In this article we shall look at how we can use sedimentology to determine the direction of paleocurrents: former currents of wind or water.
The reader may find it helpful at this point to go back and look at the main articles on rivers and on deserts to place the following article in its proper perspective.
It is possible to determine the past direction of currents of wind or water by examination of sedimentary structures in rocks.
Both ripples and dunes in stream deposits and in wind-formed sand-dunes have a shallower slope on the stoss side than the lee side, and progress in a leeward direction. This leaves the direction of the currents recorded in the cross-bedding preserved in the rocks formed from the sediments. For example in the photograph below we can see that the current that formed the cross-beds was traveling roughly from left to right.
In the case of streams and rivers, the current direction is necessarily downstream. In the case of sand, you might not suppose that there is a direction associated with the current: the wind is, after all, proverbially capricious. But it is not quite so capricious as a layperson might think. Pretty much any spot on the face of the earth has a prevailing, though not constant, direction in which the wind blows: for example in the Sahara the prevailing wind is north-easterly, and this is reflected by the motion of the dunes and the sets of cross-beds that they deposit.
On a larger scale, the direction in which a river flows is indicated by the average clast size on its bed: there is a tendency for smaller clasts to be found further downstream.
The reason for this is that if at any point in its course the river slacks off due to becoming broader and/or deeper, this will diminish its capacity to carry the larger of the clasts that it has carried up to that point, and so they will be deposited. The river may well regain speed after that point, so that it could carry clasts of that size once more; but by that time they have already been deposited further upstream.
How do we know?
We can verify that these relationships between currents, sedimentary structures, and clast size hold good in the present.
Now in this case the principle of actualism is if anything even more compelling than usual, because we are dealing here with really basic laws of physics. Is it even conceivable that in the past the laws of fluid dynamics could have been so different that (for example) the stoss side of a ripple would be steeper than the lee?
The agreement between current direction as inferred from ripples and as inferred from clast size is another reason for confidence in these methods.
In this article we shall look at how biogeography relates to climate.
Species and their preferred climates
Pretty much every species, with the noteable exception of our own, flourishes within a fairly narrow range of climatic conditions. This means that when we find representatives of that species preserved in the fossil record, we should conclude that those conditions were present at the date and location when and where they lived.
We can refine our knowlege of temperatures by looking at the overlap in the ranges of species. If one lives in regions with an average annual temperature between 20°C and 24°C , and another inhabits regions with an average annual temperature between 22°C and 26°C, then when we find them together we can conclude that they were living in a time and place where the average annual temperature was between 22 and 24°C.
How do we know?
In order for this method to work, we need to know what climatic conditions suited the species we find in the fossil record.
If the species is still extant, we can of course just look at where it lives and what the extremes of temperature are in that location. You might speculate that perhaps without any visible morphological change the species could have evolved to adapt to different climatic conditions, so that the temperature range it lived in then is not the same as that which it inhabits now. While this is possible in principle, what we observe is that species do not evolve to adapt to shifting climatic conditions; rather, they shift their locations to follow the conditions that suit them. (Even plant species will migrate; although an individual plant cannot move, the spread of their seeds allows them to migrate as a species.)
Species, of course, come and go over time; but we can also look at larger groups. For example, since out of about 2,600 species of palm trees we know of none that will grow in a cold climate without human intervention, it is reasonable to conclude that this would be also be true of an extinct species of palm tree.
Still, extinction does present us with a problem. For example, with no modern representatives at all, how are we meant to know a warm-water trilobite from a cold-water trilobite? This difficulty is not quite insurmountable in principle. For example, if we frequently find some species of trilobite in association with ooids, then we would be inclined to think that it was a warm-water species, and then if we were to find it in the absence of other climatic indicators, we would take the trilobite species as an indication of a warm climate. That being said, the use of the biogeographic method of assessing paleoclimatic conditions does in practice get less and less useful the further back into the past we wish to look.
In this article we shall discuss how the analysis of the shapes of leaves can be used to reconstruct past climates.
In the previous article, we saw that if we know the environment preferred by a species, we can use it as a climatic indicator. We also noted that this is difficult when the species is extinct, since then we can no longer find by direct observation what climate is congenial to it.
But what if the morphology of a species was an indicator of the climate it prefers? In that case, we would be able to tell what climate it inhabited just by looking at its remains in the fossil record. One morphological feature commonly used for this purpose is the leaf shapes of plants.
Plants in temperate climates tend to have leaves with serrated margins, i.e. they have jagged edges; plants in warmer and more humid climates tend to have what are known in botanical jargon as entire margins, that is, smooth and unserrated. The difference is illustrated in the photographs below.
Rather than there being a sharp cut-off between the temperate and tropical styles of leaves, there is a continuous relationship between the climate and the mix of leaf types found in it: that is, as the climate gets a little hotter and wetter, the proportion of entire margins increases a little. This means that looking at a single leaf doesn't tell us that we are looking at a temperate or tropical climate; but looking at a whole lot of species will allow us to do something a whole lot better than simply dividing climates into tropical or temperate: we can actually estimate the average annual temperature.
So once we have established the relationship between climate and the proportions of leaf shapes in the flora, we can use this relationship to find out about past climates where leaves have been preserved in the fossil record. Although leaves are not the most easily preserved material, there are some environments, such as lacustrine sediment, where they are preserved and can be used as a paleoclimatic indicator.
Other leaf characteristics
Using leaf shape as a climatic indicator began with counting the proportions of serrated to entire margins, it didn't end there. Other characteristics such as leaf length, leaf width, leaf area, the presence or absence of a pointed tip also have a relationship to the climate.
The CLAMP (Climate Leaf Analysis Multivariate Program) group has measured many more characteristics of leaves and have used sophisticated methods to produce a formula relating these characteristics to climate; their website can be found here.
This means that if we look at the ratio of leaf types in the fossil record, we can use it to produce an estimate of the temperature at the time when the plants grew.
Now, you might wonder if this relationship between leaf types and temperature really held in the past. Well, the adaptation of fauna and flora to their environment does seem to be one of the great constants of nature, so in principle we would expect the relationship to have held in the past much as it does in the present. It is after all extremely unlikely on the face of it that serrated margins would suit temperate climates in the present and yet have been more suitable to tropical climates in the past.
However, it should be noted that the exact relationship between leaf types and climate varies from continent to continent: although more serrated leaves always correspond to cooler climates on every continent, the exact numerical relationship is not the same. So a graph similar to the one above but based on Australian data rather than Asian data would still be a good approximation to a straight line sloping downwards to the right, but the exact angle of slope would be different. Now, what can vary from place to place could also in principle vary from time to time. This means that although leaf shape is a useful indicator of past climate, it would be naïve to treat it as though the relationship between climate and leaf shape was quite so exact as the relationship between temperature and a thermometer.
In this article we shall discuss the principles behind dendroclimatology. The reader may find it useful to re-read the article on dendrochronology before proceeding further.
How dendrochronology works
You should recall from the article on dendrochronology that many species of trees produce annual growth rings. The dendrochronological method depends crucially on the fact that these rings will not be of constant width, but will be thicker when the tree in question has had a good year and thinner in bad years.
Now, from the point of view of a tree, the difference between a good year and a bad year is mainly going to be determined by the weather: by the amount of sun and the amount of rainfall in its growing season.
So by using dendrochronology we can put a date on a tree-ring, and by looking at the thickness of the ring we can find out about the climate in the year of its formation. This method is known as dendroclimatology.
Difficulties of the method
Some of the difficulties of dendroclimatology are, so to speak, inherited from dendrochronology: for example, a shortage of material limits the scope of dendoclimatology in time. Also, since annual rings don't form in the tropics, the use of dendroclimatology, like dendrochronology, is limited to temperate zones.
One difficulty with dendroclimatology is that although the weather is the most likely cause of a tree having a bad year, it is not the only possible cause: disease, fire, or other external causes can have the same effect. This difficulty is not insuperable: trees affected by such events will be the odd ones out, so analysis of a sufficiently large number of trees from different locations allows us to distinguish the climatic signal from the incidental noise.
Another difficulty is that we would expect both the temperature and the amount of rainfall to affect the growth of trees. We would ideally like a proxy for one or the other, whereas when we measure the width of a growth ring we are looking at the combined effects of both. One way around this is to study what are known as limiting stands. For example, if we look at trees near the snowline of a mountain, then because storms tend to break over peaks, those trees will get all the water they need; the limiting factor on growth will be temperature, which will therefore determine the width of the growth rings.
How do we know?
We can compare the growth rings in trees with records of temperatures going back in some cases hundreds of years.
Mostly, there is agreement between direct measurement and climatological data. However, there is a discrepancy between the most recent records and recent tree ring growth. As the reader will doubtless be aware, global temperatures have been rising over the last few decades, something we know by consulting actual thermometers; but the dendroclimatological proxy has not kept pace with the change in climate. This leaves a question-mark hanging over dendroclimatology. If it doesn't tell us about events which we know have happened, then what isn't it telling us about events that we don't know about?
So the same kind of data that tells us that dendroclimatology is somewhat reliable also tells us that it isn't completely reliable. Together with the other difficulties with dendroclimatology, you might think that climatologists should abandon it and rely on proxies which are based on something less capricious than the growth of trees. However, it does have its advantages: it provides us with data specific to particular locations on land which are not covered by other proxies.
In this article we shall look at how the composition of shells of marine organisms is affected by climate, and how we can therefore obtain clues about past climates by chemical and isotopic analysis of their shells. This collection of methods is known as scleroclimatology.
The ratio of the stable isotopes oxygen isotopes 16O and 18O in water (H2O) can be shown to vary with temperature, with a higher 16O/18O ratio associated with lower temperatures.
Oxygen from the seawater is used by shell-forming marine organisms when they form their calcium carbonate shells, and as the two isotopes of oxygen are chemically identical, no isotope is preferentially involved in this process, with the result that the oxygen ratios in the shells reflect the oxygen ratios in the seawater.
This means, of course, that if we look at shells in the geological record (the tests of foraminiferans being a favorite choice for this) and if the calcium carbonate in the shells hasn't been replaced by some other mineral, and if we can put a date on the shells, then we can find the temperature of the seawater at the time and place where the shells were deposited: this proxy is known as δ18O. What's more, since it is possible to distinguish between the fossils of plantonic and benthic species, we get two figures for each location: the surface temperature and the temperature at the sea floor.
Mg/Ca and Sr/Ca
The elements magnesium (Mg) and strontium (Sr) lie in the same column of the periodic table as calcium (Ca) and so possess similar chemical properties. This means that magnesium and strontium can substitute for the calcium in calcium carbonate (CaCO3). As these elements substitute more readily for calcium at higher temperatures, the proportions of megnesium and strontium substituting for calcium can be used as a temperature proxy.
Difficulties of the method
Obviously for this method to work the mineral composition of the shells must not be changed by fossilization. While this can present us with difficulties in obtaining appropriate material, it will not usually lead us to produce erroneous results; after all, if a shell has undergone mineral replacement so that it's no longer made of calcium carbonate then this is not something a geologist would easily be able to overlook.
One thing that interferes with the oxygen isotope method is that, paradoxically, actual glaciation has the opposite effect on global marine oxygen ratios than mere low temperatures have on local isotope ratios. This is because when the glaciation of the Earth increases, water that evaporates from the seas is locked up in ice sheets; now 16O, being lighter than 18O, evaporates more readily, so that the result of increasing glaciation is a decrease in the 16O/18O ratio. To make sense of the data, it is necessary to disentangle these local and global effects by reference to other data.
When it comes to the Mg/Ca and Sr/Ca methods, the local mineral composition of the seawater can act as a confounding factor: for example, if the seawater in a particular location is particularly rich in magnesium for some reason, then this will also increase the Mg/Ca ratio.
How do we know?
We can take samples of shell-forming organisms from locations with a known temperate and measure their isotopic and chemical ratios, and see how they relate to temperature. It is also possible to grow shellfish in tanks kept artificially at a controlled temperature and see what happens.
It is unlikely that the biochemistry of shell formation has changed significantly since organisms first started forming calcium carbonate shells; and it is almost unthinkable that the physics of the evaporation of water has changed at all. It is reasonable to conclude that we can take chemical and isotopic ratios in the past as proxies for past temperatures.
Finally, we can note that there is a good if not exact correlation between the paleoclimatic data obtained from shells and other paleoclimatic proxies.
In this article we shall look at alkenones and how they can be used to construct the paleothermometer known by the somewhat cryptic name of Uk'37.
Alkenones are organic molecules having a chain-like structure, as organic molecules so often do. The length of the chain is given by the number of carbon atoms in it, so that we can speak of the alkenones C37, C38, etc. Note that these are not chemical formulas for the alkenones, which contain other atoms besides carbon; they merely record the length of the alkenone.
A further variation in the structure of alkenones is that they can be either doubly or triply unsaturated, a detail depending on the nature of their carbon bonds. So we can speak of the alkenones C37:2 and C37:3.
It is not really necessary for the reader to understand the organic chemistry of alkenones in any detail, or to know what "unsaturated" means; the important thing is that they exist and occur in different varieties.
Alkenones in nature
Only a small number of species are known to produce alkenones, all lying within the group of planktonic organisms known as haptophytes: the species Gephyrocapsa oceanica, the genus Chrysotila, and most importantly the incredibly numerous coccolithophore Emiliana huxleyi.
These alkenones have two useful properties. First of all, they survive conditions that would destroy most organic molecules. Consequently, alkenones have been found in marine sediment as much as 110 million years old. Second, the different varieties of alkenone are produced in different quantities at different temperatures. The relationship between temperature and the proportions of C37 alkenones is given by:
T = 29.41 × Uk'37 - 1.15
where Uk'37 is the proportion of C37:2 and C37:3 which is C37:2, and T is the water temperature in °C; specifically, since all the organisms that produce alkenones are planktonic, T gives us the surface temperature of the water.
The longevity of alkenones in the sediment means that we can use them as a paleoclimatic proxy.
Limitations of the method
There are two main limitations on the method. Firstly, the relevant organisms don't grow in polar waters, and so obviously can't be used to indicate their temperature. Secondly, the formula breaks down for very warm water. For the proportion of C37 which is C37:2 can never be more than 100%, which means that the formula can never yield a T value of greater than 28.3°C no matter what the actual temperature of the water is.
We should also mention a couple of potential confounding factors which have been discovered experimentally (see Prahl, Wolfe & Sparrow, Physiological impacts on alkenone paleothermometry, Paleoceanography, 18(2)). First, U37k' is reduced by conditions of low nutrition; second, it is increased by conditions of prolonged darkness.
Can we be sure that this relationship held good in the past? It seems likely: if there is some reason why C37:2 is preferable in warm waters now, then why should it have been different in the past? But we could say this with more confidence if why knew why C37:2 is preferentially produced in warmer waters. As it is, biologists don't yet know why these haptophytes produce alkenones in the first place. Under these circumstances, perhaps we should be a little more skeptical of this biochemical proxy than of those based on known chemical and physical mechanisms.
Note on terminology
The subscript and superscript in Uk'37 can also be written as U37k', and when superscripts and subscripts are not available, people will write either Uk'37 or U37k'. The reader wishing to research the subject further by performing an internet search should be aware of this variation.
In this article we shall look at the TEX86 temperature proxy, how it works, and how we know that it works.
The TEX86 method is based on glycerol dibiphytanyl glycerol tetraethers (GDGTs). These come in various forms with more or fewer cyclopentane structures (the reader need neither know nor care what these actually are). The GDGTs of interest to us can be denoted as GDGT 1, GDGT 2, GDGT 3 and GDGT 4' (pronounced "four-prime"), where the numbers 1, 2, 3 and 4 correspond to the number of cyclopentanes.
(For the benefit of those readers who wish to research the TEX86 method in the technical literature, I should point out that different papers use different numbering schemes; the one used here seems most suitable, because of the correspondence between the GDGT number and the number of cyclopentanes.)
Crenarchaeota and temperature
In nature these GDGTs are produced by the group of single-celled organisms known as the Crenarchaeota. As with the alkenones discussed in the previous article, the GDGTs resist processes that destroy most organic compounds, and so can be found in marine sediment; and just as with the alkenones, the proportions of the different GDGTs produced by the Crenarchaeota varies with temperature, according to the formula:
T = 56.2 × TEX86 - 10.78
where T is the temperature in °C and TEX86 (an abbreviation of "TetraEther indeX of tetraethers consisting of 86 carbon atoms") is defined as the ratio of the sum of the quantities of GDGTs 2, 3 and 4' to the sum of the quantities of GDGTs 1, 2, 3 and 4'.
It should be noted that this relationship ceases to hold below about 5°C; below this temperature the variation in TEX86 becomes negligible and so measurements of TEX86 can't distinguish between temperatures below that point.
How do we know?
We can measure TEX86 in living organisms and recent sediments, and measure the temperature of the water in which they are found; this is how the formula given above was derived.
However, unlike the Uk'37 method, the relationship is harder to demonstrate experimentally. Experiments do show that TEX86 increases with temperature in the lab; however for some unknown reason lab-grown cultures of Crenarchaeota produce less GDGT 4' than is found in Crenarchaeota in the wild, and so the exact relationship between temperature and TEX86 can't yet be replicated in the laboratory.
In this article we shall look at how core samples from ice can be used to give us information about paleoclimates. The reader may find it useful to look back at the main article on glaciers before reading further.
At any location where snow falls but does not melt (or at least does not completely melt) before the next year's snowfall, the snow will accumulate, and as each year's snow is buried by the further snowfall of succeeding years it compacts from loose snow to permeable firn to impermeable ice, at which point it is said to undergo closure.
Ice formed from summer snow is lighter and less dense than ice formed from winter snow; as a result, if the rate of accumulation is more than about 4 cm/year we get what in effect are varves formed from snow. At least in the upper part of the glacier, these are plainly visible if we take a core sample, as you can see in the photograph below.
Ice core data
The fact that the ice cores contain visible annual layers means that just as with varves in lakes, we can count them down from the top, and assign a year to each one.
Looking at the thickness of each layer, we can quantify the amount of snowfall in the summer and winter of each year. What is perhaps more interesting, we can measure the 16O/18O ratio of the water, which acts as a climatic indicator; we can also measure the 1H/2H ratio of the water, which acts as a climatic indicator for the same reason: water containing the 2H isotope is slightly heavier, and so evaporates less readily.
Besides these data, when the permeable firn turns to impermeable ice, bubbles of the atmosphere become trapped within the ice. This allows us to analyze the past composition of the atmosphere, and quantify gasses which affect the climate, such as carbon dioxide (CO2) methane (CH4) and sulfur dioxide (SO2).
Also, as with ordinary sediment, the ice cores will contain windborne particles including volcanic ash, pollen, and loess (dust produced by the action of glaciers).
How do we know?
We can see the annual formation of layers in ice, and we also understand the mechanism behind this in terms of seasonal variation. For the layers not to form at all, there would have to be a complete cessation of snowfall, which is unlikely; for more than one layer to form per year we would need more than one warm period per year, which is almost impossible.
We can also verify that the data preserved in the layers reflects the climate by comparing the ice core record with direct measurements of climate made over the last few centuries, and with other climatic indicators.
To this we can add that the methods outlined in this article should work; they are based, after all, on very simple physical principles. Is it even conceivable that in times gone by water composed of heavy isotopes evaporated more easily than water composed of light isotopes? Or that once upon a time when firn underwent closure it trapped bubbles of something other than the atmosphere?
So the analysis of ice core data ought to work in principle, and comparison of the results with other data confirms that it does work in practice.
Difficulties with the method
Because firn doesn't turn to ice immediately, and because the air bubbles aren't trapped until it does, the year we calculate for the age of a layer will be different from the age of the atmosphere trapped in that layer. The difference (closure time) between the age of the gas and the age of the ice can be as much as 7,000 years, as is the case in ice cores from Vostok, or as little as 30 years at Law Dome.
Mathematical methods can be used to recover data on a finer scale than the closure time; however, such methods rely on knowing the closure time. Now, the climate has changed in the past (otherwise paleoclimatology would be a completely unnecessary science) and so it seems likely that closure time will have varied from time to time in the past just as it varies from place to place in the present. This introduces an element of uncertainty into the data.
Another problem is that at greater depths the annual layers become more and more indistinct, until in some cases they can't be made out at all. One can try in such cases to estimate the age of the buried ice by measuring its depth and the accumulation at the location in question, but again this would only work perfectly if the climate, or at least the snowfall, had remained constant each year. Where the ice cores contain volcanic ash, this can be dated by radiometric methods, allowing us to put correct dates on the layers in which they are found, but only up to the accuracy of the radiometric methods used.
There is one final weakness inherent in the method. By definition, any point where the annual rate of snowfall exceeds the rate it melts must be the accumulation zone of a glacier; in the case of the Greenland and Antarctic ice sheets that most interest paleoclimatologists, continental glaciers. And a glacier must flow out from its zone or zones of accumulation to a zone or zones where it ablates. In short, the record in the ice is progressively being destroyed as well as created. This happens relatively fast by geological standards, providing us with a record that can be measured in hundreds of thousands of years; compare this with proxies based on marine sediment, which is only destroyed by the much slower process of subduction.
However, ice cores remain valuable because the ice does trap air in its pore spaces, providing us with a continuous record of the composition of the atmosphere.
In this article we shall discuss what Milankovitch cycles are, their effect on climate, and how we know that this effect exists.
The reader should remember from high school that the reason why the Earth has seasons is that its axis of rotation is at an angle to the plane of its orbit. When the northern hemisphere is tilted towards the sun, then it is summer in the northern hemisphere and winter in the southern hemisphere; when the southern hemisphere is tilted towards the sun, it's the other way round.
The magnitude of this effect depends on how tilted the Earth's axis is, and this angle varies between 22.1° and 24.5° in a 41,000 year cycle.
The reader should also recall that the Earth's orbit is not perfectly circular: it is an ellipse with the Sun at one focus, meaning that the Earth is closer to the Sun during some months of the year than others. The effect of this is less than you might suppose: the Earth is five million kilometers closer to the Sun in January than in July, but this doesn't stop the Northern Hemisphere from undergoing winter.
The magnitude of this effect depends on how far the Earth's orbit deviates from being circular, and a number of factors affecting this figure add up to a cycle of about 100,000 years.
Finally, there is the precession of the Earth's axis. At present, as we have seen, the North Pole is tilted away from the Sun at the Earth's point of closest approach to the Sun. However, this too varies, in this case in a 21,000 year cycle.
These, then, are the three Milankovitch cycles; as they are of different lengths, their interaction will produce rather a complex pattern as they go in and out of phase with one another. Together, they will affect both the total annual insolation (the amount of solar radiation that reaches the earth's surface) and also season variations in insolation.
Although the idea of Milankovitch cycles as a factor in the Earth's climate was initially greeted with some suspicion by climate scientists, it is now generally accepted that Milankovitch cycles account for about 60% of past variation in climate.
How do we know?
First of all, how do we know that Milankovitch cycles exist? A short answer is that the physics of the Solar System require that they should exist; a longer answer would require an introduction to celestial dynamics which would be excessive in length and out of place in what is supposed to be an introduction to historical geology.
From a geological perspective, we can look for the effects of the cycles on the sediments and proxies used in paleoclimatology. As Milankovitch cycles can't be the only thing affecting the climate, this is not so simple as demonstrating that the climate fluctuates perfectly in synchrony with the cycles; rather, statistical analysis is necessary to sort out the "signal" of the cycles from the "noise" produced by (for example) variations in volcanic activity. Such an approach confirms that the cycles have a real effect on the climate.
One interesting effect of the cycles is that in places where the nature of sedimentation is sensitive to the climate, we can see rhythmites with a period dictated by the lengths off Milankovitch cycles, as shown in the photograph below.
There remains one outstanding puzzle. In principle, the 100,000-year cycle should have less of an effect than the 41,000-year cycle. But for the last million years or so, the 100,000-year cycle has predominated; whereas prior to that the 41,000-year cycle was indeed more important, in line with theory. This is known as the 100,000 year problem, and serves as a useful reminder that our understanding of long-term climatic change is still imperfect.
In this article we shall look at how climate models are constructed, their strengths and weaknesses, and how they can be applied to the study of plate tectonics.
Climate and climate models
First we should look at what factors should be represented in a model of the climate.
First of all, there is the insolation. This drives the climate, but it doesn't exclusively determine it, otherwise the weather at any point could be expressed as a function of time and latitude.
But in fact there is also the atmospheric circulation to be taken into account. This transports atmospheric heat and moisture from place to place. The atmospheric circulation is partly caused by the Coriolis Effect (fundamentally, because the Earth is rotating). However, the circulation is also caused by variations in the density of the atmosphere, which are caused by variations in the temperature and amount of moisture in the atmosphere; or, to put it another way, one of the main causes of the weather is the weather. It is this fact that makes modeling the weather or the climate particularly difficult; such systems are notoriously hard to model.
One important factor affecting the nature of the circulation is the location and nature of the landmasses. The ocean will absorb more heat than the land; and then the amount of heat absorbed the land will depend on the nature of the ground cover, i.e. whether the land in question is desert, forest, covered by a sheet glacier, etc. Similar remarks can be made about moisture; obviously more water will evaporate from a sea or a lake than from a desert with the same insolation.
By modeling insolation and the circulation of the atmosphere, climatologists can produce what is known as an atmospheric circulation model, or ACM.
Such a model is the simplest that's any use at all, but it still doesn't tell the whole story. The ocean has its own circulation, and this also transports heat in tandem with the atmospheric circulation. A really good model should take this into account.
Surface ocean currents can be driven by the wind. Deeper in the oceans, we have the thermohaline circulation. As the name suggests, or would suggest if we were Greek, this circulation is driven by the temperature and salinity of the oceans, both of which affect the density of seawater. The differences in density drive vast currents, extending thousands of miles and carrying 100 times more water than the Amazon River.
The map above shows the thermohaline circulation; as you can see, it can be quite complex, with less dense currents actually flowing over denser currents flowing beneath them in a different direction.
The other thing that affects the ocean circulation is, of course, the positions of the continents, which constrain the flow.
By adding the ocean circulation to an ACM, climatologists can produce a General Circulation Model, or GCM.
Even then, there are things which we would like to add to a really good model of the climate. For example, we have noted above that ground cover affects the climate. But it is also the case that the climate will to a large extent determine the ground cover from location to location, causing forests here, and grasslands there; deserts in one place and sheet glaciers in another. Once more we have a case of the climate causing the things that cause the climate, and so we have another complicated set of interactions to model --- if we can.
Accuracy of the models
The size and complexity of the climate are a challenge even to modern supercomputers, and limited processing power (in effect, how fast the computer can perform calculations) limits the accuracy and detail attainable in climate models. The most obvious problem is that of scale. No computer can be expected to simulate the behavior of each molecule of air in the atmosphere. Suppose, then, that as a coarser approximation we divided the Earth up into "cells" each one kilometer square, and used a model that assigned to each one at each step of the computer simulation a figure representing the average temperature and the average humidity of the atmosphere inside that "cell". Then our model would be inaccurate by reason of this approximation; it would also be inaccurate by ignoring the oceanic circulation, and it would still involve simulating the interaction of half-a-billion cells; and so even such a degree of approximation might well leave our model intolerably slow. Of course, scientists can always spend more time running slower models, but they rightly suppose that an imperfect model which returns results within the lifetimes of the researchers is, despite its acknowledged flaws, superior to one which delivers more accurate answers to their great-grandchildren.
Any model of the climate must therefore involve some degree of simplification and approximation; and any such model will therefore be wrong to a certain extent. It is for this reason that there are so many climate models: the researchers must choose in what way to simplify their models in order to reduce the degree of error, and it is not yet clear how best this may be achieved. As a consequence of this, the first report of the Paleoclimate Modeling Intercomparison Project (hereafter referred to as PMIP1) lists 22 climate models produced by 19 different research institutions, each in their own way striving after accuracy.
It is possible to find out how good or bad the various models are, and in which respects, by comparing their results to data. We can of course compare the climate now (as measured instrumentally) with the climate as the models say it should be; any model not producing reasonably good agreement has fallen at the first hurdle. But we can also get the models to simulate events in the past and compare them with proxy data. This was done by the PMIP, comparing models with proxies for the mid-Holocene (6,000 years ago) and the last glacial maximum (LGM, 21,000 years ago). At these dates, the climate was markedly different from the present; yet they are recent enough that the full array of proxies can be brought to bear on what the climate was actually like.
The results are such as to be encouraging to an optimist and disappointing to a pessimist. To briefly summarize the results, the models tend to be qualitatively correct: they correctly indicate the nature of the differences between the present climate and the climates of the mid-Holocene and the LGM; on the other hand, they are quantitatively inaccurate, tending to underestimate the magnitude of the differences between then and now.
It should be added that in the first phase of PMIP the research was confined to ACMs which did not take the oceanic circulation into account; most models also omitted the interaction between climate and ground cover. One would expect the results to be more accurate when more factors are taken into account; more complex models which do so will be reported on the in forthcoming second phase of PMIP.
Paleoclimatology in deep time
If we want to look back millions rather than thousands of years, then there is another factor we need to take into account: the fact that the continents have altered their positions. As their location affects both the atmospheric and oceanic circulation, modeling paleoclimates in deep time requires a reconstruction of their positions.
This leads to an interesting line of research. When we reconstruct the position of the continents at some past date, and use climate models to tell us what, in theory, the climate should have been like, does this agree with the sedimentary evidence of what the climate was like at that time?
The answer is "yes". There are a number of results obtained from climate models which are consistent with proxies and sedimentary indicators of climatic conditions. For example:
* Continents at higher latitudes are, other things equal, colder than those nearer the equator. (This is rather obvious, but it is nice to see it confirmed by, for example, evidence of continental glaciers rolling over Africa at a time when paleomagnetic data tell us it was further south.)
* Changes in the average temperature of the Earth have more effect near the poles than the equator.
Now this sort of agreement should increase our confidence in our reconstruction of continental drift, in climate models, and in the use of sediments to reconstruct climates; for if any one of these three techniques was no good, then there would be no reason for the agreement between the models and the evidence.
In this article we shall look at what an ice age is, what causes ice ages, and how we can identify the evidence for ice ages in the geological record.
The reader will probably find it useful to look back at the main article on glaciers before reading further.
Definition of an ice age
An ice age, or to use the more technical term, a glaciation, may be defined as a time when continental glaciers are present. This, of course, means that we are currently experiencing an ice age, since there are continental glaciers on Antarctica and Greenland. To be more precise, we are living in an interglacial: a warmer interval within an ice age when the glaciers have retreated towards the poles.
Glaciations in the geological record
Geologists have identified and dated a number of ice ages in the geological record:
* The Huronian glaciation (or Makganyene glaciation) extended from 2400 to 2100 million years ago
* The Cryogenian glaciation lasted from 850 to 635 million years ago.
* The Andean-Saharan glaciation was from 460 to 430 million years ago.
* The Karoo Ice Age lasted from 360–260 million years ago.
* The Pliocene-Quaternary ice age is the one now in progress, and started about 2.6 million years ago.
The methods of dating events have already been explained in other articles; later in this article we shall discuss how we identify ice ages.
Causes of ice ages
We have quite a good idea of the causes of the present ice age, which can be explained by continental drift affecting the oceanic circulation. The causes of previous ice ages are more obscure and debatable.
This is because causality as such is not preserved in the geological record. We may be able to see that event A happened and was followed by event B, but the fact (if it is a fact) that event A caused event B must be established on theoretical grounds, by understanding the relationship between the events of the type in question. For example, it is not hard to deduce a causal relationship between the feet of dinosaurs and dinosaur footprints, because we have a good general understanding of the relationship between feet and footprints. But our understanding of climatic effects is hardly as secure.
A further point to consider is this: ice ages are rare and irregular in the geological record. This suggests that there may well not be a single cause for ice ages. If there was a constant reason why ice ages happen, then they would be a permanent feature of the geological record, or at least occur as cyclic events; the fact that they are sporadic suggests that they had singular causes. The best we can do in investigating the causes of an ice age is to look at the events leading up to it and see which of these events might, in principle, be the cause of that particular ice age.
Consider, for example, the Huronian ice age. Two explanations have been proposed for this.
(1) There is evidence (as will be discussed in the article on banded iron formations) that oxygen levels rose dramatically in the time leading up to the Huronian. Abundant free oxygen would have combined with the methane (CH4) in the atmosphere, converting it to carbon dioxide and water. Methane is a potent greenhouse gas; that is, it helps to keep the climate warm. Carbon dioxide is also a greenhouse gas, but much less potent than methane. The result of the rise of oxygen and the decline of methane should, therefore, have been a decline in global temperature.
(2) There was remarkably little volcanic activity between 2.45 and 2.2 billion years ago. This can be shown by analysis of zircons, which being resistant to weathering and erosion outlast the igneous rocks in which they are formed, and which can be accurately dated as explained in the article on uranium-based dating methods. The scarcity of zircons formed between 2.45 and 2.2 billion years ago as compared to their relative abundance both before and after strongly suggests a time of low volcanic activity.
Now volcanoes are prone to emit the greenhouse gas carbon dioxide; if volcanic activity virtually ceased, while the process of chemical weathering continued to remove carbon dioxide from the atmosphere, this would lead to a reduction of the greenhouse effect and a decline in global temperatures.
There is, of course, absolutely no reason why both these mechanisms shouldn't have worked together to produce the Huronian glaciation; but while this is possible, it is also possible that one or the other was of such a greater degree of significance that it would be reasonable to call it the cause of the Huronian glaciation.
Whichever of these mechanisms was the main cause, it would be an example of a one-off event. The transition from an atmosphere with little or no free oxygen to one with abundant free oxygen happened, and indeed could happen, only once in the history of the Earth. But a quarter-billion years of low volcanic activity is equally rare in the history of the Earth. It follows that studying the possible causes of this one ice age tells us nothing about the causes of the others, nor can studying the causes of the ice age that we're currently in tell us anything about the causes of the Huronian ice age.
While it can be difficult to say why a particular ice age occurred, it is relatively easy to determine that it occurred. We shall now turn our attention to the nature of the evidence.
Ice ages: how do we know?
As explained in the main article on glaciers, we can observe sedimentary and erosional features associated with glaciers existing today, or which have melted within the time that people have been keeping records of their location: features including striations, moraines of glacial till, deposits of outwash, dropstones in lakes, etc.
When we look at the evidence for recent glacial advances, we see the same sedimentary and erosional signs, but further south than the southern limit of the ice sheets in the present interglacial. By looking at the location of the lobe-shaped terminal moraines, we can find the southern extent of the ice sheets.
These observations would really be sufficient on their own; but in addition to this we can use proxies to investigate the advance and retreat of the glaciers. Biogeographical proxies such as pollen are particularly useful; as the events in question are so geologically recent, we can recognize the pollen of modern species and know exactly under which conditions they flourish.
Another source of information is cosmogenic surface dating. Glaciers are a powerful erosive force, and, as the striations show, scrape sediment off right down to the bedrock. When the glacier retreats in an interglacial, the bedrock is exposed to cosmic rays, and so we can use cosmogenic surface dating to date the retreat of the glaciers.
One interesting relic of the recent glacial retreat is the existence of so-called "sky islands". These are mountains found in the southern United States and Mexico which are home to a distinctive flora and fauna, and which are surrounded by dry grassland or scrubland which the sky island fauna can't cross: hence the name sky islands. But if the species can't migrate from mountain to mountain, how did these species achieve their current range? In the light of paleoclimatology, the solution becomes obvious. When the climate of the southern U.S. was cooler than it is today, the climate of the plains would have been congenial to the species now found in sky islands. But when temperatures increased in the interglacial, they had nowhere to go but the cold moist refuge of the mountains.
Another indication of the recent glacial advance and retreat is isostatic rebound. Recall from our discussion of the structure of the Earth that the lithosphere is essentially floating on top of the denser athenosphere. When a region of the Earth is burdened by the weight of continental glaciers, this ought to push the lithosphere down into the athenosphere. When the glaciers retreat, the weight is lifted, and the lithosphere should very slowly bob back up. Geologists can measure just this happening, at a rate of about 1 cm/year, in regions which (according to sedimentary and other indications) were covered by sheet glaciers before the present interglacial.
When we look further into the past, we should expect many of these indications to be absent: isostatic rebound, for example, wouldn't take a quarter of a billion years, and so would not remain as evidence of the Karoo ice age. Similarly, as explained in the article on biogeographical proxies, these become less and less useful the further back into the past we wish to look.
However, we still have the moraines, the striations, the dropstones, etc. Are these sufficient to diagnose an ice age? It is reasonable to say yes: for we find these associated with present glaciers, and we also find them associated with the other indications listed above (sky islands, isostatic rebound, proxies, cosmogenic surface dating) when the deposition of the till is recent enough for these additional indications of the former presence of glaciers to have survived to the present. It would be really astonishing if in the cases where we lack this supplementary information, the sedimentary and erosional indications of glaciation were produced by some other process; it is reasonable to conclude that they are signs of an ice age, both because of their empirical association with the present ice age, and because (with the arguable exception of drumlins) we have a clear understanding of how and why glaciers produce these effects.
Changes in global sea level are important both as a symptom and as a cause of climate change. In this article we shall look at what causes changes in global sea level, and how they can be detected in the geological record.
Causes of sea-level changes
One cause of sea level variations is the formation or melting of continental glaciers. It is estimated, for example, that if the ice-sheets currently covering Greenland and Antarctica were to melt, the global sea-level would rise by approximately 70 meters.
Another factor that affects global sea level is the mid-ocean ridges. The more active they are, the larger they are, and the more water they displace. The ridges can also change in total length: when a supercontinent rifts apart, a mid-ocean ridge will be produced where it rifts; the production of the mid-Atlantic ridge by the rifting of Pangaea is the most recent example. As the ridge is in effect a mountain range 40,000 miles long, its formation clearly displaced a massive amount of water. Hotspots producing islands such as Hawaii has a similar though lesser effect.
The collision of continents can reduce sea-level by reducing the area of the continents as they compress into one another. For example, it has been estimated that the collision of India with south Asia and the compression resulting in the raising of the Himalayas would have lowered the sea-level by about 10 meters.
The position of the shoreline can change as a result of the movement of the land, as a result of tectonic events or of isostatic rebound after depression caused by glacial cover. This means that there's no use in just looking at one or two places to estimate sea levels in the past, because what looks like a global fall in sea levels might in fact be a local rise in the level of the land.
Such local effects did in fact confuse the 18th century scientist Anders Celsius, whose measurements of the sea-level along the Scandinavian coast convinced him that the seas must be shrinking. He was not to know that in fact Scandinavia was slowly rising out of the sea as a result of isostatic rebound.
To identify global changes in sea-level, it is necessary to look beyond these localized changes and see the bigger picture.
The Vail and Hallam curves
The Vail curve representing past global sea-levels is based on the work of Peter Vail and his associates. It is based on the study of unconformities in the geological record representing times of marine regression. An important part of Vail's work was establishing that these unconformities could be identified underground by the use of reflection seismology, a method which relies on studying the reflections of seismic waves artificially generated by producing small explosions.
The Vail curve has one major problem, from a methodological point of view: it is based on proprietary data of the Exxon-Mobil corporation which the general public is not allowed to see. This doesn't necessarily mean that the Vail curve is wrong, but it does mean that if some aspect of it is wrong, if the data on which it's based are incorrect in some way, or if Vail has somehow misinterpreted the data, then no-one would be able to tell. This makes it different from other results in geology, the quality of which is insured by the ability of scientists to cross-check each other's work.
The same problem does not arise with the Hallam curve, the brainchild of Anthony Hallam. This was constructed on the basis of a different methodology: Hallam calculated the area of the continents covered by the sea over the course of time by seeing where marine sediments were and were not deposited at various times, allowing him to sketch out a series of shorelines for the various continents over time; clearly when the land area shrinks globally, this corresponds to a global rise in sea level.
The Hallam and Vail curves are presented in the chart below. Note that these are "first order" curves which only show broad trends: they are not on a fine enough scale to show the fluctuations caused by recent advances and retreats of sheet glaciers.
The reader will note the broad similarities between the two curves, which suggest that Hallam and Vail are basically on the right track, especially since this agreement is produced by two different methods. But the reader will also note the differences in detail, which suggest that one curve or the other, or probably both, are not completely accurate.
Their broad accuracy is confirmed by the fact that we see in them the features that we would expect to see; for example, we see a fall in sea-level accompanying the assembly of Pangaea, and a rise in sea-level corresponding to the fracturing of Pangaea and the formation of the Mid-Atlantic Ridge. We can also see falls in sea-level corresponding to known times of glaciation.
This should really have gone back in the section on sediment, but as it also relates to climate I guess it can go here.
Banded iron formations
In this article we shall look at what banded iron formations are and how they might have formed. The reader may find it useful to look back at the main article on siliceous ooze and chert.
Banded iron formations, or BIFs are sedimentary rocks consisting of alternating bands iron-rich sediment (typically hematite (Fe2O3, and magnetite, Fe3O4) and iron-poor sediment, typically chert; the size of the bands ranges from less than a millimeter to more than a meter in thickness. The image below shows a fairly typical banded iron formation: the red bands are the iron oxides.
While BIFs have a wide geographical distribution, they are localized in time. They start to become common about 3.5 billion years ago), peak around 2.5 billion years ago, vanish about 1.8 billion years ago, make a small comeback around 1 billion years ago, and then essentially vanish from the geological record; none are being produced today.
BIFs and the rise of oxygen
With the exception of saline giants, it is usually very easy to explain the origin of sedimentary rocks, because we can see identical sediments being deposited in the present: chalk looks just like lithified calcareous ooze, tillite looks just like lithified glacial till, aeolian sandstone looks just like lithified desert sand, and so on.
In the case of BIFs, however, no BIFs are being formed in the present, nor even recently. It seems, then, as though in searching for a cause for BIFs we must be looking for an event which could only have happened at in the past.
Fortunately, one comes to mind. According to biologists, the first living organisms neither produced nor consumed oxygen. Indeed, they would not have been able even to tolerate oxygen: oxygen is a very reactive gas, and is toxic to organisms which are not adapted to its presence (for example the modern bacterium Clostridium botulinum, which can only survive in the near-total absence of oxygen).
Biologists are also agreed that in the absence of oxygen-producing organisms, the atmosphere would have been very poor in free oxygen (i.e. the molecule O2).
What does all this have to do with BIFs? Well, one of the interesting things about iron is that elemental iron (Fe) dissolves in water, whereas the various oxides of iron (as found in banded iron formations) precipitate out. The waters of the early Earth would certainly have had sources of iron, such as emissions from submarine volcanoes, and iron liberated from rocks by chemical weathering. It follows that when organisms arose that produced oxygen, iron dissolved in the oceans would combined with dissolved oxygen to form iron oxides which would then have precipitated out, producing the iron oxides that characterize BIFs.
The iron would, indeed, form an "oxygen sink"; only after the iron had been used up in this way would O2 have begun to constitute a large proportion of the atmosphere. The accumulation of oxygen in the atmosphere, which according to geological dating methods started about 2.4 billion years ago, is variously known as the Great Oxygenation Event (GOE), the oxygen catastrophe, and the oxygen crisis.
How do we know?
The scenario given above is plausible according to biologists; indeed, if they're right about the history of early life, we should expect to see this kind of geological evidence of a rise in oxygen as a result of the rise of oxygen-producing organisms. And biology aside, it is certainly chemically very plausible: iron is soluble in water and iron oxides are not, and this at least is something we can check by direct observation. It is also plausible in that it explains the localization in time of BIFs in terms of something that we would expect to happen only once.
But was there in fact a change from an oxygen-poor to an oxygen-rich atmosphere? Studies of minerals before, during, and after the GOE answer this question in the affirmative.
The fact that the production of free oxygen is indeed a plausible explanation for BIFs is a point in favor of this scenario; what is more, this explanation is borne out by the nature of the iron oxides in BIFs: they tend to be iron oxides with a low oxygen to iron ratio such as hematite (Fe2O3, and magnetite, Fe3O4) rather than, for example, goethite (FeO(OH)); which is what we would expect if they formed in conditions in which oxygen was still scarce.
However, we would be verging on circular reasoning if we explained BIFs by the advent of free oxygen, and if our evidence for this event consisted solely of BIFs. Fortunately, this is far from being the case: there are other indications of the GOE in the mineralogy of the early Earth.
For example, before, but not after, the date assigned to the GOE, the minerals uraninite (UO2) and siderite (FeCO3) can be found in river sediments (for information on how to identify such sediments, see the main article on rivers). The significance of this is that these minerals would not survive in waters containing dissolved O2, as all rivers do today; so the rivers that deposited them must have co-existed with an oxygen-poor atmosphere.
After the date assigned to the GOE, on the other hand, we see a great diversification of mineral types in the geological record, as after the GOE new minerals could then be produced from old ones by oxidization; and it is just such minerals that we find after the GOE. (For more details, see Sverjensky and Lee (2010) The Great Oxidation Event and Mineral Diversification, Elements, 6(1).)
So we have abundant data pointing to the rise of free oxygen in the atmosphere; and such a rise would explain, indeed necessitate, the extensive formation of iron oxide deposits such as are found in BIFs.
BIFs: some questions
It seems, then, as though we have a good explanation for BIFs. However, the reader should bear in mind that what I have sketched here is only a broad outline of a broad consensus. There is still controversy over details.
The reader may well already be puzzling over a couple of these details. Firstly, why is the sediment in the dark bands of BIFs so frequently chert? And why are BIFs banded at all --- why were the chert and iron oxides not deposited simultaneously as a mixture?
The chert may have been deposited by silica-forming organisms. It is true that we don't find in it the tests of diatoms and radiolarians, so it probably didn't have its origin as the sort of siliceous ooze produced by these organisms today; but one cannot rule out the possibility of other silica-producing organisms active in the pre-Cambrian and now extinct.
Alternatively, if there were no silica-producing organisms in the pre-Cambrian, then silica could just have built up in the seas until it reached saturation point and precipitated out by itself.
The rhythmic nature of deposition seems to suggest a cyclic variation in conditions. One candidate cause is a repeated cycle of ecological boom and bust. On biological grounds, it would be reasonable to suspect that the earliest oxygen-producing organisms could not tolerate high levels of oxygen. (This is not as paradoxical as it sounds: most organisms can't live in an environment full of their own waste products. Animals, for example, produce CO2 and would suffocate in an atmosphere dominated by it.)
This suggests the following scenario: oxygen producing organisms would grow and flourish until they had produced toxic levels of oxygen; the population would then collapse almost to nothing, surviving only in low-oxygen refuges; the oxygen would be removed from the atmosphere by combining with the iron in the water to produce the iron oxide bands of BIFs; in these low-oxygen conditions the oxygen-producing organisms could once more increase in number, and the cycle would begin again. This would plausibly account for the periodic precipitation of iron oxides. It is even conceivable that the oxygen-producing organisms were identical with the (hypothetical) silica-secreting organisms mentioned above.
But here we have gone into the realms of speculation and controversy. It is quite possible that these questions will remain controversial: since geologists labor under the handicap of not being able to watch BIFs form in the present, BIFs will never be understood with quite the same certainty as other sedimentary rocks.
I need to go back and put some stuff in. For example, I need to write articles about faults and folds. While I was writing this, I thought during the section on plate tectonics that I should wait 'til the section on stratigraphy, and when I was doing stratigraphy I thought it should be in the section on plate tectonics, so I never really got around to it. So I'll go back and do that.
But anything from now on is out of order, and should have been posted earlier.
So, I'm pretty much finished. Can anyone point out to me something that I should have explained and haven't? Thanks.
Ok, there now follows stuff that I should really have done early.
With faults and folds, when I was doing tectonics I thought I should do them in stratigraphy, and when I got to stratigraphy, I thought they belonged in plate tectonics. Having thought about it, they definitely fall under stratigraphy. So here's something about folds.
In this article we shall look at folds and how to recognize them in the geological record. The reader will find it useful to have read the articles on orogeny, Steno's principles, and way-up structures before continuing.
Origin and appearance of folds
The appearance of folds shows that they are produced by the landscape being pushed sideways, as described in the article on orogeny, rather than being pushed up or pulled down from below.
This explains the normal appearance of folds (prior to erosion) as a set of parallel ridges where the landscape has been folded up (anticlines) alternating with troughs where it has been folded down (synclines). The photograph below shows an anticline on the left and a syncline on the right, seen from a side-on perspective.
Sufficient lateral motion will push the folds themselves over sideways, resulting in a recumbent fold, as shown in the article on orogeny.
Folds: how do we identify them?
When, as in the photograph above, a fold is intact and we can see it from the side, it is obviously a fold.
But now consider a sequence of events such as that shown in the block diagrams below: sediments are laid down in flat layers according to the the principle of original horizontality and the principle of original lateral continuity, lithified, folded, and then eroded.
As a result of the erosion, the shape of the fold will be destroyed. Even then, if we could see the rocks from the side it would not be hard to deduce that this was an eroded fold. But geologists are not usually so lucky, and instead have to view the landscape from the top.
Nonetheless, it would not be hard to deduce what happened. Looking at the rocks, a geologist could see that the formation labeled A is the same kind of rock as a, B is the same as b, and so on. Applying the principles of original horizontality and lateral continuity, the only possible explanation would be that the rocks were folded and then eroded. The symmetric pattern of bands of different rock types across the landscape are by themselves an excellent indication of an eroded fold.
Even if the types of sediment didn't vary during deposition, there would be plenty of other evidence. Looking at the individual beds in the rocks, we would find them sloping up to the right on one side of the anticline, and up to the left on the other. Again, the principles of original horizontality and continuity would tell us that at one point the layers would have been flat and met up in the middle.
Similarly we could look at the way-up structures in the rocks: on one side of the anticline they would indicate that the original up direction had been tilted to the left, and on the other side to the right.
One interesting sign of folding is that clasts and fossils within the rocks can be stretched and deformed as the rock gradually folds. Amusingly, before this process was well-understood, early paleontologists would identify the same species as two different species according to how the specimens were stretched: a fossil fish stretched one way would appear long and thin, whereas another originally identical fish stretched the other way would look short and wide.
Indications such as these allow us to identify folds even if they have been eroded, and even if we can't see them from the side.