Bones in the Stones, Shells in the Shale: Fossils and Fossilization
The contemporary definition of fossil is the physical remains or trace of behavior preserved in the rock record. Specimens which were actually once part of a living thing (bones, teeth, skin, shell, leaves, trunk, pollen, egg, etc.) are body fossils; those that are marks in the sediment produced by the activities of living things are trace fossils.
First things first, though. To understand fossils, we need to understand rocks and how they form.
Structure of the Earth
Different researchers might recognize various subdivisions of the spheres. Here is a list, running from the most interior (deepest and densest) outwards:
The deep interior of the Earth interacts with the parts in which climate happens and we live, but generally only very slowly. Our knowledge of the interior comes almost exclusively from various forms of remote sensing: despite movies the contrary, we do not have the means to drill deep into the mantle or to the core.
The innermost part of the earth is the core, comprised largely of the metals iron and nickel. The inner core is solid, despite having temperatures over 5700 K: with pressures of 330-360 gigapascals the metals are compressed into a solid crystalline structure. The radius of the inner core is 1220 km. The inner core is surrounded by the 2260 km thick outer core, which is liquid. Motions of this vast inner sea of molten metal generates the magnetic field of our planet. The core is hot because of heat left over from the initial formation of the planet, but also (far more importantly) from radioactive decay of various isotopes and the heat of crystallization of the growing inner core.
Surrounding the core is the 2890 km thick rocky mantle. The mantle represents 85% of the Earth's volume. It is basically solid, but because it is hot and under pressure it can flow like tremendously dense silly putty. The mantle rock is very dense: much denser than the typical rocks found on the continent or the ocean floor.
Heat from the core-mantle boundary is dissipated by the formation of vast convection cells in the
This motion (moving at rates comparable to finger nail growth: a few cm per year) drives the action of shallower geology. The mantle plays a role in the long term carbon cycle, but is otherwise mostly isolated from climate actions.
Technically speaking, the lithosphere is a dynamic subdivision of the Earth, whereas the core and
mantle are compositional subdivisions. The mantle is covered by the brittle rocky part of the Earth: the crust
(which ranges from about 5 km to 50 km deep). Functionally, however, the outermost mantle shell and the crust move
as a single unit, collectively the lithosphere. The lithosphere is divided into various plates, which
move relative to each other as a result of the mantle convection cells below. Interaction between plates results
in nearly all of geological phenomena:
Such action results in the widening of oceans; the motion of continents; the loss (subduction of older oceanic crust back into the mantle; the driving of volcanoes and earthquakes; the uplift of new mountains; and more.
The lithosphere rides along a mobile asthenosphere, a portion of the mantle where temperature-pressure conditions support the presence of many molten droplets within the rock.
Compositionally, the crust is phenomenally diverse. All sorts of rocks are formed and deposited here. The lithosphere is also a region of various types of activity, continuous or episodic, small-scale to catastrophic. Some of the major ones to consider are:
Lithospheric processes thus both add and subtract material from surficial Earth systems, and these might be as slow as the erosion of a mountain range or as rapid as the eruption of a volcano.
Some data difficult to explain if the continents did not move:
It was discovered in late 19th Century that the sea floor is flat, everywhere dense volcanic rock: very different from continents with mixed rock types and much lower average density. Thus, the ocean floor does NOT represent simply submerged versions of today's continents. The submergence/emergence model of past geography was clearly rejected. Additionally, it was discovered that when mapped out, earthquakes and volcanoes tend to follow particular tracks along the margins of some continents, in the middle of oceans, and other additional patterns that called for some explanatory theory.
Continental Drift: Theory proposed by Alfred Wegener, German geophysicist and glaciologist, in 1915. His model: the light continental masses move over dense layer of oceanic crust (by analogy to motion of light glacial ice moving over bedrock below.) Volcanoes, mountain building, earthquakes caused by crumpling of continental masses as they move along. In the distant past the continents were united together, but subsequently some force broke them apart and is moving them ever since.
Resistance to continental drift was strong in the US, Canada, and the UK (although more widely accepted by Southern Hemisphere geologists.) In part, northern resistance because Wegener failed to propose causal mechanism that could be well-verified (not that their own stabilist model had a verified causal mechanism, either!) But there were ad hominem components to the rejection, too: in part, post-war Germanophobia, and in part, cross-disciplinary "snobbery". At a 1926 Meeting of American Association of Petroleum Geologists, the majority of the talks were strongly against Continental Drift. From this point on, continental drift became a fringe subject among northern hemisphere geologists
Sea-Floor Spreading: In the 1940s and 1950s some geologists (notably Arthur Holmes and Harry Hess) had proposed a mechanism to move continents: the coninents did not move OVER the oceanic crust, but carried along with it as the sea-floor itself was recycled. In post-WWII era, additional discoveries concerning depth of earthquakes, age of oceanic crust confirmed sea-floor spreading.
Plate Tectonics: models of continental drift and sea-floor spreading were combined by John Tuzo Wilson and colleagues to form plate tectonics.
Major plate boundaries, from online USGS pamphlet "This Dynamic Earth".
Plate velocities predicted by theory confirmed by GPS studies in 1990s
Heat from Earth's core moves plates, forming mountain ranges at subduction and collision boundaries. Weather erodes uplifting mountains, wind and water and ice transports sediment to depositional environments. Over time, material becomes buried.
Plates wander over Earth's surface, so continents move from tropics to poles or back. Also, action of mid-ocean ridges causes sea levels to rise up (flooding continents) or lower (draining continents). (Current situation is very low sea level).
Big change from the 1960s-1970s model: now recognize there are LOTS of little plates (terranes) rather than just a few big plates. See here for a detailed look at the changing shape of Earth's surface for the last 600 million years; and here for a close-up on North American paleogeography.
Here is a brief animation of estimates of the position of the continental masses over the last 600 million years or so (thus, the
time scale of the course):
Rocks (naturally occurring cohesive solids comprised of one or more minerals or mineraloids) are generated in one of three primary manners (which form the basis of rock classification). Or, to put it another way, every rock is a record of the environment in which it formed:
It is important to consider, however, that not all environments are environments of deposition. Many locations will be environments of erosion: these places are sources of sediment, but because material is being lost from there rather than accumulating there, they will not wind up in the geologic record. A particular location can shift between deposition and erosion ("D-world" vs. "E-world") as local environmental conditions change.
Here are some aspects of depositional environment to consider:
By observing modern environments and their sediments and sedimentary structures, we can use the clues mentioned above (as well as other aspects) to reconstruct the paleoenvironment. Major environments of deposition represented in the geologic record include:
NOTE: Present day sea level is much lower than most of Earth history; also, as new mountains are born, once shallow deposits are uplifted. Consequently, even in the middle of continents, it is common to find sedimentary rocks deposited underwater. In fact, rocks deposited in marine environments are extremely common, even in the interiormost parts of continents.
Whatever the environment of deposition, the sediment is laid down in layers (strata). Since every rock is a record of the environment in which it formed, the strata will be of the same general sort while the environment remains the same, and change to a different sort as the environment changes. Packages of similar strata formed over a region are called formations: at any given spot, if we see a section through the bedrock, we can see the transition from one formation to another, representing a transition from one environment to another.
These factors -- alone or in combination -- can bind the sediment together, transforming it into sedimentary rock.
Lithification is an example of diagenesis: post-depositional alteration of sediment. We will see diagenetic effects in another context shortly, in fossilization.
Plate tectonics ultimately drives the Rock Cycle:
Some common type of trace fossils (also called ichnofossils) include:
Trace fossils of this sort are preserved the same way that other sedimentary structures were: the features are covered over by another layer of sediment, and the lithification process locks the shapes into place.
Other types of trace fossils are products of the action of organisms, that get preserved in the same general manner as body fossils. This includes such things as:
Something to consider about trace fossils: they were produced by the organism while it was still alive. This makes them distinct from your average body fossil, which are parts of things that had died.
Necrolysis: How long was the body exposed before burial? Did scavengers get a chance to eat it (or parts of it)? How about decay organisms, or the weather? Consider that the VAST majority of dead bodies never get a chance to get buried: they wind up in the guts of other organisms, or decayed into organic material that becomes part of the soil or sediment.
Biostratinomy: In general, the faster the rate of burial after death, the more complete the fossil will be. (This is why some of the best fossils are those from bodies trapped in some form of sediment such as tree resin (future amber) or tar or mud: the biostratinomy process began before the individual was dead!).
However, the higher the energy of the environment of burial, the greater the chance the body will be torn apart by current action. Very quiet water (like the bottom of lakes, lagoons, etc.) is very good for preserving fine details (especially of small-bodied organisms), but only if that bottom environment is not populated by worms, clams, and other sediment-churning animals: for example, if the bottom-water is anoxic.
Sometimes specimens are buried in place or in situ (also called autochthonus); other times the specimen was transported before burial (allochthonous).
Diagenesis: Once the prevailing wisdom was that most fossils were "petrified": turned to stone by chemical action. In fact, it turns out the situation is more complicated than that.
For a great many fossils the original hard parts and even some of the original organic material remains. This was not appreciated until chemical analyses began to show that the cellulose was still there in most fossil plants, the hydroxylapatite and collagen fibers in fossil bone, even the chemical residues of feathers and hair and leaf. These are definitely NOT pristine material, and they have been subject to some decay and alteration by time, but they are still there.
In some cases the original material is still there, nothing else has been added to it, and the tissue and hard parts are still basically in their original dead condition, as in this freeze-dried mammoth. Such unaltered material becomes increasingly rare the farther back in time you go.
For shells made of carbonate material, it is common to find them where the original material is still there, nothing has been added to it, but the crystals have regrown at the temperature and pressures of burial. These recrystallized fossils will lose the microscopic details of more pristine ones.
The most common mode of preservation for vertebrate bones and teeth and wood (and pretty common for shelly fossils, too) is permineralization: the original hard parts and even some organics are present, but the pore space of the bone, wood, or shell has been filled in by minerals from the pore water. This is exactly comparable too (indeed, happens at the same time as) the cementation of the sediment. Because of this, fossil bone or wood is much denser than living bone or wood; breaks along cracks; and otherwise acts like rock. But the original microscopic details -- and even chemical details -- are often still there.
Carbonized fossils are ones where compression after burial has squeezed some of the tissues (chitin; hair and leaves; cellulose and lignin) to concentrate the carbon and remove some of the other chemical components. These are preserved as thin black surfaces (in the case of wood, they can be much thicker: that is, as coal). Often the original microscopic structures of these tissues can still be found. The presence of carbonized keratin (beaks, scales, etc.) is more common than once thought: older generations of paleontologists dug right through these to get to the bones.
However, there are also situations when original fossil material is lost to some degree or other. Actual replacement of the original hard or soft parts by some other mineral does occur, as in this coiled ammonoid where the carbonate was replaced by pyrite. Recent experimental evidence suggest that chemical replacement by such minerals as pyrite, silica, phosphates, and the like actually begins quite quickly after burial if the diagenetic conditions (particular combinations of acidity or alkalinity; pressure; temperature; oxygen level; etc.) are correct.
Silicified fossils are particularly useful for understanding fine details: because silica is so chemically and mechanically resistant, it is possible to take silicified specimens and basically dissolve the rock around them, and find little projections, spikes, and other features that might normal be lost through preparation.
Sometimes the fossil itself is actually lost during diagensis: acidity in the pore water dissolved the shell or bone or wood, leaving a void space or mold. Molds are hollow spaces showing the fossil in reverse. Sometimes molds are of just one surface (an impression); other times they are the complete outside surface in reverse (external mold); or sometimes an inner space filled in (an internal mold).
These can be filled in (artificially by humans, or naturally by sediment or mineral from ground water) to produced a cast: a 3-D replica of the original fossil, where not a single atom of the cast was part of the original organism. Here is a complementary set of a mold (left) and natural cast (right) of a coiled ammonoid.
Some Lagerstätten occur because they trapped a great number of individuals, and thus concentrated the amount of specimens in one area (such as the La Brea Tar Pits of Los Angeles). Others are famous because they preserved otherwise rare soft-bodied organisms, and thus give us a much more complete picture of the biodiversity of that time and space (such as the Burgess Shale of British Columbia.)
Lagerstätten are thus of great importance for our better understanding of the ancient world.