Folds and folding - Introducing Tectonics, Rock Structures and Mountain Belts - page 56

shear zone are displaced in opposite
directions in the same way as fault
movements (Figure 6.12A). Since they
are a relatively high-temperature
phenomenon, they are found only in
metamorphic rocks, and are particu-
larly important in Precambrian gneiss
complexes. It has been suggested that
large sub-horizontal shear zones may
constitute the typical structure of the
lower part of the continental crust.
Shear zones are found on all scales
from a few millimetres in width to
many kilometres. Some of the largest
shear zones of the Precambrian base-
ment of southern Greenland are up
to 50 kilometres across and comprise
both thrust-sense and strike-slip types.
Major faults in the upper brittle crust
will pass downwards into shear zones
at depth, as shown in Figure 6.12B.
In an ideal shear zone, the strain
increases from zero at one wall of the
zone to a maximum in the centre of the
zone, and decreases to zero again at
the other wall. Such ideal zones may be
found in otherwise undeformed igneous
rocks containing no previous structure,
as in Figure 6.12C. In this case, the strain
pattern can be seen by the variation in
shape and orientation of the deformed
crystals. Generally, however, shear
zones occur in previously deformed
rocks such as gneisses, where the strain
can be deduced from the bending and
thinning of the gneissose banding.

in Chapter 4, the feature that governs
whether rock behaves in a brittle or
ductile manner was strain rate (which
in turn is controlled by temperature and
pressure); that is, a more rapid strain
rate would lead to fracturing, whereas
a lower strain rate was necessary to
achieve ductile behaviour. Conse-
quently, a necessary precondition for
a rock to produce folds would be a low
stress (but still above the yield point of
the material) applied for a long period
of time, thus giving a low strain rate. An
applied stress that was too large would
merely cause the rock to fracture, which
would immediately relieve the stress. It
follows that the stress must be applied
over periods measured in years rather
than seconds, if appreciable ductile
strains are to be achieved. Large folds
will take thousands of years to form.
Another important precondition
for folding is that the rock must have a
layered structure - folds cannot form
in a homogeneous rock. Ductile strain
in an unlayered material would result
in microstructural changes spread
throughout the whole rock, ultimately
producing a fabric ( see Chapter 7).
Moreover, the layering has to fulfil
certain conditions before folding can
take place. Experimental work has
shown that buckling will only occur if
relatively strong layers are separated
by material that is very much weaker
(perhaps by a factor of 1/50) capable
of flowing in such a way as to accom-
modate to the shape of the buckled
layer ( see Figure 6.9). Chevron folding,
as we have seen, also has quite strict
requirements in terms of the physi-
cal properties of the folded layers.

Ideal similar folding, on the other
hand, requires the rock layering to
behave in a wholly passive manner, that
is for 'flow' to occur across the layering.
If any of the layers are strong enough
to buckle, the resulting folds will not
be truly similar in style. The condi-
tions for flow folding to occur are that
the ductile flow-folded material must
be constrained by a layer that behaves
actively, that is, the flow folding should
be controlled by movement of the
boundaries of the flow-folding system.
This principle is most easily demon-
strated by the example of a shear zone,
whose deformation is controlled by
the lateral movement of its walls ( see
Figure 6.13A), or the glacier folding of
Figure 6.8. The same applies to the flow
folding exhibited by weak material con-
strained between two strong buckled
layers - here, again, it is the geometry
and movement of the strong layers
that control the flow-folding pattern.
These examples illustrate an impor-
tant principle, which is that folding is
largely dependent on the relative move-
ment of blocks of rock; in other words,
it is more helpful to understand folding
in terms of a kinematic (or movement-
based) system rather than a dynamic (or
stress-driven) one. Many fold systems
are constrained by major fault surfaces
(especially thrust planes - see Chapter
10) and their geometry is controlled
by the relative movement of the fault
sheets that form their boundaries.
Going up to a larger scale, it is the con-
vergent plate movements at subduction
zones and collisional plate boundaries
that drive the deformation in fold belts.

Why do folds form?
The reasons for the formation of folds
may not be immediately obvious.
Referring back to the discussion of the
behaviour of materials under stress