i

The Geology of the Moine Thrust Zone on the eastern shores of

Loch Eriboll, Northwest Scotland.

An undergraduate mapping project by Stephen Gillham.

Declaration:

The contents of this thesis is the original work of the author and has not been

submitted for a degree at this or any other university. Other people’s work is

acknowledged by reference.

4th April 2015

Department of Earth and Planetary Sciences, Birkbeck College, University of

London.

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Table of Contents

1 Chapter 1 Introduction 1

1.1 Mapping Area 1

1.2 Geological setting 1

2 Chapter 2 Formations within the mapping area 2

2.1 Eriboll Formation 2

2.1.1 Cross-bedded Member 2

2.1.2 Pipe Rock 4

2.2 An t-Sron Formation 6

2.2.1 Alltain Beds 6

2.2.2Salterella 8

2.3 Tor Liath Formation 9

2.3.1 Kempie Dolostone 9

2.3.2 Heilam Dolostone 10

2.4 Stratigraphic and sedimentological evolution of the area 11

3 Chapter 3 Igneous rocks 12

4 Chapter 4 Metamorphic geology 17

4.1 Lewisian Gneiss 17

4.1.1 Observations 17

4.1.2 Interpretation 20

4.2 Arnaboll Mylonite 21

4.2.1 Observations 21

4.2.2 Interpretation 22

4.3 Oystershell Mylonite 23

4.3.1 Observations 23

4.3.2 Interpretation 23

4.4 Quartz Mylonite 24

4.5 Metamorphic history of the area 24

5 Chapter 5 Structural geology 24

5.1 Faults/shear zones 24

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5.1.1 Observations 24

5.1.2 Interpretation 30

5.2 Folds 31

5.2.1 Observations 32

5.2.2 Interpretation 34

5.3 Cleavage 34

5.4 Lineations on faults and stretching lineations 35

5.5 Stylolites and associated tension gashes 35

5.6 Structural history of the area 36

6 Chapter 6 Geological history of the area 37

6.1 Lewisian Gneiss 38

6.2 Cambrian sedimentology 39

6.3 Caledonian thrusting 40

6.4 Post orogenic events 42

6.5 Summary 42

7 References 45

8 Appendix 48

9 Acknowledgments 49

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List of figures

Figure 1- Location of mapping area v

Figure 2- Geological map and cross-sections vi

Figure 3- Skolithos shear indicators 25

Figure 4- Overstepping and duplex geometry 25

Figure 5- Lateral ramp 29

Figure 6- Fault propagation folds 36

List of plates

Plate 1- Pipe Rock 13

Plate 2- Stretched quartz in mylonite 13

Plate 3- S-C fabric in Oystershell Mylonite 14

Plate 4- Sheared Alltain Beds 14

Plate 5- Hangingwall anticline 15

Plate 6- The Arnaboll Thrust on a regional scale 15

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Abstract

The geology of the Moine Thrust Zone (MTZ) has been studied since the early 19th Century by

eminent geologists such as Lapworth, Murchison and Geikie to name but a few. As our

understanding of the MTZ has developed over the years, so too has our understanding of similar

thrust belts across the globe. Much of the thrusting within the MTZ was accommodated within the

Cambrian sediments of the Ardverck and Durness groups which show remarkable continuity along

the Moine Thrust from Loch Eriboll in the north to the Isle of Skye (BGS). The uniformity of the

Cambrian sediments have been particularly helpful in the elucidation of Moine Thrust tectonics.

This thesis is based on the author’s own observations and interpretations of the mapping area.

Field work consisted of 29 days in the field, on the eastern shores of Loch Eriboll (Fig.1), where

geological mapping was conducted using 1:10 000 scale base maps. The author found that the

thrust zone comprised of a mainly foreland propagating thrust sequence with some later

overstepping thrusts stepping back into the orogen. This is broadly consistent with the findings of

other authors such as Butler (2004) and Holdsworth et al. (2006).

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1

Chapter 1 Introduction

1.1 Mapping area

The mapping area covers the area immediately east of Loch Eriboll on the north

coast of Scotland (Fig.1). From north to south, the area includes: the Heilam

Peninsula; Ben Arnaboll; Bealach Mhairi; and Kempie which is to the east of

Bealach Mhairi (Fig.2).

1.2 Geological setting

The area covers the northernmost extent of the Moine Thrust Zone (MTZ), where

compression during the Caledonian Orogeny resulted in major thrusting in a

predominantly west-northwest direction. It is an area which has been extensively

studied since the late 1800’s due to the excellent exposure of the MTZ. In the early

1880’s, Charles Lapworth concluded that the structural complexity of the area was

due to “contractional folding and faulting”, what we now know as thrusting. Terms

such as “mylonite” and “thrusts” were first coined by Geikie when he studied the

area in 1884. In the early twentieth century, Peach et al. (1907) discovered that

faulting occurred in linked arrays. Further work by numerous authors has placed

better constraints and understanding on the geological history of the MTZ at Loch

Eriboll (Law et al., 1984; Wilkinson & Soper, 1975; Butler 1984, Holdsworth et al.,

2006, 2007). The sequence of thrusting is complex within the area with some

thrusting stepping back into the orogen (Butler, 2004) and remains a much

discussed topic.

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Chapter 2 Formations within the mapping area

2.1 Eriboll Formation

2.1.1 Cross-bedded Member

Observations

The Cross-bedded Member unconformably lies upon the pre-Cambrian Lewisian

Basement. Evidence of this unconformity can be seen around Bealach Mhairi [4510

5770], where a steeply dipping contact between the Lewisian basement and Cross-

bedded Member is clearly visible. The contact is sharp with no evidence of

tectonism, indicating that the contact is unconformable rather than tectonic. The

member is fine grained and is both texturally and compositionally mature,

consisting almost entirely of quartz grains approximately 2mm in diameter. Owing

to the compositional maturity of the member, it is likely that the original sediment

was the product of extensive re-cycling. Subordinate plagioclase is present but less

than 5%. The rock type is therefore a quartz arenite. There are no pore spaces

visible within the member, indicating substantial compaction or cementation.

Sedimentary structures are well preserved. Low angle trough cross-stratification is

clearly visible within many outcrops, where finer grained sediments on cross-

bedding surfaces have been preferentially weathered. Cross bedding is also

commonly picked out by fracturing which conforms to the curved cross-bed

surfaces. Typically, the beds are 30 to 40cm thick and are bounded by sub-

horizontal bedding planes. Cross-bedding is mainly unidirectional but there are

some sets of herringbone cross-bedding, which suggests a bi-modal flow regime

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consistent with an intertidal marine setting. Cross-bedding foresets were restored

on a stereonet where a dominant palaeocurrent direction of 106o east was

obtained (Appendix 1). Excellent examples of palaeo-ripples can be seen on the

northern end of Druim na Teanga [4540 5960]. The ripples are symmetrical in

cross-section, have a wavelength of approximately 40mm and an amplitude of

8mm. The symmetry of the palaeo-ripples suggests a bi-modal flow regime. The

ripples trend 188o to the south. The palaeocurrent would have been normal to the

crest axis of the ripples, so again palaeocurrent direction is in an east-west

orientation and in agreement with the foresets. There are no evidence of fossils

within the member. This is possibly due to moderately high sediment deposition

rates within a high energy environment, rendering it uninhabitable to benthic

fauna. The thickness of the Cross-bedded Quartz Member is difficult to constrain

precisely as it has been substantially tectonised, but is in the region of

approximately 75-100 metres.

Interpretation

Given the maturity of the Cross-bedded Member, it is clear that the original

sediments have undergone extensive re-working and is highly likely that they have

been exposed to polycyclic events to reach the level of maturity seen in this

member. Deposition was in a relatively high energy environment where sediment

influx was moderate. The presences of herringbone cross-bedding and

symmetrical palaeo-ripples on bedding surfaces are strong indications that they

were deposited in an inter-tidal zone, above wave base where tides ebbed and

flowed east to west. There are two possible reasons for the lack of bioturbation

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within the member: 1) fauna could not keep up with moderate sedimentation

rates; 2) benthic fauna had not yet evolved to inhabit such a palaeo-environment.

The latter is difficult to prove as paradoxically, with a lack of bio-stratigraphic

evidence it is difficult to date the sedimentary units.

2.1.2 Pipe Rock

Observations

The Pipe Rock Member lies conformably above the Cross-Bedded Member. The

contact can be observed by the road side, approximately 300m directly east of

Kempie [4485 5800] where beds are overturned, so that the older Cross-bedded

Member sits on top of the Pipe Rock Member. Evidence for the overturned beds

can be found within the cross-bedding of the Cross-bedded Member, where cross

laminations are clearly upside down. Compositionally, both members are similar.

In comparison with the Cross-bedded Member, the Pipe Rock is also a quartz

arenite. Grain size within the unit is fairly consistent at 1.5 to 2mm and there are

no relict pore spaces preserved. The member consists almost entirely of quartz.

There are no feldspars or any other detrital minerals visible. Grain interfaces are

interlocking but it is unclear whether this is due to cementation or pressure

solution. The lack of metamorphic fabric due to little lithostatic pressure within the

Alltain Beds directly above the Pipe Rock Member would suggest that quartz

overgrowth, rather than pressure solution was the main driver behind the inter-

locking fabric. This has implications for the rock type, as if the contacts are due to

quartz overgrowth the member would be classified as a quartz arenite, whereas if

pressure solution was the dominant cause of the interlocking fabric, the member

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would be deemed a quartzite. The lack of feldspars within the Pipe Rock Member is

an indication that the sediments within the member may have undergone a slightly

more evolved history than that of the underlying Cross-bedded Member.

There is a distinct lack of cross-bedding within the Pipe Rock, although some sub-

horizontal bedding horizons are visible within some outcrops. The most

distinguishing feature is the abundance of skolithos trace fossils. They generally

take the form of vertical burrows which vary in length, but typically are between

20 to 40cm long and the thickness of the burrows are pretty consistent at 12 to

15mm in diameter. They are clearly visible in cross-sectional view as the member

is often stained red or purple (Plate 1), whereas the burrows retain the unstained

white colour of the quartz arenite. On bedding planes, burrow entrance holes are

visible as distinctive pock marks. Due to tectonic activity (discussed in Chapter 5),

some burrows are elliptical. Skolithos are often the dwelling or feeding burrows

marine worms or arthropods and are common in sedimentary rocks spanning the

whole of the Phanerozoic eon (Fillion & Pickerill, 1990), particularly within the

Cambrian. Some occasional tabular cross-bedding is preserved within the Pipe

Rock Member to the north of Ben Heilam [4710 6245], where Skolithos burrows

are not present. The Pipe Rock Member is approximately 75 metres thick

Interpretation

Similarly to the Cross-bedded Member, the Pipe Rock Member has undergone

extensive re-working prior to deposition. Deposition was in a high energy

environment and it is probable that this was in a shallow marine setting, given the

high level of grain sorting and the close proximity of the member to the Cross-

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bedded Member below which is shallow marine in origin. The presence of

skolithos suggests that aggradation rates may have been slightly lower than that if

the Cross-bedded Member as benthic fauna were able to keep pace with

sedimentation rates. The presence of occasional cross-bedding lacking

bioturbation may be the result of storm events where sudden aggradation

occurred therefore burying any benthic organisms.

2.2 An t-Sron Formation

2.2.1 Alltain Beds

Observations

The Alltain Beds are heterogeneous in composition. Fine siltstones dominate but

medium to coarse sandstones are also present in the form of tabular cross-

bedding, hummocky cross-bedding, isolated lenses and occasional channel fills.

Laying conformably above the Pipe Rock Member, the Alltain sequence begins with

medium quartz sands where herring-bone cross-stratification is visible. Theses

pass up into ferruginous brown siltstones where fine ripples and sub-parallel

laminations are present. Mud drapes are visible on some ripple structures and

represent the settling out of fine sediments. Beds of vuggy grey dolomitic siltstones

are present within some horizons indicating a transition from a siliciclastic to a

carbonate dominated environment. Moving vertically through the beds, siltstones

are occasionally interrupted by medium quartz sands, typically 80cm in thickness.

They contain hummocky cross-stratification and have an erosional base. A

hummocky bedding plane is well preserved at An t-Sron [4415 5815]. Towards the

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top of the sequence, the Alltain Beds are relatively homogenous where hummocky

sands are no longer present and dolomitic siltstones dominate. Occasional lenses

of quartz sands are present but are subordinate. Bioturbation is clearly visible on

some bedding planes where planolites are abundant and parallel to bedding.

Typically they are approximately 5mm in diameter and have an unusual seaweed

appearance. Cruziana can also be observed on some bedding planes, but are

generally not very well preserved. The Alltain Beds are approximately 20 metres

thick.

Interpretation

The Alltain Beds mark a general transgressive environment, where intertidal

herringbone cross-bedding at the base is succeeded by argillaceous siltstones

deposited below wave base. The presence of small ripple structures within some

horizons suggest that there was still a weak tidal influence present. Hummocky

cross-stratified quartz sands are the result of storm events washing clastic

material from the shore face, indicating that the Alltain Beds were generally below

wave base but above storm wave base. There is a distinct change from a

siliciclastic depositional environment within the lower Pipe Rock Member to a

transitional carbonate depositional environment within the Alltain Beds. Given the

vuggy nature of the dolomitic siltstones within the beds, it is likely that the original

protolith was predominantly a carbonate siltstone but has undergone extensive

dolomitisation. The transition to a carbonate dominated environment is most

likely a climatic response to a transition from temperate to tropical latitudes. The

loss of hummocky cross-stratification and the presence of more homogenous deep

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water siltstone facies near the top of the member indicate that transgression

continued throughout its depositional history. The presence of cruziana suggests

that trilobites were in existence during this period, a tentative estimation as to the

age range of the member would be Early Cambrian to Permian.

2.2.2 Salterella

Observations

The Salterella Member is a compositionally mature sandstone which is similar to

the Pipe Rock Member. Quartz grains dominate (95%) with subordinate

plagioclase (<5%), the member is therefore a quartz arenite. Many quartz grains

are inter-locking indicative of substantial cementation, but distinctive well

rounded millet seed grains of 1 to 2mm in diameter are visible within some

outcrops. This suggests that the provenance of the sediments may stem from

terrigenous origin in the form of Aeolian sands. A distinctive feature of this

Member is the presence of small conical Salterella fossils, which is useful in

distinguishing it from the Pipe Rock Member. Within the member, Salterella are

most commonly visible as small conical voids, where the Salterella have weathered

out. They are typically 2 to 4mm long, with distribution being relatively sparse.

Originally identified as a form of cephalopod, later work by Yochelson (1977;

1983) revealed that Salterella was a member of the short lived phylum Agmata

that existed only during the late Early Cambrian (Bonnia-Olenellus zone).

Interpretation

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The Salterella Member was deposited in a high energy intertidal zone similar to

that of the Eriboll Formation, where a regressive period resulted in the deposition

of terrigenous Aeolian sands over the muddy carbonates of the Alltain Member.

Deposition was possibly within a beach or barrier island setting but it is difficult to

prove given the lack of sedimentary structure, i.e., cross-bedding; palaeo-ripples;

etc. The member can be placed accurately within the Bonnia-Olenellus zone of the

late Early Cambrian (Yochelson 1977, 1983) due to the presence of the short lived

Salterella. The member is only 10 to 12 metres in thickness.

2.3 Tor Liath Formation

2.3.1 Kempie Dolostone

Observations

The contact between the Kempie Dolostone and the underlying Salterella is

conformable, where the quartz arenites of the Salterella give way to medium sands

consolidated within a fine grained dark grey matrix. Sand grains near the base of

the Kempie Dolostone are of well-rounded quartz and similar to that of the

underlying Salterella, it is therefore probable that are of similar terrigenous origin.

The fine grained matrix is dark grey in colour and does not react to HCl (10%), it is

therefore predominantly dolomitic. The texture is crystalline and abrasive.

Following vertically up sequence from sandy dolostones at the base, there is a

transition to monotonous dark grey dolostones devoid of any sedimentary

structures or fauna, other than occasional light bands of millimetre scale. The light

bands are sparse and follow no systematic pattern.

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Interpretation

The lack of sedimentary and faunal evidence within this formation is possibly the

result of extensive dolomitisation, where much of the evidence may have been

destroyed. The process of dolomitisation is enigmatic and poorly understood, but

is thought to result from the diagenesis of carbonates post-deposition from

magnesium rich pore waters derived from saline environments such as sabkhas

through the following reaction:

2CaCO3 + Mg2+ ↔ CaMg(CO3) + Ca2+

It is difficult to define a depositional environment for this rock type due to its

altered nature. The dark grey colour suggests the presence of fine micritic muds.

This would indicate deposition in a low energy environment, possibly near the

base of a reef front, below wave base. More importantly, the presence of dolomitic

lithologies marks a distinct change from siliciclastic deposition to carbonate

deposition. This is usually in response to climatic change, where siliciclastic

deposition at temperate latitudes gives way to carbonate dominated deposition at

tropical latitudes nearer the equator. The thickness of this member is

approximately 50 metres.

2.4.2 Heilam Dolostone

Observations

The Heilam Formation bares many similarities to the Kempie Dolostone in that it is

a fine grained dolomitic rock with a crystalline abrasive texture. As with the

11

former, it is also devoid of any structures and fauna due to the extensive

dolomitisation. The only characteristic which distinguishes it from the Heilam

Formation is the lighter grey colour. The basal contact is not visible within the

mapping area, but can be observed just south of the area near Tor Liath [4415

5755]. The transition from dark to light grey is gradational.

Interpretation

Similarly to the Kempie Dolostone, assigning a facies to this rock type is difficult

due to the extensive dolomitisation. The lighter grey colour suggests that there is

distinctly less muds present when compared with the underlying Kempie member.

Deposition was therefore in a shallower environment, marking a regressional

event where carbonates were deposited within a shallow carbonate sea, probably

above wave base, within the reef flat or back lagoon. The thickness of this

formation is unknown as it is substantially tectonised and poorly represented

within the mapping area.

2.6 Stratigraphical and sedimentological evolution of the area.

The stratigraphic sequence within the mapping area started with the deposition of

marginal marine sediments onto cratonic basement rocks. This marks the

transition from a dominantly erosive regime which exposed the metamorphic

basement, to one where marine transgression was dominant. The first

sedimentary unit is the Cross-bedded Member, where the transgression took the

form of intertidal marine sands. The depositional environment was relatively high

energy and deposition rates were moderate. It is difficult to date the Cross-bedded

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Member given the lack of fauna, but given that the conformably overlying Salterella

Member is late Early Cambrian in age (Yochelson, 1977; 1983), it is likely that the

Cross-bedded Member is Early Cambrian. The deposition of the Pipe Rock Member

represents a continuation of an intertidal marine setting. The well sorted sands are

again indicative of a high energy environment, but the lack of cross bedding within

the member is an indication that sedimentation rates were lower than that of the

Cross-bedded Member.

A transgression followed, where the argillaceous dolomitic Alltain Beds were

deposited just below wave-base. This was succeeded by a regression resulting in a

return to siliciclastic deposition within the intertidal Salterella Member where

deposited. The deposition of the micritic Kempie Dolostone marked a distinct

change to a low energy carbonate dominated environment where the depositional

environment was within tropical palaeo-latitudes nearer the equator. A further

regression followed, where the Heilam Dolostones represent shallow carbonate

deposition within a reef flat or reef lagoon.

Chapter 3 Igneous Rocks

Igneous rocks within the mapping area are exclusively within the Lewisian Gneiss.

A dolerite dyke near the summit of Ben Arnaboll has been extensively deformed

and amphibolitised so is therefore discussed in Chapter 4 (Metamorphic geology).

There are other highly modified relict ultramafic bodies within the Lewisian but it

is difficult to constrain their geological origin due to their protracted and

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14

15

S

S

16

convoluted history, so any attempt based on field evidence would be purely

speculative so will not be discussed further.

Also pervading the Lewisian Gneiss are potassic pegmatites, which are relatively

ubiquitous throughout. They bare a crosscutting relationship where they clearly

cross-cut the fabric of the Lewisian Gneiss. Generally, they are coarse grained with

a typical grain size of 20mm, but some k-feldspar crystals are much larger.

Potassium feldspar is the dominant mineral giving the pegmatites a distinctive

pink colour. Quartz and plagioclase are also present, with the overall composition

being that of an alkali granite. Their emplacement is probably linked to the

existence of a larger pluton from which the pegmatites have originated, but there is

no field evidence of this. It is likely that the pluton is at depth and has not been

exposed at the current level. According to Černý et al. (2012), pegmatites usually

form from highly evolved melts that are rich in incompatible elements such as

boron, phosphorus and fluorine are intrinsically linked to the presence of H2O.

They also go on to say that pegmatites rarely fractionate from I-type granites but

are commonly fractionated from S-type and A-type granites, which could be useful

in disclosing the origin of the Lewisian Gneiss.

As discussed, the formation of pegmatite is largely dependent on the presence of

H2O. According to London & Černý (2008), shear zones can act as conduits which

can introduce fluids to evolved magmas hence aiding the formation of pegmatites.

It is possible therefore, that the emplacement of the pegmatites was associated

with a major tectonic event. This is highly likely as the pegmatite bodies

themselves have undergone deformation (Chapter 4.1.2). They are not however,

related to the post-Cambrian thrusting within the region as this was much later.

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Chapter 4 Metamorphic Geology

4.1 Lewisian Gneiss

4.1.1 Observations

The Lewisian Gneiss is the dominant rock type around the Ben Arnaboll area and

lies structurally above the Cambrian imbricate sequence. On initial inspection they

are dull and weathered and commonly covered in lichen. Outcrops are often

smooth and rounded and possess a distinctively hummocky appearance, possibly

the result of glacial activity. Fresh surfaces reveal a coarsely crystalline,

granoblastic rock with bands of mafic minerals within a largely acidic rock type.

Compositionally, felsic bands are of quartz and plagioclase. Grains are generally

subhedral but some grains seem to be stretched in parallel to the foliation of the

rock. Typical grain size is approximately 2 to 3 mm. The plagioclase is typically

subhedral and cleavage is also visible. Grain size is similar to that of the quartz at 2

to 3 mm. Mafic minerals within the felsic bands are minimal and are only 0.5 mm

in diameter. Due to their size they are difficult to identify. They may well be oxides

rather than mafic minerals.

Mafic bands are ubiquitous within most outcrops. Grains of amphibole are

identifiable as largely anhedral grains of up to 3 mm long. Generally, they have a

vitreous lustre and cleavage is visible on some grains, although it is unclear

whether cleavage is intersecting. Elongate crystals are concordant with the

gneissose fabric. This suggests that the gneiss was subjected to non-coaxial strain.

In hand specimen in is difficult to distinguish amphiboles from pyroxenes, but at

18

outcrop level, mafic bands seem to have a dark green hue which is consistent with

amphibole rather than pyroxene. Furthermore, hornblende is a common

amphibole found within many amphibolite facies metamorphic rocks, it is

therefore plausible that the amphibole in this case is Hornblende. Other mafics

include biotite, identifiable by its dark brown colour and perfect basal cleavage. On

weathered surfaces, biotite has a distinctive flaky golden appearance. Plagioclase is

present but is subordinate to amphibole and biotite. Some garnets are present

within some outcrops and are typically 3 to 4 mm in diameter. The gneissose fabric

clearly wraps around the garnets indicating that they are pre or syn-tectonic.

Modal composition within the mafic bands are: hornblende (55%); biotite (35%);

plagioclase (10%); and garnet (0.5%). Based on the modal quantities and

gneissose texture of the rock as a whole, it can be classified as an acidic

hornblende, biotite, garnet gneiss, of amphibolite facies.

There are some examples of ultramafic boudins and enclaves within the gneiss. An

excellent example is observable approximately 400 metres south of Ben Arnaboll

[4549 5856], where an ultramafic boudin of approximately 1 metre in diameter is

enclosed within the gneissose fabric which wraps around the boudin. Well-

developed quartz megacrysts are present as pressure shadows around the more

competent boudin on either side, demonstrating a “top to the south” shear

direction. The boudins are mono-mineralic with blocky, equant crystals of 5 to 10

mm. Cleavage planes are clearly visible on some faces and lustre is vitreous. It is

dark green in colour. This is probably an amphibolitised pyroxenite, the blocky

texture being mimetic of the original pyroxenite protolith.

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Much of the Lewisian Gneiss is pervaded by granitic pegmatites (discussed in

Chapter 3, p.16) that clearly cross-cut the fabric of the Lewisian. The contact

between the two bodies is sometimes diffuse, this could be due to some

metasomatic reactions between the volatiles within the pegmatites and the host

rock. The diffusive contact is also an indication that the host rock was relatively

hot during emplacement. There is no variation in grain size at the margins. The

intrusive nature of the pegmatites and their cross-cutting relationship with the

gneissose fabric illustrate clearly that they occurred at a later stage. However,

there is also field evidence to suggest that the pegmatites were also subjected to

deformation, as in some localised areas pegmatites have been folded. This

deformation was not co-genetic with earlier deformation so signifies a later

deformational event.

There is also evidence of basic intrusion within the Lewisian Gneiss.

Approximately 300 metres northeast of the summit of Ben Arnaboll [4610 5925]

lies the remnants of a dolerite dyke. As with the ultramafic boudin discussed

earlier in this chapter, the dyke has undergone extensive amphibolitisation. The

dyke has a maximum width of 7 metres, but quickly thins out to the southwest. The

significance of this dyke in the wider context of the metamorphic basement is that

its deformational vector, deduced from the stretching of serecitised feldspars is

concordant with the fabric of the Lewisian Gneiss. The implication of this is that

the dyke must post date the formation of the gneissose fabric during (D1) but was

later subjected to deformation during a later event (D2). The pegmatites were not

affected by D2, so they therefore postdate the emplacement of the dyke.

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4.1.2 Interpretation

The quartzo-feldspathic composition of the Lewisian Gneiss is consistent with the

widely accepted view that most Archean gneisses of this type are derived from

plutonic tonalite, trondjhemite, granodiorite (TTG) protoliths associated with

Archean subduction zones (Rollinson & Windley, 1980). Commonly, subordinate

ultramafics are also found in association with TTG’s and are possibly derived from

deep plutonic cumulates such as pyroxenites, peridotites and dunites (Friend &

Kinny, 2001). The amphibolitised pyroxenite boudin for example could be the

product of such cumulates.

The extensive amphibolitisation of the gneiss indicates that it reached amphibolite

facies metamorphism during D1. It is possible that this was a retrogressive

metamorphic event where granulite facies gneisses were retrogressed to

amphibolite facies. Amphibolite facies lithologies typically form at depths of

approximately 20km assuming a typical geothermal gradient of 300C/km.

Following this, a further deformational event occurred. This event (D2) occurred

after the emplacement of the dolerite dyke as the dyke is concordant with the

foliation within the gneiss. Due to the cross-cutting relationship of the pegmatites

and the gneiss, it is clear that the pegmatites postdate these events. Their origin

seems somewhat enigmatic, but their emplacement must have been driven by a

tectono-thermal event where they were the product of highly evolved water

saturated granites (Jahns & Burnham, 1969). Unfortunately there is little evidence

in the field of any other associated igneous melts that could further elucidate their

geological provenance. Further deformation of the Lewisian Gneiss occurred

following emplacement of the pegmatites. This is observable within some outcrops

21

where some minor folding is visible. In some instances, thin pegmatitic veins are

clearly deformed and possess a sigmoidal appearance. This minor deformational

event (D3) will be overprinted onto D1 and D2 fabrics but due to the complexity of

poly-phase deformation it is difficult to see. Interference patterns such as dome

and basins; or crescent and mushroom (Park, 1989) were not observable in the

field at outcrop level. The fact that some pegmatites seem unaffected by

deformation may suggest that emplacement was contemporaneous with D3 rather

than predating it.

4.2 Arnaboll Mylonite

4.2.1 Observations

The Lewisian Gneiss is separated from younger Cambrian sediments by a major

thrust zone where Pipe Rock in the footwall is overlain Lewisian basement. This is

clearly a tectonic relationship as older Archean gneiss is juxtaposed upon younger

Phanerozoic sediments. The two units are separated by a band of ultramylonite

which varies in thickness between 10cm to >100cm. The mylonite is well exposed

just north of Ben Arnaboll summit [4615 5965] and from here it is laterally

continuous in a southerly direction for over 1km. By Sibson’s (1977) definition, the

ultramylonite (>90% matrix) is almost devoid of clasts and has undergone

extensive grain size reduction through crystal plastic deformation. The mylonites

have a distinctive green colour indicating that they are sub-greenschist grade.

Minerals are segregated into millimetre scale laminated bands, where darker

bands represent mafics and lighter bands represent the more felsic minerals such

as quartz and feldspars. Due to the microscopic grain size, it is difficult to ascertain

22

mineralogy, other than it contains mafic and felsic minerals. The mylonites readily

cleave along lamination planes giving them a flaggy appearance.

Given that the ultramylonites are green in colour, it is highly likely that they are

extensively chloritised. This suggests that that were subjected to greenschist facies

metamorphism. Asymmetrical quartz grains (Plate 2), of up to 30cm are useful

shear indicators and prove useful in determining shear sense. Similar mylonites

can be found covering an extensive area to the southeast of the mapping area

around Glac an Tioraidh [460 575].

4.2.2 Interpretation

Due to the extensive grain size reduction that the Arnaboll Mylonite has

undergone, and the highly ductile, finely laminated fabric that pervades

throughout the rock, it is highly likely that the overthrusting Lewisian has travelled

from substantial depths and over considerable distances. Most of the strain was

accommodated within a thin layer where frictional heating of framework silicates

formed an ultramylonite layer. Movement along the plane would have been sudden

and spontaneous which is evident from presence of pseudotachylite. However,

there is some deformation within the Pipe Rock, which is evident from the

deformation of skolithos burrows within the member which prove to be useful

shear indicators (Chapter 5.1, p.26). Some quartz grains are preserved within the

mylonite and have undergone diffusive mass transfer during shearing, giving them

a stretched appearance. Along with mineral lineations they are useful in

determining shear sense along the plane of movement.

23

4.3 Oystershell Mylonite

4.3.1 Observations

The Oystershell Mylonites (Plate 3) are structurally above the Arnaboll Mylonites

and outcrop on the southern side of Loch a Choin-bhoirinn [456 571]. They are

easily distinguishable by their crenulated appearance and S-C fabric, which is

useful in establishing shear sense within the unit. They bare a similar colour to the

Arnaboll Mylonites but are richer in phylosilicates where muscovite is clearly

visible in hand specimen. They are fine to medium grained and are therefore

coarser than the underlying Arnaboll Mylonites. Quartz within the mylonites are

concordant with the S fabric within the rock and often have a lunate appearance.

They often look like oyster shells, hence the name. In some instances, they contain

lenses of potassium feldspar which is similar to that of the pegmatites within the

Lewisian. According to White (1982), phyllonites are typically the product of lower

strain rates than laminated ultramylonites such as the Arnaboll Mylonites.

4.3.2 Interpretation

It is clear that the Oystershell Mylonites lie structurally above the Arnaboll

Mylonites, but are probably related to the same orogenic event. In their footwall

they contain mylonitised Lewisian basement. Hangingwall lithology in not exposed

so is difficult to interpret. It is possible that they are the product of Moine

metapsammites which outcrop to the east of the mapping area as they are rich in

24

mica, although the presence of potassium feldspar suggests that they are likely to

be derived from the Lewisian Gneiss.

4.4 Quartz Mylonite

Quartz Mylonites are interleaved within the Oystershell Mylonites and are

confined to two localised areas near Glac an Tioraidh [4615 5750]. They are

arenitic in composition, similar to that of the Cross-Bedded Member. Due to their

association with the Oystershell Mylonites, it is likely that they are from the basal

unconformity where the Oystershell Mylonites represent the Lewisian and the

Quartz Mylonite represents the Cross-bedded Member. This therefore is an

indication that the thrust zone cross-cuts the basal unconformity.

4.5 Metamorphic history of the area

The history of the Lewisian Gneiss has been covered in Chapter 4.1.2. The

mylonites within the area have undergone dynamic metamorphism but in an

historical context, they are best discussed within the structural history (see

Chapter 5.6) of the area rather than the metamorphic history.

Chapter 5 Structural Geology

5.1 Faults/shear zones

5.1.1 Observations

25

26

The Arnaboll Thrust- Due to the complex tectonic history of the mapping area,

faulting is common and pervasive throughout. Major thrust zones include the

Arnaboll Thrust, which carries allocthonous Archean gneiss in the hangingwall

onto the younger Pipe Rock Member of the Cambrian succession. This relationship

is best exposed at the northern end of Ben Arnaboll [461 596]. Virtually all of the

strain is incorporated within a relatively thin band of ultramylonite (see Chapter

4.2), where the trajectory of the fault can be inferred from the asymmetrical

deformation of quartz ribbons within the ultramylonite. The underlying Pipe Rock

in the footwall also exhibits substantial strain where formerly sub-vertical

skolithos burrows (normal to bedding) have been deformed to angles of

approximately 450 to bedding proximal to the ultramylonite zone (Fig.3), this

implies a shear ratio of 1. Both the quartz ribbons and the skolithos burrows

indicate a west-northwest trajectory for the Arnaboll Thrust. Approximately 100

metres to the east, the Arnaboll Thrust is breached by three younger thrust faults

which cut across the thrust and dip to the east. These are clearly later than the

Arnaboll thrust and may be related to imbricate thrusting to the north around Ben

Heilam.

Following the thrust contact south to the un-named lochan [4610 5885], the

Arnaboll Thrust forms an anticlinal structure where Lewisian basement is cored by

Pipe Rock, this implies that the Arnaboll Thrust is folded. This is consistent with a

model proposed by Butler (2004) of a foreland propagating duplex where an early

roof thrust, in this case the Arnaboll Thrust is folded by a series of sub- surface

imbricates which join onto the main roof thrust (Fig.4a). To the west, the Arnaboll

27

Thrust cuts up section into the Cross-bedded and the Pipe Rock Members of

Cambrian age.

The Tioraidh Thrust- Lying in the hangingwall of the Arnaboll Thrust, the

Tioraidh thrust generally consists of greenschist mylonites interleaved with

sheared Lewisian Gneiss, which gently dip to the east-southeast. The thrust zone is

much wider than the Arnaboll Thrust and mylonites are less developed. At the

base, laminated greenschist mylonites progressively give way to phyllonites

(Chapter 4.3) higher up the sequence. The presence of k-feldspar within the

phyllonites indicate that they are derived from the pegmatitic Lewisian Gneiss. L-S

fabric within the mylonite generally dips 150 to the east-southeast and S-C fabric

within the Oystershell Mylonites (Plate 3) indicate that movement was to the west-

northwest.

As the mylonites within the Tioraidh Thrust are less well developed than the

Arnaboll Thrust mylonites, it is likely that they postdate the latter and were

emplaced at higher crustal levels. The lack of folding within the Tioraidh thrust

sheet also suggests that they were emplaced after the Arnaboll sheet. Subordinate

Quartz Mylonites interleaved with the Oystershell Mylonite are probably derived

from the basal Cambrian unconformity, which is compatible with the hypothesis

that the Tioraidh Thrust mylonites are predominantly from a Lewisian protolith,

rather than Moine metapelites.

To the south of the mapping area near Kempie Bay, the thrust crosscuts a series of

ductile folds within the underlying Arnaboll Sheet. The folds are tight and isoclinal

with fold axes that dip at a shallow angle to the east-southeast, similar to that of

the thrust sheet, suggesting that they are co-tectonic with the Tioraidh Thrust. This

28

is further evidence that the Tioraidh Thrust was emplaced later than the

underlying Arnaboll Sheet.

Imbricate Thrusts- The imbricate thrusts lie in the footwall of the Arnaboll Thrust

sheet and are composed of repeating successions of Cambrian formations that crop

out along the western shores of Loch Eriboll, north of Heilam [4590 6160]. The

imbricates propagate along a sole thrust which is not exposed within the mapping

area. Steeply dipping to the east-southeast, the thrusts often exploit weaker units

such as the Alltain Beds or the Heilam Dolostone. Individual slices are

approximately 20 metres thick and can be traced along strike for over 80 metres in

some parts. On the western shore, imbricates are within the Alltain and Salterella

Members but progressively change to Salterella and Heilam Members further east.

Where thrusting occurs within the Alltain Beds, a prominent shear fabric is visible

where the fissile nature of the Alltain Beds is exploited. They are often flanked by

competent sandstone beds either side (Plate 4). To the south, the imbricates climb

up section into the dolostones and die out onto a lateral thrust ramp at Ard Neakie

[4500 5990] which marks the southern limits of the imbricates. At this location,

the dolostones are southwardly dipping and have undergone sinstral shearing

within a wide shear zone (Fig.5) along the lateral ramp. The lateral ramp at Ard

Neakie adjoins onto the Arnaboll Thrust at Druim na Teanga. Walking across strike

in a south-easterly direction from the Lighthouse [4581 6178], there is a transition

from the steeply dipping imbricates in the footwall to gently dipping Pipe Rock

with occasional horizons of sheared Alltain Beds in the hangingwall.

The transition between the imbricates and the Pipe Rock is marked by a 1 metre

thick mylonitic zone, which is well exposed 400 metres to the east of the

29

Lighthouse [4640 6190]. Most of the shearing is accommodated within the

overlying Pipe Rock in the hangingwall and the presence of thick mylonite

indicates that it has travelled from considerable depth. The Pipe Rock dips gently

to the east-southeast with remarkable consistency, so it is unlikely that the

hangingwall is folded by an underlying duplex. It is therefore plausible that this

marks an overstepping geometry, where the imbricates in the footwall have been

truncated by a later, low angle thrust that steps back into the hinterland (Fig.4b).

The low angle mylonitic band that separates the hangingwall from the footwall is

consistent with a far travelled low angle thrust fault that oversteps the underlying

strata, as alluded to by Butler (2004) in his paper on the nature of roof thrusts.

Thrusting within the Pipe Rock dominated hangingwall is low angle, where fault

propagation folds have developed into hangingwall anticlines above sub-

horizontal thrust ramps. Most anticlines are not preserved but an excellent

example can be observed approximately 200 metres east of Loch a’ Choire [4686

6123]. It clearly illustrates how movement along the thrust ramp was to the west

(Plate 5). From the Heilam Cross-section (Fig.2), a tentative estimation for the

30

amount of crustal shortening within the Heilam imbricates alone can be put at

11km. This is an estimation only, as the exact number of imbricates could not be

measured in the field due to lack of exposure. The overall crustal shortening is

lightly to be much higher when movements along the Tioraidh and Arnaboll

Thrusts are taken into account.

Other Faults- The imbricates on the western Heilam Peninsula are crosscut by a

series of much later high angle faults which are downthrown to the north. These

are probably the result of thermal cooling and subsidence of the orogen in its post

orogenic state. Similar faults crosscut the Lewisian Gneiss and mylonites south of

Ben Arnaboll, where geological units are juxtaposed against one another. On the

northern end of the Heilam Peninsula, spectacular 1.5 metre wide brittle transform

faults can be observed. In the centre of the fault, lies a 15cm thick band of

cataclasite, which is flanked on either side by fault breccia. Movement along the

fault is sinstral and to the southwest and the fault obliquely cuts across open folds.

The fault remains linear across the folds so is therefore younger.

5.1.2 Interpretation

Owing to the complex nature of thrusting within the mapping area, it has

undergone a somewhat enigmatic history. This said, there are some key events

which can be placed into a temporal framework to ascertain a sequence of events:

1) Firstly, the emplacement of the Arnaboll Thrust Sheet occurred where a far

travelled thrust placed deep crustal Lewisian Gneiss onto Cambrian Pipe

Rock.

31

2) Further crustal shortening resulted in a foreland propagating thrust

sequence within the footwall of the Arnaboll Sheet, which resulted in the

emplacement of a series of imbricate thrusts within the Cambrian

sediments. This event transformed the Arnaboll Thrust into a roof thrust,

where imbricates generally cut up section and joined onto and folded the

Arnaboll Thrust.

3) The Tioraidh Thrust then emplaced mylonitic Lewisian Gneiss and

Cambrian arenites onto the ductile Arnaboll Thrust Sheet, resulting in the

formation of tight to isoclinal folds within the Arnaboll Sheet around

Bealach Mahri and the formation of the Kempie Anticline.

4) The overstepping thrust on the Heilam Peninsula must post date the

imbricates due to its crosscutting relationship with the imbricates in the

footwall. As the breaching of the Arnaboll Thrust penetrates the same

hangingwall imbricates around Ben Heilam, it is probable that the

breaching is contemporaneous with these low lying thrusts.

All structural data including lineations; S-C fabrics; and stretched quartz are

consistent with vergence from the east-southeast.

5.2 Folds

Folding within the Lewisian Gneiss will not be discussed in this section as they

have been discussed in Chapter 4. The discussion will predominantly centre round

differences between folding in the northern Heilam region and further south

around Bealach Mhairi, and the implications these differences have with regard to

the structural styles of deformation within the regions.

32

5.2.1 Observations

On the Heilam Peninsula north of the A838, folding is strongly associated with

imbricate thrusting. On the western coast, any folding associated with the steeply

dipping Heilam Imbricates have been eroded, but originally they would have

culminated in a series of hangingwall anticlines. A large scale anticline is

associated with the lateral ramp at Ard Neakie (Fig.5). This marks the

southernmost extent of the imbricates where bedding dips to the south and shear

strain is accommodated with sinstral movement along the lateral ramp.

As discussed in Chapter 4.1.1 some excellent examples of folding are preserved

within the Pipe Rock Member around Ben Heilam, where fault propagation folds

have developed into hangingwall anticlines with axial planes that dip to the east

(Plate 5). Most of the crustal shortening is accommodated along brittle shear zones

along flat lying footwall ramps. The hangingwall anticlines take the form of tight to

isoclinal TLS (Tangential Longitudinal Strain) folds where strain is predominantly

coaxial, causing extension on the outer arc of the fold and compression within the

inner arcs (Park, 1989). The extension on the outer arc is quantifiable due to the

presence of skolithos burrow entrance holes on bedding planes. On the

hangingwall anticline south of Ben Heilam for example, originally round entrance

holes are stretched into ellipses, where the long axis of the ellipse is normal to the

fold axis, indicating that the outer arc was subjected to extension. Fold axial planes

where measurable, dip to the east-southeast.

On the northern end of the Heilam Peninsula [4750 6250], there are a series of

open folds which plunge gently to the southwest. They are exclusively within the

Pipe Rock Member and are in the footwall of the overlying imbricated Pipe Rocks

33

on Ben Heilam. Their formation is somewhat difficult to explain but are probably

linked to the lateral termination of thrusts on Ben Heilam where strain rates are

much less.

To the south, the Arnaboll Thrust Sheet is folded by the underlying duplex. The

folding is difficult to quantify within the thrust sheet as there is no bedding within

the Lewisian, so folding can only be inferred from the presence of the Pipe Rock

around the un-named lochan which is flanked by Lewisian Gneiss either side, and

the folded Cross-bedded arenites which lie in the hangingwall of the Arnaboll

Thrust Sheet at Druim na Teanga [4534 5930].

Moving to the southern end of the mapping area around Bealach Mhairi [4535

5760], folding is accommodated within the Arnaboll Thrust Sheet. Deformation

here takes the form of tight east-southeast dipping folds within the Lewisian and

the overlying Cross-bedded Arenites. The contact between the Lewisian and Cross-

bedded arenites is unconformable rather than tectonic as there is no evidence of

shearing. It is evident from cross-bedding that some beds are overturned.

Deformation here is ductile and moving east towards the Tioraidh Thrust, folds

become isoclinal and are interleaved with mylonites. This implies that at this

location there is a strong relationship between the overlying Tioraidh Thrust and

the ductile deformation within the Arnaboll Thrust Sheet. Based on this evidence it

is likely that the ductile folding in the footwall and the formation of the mylonites

in the hangingwall were co-genetic.

Towards the foreland to the west, folding becomes increasingly more open, where

a large anticline and syncline (Kempie Syncline) are present. This is an indication

that there is markedly less deformation here than there is further east towards the

Tioraidh Thrust, as almost all strain is accommodated within the higher sections

34

towards the thrust zone. The Kempie Syncline is cored by a thrust fault which dips

to the east-southeast. Most of the thrusting is accommodated within the Alltain

Beds and is clearly visible at Kempie [4460 5800]. This is likely to be a splay from

the underlying Arnaboll Thrust.

5.2.2 Interpretation

Most of the folding to the north of Ben Arnaboll is strongly linked to imbricate

thrusting, where folds take the form of hangingwall anticlines, these would have

initially developed as asymmetrical fault propagation folds (Fig.6).

Fold axes reveal that vergence was from the east-southeast. Thrust planes within

the Pipe Rock are narrow and brittle indication emplacement at relatively shallow

depths. In contrast, folding to the south of the mapping area around Bealach Mhairi

is of a much more ductile nature, where tight to isoclinal folding is strongly

associated with mylonites within the overlying Tioraidh Thrust Sheet. Proximal to

the thrust, tight isoclinal folds are subjected to intense non-coaxial shearing,

becoming mylonitic in texture.

It is evident that there are distinctive differences in structural style between the

Heilam area to the north and Bealach Mhairi to the south. The Heilam area is

dominated by brittle imbricate faulting, implying that emplacement was at

relatively shallow depths. In contrast, the folding to the south around Bealach

Mhairi is highly ductile which suggests emplacement at depth.

5.3 Cleavage

Most of the lithologies within the mapping area are massive and do not contain

platy minerals such as micas. Therefore most of the mapping area is devoid of

35

cleavage. The quartz arenites of the Eriboll Formation for example are completely

devoid of platy minerals so do not develop cleavage. Some beds do possess

foliation in the form of protomylonites and cataclasites but this if formed due to

the shear stresses associated with thrusting. This is not cleavage in the strictest

sense and is best described as foliation.

5.4 Lineations on faults and stretching lineations

Within sheared Alltain Beds, slikenlines are clearly visible on fault planes. These

are best preserved on the underside of faults where they are sheltered from the

elements. Within the mylonites around Ben Heilam and Bealach Mhairi, stretching

lineations are visible on some foliated surfaces. These are predominantly

preserves within stretched quartz grains which form a weak L-S fabric within the

mylonites. Lineation data from both the Alltain Beds and mylonites are consistent

and bare little variation. Out of a total of 20 lineation measurements taken in the

field, all plunged to the east-southeast (Appendix 2). The data was plotted on a

stereonet and a mean vector of 29o towards 109E was obtained.

5.5 Stylolites and associated tension gashes.

On the northern tip of the Heilam Peninsula [4740 6252], there are a series of

quartz filled en-echelon tension gashes. Stylolites associated with the tension

gashes are less well developed. The orientation of the tension gashes indicate that

the direction of greatest principle stress (σ1) is orientated southwest to northeast.

This is favourable with the late brittle faulting that occurs in the area (Chapter

5.1.1, p.30). The origin of the stress is unknown and is only evident along the

northern extent of the peninsula.

36

5.6 Structural history of the area

37

Crustal shorting was initiated within the area by the thrusting of deep crustal

Lewisian Gneiss onto Pipe Rock of Cambrian age along the Arnaboll Thrust.

Vergence was from the east-southeast and emplacement was post-Cambrian, i.e.

later than the Pipe Rock Member. The presence of ultramylonite along the thrust

plane indicates that exhumation was from depths of at 15km. A foreland

propagating sequence of imbricate thrusting followed this event and formed a

duplex where the Arnaboll Thrust acted as a roof thrust. Some thrusting stepped

back into the orogen, where a low angle thrust emplaced shallow dipping Pipe

Rock (Ben Heilam) onto the steeply dipping imbricates. The breaching of the

Arnaboll Thrust was contemporaneous with this event. Emplacement of the

Oystershell and Quartz mylonites of the Tioraidh Thrust then followed. The

anticline around Kempie Bay and the ductile folding on Bealach Mhairi was

contemporaneous with the emplacement of the Tioraidh Thrust. All tectonic events

discussed were due to compression from the east-southeast and the amount of

total crustal shortening is likely to be in the order of at least 10’s of kilometres.

Thermal subsidence post orogen resulted in a series of brittle faults, these are best

observed on the western side of Ben Heilam. A more recent tectonic event resulted

in the brittle transform faulting observed on the northern tip of the Heilam

Peninsula. The faulting here was the result of southwest to northeast compression.

Chapter 6 Geological history of the area

This is a brief summary of the geological history of the area, and where possible,

some correlation will be made with current understanding of the geology from

previous research and/or current understanding. Each sub-section is placed in

38

chronological order to give the reader an understanding of the temporal

distribution of events.

6.1 Lewisian Gneiss

The emplacement of the Lewisian Gneiss was during the Archean and their

composition is similar to other TTG gneisses from around the globe. Friend

and Kinney (2001) assign them a protolith age of 2840-2800Ma. A

geochemical study by Goodenough et al. (2010), deduced that their most

likely origin was of parental melts from a mantle wedge setting, similar to

calc-alkaline rocks seen today.

The protolith was then buried to depths of 20km or more and resulted in

the development of gneissose texture and metamorphism up to at least

amphibolite facies and possibly granulite facies. In Kinney and Friend’s

reappraisal on terrane based nomenclature (2005), they assigned the

Lewisian in this area to the Rhiconich Terrane, which covers the area north

of the Laxford Shear Zone from Laxford Bridge.

The emplacement of the dolerite dyke on Ben Arnaboll followed. It is

believed that this is related to the emplacement of the Scourie Dyke

Complex. Emplacement pre-dates 1855Ma (Friend & Kinney 2001).

A deformational event followed (D2). This resulted in the deformation and

amphibolitisation of the dyke. D2 is poorly constrained but there is evidence

that Badcallian event continued after the emplacement of the Scourie Dykes

(Trewin, 2002). However, the event was more commonly associated with

the Assynt Terrene.

39

Emplacement of the pegmatites followed. It is likely that their emplacement

was related to the introduction of hydrous fluids through the Laxford Shear

Zone, Friend & Kinney (2001) identified this as the Laxfordian. The

commonly accepted hypothesis is that the Laxford Shear Zone separates the

Rhiconich Terrane of the north from the genetically distinct Assynt Terrane

to the south of the shear zone. This is based on the fact that the two

terranes have undergone distinctly different metamorphic histories (Coney

et al., 1980; Goodenough et al., 2010). The Laxfordian event has been dated

at c. 1705Ma, based on hornblende Ar/Ar dating (Dallmeyer et al., 2001)

and affects both the Rhiconich and Assynt Terranes.

According to Friend & Kinney (2001), the Rhiconich Terrane was subjected

to a final deformational event at c.1670Ma. This event supposedly

overprints previous deformation but no evidence for this event was

observed in the field.

6.2 Cambrian sedimentology

The Cambrian marked the start of a transgressional period where initially,

the deposition of quartz arenites of the Eriboll Formation dominated. These

were within tidal dominated environments and involved the reworking of

mature sediments. The Eriboll Formation is Early Cambrian in origin (Park

et al., 2002) and lay unconformably on the Lewisian Gneiss.

Following the Eriboll Formation, there was a transgression marked by

deposition of the Alltain Beds. Generally these were deposited in a low

energy environment below wave base, with a marginal carbonate influence.

40

The Alltain Beds are more commonly known by their official name which is

the Fucoid Beds. Following this there was a return to intertidal siliciclastic

deposits of Salterella, formally known as the Salterella Grit member (Park et

al., 2002). The Salterella Member is accurately dated within the late Early

Cambrian (Yochelson, 1977).

A transition to tropical latitudes followed, with carbonate shelf deposition

being dominant within the Tor Liath Formation. The Kempie Member was

deposited in a low energy environment, possibly near the base of a reef

front. The overlying Heilam Member was deposited within a shallower

carbonate environment, marking another regression. The above mentioned

members are known as separate formations in modern literature and go by

the name of Ghrudaidh Formation and Eilean Dubh Formation respectively

(Goodenough & Krabbendam, 2011).

6.3 Caledonian Thrusting

Following the Cambrian transgression, there was an intense period of crustal

shortening which resulted in intensive thrusting during the Caledonian Orogeny.

Compression was exclusively from the east-southeast and generally followed a

foreland propagating sequence with some instances of overstepping:

The emplacement of the Arnaboll Thrust Sheet along the Arnaboll Thrust

occurred early on in the orogenic evolution, where a low angle thrust

emplaced Lewisian Gneiss onto Cambrian Pipe Rock.

A foreland propagating thrust sequence followed, resulting in the

imbrication of Cambrian strata towards the west. This resulted in the

41

formation of a duplex where the Arnaboll Thrust acted as a roof thrust

(Plate 6).

The Tioraidh Thrust then emplaced mylonitised Lewisian Gneiss and Lower

Cambrian quartzites onto the Arnaboll Thrust Sheet. This event resulted in

the ductile deformation and the formation of tight isoclinal folds within the

Arnaboll Sheet at Bealach Mhairi. The Tioraidh Thrust sheet, commonly

known as the Lochan Riabhach Thrust Sheet (Holdsworth et al., 2007) was

not folded.

Further thrusting continued in the hinterland where a low angle thrust

(observed to the west of Ben Heilam) truncated imbricates in the foreland

and emplaced Pipe Rock onto them. The breaching of the Arnaboll Thrust is

contemporaneous with this event.

Post orogenic extensional faults formed as a result of thermal sag.

The extensive thrusting within the mapping area is related to the closure of the

Iapetus Ocean and the bringing together of Laurentia, Avalonia and Baltica during

the Caledonian Orogeny in the Silurian. The area is the northernmost extension of

the Moine Thrust Zone, which places Moine metapsammites onto the Archean

Lewisian foreland (Goodenough & Krabbendam., 2011). The exact position of the

Moine Thrust has been debated for many years. Peach & Horne in their memoirs

(Peach et al., 1907) for example placed the Moine Thrust at the base of what is now

known as the Lochan Riabhach Thrust Sheet, whereas modern interpretation

places the Moine Thrust further towards the hinterland, where Moine Schists lie in

the hangingwall (Holdsworth et al., 2006). It is now believed that thrusting within

42

the Moine Thrust is both foreland propagating and overstepping (Butler, 2004;

2010). Within the mapping area there is evidence of both types.

6.4 Post orogenic events

Following the Caladonian Orogeny, the area was subjected to some brittle faulting

resulting in the formation of cataclasites and fault breccia. The faulting was

localised within the northern Heilam area and their origin and timing are

unknown. Sculpting from glacial activity is also observable from the presence of

glacial striations (trending north-northwest) on Pipe Rock within the Heilam area.

These are probably related to the last glacial period which ended during the end of

the Pleistocene.

6.5 Summary

In summary, the MTZ has undergone a complex history of thrusting during the

Caledonian orogeny. Most authors (Freeman et al., 1998; Dallmeyer et al., 2001;

Kinney et al., 2003) have concluded that deformation occurred during the Silurian

with the amount of crustal shortening estimated at well over 100km. Thrusting

within the MTZ generally follows a foreland propagating sequence with some later

overstepping thrusts at higher levels. Kinematic data from lineations, s-c fabric,

stretched quartz grains etc. illustrate that vergence was from an east/south-

easterly direction and show remarkable consistency. Despite all the research, the

relative timings of thrusting events within the MTZ and specifically within the Loch

43

Eriboll area are still poorly constrained. This is especially true of the Lochan

Riabhach Thrust.

44

45

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48

Appendix

A.1 Restored foresets

A.2 Lineations

Original bedding:

173/22

Original foresets:

061/19

Palaeocurrent (blue

arrow): 106oE

49

9 Acknowledgements

Firstly I would like to thank my supervisor Steve Hirons for his support. It was he

who during his second year field class first inspired me to map the Moine Thrust

Zone. I would also like to thank him for his support and expertise in the field so

thanks Steve. I must also thank Rick Allmendinger for use of his “Stereonet 9” open

source software, which proved really useful in the production of stereonets. A

special mention must go to Cara and Liam who have sacrificed many an Easter

holiday while I have been conducting my fieldwork, they have always been really

supportive so thanks guys. I would like to reserve my biggest thanks though to my

wife Sian. Without her influence I would not be in the position I find myself in now,

on the verge of completing my degree. She has been the true inspiration behind all

my endeavours. I now hope that I can be as supportive to her as she been to me as

she continues on the road to becoming a fully qualified Occupational Therapist,

thanks Sian.