Using mineral chemistry to constrain the P-T conditions of
mantle xenoliths from the Kaapvaal craton, South Africa
Viesturs Smildzins
Master’s thesis Oulu Mining School
University of Oulu
2016
Abstract
Kimberlites are igneous rocks that originate by small degrees of melting of the mantle.
Notably, kimberlites carry large variety of crustal and mantle xenoliths. Geochemical data on
xenoliths can provide insights into the processes occurring in the subcontinental lithosphere
(SCLM) and deeper.
The Kaapvaal craton in South Africa hosts one of the best-studied kimberlite
populations on Earth. In this thesis, a total of 24 thin sections of peridotite xenoliths from Group
I Letlhakane, Letseng, Premier and Frank Smith kimberlites were investigated to constrain the
pressure, temperature and depth of these mantle xenoliths. To do so, olivine, orthopyroxene,
clinopyroxene, garnet and spinel were analyzed for their major element chemistry using
electron microprobe analysis (EPMA). P-T calculations were performed using the PTEXL3
spreadsheet program, which contains different geothermobarometers. Depth constraints were
fitted to the characteristic Kaapvaal craton geotherm.
According to geochemical results and rough modal mineral estimations, the majority of
the mantle xenoliths were identified as depleted harzburgites or lherzolites. Mineral major
element compositions show trends of depletion, which correlate with the corresponding mantle
xenolith sampling depth. Olivine and orthopyroxene have high average Mg# values of 92.1 and
93.0, respectively, at shallower depth ~70-160 Km. Below ~160 km, Mg# starts to drop rapidly
and transition towards a more typical asthenospheric composition. The majority of garnet
compositions fall into the G9 classification field. Titanium shows a distinct partition trend that
correlates with depletion. Garnets have well developed alteration reaction rims, especially at
shallower depths.
Geothermobarometric calculations for four-phase peridotites are comparable with the
results from other studies. However, the temperature estimates obtained by T(BKN90) are slightly
overestimated and, in contrast, the pressure estimates from P(BBG08) are slightly underestimated.
Other assemblages have considerable calculated pressure and temperature conditions and were
best fitted for the regional conductive geotherm. The mantle xenoliths show pressures ranging
from 22 to 56 kb and temperatures from 753 to 1344 0C that characterize an extensive sampling
depth range from 70 to 190 km. Three of the samples extend into the diamond stability field.
The obtained P-T data for mantle xenoliths cluster along a 44.0±2.0 mWm-2 conductive
Kaapvaal craton continental geotherm, being slightly higher than that of the average thermal
state estimate for the craton.
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Contents
1. Introduction ..................................................................................................................... 4
2. Kimberlites ...................................................................................................................... 5
2.1. Exploration history and definition of kimberlite ...................................................... 5
2.2. Geochemistry of kimberlites and orangites .............................................................. 8
2.3. Kimberlite distribution and ages .............................................................................. 9
2.4. Kimberlite pipe formation and facies ....................................................................... 12
3. Mineralogy and classification of ultramafic mantle xenoliths .................................... 17
3.1. Mineralogy ............................................................................................................... 17
3.2. Modal and textural classification ............................................................................. 19
4. Continental lithospheric geotherms .............................................................................. 21
5. Geological setting of the Kaapvaal craton .................................................................... 24
5.1. General features ........................................................................................................ 24
5.2. Crustal components of the Kaapvaal craton ............................................................. 27
6. Kimberlites sampled for this study ............................................................................... 31
6.1. Premier kimberlite .................................................................................................... 31
6.2. Letseng kimberlite .................................................................................................... 32
6.3. Letlhakane kimberlite ............................................................................................... 34
6.4. Frank Smith kimberlite ............................................................................................. 35
7. Sampling and methods ................................................................................................... 36
7.1. Samples and thin section studies .............................................................................. 36
7.2. Electron microprobe analysis ................................................................................... 36
7.3. Geothermobarometric calculations........................................................................... 37
7.4. Geothermometers ..................................................................................................... 38
7.5. Geobarometers.......................................................................................................... 41
8. Results .............................................................................................................................. 44
8.1. Petrography .............................................................................................................. 44
8.1.1. Letlhakane ............................................................................................................ 47
8.1.2. Letseng ................................................................................................................. 50
8.1.3. Premier ................................................................................................................ 51
8.1.4. Frank Smith .......................................................................................................... 52
8.2. Mineral major element chemistry............................................................................. 53
8.2.1. Olivine .................................................................................................................. 54
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8.2.2. Garnet .................................................................................................................. 56
8.2.3. Orthopyroxene ..................................................................................................... 61
8.2.4. Clinopyroxene ...................................................................................................... 61
8.2.5. Spinel .................................................................................................................... 63
8.3. Geothermobarometry ................................................................................................ 64
8.3.1. Overview .............................................................................................................. 64
8.3.2. Inter-mineral equilibrium and P-T estimates....................................................... 65
8.3.3. P-T conditions and depth of origin of studied mantle xenoliths .......................... 69
8.3.4. Composition of the subcontinental lithospheric mantle ....................................... 73
9. Conclusions .................................................................................................................... 76
10. Acknowledgments .......................................................................................................... 77
11. References....................................................................................................................... 78
Appendices
I. Chemical composition of silicate and oxide minerals
II. Pressure and temperature calculation results
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1. Introduction
A major topic in geology that is highly debated concerns subduction processes, plate
tectonics and the related change in the continental growth mechanism after the Archean eon.
So far, geochemical and geodynamic evidence suggests that subduction processes may have
operated during the Archean to some extent. Plate subduction was more episodic and controlled
by different factors compared to the Proterozoic and Phanerozoic (van Hunen & Moyen, 2012).
However, it is hard to establish a precise and uniform model for Archean subduction due to the
lack in concrete evidence, since most of the data from subcontinental lithospheric mantle
(SCLM) rocks have been gathered using xenoliths and xenocrysts from relatively small
kimberlite intrusions. In this respect, the Kaapvaal craton in South Africa is the most
extensively studied Archean terrane (Griffin et al., 2003).
One way to explore the relationship between ancient subduction and the composition of
the subcontinental lithospheric mantle is through the use of multiple sulfur isotopes (Farquhar
et al., 2000; Johnston, 2011; Kamber & Whitehouse, 2007). This is based on the fact that mass-
independent fractionation (MIF) of sulfur isotopes took place only in the Archean time, as
recorded by the isotope composition of old sedimentary rocks. Consequently, if MIF-S is
identified in mantle rocks, it would indicate subduction of sedimentary sulfur into the mantle.
Sulfide inclusions with different compositions have been recovered from peridotite and eclogite
xenoliths in the past. The main question is whether the MIF-sulfur was transferred to the mantle
by subduction or not and how do the compositions vary in the lithospheric mantle and deeper
mantle. To map the sulfur isotope compositions from various depths in the lithosphere and
underlying mantle, a more systematic approach is needed.
The main aim of this study is to constrain the pressure and temperature conditions and
depth of origin of mantle xenoliths, using the major element compositions of the main minerals
of the xenoliths. The sample material includes peridotitic xenoliths from four Group I
kimberlites in the Kaapvaal craton, named Letlhakane, Letseng, Premier, and Frank Smith. The
obtained data may be further used to tackle the problem of the potential existence of MIF-sulfur
in the sublithospheric continental mantle.
The main tasks of the study are to determine:
Major element contents in the main minerals of mantle xenolith
Calculate P-T conditions for the mantle xenoliths using appropriate geothermometers
and geobarometers
Constrain the depth of origin of the mantle xenolith samples
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2. Kimberlites
2.1. Exploration history and definition of kimberlite
Diamonds may be present in kimberlites as an accessory mineral phase and therefore
discoveries of kimberlite pipes have been closely related to the diamond exploration history.
Caravans from Middle East were trading precious stones as early as 400 B.C. At that time, these
valuable gems with intense light-reflecting properties were hand-picked from stream sediment
placer deposits, away from their primary source rocks (Levinson, 1997). India dominated as the
main diamond producer until the 1700s when new diamond fields were discovered in Brazil
along river banks in the Minas Gerais district. Since the majority of kimberlite discoveries in
Brazil were relatively small, compared to the regions in India, it did not take long to exhaust
them, and thus new sources were required to sustain the growing gem market demand around
the globe. Decades later, in 1866, the first significant diamond discovery was claimed at the
property of the De Kalk farm mine, which is located on the southern bank of the Orange River,
in a close proximity of Hopetown, South Africa. The brownish-yellow stone was eventually cut
into a 10.73 ct glittering brilliant and named Eureka (Levinson, 1997). The splendid find caused
a major diamond rush shortly afterwards. In an instant, the Vaal and Orange River banks were
flooded with prospectors seeking for fortune that was buried in unconsolidated riverbed
sediments. Mining sites became known as “wet diggings” due to the close relationship to water
sources and the visual site created by shallow mining pits.
Between 1869 and 1871, a new age opened in diamond mining in South Africa when
actual diamond host rocks were discovered and documented further inland, away from Orange
River where only few looked for prospects. The precious stones were mined in open pits from
yellowish and reddish weathered surface rock material, mainly clay, also known as “yellow
ground”, which differed from the typical unconsolidated river sediments. At depth, the
weathered ground graded into more compact hard rock material, indicating that there is a new
type of source of diamonds. Due to their distinct bluish color, the host rocks were named the
“blue ground”. The opening of major mines Bultfontein (September 1869), Dutoitspan
(October 1869), Jagersfontein (July 1870), Koffiefontein (July 1870), De Beers (May 1871)
and Kimberley, also known as the Big Hole (July 1871), followed afterwards (Levinson, 1997).
In a short time span, due to the expanding industry, the city of Kimberley was also founded in
1871 close to the Big Hole.
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Professor H. C. Lewis was the first to propose the term “kimberlite” for the newly
discovered diamond host rock in 1887. After new pipe discoveries in the 20th century, especially
in South Africa, it was clear that the kimberlite magmatism is related to a distinctive process
with a unique set of characteristics. Based on petrographic and geochemical differences, Group
I and Group II kimberlites were distinguished. Eventually, the terms “kimberlite” for Group I
kimberlites and “orangite” for Group II kimberlites were accepted by a geological committee.
According to the modified definition by Mitchel (1995), kimberlites are volatile-rich
(CO2 dominant) potassic ultrabasic igneous rocks with a distinctive inequigranular texture due
to presence of phenocrysts or xenocrysts (macro- or megacrysts), which are set in a fine-grained
matrix. The xenocryst suite assemblage consists of anhedral olivine, magnesian ilmenite, Cr-
poor titanian pyrope, sub-calcic diopside, phlogopite, enstatite, Ti-poor chromite and, in rare
occasions, diamond. The surrounding matrix dominantly consists of well-developed second-
generation olivine crystals. Other matrix minerals may include monticellite, phlogopite,
perovskite, spinel, apatite, serpentine, and late-stage poikilitic micas. Common kimberlite
accessory minerals are nickeliferous sulfides and rutile. Deuteric serpentine and calcite
alteration and replacement is a common feature in kimberlite rocks.
Kimberlites are divided into two end-members in respect to their relative differences in
the mineralogy, geochemistry, etc. – Group I (kimberlites) and Group II (orangites). Both types
show similarities with alkaline rocks, such as lamproites and ultramafic lamprophyres (Tab. 1).
The main mineralogical differences between the rock types include (O’Brien, 2015):
CO2 is the dominant volatile phase in kimberlites whereas H2O dominates in orangites.
The CO2/H2O ratio gradually decreases from kimberlites to orangites to lamproites.
Ultramafic lamprophyres tend to have a wider range of volatile compositions with poor
correlation to the trend described earlier.
Kimberlites and orangites generally contain more and relatively coarser mantle-derived
xenoliths compared to lamproites and ultramafic lamprophyres.
Olivine is the dominant mineral phase in kimberlites whereas phlogopite is
characteristic for lamproites. In orangites, the olivine and phlogopite abundances vary, but
tend to lean more towards the lamproite field composition.
From an economic perspective, kimberlites and orangites are the main rock types that
contain sufficiently high diamond grades for profitable commercial mining. Only one known
mine, Argyle in Australia, extracts diamonds from a lamproite rock (Luguet et al., 2009).
There are no known cases of ultramafic lamprophyre containing economically valuable
diamond resources.
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Table 1. Summary of mineralogical differences between kimberlites, orangites, lamproites and
ultramafic lamprophyres (O’Brien, 2015). Published with permission from Elsevier.
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2.2. Geochemistry of kimberlites and orangites
The bulk-rock geochemistry of kimberlites may vary due to contamination with
surrounding rock material that kimberlite passed through on its way to Earth’s surface. During
a rapid magma ascent, wall rocks are brecciated and incorporated into the magma as xenoliths.
Assimilated components are partly or completely dissolved altering the primary magma
composition, with the main added components being SiO2, Al2O3 and Na2O. Hypabyssal
kimberlites represent the least altered source composition while diatreme facies may contain
vast amounts of crustal components, are devolatilized during their emplacement and represent
notable geochemical deviation from the original source. In addition, brecciated rocks are
susceptible towards fluid percolation, weathering processes and alteration, thus further masking
the original makeup. Kimberlite compositions with SiO2 <35 wt% and Al2O3 <5 wt% are
regarded as free of contamination for most of the South African examples (Mitchell, 1986).
Another approach to evaluate the extent of weathering and crustal contamination is to use the
C.I. index (contamination index) defined as (Clement, 1982):
C.I. = (SiO2+Al2O3+Na2O) / (MgO+2K2O)
The C.I. index represents the proportions of clay minerals and tectosilicates relative to
olivine and phlogopite minerals. Regardless, average geochemical compositional data must not
be used as “book example” due to kimberlites’ hybrid nature and emplacement style as
described above.
Kimberlites as a whole are silica-undersaturated rocks with low SiO2 contents of 25-35
wt%, mainly due to its formation deep in Earth’s interior (Backer & Le Roex, 2006; Mitchell,
1986). Al2O3 is usually also low, <5 wt%, and most of aluminum is bound to a micaceous phase.
The MgO content is relatively high at 15-35 wt% with olivine and phlogopite as the dominant
magnesium carriers. The kimberlites that are enriched in olivine macrocrysts, especially Group
I, tend to have a higher Mg# number of >0.85 (atomic Mg/(Mg+Fe2+)) than macrocryst-poor
samples. The Na2O/K2O ratio is extremely low (<0.5), which is related to K enrichment and
relatively low Na concentrations in bulk-rock geochemical composition. Kimberlites have
relatively higher TiO2, CaO and CO2 and lower SiO2 and K2O concentrations compared to
orangites.
Kimberlites have distinct trace element compositions (Smith et al., 1985; Mitchell,
1986; Backer & Le Roex, 2005). Compatible element concentrations of Cr and Ni are controlled
by spinel and olivine fractionation, thus correlating with the MgO content. Ni in orangites tends
to be slightly higher for given MgO and the same applies to Cr. Notably, the referred elements
together with Co, Nd, Zr and Sr are significantly enriched compared to ultramafic and alkaline
9
rocks. Incompatible elements Ba, Sr, Zr, Hf, Nd, Ta, U and Th are mainly hosted by groundmass
minerals, such as phlogopite, perovskite or apatite. Orangites have higher abundances of Ba,
Sr, Zr and Hf, but show relative depletion in Ta and Nd compared to Group I kimberlites.
Group I kimberlites and orangites have well established and similar chondrite-
normalized REE patterns with light REE enrichment over heavy REE. Orangites have more
extreme light REE enrichment, suggesting a highly metasomatised mantle source with more
residual clinopyroxene compared to kimberlites, besides a characteristic feature of orangites is
the distinctly higher La/Sm and La/Yb ratios and lower Gd/Yb ratio. REE patterns indicate that
kimberlites are derived from peridotite sources via very low degrees of partial melting (<1%).
Average initial 87Sr/86Sr and 143Nd/144Nd ratios for Group I kimberlites range between
0.7030 to 0.7055 and 0.5124 to 0.5128, respectively (Smith, 1989; Woodhead et al., 2009). Pb
isotope compositions are distinctly radiogenic. The Sr and Nd isotopic systematics indicate that
Group I kimberlites are derived from undifferentiated and slightly depleted sources relative to
the bulk Earth composition.
For orangites, the average initial 87Sr/86Sr and 143Nd/144Nd ratios range between 0.707
to 0.711 and 0.5118 to 0.5123, respectively (Smith, 1989; Woodhead et al., 2009). The Pb
isotopic signatures are unradiogenic. The Sr and Nd isotopic systematics indicate that orangites
are derived from enriched mantle sources.
2.3. Kimberlite distribution and ages
Kimberlite rocks have been identified in all continents in variable tectonic and
geological settings and occur as relatively small pipes, dikes or sills, often in clusters. Notable
provinces include the Slave and Superior cratons in North America, Sao Francisco in South
America, Kaapvaal craton in South Africa, Siberian craton in Russia, Kimberley area in
Australia and Dharwar craton in India (Jelsma et al., 2009; Shirey & Shigley, 2013). Despite
scattering and intruding into genetically unrelated rocks, the majority of kimberlites are located
within or close to Precambrian terranes, especially those of an Archean age (Fig. 1). In younger
terranes, preservation of kimberlite pipes is limited, since the intrusions have small volumes,
are easily affected by weathering processes and have a high recycling potential in tectonically
active environments.
From an economic perspective, the kimberlites that are located on-craton or close to
their inner margins tend to be diamondiferous, compared to off-craton varieties, which in most
cases are barren (Shirey & Shigley, 2013). For diamond to be stable and survive emplacement
10
in kimberlites, several preservation criteria must be met: diamond xenolith sampling within the
diamond stability field exceeding the depth of 100 km; rapid ascent through a geological
environment with an apparent low heat flow; and emplacement within a tectonically stable
geological setting. Such conditions are plausible in cold (average heat flow 30-50 mWm-2)
Archean terranes that have been tectonically stable for the past 2 Ga and have thick underlying
subcontinental lithospheric mantle roots.
Figure 1. World digital elevation model with superimposed cratonic and shield domains of Precambrian
age (light gray areas on continents), and the global distribution of kimberlites (●), lamproites (○),
melonites (x) and carbonates (+) (Jelsma et al., 2009). Published with permission from Elsevier.
Group I kimberlites are derived from an undifferentiated subcontinental lithospheric
source, while Group II kimberlite melt originates from metasomatised lithospheric mantle
(Becker & Le Roex, 2006). The geochemical characteristics of kimberlites (low Si and high
REE content) suggest that kimberlite magma must be generated as a low-degree partial melt of
a mantle peridotite source under high temperature (>1400 oC) and pressure (>4 GPa) conditions
(Sparks, 2013). The main models that may explain the kimberlite melt formation include
(Jelsma et al., 2009): subduction of the oceanic lithosphere and partial melting of the overlying
mantle; mantle plume activity; thermal perturbations that are associated with major tectonic
events; and multiple origins.
When an oceanic plate is subducted beneath the continental crust, the entrapped fluids
in the slab are released at certain critical temperature conditions, thus causing partial melting in
the underlying mantle (McCandless, 1999). Heat migration upwards up the subducted plate will
further induce partial melting events and generate magma of kimberlitic composition. In this
manner, the oldest kimberlites intrude more inland and become progressively younger towards
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the subduction center trench. In particular, the South African kimberlites emplaced at about
140-60 Ma have such a trend with a progressive westward younging relative to the oldest
kimberlite intrusion (Helmstaedt & Gurney, 1997; Heaman et al., 2003).
Melting can also be promoted by a mantle plume activity. In this case, kimberlites
cluster in narrow paths with progressive younging pattern opposite to the plate movement
direction (Crough et al., 1980). Emplacement patterns of 140-110 Ma Group II kimberlites in
South Africa may be related to the mantle plume activity (Heaman et al., 2003).
Kimberlite magmatism has occurred since the Proterozoic or even earlier times. Only
few intrusions are known from the Precambrian, with the oldest preserved examples (>1.6 Ga)
being located in the Kuruman province, South Africa. Otherwise, the vast majority of identified
kimberlites are younger than 250 Ma (Fig. 2). There is an absence of kimberlite activity between
360 Ma and 250 Ma worldwide and in South Africa in the time periods of 1100-600 Ma and
500-250 Ma. Notably, kimberlite pipe emplacement ages very well correlate with major
palaeocontinent assembly and rifting events (Heaman et al., 2003; Jelsma et al., 2009).
Figure 2. Kimberlite emplacement ages related to the Gondwana assembly and break-up events. The
gray areas mark relatively calm geological activity periods when the Gondwana supercontinent was
stable (Heaman et al., 2003). Published with permission from Elsevier.
In South Africa, multiple kimberlite events took place in the Mesoproterozoic (1635-
1100 Ma), Paleozoic (510 Ma, Pan-African), Mesozoic (240-73 Ma) and Cenozoic (54-31 Ma)
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(Jelsma et al., 2009). The periods with relatively scarce or no kimberlite magmatism at all
correlate with the existence of the palaeocontinents Rodinia and Gondwana. Emplacement
timing may be linked to the assembly or break-up events of Gondwana and Rodinia (Jelsma et
al., 2009) when magmatic activity and partial melting were triggered by subduction processes,
mantle plume activity and stress release. During such events, plates tend to fracture and form
lineaments of various scales. Distinct continental-scale corridor patterns emerge at plate
boundaries that are expressed as local geological terrane boundaries, incipient continental rifts,
and fracture zones or dike swarms. Kimberlite magmas prefer to intrude weakened zones. The
relationship between Jurassic-age kimberlites and structural trend corridors of ENE-WSW and
NNW-SSE directions mark the separation of West and East Gondwana and the opening of the
Somali Basin and Weddell Sea. Cretaceous-age kimberlites with respective trend corridors
aligned NE or NW directions are directly related to the separation of South America and Africa
with the opening of the South Atlantic (Heaman, 2003; Jelsma et al., 2009). Spatial analysis of
kimberlite distribution in South Africa shows 5 major linear trend corridors with directions 35°,
175°, 130°, 75° and 100° (Vearncombe & Vearncombe, 2002). Most of the trends are parallel
to the crustal-scale fracture zones relatively distant from the main faults. More likely,
kimberlites are emplaced and preserved within the individual craton blocks.
A distinct age and location relationship exists between the Group I and Group II
kimberlites. Orangites were emplaced in a short time frame of about 200-110 Ma, in a
condensed area approximate 400 km by 1250 km located in Kaapvaal craton, South Africa
(Mitchell, 1986), while Group I kimberlites have a wide range of emplacement ages spanning
from Precambrian (Kuruman cluster) to most recent Upper Pleistocene/Holocene (The Igwisi
Hills cluster) (Brown et al., 2012). Group I kimberlite intrusions are found in all continents, as
opposed to more restricted occurrences of orangites.
2.4. Kimberlite pipe formation and facies
In the crustal scale, kimberlite pipes typically form downward tapering diatremes that
can extend from the surface crater down to several kilometers in depth before transitioning to a
much narrower root zone (Mitchell, 1986; Sparks et al., 2013). The diatreme walls are steep to
near vertical and commonly have a brittle contact with the host rocks. The pipe infill also partly
consists of brecciated country host rocks mixed with primary kimberlite clasts. At its widest
that is usually in the crater zone, a single kimberlite pipe can cover a surface area up to several
thousand square meters (Fig. 3). Clustering of small-volume pipes in a relatively confined area
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may also occur, especially if the area has undergone brittle deformation and contains a
significant number of faults. The cross sections of the pipe structures are irregular, semi-circular
or elliptic, which represent very well the explosive and destructive nature of emplacement.
Notably, the shape of the pipe is often controlled by the geological features of the surrounding
country rocks.
Figure 3. Left: Proposed kimberlite pipe model with corresponding crater, diatreme and root zones
(Mitchell, 1986). Published with permission from Elsevier; Right: Proposed pipe emplacement model
for Cretaceous kimberlite occurrences in the Kimberly area, South Africa. The 0 m mark represents
current erosion level (Field et al., 2008). Published with permission from Elsevier.
Because most of the kimberlite pipes have undergone an erosion phase and presently
there is no active modern kimberlite volcanism to observe complete stratigraphic details,
emplacement models are highly debated. Generally, proposed theories are based on
observations from collected mining and drilling data from relatively undisturbed rock varieties.
Overall kimberlite morphological characteristics show a resemblance to Maar-type eruption
processes (Lorenz, 1975) and indicate multistage emplacement cycles rather that single
magmatic event.
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The brecciated nature of kimberlitic rocks suggests that the magma ascent occurs in a
dynamic and highly disruptive manner (Sparks, 2013). Also, the ascent must have happened at
high speeds in a relatively short time period to account for the presence of preserved mantle
xenoliths and diamonds in kimberlitic rocks. Sparks et al. (2006; 2013) suggest that
fragmentation and assimilation of surrounding rocks is accomplished by volatile oversaturation
in the melt, especially CO2 that generates a large excess pressure. The saturation may be
accomplished by orthopyroxene assimilation in the melt and CO2 liberation.
Kimberlitic magma finds its way to the surface by exploiting weak structural parts in
the lithosphere; faults and veins provide less resistance for ascending magma. Volatiles tend to
be concentrated at the tip of dikes and due to high pressure, brecciate the front, thus creating
conditions for fast and disruptive ascent. By approaching the surface, factors, such as
temperature decrease, depressurization, volatile interaction with ground water etc., trigger the
formation of a kimberlite pipe. According to Sparks et al. (2006), kimberlite pipe development
can be divided into four stages:
1) Initial cratering. In the first stage, volatile-rich kimberlite magma propagates
along narrow fissures and reach the surface. Reactions with ground water, starting down to
few hundred meters below the surface, and high overpressure brecciate the cold crust and
cause explosive eruptions, resulting in the formation of the main conduct. Kimberlite magma
and any sampled components during its ascent are ejected outwards creating a crater.
2) Pipe formation. In the second stage, erosive processes are the dominant factor
that shape the kimberlite pipe. The initial crater is significantly widened and the diatreme
deepens following the water table and kimberlite magma volatile component interaction
front. Since the crater and diatreme cross-sections increase, pressure starts to drop until it
reaches equilibrium with the atmospheric pressure. Loss of volatiles reduces some of the
initial magma volume in the pipe and explosive events start to seize. At the deeper levels,
pressure still remains high and causes rock bursting and disintegration that can still deepen
the diatreme down to 2-3 km. After a certain time period, magma replenishment starts to
decline, temperature drops and magma starts to cool slowly.
3) Pipe filling. After initial eruptions seize and no more material is ejected outside
of the pipe, diatreme starts to be infilled. Weakened and fractionated pipe walls constantly
collapse inwards together with the ejected clasts from the crater facies. Layered features are
established with fine and well-preserved layering closer to the surface. Late-stage dike and
sill intrusions may accompany and disrupt the infill process. It is still possible that Stage II
and Stage III overlap creating several emplacement cycles and further destroying original
structures.
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4) Post-emplacement metamorphism and alteration. In the final stage, highly
porous infill material is subjected to hydrothermal alteration that mainly is caused by
meteoric water circulation. Serpentinization is the principal alteration product due to the high
olivine content in the kimberlite magma. Formation of carbonates and apatite is also
common.
A typical kimberlite pipe consists of three distinct zones or facies that are associated
with a particular style of magmatic activity as described earlier (Fig. 3): crater facies, diatreme
facies, and hypabyssal facies. In the early works, the kimberlite rock classification schemes
were based on textural and genetic relationships (Fig. 4) and related to each of three kimberlite
pipe facies (Clement & Skinner, 1979; Dawson, 1980; Mitchell, 1986).
Figure 4. Textural genetic classification scheme for kimberlite rocks (Mitchell, 1986). Published with
permission from Springer.
In the modern literature, the classification scheme by Cas et al. (2008; 2009) is
commonly used for kimberlite rock description. A more simplified classification of fragmental
kimberlite rocks by Sparks et al. (2006) may also be applied.
Besides the standard kimberlite mineralogy (Mitchell, 1986), the principal components
in volcanoclastic kimberlite rocks are the “pelletal” lapilli clasts. Pelletal lapilli are spherical to
elliptic clasts that consist of a single crystal (commonly olivine) or lithic core surrounded by a
fine-grained coherent kimberlite rim. The lapilli range in size from 1 to 10 mm. The origin of
pelletal lapilli is linked with intra-vent fluidization processes. Autoliths differ from lapilli in
that they are coarser (up to 8 cm) and may contain large fragments of country rock. Term
“autoliths” is not used anymore in modern literature.
According to Sparks et al. (2006; 2013), volcanoclastic kimberlites can be divided into
four principal types: massive volcanoclastic kimberlite (MVK), layered volcanoclastic
16
kimberlite (LVK), marginal wall-rock breccia, and magmatic kimberlite (MK). The MVK is a
homogenous volcanoclastic rock that is composed primarily of lithic clasts, juvenile kimberlite
pyroclasts, and crystals sampled from various stratigraphic units which the kimberlite magma
passed through. Clastic material is mostly pelletal lapilli and coarse ash. Breccias (>64 mm) are
rare or absent. The crystal suite is composed of euhedral olivine crystals or crystal fragments
with various degree of alteration. The matrix of MVK is composed of fine ash or secondary
minerals, which fill the original pore space. Heavy serpentinization may mask the primary
matrix composition. As the dominant diatreme facies, MVK is typically located in the central
part of a pipe and may cut earlier formed LVK bodies.
The layered volcanoclastic kimberlite (LVK) is distinguished from MVK by the
presence of layering and is commonly found in the crater facies, especially at higher levels.
Fine LVK layering tends to grade into MVK in the diatreme zone. LVK contains lapilli tuffs,
lapilli stones and breccias with fine- to medium-sized bedding structures. Clast layers
commonly are poorly sorted and tend to dip steeply (up to 400) towards the crater center, which
is consistent with mass flow or sliding and avalanche-like deposition. In narrow pipes, LVK
can be found even in deeper pipe parts and dominate over MVK.
The marginal wall-rock breccias are dominated by country rock fragments mixed
together with MVK or LVK facies rocks. As the term implies, breccias form a narrow transition
zone between the country rock and kimberlite pipe infill. The width of the marginal wall-rock
breccia depends on the country rock strength and depth of volatile percolation, which is the
main disruption force. Wall collapse is common in such an environment. This is why the
breccias can develop a crude layering towards the pipe center and interlay MVK facies rocks.
The root zone is dominated by magmatic kimberlite (MK), also known as hypabyssal
kimberlite, which forms coherent, irregular or ellipsoidal bodies. Although MK can also be
found in the diatreme and crater facies, as the magma with MK fragments is pushed upwards
during eruption, gradual transition into MVK is common. MK contains abundant macrocrysts
and altered lithic clasts that are set in a fine-grained groundmass composed of monticellite,
spinel, perovskite, calcite, and serpentine minerals.
17
3. Mineralogy and classification of ultramafic mantle xenoliths
3.1. Mineralogy
Olivine (Mg,Fe)2SiO4 is the dominant constituent in peridotitic rocks and mantle
xenoliths found in kimberlite pipes. Also, olivine is the main iron, magnesium and nickel carrier
in peridotites in general. Various elements, such as Ni, Ca, Na, Al, and Cr, may substitute Fe
and Mg depending on the pressure and temperature conditions. Since only limited major
element compositional variability occurs in olivine, compared to other elements and mantle
rock forming minerals, Mg# in olivine (Mg# = 100*Mg/(Mg + Fe)) also reflects the whole-rock
composition and can be related to the degree of melt depletion or Fe enrichment in the system
(De Hoog et al., 2009; Pearson et al., 2003). Since the fayalite component is preferably extracted
prior to the forsterite component, depleted peridotites have higher Mg# values compared to the
primitive or metasomatized sources with lower Mg#. A similar relationship occurs in off- and
on-craton mantle xenoliths found in kimberlites, showing Mg# values of 88-92 and 91-94,
respectively. Nickel in olivine tends to increase with increasing forsterite content for upper
mantle peridotites. On the other hand, the manganese contents decrease with increasing
forsterite content in olivine (De Hoog et al., 2009). Both element concentration relationships
with Mg# can be used as an index of melt extraction and depletion (De Hoog et al., 2009;
Pearson et al., 2003).
Garnets are sampled and brought to the surface by kimberlite magmatism in the form of
xenolith constituent, xenocrysts or macrocryst inclusions. The chemical formula of the mineral
garnet is expressed as X3Y2[ZO4]3, where site X is occupied by divalent Ca2+, Mg2+ or Fe2+
cations and site Y is occupied by trivalent Al3+, Fe3+ or Cr3+ cations. Si4+ usually constitutes the
Z site, forming the silicon-oxygen tetrahedron, but in rare cases, it can be substituted by Al, Ti
or Fe3+. In nature, garnet rarely forms pure single mineral phase, thus its composition is
expressed as a solid solution between two or more mineral end-members (Demange, 2012). The
abundance of major oxides MgO, FeO, CaO, Cr2O3 and TiO2, and their ratios, such as Mg# and
Ca# (100*Ca/(Ca+Mg) mol), in garnet solid solution are used for classification purposes. The
most commonly applied classification system is the cluster analysis-based G1-G12 scheme by
Dawson (1975) or further expanded variations by Schulze (2003) for crustal/mantle source
clarification and by Grutter (2004) for diamond exploration purposes. In general, the cluster
scheme implies that: (G1-G2) garnets belong to a megacrystic suite; (G3) are eclogite garnets;
(G4-G5) are characterized as pyroxenitic garnets; (G6-G8) garnets are dominantly kimberlite
18
derived; (G9) belong to lherzolitic paragenesis; (G10) are harzburgitic garnets; (G11) Ti-
metasomatised variation; and (G12) fall into the wehrlitic field. Term (G0) is used for
unclassified garnets that do not belong to any of the listed groups and are usually characterized
by low (<1 wt%) Cr2O3 and (<2 wt%) CaO contents.
Orthopyroxene belongs to the low-calcium (<5 wt% CaO) pyroxene group and consists
of two main mineral types, named enstatite Mg2Si2O6 and ferrosilite Fe2Si2O6. Orthopyroxene
is a common constitute in peridotitic rocks (Morimoto et al., 1988; Demange, 2012). The Mg#
value of orthopyroxene in mantle xenoliths usually is similar to that of olivine, but slightly
higher due to relative difference in the Mg-Fe partition coefficients. The CaO and Al2O3
contents are mainly controlled by temperature and pressure conditions. CaO is low at 0.2-2.0
wt% and tends to increase with increasing temperature. The Al2O3 content varies depending on
the rock type. In the garnet peridotite facies, the Al2O3 content of orthopyroxene is usually
below 2 wt%, whereas in the spinel peridotite facies, it is between 1 and 6 wt%. Cr2O3 is a
common substitute for Al2O3 in orthopyroxene (Pearson et al., 2003). Notably, the Ca and Al
content dependence of temperature are pressure conditions can be used for thermobarometry
(Brey & Kohler, 1990).
Clinopyroxene belongs to the calcium-rich pyroxene group and contains several mineral
types, of which the most common ones are diopside CaMgSi2O6 and augite (Ca,Mg,
Fe2+,Al)2Si2O6, which are also common constitutes in peridotitic rocks (Morimoto et al., 1988;
Demange, 2012). Clinopyroxene is a major host of Na, Ca, Cr and Ti in mantle rocks and related
xenoliths. Element concentrations tend to vary depending on temperature and pressure
conditions and can be related to depletion events (Pearson et al., 2003). Clinopyroxene and
orthopyroxene tend to form solid solutions, and therefore the element exchange can be used for
thermobarometry (Brey & Kohler, 1990). The presence of secondary clinopyroxene in xenoliths
usually indicates infiltration of host magma. It can be found as a minor phase around primary
minerals, especially garnet.
Oxide minerals of cubic symmetry with the general formula X2+Y23+O4 belong to the
spinel group, where the X site is occupied by elements Mg or Fe2+ and Y is occupied by Al,
Fe3+ or Cr. Term spinel also represents the end-member MgAl2O4 ‘spinel sensu stricto’ of the
spinel group or one of the three Mg-Fe2+ solid solution series within the group, named – spinel,
magnetite and chromite series (Gill, 2010). The chemical composition of spinel varies greatly
in peridotitic rocks, depending on P-T conditions (Barnes & Roeder, 2001; Pearson et al., 2003).
For example, Cr# shows a positive correlation with Fe/Mg ratio in spinel, which is also strongly
temperature dependent. Notably, Cr# directly reflects the degree of depletion of the bulk rock.
Spinels that are found in depleted xenoliths or in high pressure garnet-facies xenoliths show
19
high Cr# values, whereas spinels found in less-depleted xenoliths or low-pressure spinel-facies
xenoliths have lower Cr# values. The TiO2 content in mantle spinels is generally low (<0.5
wt%) and correlates positively with the Fe content and degree of depletion.
3.2. Modal and textural classification
Kimberlites can sample a wide range of peridotitic rock (containing >40% olivine)
xenoliths from the mantle during the magma ascent. Such xenoliths usually are coarse grained.
Figure 5. (A) IUGS classification of ultramafic rocks (modified after Le Bas & Streckeisen, 1991); (B)
P-T diagram showing stability fields of plagioclase, spinel and garnet lherzolites and two examples of
geothermal gradients with surface heat flows of 40 and 90 mW/m2 (modified after Wörner, 1999).
Figure 5A shows the IUGS modal classification system of ultramafic rocks (Le Bas &
Streckeisen, 1991). Ultramafic rocks containing <40% olivine are called pyroxenites. Peridotite
is further subdivided into four varieties based on the relative amounts of orthopyroxene,
clinopyroxene and olivine. Dunite is almost entirely composed of olivine >90% and contains
only a minor amount of combined ortho- and clinopyroxene (<10%). Rocks containing up to
5% clinopyroxene but abundant olivine and orthopyroxene are harzburgites and rocks with
minor orthopyroxene <5% but abundant olivine and clinopyroxene are called wehrlites.
Lherzolite is somewhat in between of all three extremes, containing essentially olivine >40%
and various proportions of pyroxenes. Lherzolites are commonly four-phase rocks which
contain, in addition to orthopyroxene, clinopyroxene and olivine, one aluminous phase, garnet,
spinel or plagioclase. Depending the type of the Al-rich phase, the rocks can be subdivided into
garnet, spinel and plagioclase lherzolites, which are stable under different P-T conditions. As
shown in Figure 5B, garnet lherzolites are stable at highest pressures and plagioclase lherzolites
20
at lowest pressures. Also shown in the diagram are two different geotherms. Progressive melt
depletion of mantle peridotite tend to produce a rock series lherzolite > harzburgite > dunite.
Xenolith textures are often described according to the classification scheme of Harte
(1977). He defined four principal olivine-bearing xenolith textural types based on the varying
degrees of deformation and recrystallization: coarse, porphyroclastic, mosaic-porphyroclastic,
and granuloblastic. Five subtypes of these principal textural types further describe the xenolith
in a greater detail: equant, tabular, laminated, fluidal, and disrupted.
21
4. Continental lithospheric geotherms
The lithosphere is an outer thermal layer of the Earth, in which heat is transported
dominantly by conduction compared to the asthenosphere where heat is transferred primarily
by convection. This heat flow relationship defines the thermal state of the lithosphere (Stuwe,
2009; Furlong & Chapman, 2013). Heat is constantly lost through the Earth’s surface with a
global average heat flow of 87.0±2.0 mWm-2. In the case of the continental heat flow, the
average heat loss is 65.0±1.6 mWm-2, but for the oceanic regime, the average heat loss is
significantly higher, 101.0±2.2 mWm-2, especially near the oceanic plate rift zones (Pollack &
Chapman, 1977; Stein, 1995). Observations from the heat flow variations imply that the thermal
state of lithosphere changes over time (Fig. 6), and thus Archean terranes are characterized by
a low average heat flow of 51.0±25.6 mWm-2, compared to the young oceanic lithosphere close
to rift zones.
Figure 6. Surface heat flow variation among continental terranes relative to tectonothermal age (Stein,
1995). Published with permission from John Wiley and Sons.
A geotherm is a function that describes temperature change in the lithosphere as a
function of the corresponding depth. Surface heat flow is usually used as a principal
independent variable to construct the regional geotherm profiles (Pollack & Chapman, 1977;
Stuwe, 2009). Radioactive heat production within the mantle and crust balances the heat loss,
and consequently the thermal boundary layers have a relatively constant temperature profile.
In a steady-state one dimensional model, the surface heat flow is in equilibrium with the
heat transfer into the lithosphere at its base and the radiogenic heat produced in the lithospheric
regime. Within the lithosphere, the heat production (A) varies with depth and the thermal
conductivity (k) varies according to compositional, temperature and pressure relationships.
22
Given these relationships, the steady-state one-dimensional heat conduction in a layer can be
expressed with the following equation (Furlong & Chapman, 2013):
𝑇(𝑧) = 𝑇𝑡 + (𝑞𝑡
𝑘) 𝑧 − (
𝐴
2𝑘)𝑧2
where additional Tt is temperature and qt is the heat flow at the top of the thermal layer and z is
the depth within the layer.
Thermal conductivity and radiogenic heat production at the surface (A0) are solved by
the following equations:
𝑘(𝑇, 𝑧) =𝑘0(1 + 𝑐𝑧)
1 + 𝑏(𝑇 − 20) ; 𝐴0 =
𝑃𝑞𝑏
𝑏
where conductivity k0 is a laboratory-determined value at room temperature of 20 0C and b and
c are temperature coefficient constants. If the thermal layer in question has a thickness of Δz
and the geotherm is expressed in terms of a succession of such layers, the geotherm can be
calculated by the following equation:
𝑇𝑏 = 𝑇𝑡 + (𝑞𝑡
𝑘) ∆𝑧 − (
𝐴
2𝑘) ∆𝑧2 ; 𝑎𝑛𝑑 𝑞𝑏 = 𝑞𝑡 − 𝐴∆𝑧
where Tb is temperature and qb is heat flow through the bottom of the thermal layer. Since most
of the heat is produced within the crust and there is a relatively low contribution from the
mantle, the overall heat production in the lithosphere can be expressed as an exponential
function:
𝐴(𝑧) = 𝐴0exp (−𝑧
𝑏)
According to the calculations by Pollack & Chapman (1977) and Furlong & Chapman
(2013), a set of conductive continental geotherms can be obtained as represented in Figure 7.
The continental geotherm family must satisfy the following geological constraints to be
a valid representation of thermal regime in question (Furlong & Chapman, 2013):
1. The surface temperature at zero-depth should be used as an annual average of the
particular region to which the geotherm is applied. As a default, an annual global
temperature of 15 0C can also be used for calculations.
2. Continental conductive geotherms should approach melting conditions near the base of
the lithosphere.
3. Continental conductive geotherms should characterize subsurface pressure and
temperature values inferred from crustal and mantle xenoliths or from exhumed
metamorphic rocks.
23
Figure 7. Conductive continental geotherm family according to calculations by Pollack & Chapman
(1977) and Furlong & Chapman (2013). Light blue field represents P-T estimates from mantle xenolith
data from Hasterok & Chapman (2011) work (Furlong & Chapman (2013).
For the Kalahari craton, the conductive continental geotherm is characterized by a
constant surface heat flow of 40.0±3.0 mWm-2 (Hasterok & Chapman, 2011). Xenolith pressure
and temperature estimates from the Letlhakane, Letseng, Premier and Frank Smith kimberlites
also plot on the Kalahari craton conductive continental geotherm (Boyd & Nixon, 1975;
Danchin, 1979; Steifenhofer et al., 1997; Dluda et al., 2006).
24
5. Geological setting of the Kaapvaal craton
5.1. General features
The Kaapvaal craton is the oldest, largest and best-preserved terrane in southern Africa
(de Wit et al., 1992), which hosts several globally important mineral deposits, including highly
dense diamondiferous kimberlite populations. This ancient craton is an assemblage of several
collided Archean terranes composed mainly of granitoids, gneisses and greenstone belts, which
cover an area of approximately 12,000 km2 (Fig. 8). From South Africa, the Kaapvaal craton
partly extends into Botswana, Lesotho, Swaziland, and Mozambique, which also benefit from
the vast resources available on- and off-craton.
Figure 8. Core of the Kaapvaal craton with its principal terranes, rock units and position in the
geopolitical map (Schmitz et al., 2004). Published with permission from Elsevier.
25
The complex Namaqua-Natal Belt borders the Kaapvaal craton in the south and west,
while the eastern margin follows the Mozambique border and is shaped by the Lebombo
monocline of Jurassic volcanics, a remnant from the Gondwana break-up event. Further to the
north lies the Zimbabwean craton which is separated from the Kaapvaal craton by the Limpopo
mobile belt. Furthermore, the term Kalahari Craton is used in order to describe combined
Kaapvaal and Zimbabwean cratons and the surrounding mobile belts (Jacobs et al., 2008).
The Kaapvaal craton is characterized by a very thick and complex cratonic root,
extending down a depth of 275 km (Griffin et al., 2003; Kusov & Kronrod, 2006). Based on
collected geochemical data on mantle xenoliths from Group II kimberlites (Becker & Le Roex,
2006; Giuliani et al., 2015), the underlying lithospheric mantle beneath the Kaapvaal craton has
been periodically partly or heavily metasomatised over time, especially in the north-eastern
sections of the craton. As an Archean terrane, the Kaapvaal craton has cooled down over the
years and has a low apparent continental heat flow of 35-45 mWm-2, compared to the
surrounding mobile belts that have a much higher heat flow of ~80 mWm-2. The heat flow and
geophysical data indicate that the surrounding mobile belts have a thinner lithospheric
component, which presumably does not exceed more than 120 km in thickness (Eglington &
Armstrong, 2004).
The principal evolution of the Kaapvaal craton covers a time period of approximately 1
Ga during the Archean and Paleoproterozoic eons. Two major development stages can be
distinguished (de Wit et al., 1992):
1) Kaapvaal shield formation (3.7-3.1 Ga). From 3.7-3.2 Ga, the first Kaapvaal
continental lithosphere emerged through intraoceanic subduction, thrusting and melting
processes, similar to the environment of formation observed in the modern-day ocean basins
in Western Pacific, such as the Ontong-Java Plateau. At that time, mafic and komatiitic
volcanism dominated in the oceanic back-arc basins, accompanied by chemical and clastic
sedimentation. The Barberton greenstone belt and the Ancient Gneiss Complex formed
during this stage. Initially, oceanic island arcs collided with the granitic micro-terranes and
amalgamated at 3.3-3.2 Ga to form a core for the new shield. Extensive igneous activity
continued (3.2-3.1 Ga) with oceanic crust recycling, composite batholite formation and
chemical differentiation in the upper lithosphere driven by mafic-ultramafic crust hydration
before the final Kaapvaal shield thickening and stabilization at 3.1 Ga.
2) Kaapvaal craton formation (3.1-2.6 Ga). The second stage of development
largely concerns the regional extension of the Kaapvaal shield through intra-continental and
passive continental-margin tectonic processes, such as subduction and tectonic accretion of
composite crustal fragments. In the compressive environment, S-type high-Ca granites were
26
formed. Further craton extension was influenced by the overlapping foreland Witwatersrand
sedimentary basin development at 2.7-2.9 Ga in the central part and the following
Ventersdorp volcanism and sedimentation event at 2.7-2.6 Ga. In the northern part, cratonic
extension led to continental shelf formation with passive margin sediment accumulation.
Eventually the shelf was tectonically juxtaposed between the granite-greenstone terrains of
the Kaapvaal and Zimbabwe cratons at 2.68 Ga during the Limpopo orogeny. The Kaapvaal
craton stabilized by 2.6 Ga, but active igneous activity associated with Limpopo Belt
continued until ~2.5 Ga.
In the Paleoproterozoic, the geological activity in the Kaapvaal craton was dominated
by intracratonic extensional processes. First proto-basins of the Transvaal Supergroup started
to develop as early as ~2.65 Ga (Thomas et al., 1993; Eglington & Armstrong, 2004). From
~2.4 Ga to ~2.1 Ga the Pretoria Group Basin and the Postmasburg Groups were produced.
Shortly after the deposition of the Transvaal Supergroup was terminated at ~2.1 Ga, the
emplacement of the giant Bushveld Complex took place in the central part of the Kaapvaal
craton at 2.06 Ga. At about 2.02 Ga, the Vredefort structure was formed. It is thought to be the
oldest and largest meteorite impact crater ever found and explored on Earth. In the middle
Proterozoic, the geological activity was associated with the Eburnian orogenic event, which
started at ca. 2.0 Ga and seized at ca. 1.4 Ga, resulting in the formation of the Kheis and
Richtersveld Provinces in the western Namaqa-Natal Belt. The Kibaran orogeny followed at
1.2-0.9 Ga and completed the development of the Namaqa-Natal Belt.
In the late Proterozoic the “Pan-African” tectonothermal event took place, which
resulted in the development of several intra-cratonic foreland basins including the Nama and
Natal Groups. Due to the subduction of the paleo-pacific plate underneath the Gondwana plate
the Karoo basin opened from the late Carboniferous to the middle Jurassic, producing several
sedimentary sequences. The deposition was followed by the eruption of the Karoo flood basalts
when the Gondwana supercontinent started to break apart at 200 Ma with the initial separation
of Australia. Between 180-160 Ma, the eastern and western Gondwana parts separated from
each other. South America was separated at 135 Ma, which essentially ended the Kaapvaal
craton development. Parallel to the Gondwana break-up, the majority of kimberlites were
emplaced in southern Africa.
27
5.2. Crustal components of the Kaapvaal craton
The crustal section of the Kaapvaal craton varies in thickness from 30 to 50 km,
depending on the internal composition and complexity (James et al., 2003). The average crustal
thickness is estimated to be approximately 35 km. From the geophysical surveys performed in
the cratonic and adjacent areas, seismic wave velocities and density contrasts indicate that
Kaapvaal craton crustal components have dominantly felsic to intermediate compositions. The
Moho boundary is expressed as a distinct transitional layer ranging from ca. 0.5 to 1.0 km in
thickness. Despite the collisional history between the terranes that make up the craton, the
central part of the Kaapvaal craton is relatively uniform with the crustal thickness ranging from
~38 to 41 km; towards the south-east and south-west, the crust is thinner with a thickness of
~36 km; towards the north along the Limpopo Belt, the crust thickens up to approximately 50
km. The transitions and thickness contrasts in the southern and northern Kaapvaal craton
boundaries are relatively sharp.
The Kaapvaal craton is divided into three distinct crustal segments (Fig. 8), which are
named the Kimberly, Witwatersrand and Pietersburg Blocks. Together with the Zimbabwe
craton and surrounding mobile belts, the southern African terrane is further subdivided into 19
structural units (Fig. 9) (Griffin et al., 2003; Schmitz et al., 2004).
The Southeastern Terrane (Fig. 9 IIa) of the Witwatersrand block hosts the oldest
known rock formations in the Kaapvaal craton: the Ancient Gneiss Complex, which is
dominantly composed of TTG gneisses emplaced at ~3.6-3.2 Ga, and the NE-SW-trending
Barberton greenstone belt emplaced at 3.57-3.08 Ga. Further to the south-east lies the Pongola
Supergroup (3.3-3.1 Ga), which is characterized by late granite intrusions surrounded by earlier
sandstone, shale and basalt formations. In the north-western part of the terrane lies the
Witwatersrand Supergroup, which hosts several quartzite, conglomerate and lava sequences
disturbed and uplifted by the Vredefort meteorite impact structure. The sedimentary and
intermediate to ultramafic volcanic rocks of the Ventersdorp Supergroup partly overly the
Witwatersrand strata in the west. The Ventersdorp Group rocks further extend into the Kimberly
block. At the northern margin of the Southeastern Terrane lies the Johannesburg granite dome
and the surrounding greenstone belts. The rest of the block is covered by sandstones, shales and
tillites of the Karoo Supergroup (300-180 Ma), mainly are found into the Lesotho territory. The
Letseng kimberlite cluster has intruded into the Karoo volcanics.
28
Figure 9. Principal structural units and their inferred boundaries in the southern African territory
superimposed on a shaded map of smoothed aeromagnetic anomalies (Griffin et al., 2003). Blue stars
mark the kimberlite pipes examined in this study: (1) Letseng; (2) Frank Smith; (3) Premier; (4)
Letlhakane. Published with permission from Elsevier.
The Central Terrane (Fig. 9 IIb) of the Witwatersrand block is characterized by several
Archean granitic gneiss and migmatite complexes (3.2-3.1 Ga), which are mainly present in the
eastern part of the terrane. The Archean rocks are partly overlain by rocks of the Bushveld
layered igneous complex (2.05 Ga), which largely dominate the terrane and even extend further
to the north into the Pietersburg Terrane. Rooiberg rhyolites and Lebowa granites occur in the
center of the intrusion and Rustenburg Layered Suite of mafic-ultramafic cumulates at the sides.
The Premier kimberlite is present in the western part of the intrusive complex. Further on, the
igneous complex is surrounded by the Transvaal Supergroup sedimentary formations composed
of dolomites, limestones, BIF’s, shales, and quartzites.
The Pietersburg Terrane (Fig. 9 IIc) is a small and narrow block that hosts the
Pietersburg, Murchison and Giyani greenstone belts in the east, surrounded by TTG gneisses
(2.9-2.8 Ga), granodiorites and granites (2.7-2.6 Ga). To the west lie the Waterberg Group
arkose and conglomerate formations. Layered mafic rocks and granites of the northern limb of
the Bushveld igneous intrusion are also present in the central part of the block. The Pietersburg
29
block is separated from the Witwatersrand block in the south by the Thabazimbi-Murchison
lineament and from the Limpopo Belt in the north by the Palala shear zone.
The Kimberly Block (Fig. 9 IId) constitutes the principal part of the western Kaapvaal
craton and hosts the Frank Smith kimberlite intrusion. The crystalline basement of the terrane
is poorly exposed; only small N-S-trending Kraaipan, Amalia and Madibe greenstone belts
composed of mafic to intermediate volcanic and metasedimentary rocks outcrop in the north-
east. The greenstone belts are surrounded by various TTG gneiss and granodiorite complexes,
together with the well-known Schweizer-Reneke quartz monazite dome. The rest of the
Archean and Proterozoic terrane is composed of similar granitic gneisses and granitoids that
are overlain by Ventersdorp and Karoo Supergroup volcanics and sedimentary sequences.
Further to the north lie the Cenozoic Kalahari sands.
The Archean Limpopo Belt (van Reenen et al., 1990; Griffin et al., 2003) is a high-grade
metamorphic terrain located between the Kaapvaal and Zimbabwe cratons. During the
collisional event, the Limpopo microcontinent was thrusted over both cratons, resulting in
crustal thickening and granulite-facies rock development in the area. The Central Zone (Fig. 9
IIIa) consists of granulite- and amphibole-facies gneisses composed mainly of granodiorite,
diorite, and tonalite. The terrain is separated from the remaining units by a series of shareed
zones that developed during the collisional event. The Northern (Fig. 9 IIIb) and Southern
Marginal Zones (Fig. 9 IIIc) are characterized the presence of tectonically dismembered
greenstone belts composed of ultramafic-mafic and metapelitic gneisses with sporadic BIF’s.
Both zones are intervened with tonalitic and trondhjemitic gneisses.
The majority of kimberlites emplaced in Botswana are related to the north-western
Kalahari craton passive margin of a Paleoproterozoic age (Griffin et al., 2003), which is linked
to the crustal development during the Eburnean orogeny. The NW-SE-trending fold and thrust
belt consists of three principal structural units. To the south lies the Kreis Fold Belt (Fig. 9 Va),
which is separated from the Kimberly Block in the east and from the Neoproterozoic Pan-
African orogeny-related rocks in the west by a series of N-S-trending shear zones. The Kreis
Fold Belt consists primarily of deformed sedimentary rocks (1.93-1.89 Ga), which are disturbed
by several syn- and post-tectonic granite plutons and Kalkwerf augen gneisses (1.27 Ga) in its
central part. Part of the formation is covered by Phanerozoic rocks. The Okwa Inlier (Fig. 9 Vb)
is a northern extension of the Kreis Fold Belt, located in central-western Botswana. The poorly-
exposed Paleoproterozoic basement (2.0-2.1 Ga) forms a well-defined magnetic structure,
which mainly consisting of gneisses and granites divided into four principal lithological units.
In the southeast, the Okwa Inlier is bound to the Kaapvaal craton and in the northwest, is
30
unconformably overlain or in contact with clastic sediment and limestone formations of the
Neoproterozoic Ghanzi-Chobe belt (Mapeo et al., 2006).
The Makondi Belt (Fig. 9 Vc) constitutes the northern part of the Proterozoic passive
margin that surrounds the Kalahari craton (Griffin et al., 2003) and hosts the Letlhakane
kimberlite. Deformed and metamorphosed sedimentary and volcanic rocks of the Makondi
Supergroup partly wrap and extend around the western margin of the Zimbabwe craton and are
thrusted on the Archean and Paleoproterozoic granite-greenstone basement (Treloar, 1988).
The Makondi Supergroup rocks are subdivided into three principal stratigraphic units. The
lowermost Deweras Group consists of basal conglomerates, quartzites, basaltic lavas, and tuffs,
which are overlain by Lomagundi Group orthoquartzites, dolomites, argillites, sandstones and
greywackes. The uppermost part of formation is composed of phyllites, greywackes and shales
of the Piriwiri Group. The Makondi sequence is unconformably overlain by three younger
successions: the Sijarira Group sedimentary rocks, the Makuti Group sedimentary rocks, and
the Karoo Group sedimentary and volcanic rocks.
31
6. Kimberlites sampled for this study
6.1. Premier kimberlite
The Premier Group I kimberlite (Scott & Skinner, 1979; Field et al., 2008) was
discovered in the late 19th century during a prospecting project for heavy minerals near town of
Pretoria in Transvaal Province, South Africa. Ten more kimberlite intrusions were later found
in the area. Notably, the Premier intrusion has so far the largest by areal extent of the kimberlites
that are known to have erupted in the Kaapvaal craton. The roughly elliptical kimberlite body
covers a surface area of ~32 ha, with its major axis being approximately 1 km in length. The
crater zone of the Premier kimberlite has been eroded away. The first excavation operations
from the diamondiferous kimberlite started in 1903 and sporadically continued throughout the
century. In 2003, the Premier Mine was renamed as the Cullinan Mine due the largest rough
diamond ever recovered from the site in 1905 – the 3106 carat Cullinan diamond.
The Premier kimberlite is situated in the Central terrane of the Kaapvaal craton and it
has intruded various rocks of the Transvaal Sequence (dolomite, shale and quartzite), the
Bushveld complex (peridotite, pyroxenite, norite, gabbronorite, and dolerite) and the Waterberg
Group (quartzite and conglomerate). The kimberlite diatreme zone is intersected by a late ~70-
m-thick gabbroic sill at the ~500 m level. The emplacement age of the Premier kimberlite is
estimated to be Precambrian with an age of 1179±36 Ma (Smith, 1983), making it one of the
oldest known kimberlite occurrences in the Kaapvaal craton. The pipe infill is most complex,
as there are at least eight different types of kimberlite present, of which three distinct kimberlite
rock varieties termed Grey, Brown and Black kimberlite dominate. Each of the three types
represents separate magmatic events (Scott & Skinner, 1979). The Grey and Brown kimberlites
are considered to be typical diatreme facies rocks. Both are highly brecciated and contain up to
~40% crustal inclusions from the Bushveld Complex and Waterberg Group rocks. The Black
kimberlite is a massive hypabyssal facies rock, rich in ilmenite and serpentinized olivine, with
the crustal component taking up to ~20% of the total volume. The final magmatic activity in
the diatreme is marked by the emplacement of a carbonate-rich dike (magnetite-serpentine-
calcite) in the deeper parts of the pipe, which caused country rock carbonatization and
metasomatism (Robinson, 1975). Two distinctive modes of occurrence have been recognized,
one cutting the Grey kimberlite and the other one intruding the Black kimberlite.
The Premier kimberlite contains a large number and variety of mantle xenoliths. The
ultramafic inclusion suite consists of garnet lherzolites, garnet harzburgites, garnet websterites,
32
dunites, pyroxenites, chromite peridotites, and chromite harzburgites, Eclogites and discrete
nodules consisting predominately of garnet, diopside, enstatite and ilmenite have also been
recovered (Danchin, 1979). However, due to poor recovery methods in the early mining years,
mostly abundant garnet lherzolites and garnet harzburgites have been studied in greater detail.
Garnet lherzolite xenoliths are subdivided into two types – the deformed and coarse
varieties (Danchin, 1979). The deformed, extensively recrystallized variety tends to be enriched
in Ti and K. Based on the mineral geothermobarometry, the deformed garnet lherzolite
xenoliths are derived from a depth of around 200 km, whereas the coarse xenoliths originate
from shallower depths of 110-170 km. The coarse garnet lherzolite xenoliths are more depleted
in Fe, Al, Ca and Ti relative to the deformed type.
The garnet harzburgite xenoliths are subdivided into three distinct types (Danchin,
1979). Type I xenoliths show strong similarities to the deformed garnet lherzolite variety, such
as a high TiO2 content, an equivalent enstatite Ca# (>20), and the depth of origin. On the other
hand, Type III garnet harzburgites have similarities with the coarse garnet lherzolite variety,
thus showing dominantly coarse-grained textures, TiO2-depleted garnets, corresponding
enstatite Ca# (<20), and shallower depths of origin at 140-180 km. Type II garnet harzburgites
have intermediate compositions and characteristics compared to Type I and III garnet
harzburgites.
Mantle eclogites from the Premier kimberlite belong mainly to the Group I variety based
on textural and mineral chemistry criteria (Dludla et al., 2006): interlocking and corona
textures; Na2O >0.09 wt% in garnet and K2O 0.08 wt% in clinopyroxene. The average modal
proportions of clinopyroxene and garnet minerals in eclogites are 55:45, and phlogopite is the
dominant secondary phase. Geothermobarometric calculations at an assumed pressure of 50 kb
give an average equilibrium temperature of 1102±37 0C and sampling depths of 135 to 165 km
for the Premier eclogites, relative to the shield geotherm of 40 mW/m2.
6.2. Letseng kimberlite
The Letseng Group I kimberlites (Field et al., 2008) were discovered in 1957. There are
two intrusions present, referred to as the Main pipe (17.2 ha) and the Satellite pipe (5.2 ha).
Few large diamonds have been recovered from both pipes since first mining efforts commenced
at the site in year 1957. The most famous gem quality diamond found from Letseng Mine is the
Lesotho Promise 603 carat diamond.
33
The Letseng intrusion belongs to a larger kimberlite cluster of about 60 occurrences
located in the north-eastern part of the Lesotho highlands in the Tugela terrane (Sommer et al.,
2013), at the edge of the south-eastern Kaapvaal craton boarder (South-eastern terrane of
Witwatersrand block). The Tugela terrane is subdivided into four thrust sheets, which consist
mainly of amphibole-facies gneisses and deformed interlayers of relic ophiolite sequences. The
Lesotho kimberlite cluster, including the Letseng kimberlite, has intruded into Jurassic-age
Karoo (Stormberg) basalts with an emplacement age of around ~90 Ma (Davis, 1977). There
are 8 recognized kimberlite intrusive phases present in the pipe, named K1-K8. Most attention
has been paid to the garnet-rich type K6, since it has proven to have the highest economic
potential. The K6 kimberlite is a late-stage intrusion that cuts all other types, except K5 (Field
et al., 2008).
The Northern Lesotho kimberlites, including the Letseng intrusion, contain relatively
high amounts of lower crustal and mantle xenoliths, such as garnet-bearing gneisses, granulites,
eclogites, granular lherzolites, harzburgites, and dunites. Megacrysts of olivine, enstatite,
bronzite, diopside and garnet are also present. In terms of geochemistry and petrology, most of
the Lesotho kimberlites are dominated by basic cpx+plg+grt±opx granulites and eclogites, with
intermediate to acid granulite xenoliths being less abundant. A distinct correlation occurs
between on- and off-craton crustal xenolith varieties. On-craton kimberlites have sparse
granulite suite xenoliths, while off-craton kimberlites contain rather abundant lower crustal
xenoliths. P-T estimates of granulite suite yield a temperature range of 550-850 0C and
pressures of 5-13 kb (Griffin et al., 1979; Sommer et al., 2013), which correlate very well with
a shallow sampling depth of <50 km.
Of the mantle-derived xenolith suite, garnet-bearing lherzolites and sheared dunites
have been studied in greater detail, since these xenolith populations are more abundant in the
Letseng intrusion (Boyd & Nixon, 1975; Lock & Dawson, 2013). Mantle-derived xenolith
compositions and textures vary between the two Letseng kimberlite pipes. The garnet-bearing
and spinel-bearing peridotite xenoliths originating from shallow depths of <100 km have
dominantly granular textures, compared to the xenoliths with porphyroblstic or mosaic textures
that were sampled from greater depths. Also, the degree of deformation seams to increase with
increasing depth of origin. Coarse-textured xenoliths are more abundant in the Main pipe and
xenoliths with a deformed texture are more abundant in the Satellite Pipe. Despite the textural
differences, garnet lherzolites and other xenolith suites from both Letseng pipes tend to be
depleted in TiO2, Al2O3, FeO and CaO relative to the primitive mantle composition. With an
increasing sampling depth, the degree of depletion decreases.
34
6.3. Letlhakane kimberlite
The Letlhakane Group I kimberlite was discovered in Botswana in 1968. It consists of
two separate pipes named D/K1 (11.6 ha) and D/K2 (3.6 ha), which have intruded into the
Paleoproterozoic (1.8-2.0 Ga) Makondi Mobile Belt volcanics, orthoquartzites and carbonate
sequences (Steifenhofer et al., 1997). The Makondi Mobile Belt is underlain by Archean
subcontinental lithospheric mantle belonging to the Zimbabwean craton. Both kimberlite pipes
are overlain by relatively thin (4 to 10 m) Kalahari Group sand and calcrete sediments. The
presence of Karoo-age xenoliths in both of the pipes suggests that the timing of the Letlhakane
kimberlite emplacement must be post-Karoo and pre-Kalahari. Although the Letlhakane
kimberlites have not been precisely dated, it is estimated that the emplacement age is around
~90 Ma (Davis, 1977), the same as that of Orapa (93 Ma), as the listed intrusions belong to a
larger kimberlite cluster and have common features. The Orapa kimberlite intrusion is located
approximately 40 km north-west of Letlhakane (Steifenhofer et al., 1997; Field et al., 2008).
Both Letlhakane intrusions consist mainly of weathered diatreme facies massive
volcanoclastic kimberlite (Roberts & McKinlay, 2017). There are at least two distinct main
volcanoclastic kimberlite types named VK1 and VK2 in pipe D/K1. Magmatic kimberlite (MK)
is also present in lesser amounts. Pipe D/K2 is dominated by coherent magmatic kimberlite
(MK) and breccias rich in lithic fragments (BBR and CBR). Volcanoclastic rocks are present
in lesser amounts and comprise a complex association with at least four distinct types named
VK1, VK2, VK3 and VK5.
The Letlhakane xenolith suite consists of various peridotite types, pyroxenites,
eclogites, megacrysts, MARID, and glimmerite (Stiefenhofer et al., 1997; Wainwright et al.,
2015). The peridotite xenoliths are dominated by garnet-bearing harzburgites and lherzolites
and spinel-bearing lherzolites. Most of the xenoliths have coarse equant to mosaic
porphyroclastic textures. Similarly to the Lesotho xenoliths, the Letlhakane xenoliths from
greater depths tend to display intensely deformed textures.
Many mantle xenoliths from Letlhakane, especially garnet lherzolites, show evidence
for metasomatic overprint, which is expressed as addition of secondary mineral phases, such as
phlogopite and ilmenite, trace element enrichment and depletion in Ca, Ni and Al (Stiefenhofer
et al., 1997; Achterberg et al., 2001; Wainwright et al., 2015). The spinel-bearing lherzolite
xenoliths differ from the typical Kaapvaal craton spinel assemblages (Boyd, 1989) by having
finer textures and better developed anhedral grains. Geothermobarometric calculations indicate
that the Letlhakane xenoliths were sampled from various lithospheric mantle horizons down to
a depth of 150 km and plot along a continuous 40 mW/m2 continental geotherm gradient.
35
6.4. Frank Smith kimberlite
The Frank Smith Group I kimberlite was discovered in the late 19th century
approximately 80 km north of Kimberley town, South Africa. There are two intrusions that
represent the Frank Smith kimberlite as a whole – the Main pipe and the Weltevreden pipe.
Both are interconnected by a 40-m-wide and 180-m-long dike known as the Windsor Block
(Field et al., 2008). In terms of geological setting, the Frank Smith kimberlite is located in the
central part of the Kimberly block, intruding Archean basement crystalline rocks, Ventersdorp
System rocks, Lower Karoo System sedimentary rocks and a Jurassic-age Karoo dolerite sill.
The presence of the Stormberg sequence xenoliths in the diatreme suggests that the crater facies
and a large part of the Frank Smith upper diatreme facies were eroded away. The emplacement
age of the Group I kimberlite is ~114 Ma (Smith, 1983). Van der Spuy (1984) identified five
distinct volcanoclastic kimberlite varieties and three types of hypabyssal kimberlites in the main
Frank Smith pipe and two varieties of kimberlite in the interconnecting Windsor Block. There
are also contact breccias, large floating reefs and late-stage dikes present in the pipe. Notably,
the Frank Smith kimberlite has intruded into an older micaceous dike swarm.
The Frank Smith kimberlite contains a large variety of crustal and mantle xenoliths:
various eclogite types, garnet-bearing and garnet-free peridotite types, and macrocrysts of
olivine, ilmenite and enstatite grains (Boyd, 1973a; Boyd, 1973b; Boyd & Tsai, 1979; Clarke,
1979; Pasteris et al., 1979; Rawlinson & Dawson, 1979). Olivine macrocrysts display massive,
granular and deformed textures. Based on the xenolith chemistry and texture variety, these
xenoliths may have been derived either from dunites or sheared lherzolites. The Mg number of
olivine macrocrysts ranges from 84.5-86.5 (Boyd, 1973a).
Other ultramafic xenoliths from the Frank Smith kimberlite have close similarities to
the Northern Lesotho ultramafic xenolith suite (Boyd, 1973b). Xenolith samples of the sheared
lherzolite and enstatite megacryst suite tend to have equilibration temperatures of >1100 0C and
a depth of origin at >150 km. The deep-seated enstatite macrocrysts that have equilibration
conditions of 1200 0C and 60 kb contain inclusions of Mg-ilmenite and poliphase garnets.
Mineral compositions of these xenoliths indicate an origin from a highly depleted source
(Mayer & Tsai, 1979).
36
7. Sampling and methods
7.1. Samples and thin section studies
The aim of this study was to examine peridotitic mantle xenoliths from several Group I
kimberlite pipes that were emplaced in the Kaapvaal craton. A total of 27 previously prepared
standard thin sections from 7 different localities containing peridotite xenoliths were evaluated
at the start of the project. After sorting the initial data, a sample set of 24 thin sections from four
relevant on-craton kimberlite localities were chosen for further examination. These localities
are the Group I Premier, Frank Smith, Letseng kimberlites from the Archean terrane and the
Group I Letlhakane kimberlite from the Proterozoic terrane. They cover the most important
tectonic subdivisions of the Kaapvaal craton and its surroundings.
Petrographic descriptions were carried out for the majority of the chosen thin section
samples using a Leica DM750 series binocular optical microscope at the University of Oulu.
Thin sections were photographed using a Leica EC4 digital microscope camera and processed
with the Leica Microsystems microscope imaging software. Suitable places were identified in
thin sections for further mineral chemical analysis of the main mineral phases (olivine,
orthopyroxene, clinopyroxene, garnet, and spinel) by electron microprobe (EPMA).
7.2. Electron microprobe analyses
The major and some minor element compositions of olivine, clinopyroxene,
orthopyroxene, garnet and spinel were analyzed using a JEOL JXA-8200 electron microprobe
at the Center of Microscopy and Nanotechnology (CMNT), University of Oulu. The operating
conditions were an accelerating voltage of 15 kV and a beam current of 40 nA, and ZAF
correction was applied to the analyses. The accuracy of analyses was monitored using reference
material of similar composition. The reproducibility varied by less than 2 %. Obtained mineral
geochemical data were utilized for geothermobarometric calculations to determine the depth of
origin and nature of the studied mantle xenoliths.
37
7.3. Geothermobarometic calculations
The PTEXL3 calculation spreadsheet was used to construct the P-T conditions for
mantle xenolith assemblages examined in this study. The PTEXL3 P-T calculator was
developed by T. Köehler in the middle of the 1990s and later modified by A. Girnis (Brey et
al., 2008). It contains a summary of experimentally or empirically developed 22 thermometer
and 8 barometer reference calculation formulations suitable for a wide range of peridotitic rock
types. The calculation algorithm works in the Microsoft Office Excel environment. Besides a
relatively easy data input interface, PTEXL3 provides a result optimization option and a
possibility to plot P-T data along with the graphite-diamond boundary established by Kennedy
& Kennedy (1976) in a separate Excel spreadsheet. By entering the major oxide data from
olivine, orthopyroxene, clinopyroxene, garnet and spinel mineral assemblages in a blank
spreadsheet, the PTEXL3 program computes P-T conditions, if the mineral chemistry data set
has valid input values. The PTEXL3 spreadsheet is a free application available in the internet.
The geothermobarometry applied for peridotitic systems is based on element exchange
between the site occupancies with similar properties at certain pressure and temperature
conditions. Solid solution relationships also play a crucial role in the exchange in common
pyroxene, olivine, garnet and spinel phases. For natural and simplified MAS (MgO-Al2O3-
SiO2) systems, the Fe and Mg exchange and Al partitioning reactions are applied, but for simple
CMAS (CaO-MgO-Al2O3-SiO2) systems, Ca and Na partitioning is more often used as the
calculation basis for geothermobarometry.
The following thermometers and barometers were applied to peridotitic rocks in this
study:
TBKN90: Orthopyroxene-clinopyroxene solvus thermometer (Brey & Köehler, 1990) for
garnet-bearing four-phase lherzolites
TO’NW79: Fe-Mg exchange between olivine and garnet thermometer (O’Neill & Wood,
1979) for garnet-bearing dunites and harzburgites
TOW87: Fe-Mg exchange between olivine and spinel thermometer (O’Neill & Wall,
1987) for spinel-bearing peridotitic rocks
PBBG08: Al-in-orthopyroxene barometer (Brey et al., 2008) for garnet-bearing peridotites
PNimisTaylor00: Cr-in-clinopyroxene barometer (Nimis & Taylor, 2000) for garnet
peridotites.
38
The inter-mineral equilibrium for pressure and temperature estimates were tested by
comparing results from other independent geothermometers and geobarometers listed in the
PTXL3 spreadsheet. The additional geothermometers and geobarometers used for comparison
are:
TNaPxBK90: Na partitioning between orthopyroxene and clinopyroxene thermometer
(Brey & Köehler, 1990);
TOpxBK90: single pyroxene Ca-in-orthopyroxene thermometer (Brey & Köehler, 1990);
TKB90: Ca partitioning between olivine and orthopyroxene thermometer (Köehler &
Brey, 1990);
TKrogh88: Fe-Mg exchange between clinopyroxene and garnet thermometer (Krogh,
1988);
TNimisTaylor00: enstatite-in-clinopyroxene thermometer (Nimis & Taylor, 2000);
PBKN90: Al-in-orthopyroxene barometer (Brey & Köehler, 1990);
PKB90: Ca partitioning between olivine and clinopyroxene barometer (Köehler & Brey,
1990).
The general principles and formulation of the thermometers and barometers in question
are described further below.
7.4. Geothermometers
The orthopyroxene-clinopyroxene solvus thermometer (TBKN90) was developed by Brey
& Köehler (1990) on the basis of experimental studies performed in the pressures range of 10-
60 kb and temperatures range of 900 to 1400 0C by Brey et al. (1990). The pyroxene solvus
thermometer is based on the mutual solubility of enstatite and diopside components with
coexisting clinopyroxene and orthopyroxene mineral phases. The Mg-Fe exchange related to
the enstatite component can be expressed with the following chemical reaction:
FeSiO3Opx + CaMgSi2O6
Cpx MgSiO3Opx + CaFeSi2O6
Cpx
Utilizing the site occupancy exchange reactions, the following thermometer formulation
can be written (Brey & Köehler, 1990):
𝑇𝐵𝐾𝑁90 =23664 + (24.9 + 126.3𝑋𝐹𝑒
𝐶𝑝𝑥) ∗ 𝑃
13.38 + (ln 𝐾𝐷∗ )2 + 11.59𝑋𝐹𝑒
𝑂𝑝𝑥
39
where 𝐾𝐷∗ = (1 − 𝐶𝑎∗)𝐶𝑝𝑥/(1 − 𝐶𝑎∗)𝑂𝑝𝑥 with the applied Na correction in the M2 site as
𝐶𝑎∗ = 𝐶𝑎𝑀2/(1 − 𝑁𝑎𝑀2) and 𝑋𝐹𝑒𝑝𝑥 = 𝐹𝑒/(𝐹𝑒 + 𝑀𝑔) in pyroxenes. Temperature T is in
Kelvin degrees and pressure P in kilobars.
The Fe-Mg partitioning between garnet and olivine thermometer (TO’NW79) was
developed by O’Neill & Wood (1979) based on the experimental studies performed under 30
kb pressure and 900 to 1400 0C temperature conditions.
The Fe-Mg exchange reaction between coexisting garnet and olivine can be expressed
as follows:
2Mg3Al2Si3O12 + 3Fe2SiO4 2Fe3Al2Si3O12 + 3Mg2SiO4
Garnet prefers to incorporate magnesium over iron if both iron and magnesium are
available. This element exchange preference is temperature dependent and forms the basis for
the TO’NW79 thermometer calculation. Also, the Fe/Mg ratio is strongly dependent on the Ca
content of garnet. The TO’NW79 thermometer can be expressed with the following equation:
𝑇𝑂′𝑁𝑊79 =902 + 𝐷𝑉 + (𝑋𝑀𝑔
𝑂𝑙 − 𝑋𝐹𝑒𝑂𝑙)(498 + 1.51(𝑃 − 30)) − 98(𝑋𝑀𝑔
𝐺𝑡 − 𝑋𝐹𝑒𝐺𝑡) + 1347𝑋𝐶𝑎
𝐺𝑡
𝑙𝑛𝐾𝐷 + 0.357
where: temperature T is in Kelvin degrees and pressure P in kilobars; DV is an integrated form
of equilibrium constant Ka; and KD is the partition coefficient expressed as
KD=(Fe/Mg)Gt/(Fe/Mg)Ol. The TO’NW79 thermometer is in good agreement with two pyroxene
thermometers and provides temperature estimates with a similar precision.
The Mg-Fe exchange between olivine and spinel thermometer (TOW87) was developed
by O’Neill & Wall (1987). The thermometer is based on spinel and olivine Mg-Fe2+ partitioning
from the following exchange reaction:
0.5Fe2SiO4 + MgAl2O4 = 0.5Mg2SiO4 + FeAl2O4
Utilizing the reaction, temperature for olivine-spinel Mg-Fe2+ exchange can be
expressed with the following thermometer equation:
𝑇𝑂𝑊87 =
6530 + 28𝑃 + (5000 + 10.8𝑃)(𝑋𝑀𝑔𝑂𝑙 − 𝑋𝐹𝑒
𝑂𝑙) − 1960(1 + 𝑋𝑇𝑖𝑆𝑝)(𝑋𝑀𝑔
𝑆𝑝 − 𝑋𝐹𝑒2+𝑆𝑝 )
+18620𝑋𝐶𝑟𝑆𝑝 + 25150(𝑋𝐹𝑒3+
𝑆𝑝 − 𝑋𝑇𝑖𝑆𝑝)
𝑅𝑙𝑛𝐾𝐷 + 4.705
where: temperature T is in Kelvin degrees and pressure P in kilobars; constant R is 8.31441
(JK-1mol-1); and KD is the Mg-Fe partition coefficient. The thermometer is sensitive on the
change in the Fe3+ content and thus may produce temperatures estimates with a higher
uncertainty that the other discussed thermometers in this study.
40
Another two-pyroxene thermometer (TNaPxBK90) developed by Brey & Köehler (1990)
and used for natural peridotitic systems is based on the partitioning of Na between
orthopyroxene and clinopyroxene. The barometer can be formulated as:
𝑇𝑁𝑎𝑃𝑥𝐵𝐾90 =35000 + 61.5 ∗ 𝑃
(𝑙𝑛𝐷𝑁𝑎)2 + 19.8
where DNa=NaOpx/NaCpx, temperature T is in Kelvin degrees and pressure P in kilobars. The Na-
in-pyroxene thermometer has a very low precision of ±56 0C (1σ) and thus the use of this
thermometer must be considered with caution.
The single pyroxene Ca-in-orthopyroxene thermometer (TOpxBK90) (Brey & Köehler,
1990) is based on the diopside solubility in orthopyroxene. The Ca content in diopside is largely
controlled by pressure and temperature conditions. Notably, the Ca content of orthopyroxene is
lowered by an increase in Al and tends to increase in the presence of Fe. The Ca-in-
orthopyroxene thermometer can be expressed as:
𝑇𝑂𝑝𝑥𝐵𝐾90 =6425 + 26.4 ∗ 𝑃
−ln 𝐶𝑎𝑂𝑝𝑥 + 1.843
where temperature T is expressed in Kelvin degrees and pressure P in kilobars. The Ca-in-
orthopyroxene thermometer has a slightly lower precision compared to TBKN. Similarly to the
Na-in-pyroxene barometer, the Ca-in-orthopyroxene barometer formulation is an extension of
TBKN90, developed based on experimental results from Brey et al. (1990).
The Fe2+-Mg exchange between garnet and clinopyroxene thermometer (TKrogh88) was
developed and calibrated by Krogh (1988) using the data set obtained from various laboratory
experiments on peridotitic assemblages at 600-1300 0C and 20-40 kb (Råheim & Green, 1974;
Mori & Green, 1978; Ellis & Green, 1979). The distribution of Fe2+ and Mg between coexisting
garnet and clinopyroxene phases can be expresses as the following exchange reaction:
1/3Mg3Al2Si3O12Pyr + CaFeSi2O6
Hed 1/3Fe3Al2Si3O12
Alm + CaFeSi2O6Di
Large part of the partitioning effect of Fe2+ and Mg between garnet and clinopyroxene
is influenced by the grossular content in garnet. By utilizing the exchange reaction and
temperature dependence of the Fe-Mg exchange, the TKrogh88 thermometer can be formulated as
follows:
𝑇(𝐾𝑟𝑜𝑔ℎ88) = (−6173(𝑋𝐶𝑎)2 + 6731𝑋𝐶𝑎 + 1879 + 10𝑃
𝑙𝑛𝐾𝑑 + 1.393) − 273
where temperature T is in Kelvin degrees and pressure P in kilobars and Kd is the Mg-Fe
distribution coefficient.
The enstatite-in-clinopyroxene thermometer (TNimisTaylor00) was developed by Nimis &
Taylor (2000) on the basis of experimentally synthesized clinopyroxenes at 850-1500 0C and
41
0-60 kb in the CMS and CMAS-Cr systems. The thermometer was also calibrated for more
complex peridotite systems. The single-pyroxene thermometer is based on the enstatite
component exchange reaction between orthopyroxene and clinopyroxene, which can be
expressed as follows: Mg2Si2O6Opx Mg2Si2O6
Cpx. In natural peridotitic systems, the enstatite
component activity in orthopyroxene is close to unity and rather insensitive to temperature or
compositional variations. Hence, the temperature dependence of the exchange reaction can be
rearranged for clinopyroxene alone and written as the following thermometer formulation:
𝑇𝑁𝑖𝑚𝑖𝑠𝑇𝑎𝑦𝑙𝑜𝑟00 =23166 + 39.28 ∗ 𝑃
13.25 + 15.35𝑇𝑖 + 4.5𝐹𝑒 − 1.55(𝐴𝑙 + 𝐶𝑟 − 𝑁𝑎) + (𝑙𝑛𝑎𝑒𝑛𝐶𝑝𝑥
)2
where temperature T is in Kelvin degrees and pressure P is in kilobars and the enstatite
component activity expressed as 𝑎𝑒𝑛𝐶𝑝𝑥 = (1 − 𝐶𝑎 − 𝑁𝑎 − 𝐾) ∗ (1 −
1
2(𝐴𝑙 + 𝐶𝑟 + 𝑁𝑎 + 𝐾)).
The precision of thermometer is ±30 0C (1σ).
7.5. Geobarometers
The concept of the thermodynamic calculations for the Al-in-orthopyroxene barometer
(PKBN90 and PBBG08) in simple MAS systems was largely outlined by Gasparik & Newton (1984)
and later adopted for peridotitic rock assemblages by Brey & Köehler (1990) with acceptable
and reliable results up to 60 kb pressure. The barometer was further corrected for pressures
higher than 60 kb based on extended experimental results obtained by Brey et al. (2008). The
Al-in-orthopyroxene barometer is based on the maximum Al solubility in orthopyroxene
coexisting garnet and can be expressed with the following chemical reaction:
(Mg2Si2O6+MgAl2SiO6)Opx
Mg3Al2Si3O12Grt
The PKBN90 barometer can be formulated as follows:
𝑃𝐾𝐵𝑁90 =−𝐶2 − √𝐶2
2 + 4𝐶3𝐶1 ∕ 1000
2𝐶3
𝐶1 = −𝑅𝑇𝑙𝑛𝐾𝐷 + 5510 + 88.91𝑇 − 19𝑇1.2 + 3(𝑋𝐶𝑎𝐺𝑟𝑡)2 ∗ 82458 + 𝑋𝑀𝑔
𝑀1 ∗ 𝑋𝐹𝑒𝑀1
∗ (80942 − 46.7𝑇) − 3𝑋𝐹𝑒𝐺𝑟𝑡 ∗ 𝑋𝐶𝑎
𝐺𝑟𝑡 ∗ 17793 − 𝑋𝐶𝑎𝐺𝑟𝑡 ∗ 𝑋𝐶𝑟
𝐺𝑟𝑡
∗ (1.164 ∗ 106 − 420.4𝑇) − 𝑋𝐹𝑒𝐺𝑟𝑡 ∗ 𝑋𝐶𝑟
𝐺𝑟𝑡 ∗ (−1.25 ∗ 106 − 565𝑇)
𝐶2 = −0.832 − 8.78 ∗ 10−5 ∗ (𝑇 − 298) + 3(𝑋𝐶𝑎𝐺𝑟𝑡)2 ∗ 3.305 − 𝑋𝐶𝑎
𝐺𝑟𝑡 ∗ 𝑋𝐶𝑟𝐺𝑟𝑡
∗ 13.45 + 𝑋𝐹𝑒𝐺𝑟𝑡 ∗ 𝑋𝐶𝑟
𝐺𝑟𝑡 ∗ 10.5
𝐶3 = −16.6 ∗ 10−4
42
𝐾 =(1 − 𝑋𝐶𝑎
𝐺𝑟𝑡)3 ∗ (𝑋𝐴𝑙𝐺𝑟𝑡)2
𝑋𝑀𝐹𝑀1 ∗ 𝑋𝑀𝐹
𝑀2 ∗ 𝑋𝑀𝐹𝑀2 ∗ 𝑋𝐴𝑙,𝑇𝑠
𝑀1
R = 8.3143 J/Kmol; T is in Kelvin degrees; P is in kilobars.
The formulation of the modified PBBG08 in the following (Brey et al, 2008):
𝑃𝐵𝐵𝐺08 =−𝐶2 − √𝐶2
2 + 0.004𝐶3𝐶1
20𝐶3
𝐶1 = −𝑅𝑇𝑙𝑛𝐾 + 4970 + 84.15𝑇 − 19𝑇1.2 + 3(𝑋𝐶𝑎𝐺𝑟𝑡)2 ∗ 82456 + 𝑋𝑀𝑔
𝑀1 ∗ 𝑋𝐹𝑒𝑀1
∗ (80941 − 36.3𝑇) − 3𝑋𝐹𝑒𝐺𝑟𝑡 ∗ 𝑋𝐶𝑎
𝐺𝑟𝑡 ∗ 17795 − 𝑋𝐶𝑎𝐺𝑟𝑡 ∗ 𝑋𝐶𝑟
𝐺𝑟𝑡
∗ (1.164 ∗ 106 − 335𝑇) − 𝑋𝐹𝑒𝐺𝑟𝑡 ∗ 𝑋𝐶𝑟
𝐺𝑟𝑡 ∗ (−1.25 ∗ 106 − 730𝑇)
𝐶2 = −0.533 − 1.62 ∗ 10−4 ∗ (𝑇 − 298) + 3(𝑋𝐶𝑎𝐺𝑟𝑡)2 ∗ 3.305 − 18.25𝑋𝐶𝑎
𝐺𝑟𝑡 ∗ 𝑋𝐶𝑟𝐺𝑟𝑡
+ 3.5𝑋𝐹𝑒𝐺𝑟𝑡 ∗ 𝑋𝐶𝑟
𝐺𝑟𝑡
𝐶3 = −7.2 ∗ 10−4
𝐾 =(1 − 𝑋𝐶𝑎
𝐺𝑟𝑡)3 ∗ (𝑋𝐴𝑙𝐺𝑟𝑡)2
𝑋𝑀𝐹𝑀1 ∗ 𝑋𝑀𝐹
𝑀2 ∗ 𝑋𝑀𝐹𝑀2 ∗ 𝑋𝐴𝑙,𝑇𝑠
𝑀1
R = 8.3143 J/Kmol; T is in Kelvin degrees and P in GPa.
The precision of the PBBG barometer is ±0.3 GPa (1σ).
The Cr-in-clinopyroxene barometer (PNimisTaylor00) was developed by Nimis & Taylor
(2000) on the basis of experimentally synthesized clinopyroxenes at 850-1500 0C and 0-60 kb
in the CMS and CMAS-Cr systems. The barometer was also calibrated for more complex
peridotite systems. The exchange of the Cr component between clinopyroxene and garnet can
be described with the following chemical reaction:
CaMgSi2O6di + CaCrAlSiO6
CaCrTs 0.5(Ca2Mg)Cr2Si3O12uv2kn1 + 0.5(Ca2Mg)Al2Si3O12
gr2py1
where the diopside (di) and CaCr-Tschermak’s (CaCrTs) components from clinopyroxene react
with the uvarovite (uv), grossular (gr), pyrope (py) and knorringite (kn) components in garnet.
The Cr exchange in the reaction is pressure dependent and can be formulated as a function of
temperature as a barometer as follows:
𝑃𝑁𝑖𝑚𝑖𝑠𝑇𝑎𝑦𝑙𝑜𝑟00 = −𝑇
126.9∗ ln[𝑎𝐶𝑎𝐶𝑟𝑇𝑠
𝐶𝑝𝑥 ] + 15.483 ∗ ln (𝐶𝑟#𝐶𝑝𝑥
𝑇) +
𝑇
71.38+ 107.8
where: temperature T is in Kelvin degrees and pressure P is in kilobars; CaCr-Tschermak
component’s activity in clinopyroxene is 𝑎𝐶𝑎𝐶𝑟𝑇𝑠𝐶𝑝𝑥 = 𝐶𝑟 − 0.81 ∗ 𝐶𝑟# ∗ (𝑁𝑎 + 𝐾) and Cr#
number is expressed as Cr#=Cr/(Cr+Al) with elements in atoms per 6 oxygen. The Cr exchange
barometer has a precision of ±2.3 kb (1σ) and has a temperature dependence of 1.2-2.4 kb/50
0C.
43
The Ca-in-olivine geothermobarometers TKB90 and PKB90 by Köehler & Brey (1990) are
based on the Ca concentration in olivine coexisting with clinopyroxene. The Ca exchange
between olivine and clinopyroxene is pressure controlled and can be expressed with the
following chemical reaction:
Mg2SiO4Ol + CaMgSi2O6
Cpx = CaMgSiO4Ol + Mg2Si2O6
Cpx
The experimental results from Köehler & Brey (1990) determined a non-linear
relationship for the Ca solubility in olivine at temperatures greater than 1100 0C. Therefore, the
experimental data were fitted into two separate thermobarometric equations for low- and high-
temperature assemblages:
𝑃𝐾𝐵90 = (−𝑇 ∗ 𝑙𝑛𝐷𝐶𝑎 − 11982 + 3.61 ∗ 𝑇) ∕ 56.2
for T ≥ (1275.25 + 2.827 * P)
𝑃𝐾𝐵90 = (−𝑇 ∗ 𝑙𝑛𝐷𝐶𝑎 − 5792 + 1.25 ∗ 𝑇) ∕ 42.5
for T ≤ (1275.25 + 2.827 * P)
where Dca = CaOl/CaCpx, temperature T is in Kelvin degrees and pressure P in kilobars. The
barometer has a precision of ±1.7 kb (1σ). Notably, the Ca-in-olivine barometer can be applied
for spinel peridotites.
44
8. Results
8.1. Petrography
Approximate modal mineral abundances for studied thin-section samples were
determined using a binocular petrographic microscope, thus the obtained precision is moderate.
With the exception of one sample (P6) from the Premier kimberlite, which was identified as an
olivine websterite, the remaining 23 xenolith samples belong to peridotites, of which 14 are
garnet-bearing peridotites, 4 spinel-bearing peridotites and 5 garnet-free peridotites (Fig. 10).
A summary of petrographic descriptions is listed in Table 2.
Figure 10. Classification scheme for peridotites with rough modal proportions of the studied xenolith
samples from the Letlhakane, Premier, Frank Smith and Letseng kimberlites. Diagram taken from Gill
(2010).
The definite majority of the studied xenolith samples show coarse-grained tabular or
equant textures. Two samples, 512 and 551, from the Letlhakane kimberlite have a medium- to
coarse-grained texture. Samples Frank Smith XS2 and Premier P18 are the only ones having a
mosaic-porphyroclastic texture (Fig. 11-12). Large parts of the samples are deformed or
disintegrated making it difficult to determine their original texture. In some cases, the term
relict was used to specify the rock texture.
45
Figure 11. Photomicrographs of major rock types and textures of studied mantle xenoliths. (A)
Fractured and deformed olivine in a coarse-grained garnet-bearing lherzolite from the Letlhakane
kimberlite. Garnet in bottom right is surrounded by secondary clinopyroxene and micaceous minerals.
Sample 550; (B) Highly serpentinized, coarse-grained, tabular garnet-bearing lherzolite from the
Letlhakane kimberlite. Fine spinel can be spotted along a thick garnet alteration rim. Sample 523; (C)
Mosaic-porphyroclastic texture in in garnet-bearing lherzolite from the Premier kimberlite. Major
porphyroclasts are surrounded by a thick reaction rim. Sample P18; (D) Disintegrated and serpentinized,
coarse equant olivine and altered garnet with a thick surrounding kelyphite rim in a garnet-bearing
harzburgite xenolith from the Premier kimberlite. Sample P15; (E) Spinel assemblage in garnet-bearing
harzburgite that surrounds partly depleted garnet from the Letseng kimberlite. Rest of the field is filled
with coarse, fractured olivine and medium-grained pyroxene grains. Sample LET4; (F) Typical coarse-
grained garnet-bearing harzburgite from the Letseng kimberlite. It contains coarse, tabular, fractured
and partly deformed olivine and tabular pyroxenes. Sample LET14. All photos taken with crossed polars.
46
Figure 12. Photomicrographs of major rock types and textures of studied mantle xenoliths. (A) Typical
coarse-grained, tabular texture for garnet-bearing harzburgite kimberlite xenoliths from the Frank Smith
kimberlite, containing rounded garnets with a well-developed reaction rim. Sample XS5; (B) Deep-
seated garnet-bearing harzburgite from the Frank Smith kimberlite. The xenolith has a typical mosaic-
porphyroclastic texture. Sample XS2; (C) Highly serpentinized, refractory garnet-bearing dunite from
the Frank Smith kimberlite. The xenolith has a relict coarse-grained equant texture. Sample XS7; (D)
Metasomatized olivine websterite from the Premier kimberlite with a medium- to coarse-grained equant
texture. Part of xenolith contains altered yellow to greenish olivine grains. Sample P6; (E) Interstitial
spinel assemblage from a Premier spinel-harzburgite xenolith. Sample P5; (F) Medium- to coarse-
grained spinel-bearing harzburgite from the Letlhakane kimberlite. Sample 551. All photos taken with
crossed polars.
47
Table 2. Summary of the modal and petrographic characteristics of the studied mantle xenoliths from
the Letlhakane, Letseng, Premier and Frank Smith kimberlites.
8.1.1. Letlhakane
The Letlhakane xenoliths plot in the harzburgite and lherzolite fields (Fig. 10) and have
the most diverse xenolith populations among the samples. Two of the studied xenoliths are
spinel-bearing harzburgites and one is a garnet-bearing harzburgite. There are three garnet-
bearing lherzolites, three harzburgites and one lherzolite.
Sample 502 is a garnet-bearing harzburgite with a coarse-grained tabular texture. The
xenolith is deformed and lightly serpentinized, with the interstitial matrix being partly filled
Ol Cpx Opx Gt Sp
Letlhakane
502 68 2 25 5 Gt-Harzburgite Coarse tabular
511 70 30 Harzburgite Medium to coarse equant
512 85 1 14 Harzburgite Medium to coarse tabular
513 60 5 30 5 Gt-Lherzolite Coarse equant
523 60 5 30 5 Gt-Lherzolite Coarse equant
529 70 7 23 Lherzolite Coarse tabular
530 73 2 25 Harzburgite Coarse tabular
531 80 1 18 1 Sp-Harzburgite Coarse tabular
550 63 10 17 10 Gt-Lherzolite Coarse tabular
551 72 1 25 2 Sp-Harzburgite Medium to coarse equant
Letseng
LET4 69 1 25 5 Gt-Harzburgite Coarse equant
LET8 65 30 5 Gt-Harzburgite Coarse equant
LET14 65 2 31 2 Gt-Harzburgite Coarse equant
Premier
P5 75 2 23 1 Sp-Harzburgite Coarse tabular
P6 30 20 50 Ol-websterite Coarse equant
P14 80 20 1 Sp-Harzburgite Coarse tabular
P15 75 2 20 3 Gt-Harzburgite Coarse tabular
P17B 85 15 Harzburgite Coarse tabular
P18 58 7 20 15 Gt-Lherzolite Mosaic-porphyroclastic
Frank Smith
XS1 80 3 17 Gt-Dunite Coarse equant
XS2 65 1 24 10 Gt-Harzburgite Mosaic-porphyroclastic
XS4 74 1 20 5 Gt-Harzburgite Coarse tabular
XS5 60 30 10 Gt-Harzburgite Coarse tabular
XS7 77 3 20 Gt-Dunite Coarse tabular
SampleModal abundance
Type Texture
48
with fine-grained olivine neoblasts. Olivine occurs as deformed and fractured tabular grains
ranging in size from 1 to 5 mm. Overall fracture planes have a parallel alignment. Some olivine
crystal rims are partly disintegrated. Garnet grains are extremely small with an average size of
0.5 mm. Most of the garnet grains lack any straight crystal faces, are partly disintegrated and
surrounded by a thick kelyphite reaction rim. Orthopyroxene dominantly occurs as tabular gains
ranging in size between 1 to 5 mm. Notably, orthopyroxenes <1 mm in size have disintegrated
and altered crystal rims. Clinopyroxene is sparse and occurs as small anhedral grains.
Sample 511 is an altered, medium- to coarse-grained equant harzburgite. Olivine is
moderately deformed, fractionated and occurs as anhedral equant grains. The grain size varies
between 0.5 and 4 mm, depending on the degree of disintegration. Garnet is absent in this
particular sample. Orthopyroxene is equant and ranges in size from 0.5 to 2 mm.
Sample 512 is a moderately serpentinized and deformed harzburgite. It is characterized
by a medium- to coarse-grained tabular texture. Olivine occurs as highly fractionated and
deformed grains ranging in size from 0.25 to 1 mm. Intensely fractured zones are filled with
cross-cutting serpentine veins. Most of the olivine grains exhibit semi-parallel fractures.
Orthopyroxene occurs mainly as relatively small, elongated or tabular grains. Minor fine-
grained (~0.1 mm), irregular clinopyroxene grains are also present.
Xenolith sample 513 is highly serpentinized garnet-bearing lherzolite with a relict
coarse-grained equant texture. Most of olivine is altered, fractured and disintegrated into a fine
mass of partly rounded or tabular grains, especially at contacts with orthopyroxene. Veins of
serpentine mark approximate olivine grain boundaries before serpentinization took place.
Originally, olivines formed coarse 2- to 6-mm-sized crystals. Garnet is rare and can be found
as irregular fractured grains with thick kelyphite reaction rims. The average grain size is 2 mm.
Orthopyroxene forms larger clumps consisting of irregular and tabular grains, but
clinopyroxene is sparse.
Sample 523 represents a strongly serpentinized garnet-bearing lherzolite xenolith with
a coarse tabular texture. Olivine occurs as highly fractured and serpentinized grains ranging in
size from 1 to 7 mm. Garnet can be found as fractured subhedral grains, surrounded by a
kelyphite reaction rim up to 2 mm in thickness. Orthopyroxene is characterized by coarse-
grained (4 to 10 mm), dominantly tabular and elongated grains. Clinopyroxene is sparse and
occurs as fractured irregular grains.
Sample 529 represents a coarse-grained tabular lherzolite. Olivine occurs as coarse, 3
to 10 mm-sized, strongly fractured and partly deformed grains. Some of the olivine grains are
cross-cut by serpentine veins, which have also disintegrated olivine at the contact.
Orthopyroxene can be found as rather coarse, elongated grains up to 7 mm in length or smaller
49
tabular grains ranging in size between 2 and 4 mm. Both varieties are moderately fractured and
cross-cut by serpentine veins in a similar manner as in the case of olivine. Clinopyroxene occurs
as minor scattered and irregular grains up to 2 mm in size.
Strongly deformed harzburgite sample 530 has a coarse-grained tabular texture. Olivine
in this sample is intensely deformed and fractured till a point where initial crystal habits and
sizes are hard to distinguish. Coarse grains are up to 12 mm across and have completely
disintegrated crystal rims, thus coarse grains seem rounded. Smaller olivine grains ranging in
size between 2 and 7 mm can be partly recognized trough a serpentine vein network.
Orthopyroxenes are coarse, up to 5 mm across, and mostly occur as tabular highly fractured
grains with disintegrated crystal rims. Minor partly altered clinopyroxene also can be spotted
among the major mineral phases.
Sample 531 represents a moderately serpentinized and deformed spinel-bearing
harzburgite with a coarse-grained tabular texture. Olivine is intensely deformed, fractured and
veined by serpentine. Fracture planes have a parallel alignment. Due to intense disintegration,
most of the olivine grains lack recognizable original grain boundaries. The largest grains are up
to 8 mm across. Most of the orthopyroxene grains are also fractured and range in size between
0.5 and 5 mm. Seemingly, elongated orthopyroxene grains are aligned sub-perpendicular with
olivine fracture planes. Minor spinel and clinopyroxene can be found together with olivine.
Moderately serpentinized xenolith sample 550 is a coarse-grained tabular garnet-
bearing deformed lherzolite. Olivine occurs as intensely fractured and serpentine-veined,
coarse, tabular grains ranging in size between 1 and 6 mm. Majority of olivine grains are
disintegrated and lack recognizable original grain boundaries. The orthopyroxene grains are
intensely fractured and partly disintegrated, particularly at grain boundaries. The grain size
varies from 2 to 4 mm. Garnet similarly to other major phases is strongly fractured. Anhedral
grains are coarse up to 7 mm across and have highly irregular grain boundaries. Garnet grains
are surrounded by two generations of alteration rims. The outer rim (~1 mm wide) consists of
medium- to fine-grained clinopyroxene and lesser amphibole. The inner alteration rim consists
of thin, ~0.5-mm-wide kelyphite, which further extends inwards along fracture planes.
Clinopyroxene can be found around garnet grains, as described earlier, or as tabular, highly
fractured grains ranging in size between 0.5 and 3 mm.
The moderately serpentinized xenolith sample 551 is a spinel-bearing harzburgite with
a medium- to coarse-grained equant texture. Olivine occurs as disintegrated equant grains
ranging in size between 0.25 to 1.5 mm, but rare coarser grains are also present with a size up
to 4 mm. Orthopyroxene can be dominantly found as fractured tabular grains, ranging in size
50
between 1 and 4 mm. Medium-grained spinel is mostly in contact with orthopyroxenes.
Clinopyroxene is rare.
8.1.2. Letseng
All three Letseng xenolith samples are garnet-bearing harzburgites (Fig. 10). The
garnet-bearing harzburgite sample LET4 has a coarse-grained equant texture. Olivine is
intensely fractured rather than disintegrated or deformed and occurs as dominantly coarse, 0.5-
to 4-mm-sized, tabular grains with barely recognizable grain boundaries. The degree of
deformation or disintegration of olivine increases towards crystal rims, especially at the contact
plane with garnet. Large parts of disrupted garnet grains are partly or fully altered and replaced
by fine disseminated spinel, kelyphite or amphibole minerals. Garnet cores that are still intact
range in size between 0.1 to 1.5 mm. Remnants of original subhedral garnet crystal habits 2 to
4 mm in size are still visible. Orthopyroxene is rather abundant (~25%) and mostly occurs as
tabular, subhedral or anhedral grains varying in size from 0.5 to 5 mm. Similarly to olivine,
orthopyroxene is intensely fractured. Clinopyroxene occurs as rare small, tabular grains.
Sample LET14 is a garnet-bearing harzburgite with a coarse-grained equant texture.
Olivine and orthopyroxene dominate with relative proportions of 65% and 30%, respectively,
with the rest 5% of the volume being occupied by garnet and clinopyroxene. Olivine is intensely
fractured and partly deformed. Rare coarse grains in size from 1 to 4 mm can be recognized;
otherwise most of olivine occurs as fine, disintegrated, anhedral grains. Orthopyroxene mainly
occurs as tabular 0.5- to 3-mm-sized grains. Garnet is rare, occurring as highly fractured,
anhedral grains with an average size of 2 mm and being partly replaced by fine spinel and
surrounded by an up to 0.5 mm wide kelyphite rim. Clinopyroxene can be found as small,
irregular grains either around garnet grains or other alteration minerals.
Sample LET8 is a garnet-bearing harzburgite with a coarse-grained equant texture with
similar characteristics to those of the rest of the Letseng xenolith samples.
51
8.1.3. Premier
The Premier xenoliths contain two spinel-bearing harzburgites; garnet-bearing
harzburgite and lherzolite; garnet-free harzburgite; and one olivine websterite (Fig. 10).
Sample P5 represents an altered coarse-grained tabular spinel-bearing harzburgite
xenolith. Olivine can be found as strongly fractured, equant crystals ranging in size between 1
and 4 mm. Orthopyroxene dominantly occurs as elongated, irregular grains up to 5 mm long
and 2 mm wide. Clinopyroxene is highly altered and partly replaced by secondary minerals.
Sample P6 is a medium- to coarse-grained equant olivine websterite that has been
enriched in Fe due to metasomatism that is reflected in the mineral compositions of the sample.
Two types of olivine can be found. The first one is tabular with an average grain size of 1.5 mm
and abundance of ~30%. The second type is highly altered and has a yellowish color. The
sample contains high amounts of orthopyroxene (~50%), especially ferrosilite. The rest of the
mantle xenolith volume is occupied by ~20% of clinopyroxene and minor phlogopite.
The spinel-bearing harzburgite sample P14 has a coarse-grained tabular texture. Olivine
occurs as highly disintegrated, 0.1- to 0.5-mm-sized grain parts, which were originally part of
larger grains up to 6 mm in size. The main minerals are set in a veined calcite and minor
serpentine matrix. Orthopyroxene ranges in size between 2 to 7 mm and occurs as irregular or
tabular grains with partly disintegrated rims. Garnet is extremely rare and completely replaced
by minor spinel or other alteration minerals.
The garnet-bearing harzburgite sample P15 is highly altered and has a relict coarse-
grained tabular texture. Only rare individual tabular olivine grains can be recognized; otherwise
olivine is completely fractured and disintegrated and appears as a sparse grain mass enclosed
in a calcite and minor serpentine matrix. Garnets in sample P15 can be found as very small,
0.1- to 0.5-mm-sized, altered grains surrounded by a thick (~0.5 mm) reaction rim consisting
of kelyphite, amphibole and phlogopite. Reaction rims around garnets show up to three distinct
zones. Orthopyroxene dominantly occurs as elongated, irregular grains 0.5 to 6 mm long and
up to 2 mm wide with no particular orientation. Clinopyroxene is rare.
Sample P17B represents a deformed harzburgite, which has a coarse-grained tabular
texture. Olivine grains are strongly fractured, range in size between 2 and 8 mm and have rather
clear and rounded grain boundaries. The interstitial matrix is dominantly filled with calcite and
minor serpentine. Orthopyroxene occurs as coarse, tabular grains up to 4 mm across. Garnet is
completely replaced by fine spinel and other alteration minerals.
Sample P18 is a garnet-bearing lherzolite with a distinct mosaic-porphyroclastic texture.
Dominantly tabular and minor equant olivine grains with an average size of 0.1 mm fill the
52
fine-grained interstitial matrix. Notably, matrix olivine forms irregular patches that consist of
either relatively fine or coarse olivines. Some relict altered, coarse (~3 mm), disintegrated
olivine grains are also present. Garnet is coarse and ranges in size from 4 to 7 mm. Most of the
garnet grains are anhedral, as only rare grains show signs of a subhedral crystal habit. Garnet is
moderately fractured and has two generations of rather similar 0.1- to 0.3-mm-wide reaction
rims consisting of an outer coarser pyroxene-dominated layer and inner finer kelyphite layer.
Similarly to the garnet grains, the orthopyroxene grains have developed two generations of
outer reaction rims. Most of orthopyroxene occurs as aligned, tabular or irregular grains ranging
in length between 2 and 7 mm.
8.1.4. Frank Smith
The Frank Smith peridotite xenolith samples plot in the harzburgite-dunite fields with
varying olivine-orthopyroxene proportions and rather constant and low amounts of
clinopyroxene. Two of the samples are garnet-bearing dunites and three were identified as
garnet-bearing harzburgites (Fig. 10).
The garnet-bearing dunite sample XS1 is extensively serpentinized and has a relict
coarse-grained equant texture. Olivine is completely fractured and disintegrated into fine
anhedral pieces, although rare remnants of fractured larger grain clumps up to 5 mm across can
be still recognized. Fine secondary olivine and serpentine form the mineral interstitial matrix
taking up ~40% of the total xenolith volume. Large part of the garnet grains have subhedral
crystal habits, are moderately fractured and surrounded by a very fine kelyphite reaction rim.
The grain size varies between 0.5 and 5 mm. The rest of the garnet population forms irregular,
highly deformed and partly disintegrated clumps with similar outer reaction rims. The
orthopyroxene grains that are found in the sample are tabular with curved grain boundaries.
Sample XS2 represents a moderately serpentinized garnet-bearing harzburgite xenolith
and has a mosaic-porphyroclastic texture. The fine-grained matrix, which dominated by
dominantly tabular olivine grains, takes up ~60% of the total xenolith volume. Olivine grains
show week signs of general alignment with other minerals. Garnet occurs as rounded to
subhedral, moderately fractured grains ranging from 0.5 to 3 mm in size. Most of the garnet
grains are surrounded by a very thin, 10- to 25-μm-wide kelyphite rim. Orthopyroxene
dominantly occurs as tabular or irregular, oriented coarse grains, ranging in size from 0.75 to 6
mm. Similarly to the garnet grains, the orthopyroxenes grains are surrounded by a thin reaction
rim. Clinopyroxene is rare.
53
The moderately serpentinized garnet-bearing harzburgite sample XS4 has a coarse-
grained tabular texture. Olivine occurs as highly fractured and partly disintegrated grains with
barely recognizable grain boundaries. The olivine grain size varies between 1 and 4 mm. Garnet
occurs as relatively small, 0.25- to 2-mm-seized and mostly rounded grains, compared to the
garnet in the other Frank Smith xenolith samples. The presence of a narrow kelyphite reaction
rim around garnet grains is a characteristic feature. Orthopyroxene mainly occurs as irregular,
0.5- to 4-mm-sized grains with partly rounded crystal edges and a surrounding narrow reaction
rim. Clinopyroxene is rare and occurs as very small (0.5 mm), anhedral grains.
Sample XS5 is a moderately serpentinized, coarse-grained tabular garnet-bearing
harzburgite. Olivine occurs as large, tabular fractured grains 3 to 7 mm in size, which are veined
with serpentine. At the contact with orthopyroxene grains, olivine is highly deformed,
disintegrated and serpentinized. The interstadial mineral matrix is filled by calcite. Garnet
ranges in size between 1 to 6 mm and shows subhedral or completely rounded crystal habits.
Most of the grains have well preserved cores, and the rest of the mineral grains are fractured
and disintegrated at the rim. The amount of fractures seems to progressively increase from the
inner core outwards. Garnet grains are surrounded by a kelyphite reaction rim, which spreads
towards the core. Orthopyroxene occurs as coarse, tabular grains 2 to 9 mm in size with highly
disintegrated crystal rims.
Sample XS7 is an intensely serpentinized garnet-bearing dunite with a coarse-grained
tabular texture. Olivine occurs as altered remnants of fractured 2- to 10-mm-sized, dominantly
tabular, coarse grains with highly disintegrated crystal rims. The interstitial matrix is filled with
serpentine and fine neoblastic olivine. Garnet mostly is rounded, fractured and enclosed in a
thin kelyphite rim, though tabular variants are also present. The grain size is rather uniform and
averages 2 mm. Some garnets form larger clumps. Orthopyroxene is rare.
8.2. Mineral major element chemistry
The dominant mineral assemblages that were analyzed for major element chemistry
from the mantle xenoliths suite in this work are listed in Table 3. A full major element data set
is listed in Appendix I. Geochemical characteristics of each mineral are described and discussed
further below.
54
Table 3. Mineral phases analyzed for major element chemistry from each sample.
8.2.1. Olivine
All 24 mantle xenolith thin-sections contain abundant olivine grains (Table 3). Overall,
the Mg# values in olivine range from 88.8 to 93.6 with an average of 92.1, which are in a good
argument with olivine compositions in depleted cratonic mantle xenoliths found from the
Kaapvaal craton (Bernstein et al., 2007; Griffin et al., 2003; Pearson et al., 2003). However,
major differences in Mg# have been noted between observed localities and may represent
various sources.
Olivines from the Frank Smith xenolith suite show variable Mg# values with extremes
of 88.8 in garnet-bearing dunite and 93.6 in garnet-bearing harzburgite, but at the same time,
Ol Cpx Opx Gt Sp
Letlhakane
502 X X X X
511 X X
512 X X X
513 X X X X
523 X X X X
529 X X X
530 X X X
531 X X X X
550 X X X X
551 X X X X
Letseng
LET4 X X X X
LET8 X X X
LET14 X X X
Premier
P5 X X X X
P6 X X X
P14 X X X
P15 X X X X
P17B X X
P18 X X X X
Frank Smith
XS1 X X
XS2 X X X X
XS4 X X X X
XS5 X X X
XS7 X X
Mineral phases analyzedSample
55
has the lowest average Mg# value of 90.6, relative to other localities. On the other hand, olivines
from the Letseng kimberlite have the highest and uniform Mg# values with an average of 93.4.
The Letlhakane and Premier olivines have similar and rather uniform Mg# compositions with
corresponding averages of 92.1 and 92.7.
For minor elements in olivine, the following average concentrations (wt%) have been
measured: 0.025% Na2O; 0.03% CaO; 0.37% NiO; 0.01% K2O; 0.1% MnO; 0.02% TiO2;
0.01% Al2O3; 0.025% Cr2O3; and 0.01% V2O3. The NiO concentrations in olivine fall in a range
of 0.27-0.43 wt%, being approximately the same in each locality, and the MnO concentrations
vary in a range of 0.06-0.15%. NiO shows a weak positive and MnO a weak negative correlation
with Fo, with the trends indicating progressive melt depletion in the source mantle (Fig. 13).
Figure 13. (A) NiO and Mg# (Fo) relationship in olivine with a week positive trend; (B) MnO and Mg#
(Fo) relationship in olivine with a scattered negative trend.
The Ca and Na contents in olivine show large variations in concentrations and are very
sensitive towards temperature conditions, relative to Fe, Mg and Ni (De Hoog et al., 2009;
Pearson et al., 2003). On the other hand, the large scatter may be due the low abundance levels
that are close to the detection limits. CaO and Na2O tend to decrease with increasing Mg# in
olivine (Fig. 14).
Olivine from the Premier xenolith suite sample P6 is considered exceptional due to large
deviations in the chemical composition relative to other localities and was not included in
average calculations. There are two recognizable olivine generations in sample P6 (see Chapter
8.1.3). The Mg# of typical P6 olivine is extremely low at 79.3, which significantly differs from
a common fertile upper mantle peridotite olivine composition. The NiO content of 0.75 wt% is
significantly elevated compared to olivines in other locations. The second type of olivine
shows enrichment in Na2O (0.11 wt%), CaO (0.88 wt%), K2O (0.32 wt%) and Al2O3 (0.87
wt%). Such a composition in P6 olivine may indicate a metasomatized origin.
56
Figure 14. (A) Na2O and Mg# relationship in olivine with a negative; (B) CaO and Mg# relationship in
olivine with a negative trend.
8.2.2. Garnet
Garnet is abundant in 13 thin sections. Overall, the major oxide contents in the garnet
populations are broadly similar and don’t significantly deviate from the average values
determined for Kaapvaal craton peridotite xenoliths (Dawson & Stephens, 1975; Pearson et al.,
2003). The garnet analyses of this study yielded the following average compositions (wt%):
SiO2 42.0%, Al2O3 20.27%, MgO 21.3%, FeO 6.98%, CaO 5.05%, MnO 0.32%, Na2O 0.07%,
NiO 0.03%, K2O 0.01%, and V2O3 0.05% . However, several exceptions were noted among
the localities, such as the Frank Smith XS1, XS2 and XS7 garnets with highly elevated TiO2
contents of 1.27%, 0.99% and 1.29%, respectively, and the Premier P18 sample with a moderate
TiO2 content of 0.53%. Garnet compositions from Letlhakane and Letseng have relatively
higher average Cr2O3, especially in samples 523 (6.75 wt%), 550 (5.36 wt%) and LET4 (6.05
wt%). Notably, sample XS5 from the Frank Smith kimberlite also shows an elevated Cr2O3
content of 5.99 wt% deviating significantly from the average value.
Analyzed garnet grains contain a wide range of Cr2O3 contents, 1.5-6.7 wt%, and rather
uniform MgO and CaO contents of 20.4-22.1 wt% and 5.6-5.6 wt%, respectively. None of
garnet compositions fall into the sub-calcic range, due to their elevated CaO concentrations.
There is no apparent correlation between MgO and Cr2O3 contents (Fig. 15A), only CaO and
Cr2O3 show a week positive correlation together with a systematic Al2O3 decrease, which is a
common indicator of melt depletion (Pearson et al., 2003). Garnets have an average Mg# value
of 84.5 with a wide range of 81.6-87.1 among the localities. The average Cr# is 12.0 and all
garnet compositions exhibit an extreme Cr# variation of 4.53-19.44 between the localities. They
have positive correlation between Cr# and Mg# (Fig. 15B).
57
Figure 15. (A) Distribution of MgO and Cr2O3 in analyzed garnet grains; (B) Positive trend between
Mg# and Cr# in garnet compositions.
According to classification scheme by Schulze (2003), garnets derived either from
crustal or mantle source can be distinguished based on their Mg#, Ca# and Cr2O3 relationships
(Fig. 16).
Figure 16. Flow chart illustrating Mg#, Ca# and Cr2O3 relationships used to distinguish between crustal
or mantle derived garnets (Schulze, 2003). Published with permission from Elsevier.
All analyzed garnets are mantle derived (Fig. 17A). Most of the garnet compositions
fall into the lherzolitic paragenesis field, as only the Letlhakane samples 550 and 513 belong to
the harzburgite suite. Samples 523 and XS5 are in the borderline between the harzburgite and
lherzolite parageneses. None of the analyzed garnet grains are derived from a wehrlitic source
(Fig. 17B). However, based on the mineral abundances in thin sections (Chapter 8.1), garnets
from the Frank Smith XS1 and XS7 samples are from garnet-bearing dunite, since these samples
lack sufficient amounts of pyroxenes to be classified as harzburgite, lherzolite or wehrlite. A
similar problem occurs with the Letseng sample LET8 and Frank Smith sample XS5, which that
lack clinopyroxene, thus falling into the harzburgite field.
58
Figure 17. (A) Mg# vs. Ca# diagram discriminating between crustal garnets and mantle-derived garnets;
(B) classification of lithological associations based on the CaO and Cr2O3 contents in garnet. Diagrams
taken from Schulze (2003).
None of the analyzed garnet grains fall into the eclogite or pyroxenite (G3-G5) fields
(Fig. 18), which are defined by Mg#, CaO, Cr2O3 and TiO2 concentration relationships (Schulze,
2003; Grutter et al., 2004). Otherwise, sample 550 belongs to (G10) harzburgite suite, and the
rest of the analyzed garnet grains are distributed between the (G9) lherzolite and (G11) high-
TiO2 peridotitic sources.
Figure 18. (A) Analyzed garnet population paragenesis according to the G0-G12 classification scheme
by Schulze (2004); (B) The relationship between G1/G5/G9 suites among the analyzed garnet
populations. The G5 boundary is expressed with Mg# being bellow 70 (Grutter et al., 2004).
The eclogitic paragenesis is separated from the peridotitic paragenesis using the
following conditions for garnet compositions: Mg# <70, Cr2O3 <1 wt%, and TiO2 <0.5 wt%
(Schulze, 2003) or by the function TiO2%=213-210*Mg# (Grutter et al., 2004). The same
ranges apply to pyroxenite garnets, with the exception of a higher Mg# range of <90.
Considering the (G1) Cr-poor megacryst field overlap as the main divide between (G9) and
(G3-G5) groups, the following conditions are used distinguish the groups: TiO2>0.5 wt% and
59
Cr2O3 <4 wt% (Schulze, 2003) or function of TiO2(wt%) = 213-210*Mg# and Cr2O3 <4 wt%
(Grutter et al., 2004). Since none of the overlapping garnet samples from this study meet
conditions for (G1) Cr-poor garnet megacryst suite, samples P18, XS1, XS2, and XS7 are
considered to belong to the (G11) high-TiO2 peridotitic paragenesis.
8.2.3. Orthopyroxene
All analyzed orthopyroxenes are classified as enstatites (Fig. 19) and contain only minor
ferrosilite and wollastonite components, based on the En-Fs-Wo end-member proportions
(Morimoto et al., 1988; Demange, 2012). Orthopyroxenes ranges in composition between En81
and En94 with an average value of En91.6 and forms a rather uniform compositional cluster with
insignificant differences between observed kimberlite localities. The only exception is the
analyzed P6 orthopyroxene from the Premier kimberlite, which has a rather elevated ferrosilite
component Fs17.9 and stands out from the cluster. Otherwise, the average ferrosilite and
wollastonite abundance in orthopyroxene is Fs7.4 and Wo1.0, respectively.
Figure 19: Orthopyroxene compositions plotted in the lower part of the enstatite-ferrosilite-wollastonite
ternary diagram. The 5% wollastonite line marks the orthopyroxene-clinopyroxene boundary. All
analyzed samples fall into the enstatite field, forming a single cluster, with the exception of sample P6
from the Premier kimberlite. Diagram taken from Morimoto et al. (1988).
Mg# of orthopyroxene is rather constant among the localities and varies between 90.2
and 94.4 with an average of 93.0, except for the Premier sample P6, which has an anomalously
low Mg# value of 82.0 and was therefore been excluded from graphical representations.
Compared to Mg# of olivine (Fig. 20A), orthopyroxene has slightly greater values averaging
0.70. Such relative Mg# differences correspond to equilibrium partitioning of Mg and Fe
between orthopyroxene and olivine (Brey & Kohler, 1990). Samples 512 and P14 have slightly
lower Mg# in orthopyroxene compared to corresponding Mg# of olivine.
60
Figure 20. (A) Mg# in orthopyroxene and olivine. The majority of orthopyroxenes, with two exceptions,
have slightly higher Mg# compared to olivine; (B) Positive correlation between Mg# of orthopyroxene
and olivine.
An increase in Mg# in orthopyroxene very well correlates with the corresponding Mg#
increase in olivine (Fig. 20B). The Letseng orthopyroxene compositions have the highest Mg#,
followed by the orthopyroxene compositions from Premier with slightly lower and clustered
Mg#. The Letlhakane and Frank Smith orthopyroxene compositions have a wide range of Mg#.
There is no apparent correlation with Mg# of orthopyroxene and corresponding rock types or
textures, except for spinel-bearing harzburgites that have very uniform Mg# in both
orthopyroxene and olivine, ranging from 92.9 to 93.5 for orthopyroxene and from 92.5 to 93.6
for olivine. Cr# varies widely from 7.9 to 36.6 with no apparent correlation with Mg#.
The analyzed orthopyroxene grains are poor in minor elements, having the following
average concentrations (wt%): CaO 0.55%, Al2O31.17%, TiO20.07% and Na2O 0.14%. All
minor oxides correlate roughly with Mg#, although there are several exceptions (Fig. 21).
Compared to other rock types, the spinel-bearing harzburgite samples P5, P14, 531 and
551 have elevated Al2O3 contents of >1.5% in orthopyroxene. Sample 512 also has a slightly
elevated Al content, despite that no spinel association was identified. Otherwise, Al2O3 clusters
below average with no insignificant variations as a function of Mg# or the depth of origin.
61
Figure 21. Minor oxide concentrations in orthopyroxene as a function of Mg#.
In the case of Na2O, the spinel-bearing harzburgites have extremely low concentrations
of >0.1 wt%, compared to other rock types. Notably, sample 512 also shows a similar Na2O
concentration to that of the cluster from the spinel-bearing association. The TiO2 contents tend
to decrease in a similar manner as Na2O, with the spinel-bearing association having the lowest
TiO2 concentrations of <0.05 wt% and roughly correlate with increasing Mg#. The CaO content
in orthopyroxene is low and highly variable, falling in the range of 0.17-1.27 wt%, both between
individual localities and rock types.
8.2.4. Clinopyroxene
All Letlhakane xenolith samples, with the exception of sample 511, contain
clinopyroxene. In other localities, most of the samples lack clinopyroxene or mineral gains were
not suitable for chemical analysis. In total, analysis was carried out for 17 samples: Letlhakane
(9), Letseng (2), Premier (4), and Frank Smith (2).
In terms of the main Mg-Fe-Ca components, analyzed clinopyroxenes are classified as
augites or diopsides (Fig. 22). Letlhakane and Premier clinopyroxenes extend into diopside and
augite fields, while Letseng and Frank Smith clinopyroxenes are all augites.
62
Figure 22. En-Fs-Wo classification scheme of clinopyroxene. Analyzed clinopyroxenes fall in to the
augite and diopside field. Diagram taken from Morimoto et al., (1988).
Mg# of clinopyroxenes varies from 85.9 to 95.9, with an average of 92.5, which is
generally slightly higher than Mg# of olivine. There is a week positive correlation between Mg#
in clinopyroxene and olivine (Fig. 23). Notably, the difference between the Mg# values
correlates linearly with increasing Mg# in clinopyroxene. Sample P6 from the Premier
kimberlite was excluded from the graphical representation due to an anomalously low Mg# in
olivine.
Figure 23. Correlation of Mg# between olivine and clinopyroxene.
The Cr# value of analyzed clinopyroxenes varies greatly from 14.2 to 49.9 with an
average of 29.3. There is no apparent correlation with Mg# of clinopyroxene. The CaO content
varies between 16.5 and 24.0 wt% and averages is 19.3 wt%. The Frank Smith clinopyroxenes
contain relatively low CaO concentrations, compared to other localities. Al2O3 in clinopyroxene
varies from 1.0 to 4.3 wt%, with an average of 2.7 wt%, and is considered to be rather low. The
63
Letlhakane xenoliths tend to have higher Al2O3 concentrations. Overall, Al2O3 only correlates
positively with increasing Na2O, as there is no noticeable trend with other major oxides. The
Na2O content varies greatly, ranging from 0.6 to 3.7 wt%, and has an average value of 2.0 wt%.
TiO2 in clinopyroxene is also low and ranges from 0.01 to 0.54 wt%.
The temperature and pressure sensitive oxides Na2O, CaO, Al2O3 and TiO2 were
compared with Mg# of clinopyroxene to identify any correlation (Fig. 24). Na2O and TiO2
correlate negatively with Mg#, which is in good argument with a common depletion trend,
whereas CaO has a definite positive correlation with Mg#. Instead, Al2O3 shows no apparent
correlation with Mg#.
Figure 24: Concentrations of temperature- and pressure-sensitive oxides Al2O3, CaO, Na2O and TiO2
as a function of Mg# in analyzed clinopyroxene grains.
8.2.5. Spinel
Spinel grains were analyzed from two Letlhakane samples, 531 and 551, and two
Premier samples, P5 and P14. Both Letlhakane samples yielded low Cr# and high Mg# values
that slightly differ: 28.6 Cr# and 73.8 Mg# in sample 531, 22.9 Cr# and 75.8 Mg# in sample
551. The Letlhakane spinel grains have very low TiO2 contents of <0.1% and relatively high
64
NiO concentrations of 0.17% and 0.24%. The Premier P5 spinel has the same characteristics as
the Letlhakane spinel, but sample P14 differs by having higher concentrations of Cr2O3 (44.4
wt%) and FeO (13.1 wt%) and lower concentrations of Al2O3 (26.8 wt%) and MgO (15.4 wt%).
8.3. Geothermobarometry
8.3.1. Overview
In this study, pressure and temperature conditions for mantle xenolith assemblages were
calculated using the PTXL3 spreadsheet for all available barometers and thermometers with a
starting preset values of T=1000 0C and P=40 kb. Further on, only results obtained by the
thermometers and barometers discussed in Chapter 7 were chosen to present the best P-T fit for
the Kaapvaal craton regional geotherm.
Samples 511, and P17B were excluded from the geothermobarometric calculations, due
to insufficient information on mineral geochemical data. The mentioned mantle xenoliths were
analyzed only for olivine and orthopyroxene chemistry, as garnet and clinopyroxene were
absent or not suitable for chemical analysis. Besides, most of the thermometers and barometers
presented in PTXL3 rely on the presence of both orthopyroxene and clinopyroxene in the
peridotite suite. Furthermore, P-T relationships between thermobarometers need to be properly
paired, and thus PTXL3 does not calculate P-T values without the missing mineral assemblages.
Representative thermometer and barometer pairs for each mantle xenolith suite are listed in
Table 4 and discussed below. Chosen representative P-T pairs could also be changed depending
on the obtained accuracy or validity, compared to similar rock assemblages or
geothermobarometers. Full results are summarized in Appendix II.
For four-phase garnet bearing peridotites (9 samples), a combination of orthopyroxene-
clinopyroxene solvus thermometer T(BKN90) and Al-in-orthopyroxene barometer P(BBG08) was
used. Preset P-T values were gradually integrated to obtain best available results that would
match with the integrated presets. P-T conditions for garnet-free peridotites (5 samples) were
calculated by pairing orthopyroxene-clinopyroxene solvus thermometer T(BKN90) and Cr-in-
clinopyroxene barometer P(NimisTaylor00). Pressure was integrated in a similar manner as in case
of four-phase garnet bearing peridotites, but temperature estimates were calculated on the preset
T=1000 0C basis, for otherwise the calculations produced temperatures with high uncertainties.
65
Table 4. Thermometer and barometer pairs used for P-T calculations.
Sample number Mineral
assemblage P-T pair
P preset,
kb
T preset, 0C
502, 513, 523, 550, LET4,
P15, P18, XS2, XS4
Ol, Opx,
Cpx, Gt
P(BBG08) - T(BKN90) Integr. Integr.
512, 529, 530, LET14, P6 Ol, Opx,
Cpx
P(NimisTaylor00) -
T(BKN90)
Integr. 1000
LET8, XS5 Ol, Opx, Gt P(BBG08) –
T(O’NW79)
40 1000
XS1, XS7 Ol, Gt P(Interpolated) –
T(O’NW79)
40 1000
531, 551, P5, P14 Ol, Opx,
Cpx, Sp
P(NimisTaylor00) –
T(OW87)
40 1000
511, P17B Ol, Opx - - -
For clinopyroxene-free peridotites (2 samples), the Fe-Mg exchange between garnet and
olivine thermometer T(O’NW79) coupled with the Al-in-orthopyroxene barometer P(BBG08) was
used. For spinel-bearing peridotites (4 samples), a combination of the Mg-Fe exchange between
olivine and spinel thermometer T(OW87) and the Cr-in-clinopyroxene barometer P(NimisTaylor00)
was applied. Calculated P-T estimates are based on the starting preset values, since the
integration was not possible due to high variations in the obtained temperature and pressure
values.
In the case of garnet-bearing dunites from the Frank Smith kimberlite, temperature was
calculated using the Fe-Mg exchange between garnet and olivine thermometer T(O’NW79). None
of the barometers available in PTXL3 provided usable pressure data, and therefore pressure
conditions for garnet-bearing dunites were linearly interpolated based on the calculated pressure
data from other Frank Smith xenolith suite samples.
Calculated P-T conditions using the PTEXL3 geothermobarometers for mantle xenoliths
cover a wide temperature and pressure ranges of 642-1356 0C and 27-56 kb, respectively, and
differ in several ways: The inter-mineral equilibrium temperatures and pressures differ based
on the used calculation method; temperature and pressure estimates differ between localities;
and temperature and pressure estimates vary between certain lithologies.
8.3.2. Inter-mineral equilibrium and P-T estimates
To test inter-mineral equilibrium, the calculated P-T values were compared with
different independent thermometers for a given barometer. Only four-phase garnet-bearing
assemblages were tested for inter-mineral equilibrium, since temperature and pressure data for
66
these samples were available from all PTXL3 listed geothermobarometers and integrated preset
values were in good argument with the calculated results. For the remaining samples, there
either was only partly supportive data available or there was not enough information on
geothermobarometric data to assess inter-mineral equilibrium.
In the case of four-phase garnet-bearing xenoliths, the calculated results from the
thermobarometric pair T(BKN90) - P(BBG08) were compared with the following thermometers:
T(KB90); T(OpxBK90); T(NaPxBK90); T(NimisTaylor00); and T(Krogh88). Samples 523 and 550 from the
Letlhakane kimberlite and sample LET4 from the Letseng kimberlite show the highest
uncertainties in the calculated temperatures compared to other thermometers (Table 5; Fig. 25).
Variations in the T estimates between individual localities reflect differences in the mineral
chemistry and possible influences of alteration, which may add or remove certain elements to
or from the system. The chosen calculation method may also be a reason for high ΔT deviations
between localities and individual samples, and thus a different P-T calculation approach should
be tested and used. The rest of the mantle xenolith samples show only little or average
deviations in ΔT and are in good agreement with the geothermobarometric pair T(BKN90) -
P(BBG08), concerning especially samples from the Premier and Frank Smith kimberlites.
Table 5. Results of testing inter-mineral equilibrium in four-phase garnet-bearing xenoliths. The
calculated numbers represent temperature differences (ΔT) relative to the T(BKN90) geothermometer. ΔT\Sample Nr. 502 513 523 550 P15 P18 XS2 XS4 LET4
T(KB90)-
T(BKN90) 34 11 -120 -146 -59 0 -84 -83 -170
T(opxBK90)-
T(BKN90) 40 -150 -88 -115 -3 -18 -38 -24 -194
T(NaPxBK90)-
T(BKN90) -23 57 -36 -64 160 37 24 102 -197
T(NimisTaylor00)-
T(BKN90) -6 -89 -161 -174 -80 -24 -59 -56 -41
T(Krogh88)-
T(BKN90) -55 -32 60 51 69 -86 -56 33 -93
Overall, most of the selected thermometers underestimate the temperature compared to
T(BKN90). This concerns especially T(NimisTaylor00), which yields a ΔT range from -6 to -174 0C
with an average ΔT of -77 0C (Table 5; Fig. 25E). Results from ΔT(NaPxBK90) are scattered and
have the highest underestimate and overestimate of ΔT values, compared to other
geothermobarometric pairs. Deviations may also be influenced by the presence of Fe3+, since
most of the geothermobarometers use total Fe for P-T calculations. Nevertheless, most of the
calculated results lie within the error limits of the compared method.
67
The Fe2+-Mg exchange thermometer T(Krogh88) is in good agreement with T(BKN90), where
ΔT shows a week negative correlation with temperature, compared to other pairs. Notably, the
average ΔT value is close to 0.
Figure 25. T(KB90), T(OpxBK90), T(NaPxBKN90), T(NimisTaylor00) and T(Krogh88) thermometer comparison with
T(BKN90) for four-phase garnet-bearing xenoliths; (A) Summary of maximum, minimum and average ΔT
values for compared thermometers; (B-F) ΔT between T(BKN90) and other thermometers.
Regarding pressure calculations, the majority of independent barometers for garnet-
bearing peridotite assemblages are restricted to P(BBG08), P(BKN90), P(KB90) and P(NimisTaylor00).
Listed barometers were compared for inter-mineral equilibrium on the basis of T(BKN90)
temperature preset (Table 6, Fig. 26).
As in the case of temperature, samples LET4, 523 and 550 show the highest ΔP from all
tested barometers. Otherwise, pressure calculations are in good agreement within their
68
respective errors. Overall, the geobarometers P(NimisTaylor00) and P(BKN90) give slightly higher
pressure estimates and are rather similar among the tested samples. Compared to two other
barometers, P(KB90) provided relatively higher pressures and the pressure values are more
scattered between samples and their corresponding localities.
Table 6. Results of testing inter-mineral equilibrium in four-phase garnet-bearing xenoliths. The
calculated numbers represent the pressure differences (ΔP) relative to the P(BBG08) geobarometer. ΔP\Sample Nr. 502 513 523 550 P15 P18 XS2 XS4 LET4
P(BKN90)-
P(BBG08) 7 4 4 4 2 5 3 4 6
P(KB90)-
P(BBG08) -4 -1 19 23 9 1 15 16 28
P(NimisTylor00)-
P(BBG08) 6 0 8 5 8 6 7 4 -10
The observed differences between the pressure estimates can be explained by the error
of the chosen method, potential effect of the iron content or the low element abundances in
some of the minerals used for the pressure calculation.
Figure 26. Comparison of the P(NimisTaylor00), P(BKN90) and P(KB90) barometers with P(BBG08) in case of four-
phase garnet-bearing xenoliths; (A) Summary of maximum, minimum and average ΔP values for
compared barometers; (B-D) ΔP result summary for compared barometers with P(BBG08).
69
8.3.3. P-T conditions and depth of origin of studied mantle xenoliths
For samples that yielded incomplete pressure or temperature estimates from the
PTEXL3 calculations the equilibrium conditions were chosen based on relative differences
between the results from other available thermobarometers or on similarities within individual
localities or corresponding rock types. Also, temperature and pressure results were best-fitted
with the general Kaapvaal craton lithospheric mantle thermal profile. Full temperature, pressure
and depth estimates are summarized in Table 7 and discussed further below.
Table 7. Representative pressure, temperature and depth estimates for studied mantle xenoliths.
The equilibrium pressures of 22-56 kb and temperatures of 753-1344 0C obtained for
the Letlhakane, Premier, Letseng and Frank Smith kimberlite xenoliths suggest that the
Sample T, 0C P, kb H, km Thermometer Barometer
Letlhakane
502 1254 53 177 T [KB90] P [BBG08]
511 x x x x x
512 983 32 106 T [BKN90] P [NimisTaylor00]
513 1072 41 136 T [BKN90] P [BBG08]
523 1013 36 120 T [BKN90] P [BBG08]
529 1171 47 155 T [KB90] P [NimisTaylor00]
530 1246 45 149 T [BKN90] P [NimisTaylor00]
531 986 36 121 T [KB90] P [NimisTaylor00]
550 1065 41 138 T [BKN90] P [BBG08]
551 841 28 93 T [OpxBK90] P [BBG08]
Letseng
LET4 1048 41 137 T [OpxBK90] P [NimisTaylor00]
LET8 1051 36 121 T [O'NW79] P [BBG08]
LET14 1119 45 151 T [KB90] P [NimisTaylor00]
Premier
P5 883 27 92 T [KB90] P [NimisTaylor00]
P6 1021 37 122 T [OpxBK90] P [NimisTaylor00]
P14 753 22* 73 T [O'NW79] *
P15 892 29 96 T [BKN90] P [BBG08]
P17B x x x x x
P18 1344 54 179 T [BKN90] P [BBG08]
Frank Smith
XS1 1249 48* 160 T [O'NW79] *
XS2 1344 56 186 T [BKN90] P [BBG08]
XS4 1296 51 169 T [BKN90] P [BBG08]
XS5 1051 37 124 T [O'NW79] P [BBG08]
XS7 1266 49* 163 T [O'NW79] *
*Assumed pressure value based on linear interpolation
70
xenoliths were sampled by kimberlite magma in the subcontinental lithospheric mantle region
beneath the Kaapvaal craton from approximately depths of 70 to 190 km (Fig. 27). The depth
of origin varies depending on the kimberlite location within the Kaapvaal craton. All calculated
P-T results fit in a thermal range between the conductive continental geotherms of 40 mWm-2
and 50 mWm-2. More precisely, the xenolith P-T estimates cluster around a conductive
geotherm of 44.0±2.0 mWm-2, which is slightly higher than the average Kaapvaal craton
geotherm with a heat flow of 40±2.0 mWm-2 (Hasterok & Chapman, 2011).
Fig. 27. Geotherm plot for Letlhakane, Letseng, Premier and Frank Smith kimberlite xenoliths.
Graphite-diamond stability field is taken from Kennedy & Kennedy (1976) and the conductive
continental geotherms of 40 mW/m2 and 50 mW/m2 from Pollack & Chapman (1977). Approximate
spinel association identified in this study is marked by the blue oval line.
The relative shift from the average continental conductive geotherm of the Kaapvaal
craton towards a higher thermal regime for obtained xenolith P-T estimates is more likely
71
caused by uncertainties from the pressure and temperature calculations rather than from melt-
related introduction in the system that would change mineral equilibrium conditions. However,
Rudnick & Nyblade (1999) estimated that the present-day conductive geotherm beneath the
Kalahari craton is characterized by an average surface heat flow of 47±2.0 mWm-2. For one
reason to back the argument, temperature estimates obtained by T(BKN90) thermometer in most
cases provided up to ~50 0C higher temperature results compared to other thermometer
alternatives. Most of thermobarometers are sensitive on very low abundances of chemical
elements used as the basis for the calculations or are required for additional mineral
assemblages to be present. From the geochemical data, , it is evident that elements like Ca, Na
and Al are relatively depleted, especially in olivine, which would cause deviations in
calculations.
The Letlhakane xenoliths have equilibrated under a considerable range of T and P
conditions, ranging from 841 to 1254 0C and 28 to 53 kb, respectively (Fig. 27). As expected,
the spinel-bearing associations (samples 531 and 551) have the lowest calculated temperature
and pressure values, ranging from 841 to 986 0C and 28 to 36 kb, which correspond to sampling
depths of 93 and 121 km. The harzburgite sample 512 yielded a rather low temperature of 983
0C and pressure of 32 kb, being in this sense very similar to the Letlhakane spinel-bearing
association. The absence of garnet and geochemical results also support the fact that the
harzburgite sample should belong to the spinel association. Further investigation of this
particular sample would be necessary to fully clarify its origin and equilibrium conditions. The
garnet-bearing harzburgite xenolith sample 502 has equilibrium conditions that plot in the
diamond stability field (Kennedy & Kennedy, 1976). The remaining samples are scattered along
the 44.0±2.0 mWm-2 conductive geotherm and have equilibrium conditions that match a
sampling depth of 120-149 km. Compared to the previous studies of the Letlhakane xenoliths
(Stiefenhofer et al., 1997) where deformed varieties tend to have higher P-T equilibrium
conditions than the coarse rock types, there is no clear correlation among the Letlhakane
peridotitic xenolith types or their textures examined in this study.
The Letseng garnet-bearing harzburgite xenoliths yield generally rather uniform
calculated temperature and pressure conditions, ranging from 1048 to 1119 0C and 36 to 45 kb
respectively. The depth of origin for the Letseng cold xenoliths lies between 121 and 151 km,
being consistent withnthe previous results from the Lesotho kimberlite xenoliths (Boyd &
Nixon, 1975; Simon et al., 2003). Samples LET4 and LET8 gave more similar and slightly lower
calculated P-T values compared to sample LET14. Compared to he spinel-bearing associations
from the Letlhakane and Premier kimberlites, the Letseng xenoliths show slightly higher
72
temperature and pressure values. Notably, the Letseng xenoliths contain minor spinel that is
associated with garnet. There is no concrete textural difference among the samples. Coarse
xenoliths tend to have a more shallow origin, which is also the case for the Letseng xenoliths
studied in this thesis.
Despite the small sample pool, the Premier xenoliths cover the most extensive
equilibrium temperature and pressure range among all studied localities (Fig. 27). The spinel-
bearing association of samples P14 and P5 indicates a shallow 73-92 km depth of origin with
corresponding temperatures of and pressures of 22-27 kb. Notably, sample P15 also falls into
the spinel field, based on the obtained P-T results. It could be such a case, despite that there
were no spinel minerals analyzed in this garnet harzburgite. Garnets in the P15 sample are
strongly depleted, having very thick reaction rims that partly consist of fine spinel. Compared
to the Letlhakane spinel association, it could be stated that the samples with coexisting garnet
and spinel with signs of depletion belong to the spinel-garnet facies transition zone (Pearson et
al., 2003).
The metasomatised olivine websterite xenolith sample P6 yielded an equilibrium
temperature of 1021 0C and pressure of 37 kb according to T(OpxBK90)- P(NimisTaylor00) pair. The
sampling depth for olivine websterite is estimated to be 122 km. It is known that the Premier
kimberlite contains extensively metasomatized rock sections (Griffin et al., 2003), but their
high abundance is constrained to much deeper parts (>180 km) of the Kaapvaal craton
subcontinental mantle, compared to P6 sample. Similar correlation was also noted in other
kimberlites bound to the Kaapvaal craton, although shallow metasomatized xenoliths may also
occur in smaller quantities. Thus, sample P6 reveals a rather unusual sampling depth.
Nevertheless, it is not excluded that metasomatic melt could have percolated from deeper parts
and affected shallower sections of the mantle.
The deepest seating was calculated for the deformed garnet lherzolite sample P18. The
Xenolith originates from a depth of 179 km which is located in the diamond stability field and
corresponds to P-T equilibrium conditions of 1344 0C and 54 kb. The mosaic-porphyroclastic
texture of sample P18 is also characteristic for deep-seated xenoliths.
The calculated temperature and pressure conditions for the Frank Smith xenoliths differ
generally from other three localities in a way that the equilibrium temperatures and pressures
are higher, except for sample XS5, which yielded similar P-T values to those of the Letseng
xenoliths. For the remaining Frank Smith peridotites, the temperature estimates range from
1249 to 1344 0C and those for pressure from 48 to 56 kb, which correspond to sampling depths
from 160 to 189 km. The garnet-bearing dunite xenoliths gave lower pressure and temperature
estimates relative to the garnet-bearing harzburgites, besides that the sample XS2 xenolith plots
73
within the diamond stability field. There is a correlation between coarse-grained and mosaic-
porphyroclastic varieties. The mosaic-porphyroclastic sample XS2 has the highest calculated P-
T values compared to all other samples from Frank Smith, being similar to the values of the
Premier sample P18. Cleary, the Frank Smith xenoliths have a more deep-seated origin
compared to other localities.
8.3.4. Composition of the subcontinental lithospheric mantle
The uppermost mantle is mainly characterized by fertile peridotite of a lherzolitic
composition. In such a rock type, olivine is the dominant mineral phase covering 50-60 % of
the total rock volume, followed by orthopyroxene constituting roughly 20-25 % of the rock
volume. Clinopyroxene and the pressure-dependent aluminous minerals, plagioclase, spinel or
garnet, are the remaining minor phases present. Plagioclase is stable at pressures up to
approximately 10 kb at the solidus (Fig. 5) and is replaced at higher depths by more stable
spinel, which is stable under pressure conditions from ~10 to 25 kb. At greater pressures than
25 kb, garnet becomes the stable aluminous phase in a peridotitic rock (Walter, 2003). Other
main features of a fertile mantle composition are a low bulk Mg# value of 89.0±0.5 and
relatively high Al, Ca, Fe and REE contents. Rare earth element distribution patterns are not
considered in this study, since there are no trace element data available for the studied samples.
The obtained mineral chemical data from the studied samples, combined with the
geothermobarometric calculations, differ from the typical fertile mantle composition. Instead,
it can be stated that the majority of xenoliths were sampled from a depleted peridotitic source
within the Kaapvaal craton subcontinental mantle. For instance, Mg# in olivine, which is
commonly used as an indicator for melt depletion, indicates a progressive decrease with
increasing depth towards more typical values in an asthenosphere composition (Fig. 28) (Gaul
et al., 2000; Griffin et al., 2003; Pearson et al., 2003; Walter, 2003). The same correlation
applies for other major elements, such as Al, Ca, which are extremely low in olivine, as
discussed in Chapter 8.2.1. Another indicator that may be used to identify possible melt
extraction is the TiO2 content in garnet or olivine as titanium prefers to enter the melt phase
during melt extraction from mantle rock.
As seen from Figure 28, the TiO2 content in olivine and garnet is reversely correlated
with Mg# of olivine and is constantly decreasing towards shallower depths of origin. Down to
~160 km (P 48 kb), Mg# is relatively uniform in all locations, but with a further increase in
depth, Mg# starts to drop rapidly, especially in the Frank Smith xenolith samples XS1, XS2 and
74
XS7, which also show the highest range in Mg# (88.8-89.6) and highest TiO2 content compared
to other samples. Possibly, xenoliths XS1, XS2 and XS7 may represent the deepest setting,
though there were no concrete pressure calculations available and the pressure was assumed
based on the overall Frank Smith xenolith distribution along the continental geotherm.
Figure 28. (A) Distribution of Mg# of olivine with depth (pressure). Progressive increase in Mg# with
decreasing sampling depth indicates possible melt extraction from the host rock; (B) TiO2 (wt%) in
garnet as a function of depth; (C) TiO2 (wt%) in olivine as a function of depth.
In terms of modal abundances, mineral phases of a fertile lherzolite are progressively
depleted during partial melting and the first mineral starting to be consumed is an aluminous
pressure-dependent mineral. With increasing temperature, further clinopyroxene,
orthopyroxene and finally olivine is consumed. By progressively consuming aluminous phases
and pyroxenes, the proportion of olivine in the host rock increases, leaving behind a refractory
depleted harzburgite or dunite (Pearson et al., 2003; Walter, 2003). The majority of the studied
samples, as discussed earlier, are harzburgites or dunites with a high modal olivine, moderate
orthopyroxene and low clinopyroxene abundances, which are typical for depleted residues.
Besides, garnet is present generally in small quantities and has very well-developed alteration
rims that indicate an aluminous phase consumption. From the statistical data on garnet
xenocrysts obtained from more than 50 Kaapvaal craton kimberlites, Griffin et al. (2003) noted
that the Kaapvaal craton SCLM is most complex and heterogeneous and for a large part, is
75
dominated by highly depleted peridotitic rock types. In many cases, especially at the deepest
settings, SCLM has been metasomatically refertilized. The obtained results of this study are
generally in good argument with the Kaapvaal craton SCLM model. Xenolith populations are
heterogeneous and vary between localities both in a vertical and horizontal direction. However,
the sample set is too small to construct a representative cross-section of the Kaapvaal craton.
The Frank Smith xenoliths XS1 and XS7 are garnet-bearing dunites and may represent the
continental root zone that has been metasomatized and partly refertilized. Enrichment of Fe and
Ti in the root zone may have been caused by the interaction with an asthenospheric melt. The
deformed Premier and Frank Smith xenoliths likely represent the fertile root zone boundary.
76
9. Conclusions
The mantle xenoliths from Letlhakane, Letseng, Premier and Frank Smith kimberlites
studied in this thesis are peridotites, mainly garnet-bearing harzburgites. One sample from the
Premier kimberlite (P6) was identified as olivine websterite. The average olivine abundance in
the xenolith samples is relatively high (69%) and that orthopyroxene is moderate (23%), while
garnet (8%) and clinopyroxene (4%) are the remaining minor phases. Four of the samples
contain minor spinel. Major element geochemical data indicate various depletion trends for the
analyzed minerals, especially for olivine and orthopyroxene, which correlate with the
corresponding sampling depth. Olivine and orthopyroxene have low Ca, Al and Ti contents and
high Mg# with corresponding average values of 90.1 and 92.5.
Mantle xenoliths have equilibrated at pressures ranging from 22 to 56 kb and
temperatures from 753 to 1344 0C, being equivivalent to an extensive sampling depth range
from 70 to 190 km, which also extends into the diamond stability field. Mantle xenoliths cluster
along a conductive continental geotherm of 44.0±2.0 mWm-2, which is slightly higher than that
of the average thermal state estimate for the Kaapvaal craton. Likely, the relative shift from the
average continental conductive geotherm towards a higher thermal regime is caused by
uncertainties and overestimates in the P-T calculations.
Based on the obtained results, it can be stated that the majority of the studied xenoliths
were sampled from a depleted peridotitic source within a heterogeneous Kaapvaal craton
subcontinental mantle. Lowermost xenoliths were possibly sampled from the continental root
zone or a metasomatically fertilized peridotite domain.
77
10. Acknowledgments
Special thanks to Shenghong Yang and Eero Hanski efforts for giving the opportunity
to work on this specific topic, providing guidance and helping to structure and edit the content
of the Master’s thesis. Also, special thanks to the staff from Center of Microscopy and
Nanotechnology (CMNT), University of Oulu, for providing guidance with the microprobe
analysis. Redistribution rights and necessary permits were acquired for quoted figures and
tables from the respective publishers.
78
11. References
Achterberg, E., Griffin, W.L., Steinfenhofer, J., 2001. Metasomatism in mantle xenoliths from
the Letlhakane kimberlites: estimation of element fluxes. Contributions to Mineralogy and
Petrology 141, 397-414.
Backer, M., Le Roex, A.P., 2006. Geochemistry of South African on- and off-craton, Group I
and Group II kimberlites: Petrogenesis and source region evolution. Journal of Petrology 47,
673-703.
Barnes, S.J., Roeder, P.L., 2001. The range of spinel compositions in terrestrial mafic and
ultramafic rocks. Journal of Petrology 42, 2279-2302.
Bernstein, S., Kelemen, P.B., Hanghoj, K., 2007. Consistent olivine Mg# in cratonic mantle
reflects Archean mantle melting to the exhaustion of orthopyroxene. Geology 35, 459-462.
Boyd, F.R., 1973a. Olivine megacrysts from the kimberlites of the Monastery and Frank Smith
Mines, South Africa. Carnegie Institution of Washington Year Book 73, 282-285.
Boyd, F.R., 1973b. Ultramafic nodules from the Frank Smith kimberlite pipe, South Africa.
Carnegie Institution of Washington Year Book 73, 285-294.
Boyd, F.R., Nixon, P.H., 1975. Origins of the ultramafic nodules from some kimberlites of
Northern Lesotho and the Monastery Mine, South Africa. Physics and Chemistry of the Earth
9, 431-454.
Boyd, F.R., Tsai, H., 1979. A unique enstatite megacryst with coexisting Cr-poor and Cr-rich
garnet, Weltevreden Floors, South Africa. In: Boyd, F.R., Mayer, H.O.A. (eds.) The Mantle
Sample: Inclusions in Kimberlites and Other Volcanics. Proceedings of Second International
Kimberlite Conference, American Geophysical Union, Washington, p. 279-291.
Boyd, F.R., 1989. Kaapvaal spinel peridotites: evidence of craton origin. Carnegie Institution
of Washington, Annual Report of the Director of the Geophysical Laboratory 1988-1989, p.
3-6.
Brey, G.P., Köhler, T., Nickel, K.G., 1990. Geothermobarometry in four-phase lherzolites I.
Experimental results from 10 to 60 kb. Journal of Petrology 31, 1313-1352
Brey, G.P. and Köhler, T., 1990. Geothermobarometry in four-phase lherzolites II. New
thermobarometers, and practical assessment of existing thermobarometers. Journal of
Petrology 31, 1353-1378.
Brey, G.P., Bulatov, V.K., Grinis, A.V., 2008. Geobarometry for peridotites: Experiments in
simple and natural systems from 6 to 10 GPa. Journal of Petrology 49, 3-24.
Brown, R.J., Manya, S., Buisman, I., Fontana, G., Field, M., Niocaill, C.M., Sparks, R.S.J.,
Stuart, F.M., 2012: Eruption of kimberlite magmas: physical volcanology, geomorphology
and age of the youngest kimberlite volcanoes known on earth (the Upper
Pleistocene/Holocene Igwisi Hills volcanoes, Tanzania). Bulletin of Volcanology 74, 1621-
1643.
Cas, R.A.F., Porritt, L., Pittari, A., Hayman, P.C., 2009. A practical guide to terminology for
kimberlite facies: A systematic progression from descriptive to genetic, including pocket
guide. Lithos 112, 183-190
79
Cas, R., Porritt, L., Pittari, A., Hayman, P., 2008. A new approach to kimberlite facies
terminology using a revised general approach to the nomenclature of all volcanic rocks and
deposits: Descriptive to genetic. Journal of Volcanology and Geothermal Research 174, 226-
240.
Clarke, D.B., 1979. Synthesis of nickeloan djerfisherites and the origin of potassic sulphides at
the Frank Smith Mine. In: Boyd, F.R., Mayer, H.O.A. (eds.) The Mantle Sample: Inclusions
in Kimberlites and Other Volcanics. Proceedings of Second International Kimberlite
Conference, American Geophysical Union, Washington, p. 300-308.
Clement, C.R., Skinner, E.M.W., 1979. A textural genetic classification of kimberlite rocks.
Kimberlite Symposium II. Cambridge, Extended Abstracts.
Clement, C.R., 1982. A comparative geological study of some major kimberlite pipes in the
Northern Cape and Orange Free State. Ph.D. thesis (2 vols.), University of Cape Town, 432
p.
Crough, S.T., Morgan, W.J., Hargraves, R.B., 1980. Kimberlites: their relation to mantle
hotspots. Earth and Planetary Science Letters 50, 260-274.
Danchin, R.V., 1979. Mineral and bulk chemistry of garnet lherzolites and garnet harzburgite
xenoliths from Premier Mine, South Africa. In: Boyd, F.R., Mayer, H.O.A. (eds.) The Mantle
Sample: Inclusions in Kimberlites and Other Volcanics. Proceedings of Second International
Kimberlite Conference, American Geophysical Union, Washington, p. 104-126.
Davis, G.L., 1977. The ages and uranium contents of zircons from kimberlites and associated
rocks. Carnegie Institute Washington Yearbook 76, 631-635.
Dawson, J.B., 1980. Kimberlites and Their Xenoliths. Springer-Verlag, New York, 250 p.
Dawson, J.B., Stephens, W.E., 1975. Statistical classification of garnets from kimberlite and
associated xenoliths. Journal of Geology 83, 589-607.
De Hoog, J.C.M., Gall, L., Cornell, D.H., 2009. Trace-element geochemistry of mantle olivine
and application to mantle petrogenesis and geothermobarometry. Chemical Geology 270, 196-
215.
Demange, M., 2012. Mineralogy for Petrologists: Optics, Chemistry and Occurrences of Rock-
forming Minerals. Taylor & Francis Group, London, 201 p.
de Wit, M.J., Roering, C., Hart, R.J., Armstrong, R.A., de Ronde, C.E.J., Green, R.W.E.,
Tredoux, M., Peberdy, E., Hart, R.A., 1992: Formation of an Archean continent. Nature 357,
553-562.
Dludla, S., le Roex, A.P., Gurney, J.J., 2006. Eclogite xenoliths from the Premier kimberlite,
South Africa: geochemical evidence for a subduction origin. South African Journal of
Geology 109, 353-368.
Eglington, B.M., Armstrong, R.A., 2004. The Kaapvaal Craton and adjacent orogens, southern
Africa: a geochronological database and overview of the geological development of the
craton. South African Journal of Geology 107, 13-32.
Ellis, D.J., Green, D.H., 1979. An experimental study of the effect of Ca upon garnet-
clinopyroxene Fe-Mg exchange equilibria. Contributions to Mineralogy and Petrology 71, 13-
22.
Farquhar, J., Bao, H.M., Thiemens, M., 2000. Atmospheric influence of Earth's earliest sulfur
cycle. Science 289, 756–758.
Field, M., Stiefenhofer, J., Robey, J., Kurszlaukis, S., 2008. Kimberlite-hosted diamond
deposits of southern Africa: A review. Ore Geology Reviews 34, 33-75.
80
Furlong, K.P., Chapman, D.S., 2013. Heat flow, heat generation, and the thermal state of the
lithosphere. Annual Review of Earth and Planetary Sciences 41, 385-410.
Gasparik, T., and Newton, R.C., 1984. The reversed alumina contents of orthopyroxene in
equilibrium with spinel and forsterite in the system MgO-Al2O3-SiO2. Contributions to
Mineralogy and Petrology 85, 186-196.
Gaul, O.F., Griffin, W.L., O’Reilly, S.Y., Pearson, N.J., 2000. Mapping olivine composition in
the lithospheric mantle. Earth and Planetary Science Letters 182, 223-235.
Gill, R., 2010. Igneous Rocks and Processes. A Practical Guide. Wiley-Blackwell, Chichester,
428 p.
Giuliani, A., Phillips, D., Woodhead, J.D., Kamenetsky, V.S., Fiorentini, M.L., Mass, R.,
Soltys, A., Armstrong, R.A., 2015. Did diamond-bearing orangites originate from MARID-
veined peridotites in the lithospheric mantle? Nature Communications, DOI:
10.1038/ncomms7837
Griffin, W.L., Carswell, D.A., Nixon, P.H., 1979. Lower-crustal granulites and eclogites from
Lesotho, Southern Africa. In: Boyd, F.R., Mayer, H.O.A. (eds.) The Mantle Sample:
Inclusions in Kimberlites and Other Volcanics. Proceedings of Second International
Kimberlite Conference, American Geophysical Union, Washington, p. 59-86
Griffin, W.L., O’Reilly, S.Y., Napatov, L.M., Ryan, C.G., 2003. The evolution of lithospheric
mantle beneath the Kalahari Craton and its margins. Lithos 71, 215-241.
Grutter, H.S., Gurney, J.J., Menzies, A.H., Winter, F., 2004. An updated classification scheme
for mantle-derived garnet, for use by diamond explorers. Lithos 77, 841-857.
Harte, B., 1977. Rock nomenclature with particular relation to deformation and recrystallization
textures in olivine-bearing xenoliths. Journal of Geology 85, 279-288.
Hasterok, D., Chapman, D.S., 2011. Heat production and geotherms for the continental
lithosphere. Earth and Planetary Science Letters 307, 59-70.
Heaman, L.M., Kjarsgaard, B.A., Creaser, R.A., 2003. The timing of kimberlites magmatism
in North America: implications for global kimberlite genesis and diamond exploration. Lithos
71, 153-184.
Helmstaedt, H., Gurney, J.J., 1997. Geodynamic controls of kimberlites – What are the roles of
hotspot and plate tectonics? Russian Geology and Geophysics 38, 492-508.
Jacobs, J., Pisarevsky, S., Thomas, R.J., Becker, T., 2008. The Kalahari Craton during the
assembly and dispersal of Rodinia. Precambrian Research 160, 142-158.
James, D.E., Niu, F., Rokosky, J., 2003. Crustal structure of the Kaapvaal craton and its
significance for early crustal evolution. Lithos 71, 413-429.
Jelsma, H., Barnett, W., Richards, S., Lister, G., 2009. Tectonic setting of kimberlites. Lithos
112, 155-165.
Johnston, D.T., 2011. Multiple sulfur isotopes and the evolution of Earth’s surface sulfur cycle.
Earth Science Reviews 106, 161–183.
Kamber, B.S., Whitehouse, M.J., 2007. Micro-scale sulfur isotope evidence for sulfur cycling
in the Late Archean shallow ocean: Geobiology 5, 5−17.
Kennedy, C.S., and Kennedy, G.C., 1976. The equilibrium boundary between graphite and
diamond. Journal of Geophysical Research 81, 2467-2470.
81
Köehler, T.P., Brey, G.P., 1990. Calcium exchange between olivine and clinopyroxene
calibrated as a geothermobarometers for natural peridotites from 2 to 60 kb with applications.
Geochemica et Cosmochimia Acta 54, 2375-2388.
Krogh, E.J., 1988. The garnet-clinopyroxene Fe-Mg geothermometer – a reinterpretation of
existing experimental data. Contributions to Mineralogy and Petrology 99, 44-48
Kuskov, O.L., Kronrud, V.A., 2006. Composition, temperature, and thickness of the lithosphere
of the Archean Kaapvaal Craton. Physics of the Solid Earth 43, 42-62.
Le Bas, M.J., Streckeisen, A.L. 1991. The IUGS systematics of igneous rocks. Journal of the
Geological Society 148, 825–833.
Levinson, A.A., 1997. Diamond sources and their discovery. In: Harlow, G.E. (eds.). The
Nature of Diamonds. Cambridge University Press, Cambridge, p. 72-104.
Lock, N.P., Dawson, J.B., 2013. Contrasting garnet lherzolite xenolith suites from Letseng
kimberlite pipes: Inferences for the northern Lesotho geotherm. In: Pearson, D.G., Grutter,
H.S., Harris, J.W., Kjarsgaard, B.A., O’Brien, H., Rao, N.V.C., Sparks, S. (eds.) Proceedings
of 10th International Kimberlite Conference, Vol. I, Springer, Berlin, Heidelberg, p. 29-42.
Lorenz, V., 1975. Formation of phreatomagmatic maar-diatreme volcanoes and its relevance to
kimberlite diatremes. Physics and Chemistry of the Earth 9, p. 17-27.
Luguet, A., Jaques, A.L., Pearson, D.G., Smith, C.B., Bulanova, G.P., Roffey, S.L., Rayner,
M.J., Lorand, J.-P., 2009. An integrated petrological, geochemical and Re-Os isotope study
of peridotite xenoliths from the Argyle lamproite, Western Australia and implications for
cratonic diamond occurrences. Lithos 112, 1096-1108.
Mapeo, R.B.M., Ramokate, L.V., Corfu, F., Davis, D.W., Kampunzu, A.B., 2006. The Okwa
basement complec, western Botswana: U-Pb zircon geochronology and implications for
Eburnean processes in southern Africa. Journal of African Earth Sciences 46, 253-262.
McCandless, T.E., 1999. Kimberlites: mantle expressions of deep-seated subduction. In:
Gurney, J.L., Pacsoe, M.D., Richardson, S.H. (eds.) Proceedings of the Seventh International
Kimberlite Conference, Vol. 2, p. 545-549.
Mitchell, R.H., 1986. Kimberlites – Mineralogy, Geochemistry, and Petrology. Plenum Press,
London, 442 p.
Mitchell, R.H., 1995. Kimberlites, Orangites, and Related Rocks. Plenum Press, London, 410
p.
Mori, T., Green, D.H., 1978. Laboratory duplication of phase equilibria observed in natural
garnet lherzolites. Journal of Geology 86, 87-97.
Morimoto, N., Fabries, J., Ferguson, A.K., Ginzburg, I.V., Ross, M., Seifert, F.A., Zussman, J.,
Aoki, K., Gottardi, G., 1988. Nomenclature of pyroxenes. American Mineralogist 73, 1123-
1133.
Nimis, P., Taylor, W.R., 1999. Single clinopyroxene thermobarometry for garnet peridotites.
Part I. Calibration and testing of a Cr-in-Cpx barometer and an enstatite-in-Cpx thermometer.
Contributions to Mineralogy and Petrology 139, 541-554.
O’Brien, H., 2015. Kimberlite-hosted diamonds in Finland. In: Maier, W.D., Lahtinen, R.,
O’Brien, H. (eds.) Mineral Deposits of Finland. Elsevier, Amsterdam, p. 345-375.
O’Neill, H.St.C., Wall, V.J., 1987. The olivine-orthopyroxene-spinel oxygen geobarometer, the
nickel precipitation curve, and the oxygen fugacity of the Earth’s upper mantle. Journal of
Petrology 28, 1169-1191.
82
O’Neill, H.St.C., Wood, B.J., 1979. An experimental study of Fe-Mg partitioning between
garnet and olivine and its calibration as a geothermometer. Contributions to Mineralogy and
Petrology 70, 59-70.
Pasteris, J.D., Boyd, F.R., Nixon, P.H., 1979. The ilmenite association at the Frank Smith Mine,
R.S.A. In: Boyd, F.R., Mayer, H.O.A. (eds.) The Mantle Sample: Inclusions in Kimberlites
and Other Volcanics. Proceedings of Second International Kimberlite Conference, American
Geophysical Union, Washington, p. 265-278.
Pearson, D.G., Canil, D., Shirey, S.B., 2003. Mantle samples included in volcanic rocks:
Xenoliths and diamonds. In: Holland, H., Turekian, K. (eds.) Treatise on Geochemistry
Volume 2, Elsevier, Amsterdam, p. 171-275.
Pollack, H.N., Chapman, D.S., 1977. On the regional variation of heat flow, geotherms, and
lithospheric thickness. Tectonophysics 38, 279-296.
Råheim, A., Green, D.H., 1974. Experimental determination of the temperature and pressure
dependence of the Fe-Mg partion coefficient for coexisting clinopyroxene and garnet.
Contributions to Mineralogy and Petrology 48, 179-203.
Rawlinson, P.J., Dawson, J.B., 1979. A quench pyroxene-ilmenite xenolith from kimberlite:
Implications for pyroxene-ilmenite intergrowths. In: Boyd, F.R., Mayer, H.O.A. (eds.) The
Mantle Sample: Inclusions in Kimberlites and Other Volcanics. Proceedings of Second
International Kimberlite Conference, American Geophysical Union, Washington, p. 292-299.
Roberts, M., McKinlay, T., 2017. Desert Gems – Botswana’s Major Mines. 11th International
Kimberlite Conference, Field Trip (1) Guide.
Robinson, D.N., 1975. Magnetite-serpentine-calcite dykes at Premier Mine and aspects of their
relationship to kimberlite and to carbonatite of alkalic carbonatite complexes. Physics and
Chemistry of the Earth 9, 61-70.
Rudnick, R.L., Nyblade, A.A., 1999. The thickness and heat production of Archean lithosphere:
constraints from xenolith thermobarometry and surface heat flow. In: Fei, Y., Bertka, C.M.,
Mysen, B.O. (eds.) Mantle Petrology: Field Observations and High Pressure Experimentation:
A Tribute of Francis R. (Joe) Boyd. The Geochemical Society, Special Publication 6, p. 3-12.
Schmitz, M.D., Bowring, S.A., de Wit, M.J., Gartz, V., 2004. Subduction and terrane collision
stabilize the western Kaapvaal craton tectosphere 2.9 billion years ago. Earth and Planetary
Science Letters 222, 363-376.
Schulze, D.J., 2003. A classification scheme for mantle-derived garnets in kimberlite: a tool for
investigating the mantle and exploring for diamonds. Lithos 77, 195-213.
Scott, B.H., Skinner, E.M.W., 1979. The Premier kimberlite pipe, Transvaal, South Africa.
Kimberlite Symposium II, Cambridge, Extended Abstracts.
Shirey, S.B., Shigley, J.E., 2013. Recent Advances in Understanding the Geology of Diamonds.
Gems & Gemology 49, 188-222.
Simon, N.S.C., Irvine, G.J., Davies, G.R., Pearson, D.G., Carlson, R.W., 2003. The origin of
garnet and clinopyroxene in “depleted” Kaapvaal peridotites. Lithos 71, 289-322.
Smith, C.B., 1983. Pb, Sr and Nd isotopic evidence for sources of southern African Cretaceous
kimberlites. Nature 304, 51-54.
Smith, C.B., Gurney, J.J., Skinner, E.M.W., Clement, C.R., Ebrahim, N., 1985. Geochemical
character of South African kimberlites: A new approach based on isotopic constraints.
Transactions of the Geological Society of South Africa 88, 267-280.
83
Sommer, H., Wan, Y., Kröner, A., Xie, H., Jacob, D.E., 2013. SHRIMP zircon ages and
petrology of lower crustal granulite xenoliths from the Letseng-la-Terae kimberlite, Lesotho:
further evidence for a Namaqua-Natal connection. South African Journal of Geology 116,
183-198.
Sparks, R.S.J., Baker, L., Brown, R.J., Field, M., Schumacher, J., Stripp, G., Walters, A., 2006.
Dynamic constraints on kimberlite volcanism. Journal of Volcanology and Geothermal
Research 155, 18-48.
Sparks, R.S.J., 2013. Kimberlite volcanism. Annual Review of Earth and Planetary Sciences,
41, 497-528.
Steifenhofer, J., Viljoen, K.S., Marsh, J.S., 1997. Petrology and geochemistry of peridotite
xenoliths from Letlhakane kimberlites, Botswana. Contributions to Mineralogy and Petrology
127, 147-158.
Stein, C., 1995. Heat flow of the Earth. Global Earth Physics, A Handbook of Physical
Constants. AGU Reference Shelf I, American Geophysical Union, Washington p. 144-158.
Stuwe, K., 2009. Geodynamics of the Lithosphere: Energetics: Heat and Temperature.
Springer-Verlag, Berlin, Heidelberg, 493 p.
Thomas, R.J., von Veh, M.W., McCourt, S., 1993. The tectonic evolution of southern Africa:
an overview. Journal of African Earth Sciences 16, 5-24.
Treloar, P.J., 1988. The geological evolution of the Magondi Mobile Belt, Zimbabwe.
Precambrian Research 38, 55-73.
Van der Spuy, D., 1984. The geology of the Frank Smith kimberlite pipe, Barkly West District.
Unpublished B.Sc. Hons. Thesis, University of Cape Town, Cape Town.
van Hunen, J., Moyen, J.F., 2012. Archean subduction: Fact or fiction? Annual Review of Earth
and Planetary Sciences 40, 195-219.
van Reenen, D.D., Roering, C., Smit, C.A., Barton, J.M., Jr., 1990. The granulite-facies rocks
of the Limpopo Belt, southern Africa. In: Vielzeuf, D., Vidal Ph. (eds.) Granulites and Crustal
Evolution. Kluwer, Dordrecht, the Netherlands, p. 257-289.
Vearncombe, S., Vearncombe, J.R., 2002. Tectonic controls on kimberlite location, southern
Africa. Journal of Structural Geology 24, 1619-1625.
Wainwright, A.N., Luguet, A., Fonseca, R.O.C., Pearson, D.G., 2015. Investigating
metasomatic effects on the 187Os isotopic signature: A case study on micrometric base metal
sulphides in metasomatised peridotite from the Letlhakane kimberlite (Botswana). Lithos 232,
35-48.
Walter, M.J., 2003. Melt extraction and compositional variability in mantle lithosphere. In:
Holland, H., Turekian, K. (eds.) Treatise on Geochemistry. Volume 2, Elsevier, Amsterdam,
p. 363-394.
Woodhead, J., Hergt, J. Phillips, D., Paton, C., 2009. African kimberlites revised: In situ Sr-
isotope analysis of groundmass perovskite. Lithos 112, 311-317.
Wörner. G. 1999. Lithospheric dynamics and mantle sources of alkaline magmatism of the
Cenozoic West Antarctic Rift System. Global and Planetary Change 23, 61–77.
Appendices
Appendix I: Major element chemistry results (in wt. %) for Letlhakane, Letseng, Premier and
Frank Smith xenoliths.
Appendix
I
Ta
ble
I.2
Ma
jor
ele
men
t ch
em
istr
y r
esu
lts
(in
wt.
%)
for
Letl
ha
ka
ne,
Lets
en
g,
Pre
mie
r a
nd
Fra
nk
Sm
ith
xen
oli
ths:
Ort
ho
pyro
xen
e
Sam
ple
N
a2O
CaO
NiO
K2O
MgO
FeO
MnO
TiO
2
A
l2O
3 C
r2O
3
SiO
2
V
2O
3
Tota
l
Mg#
Cr#
Letl
ha
ka
ne
502-O
PX
0.0
61.0
40.1
30.0
035.0
85.1
80.1
10.0
50.6
30.2
257.5
50.0
2100.1
92.3
19.0
511-O
PX
0.1
90.4
90.0
60.0
134.6
66.6
90.1
50.1
70.3
60.3
157.8
30.0
0100.9
90.2
36.6
512-O
PX
0.0
60.2
40.0
90.0
036.1
24.8
90.1
30.0
21.6
10.3
557.4
30.0
1101.0
92.9
12.7
513-O
PX
0.1
70.3
20.1
20.0
136.0
65.1
90.1
20.0
70.7
00.3
457.7
20.0
4100.8
92.5
24.5
523-O
PX
0.1
60.3
70.1
10.0
036.3
74.4
30.1
00.0
70.7
00.3
657.7
40.0
3100.4
93.6
26.0
529-O
PX
0.1
90.3
70.0
80.0
036.4
44.3
50.1
20.0
80.6
40.4
457.7
60.0
3100.5
93.7
31.7
530-O
PX
0.2
40.5
70.1
00.0
035.6
84.7
70.1
20.1
10.7
00.3
157.9
90.0
0100.6
93.0
23.1
531-O
PX
0.0
60.1
70.0
70.0
135.5
54.7
70.1
30.0
52.1
80.3
257.0
50.0
1100.4
93.0
9.1
550-O
PX
0.1
70.3
70.0
50.0
036.1
44.2
90.1
20.1
30.6
90.3
258.2
10.0
1100.5
93.8
23.8
551-O
PX
0.0
20.2
10.0
50.0
135.8
14.8
90.1
50.0
12.3
70.3
057.0
80.0
0100.9
92.9
7.8
Lets
en
g
LE
T4-O
PX
0.0
60.4
90.1
30.0
136.7
53.8
60.1
20.0
10.7
00.3
458.1
30.0
2100.6
94.4
24.7
LE
T8-O
PX
0.1
80.5
10.1
20.0
136.5
94.0
60.1
00.0
40.8
30.2
858.2
20.0
2101.0
94.1
18.5
LE
T14-O
PX
0.1
50.4
10.1
40.0
136.2
54.1
80.1
10.0
60.7
60.3
458.2
50.0
0100.7
93.9
23.3
Pre
mie
r
P5-O
PX
0.0
90.8
00.0
70.0
334.2
64.6
40.1
20.0
13.9
30.7
155.5
80.0
2100.3
92.9
10.8
P6-O
PX
0.0
80.5
50.1
60.0
130.3
111.9
00.2
40.0
91.2
80.2
155.3
80.0
2100.2
82.0
9.8
P14-O
PX
0.0
30.3
20.0
50.0
135.8
84.4
30.1
20.0
11.9
30.5
557.1
30.0
0100.4
93.5
16.1
P15-O
PX
0.1
30.3
60.1
30.0
236.3
34.3
20.0
80.0
21.0
20.4
057.4
50.0
2100.3
93.7
20.8
P17B
-OP
X0.1
90.3
60.0
70.0
336.1
04.6
40.1
20.0
51.1
30.3
857.8
10.0
0100.9
93.3
18.6
P18-O
PX
0.2
01.2
70.1
30.0
134.7
35.1
20.1
00.0
90.9
10.2
657.6
00.0
0100.4
92.4
16.3
Fra
nk
Sm
ith
XS
2-O
PX
0.2
51.1
50.1
00.0
034.1
76.1
50.1
10.2
51.0
00.1
557.0
70.0
2100.4
90.8
8.9
XS
4-O
PX
0.2
71.1
30.1
40.0
034.8
74.9
40.1
00.1
01.0
20.3
457.6
60.0
1100.6
92.6
18.0
XS
5-O
PX
0.1
60.4
90.0
80.0
036.8
24.1
30.1
00.0
10.6
30.3
458.0
80.0
2100.9
94.1
26.8
Appendix
I
Tab
le I.
3
Ma
jor
elem
ent
chem
istr
y re
sult
s (i
n w
t. %
) fo
r Le
tlh
aka
ne,
Let
sen
g, P
rem
ier
an
d F
ran
k Sm
ith
xen
olit
hs:
Clin
op
yro
xen
e
Sam
ple
N
a2O
CaO
NiO
K2O
MgO
FeO
Mn
O
Ti
O2
A
l2O
3
Cr2
O3
Si
O2
V2O
3
To
tal
Mg#
Cr#
Letl
ha
kan
e
502-
CP
X0.
7420
.11
0.08
0.04
19.6
42.
890.
100.
050.
990.
6354
.73
0.01
100.
092
.429
.8
512-
CP
X2.
8518
.76
0.02
0.00
15.6
71.
880.
060.
164.
302.
4853
.86
0.06
100.
193
.728
.0
513-
CP
X2.
1919
.23
0.05
0.01
16.7
52.
570.
090.
252.
502.
6654
.06
0.05
100.
492
.141
.7
523-
CP
X3.
5217
.38
0.04
0.02
14.9
72.
080.
110.
243.
133.
5954
.33
0.06
99.5
92.8
43.5
529-
CP
X3.
7317
.18
0.05
0.03
15.0
82.
040.
070.
302.
844.
2154
.40
0.03
100.
093
.049
.9
530-
CP
X1.
9317
.84
0.04
0.03
18.0
03.
720.
100.
542.
301.
3354
.25
0.08
100.
289
.628
.0
531-
CP
X0.
9023
.03
0.08
0.01
17.3
41.
350.
070.
171.
741.
6353
.65
0.04
100.
095
.838
.6
550-
CP
X3.
6917
.11
0.02
0.01
15.2
02.
150.
080.
403.
663.
0854
.96
0.04
100.
492
.636
.1
551-
CP
X1.
5522
.54
0.02
0.02
16.3
01.
450.
060.
033.
690.
9154
.19
0.04
100.
895
.314
.2
Lets
eng
LET4
-CP
X1.
0919
.49
0.02
0.02
19.2
72.
400.
110.
372.
620.
7954
.14
0.03
100.
393
.516
.8
LET1
4-C
PX
3.04
18.2
30.
060.
0216
.15
2.20
0.07
0.18
2.87
2.79
54.5
90.
0810
0.3
92.9
39.4
Pre
mie
r
P5-
CP
X0.
5624
.00
0.04
0.01
17.2
81.
310.
050.
012.
700.
7553
.64
0.01
100.
495
.915
.6
P6-
CP
X2.
5220
.40
0.06
0.00
14.4
44.
230.
100.
443.
670.
9253
.38
0.01
100.
185
.914
.4
P15
-CP
X2.
0620
.68
0.03
0.02
16.6
91.
960.
040.
042.
442.
0954
.62
0.03
100.
793
.836
.5
P18
-CP
X1.
2017
.23
0.07
0.06
20.3
03.
150.
110.
131.
650.
7855
.13
0.03
99.8
92.0
24.2
Fra
nk
Smit
h
XS2
-CP
X1.
6216
.49
0.07
0.05
19.5
84.
050.
120.
462.
190.
5554
.47
0.05
99.7
89.6
14.3
XS4
-CP
X1.
4617
.82
0.07
0.05
19.3
53.
310.
110.
271.
981.
1054
.30
0.04
99.9
91.2
27.2
Appendix
I
Tab
le I.
4
Ma
jor
elem
ent
chem
istr
y re
sult
s (i
n w
t. %
) fo
r Le
tlh
aka
ne,
Let
sen
g, P
rem
ier
an
d F
ran
k Sm
ith
xen
olit
hs:
Ga
rnet
Sam
ple
N
a2O
CaO
NiO
K2O
MgO
FeO
Mn
O
Ti
O2
A
l2O
3
Cr2
O3
Si
O2
V2O
3
To
tal
Mg#
Cr#
Letl
ha
kan
e
502-
GR
T0.
025.
590.
010.
0020
.82
7.15
0.31
0.23
19.9
04.
4641
.87
0.07
100.
483
.813
.1
513-
GR
T0.
064.
750.
020.
0120
.76
7.77
0.35
0.25
20.1
64.
7241
.95
0.02
100.
882
.613
.6
523-
GR
T0.
095.
340.
040.
0120
.36
6.73
0.39
0.26
18.7
46.
7441
.59
0.04
100.
384
.319
.4
550-
GR
T0.
084.
660.
000.
0021
.58
6.53
0.38
0.40
19.7
85.
3641
.98
0.03
100.
885
.515
.4
Lets
eng
LET4
-GR
T 0.
005.
410.
020.
0121
.62
6.22
0.36
0.00
19.8
76.
0542
.32
0.03
101.
986
.117
.0
LET8
-GR
T0.
045.
400.
040.
0221
.37
6.00
0.35
0.03
20.6
74.
6242
.38
0.05
101.
086
.413
.0
Pre
mie
r
P15
-GR
T0.
035.
220.
060.
0221
.29
6.69
0.36
0.03
21.5
73.
4741
.91
0.05
100.
785
.09.
7
P18
-GR
T0.
085.
050.
040.
0121
.82
6.60
0.29
0.53
20.1
03.
8641
.97
0.02
100.
485
.511
.4
Fra
nk
Smit
h
XS1
-GR
T0.
054.
940.
020.
0121
.00
8.47
0.29
1.26
20.7
91.
7141
.74
0.04
100.
381
.55.
2
XS2
-GR
T0.
134.
600.
020.
0321
.31
8.07
0.24
0.99
21.2
01.
6342
.03
0.06
100.
382
.54.
9
XS4
-GR
T0.
114.
860.
030.
0122
.08
6.50
0.25
0.44
20.4
13.
8042
.19
0.06
100.
785
.811
.1
XS5
-GR
T0.
045.
120.
010.
0021
.64
5.70
0.27
0.03
19.4
35.
9941
.76
0.04
100.
087
.117
.1
XS7
-GR
T0.
164.
970.
030.
0321
.21
8.34
0.25
1.29
20.9
21.
4842
.17
0.08
100.
981
.94.
5
Tab
le I.
5
Ma
jor
elem
ent
chem
istr
y re
sult
s (i
n w
t. %
) fo
r Le
tlh
aka
ne,
Let
sen
g, P
rem
ier
an
d F
ran
k Sm
ith
xen
olit
hs:
Sp
inel
Sam
ple
N
a2O
CaO
NiO
K2O
MgO
FeO
Mn
O
Ti
O2
A
l2O
3
Cr2
O3
Si
O2
V2O
3
To
tal
Mg#
Cr#
Letl
ha
kan
e
531-
SP0.
050.
000.
170.
0218
.21
11.5
20.
170.
1043
.42
25.8
70.
010.
1299
.773
.828
.6
551-
SP0.
000.
000.
240.
0119
.30
11.0
00.
160.
0148
.74
21.5
50.
010.
1010
1.1
75.8
22.9
Pre
mie
r
P5-
SP0.
020.
000.
120.
0318
.11
11.3
80.
110.
0043
.55
26.7
40.
070.
0610
0.2
P14
-SP
0.01
0.02
0.07
0.05
15.4
413
.14
0.26
0.03
26.7
844
.44
0.04
0.13
100.
4
Appendix II: Temperature (T, 0C) and pressure (P, kb) calculations for Letlhakane, Letseng,
Premier and Frank Smith xenoliths.
Appendix II
Table II.1
Locality
Sample 502 511 512 513 523 529 530 531
Type Gt-Hzb Hzb Hzb Gt-Lhz Gt-Lhz Lhz Hzb Sp-Hzb
ParagenesisOl, Gt,
Opx, CpxOl, Opx
Ol, Opx,
Cpx
Ol, Gt,
Opx, Cpx
Ol, Gt,
Opx, Cpx
Ol, Opx,
Cpx
Ol, Opx,
Cpx
Ol, Opx,
Cpx, Sp
TPRESET 1229 1000 1000 1078 1020 1000 1000 1000
PPRESET 53 40 32 41 36 47 45 40
T [BKN90] 1219 983 1072 1013 1019 1246 767
T [KB90] 1254 730 1082 893 1171 996 986
T [Krogh88] 1164 1040 1073
T [KroghRavna00] 1212 1058 1063
T [O'NW79] 1230 977 927
T [Harley84] 1142 981 933
T [EG79] 1193 1101 1115
T [Powell85] 1180 1086 1102
T [Wells77] 1107 891 953 876 868 1075 817
T [BM85] 1162 937 999 944 945 1151 770
T [OpxBK90] 1260 831 922 925 975 1056 812
T [NaPxBK90] 1197 782 1129 976 1030 1269 1078
T [OW87] 701 698
T [WS91] 791 846
T [Berman95] 1075 1006 997
T [BermanMod] 1101 1066 1091
T [Ballhaus91] 664 670
T [Taylor98] 1300 1058 1136 1038 1020 1262 954
T [Canil99] 1169 1376 1530
T [NimisTaylor00] 1213 983 852
T [NiRGP96] 1273 1606 2077
T [ZnRGP96]
P [BKN90] 60 45 40
P [BBG08] 53 41 36
P [KB90] 49 77 40 55 21 45 42
P [NG85] 59 47 44
P [MC74] 62 51 47
P [Al in Ol]
P [NimisTaylor00] 59 32 41 44 47 45 36
P [RGP96] 54 43 44
Temperature (T, 0C) and pressure (P, kb) calculations for Letlhakane, Letseng,
Premier and Frank Smith xenoliths
Letlhakane
Appendix II
Table II.1 (continued)
Locality
Sample 550 551 LET4 LET8 LET14 P5 P6 P14
Type Gt-Lhz Sp-Hzb Gt-Hzb Gt-Hzb Gt-Hzb Sp-Hzb Ol-Wbs Sp-Hzb
ParagenesisOl, Gt,
Opx, Cpx
Ol, Opx,
Cpx, Sp
Ol, Gt,
Cpx, Opx
Ol, Opx,
Gt
Ol, Cpx,
Opx
Ol, Opx,
Cpx, Sp
Ol, Opx,
Cpx
Ol, Opx,
Sp
TPRESET 1071 1000 1251 1000 1000 1000 1000 1000
PPRESET 41 40 51 40 45 27 37 40
T [BKN90] 1065 664 1242 1041 1048 687 642
T [KB90] 919 667 1072 1061 1119 883 856
T [Krogh88] 1116 1149 1178
T [KroghRavna00] 1149 1196 1191
T [O'NW79] 989 1024 1051
T [Harley84] 995 1009 1000
T [EG79] 1168 1184 1204
T [Powell85] 1157 1171 1196
T [Wells77] 900 736 1103 909 907 791 685
T [BM85] 990 680 1183 967 972 707 918
T [OpxBK90] 950 841 1048 1008 985 1057 1021
T [NaPxBK90] 1001 629 1045 1063 1013 1294 901
T [OW87] 687 753
T [WS91] 845 918
T [Berman95] 1050 1053 1105
T [BermanMod] 1164 1106 1199
T [Ballhaus91] 665 703
T [Taylor98] 1062 843 1288 1076 1075 927 802
T [Canil99] 969 1290 1599
T [NimisTaylor00] 891 1201 902
T [NiRGP96] 852 1448 2206
T [ZnRGP96]
P [BKN90] 45 57 39
P [BBG08] 41 51 36
P [KB90] 64 102 79 32 30 44 59
P [NG85] 48 58 41
P [MC74] 51 62 43
P [Al in Ol] 63 45
P [NimisTaylor00] 46 28 41 45 45 27 37
P [RGP96] 46 52 42
Letlhakane Letseng Premier
Appendix II
Table II.1 (continued)
Locality
Sample P15 P17B P18 XS1 XS2 XS4 XS5 XS7
Type Gt-Hzb Hzb Gt-Lhz Gt-Dun Gt-Hzb Gt-Hzb Gt-Hzb Gt-Dun
ParagenesisOl, Opx,
Cpx, GtOl, Opx
Ol, Opx,
Cpx, GtOl, Gt
Ol, Opx,
Cpx, Gt
Ol, Opx,
Cpx, Gt
Ol, Opx,
GtOl, Gt
TPRESET 897 1000 1350 1000 1345 1300 1000 1000
PPRESET 29 40 54 40 56 51 40 40
T [BKN90] 892 1344 1344 1296
T [KB90] 834 1344 1260 1213
T [Krogh88] 961 1258 1288 1328
T [KroghRavna00] 925 1339 1393 1414
T [O'NW79] 928 1286 1249 1281 1244 1051 1266
T [Harley84] 903 1245 1232 1211 1050
T [EG79] 1024 1282 1316 1342
T [Powell85] 1008 1273 1309 1339
T [Wells77] 864 1215 1185 1148
T [BM85] 848 1313 1296 1238
T [OpxBK90] 890 1326 1306 1272
T [NaPxBK90] 1052 1381 1368 1398
T [OW87]
T [WS91]
T [Berman95] 944 1192 1232 1303
T [BermanMod] 1006 1242 1301 1363
T [Ballhaus91]
T [Taylor98] 1018 1384 1355 1332
T [Canil99] 1723 1557 1385 1363 1401 1108 1515
T [NimisTaylor00] 812 1320 1285 1240
T [NiRGP96] 2660 2072 1621 1565 1744 1153 1820
T [ZnRGP96]
P [BKN90] 31 59 59 54 41
P [BBG08] 29 54 56 51 37
P [KB90] 38 55 71 66
P [NG85] 33 58 57 55
P [MC74] 34 65 63 59 47
P [Al in Ol] 79 7 56 70 23
P [NimisTaylor00] 37 60 63 54
P [RGP96] 31 56 53 51 44
Premier Frank Smith
