Lecture 2a: Igneous classification, mid-ocean ridges
• Questions
– How many igneous rock names should you learn, and what
do they mean?
– What is a mid-ocean ridge and how does it work?
– How does plate spreading produce the oceanic crust?
– How do rocks observed at ridges constrain the state of the
upper mantle?
• Tools
– Field and shipboard geology and analytical chemistry
– Thermodynamics
– Fluid dynamics
• Reading: Grotzinger et al., chapter 4; Albarède, chapter 8
1
Igneous Classification
• Igneous rocks can be classified according to composition,
mineralogy, texture and/or locality(!).
• The first distinction is between volcanic and plutonic rocks.
– Volcanic rocks are erupted at the Earth’s surface and cool
very quickly. There is insufficient time to grow large
crystals. This leads to formation of glass or very finegrained rocks, or to phenocrysts (crystals that grew before
eruption) in a fine groundmass.
– Plutonic rocks crystallize at some depth, and therefore lose
heat relatively slowly. Crystals have time to grow after
nucleation, and the resulting rocks generally have individual
crystals large enough to see unaided.
– Rocks of exactly the same composition and mineralogy get different
names in their volcanic and plutonic forms, because they look different!
2
Plutonic vs. Volcanic
3
Classification by mineralogy
• The standard classification scheme uses the mineralogy of the
rock (how much quartz, how much plagioclase, etc.)
– There is one important twist…for volcanic rocks you usually
cannot measure the actual minerals present (or it may be a
glass and there are no minerals present).
– In this case, instead of the actual minerals, you classify
based on normative mineralogy
• The norm is a calculation based on the bulk composition
of a volcanic rock, for what minerals would be present if
it were fully crystallized.
• The standard norm calculation is called the CIPW norm,
after Cross, Iddings, Pirsson, and Washington (1902).
4
Classification by
mineralogy
Mineral content (actual or
normative) of the rock by
volume is divided into
Quartz (Q), Alkali feldspar
(A), Plagioclase (P),
Feldspathoids (F), and
Mafic minerals like
amphibole, biotite,
pyroxenes, and olivine (M).
For rocks with Q+A+P+F > 10%
Quartz
Quartz-rich
granitoids
Alkali Feldspar
Granite (Rhyolite)
Alkali
Feldspar
For rocks with M < 90%,
the Streckeisen doubletriangle is used. It shows
names defined by Q-A-P-F
recalculated to 100%.
Many of these names are
really obscure; don’t try to
learn all of them.
Granite
(Rhyolite)
Quartz Alkali
Feldspar Syenite
(Trachyte)
Alkali Feldspar
Syenite (Trachyte)
Nepheline-bearing
Alkali Feldspar
Syenite (Trachyte)
Tonalite
Trondhjemite
Plagiogranite
Granodiorite
(Dacite)
Quartz
Monzonite
Quartz
Syenite
(Quartz Trachyte)
(Quartz Latite)
Quartz Diorite
(Quartz Andesite)
Quartz
Monzodiorite
( Andes ite)
Syenite (Trachyte) Monzonite (Latite) Monzogabbro
Nepheline-brg Nepheline-brg Nepheline-brg
Syenite (Trachyte) Monzonite (Latite) Monzogabbro
Nepheline
Monzosyenite
(Tephritic
Phonolite)
Nepheline
Monzodiorite,
Monzogabbro
(Phonolitic
Tephrite,
Basanite)
Diorite (Andesite),
Anorthosite, Gabbro,
Norite (Basalt)
Plagioclase
Nepheline-bearing
Diorite (Andesite)
or Gabbro (Basalt)
Nepheline Diorite,
Gabbro (Tephrite,
Basanite)
Nepheline Syenite
(Phonolite)
Ijolite,
Nephelinite,
Leucitite
Feldspathoids
5
Classification by mineralogy
For rocks with mostly mafic minerals, a different scheme is used. The proportion of
olivine, orthopyroxene, clinopyroxene, and plagioclase locate a rock using the
appropriate Streckeisen ternary diagram.
plagioclas e
anorthosite
90
(leuco-)
dunite
65
gabbro or
norite
90
troctolite
olivine gabbro
or norite
harzburgite
35
wehrlite
peridotites
lherzolite
(mela-)
plag-bearing ultramafic rocks
pyroxene s
40
olivine
orthopyroxenite
10
olivine websterite
olivin e
to plagioclas
e
orthopyroxenite
gabbro or norite
norite
olivin e
10
orthopyroxen
websterite
e
olivine
clinopyroxenite
pyroxenites
clinopyroxenite
clinopyroxen e
gabbro
cpx norite
opx gabbro
plag-bearing ultramafic rocks
clinopyroxen e
orthopyroxen e
6
Classification by composition
• There are several classifications, of individual rocks or rock suites.
• By silica percentage:
%SiO2
>66
52-66
45-52
<45
Designation
Acid
Intermediate
Basic
Ultrabasic
% Dark Minerals
<40
40-70
70-90
>90
Designation
Felsic
Intermediate
Mafic
Ultramafic
Example rocks
Granite, rhyolite
Diorite, andesite
Gabbro, basalt
Dunite, komatiite
• By alumina saturation (this controls which dark minerals show up):
Chemistry
Designation
Al2O3>Na2O+K2O+CaO Peraluminous
Na2O+K2O+CaO>Al2O3 Metaluminous
& Al2O3 > Na2O+K2O
Al2O3 ~ Na2O+K2O
Subaluminous
Al2O3 < Na2O + K2O
Peralkaline
Distinctive Minerals
Muscovite, biotite, topaz,
corundum, garnet, tourmaline
Melilite, biotite, pyroxene
hornblende, epidote
Olivine, pyroxenes
Sodic pyroxenes & amphiboles
7
Classification by composition
• By Alkali-Lime index: for a suite of rocks, CaO and Na2O+K2O are
plotted against SiO2. Generally, CaO decreases with increasing SiO2
while Na2O+K2O increases. Suites are classified by the SiO2 where the
intersection occurs:
Rock Suite
Calcic
Calc-alkaline
Alkali-calcic
Alkaline
Alkali-Lime Index
>61 %SiO2
56-61%
51-56%
<51%
Illustrative rock series
Mid-ocean ridge basalts
Continental margin arc series
Some intraoceanic island arcs
Intraplate continental melts
8
Ocean crust geology
• Recall the typical sequence
of rocks observed in
ophiolite exposures and in
drilling the ocean crust:
• Deep-sea marine sediments
• Massive sulfide deposits
• Pillow basalts
• Sheeted basaltic dikes
• Layered gabbro
• Serpentinized peridotites
• This sequence is consistent
with the seismic velocity
profile of oceanic crust:
9
Ocean Crust Geology
Modern and ancient pillow
basalts, photographed by
submersible and by field
geologist, respectively.
10
Ocean Crust Geology
Modern and ancient sheeted
dike complexes observed by
seismology and by field
geologist, respectively.
11
Ocean Crust Geology
Modern and ancient
layered gabbro (oceanic
layer 3) complexes
observed by ocean drilling
and by field geologist,
respectively.
12
Ocean Crust Geology
Modern and ancient
harzburgite/dunite
uppermost mantle
assemblages:
An abyssal peridotite and
the Muscat massif in
Oman
13
Mid-ocean Ridges
Now let’s discuss the origin of this rock sequence in the detailed mechanisms and
variations that occur along ridges.
• Fluid dynamics: plate-driven flow, internal buoyancy
– A ridge is a viscous fluid asthenosphere overlain by an
initially very thin rigid lithosphere that is being pulled
laterally apart along a linear rift by externally-imposed
forces. Approximately two-dimensional flow.
– Flow is slow enough to neglect inertia (tiny Reynolds
number); hence the 2-D stream function y (a scalar function
whose contours are parallel to the flow and whose spacing is
proportional to velocity) for incompressible flow satisfies
the biharmonic equation:
 y 0
4
14
Mid-ocean ridge dynamics
• For uniform viscosity, the steady solution to this corner flow is
given by Batchelor (1967)
Half-spreading rate -U
Half-spreading rate U
Stream function for constant viscosity corner flow
vertical exaggeration 10:1
symmetry boundary condition
15
Mid-ocean ridge dynamics
• Things to note about this flow field:
– It fills all of half-space, out to infinite depth and infinite
lateral extent.
– The pressure goes to –∞ at the corner!
– Flow under the ridge axis is near vertical, flow far out to the
side is nearly horizontal, but there is a positive upward
vertical component everywhere.
• There are real complexities superposed on this simple solution
that should be noted right away:
– The lithosphere thickens away from the ridge axis with the
square root of time, so the upper boundary of the
asthenospheric flow is not horizontal. Streamlines get
gradually incorporated into the lithosphere and the depth of
material points becomes fixed.
16
Mid-ocean ridge dynamics: More complexities
• The viscosity is NOT constant. Viscosity variations are
due to:
 Temperature – to first order this is the difference between
lithosphere and asthenosphere, but it is not really a sharp
cutoff.
 Strain Rate – the viscosity may be strain or stress-dependent,
so that areas flowing fast are weaker and tend to concentrate
strain ever more.
 Melting – the presence of melt above the solidus may
weaken the partially molten region
 Water – ~100 ppm H2O in mantle olivine lowers viscosity
by a factor of ~500. This H2O is removed to the melt phase
when melting begins, so melting may increase the viscosity!
• The viscosity of mantle flow in and near the melting
region is somewhere in the range 1018 to 1023 Pa-s
17
Mid-ocean ridge dynamics: More complexities
• Drag from upper boundary condition is not the only
driving force for flow. There is also internal buoyancy.
due to:
 Presence of retained melt in the melting region. Liquids are
less dense than mantle minerals at these pressures. The more
melt is retained, the more buoyant the rock.
 Change in composition of the residue with progressive melt
extraction. Melting removes Fe and garnet-forming
components, so solid residue becomes less dense even as
low density melt is being extracted!
• [This is not non-physical; there is no law of conservation
of volume]
• The flow is not incompressible – both melt production and
melt extraction cause changes in the volume of the “fluid”
(really a multiphase medium).
18
Mid-ocean ridge dynamics: More complexities
Here is a model incorporating these complexities, showing cases
dominated by plate spreading and by internal buoyancy. The R
parameter depends on spreading rate, viscosity, and buoyancy of
retained melt. Color is melt-filled porosity, white streamlines are solid
flow, black streamlines are melt flow.
19
Melting and melt extraction
• When mantle is dragged upwards towards the ridge by plate
spreading, it is initially solid. Why does it melt? How does it
melt?
– (see my chapter from the Encyclopedia of Volcanoes)
• The upwelling path is (to very good approximation) adiabatic
and reversible and therefore isentropic. The drop in temperature
per unit pressure at constant entropy in the absence of phase
changes is
T   TV
PS Cp
– For upper mantle this slope is ~10 K/GPa.
– The solidus of fertile peridotite has a slope of ~130 K/GPa
• It follows that unless the mantle is really cold (and it is not), the
adiabat will intersect the solidus upon decompression
20
Melting and melt extraction
• The intersection of adiabat and solidus depends on the potential
temperature of the mantle (the temperature where the adiabat
would reach 1 atmosphere if no melting took place).
• For ordinary regions, this is close to 1350°C (~1250 to ~1500
global range) and the solidus is crossed at about 2.5 GPa (~1.5
GPa to ~6 GPa global range).
1475°C (plume)
1500
Partially
Molten
Peridotite
1400
1350°C (normal)
Subsolidus
Adiabats
(stable)
Metastable
Subsolidus
Adiabats
Temperature (°C)
1300
1250°C (cold)
Peridotite Solidus
(schematic)
1200
1100
Potential
Temperature
0
1
2
Pressure (GPa)
Solid
Peridotite
3
4
21
Melting and melt extraction
• Once melting begins on an upwelling streamline, it will
continue until either
– upwelling ceases
– conduction to the surface cools the system
– internal thermochemistry of the residue slows or stops melting
• The cartoon version of the melting regime under the ridge axis
is then drawn on the cartoon version of the flow field as some
generally triangular shape in cross-section:
ridge axis
solidus
22
vertical exaggeration 10:1
The mid-ocean ridge Melting Regime
• The bottom of the melting regime is horizontal if the potential
temperature is locally uniform
• The upper edges of the triangle either represent where the flow
lines become effectively horizontal or where cooling from the
surface has penetrated to the relevant depth
• The extent of melting increases upwards along each streamline
from the solidus intersection to the exit of the melting regime
(indicated by shading in the figure)
ridge axis
solidus
23
vertical exaggeration 10:1
The mid-ocean ridge Melting Regime
• The erupted melt is going to be some average of the
melts produced throughout this melting regime
– Several models are possible of how and where the melt is
extracted and what happens to it during transport
• This average melt is primary mid-ocean ridge basalt
(MORB).
The melting regime is wide
(hundreds of kilometers), but
eruptions are focused in a
neovolcanic zone only a few
km wide – melts have to be
focused somehow to the
ridge axis.
This figure is the result of a
large project to seismically
image the melting regime on
the East Pacific Rise.
24
Melt productivity
0.3
After melting begins,
the P-T path and the
amount of melt
production can be
inferred from
conservation of
entropy.
At constant total
entropy, the only
way for the system to
find the entropy of
fusion is by cooling
more steeply than the
slope of the solid
adiabat.
Adiabatic
Melting Paths
0.2
(exhaustion of
clinopyroxene)
Extent of
Melting (F)
0.1
0
0
1
3
4
1500
Adiabatic
Melting Paths
(stable)
1400
Subsolidus
Adiabats
(stable)
Temperature (°C)
1300
Peridotite Solidus
(schematic)
1200
25
2
Pressure (GPa)
1100
0
1
2
Pressure (GPa)
3
4
Melt Migration
• There are multiple mechanisms to get the partial melt out of the
residue and deliver it to the crust at the ridge axis.
• Porous flow
– The melt phase quickly forms an interconnected network along the grain
boundaries of the rock (mantle source rocks are made of crystals about
1mm-1cm in diameter).
– Interconnection occurs because the energy of
a melt-crystal interface is low enough that the
melt adopts a high-surface area geometry.
– This geometry implies that the partially
molten rock is permeable to flow of the melt
relative to the solid.
– Permeable flow is governed by D’Arcy’s Law
v
k
mf
P
where P is pressure, k is permeability, f is porosity or
melt fraction, m is melt viscosity, and v is the velocity
of melt relative to solid.
26
Melt Migration
• Pressure gradients driving porous flow of the melt arise from
buoyancy of the melt (density of mantle ~ 3300 kg m-3, density
of basaltic liquids ~2700 kg m-3) and from shear and bulk
deformation of the solid matrix.
• Permeability is an increasing function of melt-filled porosity
and grain size.
– Exact function is unknown, but it may be k ~ d2f2 or d2f3.
– For 1 mm grains and a melt-filled porosity of 1%, the
permeability is ~10–11 cm2.
– For basaltic melt with viscosity ~ 10 Pa s, v ~10 mm/yr, not
much faster than solid flow rates!
– We know melt somehow eventually moves much faster than
this, so there must be other melt flow mechanisms.
27
Melt Migration
• How do we know that the melt-filled porosity resulting from
melting and melt migration is ~1%?
– Cannot guess from balancing melting rate and migration
rate, since permeability function is poorly known
– But we have seismology, which does not show the
extraordinarily low velocities that would be expected for
large melt fractions.
And we have
chemistry of residues
(There is a homework
question on this…)
28
Melt Migration
• During porous flow at low melt fractions, melt and solid remain
in chemical equilibrium. But primary MORB is not in
equilibrium with uppermost oceanic mantle. So again,
something else must happen.
• Fast porous flow
– Porosity waves: The equations for porous flow in a viscous two-phase
system allow solutions in the form of solitary waves that may propagate
much faster than the D’Arcy velocity and speed up melt transport
(contours of porosity; gridded area is <1%; this figure shows the collision
of two magma solitons moving upwards in a deformable porous medium)
29
Melt Migration
• High-porosity zones, reactive transport channels: focusing the melt flow
This may arise due to
reactive infiltration
instability: porous flow of
liquid breaks up into
channels if it can dissolve the
pyroxenes, leaving high melt
fraction and olivine residue.
The evidence for this process in nature is
veins of dunite (pure olivine) in
ophiolites: if melt flows through the
dunite channels, this may explain the
observation that MORB is not in
equilibrium with abyssal peridotite
(when you find dunites, these are in
equilibrium with MORB).
30
Formation of the crust, magma chambers, differentiation
• End result of mantle melting and melt migration is delivery of
primary magma into the crust.
• Too dense to erupt, forms magma chambers instead.
• Cools by conduction and hydrothermal convection and begins
to crystallize, with two essential consequences:
– the formation of a gabbroic lower crust from the crystallized fraction
– the compositional evolution of the remaining melt before eruption.
Pprimary MORB crystallizes olivine, then
olivine+plagioclase, then olivine+
plagioclase+clinopyroxene. These are the
minerals of the oceanic lower crust and also
the phenocryst phases in erupted MORB.
31
Evolution of oceanic crust
Olivine+plagioclase
fractionation
Troctolite:
24% Al2O3
13% MgO
Olivine+plagioclase
+clinopyroxene
fractionation
Gabbro:
17% Al2O3
11% MgO
Olivine
fractionation
Dunite:
Primary
Liquid
+titanomagnetite or
ilmenite
0% Al2O3
50% MgO
Formation of the crust, magma chambers, differentiation
• Where does the differentiation take
place?
– Phase equilibria indicate low pressure
(0-2000 bars)
– Seismic images show melt lens ~1 km
below the axis of the East Pacific Rise
– Crystallization takes places in the
shallow melt lens or in mush zone in
lower crust
33
Primary liquids?
• If the primary liquid never erupts, what can we say about
melting conditions in the mantle? How do we see back through
the filter of differentiation?
• One solution is to pick an arbitrary MgO content at which to
compare data and models. MgO decreases throughout
differentiation, so it is a good index to normalize against.
3.5
Here is the correction of data to 8%
MgO. Note that variation along the
3.0
liquid line of descent is the first
principal component of variability
Na2O
within basalt samples from a given
2.5
location, but there remains local
variability at 8% MgO, and when each
2.0
region is averaged there are systematic
differences between regions
Kane
FAMOUS/AMA
R
1.5
6
34
7
8
Mg O
9
10
Global systematics
• The values of regionally-averaged Na8 (i.e., Na2O concentration
corrected to 8% MgO), Fe8, water depth above the ridge axis, and
crustal thickness show significant global correlations.
– Where Na8 is high, Fe8 is low
– Where Na8 is high, the ridges are deep
– Where Na8 is high, the crust is thin
3.5
3.0
Na8.0
2.5
2.0
1.5
6
7
8
9
Fe8.0
10
11
12
35
Global systematics
• What is the interpretation of the global correlations in Na8, Fe8,
axial depth, and crustal thickness?
•
•
Answer: Na8, an incompatible element, is an indicator of mean extent of
melting. Fe8 is an indicator of mean pressure of melting. Axial depth is an
indicator of mantle temperature, extent of melting, and crustal thickness
combined.
So the global correlation implies that extent of melting and pressure of melting
are positively correlated, on a global scale.
4
Regional avg. MORB data at 8% MgO
3
Iceland
2 Na2O
DNa =0
1
Forsyth
(1992) DNa =0.06
(wt %)
1
10
Crustal Thickness (km, log scale)
0
100
36
Synthesis of global systematics
• The correlation of extent of melting with pressure of melting
requires that the first-order control on variation among ridge
segments is potential temperature.
• If melting continues under the axis to the base of the crust
everywhere, then high potential temperature means: long melting
column  high mean extent of melting  low Na8 and high
crustal thickness  shallow axial depth; high mean pressure of
melting  high Fe8. Cold mantle yields the opposite.
Hot mantle
Cold mantle
sea level
axial depth
crust
mean
P
25%
20%
15%
10%
5%
40%
F
mean
F
35%
solidus 1.5 GPa
30%
F
mean
P
solidus 4.5 GPa
25%
20%
15%
10%
5%
mean
F
37