (U Washington), and then Larry Sloss (Northwestern U) be- gan to recognize the importance of these unconformities, both for mapping purposes and as time markers. In short, they began to break out lithological units based on their bounding unconformities, which set the stage for modern sequence stratigraphy. Wheeler developed the technique of plotting sedimentation in space (on the x axis) as a function of time (y axis), which highlights the importance of the unconformities. These plots are now known as Wheeler diagrams. Sloss famously subdivided the Paleozoic- sedimentary succession of eastern North American into the four megasequences: Sauk (Cambrian), Tippecanoe (Or- dovician), Kaskaskia (Silurian-Devonian), and Absaroka (Late Carboniferous to Permian). To a large extent, these early eﬀorts recognized the importance of unconformities and time in interpreting and subdividing the lithostratigraphic record, and these sequences are in eﬀect tectonostratigraphic packages in that the bounding unconformities are related to ma- jor tectonic events. The sequences themselves were vertical successions of facies arranged in a predictable and coherent manner (think Walther’s Law). One shortcoming of their approach is that it really only worked well on a basin-scale for these major boundaries. Smaller unconformities, which might be well developed on the edges of the continents due to smaller changes in sea level, disappear basinward where no exposure occurred. Hence, where multiple higher order sequences might be deﬁned on the basin margin, perhaps only a single sequence could be distinguished towards the basin center.
If it was the early insight of Wheeler and Sloss that set development of sequence stratig- raphy in motion, it was the development and reﬁnement of seismic stratigraphy in the petroleum industry and access to vast amounts of data (Gradstein et al., 2013) that pushed the accelerator. Petroleum geologists recognized early on that sedimentary packages on continental margins were separated by major, basin-wide reﬂectors and they subdivided strata into broadly transgressive-regressive packages. Distinct stratal terminations helped them recognized the boundaries to these packages (later to be deﬁned as systems tracts). By the 1970’s, sequence stratigraphy was beginning to take hold and revolutionizing how geologists study and reconstruct the stratigraphic record. Sequence stratigraphers further pushed the envelope by developing global eustatic sea level curves based on sequence strati-
27 Chapter 3: Sequence Stratigraphy 28 graphic data. Whereas these global eustatic curves must be weighed against evidence for local tectonics and mantle-induced dynamic topography (Moucha et al., 2008), which can play a signiﬁcant role and signiﬁcantly skew both eusatic and tectonic interpretations (Fig. ??.)
Figure 3.1: A prediction of the dynamic topography experience by the New Jersey Shelf to- day versus 30 million years ago from Moucha et al. (2008). Such dynamic eﬀects (driven by mantle upwelling and downwelling) can be signiﬁcant, and this particular result is impor- tant because global sea level curves have been developed in part on sequence stratigraphic records from this margin.
Although sequence stratigraphy employs some methods from traditional lithostratigraphy, it emphasizes facies relationships and the the geometry of time-equivalent stratal units. Hence, an overarching theme and aim in sequence stratigraphy is to develop stratigraphic models or frameworks that delineate chronostratigraphic boundaries such that sedimentary basin ﬁll can be interpreted in terms of its spatio-temporal evolution.
The sequence stratigraphic framework of a basin develops as the result of the interplay of generation (or destruction) of accommodation space and sediment supply. Speciﬁcally, the three processes that control the sequence stratigraphic evolution of a basin are
• rate of crustal subsidence or uplift
• eustatic sea-level change
• rate of supply of sediments
The slope of the basin margin also inﬂuences the geometry of sequence stratigraphic pack- ages. The interplay between these eﬀects gives rise to the development of sequences, which were historically deﬁned as ”a relatively conformable successions of genetically related strata bounded by unconformities and their correlative conformities.” However, as de- scribed in more detail below, the deﬁnition of what a sequence is can and should be broadened. Chapter 3: Sequence Stratigraphy 29
3.2 Fundamental concepts
3.2.1 Sea level, base level, and accommodation space We often interchange the terms sea level and base level when discussing sedimentary pro- cess, particularly those that occur near basin margins and are inﬂuenced by ﬂuctuations in sea level (by they tectonically or ecstatically driven).
Sea level is deﬁned relative to the centre of the earth. Hence, it can only rise and fall as a consequence of eustatic process. Local tectonics alone do not eﬀect sea level.
Tectonic uplift and subsidence are deﬁned relative to the centre of the Earth as well.
Base level is the equilibrium surface above which erosion takes place and below which sediment accumulation takes place. For many purposes, it closely approximates sea level, although in detail, it is distinct. Its equivalent in ﬂuvial systems (the equilibrium point between erosion and sediment accumulation) is referred to as the graded proﬁle.
Relative sea level is sea level deﬁned locally relative to a horizontal datum in a basin, such as the depth to basement and is the distance between this datum and base level. Hence, a rise in base level equates to a rise in relative sea level and vice versa. Importantly, this means it is insensitive to the eﬀects of sediment ﬁll on water depth. It is controlled by the combination of eustasy and tectonics, and it is the variations in relative sea level that are important in sequence stratigraphy.
Water depth is simply the distance from sea level to the basin ﬂoor. It is important insofar as it controls the depositional processes that ﬁll the accommodation space. Water depth is commonly confused with relative sea level.
Accommodation space is the physical space in which sediments can accumulate. It is the zone between base level and the sea ﬂoor. Initial accommodation space is generated either by subsidence or eustatic rise in sea level, although sediment compaction also plays an important role as thick packages of sediments accumulate. Accommodation space is lost either by tectonic uplift, eustatic drop in sea level, or inﬁlling of that space by sediments.
3.2.2 Transgression and regression Transgression occurs when a shoreline migrates landward as a result of a relatively rapid rise in base level. This happens when new accommodation space is produced more rapidly than sedimentation can keep up with it. It is marked everywhere by a shift to deeper water environments and gives rise to a shoreline and associated environments that retrograde.
Regression occurs when a shoreline and its laterally associated environments migrate bas- inward, resulting in progradation. Regression need not involve a drop in base level, as will be discussed below. Chapter 3: Sequence Stratigraphy 30
3.3 Shoreline trajectories and sediment stacking patterns
Sequence stratigraphy can be described in many diﬀerent ways, but ultimately it comes down to the sedimentary stacking patterns that result from the interplay of accommoda- tion space and sedimentation, which drive shifting shorelines. For any given place in time, one of ﬁve diﬀerent depositional (that is, stacking) or non-depositional scenarios is possible:
Aggradation is vertical stacking and occurs where a given facies stacks directly on top of itself–that is, with no lateral migration of facies belts. Aggradation requires both a perfect balance between accommodation and sedimentation (i.e., A = S), which is rare. Gives rise to an up-stepping stacking pattern.
Erosion is simply removal of underlying rock or sediment without the addition of new sediment (i.e., no new sediment supply).
Retrogradation occurs where there is net production of accommodation space (A+) and A>S. Retrogradation gives rise to a back-stepping stacking pattern.
Progradation occurs, as mentioned, during regression. Regression arises in two diﬀerent ways.
• Normal regression, where accommodation space is being produced (A+), but sedi- mentation exceeds it (S>A). This results in both vertical and basinward migration of facies. This gives rise to a fore-stepping stacking pattern.
• Forced regression, where accommodation space is lost (A-), no matter what S is. This gives rise to a down-stepping stacking pattern.
Because shorelines are inherently dynamic and pure aggradation is very rare (leastwise, in siliciclastic settings), there are eﬀectively just three stacking patterns that occur in coast marine systems: back-, fore-, and down-stepping. The geometry of the stacking patterns is fundamentally shaped by sedimentation, base level ﬂuctuation, and environmental energy. Consider a shoreface environment, which will have a low-dip, coastal plane, followed by a concave upward shoreline proﬁle shaped by wave energy. Although wave energy does not control which stacking pattern is formed, it does play an important role in controlling sediment distribution on the coast-line and weather net or erosion or deposition is taking place (Fig. 3.2.
3.3.1 Back-stepping stacking pattern As you now know, the back-stepping stacking pattern develops during transgression, when the rate of base level rise exceeds that of sediment accumulation. The coastal plain prism stacks backwards and upwards, resulting in coastal onlap onto the older subaerial uncon- formity surface. This production of accommodation space atop the coastal plain traps incoming sediments, starving the outer part of the basin. As transgression proceeds, an erosional surface known as a wave ravinement surface develops (Fig. 3.3). This is the result of the high energy of the beach system, which reworks existing sediments as it transgresses the continental margin. This may produce both a sharp surface of marine erosion, as well as a transgressive lag, where only the coarsest sediments escape being redirected by the Chapter 3: Sequence Stratigraphy 31
Figure 3.2: Depiction of the role of wave energy in shaping the shore face environment. Where wave energy and sedimentation are not balanced, the shoreline will migrate, either shoreward (wave energy dominates) or basinward (sedimentation dominates). From et al. (2009).
high energy shoreline; these accumulate in a distinct, coarse-grained deposit which marks the transgression.
3.3.2 Down-stepping stacking pattern Down-stepping occurs when there is a drop in base level, irrespective of the magnitude of supply of sediment. Here, the shoreline system both migrates basinward and drops along with base level. A subaerial unconformity develops, which migrates seaward so long as base level continues to fall. This produces downlap of stratal terminations and is the product of forced regression. Chapter 3: Sequence Stratigraphy 32
3.3.3 Fore-stepping stacking pattern Fore-stepping involves both aggradation and progradation due to base level rise and rel- atively high sediment accumulation rates. The coastal prism builds up, onlapping the underlying strata. The shore face wedge builds upward and outward in a series of curved clinoforms. The tangential toes to these clinoforms oﬄap the underlying surface. Fore- stepping is the result of normal regression.
Figure 3.3: The tree main types of stacking patterns resulted from normal regression, forced regression, and transgression. From et al. (2009).
3.4 Base-level changes and the development of systems tracts
During a given phase of normal regression, forced regression, or transgression, a suite of genetically related sediments are deposited from the onshore to oﬀshore environment. This package of sediments is known as a systems tract, and there are four types of systems tracts corresponding to the four phases in base level change: falling stage systems tract (forced regression), lowstand systems tract (normal regression), transgressive systems tract, (trans- gression) and highstand systems tract (regression).
3.4.1 Falling stage systems tract (FSST) Consider a simple scenario where a continental margin is subject to cyclical ﬂuctuations in sea level (say, a sine wave)(Fig. 3.4. Let’s start with the interval of falling base level (decreasing relative sea level), which removes accommodation space. This will result in the development of a fore-stepping stacking pattern of the the marginal marine environment. Chapter 3: Sequence Stratigraphy 33
Figure 3.4: Connection between base-level changes, sedimentation rates, and shoreline shifts. From Catuneanu (2006).
This prograding system is the product of forced regression—that is, progradation resulting from loss of accommodation space. The suite of sediments deposited during this interval of base level fall are known as the falling stage systems tract. They comprise the down- stepping shoreface sediments and those sediments deposited more distally at the same time, which may include basin fans comprising sediment eroded from the exposed proximal part of the basin and deposited by gravity ﬂow processes.
The extent of development of a FSST fan of deepwater coarse sediments as well as the exact pattern of the FSST will depend in part on the sediment supply rate. Where sediment supply is low, for example on a wave-dominated coastline (Fig. 3.5) and downward incision dominates over outward progradation, then a sharp marine erosional surface will develop that juxtaposes coarse, shore face deposits and older, ﬁner-grained oﬀshore deposits. At the same time, the courses sediments will be eroded and redeposited in the deeper sea environment. In contrast, where sediment supply rates are high, then the oﬄapping shore face systems will form gradational delta front deposits that converge down dip (Fig. 3.5).
3.4.2 Lowstand systems tract (LST) When the base level curve switches from dropping to rising, accommodation space is pro- duced again. This is the end of forced regression. But initially, assuming that rate of rise is low, as is the case with cyclically varying base level (Fig. 3.4), the rate of base level rise is not suﬃcient to keep up with sediment supply. The result is that progradation continues, but in this case, so does aggradation. That is, the shoreline fore-steps, building up and out, as a result of normal regression. Coastal plain sediments aggrade and onlap the coast while shoreface sediments build up and out. The end of this phase of normal regression, hence the upper bounding surface to the LST is the maximum regressive surface. Chapter 3: Sequence Stratigraphy 34
Figure 3.5: The diﬀerent oﬄapping stacking patterns that develop in high-sediment versus low-sediment supply margins in the falling stage systems tract. From Catuneanu (2006).
3.4.3 Transgressive systems tract (TST) Eventually, the rate of rise in base level may exceed the rate of sediment supply. At this point, the shoreline transgresses and the coastal marine prism retrogrades. The result is a systems tract that comprises a combination of aggradational and onlapping, coastal ma- rine sediments, backstepping shoreface system, and sediments eroded from the shoreline by wave energy and redistributed out to sea. Where suﬃcient sediment is eroded from the shoreline and redeposited in the deeper water environment on lap this erosional wave ravinement surface and are referred to as the healing phase edge. Otherwise, ﬂuvially de- rived sediments are largely trapped in the coastal environment and the deeper basin is starved, resulting in a zone of sediment starvation. The end of the transgression is known as the maximum ﬂooding surface and should be marked by the ﬁnest grained sediments.
3.4.4 Highstand systems tracts (HST) The highstand systems tract begins before the end of base level rise, but when the rate of rise drops below the sediment supply. The HST inherits high relative sea level from the transgression, but experience normal regression, with the sediments again building up and out. As clinoforms again develop, they prograde and downlap on the formerly sediment Chapter 3: Sequence Stratigraphy 35 starved seaﬂoor. Hence the HST is distinguished in seismic by the progradation of the shoreline/delta system with progradation of clinoforms that down lap onto the MFS.
3.4.5 The sequence cycle These systems tracts are by deﬁnition bound above and below by sequence boundaries.A full sequence will begin and end at the same point on the cyclical curve. A sequence may contain up to all four systems tracts, but need not have all of the system tracts. Consider cyclical sea level ﬂuctuations added to a long-term rise in base level due to tectonic sub- sidence. Depending the rate of subsidence and the periodicity and eustatic change, base level may not ever drop, such there are no FSST’s. Or, imagine a system with a a high sediment supply, such that transgression does not occur. In deep water systems, all four systems tracts may occur during a given sequence cycle, but may not be identiﬁable due to continuous ﬁne-grained sedimentation.
Importantly, a sequence cannot contain a single systems tract more than one. However, note that it can contain both LST and HST, which are normal regressive systems and have the same stacking pattern. Indeed, a single sequence may simply be an alternation between LST’s and HST’s, and hence comprise entirely fore-stepping stacking. Conveniently, there is way to distinguish between LST and HST based on the pattern of fore-stepping alone. That is because LST’s form during accelerating base level rise, which results in a concave up shoreline trajectory (Fig. 3.6); HST’s form during decelerating base level rise, resulting in a convex up shoreline trajectory (et al., 2009).
Figure 3.6: Forestepping stratal stacking patterns for two types of normal regression: low- stand systems tract and highstand systems tract. Note the slight diﬀerence in trajec- tory (convex up versus down), which can be used to distinguish between the two systems tracts.et al. (2009).
You can think of a systems tract as containing all of the sediments deposited during a particular phase of shoreline trajectory. Hence, these are time correlative units (facies) that represent laterally linked depositional environments. Chapter 3: Sequence Stratigraphy 36
3.5 Sequence boundaries
Sequence boundaries are stratigraphic surfaces that develop in response to changing weight of inﬂuence of subsidence (or uplift), eustatic sea-level change, and sediment supply rates. These eﬀects are manifested in shifts between transgression and regression of the shoreline, or a shift from forced to normal regression, or vice versa. That is, sequence boundaries separate separate systems tracts. You are already familiar with several of these surface, which include the subaerial unconformity, wave ravinement surface, and maximum ﬂooding and regressive surfaces.
Sequence stratigraphic surfaces are much better (but in most cases, not perfect) approxi- mations of time lines than lithological boundaries (that is, boundaries between facies), and for the most part, are recognizable in seismic sections (Fig. 3.7.4). Some of these surfaces develop throughout the course of a systems tract, while other are nearly isochronous. The seven sequence boundaries can be subdivided into those associated with base level fall and those associated with base level rise.
3.5.1 Sequence boundaries developed during base level fall Subaerial unconformity A subaerial unconformity (SU) obviously develops when a formerly deposited sediment is exposed to subaerial erosion or non-deposition. Within the sequence stratigraphic frame- work, these develop during the phase of falling relative sea level: the FSST. As relative sea level drops, the subaerial unconformity follows the down-stepping coastal sequence. The result is that the SU is both diachronous and variable in its duration. That is, basinward, the duration of exposure decreases. Although SU’s are logical and important stratigraphic boundaries insofar as the often separate genetically distinct packages of sediment, the variable hiatus and diachroneity of the SU must be born in mind when considering the time-evolution of those sediments.
Subearial unconformities are commonly associated with ﬂuvial incision, which might show up distinctly on seismic sections or in borehole data. Such incision can sometimes be traced out and mapped in detail in the ﬁeld (e.g. Christie-Blick et al., 1988a). Other times, the subearial unconformity may be distinguished by the development of a soil proﬁle (paleosol) or distinct trace fossils. They may also show up as a sharp juxtaposition of facies, in particular marine and non-marine. Often, however, it is not so easy to pick out because it may be simply a surface of non-deposition or deﬂation.
Correlative conformity The correlative conformity (CC) is the surface that extends underwater from the basinward extent of the subaerial unconformity. It is conformable and is includes the upper surface of the youngest clinoform developed during the FSST. The CC is particularly important because it links the SU to deeper water sediments. It is also a low diachroneity surface that corresponds to the end of base level fall (Fig. 3.4) and shows up as a distinct reﬂector in Chapter 3: Sequence Stratigraphy 37 seismic proﬁles. It is down lapped by prograding clinoforms within the LST. However, one problem with the CC is that it is diﬃcult to distinguish in outcrop and well logs. Hence, it has been the source of some controversy.
Basal surface of forced regression The basal surface of forced regression is a minor and relatively rarely well preserved, low diachroneity, conformable surface. It connects the SU proximally, and laterally is equivalent to the lower bounding surface of the oldest clinoform development during the FSST. Hence, in the marine environment, it separates the HST below from the FSST above. In systems with relatively high sediment supply (Fig. 3.5, it continues oﬀshore and converges withe SU. In low-sediment supply, erosive systems, where it is not eroded, it links the SU with the regressive surface of marine erosion, discussed below.
Regressive surface of marine erosion Where sediment supply is not signiﬁcant during a forced regression and the shoreface cuts down into underlying sediments as sea level falls (Fig. 3.5), a marine erosion surface de- velops. This regressive surface of marine erosion (RSME) juxtaposes ﬁner-grained shelf sediments of the underlying HST below from coarser-grained shore face sediments of the FSST above (Fig. 3.5). It forms the base of all of the down-stepping clinoforms of the FSST, and hence is a highly diachronous surface.
Further oﬀshore, the RSME manifests as a typically sharp (but not necessarily erosive) contact between ﬁne-grained pelagic sediments below and coarser gravity ﬂow sediments (usually turbidites) above that form the FSST basinal fan.
3.5.2 Sequence boundaries developed during base level rise Maximum regressive surface The maximum regressive surface (MRS) occurs at the end of regression and the onset of transgression, and so separates prograding from retrograding strata. This occurs during base level rise, and so is temporally oﬀset from the end of the FSST (the intervening time being the LST). It is a low diachroneity surface. It commonly shows up in stratigraphic proﬁles as the coarsest-grained sediments, and hence has a low GR signature in well logs. On continental slopes, it marks the boundary between the youngest prograding clinoform below and the healing phase deposits above.
Maximum ﬂooding surface The maximum ﬂooding surface (MFS), as its name implies, marks the end of transgression. Hence it forms the boundary between a the TST and the HST. It is a highly useful, low diachroneity, conformable surface. It is overlain by prograding strata, and so is a down lap surface that shows up in seismic data. It is perhaps the easiest boundary to distinguish in marginal marine systems because it marks a zone of sediment condensation and commonly shows up as the ﬁnest-grained sediments. Because of sediment starvation, the MFS may be associated with carbonates or concentrations of glauconite. Chapter 3: Sequence Stratigraphy 38
Wave ravinement surface During transgression, the high energy shoreline transgresses the continental margin, eroding that shoreline. This erosional surface, is known as a transgressive ravinement surface (TRS), or wave ravinement surface (WRS) where it is shaped wave erosion. it is sharp, unconformable, and diachronous. It may replace the maximum regression surface near the shift from prograding to retrograding, and depending on the extent of aggradation of the coastal plain sediments during transgression, it may completely remove the transgressive sediments beneath it and even the subaerial unconformity below. The wave ravinement surface shows up in well logs as shift to a ﬁning upward pattern.
3.6 Stratal terminations
Strata commonly terminate or appear to terminate against surfaces. These surfaces are commonly stratigraphic surfaces and the geometry of the surfaces relative to the termi- nating strata is informative in determining the nature of those surfaces and the strata terminating against them. The accumulation of sequences results in the development of distinct strata terminations (Catuneanu, 2006). Stratal terminations are easiest to iden- tify in seismic data, but even where they are not easy to see in well log data or outcrop, recognition of their ubiquity and origin is necessary for correct sequence stratigraphic in- terpretation of any stratigraphic data.
• truncation: termination of strata against an overlying erosional surface
• toplap: termination of clinoforms against an overlying lower angle surface mainly as a result of non-erosion (sediment bypass)
• onlap: termination of low-angle strata against a steeper surface
– marine onlap on continental slopes during transgression – coastal on lap onto revilement surfaces during transgression – ﬂuvial on lap is landward shift of upstream end of coastal plain during trans- gression or normal regression
• downlap: termination of inclined surfaces against a lower-angle surface
• oﬄap: progressive oﬀ-shore shift of up-dip terminations of sedimentary surfaces (dur- ing forced regressions)
3.7 Types of Sequences
Diﬀerent sequence boundaries can be used to deﬁne the upper and lower boundaries of an individual sequence. For example, in the original deﬁnition of depositional sequences, the boundaries are the subaerial unconformity and correlative conformity. In the Transgressive- Regressive (T-R) designation popularized by Alan Embrey (see, e.g., Embrey, 2009–2010), bounding surface is the maximum regressive surface (MRS). The point that et al. (2009) made is that these approaches are all compatible with one another, so long as the sys- tems tracts are consistent, and that certain approaches are more appropriate under certain Chapter 3: Sequence Stratigraphy 39
Figure 3.7: Types of stratal terminations. From Catuneanu (2006).
circumstances. For example, depositional sequences make sense in the realm of seismic interpretation, because all of the key boundaries are readily distinguished through a com- bination of stratal terminations. However, in outcrop, particularly where subaerial uncon- formities are subtle or not preserved, identifying correlative conformities and basal surfaces of forced regression can be diﬃcult if not impossible. In this case, T-R sequences may be more appropriate.
3.7.1 Depositional Sequence The depositional sequence was the ﬁrst type of sequence deﬁned when the Exxon guys ﬁrst published theory model for sequence stratigraphic applications in the 1970’s. In this model, the bounding surface to sequences is the subaerial unconformity or the equivalent correlative conformity. The designation of the lowstand systems tract and falling stage systems tract and the location of the CC unconformity relative the base level ﬂuctua- tions underwent some modiﬁcation as the theoretical framework for sequence stratigraphy evolved from the seismic stratigraphic framework. Namely, the question arose weather the correlative conformity should bound the forced-regressive package above or below. The general consensus now is to place it above, such that forced regressive deposits form a wedge between the basal surface of forced regression/regressive surface of marine erosion and the correlative conformity, but the overall gist of the approach remainder the same.
Depositional sequences are commonly applied both to seismic stratigraphy and outcrop (e.g. Christie-Blick et al., 1988b), but also can be applied to well logs. A key argument in favor of applying the depositional sequence model is that the sequences ﬁt the deﬁnition of between genetically related packages of strata (Catuneanu, 2006). .
3.7.2 Genetic Sequences An alternative to the depositional sequence was developed by Bill Galloway based on use of well log data (Galloway, 1989) and makes use of the maximum ﬂooding surface as the sequence bounding unit, but otherwise subdivides a sequence into the same systems tracts as a depositional sequence. Although this method can be seen as somewhat inconsistent with the idea of an individual sequence being entirely genetically related, it has a conceptual advantage insofar as the maximum ﬂooding surface marks the time of greatest shift in the shoreline. Importantly, the MFS is the most consistently identiﬁable stratigraphic surface Chapter 3: Sequence Stratigraphy 40 in many well logs in both continental and marine environments and can commonly be correlated with conﬁdence across a basin. Even where sequence stratigraphers do not adhere to the genetic sequence model, they often begin correlations with the MFS.
3.7.3 Transgressive-Regressive Sequences The transgressive-regressive (T-R) sequence approach employs a composite stratigraphic surface that comprises the subearial unconformity on the basin margin and the maximum regressive surface in the deeper water part of the basin. As originally deﬁned, the sequence itself is then separated into the transgressive (retrograding) systems tract (TRT) and re- gressive (prograding) systems tract (RST) (Embry and Johannessen, 1992). An advantage of the T-R approach is that these are readily identiﬁable in outcrop. At the same time, it preserves the use of the subearial unconformity as a sequence boundary. On the other hand, it also subsumes all prograding systems tracts into a single system tract, and in doing so, may conceal important information about base level ﬂuctuations. In many cases, this turns out not to be a problem, because falling stage systems tracts (FSST) do not occur or are diﬃcult to identify, in which case the RST is a composite between normal regressive LST and HST. Where there is no FSST, then the prograding portion of the systems tract can simply be thought of as the HST.
Importantly, at least in my opinion, one can retain the T-R approach while still discerning FSST, LST, and HST, where possible
3.7.4 Parasequences Stratigraphers often simply their sequence stratigraphic analyses to subdividing sections into parasequences, which are distinct, coarsening upward cycles, most commonly identi- ﬁed in the coastal marine environment (Catuneanu, 2006). Van Wagoner (1995) deﬁned a parasequences as ’a relatively conformable succession of genetically related beds or besets bounded by ﬂooding surfaces.’ A ﬂooding surface is an abrupt shift from sediments in- ferred to have been deposited in relatively shallow-water conditions to sediments deposited in deeper water conditions. Generally, ﬂooding surfaces are identiﬁed by a sharp decrease in grain size and are higher order elements of a larger-scale sequence.
Subdivision of successions into parasequences is a popular approach because these coars- ening upward packages are common in the sedimentary record, and it provides a relatively robust means of correlating strata that is not lithofacies dependent and often, if not al- ways, tied to ﬂuctuations in base level. However, one criticism of parasequences is that their boundaries, while recognizable in the ﬁeld, are not strict sequence boundaries. In- deed, a ﬂooding surface may correspond to any of a number of sequence boundaries, and hence it is not always easy to interpret. One way around this problem is to deﬁne higher or- der sequences as transgressive-regressive sequences (T-R sequences), where the boundaries maximum regressive surfaces (or, you could identify them based on the maxim ﬂooding surface).
Parasequences often form distinct stacking patterns of their own, where they thin or thicken upwards. Such thinning or thickening commonly points to a lower order cyclicity in water depth, as will be discussed further in a subsequent section. Chapter 3: Sequence Stratigraphy 41
Figure 3.8: The classic Exxon sequence stratigraphic model (from Vail, 1987). A. is a schematic dip proﬁle over a migrating coastline system (showing deposition as function of height and distance from the coastline), which shows key surfaces and depositional sequences, and B. is a Wheeler diagram, showing deposition as a function of time and space. That is, the Wheeler diagram demonstrates both where sedimentation is taking place and where non deposition or erosion is occurring through time. Chapter 3: Sequence Stratigraphy 42
3.8 Sequence Stratigraphy of Carbonate Platforms
Many of the concepts and rules developed for sequence stratigraphy in siliciclastic systems hold for carbonate and mixed carbonate-clastic systems. The diﬀerence arise because of the source of sediments; carbonate platforms produce their own sediment supply, and this production (the carbonate factory) is tied to water depth and the area of the platform.
Carbonate platforms may be either ramps and rimmed platforms. Let’s begin by consider- ing a rimmed system, where the outer edge of the platform is a reef system that has built up to sea level (Fig. 3.9-1), and the region behind it is a quiet region of shallow water. Gradual rises in sea level will typically be balanced by sediment precipitation and accu- mulation rates such that the platform itself my experience a high degree of aggradation. The edge of the platform will be a reef built of some consortium of reef-builders, including organisms such as corals, sea grasses, sponges, bryozoa, mollusks, and stromatolites. A sharp platform edge (slope) develops, and the escarpment can be many 10s of meters. De- bris sheds oﬀ the escarpment as talus or other sorts of gravity ﬂows (debris ﬂows, turbidity currents, grain ﬂows), such that calcareous sediments accumulate on the toe of the slope and margin of the deep basin, often interbedded with pelagic, basinal muds. This high- stand shedding delivers a large supply of sediment to the deeper basins, the result being pronounced progradation and overstepping of the reef system.
Because the platform is largely at sea level, even a small drop in base level can lead to exposure of the main platform and the shoreline system will drop onto the slope. This results in a signiﬁcant narrowing of the zone of carbonate production and accumulation and sediment starvation in the deep basin. The exposed carbonate platform is prone to karstiﬁcation and meteoric diagenesis. The FSST here should show up distinctly in the juxtaposition of deep water sediments and shallow water carbonates, such as cross-bedded oolites and interclast breccias.
A subsequent graduate rise in sea level permits the entire platform to resume carbonate productivity, and the bounding reef system may reestablish itself and highstand shedding will resume as soon as the transgression ends.
An abrupt rise in sea level can have a profound eﬀect on a carbonate platform. If suﬃ- ciently rapid that carbonate precipitation cannot keep up, the development of deep water conditions on the platform drastically decrease or shuts oﬀ the carbonate factory. In this case, a drowning unconformity develops, which maybe be marked as either a hardground (that is, carbonate sediment that is cemented on the seaﬂoor), or simply as a cessation of carbonate precipitation and shift to mud deposition, which must accumulate very slowly unless the rise in base level also corresponds to an increased ﬂux of clastic sediments. At the same time, lacking the ongoing source of encrusting organisms and buttress of prograd- ing clinoforms, the platform edge may destabilize and erode subaqueously.
Following the rise in sea level, ﬁne-grained clastic sediments may give way to a deltaic sys- tem, if there is a source of epiclastic systems. Or it may transition upward into increasingly carbonate-rich sediments as the carbonate factory reestablishes itself. Once this happens, highstand progradation may proceed very rapidly and reef systems may reestablish. Chapter 3: Sequence Stratigraphy 43
Figure 3.9: Illustration of generalized systems tract styles developed during variable phases of base level ﬂuctuation on a rimmed carbonate platform. From (Catuneanu, 2006). Chapter 4
Learning to Tell Geological Time
Time is a key dimension in geology, particularly in deciphering Earth’s history. Of course, ages are important to develop the proper sequential order of events. But time is also crit- ically important for deterring processes. For example, in determining the cause a major carbon isotope anomaly in the oceans, it is critical to know if that anomaly occurs and recovers on time scales of thousands versus hundreds of thousands of years. Time is also important for integrating information. To the extent that we can correlate time-equivalent geological units from across sedimentary basins and across the globes, we can produce snapshots of paleogeographies, the global environment, and the biosphere at speciﬁc times in the past. However, telling time in the geological record is challenging. Much of what ﬁeld geologists do is sort out the relative timing of events. If the basic principles that we rely upon to do so are not so complicated, the actual process of doing so can be, because it involves mapping out geological relationships, which are often complex or poorly exposed on the Earth’s surface.
Given a robust ordering of geological events (roughly, chronostratigraphy), then we can ap- ply radiometric and other techniques to date events. This is essentially what the geological time scale is—a partially dated global stratigraphic record. Once a speciﬁc stratigraphic boundary has been dated, then dates can be applied indirectly via correlation of that boundary. For example, the Precambrian-Cambrian boundary is rather precisely dated at 541 Ma, and so if that boundary can be identiﬁed reliably in any succession of rocks, then a date of 541 Ma can be appled reliably. Unfortunately, the quality and quantity of dates is not evenly spread across the Geological Time Scale and so some intervals of geological time are more diﬃcult to correlate and approximate an age for. This is particularly acute for the Precambrian, but also an issue for intervals between dated stratigraphic boundaries.
Often, we need or at least want to present stratigraphic data on a temporal scale. This involves developing a model for transforming stratigraphic height, depth, or thickness into time. Again, this is no simple process, but it is one for which there are a variety of tools that can be applied. Indeed, a large proportion of stratigraphic research is applied for the purpose of developing time scales. These time scales are then used to interpret strati- graphic data, such as diversiﬁcation rates and patterns, shifts in the chemical composition of seawater as recorded in limestones or other marine precipitates, and rates of continental
44 Chapter 4: Geological Time 45
INTERNATIONAL CHRONOSTRATIGRAPHIC CHART www.stratigraphy.org International Commission on Stratigraphy v 2013/01
numerical numerical numerical Eonothem numerical
Series / Epoch Stage / Age Series / Epoch Stage / Age Series / Epoch Stage / Age
GSSP GSSP EonothemErathem / Eon System / Era / Period age (Ma) EonothemErathem / Eon System/ Era / Period age (Ma) EonothemErathem / Eon System/ Era / Period age (Ma) / Eon Erathem / Era System / Period GSSA age (Ma) present ~ 145.0 358.9 ± 0.4 ~ 541.0 ±1.0 Holocene Ediacaran 0.0117 Tithonian Upper 152.1 ±0.9 Famennian ~ 635 0.126 Upper Kimmeridgian Neo- Cryogenian Middle 157.3 ±1.0 Upper proterozoic Pleistocene 0.781 372.2 ±1.6 850 Calabrian Oxfordian Tonian 1.806 163.5 ±1.0 Frasnian Callovian 1000 Quaternary Gelasian 166.1 ±1.2 2.588 Bathonian 382.7 ±1.6 Stenian Middle 168.3 ±1.3 Piacenzian 3.600 Bajocian 170.3 ±1.4 Givetian 1200 Pliocene Middle 387.7 ±0.8 Meso- Zanclean Aalenian proterozoic Ectasian 5.333 174.1 ±1.0 Eifelian 1400 Messinian Jurassic 393.3 ±1.2
7.246 Toarcian Devonian Calymmian Tortonian 182.7 ±0.7 Emsian 1600 11.62 Pliensbachian Statherian Serravallian Lower 407.6 ±2.6 13.82 190.8 ±1.0 Lower 1800 Miocene Pragian 410.8 ±2.8 Proterozoic Neogene Langhian Sinemurian Orosirian 15.97 199.3 ±0.3 Lochkovian Paleo- 2050 Burdigalian Hettangian 201.3 ±0.2 419.2 ±3.2 proterozoic 20.44 Mesozoic Rhaetian Pridoli Rhyacian Aquitanian 423.0 ±2.3 23.03 ~ 208.5 Ludfordian 2300 Cenozoic Chattian Ludlow 425.6 ±0.9 Siderian 28.1 Gorstian Upper Norian 427.4 ±0.5 2500 Oligocene Precambrian Rupelian Wenlock Homerian 430.5 ±0.7 Neo- 33.9 ~ 227 Sheinwoodian 433.4 ±0.8 archean
Priabonian Carnian Silurian Telychian 2800 38.0 riassic ~ 237 Llandovery 438.5 ±1.1 Meso- Bartonian T 41.3 Ladinian Aeronian 440.8 ±1.2 archean Eocene Middle ~ 242 Rhuddanian Lutetian Anisian 443.4 ±1.5 3200 47.8 247.2 Hirnantian 445.2 ±1.4 Paleo-
Olenekian Archean Ypresian Lower 251.2 archean Paleogene Induan 252.17 ±0.06 Upper Katian 56.0 Changhsingian 453.0 ±0.7 3600 Thanetian 254.14 ±0.07 59.2 Lopingian Sandbian Eo- Wuchiapingian Paleozoic Paleocene Selandian 259.8 ±0.4 458.4 ±0.9 Phanerozoic 61.6 Phanerozoic Phanerozoic archean Capitanian Darriwilian 4000 Danian 265.1 ±0.4 Middle 66.0 467.3 ±1.1 Guadalupian Wordian Dapingian Hadean Maastrichtian 268.8 ±0.5 470.0 ±1.4 Roadian Ordovician ~ 4600 72.1 ±0.2 272.3 ±0.5 Floian Campanian Kungurian Lower 477.7 ±1.4 Units of all ranks are in the process of being defined by Global Boundary Stratotype Section and Points (GSSP) for their lower 83.6 ±0.2 Permian 283.5 ±0.6 Tremadocian Santonian 485.4 ±1.9 boundaries, including those of the Archean and Proterozoic, long Upper 86.3 ±0.5 Artinskian defined by Global Standard Stratigraphic Ages (GSSA). Charts and Cisuralian 290.1 ±0.26 Stage 10 Coniacian ~ 489.5 detailed information on ratified GSSPs are available at the website 89.8 ±0.3 Sakmarian http://www.stratigraphy.org. The URL to this chart is found below. 295.0 ±0.18 Furongian Jiangshanian Turonian ~ 494 Asselian Numerical ages are subject to revision and do not define units in 93.9 298.9 ±0.15 Paibian ~ 497 Cenomanian Gzhelian Guzhangian the Phanerozoic and the Ediacaran; only GSSPs do. For boundaries 100.5 Upper 303.7 ±0.1 ~ 500.5 in the Phanerozoic without ratified GSSPs or without constrained Kasimovian numerical ages, an approximate numerical age (~) is provided.
Paleozoic 307.0 ±0.1
aceous Series 3 Drumian Albian ~ 504.5 Middle Moscovian Numerical ages for all systems except Permian,Triassic, Cretaceous 315.2 ±0.2 Stage 5 Mesozoic ~ 113.0 and Precambrian are taken from ‘A Geologic Time Scale 2012’ by Cre t ~ 509 Lower Bashkirian Gradstein et al. (2012); those for the Permian, Triassic and
Aptian Pennsylvanian Stage 4 323.2 ±0.4 Cretaceous were provided by the relevant ICS subcommissions. ~ 125.0 Series 2 ~ 514 Upper Serpukhovian Cambrian Stage 3 Barremian 330.9 ±0.2 Lower ~ 129.4 ~ 521 Hauterivian Coloring follows the Commission for the ~ 132.9 Middle Visean Stage 2 Geological Map of the World. http://www.ccgm.org Carboniferous ~ 529 Valanginian 346.7 ±0.4 Terreneuvian Chart drafted by K.M. Cohen, S. Finney, P.L. Gibbard ~ 139.8 Fortunian (c) International Commission on Stratigraphy, January 2013 Berriasian Mississippian Lower Tournaisian ~ 145.0 358.9 ±0.4 541.0 ±1.0 http://www.stratigraphy.org/ICSchart/ChronostratChart2013-01.pdf
Figure 4.1: The most recent Geological Time Scale (Gradstein et al., 2012) from the International Commission on Stratigraphy, showing locations of GSSP’s.
4.2 The Geological Time Scale
The Geological Time Scale might seem like a static and boring construction, pieced to- gether by geologists and palaeontologists with nothing better to do with their time. But this is far from the truth. Rather, the time scale is an evolving framework through which scientists decode the history of our planet Gradstein (2012). It represents the distillation of all of the geological knowledge of Earth’s history that has been acquired over the past two centuries of geological scholarship. So while it can be boiled down to a chart with numbers and names of eras, periods, and stages that is invaluable in its own right is a prop to teach geology, it is also a complex and dynamic construct by which we map Earth’s history onto a linear, temporal axis. Earth’s history cannot be interpreted without time.
One reason that the Geological Time Scale is in fact so complex is that it represents the join between the actual rock record, with its constituent stratigraphic divisions, and tech- Chapter 4: Geological Time 46 niques for measuring geological time. The former is referred to as the chronostratigraphic scale, and the latter the chronometric scale (Gradstein, 2012). The units and divisions of the chronostratigraphic scale are relative and are a function of the rock record: they reﬂect the vagaries of Earth history. These subdivisions include the formal, standardized units, such as the ‘Cambrian’ and ‘Neogene’, polarity chrons (based on reversals of Earth’s magnetic ﬁeld), and biozones established by fossils. The original, international geological time scale was developed in the late 1800’s based simply on resolving the relative order of chronostratigraphic units deﬁned in various parts of the world (but mostly in Europe).
The chronometric time scale is time axis and is heavily dependent on radiometric dating techniques for establishing actual ages for events and chronostratigraphic boundaries. The following chapter will provide a more detailed description of they key radiometric tech- niques applied to dating the stratigraphic record. Another chronometric tool, which is particularly powerful for dating and calibrating the geological record over the past several millions of years is the astronomical scale. Based on the realization that Earth’s climate is inﬂuenced by systematic and predictable orbital cycles and other astronomical forcing, i.e. the Milankovitch cycles, and the observation that this cyclicity leaves an imprint in the geological record, scientists have been able to ‘tune’ the geological time scale based on these cycles Shackleton (e.g. 2000). This chronometric technique will also be described in a separate chapter devoted to sedimentary cycles.
The combination of these techniques was pioneered by a young Arthur Holmes, who used a combination of 5 Phanerozoic U-He ages from uraniferous minerals and maximum thick- ness for each geological period to extrapolate ages for each period boundary (1913, 2nd edition in 1937). This approach was based on the assumption of similar and constant sediment accumulation rates for each of the period, which while unrealistic, nevertheless produced an age of 600 million years for the age of the Cambrian (Gradstein, 2012). In fact, geologists continue to apply the same technique to develop age models for successions with limited age constraints, only over much more limited durations/thicknesses.
4.2.1 Stratotypes A stratotype is a type section of strata which is referenced to the Geological Time Scale. Unit stratotypes are type sections that span a particular rock-time interval—typically a stage—which is used as the reference point for deﬁning and characterizing a particular rock unit. The Geological Time Scale used to be largely subdivided based on such unit stratotypes. This approach led to many problems, foremost of which was the fact that many intervals of the Geological Time Scale were left unrepresented, while other strato- types partially or wholly overlapped. Even worse, some stratotypes turned out to be facies equivalents of previously deﬁned stratotypes rather chronostratigraphically distinct units. This rather inverted approach to building the Geological Time Scale led to signiﬁcant con- fusion and conﬂict.
Now the International Commission on Stratigraphy advocates the approach of deﬁning boundary stratotypes, which are type section that span a speciﬁc boundary and provide a template for identifying that boundary globally. Standardization in how we deﬁne bound- ary stratotypes has resulted in development of Global Stratotype Sections and Point, or Chapter 4: Geological Time 47
Figure 4.2: Illustration of the formation of the Geological Time Scale, as produced via the synthesis of chronometric and chronostratigraphic scales. From Gradstein (2012).
GSSP, which represent the formal boundary between two chronostratigraphic intervals in the Geological Time Scale. This points (the so-called ‘golden spikes’) are placed in well exposed and represented, usually reasonably accessible, stratigraphic sections which repre- sent continuous sedimentation across the boundary. The choice of a GSSP is not trivial and takes into account the ensemble of features, based mainly on fossils but also other unique stratigraphic features, that allow the precise and consistent deﬁnition of that boundary. The selection of a GSSP involves a committee of individuals that is familiar with that interval of time and usually requires several years and many meetings and ﬁeld trips.
4.3 Telling Time in the Stratigraphic Record
We have already established some of the fundamental tools for determining time and age relationships in the geological record, beginning with relative time and the application of Steno’s Laws. Radiometric dating techniques, biostratigraphy, magnetostratigraphy, astrochronology, and diverse chemical stratigraphic techniques will be discussed in subse- quent chapters. Beyond these key and accepted techniques, there are other methods for telling time in the geological record. Chapter 4: Geological Time 48
4.3.1 Layer counting One rather straightforward way of producing a chronology is to count layers that were deposited at a known interval. This requires an important assumption that the period of said interval is known with certain, but this can be done in some cases. Ice layers in glaciers and tree rings are two obvious techniques that are used, although these are not applicable to the stratigraphic record. In sediments, lake varves can also be counted and reliably yield chronologies spanning hundreds to thousands of years in some cases. Tidal rhythmites can also be counted. These are not common in the sedimentary record, but the day occur. Where they do occur, then often only provide a limited record, but this record can be used to determine the length of days and years in the past (e.g Williams, 1998).
Layer counting is not entirely diﬀerent from the technique of maximum sediment thickness employed by Arthur Holmes. Various other geologists have employed similar methods for determining the time spans of certain intervals of rocks. For example, Charles Walcott estimated the age of the Paleozoic to be 80 m.y. based on stratigraphic thicknesses and sedimentation rates determined in modern environments. However, as soon as one begins to deal with strata that were not deposited by some periodic process, the reality that sediments accumulate in ﬁts and pulses becomes a major problem.
4.3.2 Cosmic ﬂuxes Cosmic dust is constantly being loaded on the atmosphere and a fraction of it subsequently settles through the water column of the ocean and is incorporated into sediments. Cosmo- genic dust contains a relatively high concentration certain elements, namely the platinum group elements (PGEs), that are in very low abundance on Earth’s surface. If it is assumed that all of or some known fraction of one of the PGE’s in a sediment is derived from cosmic dust, then variations in the concentration in the sediment reﬂect variable sedimentation rates. Conversely, these concentrations can be used to determine a timescale for deposition of these sediments.
When Walter Alvarez discovered the Ir spike at the K-Pg boundary in Gubbio, Italy, this is exactly what he was doing. That is, he was measuring Ir concentrations as a proxy for sedimentation rate, and hence a time scale he could use to determine the rates of continent drift of Italy spanning the K-Pg boundary. The spike in Ir did not reﬂect a major condensation of section, but rather a sudden loading of cosmic Ir into the oceans from the collision of a meteorite.
4.3.3 Subsidence modelling One technique that is employed widely by the petroleum industry and in the study of deep sea cores is subsidence analyses. This technique is based on the premise that subsidence, and hence net sediment accumulation, in many basins is governed over long time scales by the mechanism of subsidence. Thermal (Fig. 4.3) and to a lesser extent, ﬂexural subsi- dence can be modelled quite reliably, though with many simplifying assumptions. Thermal subsidence curves are based on the solution to the heat ﬂow equation, and simply reﬂex the cooling of the lithosphere after extensional thinning, and the resulting isostatic adjustment. Chapter 4: Geological Time 49
With a few ages for an interval of sediments, along with with some geological constraints on the mechanisms of subsidence and corrections for water depth, sediment compaction, and isostatic adjustment, a model curve can be developed for the generation of accommodation space. This can then be used to convert sediment height or depth into time or age.
Figure 4.3: Thermal subsidence curves generated using code from Allen and Allen (2005) for subsidence resulting from cooling of the lithosphere following instantaneous 1D thinning (extension).
4.3.4 Optically Stimulated Luminescence dating Optically stimulated luminescence (OSL) is one type of luminescence dating. Luminescence results from the gradual accumulation of radiation damage to mineral grains or other ma- terials that accumulate over time due to bombardment by ionizing radiation emitted from radioisotopes. The extent of the damage, which is manifested in part by the formation of structurally unstable electron traps within the material, which when stimulated by the appropriate wavelength of light (blue, green, or infrared), luminesces as the stored unstable energy from the electron traps is released.
The total radiation damage accumulates over time as a function of the amount of ionizing radiation to which it is exposed (the dose). OSL is applied speciﬁcally to mineral grains, most commonly quartz. The reason OSL works as a technique for dating sediments is that exposure to sunlight resets the luminescence signal. So by measuring the luminescence of Chapter 4: Geological Time 50 a grain, as well as the dose of radiation to which it was exposed, the time since that grain was exposed to sunlight can be measured. In eﬀect, this date is the age of the sediment. In detail, the age is requires ﬁrst determining ﬁrst the equivalent dose, which is the dose of radiation required to produce the luminescence signal in the sample. This is because for any given sample, the OSL that results from the natural dose will vary depending on properties of the material or mineral. The equivalent dose can be determined by measuring the OSL response in splits of natural sample that have been exposed to varying amounts of radiation in the lab, such that extrapolation back to 0 radiation gives the natural dose.
Given the equivalent dose and the dose rate (which can simply be measured from the sample), the age is determined as:
Equivalent − dose Age = (4.1) Dose − rate This technique is generally applied to quartz and feldspar minerals. Quartz has a somewhat lower saturation dose, which limits the upper range of dating to about 100 kyr. Felspar, on the other hand, has a higher saturation dose and allows dating on materials up to 200–300 kyr (Zoeller and Wagner, 2014).
OSL dating has been applied in a variety of depositional environments. Aeolian and ﬂuvial environments are logical targets for OSL, but coastal marine and colluvial sediments have also been targeted. Other similar dating techniques include thermal luminescence (TL), infrared luminescence (IRSL), and radioﬂuorescence dating. Chapter 5
Radiometric dating techniques are quintessential to geology, and of course, absolutely nec- essary in the calibration of the Geological Time Scale. Arthur Holmes initiated the ap- plication of radiogenic isotopes to dating geological materials, and even though the initial results he obtained may seem wildly oﬀ by today’s standards, they fundamentally changed geology. Radiometric dating began to take oﬀ in the 1960’s and continues to be a rapidly evolving ﬁeld, with important progress being made analytically, in particular in the devel- opment of extremely well calibrated and widely distributed isotopic tracers (required for most radiometric dating methods), improved instrumentation, and improved blanks as the sources of contamination during the analytical process are weeded out.
Some techniques that were more popular and widely applied previously, namely Rb-Sr dating of diagenetic minerals, have gone out of fashion, while other methods, such as Re- Os have come to the forefront. There are a widely variety of isotope systems that can be applied to dating the sedimentary record and a full course could easily be given covering these. Here we will focus on a small handful of isotope systems that are particularly important for dating sedimentary rocks at both long (>107 years) and short (<106 years) time scales. These include the U-Th-Pb, K-Ar, Re-Os, and radiocarbon systems. Out of these systems, the U-Pb system applied to zircon crystals and the K-Ar (40Ar-39Ar) method applied to sanidine felspar are the preferred geochronometers for high-precision dating (Schmitz and Kuiper, 2013).
Table 5.1: List of decay constants used in the most recent version of the Geological Time Scale (GTS2012, Gradstein et al., 2012), as reported in (Schmitz, 2013).
Radio-isotope Decay constant in years−1) 238U (to 206Pb) 1.55125×10−10 ± 0.00016 235U (to 207Pb) 9.8571×10−10 ± 0.00012 40K (total) 5.463×10−10± 0.0107 187Re 1.6689×10−11 ± 0.000031 14C 1.2908×10−4
51 Chapter 5: Radiometric Techniques 52
5.2 U-Th-Pb System
The uranium-thorium-lead decay scheme is the most powerful radiogenic isotopic system available for dating geological materials. Many diﬀerent dating techniques emerge from this system, owing in part to the paired 238U and 235U decay pathways and to the large number of intermediate radioactive isotopes between U and Pb that have diverse chemical properties and half-lives. Of course, only some of these techniques are directly applicable to dating the sedimentary record, and it those systems that are most relevant that will be discussed in this chapter. The gold standard of all radiometric dating techniques is the U-Pb technique applied to zircons (and other similar minerals). This method is most powerful where it is applied to primary populations of zircons (versus inherited or reworked zircons), which are formed mainly in igneous systems. Both primary zircons and secondary growths on older zircons can also form during metamorphism, allowing dating of metamor- phic events. Fortunately for stratigraphers, volcanic beds are strata and dateable volcanic horizons—mainly felsic tuﬀs and ﬂows—occur in the stratigraphic record. It is the dating of such volcanic horizons that is largely responsible for the calibration of the Geological Time Scale beyond the Cenozoic Era (Gradstein et al., 2012).
5.2.1 U-Pb zircon dating U-Pb zircon data is especially powerful for several reasons. First, the dual decay schemes of U allow dates to be calculated based on the 206P b/238U, 207P b/235U, and 207U/206Pb ratios. This allows geochronologists to verify the assumption of closed system behavior of the zircon (or other mineral that is dated), since the ages should all be concordant (that is, consistent with one another). In this case, the age of the zircon should be dateable by either the 235U or 238U decay scheme:
206P b∗ = eλ238t − 1 (5.1) 238U and 207P b∗ = eλ235t − 1 (5.2) 235U where * denotes radiogenic lead.
For most minerals, this assumption does not easily hold up. Hence a second reason that zircons are such coveted minerals for geochronology is that they are unusually robust, resis- tant to both physical and chemical degradation under many conditions. Additionally, they are rather ubiquitous minerals. Limitations in the radiometric dates that can be produced on zircons derive from complicated zircon populations, blank levels, and quality of isotopic tracers and standards. Zircons that crystallize from U-rich magmas may also incorporate so much U that they are prone to radiation damage.
Zircons are dated via two very diﬀerent approaches. The classical approach and that which is still best suited for high precision analyses is via isotope-dilution and solid source mass spectrometry. However, in situ dating techniques via Sensitive High Resolution Ion Microprobe (SHRIMP) (as well as regular ion microprobe) analysis and laser ablation (LA) ICP-MS have become increasingly popular and are particularly powerful for dealing with complex or detrital zircon populations. Chapter 5: Radiometric Techniques 53
Figure 5.1: The U-Pb concordia diagram shows the evolution of U-Pb ratios through time. A discordia is a line deﬁned by an array of zircons that fall oﬀ of the concordia, but share the same crystallization age and disturbance age.
CA-TIMS The standard but time-consuming technique for analyzing zircons involves dissolving zircon grains (where possible individual grains) and analyzing U, Th, and Pb ratios by thermal ionization mass spectrometry (TIMS). Samples are ﬁrst spiked with an isotopic tracer, such that the actual ratios measured on the mass spectrometer reﬂect a carefully controlled mix- ture between the tracer and the natural sample (isotope dilution; Schmitz, 2013). In the case of U-Pb analysis, the synthetic isotopes 205Pb and 233U are commonly used. Isotope dilution eliminates eﬀects from fractionation of the ratio of daughter and parent isotopes during the dissolution and chemical separation procedure, since both the natural and the tracer parent-daughter ratios should fractionate equivalently (Schoene et al., 2013). Fol- lowing chromatographic separation, by which U and Pb are separated from their matrixes and each other, the samples are ready for measurement by TIMS. From the measured isotope ratios, the original isotopic abundances can then be back-calculated based on the amount of tracer added and the measured ratios.
Despite being ideal crystals, zircons do have ﬂaws, which may include metamorphic rims and damaged zones resulting from radiation or geological abuses, such as reworking and metamorphism. Consequently, simple dissolving zircons and analyzing them may yield in- accurate ages or signiﬁcant scatter (i.e. poor precision). Up until 10 years ago, the preferred method for cleaning up zircons was via a physical air abrasion technique—basically wearing down the edges and week zones of zircons much like one would smooth a rough stone or mineral—to remove potentially miscreant components of the zircon (Krogh, 1982). Some zircons would simply fall apart during abrasion, and these would be those poorly suited for analysis anyway. Chapter 5: Radiometric Techniques 54
Although this approach signiﬁcantly reduced the external error in dating zircons, it turned out not to be suﬃcient for resolving certain ages at the level of precision required by geologists. The classic example is in the dating of the Permo-Triassic boundary, where ages produced by diﬀerent labs following the same technique (and, 1998; Mundil et al., 2001) actually varied beyond the quoted uncertainties of the measurements (Schmitz and Kuiper, 2013). A signiﬁcant breakthrough in U-Pb zircon dating via TIMS was made by Jim Mattinson 2005 of UCSB who applied a chemical abrasion technique after ﬁrst anneal- ing (heating up the zircons to 800–900◦C), which migrates lattice defects out low of low-U domains. Subsequently, the zircons are subjected to sequential leaching in hydroﬂuoric acid, which preferentially dissolves the damaged, high-U zones of the zircon the are the most likely suﬀered lead loss.
Another huge advance in U-Pb zircon geochronology came in the development of a common U-Pb tracer that was developed through the Earthtime initiative. This new tracer was calibrated gravimetrically, made from certiﬁed NIST standards, and distributed to labs globally. Not only did this initiate generate the highest quality isotope tracer possible, it levelled the playing ﬁeld insofar as eliminated inter-laboratory error resulting from the use of diﬀerent standards. The results have been staggering; Paleozoic zircons populations can now be dated with a precision of 0.1 m.y., are even slightly better in some cases.
In situ techniques Development of the SHRIMP and subsequent application of coupled laser ablation-mass spectrometry revolutionized U-Pb geochronology. The distinct advantage of these tech- niques is that they permit rapid, high spatial resolution analysis. This, of course, is ap- pealing on many levels. It allows geochronologists to deal with complexly zone zircons, in particular those with inherited cores, hence extracting more and often more accurate in- formation than could have been ascertained from the ID-TIMS techniques. It also permits dating a lot more zircons than is possible by ID-TIMS, which requires timing-consuming, meticulous clean lab geochemistry to achieve suitable results. This capability has opened up new applications of U-Pb zircon geochronology, as well as the application of in-situ dating techniques to a variety of other minerals, such as monazite, titanite, and baddel- lyite. LA-ICP-MS also allows measuring a whole series of other elements at the same as the U-Pb ratios, which is a powerful way to characterize and discriminate zircons from diﬀerent magmatic sources.
A drawback of in situ techniques is that they do not achieve the high level of precision possible with ID-TIMS; the precision on individual populations is typically at least an order of magnitude better by ID-TIMS than by SHRIMP, and the spread is even greater between ID-TIMS and LA-ICP-MS. Another shortcoming is that LA-ICP-MS is not able to assess the contribution of common lead (that is, lead that is present in the zircon or other mineral that is not derived from U-decay). With ID-TIMS, common lead is measured by monitoring 204Pb during measurement of the other Pb isotopes. However, because 204Hg interferes with 204Pb, a correction must be made for this interference, which is done by measuring 202Hg simultaneously. With LA-ICP-MS, however, 202Hg cannot be reliably measured and so this common Pb contributions to the 207Pb and 206Pb signal cannot be reliably assessed. However, this problem is overcome by measuring the 204 signal and Chapter 5: Radiometric Techniques 55 simply omitting any samples that have a mass 204 component (be it Hg or Pb).
So each technique has its advantages and disadvantages. Clearly, certain techniques are more suitable than the others depending on the problem. High-precision dating for calibrat- ing the time scale or rates of important geological processes is best achieved via ID-TIMS. Detrital zircon geochronology (see below), dating of complex populations or geological terranes, and screening of zircons is commonly best accomplished via SHRIMP or LA- ICP-MS. An important advance for U-Pb geochronology has been the coupling of in situ and ID-TIMS analyses, where by zircon populations can ﬁrst be screened via SHRIMP or LA-ICP-MS to survey diﬀerent zircon populations and assess the degree of heterogeneity both within and between populations.
5.2.2 Detrital zircon analysis The ability to analyze large numbers of zircons in a single analytical session has eﬀectively forged a new sub-discipline within geochronology: detrital zircon analysis. This involves analyzing the zircons present in a sandstone or other epiclastic sediment or rock, usually by SHRIMP or LA-ICP-MS. In this was, age spectra can be generated that provide a ﬁrst-order and often highly valuable window into the source terranes of those sediments. Detrital zircon data are now applied to a broad range of problems, including provenance studies, tectonics (because major tectonic events commonly result in a change in source area), and correlation of formerly conjugate or otherwise linked sedimentary basins, with obvious applications to paleogeography and plate tectonic reconstructions. Unfortunately, the data that can be generated or so tantalizing for the relative ease with which they can be produced and the information that they yield that users of detrital zircon data sometimes overlook both the geological constraints that can and should be used to extract the most possible information from those data, and the limitations on those data imposed by the analytical technique.
A sometimes useful application of detrital zircon geochronology is the use of the youngest obtained zircons within a sediment as a maximum age constraint on those sediments (think principle of inclusion).
5.2.3 U-Th disequilibrium Uranium-thorium disequilibrium dating is distinct from many radiometric dating tech- niques insofar as it does not calculate an age based on the abundance of parent and daugh- ter isotopes. This is because thorium is itself radioactive and so decays. The U-Th dating technique instead measures the degree to which 234U (a product of the 238U decay chain) and its daughter product, 230U, have attained secular equilibrium. This technique is ap- plied to carbonates precipitated from seawater or other natural waters and works because thorium is highly insoluble under Earth surface (or near surface) conditions. As a conse- quence, U, but not Th is incorporated into the carbonate crystal lattice.
As 230Th begins to accumulate via the decay of 234U (half-life = 2.455 ×105 years), some of the 230Th then decays to 226Ra (half-life = 7.538 ×104 years). Initially, the system is out of equilibrium, such that 230Th gradually increase. Secular equilibrium is reached when the input of 230Th from decay of 234U is matched by output of 230Th through decay to Chapter 5: Radiometric Techniques 56
226Rd. Hence an age can be measured based on the abundance of 234U and 230Th.
For samples that contain U and negligible amounts of Th (some Th will be present in dust and other insoluble residue, as well as the analytical blank), ages can be determined using the equation (Zhao et al., 2009):
230T h 234U λ −λ230t 230 λ234−λ230t 1 − 238 = e − 238 − 1 × (1 − e ) (5.3) U U λ230 − λ234 where t is the age of the sample. Notice that 238U appears in the equation; this is because it is also a source of 238. This validity of this equation is dependent on assumptions:
• The system remains closed with respect to the U-Th system from the time the car- bonate mineral formed
• The dated material contains no initial 230Th
However, a correction for initial 230Th can be made by measuring the 230Th/232Th of the sample and assuming an initial 230Th/232Th ratio. U-Th ratios are commonly measured by thermal ionization mass spectrometry (TIMS), but multi-collector inductively coupled plasma mass spectrometry (MC-ICP-MS) is being used increasingly.
5.2.4 210Pb and 137Cs dating Lead-210 (210Pb) is a natural, radioactive decay product of in the 238U decay series. Un- supported 210Pb (that is, where it has been separated physically from its U-Th source) can be used to date sediments that are up to a few hundred years old. The basis for this technique is the atmospheric ﬂux of radiogenic gas 222Rn from the upper crust, where it is directed from the radioactive decay of 226Ra. The 222Rn in turn decays to 210Pb, which then settles from the atmosphere. If constant 210Pb deposition rates are assumed, then by measuring the 210Pb activity in sediments, one can determine sedimentation rates, and hence a time scale for those sediments. The decay rate of 210Pb to 206Pb is about 22 years, so this technique is only applicable for sediments dating back a few hundred years. In an ideal situation where sedimentation rates and Pb-loading rates are constant, then a curve of 210Pb with depth will decrease logarithmically.
There are a few methods of applying this technique, but the most common is the CRS model (Appleby, 2002), which is based on the assumption of a constant ﬂux of unsupported 210Pb to a sediment. If that ﬂux, F, can be determined, then sediment age (t) at depth x can be determined by the equation 1 λ t = e−λt1 + A (5.4) λ F x 210 where Ax is the excess Pb inventory between the top of the core and depth x. F can either be assumed or calculated if one age in the sediment core is already known. A common technique for determining this age is 137Cs dating, which exploits the production of this isotope beginning with the ﬁrst nuclear tests in 1955. 137Cs is radioactive isotope that forms as the result of nuclear ﬁssion of 235U, and so its lowermost presence in sediments generally assumed to mark the year 1955 in those sediments. Given a 137Cs or other date for a level in the sediment core, then F can be calculated using the equation Chapter 5: Radiometric Techniques 57
λ∆A F = (5.5) 1 − e−λt1 where ∆A is the total 210Pb excess between the top of the sediment core and the horizon of known age, t1.
5.3 K-Ar System
Potassium-Argon radiometric dating is now done using the 40Ar/39Ar method, but both techniques utilize the decay of 40K to 40Ar via electron capture. Importantly, this is only one of two decay pathways for 40Ar, which also decays to 40Ca via beta decay. However, this does not impose any great diﬃculty in terms of the formulate for dating via this method:
λ 40 40 40 K−Ar λK−Ar+λK−Ca Ar(t) = Ar0 + K(t) e − 1 (5.6) λK−Ar + λK−Ca where λK−Ar and λK−Ca are the decay constants for the K–Ar and K–Ca decay schemes, respectively. Because argon is a noble gas, there is typically no or negligible initial argon 40 in minerals used for K-Ar dating. Hence the Ar0 term can generally be neglected.
A problem with the K-Ar method is that it requires determination of the K content and the Ar isotope ratios on by separate techniques on separate aliquots of a sample. The Ar-Ar method, which is now widely favoured, is able to avoid this problem through irradiation of the sample, which converts all 39K to 39Ar. In this way, K concentration and Ar ratios can be determined by a single analysis of Ar isotope ratios (which are inherently easier to measure precisely than absolute concentrations), which greatly reduces the error in the measurements. The age of a sample can now simply be calculated via the following equation:
1 40Ar∗ t = × J 39 + 1 (5.7) λK−Ar Ar where 40Ar* is equal to the right-hand side of the previous equation and J is the irradi- ation eﬃciency of the conversion of 39K to 39Ar which can be determined empirically by analyzing a constant of known age that was irradiated at the same time as the sample.
Dating by the Ar-Ar is well suited on minerals such as hornblende, mica, and feldspar, all of which are common igneous minerals. However, because Ar tends to diﬀuse easily, the closure temperature for this minerals with respect to the Ar decay product is much lower than the crystallization temperature. The consequence is that the age determined on igneous rocks is a cooling age rather than a crystallization age. This is particularly a problem for intrusive igneous rocks. However, for volcanic rocks, it should not be as sig- niﬁcant of a problem since the cooling age should closely approximate the crystallization age.
Because Ar is easily disturbed by subsequent tectonothermal events, the Ar-Ar method is not suitable for dating the protolith ages of metamorphic rocks. On the other hand, is an excellent tool for dating unmetamorphosed volcanic rocks. Because K is so much more Chapter 5: Radiometric Techniques 58 abundant than U, Ar-Ar dating is generally regarded as the most eﬀective tool for dating young volcanic rocks. Sanidine is the mineral of choice and is abundant in felsic igneous rocks because it crystallizes at relative high temperatures, more comparable to zircon that other feldspars. In maﬁc rocks, biotite and hornblende are also commonly dated and are particularly important in basalts, which are otherwise typically diﬃcult to date by the U- Pb since they usually lack zircons. Ar-Ar is an important dating technique for calibrating the Cenozoic and even Mesozoic time scales and is widely used.
5.3.1 Analtyical considerations Following irradiation, minerals can be directly measured by gas source mass spectrome- try coupled to some device that heats the mineral grains. Samples are typically heated up gradually to release the argon gas over the course of a run through stepwise heating. These days, this is commonly doing using a laser. Because the Ar ratios can be measured continuously, ages can be calculated for diﬀerent stages of release. In an ideal sample that has experienced no Ar loss or gain, this technique should yield a ﬂat spectra of ages for each interval of gas analyzed. However, this is often not the case, and geochronologists will commonly integrate only those ages from a well-deﬁned plateau in the age spectrum.
One signiﬁcant problem in the Ar-Ar dating was that it tended to yield results that were at odds with ages determined by the U-Pb zircon technique. This discrepancy causes quite some strife, but just as the Earth Time initiative aimed to iron inter laboratory diﬀerences in U-Pb analyses, so also did it aim to solve this systematic oﬀset in K-Ar and U-Pb ages, which resulted in large part from the more poorly constrained decay constant for K-Ar. Calibration of the two techniques based on zircons and sanidines in the c. 28.2 Ma Fish Canyon Tuﬀ has now largely resolved this issue, although this was not a simple matter given the diﬀerent crystallization and closure temperatures for these two minerals and the inherent complexity in volcanic systems. However, now that these systems are inter cali- brated, it means that even if the decay rates require further revision in the future, so long as the decay constants are intercalibrated, thet should at least oﬀset proper ages in the same amount and in the same direction, such that rates and timescales will not be aﬀected, even if absolute ages are slightly erroneous.
5.3.2 Detrital muscovite dating Detrital muscovites can be detailed and used for provenance and maximum age constraints in much the same way as zircons. Although muscovite is nowhere near as robust as zircon, it can nevertheless be a useful mineral under some circumstances. For example, muscovite is less prone to surviving multiple recycling events, such that the age spectra of detrital muscovites should be expected to much narrower than that for zircons and more closely reﬂect local provenance. Moreover, because the age on an igneous muscovite is a cooling rather than crystallization age, it can be signiﬁcantly younger than zircon crystals that formed from the same magma body. Hence, muscovites can provide tighter maximum age constraints than zircons. Chapter 5: Radiometric Techniques 59
5.4 Re-Os System
The rhenium-osmium (Re-Os) dating technique exploits the beta decay of 187Re to 187Os, which has an exceedingly long half-life relative to most of the other techniques used. Re and Os are also platinum group metals, and as such occur in relatively low abundance in most minerals. The are concentrated in molybdenum ores, which are routinely dated via this technique, but with regards to the stratigraphic record, the important rock type is organic- rich shales and, to a lesser extent, carbonates. Both Re and Os tend to adsorb to organic matter under hypoxic to anoxic conditions and are thus concentrated in organic-rich sedi- ments. To the extent that only the hydrogenous (i.e., that part derived from seawater) Re and Os can be separated from a rock, then it can be dated using the Re-Os geochronometer.
A signiﬁcant challenge in Re-Os geochemistry has been the development of appropriate selective extraction techniques that collect the hydrogenous Re and Os fractions but do not incorporate detrital Re and Os. Formerly, inverse aqua regia (an HCl-HNO3 mixture) was applied in Re-Os dating, but chromate/sulfuric acid solution is now more commonly applied (Kendall et al., 2004).
The Re-Os goechronometer has potentially powerful applications to calibrating the Pro- terozoic and Paleozoic time scales, in particular for parts of the stratigraphic record where volcanic rocks are rare or absent. One such area that has been notoriously diﬃcult to establish precise ages for is the Neoproterozoic successions in Australia. Brian Kendall and colleagues 2006; 2009 dated black shales from what are interpreted to be post-Sturtian glacial sequences in several of the Australian basins. The result was a series of ages that were similar to each other, which seemed to validate the method. However, the ages gener- ally seemed slightly too young given what was assumed to be the actual age of these rocks. Hence, skepticism remained as to whether or not Re-Os could truly be applied with con- ﬁdence too dating and calibrating the ancient sedimentary record. More recently, Rooney et al. 2014 published a pair of Re-Os from rocks bracketing the Sturtian the glaciation in northwestern Canada based on somewhat improved methods (Fig. 5.2. These ages are consistent with known age constraints on these rocks and seem to validate the technique.
5.5 Radiocarbon Dating
Radiocarbon dating is the most widely applied radioisotopic method for dating Quaternary sediments due to the ubiquity of carbon and a suitable half life (∼ 5730 years). Radiocarbon is produced in the upper atmosphere through bombardment by cosmic radiation, which produces neutrons, which bombard nitrogen: 14 1 14 1 7 N +0 n →6 C +1 H (5.8) Carbon-14 subsequently converts back to nitrogen through beta decay. Even though there is some geographic variation in the production rates of radiocarbon, rapid mixing and dif- fusion in the atmosphere eliminate spatial variability in the ratio of 14C to total carbon. The radiocarbon reservoir in secular equilibrium in the atmosphere, where production rate is balance by loss through decay and incorporation into plants and diﬀusion into water. That is, the atmosphere was in equilibrium prior to atomic bomb explosions, which greatly increased the input of 14C to the atmosphere. Chapter 5: Radiometric Techniques 60
Figure 5.2: Re-Os isochrons from organic-rich samples below and above the Sturtian- equivalent (Rapitan) glacial deposits in the Mackenzie Mountains, Northwest Territories (Rooney et al., 2014).
The remains of land plants and mammals are the most straightforward materials to date radiometrically because they sample the atmosphere. Because living organisms continue to incorporate 14C, radiocarbon dating eﬀectively captures the time of death of those organ- isms. Radiocarbon dating can be applied to marine or lacustrine organic matter, as well as carbonate. However, there are usually additional complications associated with dating such materials. The oceans contain a signiﬁcant of old carbon, such that a radiocarbon age on, say, a benthic marine organism will not provide a precise date on when it died. Lakes and streams are subject to he input of inert carbon from the weathering of old soil organic carbon or carbonate rocks, such that incorporation of some of this carbon into a tissue or shell will yield an apparent age that can be dramatically older than their actual age. Furthermore, the combustion of fossil fuels, which are also inert, has changed the isotopic ratios of both 13C/12C and 14C/12C in the atmosphere, such that organic material that is less than a few hundred years old will yield an erroneously old age. This fossil fuel contribution to the atmospheric carbon isotope ratios is known as the Seuss eﬀect.
The oceanic carbon reservoir is out of equilibrium with the atmosphere due the eﬀect of circulation and the isotopic fractionation associate with diﬀusion across the sea-air inter- face. The average apparent (reservoir) age of low latitude surface waters is about 400 years, and this ﬁgures increases up to 1000–1200 years in the high latitudes. Consequently, any oceanic material from the surface ocean requires a signiﬁcant correction. That is, there is a signiﬁcant radiocarbon gradient between the low and high latitudes. Because it is diﬃcult to ascertain the gradient for the past, the correction required becomes increasingly uncertain for older specimens.
Another challenge in dating materials is the incorporation or contamination of a sample by modern carbon can give rise to a younger apparent age. This is particularly a problem for older materials, that have a low radiocarbon abundance in the ﬁrst place.
Yet another diﬃculty in applying radiocarbon ages is that the assumption that the radiocar- Chapter 5: Radiometric Techniques 61 bon reservoir has remained constant is in fact and not surprisingly not valid. Radiocarbon abundances were shown to vary systematically in tree ring samples. Fortunately, however, tree rings provide a completely independent Quaternary chronology. By coupling radio- carbon measurements of tree rings to a tree ring chronology, as well as to other countable records (Hughen et al., 2004), the ﬂuctuations in atmospheric radiocarbon abundance have been calibrated (Fig. 5.3).
Figure 5.3: Calibration of the radiocarbon record over the past 50Kyr, as reviewed and compiled by Fairbanks et al. (2005).
Initially, radiocarbon activity was measured through the detection of β particles; however, radiocarbon dating is now typically accomplished via measurement of 14C, 13C, and 12C isotopic abundances by accelerator mass spectrometry (AMS). The advantage of AMS over conventional mass spectrometry is that it can eﬃciently separate 14C from 13C, which due to their much diﬀerent abundances, can be a problem. Because isotopic fractionation accompanies photosynthesis and incorporation of atmospheric carbon into tissues, both 14C/12C and 13C/12C must be measured. Conveniently, δ14C is approximately equal to 2×δ13C, where there has been no radiocarbon loss due to decay. Hence, the age of a sample is eﬀectively calculated based on the oﬀset between 2×δ13C and measured δ14C. Chapter 6
Magnetic stratigraphy, or magnetostratigraphy, is one of several geochronometric tech- niques that has seen wide application to dating the geological record. Magnetostratigra- phy as a geochronological tools entails the documentation and calibration of the global geomagnetic polarity sequence (GGPS), which is simply the temporal record of normal versus reversed polarity (Ogg, 2013). This record can then be used for high resolution cor- relation, and to the extent that it is well calibrated, indirect dating. Due to a signiﬁcant recent eﬀorts, the Cenozoic portion of the GGPS is now quite precisely dated, with the most recent part of the record being astronomically tuned (see Chapter 8).
Quite simply, Earth’s magnetic ﬁeld ﬂips polarity through time. Whereas when reversed polarities were ﬁrst discovered, it was thought that reversals should be periodic, it was subsequently demonstrated that reversal rates are stochastic. On average, reversals occur about 3 times per million years during the Cenozoic, but they vary in duration from as little as 30 thousand years to up to 40 m.y., in the case of the Cretaceous superchron. Polarity chrons are zones of a predominant polarity, and tend to be numbered back in time from the present, with the most recent also being named after early pioneers in magnetostratigraphy (Fig. 6.1).
A hierarchy of chron nomenclature also exists, with chrons being lumped into superchrons, and subdivided into sub-chrons, which are a common feature of the GGPS. Additionally, Earth’s magnetic ﬁeld undergoes episodic excursions, which are distinct from both reveals and the usual secular variation in the location of the magnetic north pole. Both subchrons and the excursions provide the potential of more precise dating of a rock sequence by cor- relation with chron boundaries alone.
The magnetic polarity is recorded in sedimentary grains that align with the magnetic ﬁeld during deposition and in magmatic minerals that lock in their magnetic signatures when the cool beyond the Curie temperature. These minerals are almost exclusively iron-bearing min- erals, with iron oxides and oxyhydroxides having the strongest magnetic signatures. The preservation of the ambient magnetic ﬁeld within a rock is known as the natural remnant magnetization (NRM). There are three types of NRM: thermal remanent magnetization (TRM; when minerals cool below Curie temperature), chemical remanent magnetization (CRM; when magnetic minerals grow beyond a critical locking size where it can no longer adjust to the magnetic ﬁeld), and detrital remanent magnetization (DRM, in detrital sed-
62 Chapter 6: Magnetostratigraphy 63
Figure 6.1: Evolution of the GGPS for the past 8 million years. The current GGPS is now calibrated astronomically over this period, yielding very precise ages for the boundaries of the chrons. From (Langereis et al., 2010).
Remanent magnetization can be preserved in both sedimentary and igneous rocks. How- ever, both rock types are subject to overprinting of subsequent magnetic ﬁelds, particularly during diagenetic or thermal events. Consequently, the NRM as measured in a rock is com- monly a vector sum of diﬀerent magnetic components. As a result accurate determination of original magnetic signatures requires a step-wise demagnetization process (usually through stepwise heating) procedure whereby overprinting magnetic signatures can be progressively stripped from the overall magnetic signature. The results of the stepwise measurements are commonly plotted in Zijderveld diagrams, where the declination and inclination com- ponent of the vectors measured from each stage of demagnetization are projected onto a horizontal and vertical projection, respectively (Fig. 6.2). Chapter 6: Magnetostratigraphy 64
Figure 6.2: Illustration of a Zijderveld diagram for representing paleomagnetic data, from (Langereis et al., 2010). The horizontal (declination) component of vector, commonly depicted with ﬁlled points, is project onto NS-EW plane, where the line connecting the origin and the point represents that horizontal component. Hence the length of that line reﬂects the intensity and the angle represents the declination. Here, the declination is about 5◦ west of north. The vertical (inclination) component is commonly depicted with open symbols and projected on to the north-south-down plane, where the angle between the projected vector and the N-S line reﬂects the inclination. A stable primary component is visualized as a series of data points that plot on a line that connects the origin. Secondary components can also be determined, where the declination and inclination are the angles between a line parallel to the N-S axes and the line deﬁned by the data points. The intensity of the secondary component is measured as the distance between the two end member points deﬁning a straight line.
The polarity scale as recorded in the stratigraphic record are closely inter-calibrated with the seaﬂoor record for the past 180 m.y. The magnetostratigraphic record is best studied and calibrated in deep sea sediments which experience steady sediment accumulation and are not prone to the unconformities and diastems that are typical of shallow water sedi- mentary successions. Older stratigraphic successions are also much less inclined to have suﬀered magnetic overprinting. Consequently, the pre-Jurassic GGPS is much more poorly calibrated than the post-Jurassic GGPS. Nevertheless, magnetostratigraphy has important applications in older strata. Local composite polarity scales can be developed and utilized for basin-wide correlation. Some paleomagnetists have even used apparent polar wander paths (AAPWs) as a way to date rocks indirectly. That is, if the AAPW is reasonably calibrated over a certain interval of time, placing the apparent pole for a particular interval of rock provides a way to estimate the age of that rock. This is, however, not a particularly robust technique for dating strata. Chapter 7
Biostratigraphy is a subdivision of sedimentary geology that emphasizes the physical zona- tion of strata based on the fossils they contain. The underling global of of biostratigraphy is to establish bases for correlating age-equivalent strata between diﬀerent geographic lo- calities. Such correlations allow both relative dating of the strata, and where the biozone of interest has been dated reliably elsewhere, and indirect dating method.
Biostratigraphy is not straightforward for several reasons. First, it is built upon the study of what are typically extinct organism, whose evolutionary and environmental complexity can only be surmised. Furthermore, the vagaries of sedimentation, evolving paleoenvirnoments, and taphonomy (that is, preservation) make applying a ﬁxed time scale to fossil ranges a tremendous challenge. In general, the most useful or reliable fossil taxa are those that are short-lived and wide-ranging, while also being prone to preservation in the sedimentary record. These index fossils can constrain the age of a rock in which it is found to a narrow range. As a general rule, the best index fossils are that that live in the open water column. Both nektonic and planktonic taxa tend to be rapidly evolving and widely distributed.
7.1.1 Biostratigraphic zonation 7.2 Biochronology
Biochronology refers to the application of fossil record as a geochronological technique. Biochronology is less concerned with establishing zones and stages than recognizing key fossil or paleontological events (Gradstein, 2012) that can be used as a means for high- resolution correlation and dating. In this regard, biochronology exploits the unidirection- ality of evolution. Two logical classes of fossil events are the First Appearance Datum (FAD) and Last Appearance Datum of a taxon in the fossil record, or alternatively, the ﬁrst or last consistent appearance of a taxon, which compensates for the possibility and likelihood that the ﬁrst or last appearance can be masked by vagaries in the fossil record.
Biochronology developed as a science as a way of dealing statistically with the ﬁrst and last appearance of various ocean plankton and relating them to geomagnetic reversals Gradstein13. What truly distinguishes biochronology from biostratigraphy is that the
65 Chapter 7: Biochronology 66 former attempts to quantify the fossil events—that is, place both date the events and place an uncertainty on that age. Chapter 8
Cyclcity in the stratigraphic record has been recognized since soon after geology became established. Larry Sloss and Harry Wheeler famously begin to subdivide the geological record into large scale cycles, or sequences, which became the seeds for sequence stratig- raphy. In 1932, Harold Wanless coined the term cyclothems to described cycles of marine and non-marine sediments, often with coal seems in the middle and common in the Car- boniferous.
Cyclical sedimentological and stratigraphic patterns are rife in the geological record and occur at many diﬀerent scales. Whereas some of these are truly cyclical, in the sense they reﬂect some sort of periodic or quasi-periodic forcing, many other examples are not true cycles. Some of the most robots periodic and quasi-periodic patterns do not necessarily show up obviously in the lithological or facies stacking patterns, but rather appear in geo- chemical or other high resolution records, such as magnetic susceptibility. Indeed, where the study of terrestrial Pleistocene glacial deposits failed to capture the richness of glacial cycles, oxygen isotope stratigraphy in deep marine cores did. Subsequent study of these oxygen isotope records revealed that they reﬂected orbitally-driven, quasi-periodic ﬂuctua- tions in climate (Hays et al., 1976)—the Milankovitch cycles. The ﬁeld of cyclostratigraphy is now largely dedicated to trying to recognized and analyze orbital cycles in the geological record and use them as a means of tuning the geological time scale.
8.2 Stratigraphic cycles
Cycle stratigraphy is the study of cyclic or quasi-cyclic patterns in the stratigraphic record, which implies that there is a time component to the cyclical patterns. The term is some- what confusing because we often use the term cycles interchangeably with parasequences or simply higher order sequences, which need not imply a true periodic forcing. The term cycles is also often used to describe ostensibly non-cyclic sedimentary patterns. There is also an ongoing discussion over the driving mechanisms behind cyclic patterns, in particu- lar with respect to shallow carbonate cycles (e.g. Drummond and Wilkonson, 1993). That is, internal, basinal processed, related to the sediment supply and sediment routing can give rise to apparent cyclicity, but which, when analyzed, is clearly aperiodic.
67 Chapter 8: Cyclostratigraphy 68
Whether truly cyclic or not, sedimentary patterns resembling cycles may driven by pro- cesses internal or external to the basin. External or allocyclic processes include seasonal or climatic changes, sea level ﬂuctuations, or tectonics. Climatic ﬂuctuations tend to show broad lateral continuity and may be expressed in many ways, several of which are inde- pendent of sea level. Tectonic ﬂuctuations tend to be more periodic—that is, with quite well deﬁned peaks or events (say, e.g., sudden deepening).
Autocyclic processes are conﬁned to the sedimentary basin and tend to be related to the sources and redistribution of sediments.
8.3 Climate-independent cyclic sedimentary patterns
8.3.1 Varves Varves are annual laminae, which may form in lakes or marine settings through a variety of processes tied seasonality. The variations that deﬁne the varies may therefore result from ﬂuctuations in organic carbon content, chemical precipitation, or grain size. Although climatic ﬂuctuations can and do eﬀect the expression of the varves, they are fundamentally driven by seasonality.
8.3.2 Tidal rhythmites Tidal rhythmites, which look superﬁcially like varves, reveal ﬂuctuations in laminae thick- ness due to waxing and waning of tidal intensities, tied to the lunar orbit. Hence the rhythmites, reﬂecting a daily or twice daily cycle of tides is in turn modulated by the longer lunar cycles. Each bundle of laminae within a lunar cycle is referred to as a tidal bundle, and these tidal bundles are in turn arranged in annual packages. Analysis of tidal bundles can be used to deduce the length of days and years in Earth’s history and have been shown to record the gradual increase in the distance between Earth and the moon (Williams, 1998).
8.4 Sedimentary parameters linked to climate change
Periodically or quasi-periodically ﬂuctuating climate may eﬀect sedimentary parameters and patterns, such as geochemistry, mineralogy, and grain size in a variety of diﬀerent ways, some of which are well understood and others less so. These variations may be either extrinsic, meaning they are independent of sedimentation rate, or intrinsic, meaning they are directly related to or inﬂuenced by sedimentation rate (Hinnov and Hilgen, 2012).
8.4.1 Extrinisic sedimentary patterns Oxygen isotopes Whereas there are probably quite a few diﬀerent geochemical proxies that can be used to measure cyclicity, oxygen isotopes are one of the most powerful tools. Oxygen isotope compositions are typically measured in carbonates, where the 18O/16O ratios reﬂect a com- bination of the temperature of the water and the oxygen isotope composition of the water. In the marine realm, the δ18O composition of carbonates increases during glacial maxima, when 18O-depleted ice is locked up on land (hence resulting in 18O-enriched seawater) and Chapter 8: Cyclostratigraphy 69 when the water from which carbonates precipitate (mostly measured on forams) decreases, which results in a greater oxygen isotope fractionation between carbonate and water (and hence higher δ18O carbonates, since carbonate is 18O-enriched relative to water).
Because oxygen isotopes are so sensitive to climate, climatic variations are also recorded in lacustrine and other terrestrial carbonates. However, the degree of ﬂuctuation and the potential drivers of this ﬂuctuation are greater and more diverse in the terrestrial realm, and linking the ﬂuctuation to speciﬁc environmental changes is not always so easy.
Clay assemblages Clays are a byproduct of chemical weathering. Hence, the type and abundance of clays that form during weathering are dependent on climate (both precipitation and temper- ature). Clays tend to alter during erosion, such that signiﬁcant erosion can change the clay content. Speciﬁcally, kaolinite is the common weathering product of granitic rocks in high rainfall regions, whereas mixed smectite and kaolinite is more common in more arid regions (Deepthy and Balakrishnan, 2005). As a consequence, the assemblages of clays in a ﬁne-grained sedimentary rock my vary systematically as a result of changes in surface hydrology, hence climate.
Magnetic Susceptibility Most sediments contain some magnetic minerals. The magnetic susceptibility of a material is the degree of magnetization is exhibits when subjected to an applied magnetic ﬁeld, reported as a dimensionless number scaled to the strength of the applied magnetic ﬁeld. Magnetic susceptibility tends to be highly variable in sediments and is sensitive to climate change through its inﬂuence on weathering and eroding magnetic minerals. Magnetic susceptibility can be measured quickly and at high resolution, hence providing an excellent data set for time series analysis.
Wireline logs Just as magnetic susceptibility may be a useful parameter for detective cyclical variations, so also are gamma ray logs. Because gamma ray signatures reﬂect the abundances of K, U, and Th, it is closely tied to mineralogy and grain size. Mineralogy and grain size, in turn, might be expected to reﬂect environmental parameters beyond simple sea level ﬂuctuations, such the intensity of weathering and strength of erosion.
8.4.2 Intrinsic sedimentary patterns Laminated evaporites Laminated gypsum or anhydrite, both in lacustrine and margin marine settings may ex- hibit cycles related to terrigenous input, which in places have been shown to be cyclic. Those of you that went on the ﬁeld trip to Spain saw an excellent example of one type of evaporite cyclicity. Here, the base of a cycle (depending on how you deﬁne the base), was represented by a continuous gypsum crust, comprising palisade selenite. These tran- sitioned upward into the gypsum supercones, which formed bush-like structures that in places trapped laminated sediments in hollows between branches. This facies represents Chapter 8: Cyclostratigraphy 70 increased input of detrital material in the restricted basin. Freshening of the basin and increased input of marl eventually diminishes and drowns out the gypsum facies, and it is replaced by pure marl. The marl, in turn, is overlain by another seaﬂoor gypsum crust, marking the onset of the next cycle.
A reasonable interpretation of these cycles is that the gypsum crust at the base represents highly evaporative, restricted conditions in the basin: hot, dry weather. Increased pre- cipitation both freshens the basin and increases the input of ﬁne-grained sediments (i.e., erosion related to increased rainfall in the drainage basin).
Marl-limestone and shale-marl Like laminated evaporites, cycles may also be manifest in periodic ﬂuctuations in the carbonate continent of sediments, which can be easily quantiﬁed. This can give rise to either shale-marl or marl-limestone samples. Indeed, some of the best studied sedimentary cycles are documented in this way. The forcing mechanisms behind such cycles is likely quite similar to that behind laminated evaporite cycles—that is, controlled by precipitation and detrital sediment supply, which tends to drown out the carbonate content of the sediments.
Biogenic silica The abundance of silica (speciﬁcally, biogenic silica), is closely tied to biological produc- tivity in some sediments, in particular in large lakes. Fluctuations in productivity can be directly related to climatic ﬂuctuations. For example, Williams et al. (1997) have demon- strated variations in biogenic silica in Pliocene-Pleistocene sediments in Lake Baikal that mirror the marine oxygen isotope record. These same variations are mimicked by magnetic susceptibility (Williams et al., 1997).
Color and grayscale Both colour and grayscale may vary systematically in sediments as a result of ﬂuctuation mineralogy, redox, grain size, and other properties. Spectacular examples include the cycles in the Eocene Green River Formation in Wyoming and the middle Neoproterozoic Chuar Group in Death Valley. Well developed cycles at two diﬀerent frequencies have been detected in sediments from Mars as picked out by variations in grayscale (brightness) (Lewis et al., 2008).
8.4.3 Dyscyclic sedimentary patterns Many types of sedimentation give rise to sedimentary patterns that resemble cycles, but which are not in fact periodic or quasi-periodic in nature. A common example are deltaic cycles, where the sweeping back and forth of the active delta front make give rise to parase- quences that appear cyclic in nature, but are not in fact cyclic at all.
Many well bedded and laminated sediments might at ﬁrst glance appear cyclic. Indeed, ﬁne turbidites are commonly called rhythmites, which implies some sort of cyclicity. But these are just a ﬁne-grained and laminated equivalent of turbidites, which are demonstrably deposited by stochastic processes. Discrimination between cyclic and stochastic controls on sedimentation can be determined statistically. For example, it has been shown that Chapter 8: Cyclostratigraphy 71 turbidite layer thicknesses follow a power law distribution, where the log of layer thickness plotted against the log of the frequency of layer that are greater than or equal to that thickness deﬁnes a straight line. That is, it follows a power law:
N(h) = ah−β (8.1) where h is bed thickness and N(h) is is the number of layers with thickness greater than h, B is the scaling exponent and a is a constant (Rothman et al., 1994).
Shallow-water carbonates Meter-scales cycles are a common feature in shallow-water carbonate successions, in par- ticular those deposited in peritidal environmentals (Ginsburg, 1971; Goldhammer et al., 1993). Typically these are recorded by facies that suggest an abrupt increase in relative water depth. Whereas these cycles were once commonly thought to relate to eustatic ﬂuctuations (Grotzinger, 1986; Goldhammer et al., 1993, e.g.), this notion has been chal- lenged. Drummond and Wilkonson (1993) have argued that many carbonate cycles are best described by stochastic Possonian processes, related to non-linear sedimentary dynamics related to carbonate production and redistribution (Ginsburg, 1971). Alternatively, such cycles could reﬂect (autocyclic) tectonic forcing on base level and water depth Bosencee- tal09.
Modeling of shallow water carbonate deposition under the inﬂuence of ecstatically varying sea level has demonstrated that it is very diﬃcult to parse autocyclic versus allocyclic controls on carbonate cycles. Hence, at present, although tempting and potentially correct to interpret shallow-water carbonate cycles as being environmentally forced, statistical analyses of such cycles is suspect at best (Hill et al., 2012).
8.5 Astronomical Forcing
Milankovitch famously postulated that quasi-periodic variations in solar insolation driven by astronomical forcing could drive systematic climatic ﬂuctuations on Earth (Milankovitch, 1941). However, this extraordinary hypothesis came during World War 2 and well before any robust methods of actually detecting such variations over geological time had been de- vised. As a consequence, Milankovitch’s theory, although grounded in robust calculations, did not gain much traction for decades. It was not until high resolution oxygen isotope records from deep sea cores were developed through the Deep Sea Drilling Program (DSDP) that the ﬁrst hint that Pleistocene glacial cycles were in fact modulated by variations in orbital parameters.
8.5.1 Eccentricity Eccentricity is the measure of the non-circularity of Earth’s orbit around the sun, which is typically quantiﬁed as the dimensionless number, which varies between 0 and 1, with 0 corresponding to a perfectly circular orbit. Earth’s orbit has varied between 0 and 0.07 over the past 10 million years (Fig. 8.1). The frequencies of this variation lie dominantly between 95–99 kyr, 124–131 kyr, and at 405 kyr. The changes are driven by gravitation attractions among the plants, which Jupiter exerting the strongest inﬂuence. Earth’s average distance Chapter 8: Cyclostratigraphy 72
Figure 8.1: Variations in Earth’s three main orbital parameters based on the La2004 astronomical model (Laskar et al., 2004): (a) orbital eccentricity, (b) obliquity, and (c) precession index. From (Hinnov and Hilgen, 2012).
from the sun remains the same no matter what the eccentricity is. What eccentricity does is vary the intensity of the seasons be aﬀecting how close Earth is to the sun during summers and winters.
8.5.2 Obliquity Earth’s spin axis tilts relative to the plane of the ecliptic by bout 22.5◦ at present. This obliquity varies on time scales by a about a total of 2 degrees, such that the angle of tilt oscillates between 22.2 and 24.4◦ (Fig. 8.1). These changes in obliquity are driven by planetary motions in the solar system. The dominant frequency lies at about 40 kyr, but smaller periods occur at about 55 and 29 kyr. Chapter 8: Cyclostratigraphy 73
Figure 8.2: Spectral analysis of Earth’s astronomical parameters as shown in Figure 8.1: (a) eccentricity, (b) obliquity, and (c) precession index. From (Hinnov and Hilgen, 2012). Chapter 8: Cyclostratigraphy 74
Changes in obliquity aﬀect the severity of the seasons, such that higher obliquity increases seasonality and lower obliquity decreases seasonality.
8.5.3 Precession Precession is the wobble of the Earth’s spin axis, which results in the clockwise rotation of where the seasons occur during Earth’s orbit around the sun. That is, the timing of the solstices with respect to the calendar year changes. The frequencies of this variation lie dominantly at about 19 kyr and 23 kyr.
Precession only inﬂuences climate when Earth’s orbit is not circular. When it isn’t circular, it aﬀects climate because it determines where relative to the apogee (point at which the Earth is farthest from the sun) and perigee (the point at which the Earth is closest to the sun). So, for example, the northern hemisphere will experience its hottest summers when the north hemisphere is tilted towards the sun (northern hemisphere summer) when it is at the perigee. By the same token, the northern hemisphere will also experience cold winters in this arrangement, because it will be at the apogee in winter. The precession index, shown in Fig. 8.1, takes into account the inﬂuence of the eccentricity cycle on precession’s eﬀect on solar insolation.
8.5.4 Eﬀect on Insolation and Climate The three orbital parameters inﬂuence Earth’s climate through changing the intensity and timing of insolation across Earth’s surface. Milankovitch himself recognized the importance of solar insolation in the high northerly latitudes, due to its aﬀects on ice sheets, which are a strong climatic ampliﬁer through the albedo eﬀect. Hence, by convention, the so-called Milankovitch cycles are shown for 65◦N.
Just exactly how these ﬂuctuations might aﬀect climate might not be at ﬁrst intuitive. Whereas you might expect that the coldest winters in the high northerly latitudes (think about when these might be) to correspond to glacial expansion, this is not in fact the case. Cold winters result in less precipitation. But even more detrimental to the growth of ice sheets, where there are cold winters, there are hot summers. Melting during hot summers is not propitious for long-term growth of ice sheets. Rather, expansion of northern hemi- sphere ice sheets corresponds to times of milder winters and summers, which result in more snowfall and less melting.
The total change in solar insulation resulting from the Milankovitch cycles is seemingly minuscule, amounting to at most a few percent (and of course, globally average, it does not eﬀectively change). So how then do these small changes in insolation drive large global climatic ﬂuctuations that result in the growth and decay of continental ice sheets? One answer is those ice sheets themselves, which act as a positive feedback on temperature. A second factor is green house gases. Although the exact reason why remains debated, atmospheric CO2 and CH4 concentrations are closely coupled to glacial cycles, and hence help to account for changes in global climate during the Milankovitch cycles. Chapter 8: Cyclostratigraphy 75
8.5.5 The Astronomical Time Scale The gravitational mechanisms of the solar system are suﬃciently well understood and con- strained that the ﬂuctuations in Earth’s orbital parameters can be modelled quite precisely both into the past and into the future, with an anchor at the present time. This has been done and the combined eﬀects on solar insolation have been determined with a high degree of conﬁdence back in time to 50 million years ago (Laskar et al., 2004). Due to long term secular changes in astrodynamical factors that eﬀect the orbital cycles, such as chaotic diﬀusion of the solar system, tidal dissipation of the Earth-Moon angular momentum, and other unknown aﬀects, the orbital cycles have not yet been extended beyond 50 million years. The time predictability of the astronomical parameters provides an opportunity for very high resolution geochronometry on rocks which display orbitally driven ﬂuctuations. If the age of those rocks is known broadly well enough to link the cyclostratigraphic cycles to the proper location on the astronomical time scale, then this astronomical time scale can be used to establish precise ages through the stratigraphic interval. Such tuning was originally applied to the oxygen isotope record [ e.g.]Shackleton00 and is increasingly being applied to older rocks. Much of the late Neogene time scale, for example, has been astronomically tuned based on the exceptional cycles preserved in the Mediterranean area, including Spain (Krijgsman et al., 2001; Hilgen et al., 2007; Manzi et al., 2013). The astronomically tuned age models are now incorporated into the Geological Time Scale, which is astronomically tuned back to 34 Ma(Gradstein et al., 2012).
For sedimentary rocks older than 50 Ma, directly linking to the AST is not possible. Even for sequences that are younger than 50 m.y., i available age constraints might be insuﬃcient to link a cyclostratigraphic record to the ATS. Nevertheless, assuming that some age constraints are available, cyclostratigraphic analysis is still possible and can be used to tune so long as some anchor date is available, such as a radiometric age or magnetochron boundary. Questions remain as to how orbital cycles could manifest themselves during non- glacial time periods, but accumulating evidence suggests that they can. Orbitally forced cycles have been recognized in strata of various ages. For example, Boulila et al. (2014) applied astronchronology to constraining the duration of the Toarcian Ocean anoxic event and Blackburn et al. (2013) have used high precision U-Pb geochronology to support the long held contention that cycles in the early Triassic Newark rift basin are orbital.
8.5.6 The 405-kyr Metronome Most of the frequencies of the astronomical cycles, in particular the ∼100 kyr cycle, change in time due to evolution of the solar system and the parameters that govern these cycles. However, the 405-kyr eccentricity cycle is relatively well behaved and has remained stable for at least the past 250 myr (Laskar et al., 2004) due to its dependence on the large mass of Jupiter. As a consequence, the 405 kyr cycle can be used to calibrate pre-50 Ma records. Boulila et al. (2010) have demonstrated that some of the well developed 3rd-order cycles in the Mesozoic are driven by the 405-kyr metronome. Chapter 9
Chemical stratigraphy, or chemostratigraphy, is the application of chemical signatures pre- served in sediments and sedimentary rocks for the purpose of correlation, dating, or inter- preting past environments and environmental change (Weissert et al., 2008; Halverson et al., 2010). Marine isotope stratigraphy is the most important subdivision of chemostratigra- phy and encapsulates diverse stable and radiogenic isotope systems that have been applied to studies of marine strata spanning from Archean to Recent in age. The list of isotope systems that have been investigated is large and growing, spurred in part by signiﬁcant recent developments in gas source (IRMS) and multi-collector inductively coupled plasma mass spectrometry (MC-ICP-MS). However, the focus here is on a small subset of estab- lished isotope proxies that are widely utilized and particularly useful for chronostratigraphic purposes. Speciﬁcally, the light stable isotope systems of oxygen (δ18O), carbon (δ13C), sulfur (δ34S) and the radiogenic isotope systems of neodymium (143Nd/144Nd), strontium (87Sr/86Sr) and osmium (187Os/186Os) are brieﬂy reviewed.
9.2 Underlying Principles
An overarching principle in the application of marine isotope stratigraphy is that targeted materials for analysis, such as shells or organic matter in marine sediments, can be regarded as reliable proxies of the isotopic composition of the seawater in which they formed. Where these data are to be applied for correlation or dating, it is paramount that the residence time of the isotope system in seawater is suﬃciently large relative to the mixing time of the oceans (∼103 years) that the isotope proxy retains a global seawater signature. If these conditions are met and the isotope systems have not been signiﬁcantly disturbed by subse- quent diagenesis or other alteration, then ﬂuctuations, or in some cases absolute values, in the isotope ratios can be used to correlate strata deposited in diﬀerent parts of the global ocean or assign an approximate age based on comparison with a reference marine isotope curve (see also entry on the Seawater Sr Curve). Because seawater compositions of most commonly applied isotope ratios have varied secularly over Earth history, a simple isotope ratio or proﬁle is typically insuﬃcient to assign an age, and other chronostratigraphic tools, such as biostratigraphy, magnetostratigraphy, or absolute ages on bounding strata must be used ﬁrst to narrow down the possible age range of a given sample or stratigraphic section. The most useful proxies and time intervals in which they are applied are those
76 Chapter 9: Chemostratigraphy 77 that show rapid changes in seawater composition (McArthur et al., 2012). Fortuitously, isotopic shifts and excursions (Figure 9.1) are common across major geological events, such as mass extinctions, glaciations, and episodes of abrupt warming.
non-steady state (excursion)
Isotope RatioIsotope steady state (shift)
A B Time
Figure 9.1: Schematic illustration of the diﬀerence between steady state and non-steady state changes in marine isotope ratios. A steady state shift (blue line) results from a change in the isotope ratio of the input to or output from the ocean of a given system. For example, a shift in 87Sr/86Sr could result from the extrusion of a large continental ﬂood basalt (time A) and resulting decrease in average weathering input. The timescale over which the shift occurs is proportional to the residence time of that element in the ocean. A non-steady state perturbation (red line) results in an excursion, commonly driven by a relatively rapid input of the element of interest to the ocean reservoir (at time A), such as a catastrophic release of low δ13C methane from seaﬂoor clathrate reservoirs. The magnitude of the excursion is a measure of the relative size of the non-steady state addition (or subtraction) of the element to seawater, and the timescale over which the system recovers is governed by the residence time.
9.2.1 Stable isotopes Temporal ﬂuctuations in stable isotope compositions of seawater result from the tendency of isotopes to fractionate at low temperatures due to their relative mass diﬀerences. These mass discrepancies give rise to subtle but important diﬀerences in the strength of chemical bonds formed in molecules by the diﬀerent isotopes of a given element. The result is the Chapter 9: Chemostratigraphy 78 preferential diﬀusion, reaction or biological uptake of one isotope over another, resulting in fractionation of isotope ratios between diﬀerent molecules or phases. This fractionation generates geological or oceanic reservoirs with contrasting isotopic ratios. For example, the exit channels of carbon from the seawater dissolved inorganic carbon (DIC) pool can be subdivided into inorganic (i.e. carbonate; relatively 13C-enriched) and organic sedimentary carbon (13C-depleted) sinks. The relative importance of these sinks controls the 13C/12C ratio of the DIC pool at steady state.
Many biogeochemical and geological processes can drive the secular variation in isotope ratios in seawater, but those changes that occur at time scales useful in stratigraphy are commonly the result of changing paleoenvironments. Where these changes occur at steady state (that is, the mass of the seawater reservoir of the element of interest does not change in time), the resulting isotopic change is gradual. However, the most useful isotopic ﬂuc- tuations are commonly in the form of relatively brief excursions (Fig. 1) that result from sudden or catastrophic events, such as glaciation (see section on oxygen isotopes below), ﬂood basalt volcanism, and mass extinctions. Such events often give rise to precipitous (non-steady state) changes in multiple isotope proxies simultaneously (e.g. ﬂuctuations in carbon and oxygen isotope ratios are commonly correlated or anti-correlated). Impor- tantly, the time scale of these changes may diﬀer due to variable oceanic residence times and the proximal driving mechanism for the perturbation.
9.2.2 Radiogenic isotopes Rapid changes in the ratios of radiogenic isotopes in seawater are fundamentally the result of the fractionation of two elements that include the parent and daughter isotopes of a given radiogenic system (e.g. Rb and Sr, where 87Rb decays to 87Sr) between geological reservoirs, rather than isotopic fractionation. Coupled with progressive decay of the ra- diogenic isotope, this elemental fractionation produces reservoirs, such as the mantle and the continental crust, with distinct radiogenic isotope ﬁngerprints. Progressive decay of the radiogenic isotope results in a ﬁrst-order, long-term rise in the radiogenic isotope ratio over Earth history. Changes in the relative weight of the ﬂux to the ocean of one of these elements from the diﬀerent reservoirs results in a steady state shift in the isotope ratio in seawater, over a time scale determined by the residence time of that element in seawater (Figure 9.1) and the time span over which the change in ﬂux occurs. In some cases, sharp excursions may occur.
9.3 Development and Application of the Method
The origin of marine isotope stratigraphy can be traced to the pioneering work of Harold Urey and his research group at the University of Chicago in the 1950s who analyzed car- bon and oxygen isotope ratios in modern and ancient carbonate shells and other biogenic materials. The initial motivation of Ureys group was in reconstructing paleoenvironments, namely past seawater temperatures (Epstein et al., 1951; Urey et al., 1951). However, Emiliani’s 1955 documentation of systematic stratigraphic (hence, temporal) variations in oxygen isotope ratios in foraminifera tests recovered from Pleistocene sediment cores established the great potential of oxygen isotope stratigraphy for correlation and dating purposes (Emiliani, 1958). Oxygen isotopes would subsequently prove to be an invaluable and much more eﬀective means of reconstructing the tempo and magnitude of late Cenozoic Chapter 9: Chemostratigraphy 79 glacial advances and retreats than the continental record of glacial drift (e.g. Shackleton and Opdyke, 1973; Hays et al., 1976; Lisieki and Raymo, 2005).
Although Wickman (1948) ﬁrst speculated that strontium isotopes were a potentially useful dating tool in marine carbonates and evaporites well before analytical techniques became routine, the full potential of strontium isotope stratigraphy was not realized until Burke et al. (1982) published the hallmark record of Phanerozoic seawater Sr isotope composi- tions, which demonstrated substantial variations in 87Sr/86Sr. Similarly, carbon isotope stratigraphy did not emerge as an eﬀective and eﬃcient chronostratigraphic tool until the late 1970s and early 1980s (e.g. Berger et al., 1978; Scholle and Arthur, 1980), but is now a standard technique when working on sedimentary carbonates of all ages.
9.4 Stable Isotope Systems
Due to the relatively small absolute variations in stable isotope ratios, such as 13C/12C and 18O/16O, most commonly measured stable isotope ratios are expressed as per mil (parts-per-thousand) units in delta (δ) notation, relative to an international standard:
R − R δ( ) = 1000 × SAM STD (9.1) RSTD h where RSAM and RSTD refer to the isotope ratio of interest of a measured sample and standard, respectively. Many diﬀerent isotope ratios are measured on sediments and sed- imentary rocks, but of these, few have proven consistently eﬀective in chronostratigraphy. Several other systems have very speciﬁc environmental applications, such as boron iso- topes (paleo-pH) and the clumped C-O isotopes (paleo-temperature) in carbonates. Other proxies, such as Mg, Ca, and Li isotopes in carbonates, have emerged as potentially use- ful chemostratigraphic tools, but are either not yet well established or are not eﬀective replacements for established isotope systems.
9.4.1 Oxygen isotopes Oxygen isotope stratigraphy has proven invaluable in understanding Cenozoic paleoclimate (Emiliani, 1958; Hays et al., 1976; Zachos et al., 2001; Lisieki and Raymo, 2005), and has enjoyed wide application to rocks and sediments of all ages. Oxygen isotope ratios are normalized to two diﬀerent standards: PDB (named for belmnites in the Cretaceous Peedee Formation, which was the original laboratory carbonate standard developed by Ureys group at the University of Chicago, and subsequently recalibrated as VPDB) and SMOW (Standard Mean Ocean Water standard, recalibrated as VSMOW). Oxygen isotope values for a given material can be converted between these two standardizations using the equation
18 18 δ OVSMOW = 1.03091 × δ OVPDB + 30.91 (9.2) By convention, silicate and phosphate minerals are standardized to VSMOW and carbon- ate minerals are standardized to VPDB. Here, we are only concerned with those miner- als deposited from seawater, most importantly aragonite and calcite, although oxygen in phosphate biominerals is also used for chemostratigraphy. The original oxygen isotope composition of these minerals is determined by temperature at which the mineral formed Chapter 9: Chemostratigraphy 80 and the δ18O of the oxygen source for that mineral, which in the case of carbonates, is the oxygen bound to carbon in the dissolved inorganic carbon (DIC) pool (Grossman, 2012). The δ18O of DIC, in turn, is governed by the δ18O of seawater. Hence, oxygen isotope ratios can be used as a paleothermometer, assuming the composition of the waters from which the carbonate was precipitated is known (Epstein and Mayeda, 1953).
Carbonates are strongly enriched in 18O relative to the waters from which they precipi- tate. Because this fractionation is controlled by temperature, with larger fractionations occurring at lower temperatures, if seawater composition is held constant, then increas- ing temperatures result in carbonates with lower δ18O values, and vice versa. However, seawater composition does vary, speciﬁcally as a consequence of the growth and decay of continental ice sheets, which are strongly depleted in 18O compared to seawater. The mag- nitude of this shift between the last glacial maximum and today was about 0.9 (Schrag et al., 1996). Fortunately, this ice volume eﬀect drives the δ18O of seawater to higher values during glacial maxima and lower values during glacial minimathat is, in the same direc- tion as the temperature eﬀect. Hence oxygen isotopes are invaluable in reconstructing late Cenozoic (PliocenePleistocene) glaciation, which is now widely accepted to be modulated by orbital (Milankovitch) cycles (Hays et al., 1976). Indeed, both marine carbonate and ice core δ18O records over the past million to 5 million years are commonly tuned to os- cillations in insolation (typically June 21 insolation at 65◦N; e.g. Shackleton, 2000; Zachos et al., 2001).
Figure 9.2 shows the marine benthic foraminifera (i.e. deep water) δ18O record (LR04) for the past 3 m.y., as compiled and orbitally tuned by (Lisieki and Raymo, 2005). Benthic forams are preferred over planktonic forams for generating seawater curves because the latter are sensitive to local temperature or salinity aﬀects, whereas benthic forams better capture whole ocean values and trends. The LR04 record captures many key features, including a general increase in δ18O of marine carbonates through the PliocenePleistocene that is presumably related to gradual intensiﬁcation of continental glaciation, the well- developed, asymmetric sawtooth 100 k.y. glacial-interglacial cycles over the past 800 k.y., and the prevalence of 40 k.y. glacial cycles prior to 800 k.a. The peaks and troughs in the marine isotope record are identiﬁed as numbered marine isotope stages, beginning with the present low (interglacial) stage (MIS 1) and extending back through the Pliocene Lisieki and Raymo (2005). Correlating stages between a given oxygen isotope record and the LR04 or other compilation provides a means of dating sediment (or ice) cores and enables direct correlation and comparison of equivalently aged strata, which remains valid even if the timescale of the compilation is subsequently revised. This technique is most reliably applied to deep-sea strata that experience consistent, uninterrupted sedimentation.
Oxygen isotope stratigraphy is also widely applied to older strata and successfully delin- eates many key climatic events and trends over the Cenozoic Era, including the Paleocene- Eocene Thermal Maximum (Fig. 9.3A), the onset of Antarctic glaciation (Eocene-Oligocene boundary), and mid-Miocene climatic optimum (Zachos et al., 2001). The utility and reli- ability of this proxy decreases back in time with diminished deep-sea sediment records and increasing degree of alteration of most carbonate sediments and rocks. Nevertheless, δ18O anomalies and trends that are consistent with known events such as glaciations have been documented in rocks spanning the Phanerozoic Grossman (2012). Chapter 9: Chemostratigraphy 81
δ18O (‰ VPDB) 0.0 2 1 6 5 0.2 10 9 0.4 13 14 Brunhes 0.6 17 18 0.8 22 21 25 Mat. 1.0 26 29 J 33 1.2 37 43 1.4 49 1.6 55 61 1.8 67
73 Old 2.0 79
2.2 85 Age (Millions of years before present) (Millions before Age of years 2.4
2.8 Gauss 3.0 5 4 3
Figure 9.2: The benthic foraminifera oxygen isotope record (LR04) for the past 3 billion years, as compiled by (Lisieki and Raymo, 2005) from a stack of 57 sediment core records. Red numbers refer to marine isotope stages (MIS), although note that only selected stages are labeled. Glacial maxima occur at higher δ18O values (positive MIS; see text for expla- nation), hence the inverse values on the x-axis. The boxes on the right delineate normal (blue) and reversed (white) magnetic polarity chrons. Mat. = Matuyama; J = Jaramillo; Old = Olduvai. Chapter 9: Chemostratigraphy 82
-1 O (‰ VPDB) 18 δ North Atlantic -2 Paci c Southern Ocean
C (‰ VPDB) 0 13 δ -1
Paleocene Eocene 57.0 56.5 56.0 55.5 55.0 Age (Millions of year before present)
Figure 9.3: Compilation of oxygen (A) and carbon (B) isotope data from benthic foraminifera, spanning the Paleocene-Eocene Thermal Maximum (PETM) event, which is attributed to a catastrophic input of 13C-depleted carbon to the ocean-atmosphere sys- tem (see Zachos et al., 2007 for a review of possible mechanisms and consequences of this event). Carbon and oxygen isotope records (extracted from the database of Cramer et al., 2009) from individual cores from the North Atlantic (ODP1051), Paciﬁc (ODP865), and Southern Oceans (ODP690) are shown for comparison.
9.4.2 Inorganic carbon isotopes The carbon isotope ratio (13C/12C) of the marine dissolved inorganic carbon (DIC) pool has varied secularly over Earth history. The steady state δ13C value of DIC reﬂects the composition of carbon input to the ocean (commonly assumed to be 5 to 6 over long time scales; Kump and Arthur, 1999), the relative partitioning of that carbonh into inorganic (carbonate; δ13Ccarb) carbon and sedimentary organic carbon (δ13Corg) that is buried within marine sediments (forg = fraction of carbon buried as organic matter), and the isotopic diﬀerence between these two reservoirs (δ13Ccarb δ13Corg ≈ δ13CDIC δ13Corg = carb−org) (Fig. 9.4). Because δ13Ccarb is close to value of the DIC reservoir and the fractionation between the two species is eﬀectively insensitive to temperature, the carbon Chapter 9: Chemostratigraphy 83 isotope composition of marine carbonates is treated as an approximation of the seawater from which it derives.
Volcanic Atmospheric CO2 emissions + continental weathering (≈ –5‰) Dissolved Inorganic Carbon (δ13C ≈ 1‰)
Carbonate sedimentary sediments organic carbon (≈ 1‰) (≈ -21‰)
Figure 9.4: A simple schematic model of the coupled atmosphere-ocean carbon cycle at present, which illustrates the carbon isotope composition of the major ﬂuxes and reservoirs that aﬀect the δ13C value of dissolved inorganic carbon (DIC) in seawater and sedimentary carbon over geological time scales. Steady state changes in the δ13C value of DIC can be eﬀected by 1) a change to the composition of the input to the coupled ocean-atmosphere system, 2) a change in the relative amount of carbon removed from the ocean-atmosphere system and buried in sediments, 3) δ13C composition of organic carbon. A non-steady excursion can be brought about by a sudden input of isotopically distinct carbon to the DIC-atmosphere reservoir.
This approximation is most applicable to pelagic carbonates, which are deposited in the open ocean. The predominant sources of pelagic carbonate are calcareous nannoplankton (e.g. Coccolithophora) and planktonic foraminifera. Although these two groups evolved earlier, neither was widespread in the global ocean until the early Cretaceous (Saltzman and Thomas, 2012). Consequently, most middle MesozoicCenozoic marine carbon isotope compilations are derived from deep-sea sediment cores. These may be either bulk car- bonate records, which dominantly comprise planktonic nannofossils (Milliman, 1993), or records developed from individual planktonic or benthic foraminifera (e.g. Fig. 9.5). Be- cause the DIC of the surface mixed layer may be heavily modiﬁed by primary productivity (photosynthesis), planktonic and bulk records may vary considerably from the benthic foraminifera records, which are a more faithful proxy for whole ocean DIC composition. Combined planktonic and benthic foraminifera records may be used to quantify the eﬀect of the biological pump, which produces a surface-to-deep δ13C gradient where surface wa- Chapter 9: Chemostratigraphy 84 ters are relatively 13C-enriched.
Even benthic foramifera records may not precisely record the carbon isotope composition of the integrated oceanic DIC pool. As shown in Figure 9.5, the present Paciﬁc and North Atlantic basins are oﬀset from one another by about 1 . This carbon isotope gradient between the two ocean basins ﬁrst developed in the middleh Miocene. For the remainder of the Cenozoic Period, the North Atlantic and Paciﬁc records align reasonably well, and both capture major shifts in δ13C, including a large decline across in the PaleoceneEocene boundary (punctuated by the PETM event; see below), the ca. 17 Ma Monterrey carbon burial event, and a steady decline starting about 8 million years ago and presumably re- lated to the expansion of C4 plants (Fig. 9.5; Zachos et al., 2001). Such salient trends are important from the perspective of chronostratigraphy because they permit correlation and approximate dating of poorly age-constrained sections by comparison with a reference seawater δ13C curve. However, because both positive and negative steady state shifts are relatively common occurrences in the geological record, they typically require additional chronological constraints or a long record with enough structure that it can be conﬁdently correlated to the reference curve.
Non-steady state perturbations to the seawater DIC isotopic composition, known as car- bon isotope excursions, may provide a unique solution when correlating and dating marine sediments. Perhaps the best-studied example of such a non-steady state excursion is the carbon isotope anomaly that spans the Paleocene-Eocene Thermal Maximum (PETM; Fig. reﬃg:Fig3). This anomaly features an abrupt decline in the δ13C composition of benthic foraminifera of 23 , followed by a gradual recovery over a few hundreds of thousands of year; this timescaleh is governed by the c. 100200 thousand-year residence time of DIC in the oceans (Fig. 9.3B). It is widely accepted that this excursion must have resulted from the geologically instantaneous addition of a large mass of 13C-depleted carbon to the marine DIC pool. However, the actual source of this carbon and its precise link to the cor- responding temperature change evinced by the oxygen isotope record (Fig. 9.3A) remain debated (Higgins and Schrag, 2006).
Carbonates deposited on continental margins tend to show greater variability than pelagic carbonate records. One reason for this discrepancy in signals is that they are more sus- ceptible to local and regional aﬀects that generate heterogeneity in seawater DIC, such as input of 13C-depleted surface and groundwater or elevated bioproductivity and removal of 13C-depleted organic matter. This problem is particularly acute in restricted water bod- ies, such as epeiric seas, where local carbon cycling processes can strongly inﬂuence the δ13C record preserved in carbonate sediments (e.g. Holmden et al., 1998; Panchuk et al., 2005). Another reason for the variability is that platform carbonates are more susceptible to overprinting by meteoric diagenesis, particularly during glacial epochs where they are periodically exposed during glacioeustatic lowstands (Swart, 2008). Nevertheless, because deep-sea pelagic carbonates are absent from the pre-Mesozoic record and other deepwater records are sparse, carbon isotope stratigraphy is heavily dependent on records derived from shallow water carbon platforms (Saltzman and Thomas, 2012). The reliability of these archives for chemical stratigraphy remains highly debated (Derry, 2010). Chapter 9: Chemostratigraphy 85
North Atlantic Paci c 3
0 C (Benthic Foraminera) 13
-2 PETM Oi 1 Glaciation Monterey carbon C4 plants burial Paleocene Eocene Oligocene Miocene Pli. Q 65 55 45 35 25 15 5 Age (Millions of year before present)
Figure 9.5: Compilation of carbon isotope composition of benthic foraminera over the past 65 million years (Cenozoic). Blue data points represent all data from deep-sea sediment cores, as compiled by Cramer et al. (2009), adjusted to the time scale of Gradstein et al. (2004). White swaths indicate intervals with extensive data coverage. Black and grey lines are ﬁts to the North Atlantic and Paciﬁc data, respectively Cramer et al. (2009). Several carbon isotopic features are apparent in this curve, including the Paleocene-Eocene Thermal Maximum (PETM) negative excursion (see also Figure 9.3), the Monterey Carbon Burial Event, and a decline in marine δ13C corresponding to the expansion of C4 plants beginning about 8 million years ago (Zachos et al., 2001). Note the gap between Paciﬁc and North Atlantic δ13C beginning in the middle Miocene, indicating an increasing isotopic gradient between ocean basins.
9.4.3 Organic carbon isotopes The carbon isotopic composition of marine sedimentary organic matters (δ13Corg) also yields important paleoenvironmental information, and in some cases is a useful chemical stratigraphic tool. Most carbon isotope data produced on sediments and sedimentary rocks are measured on bulk total organic carbon (TOC). To the extent that this TOC reservoir is largely derived from primary biomass and has not been heavily altered isotopically during diagenesis and metamorphism, it should closely parallel the original DIC pool from which that carbon was ﬁxed (Hayes et al., 1999). For this reason, the carbon isotopic composition of bulk sedimentary organic carbon (δ13Corg ≈ -20 to -30 ) has been widely applied in place of or as a supplement to carbonate carbon isotopesh (where carbonate is scarce) to capture secular evolution and excursions in the DIC reservoir (e.g. Knoll et al., 1986).
Paired carbonate and organic carbon data can additionally be used to calculate the net isotopic fractionation between the inorganic and organic carbonate fractionations (TOC ), which is an approximation of the initial fractionation between DIC and the primary (pho- Chapter 9: Chemostratigraphy 86 tosynthetic) organic carbon pool (Hayes et al., 1999). This ﬁgure can vary considerably due to a several factors, but mainly to diﬀerences in the initial fractionation between dissolved CO2 and primary biomass (P ) which is imparted by primary producers and principally controlled by dissolved CO2 concentrations, growth rate, and cell geometry (Popp et al., 1998). Not surprisingly, TOC has varied considerably over geological time, with highs of up to 34 during the Neoproterozoic (100–541 Ma) and current values below 22 (Hayes et al., 13 TOC 13 carb 1999). As a consequence, δ C cannot simply be used as a replacementh for δ C , even though it commonly captures the major excursions and trends in the DIC reservoir and hence can be applied to chronostratigraphy.
9.4.4 Sulfur isotopes Sulfur isotope stratigraphy has been applied to a wide range of paleoenvironmental ques- tions, for which the ratio 34S/32S is commonly used. Sulfur occurs in seawater mainly in 2− the oxidized form of sulfate (SO4 ), but locally and at times in the past has also been abundant in its reduced form of hydrogen sulﬁde (HS-). Sulfate is delivered to the ocean mainly through continental weathering (δ34S = 0–10) and removed as both sulfate and sulﬁde. The sulfate is sequestered in evaporite minerals (gypsum and anhydrite), barite (CaSO4), and as a trace constituent in carbonate minerals (carbonate associated sulfate), all of which may be used as proxies for seawater sulfate isotopic compositions. Under anoxic conditions, sulfate is also used as the electron acceptor during bacterial sulfate reduction (BSR). This metabolic pathway, along with complex redox recycling of intermediate sul- fur species, produces sulﬁde that is depleted in 34S relative to seawater sulfate typically by 34 ∼30–60 . The highly variable net fractionation results in a wide range of δ Ssulfide values 34 in the sedimentaryh record (Canﬁeld, 1998). For this reason, δ Ssulfide is not an especially 34 useful tool in chronostratigraphy, although distinct ranges in δ Ssulfide are characteristic of certain time intervals over Earths history (e.g. Canﬁeld, 1998).
The generally highly depleted values of sedimentary sulﬁde are responsible for the rela- tively high δ34S of sulfate in seawater. Sulfate is the second most abundant anion (after Cl-) in modern seawater and has a residence time of nearly 10 million years. Hence, sulfate 34 is eﬀectively uniform in the modern ocean (δ SSO4 = 21.0 ± 0.2 in modern seawater; Rees et al., 1978) and its isotopic composition is expectedh to remainh relatively stable over time scales of millions of years (Paytan and Gray, 2012). Nevertheless, δ34SSO4 has varied considerably over Earths history, in large part as a result of signiﬁcant changes in the size of the sulfate reservoir and hence its sensitivity to perturbations by changes in other pa- rameters controlling its isotopic composition, namely the net fractionation accompanying BSR and sulﬁde burial and the fraction of sulfur buried as sulﬁde. Due to high intrin- sic variability of δ34SSO4 values recorded in pre-Mesozoic sedimentary rocks and diﬃculty in establishing their ﬁdelity as seawater proxies, sulfur isotope stratigraphy has limited application as a dating tool throughout most of Earth history. However, in some cases, sulfur isotope chemostratigraphy can be used to apply approximate age constraints. For example, distinctly 34S-enriched values (δ34SSO4 > 30 ) characterize the latest Precam- brian and Cambrian sulfur isotope record (Kampschulteh and Strauss, 2004). The sulfur isotope record for the past 130 million years shows many distinct features, notably a 5 de- cline between 130 and 120 Ma and a similar rise around 55 Ma (Figure 9.6; Paytan et al., 2004; Wortmann and Paytan, 2012). This unique structure to the seawater δ34SSO4 record presents potential of using sulfur isotopes for age determinations in sedimentary rocks of Chapter 9: Chemostratigraphy 87 middle Mesozoic and younger age, although in many cases, other proxies would be more eﬀective.
18 S (‰ CDT) 34
15 Early Cretaceous Late Cretaceous Pal. Eocene Olig. Miocene P Q 140 120 100 80 60 40 20 0 Age (Millions of year before present)
34 Figure 9.6: A LOWESS ﬁt sulfur isotope curve to the marine barite δ SSO4 record for the past 130 million years (Paytan et al., 2004), modiﬁed from (Paytan and Gray, 2012). Barite, which is a trace constituent in many deep-sea sediments, is presumed to preserve primary seawater sulfate isotope ratios and is robust during diagenesis, which makes it a well suited mineral for reconstructing ancient seawater composition.
9.4.5 Lithium isotopes Lithium is an alkali element which behaves as a conservative cation in the ocean, where its residence time is ∼1.2 million years and δ7 composition about 31 (Hathorne and James, 7 6 2006; Misra and Froelich, 2012). Lithium has two stable isotopesh ( Li, Li), and the large relative mass diﬀerence results in very large fractionations at surface temperatures. Li isotope values on the earth surface vary from about 0 to 42 (von Strandmann et al., 7 2013). The result is signiﬁcant ﬂuctuations in δ Li of seawater,h as recorded in marine carbonates, over at least the Cenozoic (Fig. 9.7). Due to the long residence time and propensity to vary signiﬁcantly, Li is a potentially useful chemostratigraphic tool. Among the most pronounced fractionations to aﬀect the Li cycle are related to continental weathering. At present, the dissolved lithium load in seawater is about 21 heavier that average continental crust (Fig. 9.8) due to the large fractionation involvedh in dissolution of silicate minerals as well as the overall weathering intensity (Kisak¨ureket al., 2005). High fractionations are only expressed where silicate weathering is incongruent. In this case, the preferentially uptake of isotopically light Li during the formation of secondary aluminosilicate clay minerals and oxides results in elevated δ7Li in the riverine dissolved load (Vigier et al., 2008). In contrast, where weathering intensity is high and dissolution Chapter 9: Chemostratigraphy 88
Figure 9.7: The Cenozoic seawater Li isotope record as reconstructed from planktonic foraminifera. Modiﬁed from (Misra and Froelich, 2012).
is congruent, all Li is released and the dissolved load is much lower in δ7Li. The net result is a large range in riverine δ7Li ratios, where this variation is not driven by diﬀerences in bedrock but rather diﬀerences in weathering intensity. In short, where silicate weathering is limited by the physical supply of crustal silicate minerals (i.e., thick soils and equilibrium with clays), the δ7Li of river waters will be low—similar to bedrock values. In contrast, where weathering is limited by temperature and precipitation (thin soils), tδ7Li will be high.
The riverine source of Li to seawater accounts for somewhat more than a third of total Li input. Additional sources include hydrothermal ﬂuids, resulting from high temperature alteration of the seaﬂoor, and reﬂuxed Li from subduction zones (Fig. 9.8). Although Li is preserved as a trace constituent of carbonates, the main Li sink from seawater is from low-temperature alteration of oceanic crust, where Li is incorporated into altered Mg silicates, and reverse weathering, which produces authigenic aluminosilicate clays (Misra and Froelich, 2012). On average, these processes remove Li that is fractionated relative 7 to seawater by 15 (that is, ∆ Liwater−clay ≈ 15 ). Lithium incorporation in carbonate entails a small fractionation,h such that carbonateh is about 3–5 depleted relative to sea- h Chapter 9: Chemostratigraphy 89
Figure 9.8: The Li cycle of modern seawater, taken from Misra and Froelich (2012).
water (von Strandmann et al., 2013), but the removal of Li into carbonates is too small to have any measurable aﬀect on Li isotope mass balance in seawater.
The lithium isotope record of seawater (Fig. 9.7) shows a ﬁrst order rise of about 9 during the Cenozoic, as well as a series of prominent ﬂuctuations. At face value, the substantialh rise appears to reﬂect in increase in incongruent weathering (that is, physical weathering dominating over chemical weathering), consistent with the inﬂuence of Himalayan uplift. A decline of several permit that is closely associated with intrusion of the Columbia River Flood Basalts, rather than reﬂecting the increased weathering component of basalt, as is the case for Sr (see below), most likely reﬂects generally more intense weathering resulting from higher CO2 and accompanying run-oﬀ.
9.5 Radiogenic isotope systems
In addition to their direction applications in geochronology, several diﬀerent radiogenic isotopes systems serve as important chemical stratigraphic tools. Although many diﬀerent radiogenic isotope systems have been applied to marine sedimentary records for multiple purposes, only a limited number of these techniques enjoy routine application with chronos- tratigraphic relevance. Radiogenic isotope ratios are conventionally reported with a stable isotope with a radiogenic contribution of a given element as the numerator (e.g. 87Sr) and a stable isotope with no or a negligible radiogenic component as the denominator (e.g. 86Sr). Chapter 9: Chemostratigraphy 90
9.5.1 Neodymium isotopes Nd isotopes, speciﬁcally the ratio 143Nd/144Nd are applied to both recent and ancient and carbonate and clastic sediments and sedimentary rocks. Neodymium isotope ratios in sedi- mentary rocks are commonly reported in epsilon notation (Nd(t)), where the 143Nd/144Nd ratio is normalized against the chondritic uniform reservoir (CHUR) and both are corrected for radiogenic ingrowth of 143Nd (DePaolo and Wasserburg, 1976):
" 143 # ( Nd ) = 144Nd sample(t) − 1 × 10000 (9.3) Nd(t) 143Nd ( 144Nd )CHUR(t) Like the delta notation in stable isotopes, this convention simply makes dealing with Nd isotope ratios more tangible.
Neodymium isotope stratigraphy owes its utility to the fact that Sm is preferentially re- tained in maﬁc magmas relative to Nd, leading to a greater contribution of radiogenic 143Nd resulting from alpha decay of the parent isotope 147Sm. This elemental fractiona- tion gives rise to relatively radiogenic (higher 143Nd/144Nd) maﬁc rocks and less radiogenic felsic rocks, which makes the 143Nd/144Nd ratio a powerful tracer of sediment and solute sources. However, due to the low solubility of the rare earth elements (which include both Sm and Nd) in seawater, their tendency to adsorb onto settling particles, and isotopic exchange between sediments and seawater, the eﬀective isotopic residence time of Nd in the oceans is small: ∼200 to 1000 years (Tachikawa et al., 1999). Consequently, seawa- ter is not uniform with respect to 143Nd/144Nd ratios, and no unique seawater Nd curve can be generated. Rather, Nd isotopes, as typically measured in carbonates, phosphates, and ferromanganese nodules, are more commonly used to track the sources and circulation of ancient water masses (Banner, 2004). Figure 9.9 shows the Nd isotopic evolution of deep waters in the Paciﬁc, Indian, and Atlantic oceans over the past 50 m.y. as recorded in ferromanganese nodules Frank et al. (1999). The stark diﬀerence between the three ocean basins reﬂects the diﬀerence is relative contribution of juvenile maﬁc crust and old continental crust to the weathering input to each basin. The decreasing Nd values in the Atlantic presumably reﬂected decreasing input from maﬁc igneous rocks such as the Central Atlantic Magmatic Province which accompanied break-up of Pangaea and initial opening of the Atlantic ocean basin.
Neodymium isotope ratios measured in ﬁne-grained siliciclastic sediments can be used to track tectonic, volcanic, and geographic evolution of the drainage basins supplying those sediments. To the extent that shifts in the initial 143Nd/144Nd composition of sediments can be linked to speciﬁc events, such as basement uplift or an episode of ﬂood basalt magmatism, the Nd isotope record can have important chronostratigraphic applications to poorly dated sedimentary sequences.
9.5.2 Strontium isotopes Strontium isotope stratigraphy is one of the most powerful and widely applied chemostrati- graphic tools to determine ages in otherwise undated marine rocks (see also entry on the Seawater Sr Curve). Similar to the Nd isotope system, strontium isotope ratios, namely 87Sr/86Sr, are useful because of the elemental fractionation between felsic and maﬁc rocks. In this case, Rb is concentrated in felsic rocks, resulting in higher 87Sr/86Sr ratios in felsic Chapter 9: Chemostratigraphy 91
Figure 9.9: Nd(t) evolution of Paciﬁc, Indian, and Atlantic deep waters over the past 50 m.y. From Frank et al. (1999).
continental crust from the decay of 87Rb. In contrast, maﬁc continental and oceanic crust have low (unradiogenic) 87Sr/86Sr. Distinct from Nd, however, the Sr budget of seawater (Fig. 9.10) includes not only a riverine input reﬂecting typically radiogenic continental sources, but also a large unradiogenic component from the hydrothermal alteration of oceanic crust. The residence time of strontium is also much higher than neodymium: 4–5 million years, meaning that oceans are uniform with respect to 87Sr/86Sr, except in rare cases where local water masses are heavily inﬂuenced by isotopically distinct freshwater or hydrothermal input.
Because the Sr2+ cation readily substitutes for Ca2+, marine strontium isotope ratios can be preserved in evaporite, carbonate, and phosphate minerals precipitated directly from seawater. However, Sr isotope ratios are susceptible to diagenesis, during which these min- erals are commonly subjected to loss of total Sr and input of radiogenic strontium from the detrital sediment fraction or meteoric ﬂuids. Hence, constructing the seawater Sr curve (Fig. 9.11) and using Sr isotopes as a correlation and dating tools requires the selection of high quality samples and a combination of petrographic and geochemical screening for the Chapter 9: Chemostratigraphy 92
Figure 9.10: The simpliﬁed global strontium cycle.
aﬀects of diagenesis (McArthur et al., 2012).
0.7070 Sr/ Sr
0.7050 NEOPROTEROZOIC PALEOZOIC MESOZOIC Cenozoic Tonian Cryogenian Ediacaran Camb Ord Sil Devon Carbon Perm Trias Jura Cretaceous 1000 800 600 400 200 0 Millions of years ago
Figure 9.11: A LOWESS ﬁt Sr isotope curve for the Ordovician to present (McArthur et al., 2012), plus the Sr isotope ﬁt to data for the Cambrian and Neoproterozoic, adapted from Shields and Veizer (2002) and (Halverson et al., 2010). See entry on the Seawater Sr Curve for a more detailed discussion of this record and its application to chronostratigraphy.
The secular evolution of seawater 87Sr/86Sr (Fig. 9.11) is conventionally interpreted in terms of competing inﬂuence of continental weathering (riverine) and seaﬂoor hydrothermal Sr inputs. However, other factors also inﬂuence the long-term evolution of seawater Sr Chapter 9: Chemostratigraphy 93 isotope ratios, including the progressive decay of 87Rb, which imparts a long-term rise in 87Sr/86Sr, and both tectonic (e.g. rifting or mountain building) and climatic eﬀects. The result is a seawater Sr isotope curve with structure at multiple diﬀerent time scales (Burke et al., 1982). Strontium isotope stratigraphy is most useful as a correlation and dating tool during intervals where 87Sr/86Sr changes rapidly (DePaolo and Ingram, 1985), such as across the Permo-Triassic boundary and in the latter half of the Cenozoic (Figs. 9.11, 9.13). Due to the considerable ﬂuctuations in 87Sr/86Sr over the past billion years (McArthur et al., 2012, Fig. 9.11), obtaining a single, precise and accurate 87Sr/86Sr value from a marine sedimentary rock is not suﬃcient to establish a date. Additional age or chemostratigraphic constraints or a distinct 87Sr/86Sr proﬁle are necessary to link to the composite strontium reference curve (Fig. 9.11). However, where this connection can be made, strontium isotope stratigraphy can be used in many cases to establish relatively precise (<1 m.y.) ages (McArthur et al., 2012).
The strontium isotope seawater curve (1.0–0 billion years) Since (Burke et al., 1982) published their record of Phanerozoic seawater 87Sr/86Sr, a great deal of work has gone into both developing high resolution records for speciﬁc intervals of the Phanerozoic (e.g. Koepnick et al., 1990; Fig. 9.12) and in producing compilations of the highest quality and best dated samples (Veizer et al., 1999; McArthur et al., 2001). The most comprehensive compilation of Phanerozoic 87Sr/86Sr data is found in the database of (McArthur et al., 2001), which has been progressively updated (e.g., McArthur et al., 2004, 2012), and now comprises over 4000 data points from some 50 sources (McArthur et al., 2012). The construction of such a record requires the application of age models to speciﬁc data sets, the integration of these data sets, and a technique for ﬁtting a curve to the data, which can then be used to calculate an age and appropriate error. For the relatively well-dated Phanerozoic record (Fig. 9.11), age models are reasonably robust and there exist high-resolution records, resulting in typical 95% conﬁdence intervals about the best-ﬁt curve of 0.00001 to 0.00004 (McArthur et al., 2012).
The most complete database for Precambrian strontium isotopes is from Shields and Veizer (2002). However, the age constraints on Precambrian samples are deﬁcient, and the quality of data is lower and more diﬃcult to verify than for the Phanerozoic. Here, we include only data from 1.0 billion years onward because the pre-1.0 Ga record is poor and shows only limited change, whereas the NeoproterozoicCambrian (1000–485 Ma) Sr isotope record has been extensively studied and shows signiﬁcant variability. Because the resolution and quality of the data from this time interval is low, rigorous ﬁtting techniques as used for the Phanerozoic record are not yet applicable.
The Proterozoic-Phanerozoic strontium isotope seawater curve (Fig. 9.11) reveals many striking features. The ﬁrst order trend is one of increasing 87Sr/86Sr, which results from the progressive decay of 87Rb (Wickman, 1948; Shields, 2007). The second order pattern is one of peaks at present and c. 500 million years ago and troughs at about 180 and 1000 million years ago, which may reﬂect the cycle of supercontinental break-up and assembly Halversonetal07. Superimposed on this pattern are prominent ﬂuctuations, some by as much as 0.001, over intervals of a few to 25 million years, many of which closely correspond to tectonic events. Chapter 9: Chemostratigraphy 94
The main parameters controlling the seawater Sr isotope composition are the magnitude of the seaﬂoor hydrothermal ﬂux (both mid-ocean ridge volcanism and oﬀ-axis alteration of the seaﬂoor) and continental weathering and their respective 87Sr/86Sr values. The composition of the seaﬂoor ﬂux ( 0.7037; Banner, 2004) is approximately ﬁxed by mantle composition, whereas the 87Sr/86Sr of continental weathering is presently ≈0.712, but presumably varied considerably through time. Because the residence time of Sr in the oceans is ∼5 million years and there are no plausible mechanisms to increase or decrease the ﬂux into or out of the oceans rapidly, changes in 87Sr/86Sr are steady state (i.e. gradual; see also entry on Marine Isotope Stratigraphy). In principal, changes in any one of the four major inﬂux parameters can drive a change in seawater 87Sr/86Sr, but due to the large isotopic diﬀerence between the hydrothermal and continental weathering ﬂuxes, many researchers invoke changes in the relative contribution of these two ﬂuxes to explain changes in seawater 87Sr/86Sr. On the other hand, although the continental ﬂux is buﬀered by the weathering of marine carbonates, the huge variations between diﬀerent rivers (Gaillardet et al., 1999) suggests that the continental weathering ﬂux should have changed substantially through time as a result of paleogeographic evolution of the continents and paleoclimate.
Use as a geochronological tool Given a reference seawater 87Sr/86Sr curve, dating a marine rock or sediment of poorly constrained age may be accomplished by comparing its strontium isotope composition to a reference curve. However, the eﬀectiveness of this tool is contingent upon the age and available age constraints on that sample, whether or not there are strontium isotope data from overlying or overlying strata, and the robustness of the curve. For example, a marine carbonate sample that is known only to be Jurassic in age and has an 87Sr/86Sr value of 0.70720 could equally be Pleinsbachian (190.8–182.7 Ma), Toarcian (182.7–174.1 Ma), Bojacian (170.3–168.3 Ma), or Tithonian (152.1–145.0 Ma) in age (Fig. 3A). However, if the stage of the sample is known through biostratigraphy or other chronostratigraphic constraints, then the age can be speciﬁed more precisely (Fig. 9.12)), with the precision being a measure of the conﬁdence interval for the Sr isotope curve at that age (Fig. 3) and the uncertainty in the age model upon which the reference Sr isotope curve is constructed (McArthur et al., 2012).
In the absence of additional age constraints that eliminate some of the possible matches between the sample and the reference curve, it helps to match both the absolute value and stratigraphic (i.e. temporal) trends. For example, if multiple samples from diﬀerent stratigraphic levels are analyzed from the section containing the original sample of interest, then these can be used to determine the trajectory in 87Sr/86Sr. If the samples show a trend of decreasing 87Sr/86Sr up-section (that is, in progressively younger rocks), then this pattern eliminates either a Toarcian or Berriasian age, both of which correspond to intervals of increasing 87Sr/86Sr (Fig. 9.12a) Whereas the slope of the declining trends (d[87Sr/86Sr]/dt) diﬀers between the Pleinsbachian and Bajocian, it would be virtually impossible to use the distinction alone in a poorly dated section of strata to discriminate between these two ages. Rather, a more complete 87Sr/86Sr proﬁle that captured, for example, the late Pleinsbachian trough (0.707075) or the continued decline to 87Sr/86Sr values lower than 0.707075 in the Bathonian (Fig. 9.12b), would be required. Alternatively, additional chemostratigraphic data, such as carbonate carbon isotopes or sulfate sulfur isotopes (see entry on Marine Isotope Stratigraphy) or magnetic polarity data may provide Chapter 9: Chemostratigraphy 95
0.7078 A 0.7073 B 0.70725
0.70705 Sr 170.0 169.8 169.6 169.4 169.2 169.0 86 Sr/
0.7068 Sinemurian Pleinsbach. Toarcian Aal. Bj. Bt. Cal. Oxford. Kimm. Tithonian Berrias. 200 190 180 170 160 150 140 Millions of Years Ago
Figure 9.12: A. The LOWESS ﬁt Sr isotope curve for the time interval 200–140 million years ago (Jurassic stages in blue and grey: Aal. = Aalenian; Bj. = Bajocian; Bt. = Bathonian; Cal. = Callovian; Kimm. = Kimmeridgian. Cretaceous Berriasian stage is in green), from McArthur et al. (2012). The red line is the LOWESS ﬁt, and the blue lines mark the upper and lower 95% conﬁdence intervals. Dashed line at 87Sr/86Sr = 0.70720 demonstrates the non-uniqueness of this value during the Jurassic, where it occurs during four distinct stages. B. Inset is a zoom-in of the interval from 170–169 million years. If a sample with a Sr isotope composition of 0.70720 is known to be Bajocian in age (e.g. based on biostratigraphy or other chemostratigraphic data), then an age can be determined by comparison with the reference curve. In this case, a value of 0.70720 yields an age of 169.60 ± ∼0.04 Ma (note that the error does not factor in uncertainty related to the calibration of the time scale).
suﬃcient constraints to narrow down the appropriate age for the 0.70720 sample.
9.5.3 Osmium isotopes Osmium isotope ratios, which owe their heterogeneity and utility to ongoing radioactive decay of 187Re to 187Os, are commonly reported as 187Os/188Os ratios. Osmium has a residence time in the modern oceans of 10–50 thousand years, and its isotopic composition (187Os/188Os = 1.06 ± 0.005) reﬂects the balance between an unradiogenic mantle compo- nent (187Os/188Os≈ 0.13) and a radiogenic continental component (187Os/188Os≈ 1.4), with a small contribution from the dissolution of unradiogenic cosmic dust Peucker-Ehrenbrink and Ravizza (2000, 2012). Hydrogenous osmium is removed from seawater dominantly in metaliferous and organic-rich sediments. Whereas the former serves as an important Chapter 9: Chemostratigraphy 96 archive for Cenozoic Os isotope ratios (Fig. 9.13), organic-rich sediments are important in determining pre-Cenozoic 187Os/188Os. However, because organic-rich sediments also concentrate Re, this requires an isochron approach to measure initial 187Os/188Os ratios.
As seen in Figure 9.13, the Cenozoic marine Os isotope record broadly resembles the Sr isotope record with both displaying a prominent, ﬁrst-order rise towards more radiogenic ratios. However, due to the much shorter residence time of Os in seawater, it is sensitive to certain perturbations where the Sr isotope system is buﬀered. Speciﬁcally, prominent low 187Os/188Os excursions occur across the CretaceousPaleogene and Eocene-Oligocene boundaries. The sensitivity of the Os isotope system to events such as ﬂood basalt volcan- ism (Cohen and Coe, 2002; Turgeon and Creaser, 2008), abrupt warming (Ravizza et al., 2001; Schmitz et al., 2004), and large meteorite impacts (Paquay et al., 2008) highlight its importance as a chronostratigraphic tool. For example, Ravizza and Peucker-Ehrenbrink (2003) used Os isotope stratigraphy on deep sea cores to resolve the timing of Deccan Trap volcanism from the Cretecous-Paleogene boundary, and hence, link the main mass extinc- tion event to the Chicxulub bolide impact. Similarly, to the extent that events that were large enough to leave an imprint in the seawater Os isotope record are well dated, then the Os isotope system provides an indirect but robust way of dating sedimentary records. Chapter 9: Chemostratigraphy 97
0.7086 Sr 86 0.7082 Sr/
87 1.2 0.7078 1.0 187 0.7074 Os/ 0.8 188
Cret. Paleocene Eocene Oligocene Miocene Pli. Q 65 55 45 35 25 15 5 Age (Millions of year before present)
Figure 9.13: Compilations of seawater 87Sr/86Sr and 187Os/188Os data spanning from 70 million years ago until the present, replotted from data in Misra and Froelich (2012). Note the short-lived anomalies in 187Os/188Os across the Cretaceous-Paleogeone and Eocene- Oligocene boundaries (dashed vertical lines), which are not mirrored in the 87Sr/86Sr record. Bibliography
Allen, P. A., Allen, J. R., 2005. Basin Analysis: Principles and Applications, 2nd Edition. Blackwell. and, S. A. B., 1998. U/Pb zircon heochronology and tempo of the rnd-Permian mass extinction. Science 280, 1039–1045. Appleby, P. G., 2002. Chronostratigraphic techniques in recent sediments. In: Tracking Environmental Change Using Lake Sediments. Springer, pp. 171–203. Banner, J. L., 2004. Radiogenic isotopes: systematics and applications to earth surface processes and chemical stratigraphy. Earth-Science Reviews 65, 141–194. Berger, W. H., Killingley, J. S., Vincent, E., 1978. Stable isotopes in deep-sea carbonates: Box core ERDC-92, west equatorial Paciﬁc. Oceanological Acta 1, 203–216. Bjorlykke, K., 2010. Well logs: A brief introduction. Springer-Verlag, Berlin-Heidelberg, Ch. 16, pp. 361–363. Blackburn, T. J., Olsen, P. E., Bowring, S. A., McLean, N. M., Kent, D. V., Puﬀer, J., McHone, G., Rasbury, E. T., Et-Touhami, M., 2013. Zircon U-Pb Geochronology Links the End-Triassic Extinction with the Central Atlantic Magmatic Province. Science 340, 941–944. Boulila, S., Galbrun, B., Hinnov, L. A., Collin, P.-Y., Ogg, J. G., Fortwengler, D., Marc- hand, D., 2010. Milankovitch and sub-milankovitch forcing of the oxfordian (late juras- sic) terres noires formation (se france) and global implications. Basin Research 22 (5), 717–732. Boulila, S., Galbrun, B., Huret, E., Hinnov, L. A., Rouget, I., Gardin, S., Bartolini, A., 2014. Astronomical calibration of the toarcian stage: implications for sequence stratig- raphy and duration of the early toarcian oae. Earth and Planetary Science Letters 386, 98–111. Burke, W. M., Denison, R. E., Hetherington, E. A., Koepnik, R. B., Nelson, M., Omo, J., 1982. Variations of seawater 87Sr/86Sr throughout Phanerozoic shales. Geology 10, 516–519. Canﬁeld, D. E., 1998. A new model for Proterozoic ocean chemistry. Nature 396, 450–452. Catuneanu, O., 2006. Principles of Sequence Stratigraphy. Elsevier. Christie-Blick, N., Grotzinger, J. P., von der Bosch, C. C., 1988a. Sequence stratigraphy in Proterozoic successions. American Journal of Science 290, 295–332.
98 Chapter 9: Chemostratigraphy 99
Christie-Blick, N., Grotzinger, J. P., von der Bosch, C. C., 1988b. Sequence stratigraphy in Proterozoic successions. Geology 16, 100–104.
Cohen, A. S., Coe, A. L., 2002. New geochemical evidence for the onset of volcanism in the Central Atlantic Magmatic Province and environmental change at the Triassic-Jurassic boundary. Geology 30, 267–270.
Cramer, B. S., Toggweiler, J. R., Wright, J. D., Katch, M. E., Miller, K. G., 2009. Ocean overturning since the Late Cretaceous: Inferences from a new benthic foraminiferal iso- tope compilation. Paleoceanography 24, doi:10.1029/2008PA001683.
Deepthy, R., Balakrishnan, S., 2005. Climatic control on clay mineral formation: Evidence from weathering proﬁles developed on either side of the Western Ghats. Journal of Earth System Science 114, 545–556.
DePaolo, D. J., Ingram, B. L., 1985. High-resolution stratigraphy with strontium isotopes. Science 227, 938–941.
DePaolo, D. J., Wasserburg, G. J., 1976. Nd isotopic variations and petrogenetic models. Geophysical Research Letters 3, 249–252.
Derry, L. A., 2010. A burial diagenesis origin for the Ediacaran-Shuram Wonoka carbon isotope anomaly. Earth and Planetary Science Letters 294, 152–162.
Drummond, C. N., Wilkonson, B. H., 1993. Aperiodic accumulation of cyclic peritidal carbonate. Geology 21, 1023–1026.
Dumas, S., Arnott, R. W. C., 2006. Origin of hummocky and swaley cross-stratiﬁcation-The controlling inﬂuence of unidirectional current strength and aggradation rate. Geology 34, 1073–1076.
Edwards, M., 1986. Glacial environments. In: Reading, H. G. (Ed.), Sedimentary Environ- ments and Facies. Blackwell, pp. 445–470.
Embry, A. F., Johannessen, E. P., 1992. T-R sequence stratigraphy, facies analysis and reservoir distribution in the uppermost Triassic-Lower Jurassic succession, western Sver- drup Basin, Arctic Canada. In: Vorrn, T. O., Bergsager, E., Dahl-Stamnes, O. A., Holter, E., Johansen, B., Lie, E., Lund, T. B. (Eds.), Arctic Geology and Petroleum Potential. Vol. 2 of Special Publications. Norwegian Petrolum Society, pp. 121–146.
Emiliani, C., 1955. Pleistocene temperatures. The Journal of Geology 63, 538–578.
Emiliani, C., 1958. Paleotemperature analysis of Core 280 and Pleistocene glaciations. The Journal of Geology 66, 264–275.
Epstein, S., Buchsbaum, R., Lowenstam, H., Urey, H. C., 1951. Carbonate-water isotopic temperature scale. Geological Society of America Bulletin 62, 417–426.
Epstein, S., Mayeda, T., 1953. Variation in O18 content of waters from natural sources. Geochimica et Cosmochimica Acta, 213–224. et al., O. C., 2009. Towards a standardization of sequence stratigraphy. Earth-Science Reviews 92, 1–33. Chapter 9: Chemostratigraphy 100
Fairbanks, R. G., Mortlock, R. A., Chiu, T.-C., Cao, L., Kaplan, A., Guilderson, T. P., Fairbanks, T. W., Bloom, A. L., Grootes, P. M., Nadeau, M.-J., 2005. Radiocarbon calibration curve spanning 0 to 50,000 years bp based on paired¡ sup¿ 230¡/sup¿ th/¡ sup¿ 234¡/sup¿ u/¡ sup¿ 238¡/sup¿ u and¡ sup¿ 14¡/sup¿ c dates on pristine corals. Quaternary Science Reviews 24 (16), 1781–1796.
Frank, M., O’Nions, R. K., Hein, J. R., Banakar, V. K., 1999. 60 Myr records of major elements and Pb–Nd isotopes from hydrogenous ferromanganese crusts: Reconstruction of seawater paleochemistry. Geochimica et Cosmochimica Acta 63, 1689–1708.
Gaillardet, J., Dupr´e,B., Louvat, P., Allegre, C., 1999. Global silicate weathering and co¡ sub¿ 2¡/sub¿ consumption rates deduced from the chemistry of large rivers. Chemical Geology 159 (1), 3–30.
Galloway, W. E., 1989. Genetic stratigraphic sequences in basin analysis; I, Architecture and genesis of ﬂooding-surface bounded depositional units. American Association of Petroleum Geologists Bulletin 73, 125–142.
Ginsburg, R. N., 1971. Landward movement of carbonate mud—new model for regressive cycles in carbonates. American Association of Petroleum Geologists Bulletin 55, 340.
Gluyas, J., Swarbrick, R., 2004. Petroleum Geoscience. Blackwell.
Goldhammer, R. K., Lehmann, P. J., Dunn, P. A., 1993. The origin of high-frequency plat- form carbonate cycles and 3rd-order sequences (Lower Ordovician El-Paso Gp, West Texas)—constraints from outcrop data and stratigraphic modeling. Journal of Sedimen- tary Petrology 63, 318–359.
Gradstein, F. M., 2012. Introduction. In: Gradstein, F. M., Ogg, J. O., Schmitz, M., Ogg, G. (Eds.), The Geological Time Scale 2012. Elsevier, Ch. 1, pp. 1–29.
Gradstein, F. M., Ogg, J. G., Schmitz, M. D., Ogg, G. M., 2012. The Geological Time Scale 2012. Elsevier.
Gradstein, F. M., Ogg, J. G., Schmitz, M. D., Ogg, G. M. (Eds.), 2013. The Geological Time Scale 2012. Elsevier, Ch. 13, pp. 239–267.
Gradstein, F. M., Ogg, J. O., Smith, A. G. (Eds.), 2004. A Geological Time Scale 2004. Cambridge University Press.
Grossman, E. L., 2012. Oxygen isotope stratigraphy. In: Gradstein, F. M., Ogg, J. O., Schmitz, M., Ogg, G. (Eds.), The Geological Time Scale 2012. Elsevier, pp. 181–206.
Grotzinger, J. P., 1986. Cyclicity and paleoenvironmental dynamics, Rocknest platform, northwest Canada. Geological Society of America Bulletin, 1208–1231.
Halverson, G. P., Wade, B. P., Hurtgen, M. T., Barovich, K., 2010. Neoproterozoic chemostratigraphy. Precambrian Research 182, 337–350.
Hathorne, E. C., James, R. H., 2006. Temporal record of lithium in seawater: a tracer for silicate weathering? Earth and Planetary Science Letters 246, 393–406. Chapter 9: Chemostratigraphy 101
Hayes, J. M., Strauss, H., Kaufman, A. J., 1999. The abundance of 13C in marine organic matter and isotopic fractionation in the global biogeochemical cycle of carbon during the past 800 Ma. Chemical Geology 161, 103–125.
Hays, J. D., Imbrie, J., Shackleton, N. J., 1976. Variations in Earth’s orbit: Pacemaker of the ice ages. Science 194, 1121–1131.
Higgins, J. A., Schrag, D. P., 2006. Beyond methane: Towards a theory for the Paleocene- Eocene Thermal Maximum. Earth and Planetary Science Letters 245, 523–537.
Hilgen, F. J., Kuiper, K., Krijgsman, W., Snel, E., van der Laan, E., 2007. Astronomical tuning as the basis for high resolution chronostratigraphy: the intricate history of the Messinian salinity crisis. Stratigraphy 4, 231–238.
Hill, J., Curtis, R. W. A., d. M. Tetzlaﬀ, 2012. Preservation of forcing signals in shallow water carbonate sediments. Sedimentary Geology 275-276, 79–92.
Hinnov, L. A., Hilgen, F. J., 2012. Cyclostratigraphy and astrochronology. In: Gradstein, F. M., Ogg, J. O., Schmitz, M., Ogg, G. (Eds.), The Geological Time Scale 2012. Elsevier, Ch. 4, pp. 63–83.
Hoﬀman, P. F., 2011. Strange bedfellows: glacial diamictite and cap carbonate from the Marinoan (635 Ma) glaciation in Namibia. Sedimentology 58, 57–119.
Holmden, C., Creaser, R. A., Muehlenbachs, K., Leslie, S. A., Bergstrom, S. M., 1998. Isotopic evidence for geochemical decoupling between ancient epeiric seas and bordering oceans: Implications for secular curves. Geology 26, 567–570.
Hughen, K. A., Lehman, S., Southon, J., Overpeck, J., Marchal, O., Herring, C., Turnbull, J., 2004. 14C activity and global carbon cycle changes over the past 50,000 years. Science 303, 202–207.
Kampschulte, A., Strauss, H., 2004. The sulfur isotopic evolution of Phanerozoic seawater based on the analysis of structurally substituted sulfate in carbonates. Chemical Geology 204, 255–286.
Kendall, B., Creaser, R. A., Calver, C. R., Raub, T. D., Evans, D. A. D., 2009. Correlation of Sturtian diamictite successions in southern australia and northwestern Tasmania by Re-Os black shale geochronology and the ambiguity of ”Sturtian”-type diamictite-cap carbonate pairs as chronostratigraphic marker horizons. Precambrian Research 172, 301– 310.
Kendall, B., Creaser, R. A., Selby, D., 2006. Re-Os geochronology of postglacial black shales in Australia: Consequences for timing of the Sturtian glaciation. Geology, 729–732.
Kendall, B. S., Creaser, R. A., Ross, G. M., Selby, D., 2004. Constraints on the timing of Marinoan snowball Earth glaciation by 187Re-187Os dating of a Neoproterozoic, post- glacial black shale in western Canada. Earth and Planetary Science Letters 222, 729–740.
Kisak¨urek,B., James, R. H., Harris, N. B. W., 2005. Li and δ7Li in Himalayan rivers: Proxies for silicate weathering? Earth and Planetary Science Letters 237, 387–401. Chapter 9: Chemostratigraphy 102
Knoll, A. H., Hayes, J. M., Kaufman, A. J., Swett, K., Lambert, I. B., 1986. Secular variation in carbon isotope ratios from Upper Proterozoic successions of Svalbard and east Greenland. Nature 321, 832–837.
Krijgsman, W., Fortuin, A. R., Hilgen, F. J., Sierro, F. J., 2001. Astrochronology for the Messinian Sorbas basin (SE Spain) and orbital (precessional) forcing for evaporite cyclicity. Sedimentary Geology 140, 43–60.
Krogh, T. E., 1982. Improved accuracy of U-Pb zircon ages by the creation of more con- cordant systems using an air abrasion technique. Geochimica et Cosmochimica Acta 46, 637–649.
Kump, L. R., Arthur, M. A., 1999. Interpreting carbon-isotope excursions: carbonates and organic carbon. Chemical Geology 161, 181–198.
Kunzmann, M., Halverson, G. P., Macdonald, F. A., Hodgskiss, M., Sansjofre, P. D., Schumann, D., Rainbird, R. H., 2014. The early Neoproterozoic Chandindu Formation of the Fifteenmile Group in the Ogilvie Mountains. In: McFarland, K. (Ed.), Yukon Exploration and Geology 2013. Yukon Geological Survey, Whitehorse, p. in press.
Langereis, C. G., Krijgsman, W., Muttoni, G., Menning, M., 2010. Magnetostratigraphy— concepts, deﬁnitions, and applications. Newsletters on Stratigraphy 43, 207–233.
Laskar, J., Robutel, P., Joutel, F., Gastineau, M., Correia, A., Levrard, B., et al., 2004. A long-term numerical solution for the insolation quantities of the earth. Astronomy & Astrophysics 428 (1), 261–285.
Lewis, K. W., Aharonson, O., Grotzinger, J. P., Kirk, R. L., McEwen, A. S., Suer, T.- A., 2008. Quasi-periodic bedding in the sedimentary rock record of Mars. Science 322, 1532–1535.
Lisieki, L. E., Raymo, M. E., 2005. A Pliocene-Pleistocene stack of 57 globally distributed benthic d18O. Paleoceanography 20, doi:10.1029/2004PA001071.
Manzi, V., Gennari, R., Hilgen, F., Krijgsman, W., Lugli, S., Roveri, M., Sierro, F. J., 2013. Age reﬁnement of the messinian salinity crisis onset in the mediterranean. Terra Nova 25 (4), 315–322.
Mattinson, J. M., 2005. Zircon U-Pb chemical abrasion (“CA-TIMS”) method: Combined annealing and multi-step partial dissolution analysis for improved precision and accuracy of zircon ages. Chemical Geology 220, 47–66.
McArthur, J., Banerjee, D., Hudson-Edwards, K., Mishra, R., Purohit, R., Ravenscroft, P., Cronin, A., Howarth, R., Chatterjee, A., Talukder, T., et al., 2004. Natural organic matter in sedimentary basins and its relation to arsenic in anoxic ground water: the example of west bengal and its worldwide implications. Applied Geochemistry 19 (8), 1255–1293.
McArthur, J., Howard, R. J., Shields, G. A., 2012. Strontium isotope stratigraphy. In: Gradstein, F. M., Ogg, J. O., Schmitz, M., Ogg, G. (Eds.), The Geological Time Scale 2012. Elsevier, pp. 127–144. Chapter 9: Chemostratigraphy 103
McArthur, J., Howarth, R., Bailey, T., 2001. Strontium isotope stratigraphy: Lowess ver- sion 3: best ﬁt to the marine sr-isotope curve for 0–509 ma and accompanying look-up table for deriving numerical age. The Journal of Geology 109 (2), 155–170.
Milankovitch, M., 1941. Kanon der Erdebestrahlung und seine anwendung auf das eiszeit- enproblem. K¨oniglich Serbische Akademie.
Milliman, J. D., 1993. Production and accumulation of calcium carbonate in the ocean: Budget of a nonsteady state. Global Biogeochemical Cycles 7, 927–957.
Misra, S., Froelich, P. N., 2012. Lithium isotope history of Cenozoic seawater: Changes in silicate weathering and reverse weathering. Science 335, 818–823.
Mondol, N. H., 2010. Seismic exploration. Springer-Verlag, Berlin-Heidelberg, Ch. 17, pp. 375–402.
Moucha, R., Forte, A. M., Mitrovica, J. X., Rowley, D. B., Qu´er´e,S., Simmons, N. A., Grand, S. P., 2008. Dynamic topography and long-term sea-level variations: There is no such thing as a stable continental platform. Earth and Planetary Science Letters, 101–108.
Mundil, R., Metcalfe, I., Ludwig, K. R., Renne, P. R., Oberli, F., Nicoll, R. S., 2001. Timing of the Permian-Triassic biotic crisis: implications from new zircon U/Pb age data (and their limitations). Earth and Planetary Science Letters 187, 131–145.
Nagy, J., Bjorlykke, K., 2010. Petroleum Geoscience: From Sedimentary Environments to Rock Physics. Springer-Verlag, Ch. 7, pp. 213–224.
Ogg, J. G., 2013. Geomagnetic polarity time scale. In: Gradstein, F. M., Ogg, J. O., Schmitz, M., Ogg, G. (Eds.), The Geological Time Scale 2012. Elsevier, Ch. 5, pp. 85–113.
Panchuk, K. M., Holmden, C., Kump, L. R., 2005. Sensitivity of the epeiric sea carbon isotope record to local-scale carbon cycle processes: Tales from the Mohawkian sea. Palaeogeography, Palaeoclimatology, Palaeoecology 228, 320–337.
Paquay, F. S., Ravizza, G. E., Dalai, T. K., Peucker-Ehrenbrink, B., 2008. Determining chondritic impactor size from the marine osmium isotope record. Science 320, 214–218.
Paytan, A., Gray, E. T., 2012. Sulfur isotope stratigraphy. In: Gradstein, F. M., Ogg, J. O., Schmitz, M., Ogg, G. (Eds.), The Geological Time Scale 2012. Elsevier, pp. 167–180.
Paytan, A., Mart´ınez-Ruiz, F., Eagle, M., Ivy, A., Wankel, S. D., 2004. Using sulfur isotopes in barite to elucidate the origin of high organic matter accumulation events in marine sediments. In: Sulfur Biogeochemistry. Vol. 379 of GSA Special Paper. Geological Society of America, pp. 151–160.
Peucker-Ehrenbrink, B., Ravizza, G., 2000. The marine osmium isotope record. Terra Nova 12, 205–219.
Peucker-Ehrenbrink, B., Ravizza, G., 2012. Osmium isotope stratigraphy. In: Gradstein, F. M., Ogg, J. O., Schmitz, M., Ogg, G. (Eds.), The Geological Time Scale 2012. Elsevier, pp. 145–166. Chapter 9: Chemostratigraphy 104
Popp, B. N., Laws, E. A., Bidigare, R. R., Dore, J. E., Hanson, K. L., Wakeham, S. G., 1998. Eﬀect of phytoplankton cell geometry on carbon isotopic fractionation. Geochimica et Cosmochimica Acta 62, 69–77. Ravizza, G., Peucker-Ehrenbrink, B., 2003. Chemostratigraphic evidence of Deccan vol- canism from the marine osmium isotope record. Science 302, 1392–1395. Ravizza, G. E., Norris, R. N., Blusztajn, J., Aubry, M.-P., 2001. An osmium isotope excursion associated with the late Paleocene thermal maximum: Evidence of intensiﬁed chemical weathering. Paleoceanography 16, 155–163. Rees, C. E., Jenkins, W. F., Monster, J., 1978. The sulphur isotopic composition of ocean water sulphate. Geochimica et Cosmochimica Acta 42, 377–382. Rooney, A. D., Macdonald, F. A., Strauss, J. V., Dud´as,F. O., Hallmann, C., Selby, D., 2014. Re-OS geochronology and coupled Os-Sr isotope constraints on the Sturtian snowball Earth. Proceedings of the National Academy of Sciences 111, 51–56. Rothman, D. H., Grotzinger, J. P., Flemings, P., 1994. Scaling in turbidite deposition. Journal of Sedimentary Research A64, 59–67. Saltzman, M. R., Thomas, E., 2012. Carbon isotope stratigraphy. In: Gradstein, F. M., Ogg, J. G., Schmitz, M., Ogg, G. (Eds.), The Geological Time Scale 2012. Elsevier, pp. 207–232. Schmitz, B., Peucker-Ehrenbrink, B., Heilmann-Clausen, C., Aberg, G., Asaro, F., Lee, T. A., 2004. Basaltic explosive volcanism, but no comet impact, at the Paleocene-Eocene boundary: High-resolution chemical and isotopic records from Egypt, Spain and Den- mark. Earth and Planetary Science Letters 225, 1–17. Schmitz, M. D., 2013. Radiogenic isotope geochronology. In: Gradstein, F. M., Ogg, J. O., Schmitz, M., Ogg, G. (Eds.), The Geological Time Scale 2012. Elsevier, Ch. 6, pp. 114–126. Schmitz, M. D., Kuiper, K. F., 2013. High-precision geochronology. Elements 9, 25–30. Schoene, B., Condon, D. J., Morgan, L., McLean, N., 2013. Precision and accuracy in geochronology. Elements 9, 19–24. Scholle, P. A., Arthur, M. A., 1980. Carbon isotope ﬂuctuations in Cretaceous pelagic limestone: potential stratigraphic and petroleum exploration tool. American Association of Petroleum Geologists Bulletin 64, 67–87. Selley, R. C., 1998. Elements of Petroleum Geology, 2nd Edition. Academic Press, London. Shackleton, N. J., 2000. The 100,000-year ice-age cycle identiﬁed and found to lag temper- ature, carbon dioxide, and orbital eccentricity. Science 289, 1897–1902. Shackleton, N. J., Opdyke, N. D., 1973. Oxygen isotope and palaeomagnetic stratigraphy of equatorial Paciﬁc core V28-238: Oxygen isotope temperatures and ice volumes on a 105 year and 106 year scale. Quaternary Research 3, 39–55. Shields, G., Veizer, J., 2002. Precambrian marine carbonate isotope database: Version 1.1. Geochemistry, Geophysics, Geosystems 3, 10.1029/2001GC000266. Chapter 9: Chemostratigraphy 105
Shields, G. A., 2007. A normalised seawater strontium isotope curve: possible implications for Neoproterozoic-Cambrian weathering rates and further oxygenation of the Earth. eEarth 2, 35–42.
Swart, P. K., 2008. Global synchronous changes in the carbon isotopic composition of carbonate sediments unrelated to changes in the global carbon cycle. Proceedings of the National Academy of Sciences, 13741–13745.
Tachikawa, K., Jeandel, C., Roy-Barman, M., 1999. A new approach to the Nd residence time in the ocean: the role of atmospheric inputs. Earth and Planetary Science Letters 170, 433–446.
Turgeon, S. C., Creaser, R. A., 2008. Cretaceous oceanic anoxic event 2 triggered by a massive magmatic episode. Nature 454, 323–326.
Urey, H. C., Lowenstam, H. A., Epstein, S., McKinney, C. R., 1951. Measurement of paleotemperatures and temperatures of the Upper Crectaceous of England, Denmark, and the southeastern United States. Geological Society of America Bulletin 62, 399–416.
Vail, P. R., 1987. Seismic stratigraphy intepretation procedure. In: In Atlas of Seismic Stratigraphy. No. 27 in Studies in Geology. American Association of Petroleum Geolgists, pp. 1–10.
Van Wagoner, J. C., 1995. Overview of sequence stratigraphy of foreland basin deposits: terminology, summary of papers, and glossary of sequences stratigraphy. In: Van Wag- oner, J. C., Bertram, G. T. (Eds.), Sequence Stratigraphy of Foreland Basin Deposits. Vol. 64. American Association of Petroleum Geolgists, pp. ix–xxi.
Veizer, J., Ala, D., Azmy, K., Bruckschen, P., Buhl, D., Bruhn, F., Carden, G. A., Di- ener, A., Ebneth, S., Godderis, Y., et al., 1999. 87sr/86sr, δ13c and δ18o evolution of phanerozoic seawater. Chemical Geology 161 (1), 59–88.
Vigier, N., Decarreau, A., Millot, R., Carignan, J., Petit, S., France-Lanord, C., 2008. Quantifying Li isotope fractionation during smectite formation and implications for the Li cycle. Geochimica et Cosmochimica Acta 72, 780–792. von Strandmann, P. A. E. P., Jenkyns, H. C., Woodﬁne, R. G., 2013. Lithium isotope evidence for enhanced weathering during Oceanic Anoxic Event 2. Nature Geoscience 6, 668–672.
Weissert, H., Joachimski, M., Sarntheiin, M., 2008. Chemostratigraphy. Newsletters on Stratigraphy 42, 145–179.
Wickman, F. E., 1948. Isotope ratios: A clue to the age of certain marine sediments. Journal of Geology 56, 61–66.
Williams, D. F., Peck, J., Karabanov, E. B., Prokopenko, A. A., Kravchinsky, V., King, J., Kuzmin, M. I., 1997. Lake Baikal record of continental climate response to orbital insolation during the past 5 million years. Science 278, 1114–1117.
Williams, G. E., 1998. Precambrian tidal and glacial clastic deposits: implications for Precambrian Earth-Moon dynamics and palaeoclimate. Sedimentary Geology 120, 55– 74. Chapter 9: Chemostratigraphy 106
Wortmann, U. G., Paytan, A., 2012. Rapid variability of seawater chemistry over the past 130 million years. Science 337, 334–336.
Zachos, J., Pagani, M., Sloan, L., Thomas, F., Billups, K., 2001. Trends, rhythms, and aberrations in global climate 65 Ma to present. Science 292, 686–693.
Zachos, J. C., Bohaty, S. M., John, C. M., McCarren, H., Kelly, D. C., Nielsen, T., 2007. The Palaeocene-Eocene carbon isotope excursion: Constraints from individual shell planktonic foraminifer records. Philosophical Transactions of the Royal Society of London A365, 1829–1842.
Zhao, J., Yu, K., Feng, Y., 2009. High-precision 238U-234U-230Th disequilibrium dating of the recent past: a review. Quaternary Geochronology 4, 423–433.
Zoeller, L., Wagner, G. A., 2014. Luminescence dating, history. In: Rink, W. J., Thompson, J. (Eds.), Encyclopedia of Scientiﬁc Dating Methods. Earth Science Series. Springer.