Transcript Slide 1

Sedimentary Rocks
Francis, 2014
Sedimentary rocks form by surface processes with which we have every day experience. Anyone who has
wandered outside has observed the sedimentary processes of erosion, sediment transport, and sediment deposition
by water, especially during the spring thaw or after heavy rains. Much more information can be extracted from
sedimentary rocks in the field than in the case of igneous or metamorphic rocks, which typically require extensive
laboratory analysis before revealing their secrets. It is precisely because they record everyday surface processes
that sedimentary rocks have now become so important to the documentation of long term climate and tectonic
trends in the Earth’s past.
Sedimentary rocks cover 65 - 70 % of the Earth’s surface and constitute
approximately 8% by weight of the Earth's crust, in the proportions:
shales (<0.06 mm)
sandstones (2 - 0.06 mm)
carbonates
others
-
65 %
20 %
10 %
5%
On the continents, the thickness of the sedimentary cover ranges from 0 metres on Archean
shields to more than 20,000 metres in marginal basins, and averages approximately 1800 metres
on stable continental platforms.
1. Terrigeneous siliciclastic or epiclastic fragments transported by
gravity flows, water, wind, and ice, including:
Sedimentary rocks form as the
result of physical, chemical,
and biological processes that
can be grouped into 4 basic
categories:
breccias
conglomerates
sandstones
mudstones
diamictites
2. Biogenic precipitates, including:
limestones
dolomites
chert
phospates
coal
3. Chemical precipitates, including:
evaporites
iron formations
chert
4. Volcanoclastics, including:
ignimbites
tuffs
agglomerates
Relevance of Sedimentary Rocks and Stratigraphy
Paleo-environmental and tectonic setting indicators:
On a local scale, sequences of sedimentary rocks are the key indicators of depositional
environment and tectonic setting. For example, something as simple as the sequential
interbedding of sandstones and shales can be used to determine whether the rocks you are
looking at formed in a deep-sea fan, or in a meandering river system.
Braided
versus
Meandering
Paleo-climatic indicators
The current interest in climate change has heightened interest in the study of
sedimentary sequences on a global scale as the systematic recorders of past climate
change.
For example, the detailed stratigraphic analysis of oxygen isotope
compositions in limestone successions can be used to chart changes sea level,
temperature, and compositional in time over the Earth’s history.
Sea Level Variations
Late-Tertiary : 12 Mys
Phanerozoic : 500+ Mys
sequence
boundary
Analyses of the climate variation over
the last few 100 thousand years
reveals a strong inverse correlation
between the concentration of CO2 in
the atmosphere, δ18O, and the volume
of polar ice.
Present CO2 concentration ~ 387 ppm
Mount Sharp
Gale Crater
Mars
Lower Mound Unit
or Formation
Mount Sharp, Gale Crater
Mars
Hosts to petroleum and economic minerals reserves:
Many important types of mineral deposits occur in sedimentary rocks; including:
- both exhalative and Mississippi valley-type Pb-Zn deposits.
- red bed copper deposits.
- paleo-placer deposits of Au and U.
- Carlin-type gold deposits
- Wyoming-type U deposits.
- sources of salt, gypsum, phospates, nitrates, manganese, etc.
Sedimentary rocks form both the source rocks and the reservoirs for hydrocarbon reserves.
Indicators of metamorphic grade
We shall see in the final third of this course that mineral reactions in meta-sedimentary
rocks are particularly useful for constraining pressures and temperature histories during
metamorphism. This reflects their distinctive Al-rich composition compared to igneous
rocks, a result of weathering that originally created the sediment.
Weathering
The majority of igneous minerals formed at temperatures in excess of 700 oC and are unstable
under atmospheric or submarine conditions. The action of weathering converts these high
temperature minerals into three distinct types of material:
1) elements which are dissolved by water and carried off in solution:
- K, Na, Ca, Mg
2) materials which are ‘resistates’ and carried as clastic fragments:
- quartz, rutile, zircon, diamond, garnet, feldspar, magnetite
3) newly formed insoluble minerals and amorphous phases formed by the
breakdown of unstable phases such as feldspars and mafic silicates:
- clay minerals, hydroxides of Fe, Mn, and Al, silica
Physical Weathering:
Frost action:
Water expands 9.2 % on freezing
and is capable of generating
pressures on the order of 150 bars
within rock fractures.
The
effectiveness of this process is
highly latitude dependent.
Most effective at high altitudes
and latitudes that experience
numerous freeze–thaw cycles,
results in alpine or arctic
felsenmeer.
Physical Weathering:
Diurnal temperature variation:
Debatable effectiveness. Observations
on Egyptian monuments indicate that
erosion is greatest on permanently
shaded sides rather than those exposed
to the sun. This is thought to reflect the
action of moisture condensation on cool
surface.
Physical Weathering:
Sheeting:
Many massive rock lithologies such, as granite, develop a sheet-like jointing parallel to the surface
due to expansion following unloading by erosion. Expansion due to decreasing pressure during
unroofing also tends to disaggregate coarse grained rocks such as granitoids, forming sand size
clastic grains.
Physical Weathering:
Biologic activity:
Plants and animals exert both physical and chemical effects that act to breakup and decompose
rocks. Lichens, for example, both physically pluck minerals grains from rock surfaces and secrete
organic acids that extract soluble cations from minerals, accelerating their breakdown in the
presence of water.
Oxidation
magnetite
2Fe3O4
+
pyroxene
4FeSiO3
olivine
Fe2SiO4
+
+
1/2O2
O2
1/2O2
hematite
3Fe2O3
hematite
silica
2Fe2O3 + 4SiO2
hematite
Fe2O3
+
silica
SiO2
Hydration
pyroxene
3MgSiO3
serpentine
+
2H2O
hematite
Fe2O3
silica
Mg3Si2O5(OH)4 + SiO2
goethite
+
H2 O
2FeOOH
pyrite
goethite
2FeS2 + 5H2O + 15/2O2
2FeOOH + 4SO4= + 8H+
Hydrolysis
Exchange of H+ ions in water for a soluble cations in mineral.
This is the process by which soluble cations are leached from silicates, dissolved
and carried away by water.
orthoclase
2KAlSi3O8 + 9H2O +
kaolinite
2H+
Al2Si2O5(OH)4 + 4H4SiO4 + 2K+
pyroxene
MgSiO3 + H2O + 4H+
Mg2+ + H4SiO4
anorthite
CaAl2Si2O8 + H2O +
kaolinite
2H+
Al2Si2O5(OH)4 + Ca2+
albite
kaolinite
2NaAlSi3O8 + 9H2O + 2H+
Al2Si2O5(OH)4 + 4H4SiO4 + 2Na+
Silt-Sized Silica (SiO2)
All the above reaction produce clay minerals and silicic acid at the expense of the
original feldspars and mafic silicate minerals. Although some of the silica produced in
these reactions is removed in solution as silicic acid (H4SiO4), the low solubility of this
acid means that many of the above weathering reactions produce silica, as well as silicic
acid, which remains as silt-sized crystalline quartz.
The majority of the silt-sized silica in shales, which are mixtures of fine quartz
grains and clay minerals, is probably produced in this way.
The proportions of crystalline silica versus silicic acid is controlled by the equilibria:
SiO2 + 2H2O
H4SiO4
We will look more closely at this reaction when we discuss the formation of cherts.
Clays
Two - sheet phyllosilicates:
Kaolinite
Al2Si2O5(OH)4
Serpentine
Mg3Si2O5(OH)4
The basic 2-sheet kaolinite-layer building block of tetrahedrally and octahedrally
coordinated cations has no net charge, and thus no K+ or Na+ formally bonding the layers
together. Additional cations and water are, however, commonly loosely bound between
these sheets.
tetrahedral sheet
octahedral sheet
Three - sheet phyllosilicates:
Clays
Pyrophyllite
Al2(Si4O10)(OH)2
Talc
Mg3Si4O10(OH)2
Montmorillonite (Al1-XMgX)2Si4O10(OH)2 + XNa+.nH2O
Micas
In the micas, an Al substitutes for 1 of every 4 positions in the tetrahedrally-coordinated sheet
(AlSi3), giving the layer a net negative charge, which is balanced by large K+ and or Na+ cations
between the layers, bonding them together.
Muscovite
Paragonite
Phlogopite
KAl2(AlSi3O10)(OH)2
NaAl2(AlSi3O10)(OH)2
KMg3(AlSi3O10)(OH)2
tetrahedral sheet
octahedral sheet
tetrahedral sheet
Montmorillonite
Montmorillonite is a clay mineral that is intermediate between kaolinite and mica. Many
of the weathering reactions we looked at previously form the clay mineral
montmorillonite:
(Al1-XMgX)2Si4O10(OH)2-X + X×Na+.nH2O
as an intermediate step on the way to the formation of kaolinite. The substitution of some
Al3+ by Mg2+ gives the octahedral sheet a negative charge, which is balanced by large
low-valence cations such as Na+ and K+ between the sheets holding them together.
Commonly kaolinite layers are intermixed with montmorillonite layers on a
submicroscopic scale. The process of lithification involves the reverse reaction with a
progressive gradation from two-sheet clay minerals to three-sheet micas with increasing
temperature and pressure.
Clay Minerals
I
S
Chl
K
Q
=
=
=
=
=
illite (muscovite)
montmorillonite (smectite)
chlorite
kaolinite
Many clays
quartz
Scanning electron
microscope image
of a clay mineral
are in fact
mixtures
of
different
interlayered clay minerals.
X-ray
Diffractogram
300oC
kaolinite / montmorillonite
illite
/ muscovite
Extreme
chemical
weathering
reduces most rocks to mixtures of
kaolinite, Al(OH)3, FeOOH, and
silica, with the other elements
largely being lost to solution.
Weathering Profile Developed in Columnar Jointed Basalt
The relative loss or
enrichment of any element
during
chemical
weathering
can
be
estimated if 1 element,
such as Al is assumed to
be constant. Al, along
with Fe3+ are very
insoluble
in
surface
waters. If Al is assumed
to be unchanged by
weathering, other than by
a closure effect, then the
apparent increase in Al
during
weathering
is
simply a function of the
loss of other constituents.
Parana
Basalt
Saprolite
%
loss
Mesozoic
Granodiorite
Soil
SiO2
50.7
35.1
-67.6
67.92
55.46
TiO2
3.19
3.96
-42.0
0.55
1.53
Al2O3
12.3
26.3
0
14.70
21.37
Fe2O3
15.1
21.5
-33.5
0.91
8.65
2.61
0.37
FeO
MnO
0.20
0.10
-75.0
0.03
0.03
MgO
4.18
0.26
-97.1
0.98
0.56
CaO
7.83
0.06
-99.6
2.94
0.38
Na2O
2.53
0.02
-99.6
3.31
0.49
K2O
1.71
0.13
-96.5
4.38
6.48
P2O5
0.46
0.11
-89.1
0.18
0.20
LOI
1.70
12.4
+240
1.16
4.55
%
loss
Al2O3 :
Al2O3original = Al2O3lost + Al2O3remaining
12.3 = (0 × (1-X)) + (X × 26.3), where X is the wt. fraction of the original
rock remaining.
X = 12.3 / 26.3 = 0.467
This means that 53.3 wt.% of the original rock has been lost. The amount of
any other element remaining from 100 gm of starting fresh rock, such as Si is:
Remaining SiO2 = 0.467 × 35.1 = 16.4 gm
Lost SiO2 = 50.7 – 16.4 gm = 34.3 gm
The relative change in SiO2 is – 34.3 / 50.7 or -67.6 %
As an exercise, try calculating the % loss or gain of each element during the
weathering of the Mesozoic granodiorite in the forgoing table.
Chemical equilibria of
the
weathering
of
orthoclase to gibbsite:
The results of an interaction between
a mineral like orthoclase and surface
water is a function of the water /
rock ratio, in addition to water
composition and temperature. At
low water / rock ratios, the water is
forced to come into equilibrium with
the mineral, whereas at high water /
rock ratios, the rock is forced to
come into equilibrium with the fluid.
microcline
KAlSi3O8 + 7H2O + H+
gibbsite
Al(OH)3 + 3H4SiO4 + K+
Chemical equilibria of the weathering of
microcline to gibbsite:
If the amount of fluid is relatively small with respect to the
amount of microcline, then the fluid will change
composition as microcline is converted to gibbsite. In the
extreme case of there being an infinite amount of
microcline, the composition of the fluid will change to
eventually come into equilibrium with microcline. At E,
where orthoclase, kaolinite, and muscovite coexist, the Ph
and K activity of the fluid are fixed or buffered by the
mineral assemblage to constant values.
A to B
microcline
KAlSi3O8 + 7H2O + H+
gibbsite
Al(OH)3 + 3H4SiO4 + K+
B to C
gibbsite
2Al(OH)3 + 2H4SiO4
kaolinite
Al2Si2O5(OH)4 + 5H2O
C to D
microcline
2KAlSi3O8 + 9H2O + 2H+
kaolinite
Al2Si2O5(OH)4 + 4H4SiO4 + 2K+
kaolinite
3Al2Si2O5(OH)4 + 2K+
muscovite
2KAl3Si4O10(OH)2 + 3H2O + 2H+
microcline
3KAlSi3O8 + 12H2O + 2H+
muscovite
KAl3Si4O10(OH)2 + 6H4SiO4 + 2K+
D
D to E
Transport
Fluid transport
water
wind
Gravity flows
turbidites
grains flows
mud flows
debris flows
slumps
Ice transport
Elementary Fluid Dynamics
Fluids by their very nature are a challenge to deal with quantitatively for those used to
thinking about the physics of solids. By definition, a fluid has no yield strength, that is it
will deform in an unrecoverable way in response to any stress, this makes them complicated.
Density:
Density is one of the most important properties of fluids, controlling dynamic quantities
such as fluid momentum, force due to momentum change, and pressure. The density of a
fluid is a function of temperature, suspended solids, and dissolved salts.
Molecular Viscosity:
Molecular viscosity is a measure of a fluid’s resistance to flow. The viscosity of a fluid is
the constant that relates shear force to rate of strain. Viscosity can be thought of as the
force needed to maintain a unit velocity difference between two unit areas a unit distance
apart.
 =  / (δV/ δy)
 = shear stress : force/unit area
V = fluid velocity
y = distance
Newtonian Fluids
Newtonian fluids have no yield strength,
and a constant viscosity at constant
temperature and pressure. Most surface
waters are Newtonian, but at very high
sediment loads become Bingham fluids
with a finite yield strength (eg. mud and
grain flows).
Einstein-Roscoe equation:
Viscosity of solid - fluid mixtures:
mix = (1 - 3.5 × X) - 2.5 × o
X = volume fraction solids
Kinematic Viscosity:
v = /
 = density
Dimensionless Parameters:
Four different forces are at play in fluid dynamics:
1) Gravitational forces proportional to g × ρ
2) Bouyancy forces proportional to g ×Δρ, due to the effect of gravity on differences in
density.
3) Inertia forces proportional to V × ρ, due to the momentum of the current flow.
4) Viscous or retarding forces proportional to , due to viscosity or resistance to flow.
Many of the physical properties of fluid flow can be scaled in terms of dimensionless
parameters that are effectively ratios of these four different forces.
Rayleigh Number:
The Rayleigh number is the ratio of buoyancy forces (α × ΔT × g × d3) to inertia
forces (K ×  / ρ) :
Ra = (α × ΔT × g × d3) / K ×  / ρ
where: α = coefficient of thermal expansion
K = thermal diffusivity
d = depth
ρ = density
μ = viscosity
Fluids with Rayleigh numbers in excess of ~ 1000 will spontaneously undergo
adiabatic convection. This also applies to solids that flow plastically such as the
Earth’s mantle, and is responsible for the adiabatic melting that occurs in the
upper mantle.
Reynolds Number:
The Reynolds number is the ratio between inertial forces due to momentum (V ×
L × ρ) and viscous or retarding forces ():
Re = (V × L × ) /  = V × L / v
L
= length scale, commonly depth, but can be many other
dimensions. For example, we can speak of the Reynolds number
(and thus the onset of turbulence) of a grain in suspension if we use
the grain diameter for L.
For Re less than about 1000, fluid flow is laminar. For Re greater than about
1000, fluid flow is turbulent. Most streams are in the turbulent flow regime.
Two fluids with different velocities, densities, or linear dimensions are thought
to be dynamically similar if they possess the same Reynolds number, that is they
will develop geometrically similar streamlines in the vicinity of geometrically
similar objects.
Froude Number
The Froude number is the ratio between inertial momentum forces
(V × ρ (1gm/cc)) and gravity forces (g × L):
Fr = V / (g × L)1/2
where: V = flow velocity and
L = water depth
The Froude number is relevant to fluid flow involving a free surface. If the
Froude number is less than 1, then the velocity at which waves travel on the
surface is faster than the fluid velocity, and waves can move up stream. Stream
flow at this condition is said to be tranquil, streaming, or sub-critical. If the
Froude number is greater than 1, then surface waves can not move upstream,
and the flow is said to be rapid, shooting, or supercritical.
Bernoulli’s Equation:
½ (ρ × V2)
kinetic
energy
+
Ρ +
flow
pressure
(ρ × g × z) = constant
gravitational
potential energy
Bernoulli’s equation is simply an expression of the conservation of energy for a
flowing fluid. If the velocity increases, the resultant increase in kinetic energy is
balanced by a decrease in flow pressure. Conversely if the flow velocity decreases,
there is a corresponding increase in flow pressure. Among other things, this equation
explains everything from the flow of water through pipes to how airplanes fly.
Stokes Law:
Gives the settling velocity for a spherical
particle via laminar flow.
V = ( × g × d2 ) / 18
or
V = C × d2
C = constant =  × g / 18
In water, Stokes law works only for
particles smaller than 0.1 to 0.2 mm,
particles (smaller than fine sand). The
settling velocities of sand-sized and larger
particles are overestimated by Stokers law
because of the viscous drag of the
turbulence in the wake of passing grain.
The relationships between fluid velocity and
the sand transport capacity are better
determined experimentally because of the
theoretical complications of turbulent flow.
Re = (V × d × ) / 
Boundary Layers:
Up until now, we have been dealing with fluids having
no boundaries. However, what will interest us most
are the effects of a boundary, such as a river’s bottom
or channel walls on fluid movement. In general there
is a boundary layer in which the velocity of the fluid
is a function of distance from a bounding surface
because of friction, in which viscous and inertia forces
give rise to shear stresses. We can define a property
called shear velocity in which shear stress is written
in terms of velocity by reversing the formula for the
definition of viscosity.
Shear Velocity:
V* =
/ρ = /ρ×V/Z
Where is  is the shear stress, and ρ is the density of
the fluid,  is the viscosity, V is the flow velocity in
the center of the channel and Z is the thickness of the
boundary layer or depth of flow.
 =  / (δV/δz)
viscosity
 =×V/Z
It is important to remember that the shear velocity is not a measured property, but one that is calculated
for the particular geometry of the situation at hand, eg. channel depth and width, bed roughness, etc. As a
rule of thumb, however, shear velocity is typically ~ 1/10 channel flow velocity
The foregoing analysis is, however, incomplete in that it assumes laminar flow. Most
natural water systems have sufficiently high Reynold’s number that flow is turbulent and
highly unsteady at any give spot. Under turbulent flow conditions, the fluid movement
involves sudden bursts and sweeps. A burst is a rise of fluid from the boundary layer into
the body of the flow, commonly observed as “boils” on the water surface. A sweep
consists of a flow from the main flow down into the boundary layer. The overall result is
rather chaotic and results in large fluctuations in flow rate throughout the boundary layer,
although the average velocity still decreases across the boundary layer towards the
confining surface.
viscous sub-layer
Even under conditions of turbulent
flow, however, there typically exists
a viscous sub-layer across which
flow is laminar, with a linear
increase with distance.
The
thickness (~ 0.5 mm – 1 cm). of the
viscous sub-layer is a function of the
bed roughness and temperature. The
colder the water and the larger the
average grain size of the bed, the
thicker the viscous sub-layer.
Individual grains which protrude
above the viscous sub-layer shed
turbulent eddies and will experience
much more fluid drag than grains
that lie wholly within the viscous
sub-layer.
Transport thresholds:
Theoretically predicting the threshold velocities
for grain movement and entrainment by a fluid,
although possible in principle, is in practice
very difficult because of the numerous
complications of turbulence, grain size, viscous
layer thickness, etc.
As a result, these
thresholds are much better estimated
experimentally.
Thresholds for erosion,
transport, and deposition:
Ripples
The first sign of grain transport in
silt to sand sized sediment is the
development of ripples on the
bottom. Ripples develop due to
instabilities in the viscous sublayer and grow to heights that
exceed the viscous sub-layer.
Their wavelengths (5 - 20 cm) and
amplitudes (cm’s) scale with grain
size, viscosity, and the thickness
of the viscous sub-layer, but are
independent of water depth.
Dunes
At higher flow velocities, larger dunes
appear on the channel bottom whose
amplitude (10’s cm - metres) and
wavelength (0.5 - 10’s of metres) scale
with grain size, the thickness of the
boundary layer, and water depth. In
shallow water, the water surface is out of
phase with the topography of the dunes,
with the shallowest depths occurring over
dune crests and the deepest over troughs,
which are marked by surface “boils”.
Eolian dunes
water dunes
Bedforms
The nature of the bedforms on the bottom
surface of a flowing channel, such as ripples
and dunes, is dependent on the current
velocity and grain size of the sediment, with
a transition from ripples to dunes to plane
beds and then anti-dunes with increasing
flow velocity.
*
Antidunes are characterized by
low-angle cross bedding, but are
rarely preserved in sedimentary
rocks because they commonly
develop in regimes of net erosion.
Water surface is in phase with
antidunes.
Characteristic
of
volcanic base surge deposits.
Plane laminations of the upper
flow regime.
The foreset beds of dunes are
commonly preserved as trough
cross bedding in sedimentary
rocks, while those of ripples are
preserved as cross laminations.
Water surface is out of phase
with dunes.
Bed Forms
Eolian Dunes
Eolian Dunes
ripples ( = 0.05 - 0.2 m) cm’s
dunes ( = 0.5 - 10 m) meters
Eolian dunes ( = 10-100+ m) 10’s meters
Ripples
Eolian Ripples on Dunes
Grain-Size versus Transport
Mechanism
There is a rough correspondence between the major
grain size divisions and the transport mechanism,
which is in return responsible for their physical
separation during the fluid transport process.
• silts and clays are carried in suspension in the
‘wash load’
• sands are carried in ‘bed load’ by intermittent
saltation and suspension
• pebbles and larger are carried in the ‘bed load’ by
traction
Grain Size Statistics:
Sediment grain size distributions are
determined using sieves of different
hole sizes for fine sand to gravel and
by measuring settling rates for clays to
silts in water columns and using
Stokes Law to determine grain size.
The results are in weight units - that is
the contents of each sieve are weighed
to determine the weight portions of
the grain size range caught by the
sieves.
median
mean
standard
deviation
skewness
kurtosis
= Φ50
= (Φ16 + Φ50 + Φ84) / 3
= (Φ84 - Φ16) / 4 + (Φ95 - Φ5) / 6
= (Φ84 + Φ16 – 2Φ50) + (Φ95 + Φ5 – 2Φ50)
2(Φ84 - Φ16 )
2(Φ95 – Φ5 )
=
(Φ95 – Φ5 )
2.44(Φ75 – Φ25)