kimberlites, orangeites, and related rocks || geochemistry of orangeites
TRANSCRIPT
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GEOCHEMISTRY OF ORANGEITES
Most previous studies of the geochemistry of orangeites have been undertaken and interpreted on the premise that they are merely mica-rich varieties of kimberlite. Thus, compilations of the average trace element abundances in kimberlites, made by Wedepohl and Muramatsu (1979), Dawson (1980), Muramatsu (1983), and Mitchell (1986), are based upon combined data for both kimberlites and orangeites. Given the conclusion that kimberlites and orangeites are members of different petrological clans, such averages must now be seen as incorrect. Similarly, comparative discussions of the geochemistry of ''micaceous kimberlites" and "mica-poor kimberlites," as presented by Mitchell (1986), Fesq et al. (1975), and Kable et al. (1975), must now be viewed in a different light.
The earliest studies of the geochemistry of kimberlites and orangeites were restricted to establishing differences in their major element composition. These studies detennined that "the only apparent difference in composition between micaceous and basaltic kimberlites is that the former are somewhat richer in potassium oxide and alumina, and somewhat poorer in magnesia, as one would expect from their mineralogical constitution" (Wagner 1914, pp. 110-111). This conclusion has been reiterate~ in many subsequent studies, e.g., Dawson (1971, 1980), which have also emphasized the volatile-rich nature of both varieties of rocks and the high content of Cr and Ni coupled with high REB, Sr, and Ba contents. Tainton (1992) has provided the most recent significant discussion of the major element geochemistry ofkimberlites with respect to inter-element variation and the problems of alteration and contamination.
Detailed investigations of the trace element geochemistry of orangeites were not undertaken until the early 1970s. The studies by Fesq et al. (1975), Kable et al. (1975), Mitchell and Brunfelt (1975) and Gurney and Berg (1969) established that, compared to kimberlites, orangeites were richer in many incompatible elements and exhibited very highly fractionated REE distribution patterns. One important conclusion of the work of Kable et al. (1975), based upon NblP ratios, was the first suggestion that kimberlites and orangeites must be derived from mineralogically different mantle sources. A major conclusion ofthe initial geochemical studies was that orangeites and kimberlites represent mixtures of an incompatible element-rich melt with mantle-derived xenocrysts.
Subsequent to the isotopic studies (Smith 1983) which established "micaceous kimberlites" as a group of rocks with different origins from other "kimberlites," Smith et al. (1 985b ) proposed a new geochemical classification of "kimberlites" based upon
249 R. H. Mitchell, Kimberlites, Orangeites, and Related Rocks© Plenum Press, New York 1995
250 CHAPTER 3
isotopic criteria. The major and trace element compositions of the isotopically defined groups were shown by multivariate statistical methods to be significantly different. Isotopic group II kimberlites (orangeites) were noted to have higher abundances of Si02, K20, Pb, Rb, Ba, and light REE and lower abundances ofTi02 and Nb than isotopic group I rocks (archetypal kimberlites). This conclusion was predictable from general geochemical principles and the observed mineralogy. Smith et al.'s (1985b) study was especially valuable in that a wide spectrum of trace element abundances was obtained on a suite of hypabyssal facies rocks, carefully selected as being contamination-free and minimally altered.
Smith et al. (1985b) further divided isotopic group I kimberlites into "on craton" (IA) and "off-craton" (IB) subgroups; the latter were considered to be relatively enriched in Ti, P, Nb, Zr, Y. Unfortunately, only 33 samples were analyzed and the value of undertaking extensive multivariate statistical analysis on such a small sample is questionable. Not surprisingly, the IA-IB subdivision is not supported by the more extensive data set now available (see below).
The most significant advance resulting from the Smith (1983) and Smith et al. (1985b) studies was the conclusion that isotopic group II rocks originated from lithospheric sources, whereas isotopic group I rocks were derived from asthenospheric-like sources. Subsequent isotopic studies ofkimberlites and orangeites have confirmed these original observations (see 3.8).
Recently, the detailed studies of the Finsch (Fraser 1987, Fraser and Hawkesworth 1992) and Barkly West orangeites (Tainton 1992) have substantially increased the number of major, trace, and isotopic analyses of well-characterized suites of samples. These data sets are extremely valuable for assessing the geochemical variation within and between orangeite intrusions.
Despite the advances noted, our knowledge of the geochemistry of orangeites is still limited. Although new data are presented in this work for orangeites from Swartruggens, New Elands, and Star, there remains a paucity of data for orangeites from the Swartruggens, Kroonstad, Winburg, Boshof, and Prieksa areas. Determinations of the trace element and isotopic composition of the minerals comprising orangeites have not yet been undertaken.
This chapter summarizes the basic features of the major and trace element and isotope geochemistry of orangeites and provides comparisons with that ofkimberlites and olivine lamproites. The geochemistry of orangeites is not, compared with that of the mica-rich rocks occurring in West Greenland, termed "micaceous kimberlites" (Scott 1981, Larsen and Rex 1992). This omission is intentional, as the rocks are not orangeites and are more probably related to ultramafic lamprophyres than kimberlites.
3.1. CONTAMINATION AND ALTERATION
Investigation of the geochemistry of orangeites is subject to all of the contamination and alteration problems discussed by Mitchell (1986) and Clement (1982) with regard to kimberlites.
Because of the extremely high contents of Mg, Ni, Cr, Sr, Ba, REE, Zr, and Nb relative to their abundance in common crustal rocks, it is now considered that crustal
GEOCHEMISTRY OF ORANGEITES 251
contamination cannot be responsible for either the observed enrichment in these elements or the compositional variation within and between intrusions (Tainton 1992, Fraser 1987, Mitchell 1986). Tainton (1992) has also noted that the ZrlY (>10) and NblY (>4) ratios of orangeites are far in excess of those of possible crustal contaminants (ZrlY < 10, NblY < 3). Hence, incompatible element ratios may reflect those of their parental magma rather than mixing with crustal contaminants.
Xenoliths of all types of crustal rocks found in unevolved orangeites either show no effects of incorporation in the magma or are slightly thermally metamorphosed. The latter process leads, in limestones and shales, to the formation of concentric zones of recrystallization and slight metasomatism, recognizable by color banding. Shale and basalt xenoliths immersed in evolved orangeites commonly undergo extensive reaction and are replaced by phlogopite, amphibole, and sanidine. The formation of these minerals represents the response of the magma to xenolith assimilation by the precipitation of only the current or possible liquidus phases.
Various contamination indices have been devised to estimate the degree of contamination (or alteration) ofkimberlites (Ilupin and Lutz 1971, Clement 1982) and may be directly applied to orangeites. Of these, the most widely used is Clement's contamination index (CI), which is defined as the ratio of (Si02 + Ah03 + Na20)/(MgO + K20) wt%. For kimberlites, contamination indices close to unity are considered to indicate uncontaminated or fresh kimberlite. Clement (1982) has noted that many apparently contamination-free orangeites have CIs of 1-5. These indices are not a result of contamination but a consequence of high modal contents of phlogopite. CIs for contaminated orangeites are considered more likely to be greater than 1.5. However, CIs for apparently uncontaminated orangeites analyzed by Dawson (1987) range from 1.5 to 2.6.
Alteration of orangeites leads to the formation of chlorite and clay minerals from micas, serpentine from olivine, and the replacement of pyroxene and apatite by carbonates. Spinels and titanates may remain as residual minerals. The geochemical effects of alteration include loss of mobile alkali elements (K, Rb, Cs), decrease in the Mg/Si ratio, and increase in Ca content. The overall result is an increase in the contamination index.
Although altered rocks may have undergone significant changes in major element composition, the trace element signature of the more immobile elements may not have been substantially altered. Thus, strongly-altered rocks with high Ni and Cr contents coupled with high Ba, Sr, and REE might still be recognized as orangeite, kimberlite, or lamproite.
Geochemical characterization of altered rocks is an important component of many exploration programs for diamond-bearing rocks. Commonly, the only material available for the identification of a particular exploration target is highly-altered rock obtained by core drilling. A combination of trace element analysis and identification of relict resistant accessory minerals might be the only means of classifying such samples (Mitchell 1995).
Diatreme facies rocks are obviously particularly prone to contamination and alteration, and analysis of such material should be avoided. Studies of diatreme facies kimberlites have clearly shown that it is impossible to rid samples of microscopic crustal clasts (Dawson 1980). The bulk compositions of all diatreme facies rocks should be regarded as contaminated.
252 CHAPTER 3
3.2. PRIMARY MAGMA COMPOSITIONS
Fortunately, many orangeites occur as relatively-fresh hypabyssal intrusions, containing few crustal xenoliths. Hence, whole rock compositions can be expected to be representative of the crystal-phyric magma composition at the time of consolidation.
Petrographic studies (1.10) and major element compositional data (3.3) indicate that all orangeites have been contaminated with mantle-derived olivine xenocrysts. Rocks which are relatively xenocryst-free contain transported assemblages of macro crystal and microphenocrystal phlogopite (2.1.7) in addition to primary groundmass minerals which have crystallized in situ. The ratios of primary phlogopite to apatite, diopside, and carbonates vary widely, suggesting that pre- and post-intrusion crystal fractionation (flow differentiation, filter pressing, crystal settling) plays a significant role in controlling the final modal assemblage. It is this assemblage plus the amount ofxenocrystal olivine which controls the bulk composition of the rock.
The observation that the modal proportions of the primary groundmass minerals are not constant within a given intrusion demonstrates that this matrix cannot represent the bulk composition of the orangeite parental magma. If the matrix represented a relatively rapidly-quenched liquid, one would expect to find approximately constant proportions of primary minerals in all samples. This conclusion implies that if major and trace elements are subtracted from the whole rock composition in proportion to the amount of xenocrystal olivine present, the result will not approximate the composition of the orangeite parental magma. The "corrected" compositions can only represent those of random mixtures of microphenocrysts, ground mass, and mesostasis minerals.
It is concluded from the above observations that the bulk compositions of orangeites cannot represent those of the parental magma or of magma plus olivine xenocrysts. Unfortunately, glassy or aphyric margins to hypabyssal intrusions, which would provide a better estimate of magma compositions, are not found. Consequently, it is suggested here that we do not know the true compositions of primary or derivative orangeite magmas. This conclusion has very important consequences regarding attempts to model orangeite petrogenesis using techniques devised for rocks whose compositions approximate those ofliquids (Fraser and Hawkesworth 1992, Tainton 1992, Tainton and McKenzie 1994, Mitchell and Brunfelt 1975). The problem is discussed further in Sections 3.3.2 and 4.1.
3.3. MAJOR ELEMENT GEOCHEMISTRY
The principal features of the major element geochemistry of orangeite were established by Wagner (1914) and reiterated by Dawson (1980,1987). Subsequently, Smith et al. (1985b) demonstrated that orangeites have compositions statistically distinct from those of archetypal kimberlites. Detailed geochemical studies of the compositional variation within individual intrusions have been undertaken by Fraser (1987), Fraser and Hawkesworth (1992), and Tainton (1992). The latter study, together with that of Skinner et at. (1994), documented for the first time the geochemistry of relatively silica-rich orangeites from the BarkIy West and Prieska areas respectively.
GEOCHEMISTRY OF ORANGEITES 253
The rocks referred to in this chapter as "unevolved orangeites," are composed primarily of phlogopite, apatite, and carbonate; late-stage richterite and sanidine are absent, and olivine macrocrysts mayor may not be present. Representative examples include orangeites from Bellsbank, Sover, Swartruggens, New Elands, and Star. "Evolved orangeites" contain richterite and sanidine; examples are known from Pniel, Postmasburg, Kroonstad, and the Prieska areas (see 1.10).
3.3.1. Unevolved Orangeites
Tables 3.1 and 3.2 present the average major element composition of unevolved orangeites from several of the main occurrences. Notable aspects of their composition are the wide ranges in the concentration of many elements (MgO, CaO, P20S, K20), low Si02 and Na20 contents, and high MgO, K20, P20S, and volatile contents. As a consequence of high K20 coupled with low Ah03 and Na20 contents, the majority of the rocks are peralkaline [molar (K20 + Na20)/Ah03 > 1], ultrapotassic (molar K20/Na20 > 3), perpotassic (molar K20/Ah03 > 1), and ultrabasic.
The majority of the Fe contents have been detennined by X-ray fluorescence methods, and Fe203/FeO ratios are not well-established. However, these can be expected to be high given the common occurrence oftetraferriphlogopite in the orangeite groundmass. Given the prevalence of late-stage deuteric alteration in many orangeites, it is extremely unlikely that measured oxidation ratios will reflect the redox conditions of the parent magmas. Data given by Smith et al. (1985b) indicate that FeO and Fe203 contents of 16 diverse orangeites range from 2.88 to 5.10 wt% and 2.50 to 5.25 wt%, respectively. The majority of the samples have Fe203/FeO ratios> 1 (range 0.46-1.71, average 1.16). Orangeites from Finsch (Clement 1982) contain 3.52-4.66 wt% FeO and 3.18-7.68 wt% Fe203 and have Fe203/FeO ratios of 0.76-1.81 (average of 7 samples = 1.17). Five randomly selected orangeites analyzed by Dawson (1987) contain 3.11-4.58 wt% FeO and 1.94-5.85 wt% Fe203, with Fe203/FeO ratios of 0.50 to 1.88 and averaging 1.14.
Tables 3.1 and 3.2 demonstrate that most orangeites have high and extremely wide-ranging values ofloss on ignition (LOI), representing primarily the sum of H20 and C02. Increasing LOI commonly correlates positively with increasing CaO, reflecting the very high modal calcite and/or dolomite contents of some samples. Actual H20 and C02 contents quoted by Clement (1992), Dawson (1987), and Smith et al. (1985b) vary widely and depend upon the relative proportions of phyllosilicates to carbonates. Thus, different rocks from a given intrusion may be rich in H20 or C02. In all instances neither the volatile content (LOI) nor the H20/C02 ratio reflects that of the parent magma.
The sulfur and fluorine contents have been insufficiently investigated, and no data have been reported for chlorine or bromine. Dawson (1987) reports the F content of five orangeites from not detectable to 0.03 wt% F.
Seven determinations of S content (0.Q1-0.12 wt% S, average 0.05 wt% S) for Finsch orangeites have been given by Clement (1982). Five determinations of S03 reported by Dawson (1987) range from 0.03 to 0.74 wt% and average 0.37 wt%. The high modal contents of barite in many orangeites indicates that the parental magmas must contain significant amounts of S. The paucity of sulfides and the dominance of sulfate suggests that orangeite magmas are highly oxidizing relative to kimberlites and basaltic magmas.
Tab
le 3
.1.
Ave
rage
and
Ran
ge o
f Maj
or E
lem
ent C
ompo
siti
ons
of
Rep
rese
ntat
ive
Ora
ngei
tes
Wt%
S
war
trug
gens
F
insc
h B
ells
bank
S
over
Si0
2 36
.44
± 2.
98
30.0
0-40
.75
37.5
3 ±
3.13
27.6~1.93
33.0
2 ±
1.89
27
.59-
36.0
0 3
5.0
9±
1.7
2 32.8~.41
Ti0
2 1
.58
±0
.30
1.
28-2
.52
0.88
± 0
.20
0.60
-1.5
7 0
.74
±0
.13
0.
43-0
.97
1.0
6±
0.2
9
0.48
-1.6
4 A
I203
4
.02
±0
.87
2.
76-6
.03
3.34
±0.
90
1.62
-5.6
9 1.
64 ±
0.3
8 0.
91-2
.29
2.5
5±
0.7
0
1.30
-3.9
6 Fe
203
8.1
5±
0.8
5
6.17
-10.
30
7.99
±0.
75
5.68
-8.8
4 7
.77
±0
.37
6.
83-8
.56
7.7
8±
0.4
7
7.07
-9.2
1 M
nO
0.16
± 0
.03
0.1
2-0
22
O
.17±
0.07
0.
13-0
.46
0.16
± 0
.24
0.12
-0.2
4 0.
15 ±
0.0
3
0.09
-0.2
2 M
gO
21
.25
±4
.25
14
.80-
27.3
0 28
.18±
5.09
10
.44-
33.3
8 3
1.4
0±
4.2
7
20.9
7-39
.49
29
.02
±4
.53
20
.94-
36.4
7 Ca
o 8.
39 ±
3.7
5 2.
94-1
5.26
6.
54±
4.48
3.
29-2
4.48
6.
61 ±
2.0
7
3.49
-13.
41
6.49
± 2
.46
3.50
-11.
48
Na2
0 0
.24
±0
.10
0.
03-0
.58
0.21
±O
.l8
0.03
-0.7
4 0
.12
±0
.06
0.
01-0
.25
0.18
±0.
1O
0.03
-0.5
2 K
20
4.6
5±
1.0
0 3.4~.72
3.14
± 0
.76
0.81
-4.4
3 1
.72
±0
.55
0.
67-3
.16
2.91
± 1
.32
1.16
-5.7
9 P2
0S
1.34
± 0
.44
0.74
-2.2
5 0.
61 ±
0.19
0.
30-1
.18
1.41
±0
.52
0.
72-3
.31
0.68
± 0
.42
0.10
-1.6
2 L
OI
1O.4
5±2.
26
6.31
-15.
00
9.90
±3.
71
5.28
-21.
47
12
.99
±2
.14
9.
80-1
8.40
1
l.7
6±
1.4
2 8.
58-1
4.50
T
otal
96
.67
98.4
9 97
.58
97.6
7 PI
1.
35
Ll2
1.
26
1.35
U
Pf
12.7
5 9.
84
9.43
10
.64
PPI
1.25
1.
02
1.14
1.
24
(N)
18
30
35
31
LO
I = lo
ss o
n ig
niti
on. P
I. U
PI. a
nd P
PI a
re th
e pe
raIk
alin
ity, u
ltrap
otas
sic,
and
per
pota
ssic
indi
ces,
resp
ecti
vely
. Tot
al F
e is
exp
ress
ed a
s F~3. (
N) =
num
ber o
f sam
ples
. Dat
a so
urce
s: S
war
trug
gens
(t
his
wor
k, S
mit
h et
al.
1985
a,b)
; Fin
sch'
(Fra
ser
1987
, thi
s w
ork)
; Bel
lsba
nk (
fain
ton
1992
, Sm
ith
et a
l. 19
85a,
b, th
is w
ork)
; Sov
er (
fain
ton
1992
, thi
s w
ork)
.
~ ~ ~
GEOCHEMISTRY OF ORANGElTES 255
Table 3.2. Average and Range of Major Element Compositions of Representative Orangeites
Wt% Newlands New Elands Star
Si02 33.52 ± 1.85 29.80-36.31 36.64 ± 0.09 35.83-37.90 34.01 ± 1.09 32.10-35.30 TI02 0.62±0.13 0.47-0.91 l.32±0.16 1.08-1.43 1.27+0.18 1.01-1.61
Ah03 1.71 ±0.32 1.29-2.32 4.23±0.27 3.84-4.44 2.79 ± 0.47 2.13-3.66
Fe203 7.36±0.25 6.86-7.71 7.30± 1.45 5.13-8.22 8.59 ± 0.60 7.30-9.13 MnO 0.14±0.02 0.11-0.16 0.31 ±0.20 0.19-0.60 0.26 ±0.14 0.16-0.57 MgO 34.08 ±3.21 28.40-39.84 20.86 ± 5.79 12.30-25.04 25.34±5.66 18.80-31.70 CaO 6.12± 1.64 3.11-8.58 10.18 ± 5.71 6.15-18.65 7.92±2.65 4.90-12.30 Na20 0.11 ±0.05 0.03-0.23 0.19 ±0.05 0.13-0.25 0.17 ±0.1O 0.05-0.35 K20 1.02 ±0.36 0.52-1.65 4.73±0.36 4.27-5.12 2.95±0.27 2.51-3.48
P20S 1.13±0.24 0.65-1.52 1.22±0.18 1.00-1.41 0.82±0.23 0.39-1.16 LOI 12.42 ±2.19 8.71-16.80 11.43±2.86 9.48-15.68 12.89 ± 2.99 8.75-17.20 Total 98.23 98.98 97.01 PI 0.75 1.28 1.24 UPI 6.10 16.38 11.42 PPI 0.65 1.21 1.14 (N) 19 4 8
LO! = loss on ignition. PI, UPI, and PPI are the peralkalinity, ultrapotassic, and perpotassic indices, respectively. Total Fe expressed as Fe.!03. (N) = number of samples. Data sources: Newlands (Tainton 1992, Smith et al. 1985a, this work); New Elands (Smith et al. 1985b); Star (this work).
3.3.2. Mineralogical Controls on the Major Element Geochemistry
Petrographic examination reveals that unevolved orangeites are essentially mixtures of olivine, phlogopite, carbonate (calcite and/or dolomite), and apatite. The majority of the olivines are macrocrysts derived by the fragmentation of mantle-derived harzburgite or lherzolite xenoliths. Thus, they may be regarded as contaminants in the magma, whereas the other minerals are primary phases. Consequently, the bulk major element compositions of orangeites represent mixing lines between the composition of forsteritic olivine and orangeite primary phases. The relatively small amounts of primary olivine in many orangeites (LlO, 4.5.4) do not contribute significantly to the bulk composition. Tables 3.1 and 3.2 indicate that the P205 contents of orangeites are relatively low; hence, the bulk compositions of orangeites may be regarded as reflecting modal variations in olivine macrocrystal, phlogopite, and carbonate content.
Compositional data for unevol ved olivine macrocryst-rich orangeites from Sover and BelIsbank (Tainton 1992), when plotted in the ternary system MgO-K20-CaO (Figure 3.1), demonstrate clearly that bulk compositions are controlled by mixing of the assemblage phlogopite--carbonate with macrocrystal olivine. Different olivine control lines for each intrusion reflect their differing phlogopitelcarbonate ratios. Figure 3.1 does not reflect the variations in the calciteldolomite ratio known to occur in these suites of samples (Tainton 1992). However, the figure indicates that the presence of dolomite will move bulk compositions to relatively CaO-poor compositions for a given phlogopitelolivine ratio.
256
MgO
/ 5
CHAPTER 3
+ SOVER • BELLSBANK • SWARTRUGGENS o NEW ELANDS
35
o "0
40 \
45\
CoO CT
CoO CT
Figure 3.1. Compositions (wt%) of orangeites plotted in the ternary system MgO-K20-CaO. The diagram also shows compositional tie lines for mixtures of olivine (OL). phlogopite (PHL). calcite (Cf). or dolomite (DOL). Isocompositionallines show the ternary percentages of these minerals. e.g .• the 850L line shows the varying composition of ternary mixtures containing 85% forsteritic olivine with respect to changing calcitelphlogopite ratios. Compositions of olivine and phlogopite used in the calculations are from Table 2.25 (anal. 4C) and Table 2.1 (anal. 10). respectively. Data for Sover and Bellsbank from Tainton (1992); Swartruggens and New EIands from this work and Smith et al. (1985b).
GEOCHEMISTRY OF ORANGEITES 257
Similar relationships are evident (Figure 3.1) for relatively olivine-poor Swartruggens and New Elands orangeites. Note that several Swartruggens samples have bulk compositions that plot parallel to the phlogopite-calcite join. This agrees with petrographic observations that the rocks are essentially mixtures of phlogopite and calcite (see 1.10).
Figure 3.1 is interpreted to show that the bulk compositions of orangeites from Sover and Bellsbank cannot represent liquid compositions. The compositions found are the result of mixing of the crystal-laden magma, which formed the groundmass, with xenocrystal olivines. Addition of xenocrystal olivine implies that lherzolite-derived orthopyroxene, together with minor clinopyroxene and garnet, must have been added to the magma, unless the xenocrystal contaminants are entirely derived from dunites. Orthopyroxene xenocrysts have not been recognized in orangeites, implying that any lherwlite-derived enstatite must be completely assimilated by the magma either at its source or during transport. Assimilation of orthopyroxene will raise the silica content of the hybrid magma. The addition of the amounts of orthopyroxene typically found in lherwlite xenoliths (20-40 vol%) may lead to the formation of relatively siliceous evolved orangeites. However, lacking knowledge of the composition, volume, and temperature of the primary magmas involved, calculation of the potential compositions of hybrids is fraught with uncertainty.
Orangeites relatively-poorin olivine, such as occur at New Elands and Swartruggens, might have lost the majority of their load of xenocrystal olivines during periods of stagnation in the ascent of the magma. Alternatively, the magmas might not have been extensively contaminated at their sources. Regardless, the complex mica assemblage present in these rocks and the bulk compositional variation attributable to varying modal phlogopite/carbonate ratios, demonstrate that they were intruded as crystal-charged slurries with a minor fluid content.
The above observations suggest that the composition of the parental orangeite magma cannot be determined by the simple subtraction of olivine from the measured bulk compositions, as the proportions of phlogopite, apatite, and carbonate vary widely because of pre- and post-intrusion crystal fractionation. This conclusion has important implications regarding the interpretation of the trace element geochemistry of orangeite. For example, Fraser (1987) and Fraser and Hawkesworth (1992) have concluded that the variations in incompatible element abundances in the Finsch orangeites merely reflect the results of mixing varying amounts of peridotite contaminant with an unfractionated trace element-rich magma. However, early crystallizing apatite and diopside are common in many orangeites, and fractional crystallization processes involving these phases may playa role in determining the abundances of Sr and the REE abundances.
3.3.3. Evolved Orangeites
Table 3.3 presents the averages and ranges of composition of evolved orangeites from Sover North and Postmasburg (Tainton 1992), together with representative compositions of similar rocks from Pniel (Tainton 1992) and the Prieska area (Skinner et al. 1994). Compared to unevolved orangeites, the rocks are relatively rich in silica as a consequence of the presence of potassium feldspar and richterite. Other significant differences are their
258 CHAPTER 3
Table 3.3. Average and Range of Major Element Compositions of Representative Evolved Orangeites
Wt% 2 3 4 5
Si02 45.75 45.20-46.92 44.87 44.13-45.57 42.51 41.01 44.27 Ti02 1.87 1.79-2.03 1.29 1.23-1.39 1.18 1.24 1.91 AI20 3 5.75 5.53-6.10 7.41 7.12-7.87 4.08 6.43 6.26 FeZ03 8.20 7.92-8.40 8.80 8.71-8.98 8.26 7.67 6.26 MnO 0.12 0.11-0.13 0.17 0.15-0.20 0.13 0.11 0.14 MgO 20.92 19.74-21.96 14.89 14.54-15.47 27.99 20.40 18.23 CaO 4.86 4.42-5.16 10.36 10.12-10.79 4.04 6.15 7.02 Na20 1.12 0.82-1.56 0.83 0.61-1.06 0.60 0.36 0.68 K20 4.07 3.35-4.84 5.52 5.38-5.70 4.31 3.17 3.70 P20 S 0.78 0.55-1.07 0.71 0.41-1.09 0.43 1.06 1.09 LOI 4.25 3.12-5.09 3.29 2.64-3.95 4.26 8.42 5.51 Total 97.64 98.14 97.79 96.02 97.19 PI 1.09 0.99 1.39 0.63 0.82 UPI 2.39 4.38 4.73 5.79 3.58 PPJ 0.77 0.81 1.14 0.53 0.64 (N) 9 3 1
LOI = loss on ignition. PI, UPI, and PPJ are the peralkalinity, ultrapotassic, and perpotassic indices respectively. Total Fe is expressed as Fez03. (N) = number of samples. I = Sover North; 2 = Postmasburg 24/PK37; 3 = Pniel; 4 = Brandewynskuil; 5 = Slypsteen. Data sources: 1-3 (Tainton 1992); 4-5 (Skinner et QI. 1994).
relative enrichment in Ah03 and depletion in MgO and volatiles. Rocks from Sover North and Postmasburg have high Na20 contents which are considered by Tainton (1992) to result from low-temperature alteration ofleucite and sanidine to Na-zeolites. Data on the abundances of the individual volatiles constituting the LOI are not available. The paucity of carbonate in evolved orangeites (1.10) suggests that the major volatile component is H20.
The limited data available indicate that evolved orangeites are typically miascitic, not perpotassic, and only weakly ultrapotassic compared with unevolved orangeites (Table 3.3). The Pniel orangeite is anomalous in that it is agpaitic and perpotassic as a consequence of the high modal abundance of potassic richterite (Tainton 1992).
Tainton (1992) considers that the compositions of evolved orangeites, i.e., Si02-rich rocks, do not lie on extensions of the linear arrays defined by unevolved orangeites on plots of Si02, Ah03, CaO, and Na20 versus MgO (Figure 3.2). However, it should be realized that these plots mainly reflect variations in macrocrystal olivine content and not the compositions of evolving liquids. Hence, there is no a priori reason why any simple compositional relationship should exist between diverse orangeites in these bivariate plots. Moreover, Figures 3.2B and 3.2F appear to contradict Tainton's (1992) assertions.
3.3.4. Comparison with Kimberlites
Kimberlites show a remarkably wide range in their major element composition (Mitchell 1986, Smith et at. I985b, Gurney and Ebrahim 1973) as a consequence of differentiation and modal variations in their macrocrystaI and primary mineral contents (MitcheII 1986). Average compositions (Table 3.4) are unlikely to have any real geochemical significance but are useful for comparative purposes. MitcheII (1986) has noted
GEOCHEMISTRY OF ORANGEITES
50 -
-
~ 40-
CJ) -
30 -
2-5 -
-(\11-5 -o
i= -0-5 -
1-6
1-2 o
(\I o 0-8 Z
0-4
A
D
:~.~ 0 ••
OJ o
I I I I I I
c .i4 • •
•• • • D~.
Ql~O
I I I I I
II •
• D •
E
0-0 +---,--,----,
12
o 8 o U
4
o G
I I I I I I
10 20 30 40
MgO
8
6
2
9-5 -
-If) 8-5 -
0(\1 -
If 7-5-
-
B
o • o 0 AD ~Q 0
•• ~v~'o '·0 0
DO "" .~ .. . o
6-5~-~1-r-1-r-1-~1~1--~1
6 -
-
4-
-
2-5 -
-010 1-5-
~ 0-5 -
I I I I I I
10 20 30 40
MgO
259
Figure 3.2. Major element compositional variation of orangeites from the Barldy West (Bellsbank, Sover, Sover North, Pniel) and Postmasburg (PK35-37) regions (after Tainton 1992)_ 0 Be\lshank and Newlands;. Sover; 0 Pniel and PK35; • Sover North and PK36; II PK37_
that kimberlites may be considered to be undersaturated ultrabasic rocks (Si02 = 25-35 wt%) with low Ah03 contents «5 wt%) and low Na201K20 ratios «0.5)_ Calcite and dolomite are major minerals in most kimberlites; consequently volatile contents are high (> 10 wt%) and dominated by C02_ Kimberlites are typically potassic but not agpaitic_
Thus, the major geochemistry of kimberlites is, in many respects similar to that of unevolved orangeites_ Hence, major element compositions do not provide any simple means of distinguishing the two rock types. Smith et aI_ (l985b) have shown that K20 and Ti02 may provide the only effective discriminant when both elements exceed I wt%.
;J
Tab
le3A
. R
epre
sent
ativ
e C
ompo
siti
ons
of K
imbe
rlit
es a
nd O
livi
ne/M
adup
itic
Lam
proi
tes
Wt%
2
3 4
5 6
7 8
9
Si<
h 30
.00
31.9
9 34
.37
37.4
8 33
.86
33.9
2 42
.31
±2.
21
45.4
7 ±
1.16
39
.91
n0
2
1.72
2.
32
0.74
0.
38
1.77
1.
46
3.75
±0.
82
2.34
± 0
.32
2.89
A
h0
3
1.99
2.
68
1.04
2.
31
3.88
4.
26
3.92
±0.
87
8.8
9±
0.6
7
3.88
F
e20
3
5.23
5.
64
4.12
3.
88
10.4
8 8.
27
8.71
F
eO
3.32
3.
24
3.56
3.
40
8.27
±0.
54
5.99
± 0
.27
MnO
0.
16
0.16
0.
13
0.11
0.
17
0.16
0.
13
MgO
32
.49
32.4
4 38
.55
34.4
3 30
.67
21.9
3 24
.42±
3.56
11
.15±
0.94
27
.17
Cao
10.9
0 6.
71
7.03
2.
13
8.64
14
.60
5.0
0±
0.9
5
11.8
4 ±
1.79
5.
16
Na2
0
0.19
0.
05
0.19
0.
03
0.24
1.
00
0.50
± 0
.25
0.8
3±
0.1
5
0.32
K
20
0.
70
1.11
0.
80
0.65
0.
86
2.92
4.
01 ±
1.0
9 7.
75 ±
1.4
9 2.
69
P2
05
1.
89
1.51
1.
70
0.21
0.
80
1.43
1.
59 ±
0.4
8 2
.08
±0
.66
0.
35
LO}
10.7
1 11
.51
7.42
13
.87
8.94
9.
56
6.0
7±
1.8
8 3
.49
± 1
.19
8.63
99.3
0 99
.36
99.6
5 98
.88
100.
31
99.5
1 99
.84
99.8
3 99
.78
On-
crat
on k
imbe
rJit
es:
I D
e B
eers
; 2 W
esse
lton
; 3 D
utoi
tspa
n; 4
Jag
ersf
onte
in (1
-3; C
lem
ent
1982
; 4, S
mit
h et
al.
I 985
a). O
ff-e
rato
n ki
mbe
rlit
es: 5
Ber
seba
Res
erve
#2;
6 A
nis
Kub
ub (
5-6
Spr
iggs
19
88);
7 a
vera
ge o
f 10
5 E
llen
dale
oli
vine
Jam
proi
tes
base
d on
the
dat
a o
f Ja
ques
et
al.
(198
6);
8 av
erag
e o
f 6
Leu
cite
Hi1
Is m
adup
itic
lam
proi
tes
base
d on
the
dat
a o
f C
arm
icha
el (
l967
a);
9 ol
ivin
e-m
adup
itic
lam
proi
te, P
rair
ie C
reek
(Fra
ser
1987
). L
OI = l
oss
on ig
niti
on, t
his
is C
02
and
H2
0 in
kim
berl
ites
bu
t mai
nly
H2
0 in
lam
proi
tes.
~ to>
GEOCHEMISTRY OF ORANGElTES
t 5·0
4'0
-t. ..: 3'0 ~ N
Q 2'0 ~
I 1·0
0
• •• • • •
•
.... ~ . ••• •• • . ..
• •
KIMBERLITES • ON -CRATON • OFF-CRATON
• ORANGEITES
o 1'0 2·0 3·0 4·0 5·0 6·0 7'0
-- K20 Wt.% •
261
Figure 3.3. Ti02 versus K20 for on- and off-craton kimberlites and orangeites (after Smith et al. 1985b).
Thus, kimberlites are typically characterized by low K20 and high Ti contents, whereas orangeites exhibit the inverse relationship (Figure 3.3). The high Ti contents of kimberlites are attributable to the characteristically high modal abundances of groundmass Ti-rich spinels and perovskite. The low K contents reflect a paucity in phlogopite and the common presence ofkinoshitalite-rich groundmass micas.
As the mineralogical differences between evolved orangeites and kimberlites are so distinctive, it is not surprising that evolved orangeites are easily distinguishable from kimberlites on the basis of their higher Si02 and Ah03 and lower MgO, CaO, and C02 (LOI) contents.
3.3.5. Comparison with Lamproites
Lamproites show an exceedingly wide range in composition because of the numerous possible primary minerals coupled with extensive differentiation within the clan (Mitchell and Bergman 1991). Phlogopite and sanidine lamproites are mineralogically so different from all orangeites that comparison of bulk rock compositions is unnecessary. However, olivine and madupitic lamproites are low-silica, high-K20 rocks with some mineralogical and compositional affinities with evolved orangeites.
Dawson (1987) has noted that the compositions of unevolved orangeites are similar to those of olivine lamproites when the former are expressed on a CaC03-free basis (Table 3.4). Although this procedure does indeed result in bulk compositions resembling those of olivine lamproites, it should be realized that there are no petrogenetic grounds for undertaking this recalculation procedure. Excluding the constituents of one of the major primary groundmass minerals from the bulk composition is petrologically unsound. The inappropriateness of Dawson's (1987) approach may be realized by considering the analogous deduction of an amount of Si02 equivalent to the quartz in a granite and then claiming that granites are compositionally similar to syenites. Recalculation procedures
262 CHAPTER 3
of the type utilized by Dawson (1987) are valid only if the subtracted components are the constituents of secondary minerals introduced into a rock subsequent to consolidation.
Thus, it is suggested here that unevol ved orangeites are compositionally distinct from olivine lamproites, being poorer in Si02 and richer in CaO and C02. In contrast, evolved orangeites have bulk compositions closely resembling those of olivine lamproites (Table 3.4).
3.4. FIRST-PERIOD TRANSITION ELEMENTS
First-period transitional elements may be considered compatible trace elements in orangeites (and kimberlites) as they substitute for Fe and Mg in the principal early crystallizing primary phases: olivine (Sc, Ni, Co), phlogopite (Sc, Cr, Cu), spinel (Sc, V, Cr, Co, Zn), and pyroxene (Sc, Cr). Chromium occurs as a major element in primary groundmass magnesiochromites. Thus, Cr abundances are strongly controlled by the presence or absence of this mineral, e.g., the spinel-free Swartruggens orangeites are relatively poor in Cr compared to spinel-rich rocks from Finsch. Primary sulfides are very rare in orangeites and play no significant role in controlling the distribution of chaJcophiJe transition elements. Nickel sulfides are common in setpentinized olivine macrocrysts. However, the nickel forming these sulfides was originally present in solid solution as the liebenbergite molecule in olivine and has been merely redistributed during serpentinization.
Tables 3.5 and 3.6 show the abundance of first-period transition elements to be similar within and between orangeites. The only significant difference between unevolved and evol ved orangeites is with respect to their Ni contents. The lower Ni contents of the latter correlate to the relative paucity of olivine in these rocks.
Data presented by Tainton (1992) indicate no systematic correlations between Sc, V, Cr, Zn, and Cu with MgO. Ni correlates positively with Mg as expected, as olivine is the major host for Ni. Absolute abundances for Cr and Ni in orangeites are high compared to other mantle-derived basic magmas (Tainton 1992, Fraser 1987) and significantly higher than predicted for primary melts (Ni = 300-400, Cr = 400-500 ppm) from peridotite sources that have not been metasomatized (Fraser 1987).
Abundances of Sc, V, Co, Cu, and Zn are not very different from levels found in a wide variety of mantle-derived magmas, including kimberlite and olivine lamproite (Table 3.5; Mitchell 1986, Mitchell and Bergman 1991). Smith et aZ. (1985b) have suggested that kimberlites have, on average, lower Cr contents (1000 ppm) than orangeites (1800 ppm), a conclusion not supported by the data of Tainton (1992) or this work.
In summary, first-period transitional element abundances are of little use in distinguishing orangeites, kimberlites, and olivine lamproites. Ni abundances, being related to the presence of macrocrystal olivine, might permit estimation of the amount of contamination of orangeite magma with xenocrystal components (see 4.1.2). Abundances and/or ratios of abundances of first period transition elements, with the exception of Ti and Ni, are oflittle use in geochemical modeling of orangeite petrogenesis.
Sc
V
Cr
Ni
Co
Cu
Zn
(n)
Sc
V
Cr
Ni
Co
Cu
Zn
(n)
Tab
le 3
.5.
Ave
rage
and
Ran
ge o
f Fir
st-P
erio
d T
rans
itio
n E
lem
ent A
bund
ance
s (p
pm)
in U
nevo
lved
Ora
ngei
tes
Sw
artr
ugge
ns
20:
16-2
8 13
1: 9
1-15
2 12
07:3
15-1
424
1034
: 470
--17
42
73:5
4-96
29
:25-
34
84:7
9-88
4-
21
Sta
r
23:
18-3
5
2156
: 14
10--
2620
12
07:
895-
1570
78
:56-
89
8
Fins
ch
17:
12-2
3 13
2: 6
-285
17
65:
1100
--21
90
1214
:21-
1544
71
:61-
93
36
53 1-30
Kim
berl
ite
14:6
-38
100:
21-7
60
893:
430-
-255
4 96
5:47
1-18
00
65: 9
-125
93
: 6-1
320
69:
10--2
87
Bel
lsba
nk
Sov
er
22:
11-3
9 16
:2-2
6 72
: 41-
102
82:2
6-18
0 16
70:
1130
--225
1 18
52:9
75-2
865
1396
: 573
-202
2 12
53:6
48-9
20
96:
87-1
121
80:8
3-92
21
: 7-
50
21: 2
-49
53:4
6-62
71
: 41-
409
6-48
3-
31
WK
-01-
lam
proi
te
21: 9
-39
82:2
0--2
67
1014
: 379
-170
3 96
8: 4
01-1
500
69:3
1-92
55
:39-
93
73:5
8-10
7
New
land
s
22:
16-3
2 48
:30-
-77
1891
: 16
16-2
861
1450
:812
-174
9 71
: 65-
84
22:8
-38
43:4
0--4
8 5-
19
New
Ela
nds
23:
19-2
5 10
5:83
-148
15
14:
1430
--16
41
1036
: 902
-134
8 69
:62-
76
38:3
2-42
83
:75-
87
4
PC
-OI-
Iam
proi
te
15:
14-1
6 46
:27-
-68
1447
: 13
91-1
500
1356
: 12
85-1
443
96:9
5-97
52
:47-
57
73:7
1-74
Dat
a so
urce
s: S
war
trug
gens
(th
is w
ork,
Sm
ith e
t al.
1 985
b); F
insc
h (t
his
wor
k, F
rase
r 19
87);
Bel
lsba
nk (t
his
wor
k, T
aint
on 1
992)
; Sov
er (t
his
wor
k, T
aint
on 1
992)
; New
land
s (T
aint
on 1
992)
; New
E
land
s (S
mit
h et
aJ.
1985
b);
Sta
r (t
his
wor
k).
Ave
rage
kim
berl
ite (
Mit
chel
l 19
86);
ave
rage
Wes
t K
imbe
rley
oli
vine
lam
proi
te (
WK
-OI-
Iam
proi
te),
Jaq
ues
et a
l. (1
986)
; av
erag
e Pr
airi
e C
reek
la
mpr
oite
(P
C-O
l-Ia
mpr
oite
), F
rase
r (1
987)
.
i ~ ~ I ~
264 CHAPTER 3
Table 3.6. Abundances (ppm) of First-Period Transition Elements in Evolved Orangeites
2 3 4 5
Sc 11: 10--13 12:9-15 9 v 161: 139-182 104: 80--146 159 13I 231 Cr 1239: 1082-1399 1155: 970--1219 1965 1412 1267 Ni 495:470--524 960: 811-1146 1266 972 676 Co 77 69 Cu 25:24--26 31:23-36 31 28 54 Zn 60:51--65 77:72-92 60 76 86 (11) 3 9
1 ~ Postmasburg 24/P37; 2 ~ Sover North; 3 ~ Pniel; 4 ~ Brandewynskuil; 5 ~ Slypsteen. Data sources: 1-3 Tainton (1992); 4-5 (Skinner et al. 1994).
3.5. INCOMPATIBLE ELEMENTS
The incompatible trace elements (Sr, Ba, Zr, Nb, REE, Rb, Th, etc.) are usually defined as elements having solid/liquid distribution coefficients for common rockforming silicates of approximately zero. They are strongly partitioned into the liquid phase during partial melting of Iherzolitic sources and preferentially concentrated in derivative liquids during crystal-liquid fractionation processes. Their abundances and inter-element ratios are commonly used to infer the nature and degree of partial melting of magma sources. Much of the geochemical lore pertaining to incompatible trace elements has been deri ved from studies of basalts and related rocks. Parental magmas to these rocks are thought to be derived from simple lherzolitic mantle source rocks. Metasomatic phases enriched in incompatible elements have been postulated as being present in lherzolitic sources of magmas of more extreme compositions, e.g., melilitite, kimberlite. In such cases the incompatible element-rich phases are usually considered to be completely consumed during the partial melting episode that gave rise to the incompatible element-rich magma.
As a caveat to the above, it should be realized that if incompatible element -rich phases are not consumed during melting, then the elements in question are no longer incompatible. In these instances, extensions of hypotheses derived to explain the trace element distributions of basaltoid rocks to kimberlites and orangeites (Fraser 1987, Fraser and Hawkesworth 1992, Tainton 1992, Tainton and McKenzie 1994) are unlikely to be realistic.
In common magmas the elements may behave as incompatible elements throughout most of their crystallization interval. In the case of magmas of extreme composition early-forming liquidus phases include many minerals which have trace element solid/liquid distribution coefficients greater than zero, e.g., REE and Sr in apatite or Rb, Ba, and Cs in phlogopite. In these instances the elements in question are compatible and their abundances may be significantly affected by crystal-liquid fractionation and/or crystal accumulation. The latter process is particularly important regarding orangeites composed principally of phlogopite.
GEOCHEMISTRY OF ORANGElTES 265
In some instances incompatible elements might not be removed from the liquid until the later stages of crystallization of the groundmass. Therefore whole rock analysis may provide reasonable estimates of their abundances and ratios. The mixing of significant amounts of xenocrystal olivine with orangeite and kimberlite magmas implies that absolute abundances of incompatible elements are of little geochemical significance unless the magnitude of intra- and inter-intrusion differences are expressed against some normalizing index or are extremely large (Le., differ by a factor of 10 or more).
Importantly, studies oflamproites and kimberlites have suggested that simple garnet Iherzolites are not suitable source rocks for their parental magmas (Foley 1990, 1992a, Edgar 1987, Mitchell and Bergman 1991, Mitchell 1986, Smith et al. 1985b), and their sources may retain a residual mineralogy in which many so-called incompatible elements are actually compatible. Recognition of this aspect of magma genesis and the possibilities of compatibility during much of the crystallization interval have important ramifications regarding the interpretation of the trace element geochemistry of kimberlites and orangeites.
Tables 3.7 and 3.8 give abundances of incompatible elements in orangeites and provide some comparisons with kimberlites and lamproites. In some instances the difference in abundance is significant and useful in discriminating between rock types. However, all of these geochemical data should be regarded in the context of the above comments, realizing that abundances measured do not reflect those of the actual magma, as they have been reduced or "diluted" by mixing with macrocrystal olivine.
3.5.1. Alkaline Earths
In unevolved orangeites, Ba is hosted primarily by phlogopite, and also occurs as a major element in Ba-carbonates, hollandites, and barite. Sr is hosted primarily by apatite, perovskite, and late-stage carbonates. Ba and Sr exhibit wide ranges in their abundances within and between orangeites (Table 3.7). High and widely ranging Ba abundances, e.g., Swartruggens, reflect the presence of abundant late-stage, irregularly distributed barite. High Sr abundances, e.g., Sover, reflect the presence of Sr-carbonates and phosphates. BalSr ratios range from 0.6 to 4.8 (Table 3.9). There is no correlation between Ba and Sr abundances, and both elements are regarded as not being strictly incompatible.
Evolvedorangeites have Sr and Bacontents (Table 3.8) and BalSrratios (1-2) similar to unevolved types.
Unevolved orangeites are richer in Ba and Sr than kimberlites and poorer in Ba than olivine and madupitic lamproites (Table 3.7). Evolved orangeites are poor in both Ba and Sr compared to the latter.
3.5.2. Second- and Third-Period Transition Elements
3.5.2.1. Zirconium and Hafnium
Zr and Hf are concentrated in the groundmass of orangeites and hosted primarily by late-crystallizing zirconium silicates. There is a wide range of Zr and Hf abundances within and between orangeites (Table 3.7), and the highest contents are found in rocks in which groundmass Zr minerals are most abundant, e.g., Swartruggens and New Elands. Kable et al. (1975) report low Zr and Hf abundances in rocks of the Bellsbank Water
~
Thb
le 3
.7.
Ave
rage
and
Ran
ge o
fBa,
Sr,
Zr,
Hf,
Nb,
Ta,
Th
, an
d U
Abu
ndan
ces
(ppm
) in
Une
volv
ed O
rang
eite
s
Swar
trug
gens
F
insc
h B
eIls
bank
S
over
N
ewla
nds
New
Ela
nds
Ba
5183
: 13
90-1
6300
14
67:2
90-2
590
3439
:644
-594
2 24
42:7
92-5
480
3351
:557
-853
6 18
95:
1606
--21
67
Sr
1139
: 51(
}"28
80
738:
416-
1423
14
14: 6
01-2
484
1127
: 48(
}..6
591
1261
:545
-263
0 13
56:
1184
-148
7 Z
r 40
1:25
6-10
60
184:
11(
}"35
9 29
1: 1
20-5
98
214:
53-4
96
193:
129
-395
40
0: 3
19-4
72
Hf
10:6
-26
5: 3
-10
8:4-
15
5:3
-8
4:3
-6
12:9
-15
Nb
143:
85-
258
51: 2
-93
168:
108
-289
97
:46-
161
139:
95-
169
111:
94-1
20
Ta
8: 5
-13
3: 2
--6
14:7
-26
9:4
-14
9:
7-12
7
:5-8
T
h 26
: 16
-42
9:5-
16
45:2
0-74
30
:6-5
1 33
:21-
44
27:2
4-30
U
7:
5-17
3
:2-4
7:
3-1
4 3:
]-5
5:
3-9
5:
4--6
(n
) 17
21
-30
25-4
8 15
-31
15-1
9 4
Star
SA
K
ZK
K
IMB
K
Wes
t Kim
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ey
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cite
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s Pr
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reek
Ba
4370
: 31
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57oo
85
0 97
3: 1
95-1
652
833:
164-
2292
10
584:
3007
-413
78
9831
:431
9-14
400
1971
: 16
24-2
50
Sr
1808
: 13
6(}.
.255
0 10
20
468:
186
--829
10
54:3
46-2
428
1325
:725
-188
9 38
60: 2
581-
4904
10
94:9
72-1
284
Zr
194:
165
-277
38
5 12
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-164
27
9: 1
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717
1167
: 564
-374
0 13
02:
1152
-155
7 71
8:66
2-74
5 H
f 7:
4-1
3 3:
2.5
-3.7
6:
3-15
38
:32-
49
42:3
8-43
17
: 16
-18
Nb
134:
119
-173
21
0 15
3: 1
32-1
93
150:
37-3
46
186:
104
-309
99
: 79-
137
101:
87-
112
Ta
10:5
-18
14:1
1-19
9:
2-23
10
: 6-1
4 6
:5-7
6:
5-7
T
h 28
: 14
-50
27
15:1
4-18
19
:9-5
57
: 18
-98
37:3
3-45
12
: 11
-14
U
6 3:
3-5
2: 1
-0
9: 8
-9
3:2
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(n)
4-8
10
23
18-3
5 9-
109
3-10
3-
5
Dat
a so
urce
s fo
r or
ange
ites
as
in T
able
3.5
. SA
K =
Sout
h A
fric
an k
imbe
rlit
e (S
mith
et a
l. I 9
85b)
; ZK
:= Z
aire
kim
berl
ites
(Kam
pata
199
3); K
IMB
K =
Kim
berl
ey g
roup
kim
berl
ites
(G
emen
t 19
82,
this
wor
k);
aver
ages
for
the
Wes
t Kim
berl
ey, L
euci
te H
ills
, and
Pra
irie
Cre
ek la
mpr
oite
s fr
om M
itch
ell
and
Ber
gman
(19
91).
~ c.>
GEOCHEMISTRY OF ORANGElTES 267
Table 3.8. Abundances (ppm) ofIncompatible Elements in Evolved Orangeites
2 3 4 5
Ba 2234: 1080--29292002: 1902-2077 1354 1929 1479 Sr 1293: 1404-1500 795: 624-1120 624 1445 1072 'Ix 554:516-550 305:233-425 171 297 191 Hf 19 Nb 63: 62-70 54:30--78 30 67 75 Ta 5:5-6 4:3-5 2 Th 17: 12-17 8: 7-10 4 U 3:2-5 1: 1-2 2 (n) 1-9 3
1 = Sover North; 2 = Postmasburg 24/PK37; 3 = Pniel; 4 = Brandewynskuil; 5 = Slypsteen. Data sources: 1-3. Tainton (1992); 4-5. Skinner et al. 1994.
Table 3.9. Averages and Ranges of Ratios of Incompatible Elements in Orangeites
Swartruggens Fiosch Bellsbank Sover
Rb/Sr 0.21: 0.05-0.36 0.20: 0.10--0.33 0.08: 0.03-0.20 0.18: 0.08-{).40 ZrlNb 2.75: 2.23-4.11 3.70: 2.27-5.69 1.71: 0.76-2.85 2.16: 0.13-3.75 BalSr 4.76: 1.55-9.67 1.97: 0.39-3.06 2.62: 0.39-5.53 2.85: 0.65-7.46 PblCe 0.08: 0.05-0.09 0.09: 0.04-0.14 0.06: 0.04-0.11 0.05: 0.02-0.08 NbIU 20.9: 15.2-28.15 18.7: 13.3-27.3 24.9: 11.7-38.9 29.3: 16.8--58.1 NbIY 7.49: 5.19-11.58 5.28: 3.78-8.86 10.3: 6.1-18.6 6.18: 4.08-8.82 KlRb 203: 169-255 197: 126-342 134:83-167 149: 101-166 ZrlHf 40.9: 33.8-43.9 38.2: 34.1-41.3 39.2: 20.4-50.1 34.1: 24.5-43.0 Nb/Ta 18.4: 15.6-21.6 14.4: 12.1-17.5 14.0: 7.8-23.3 11.6: 6.2-18.9 KlNb 283: 177-439 565:243-937 81:44-129 243: 142-334 KlBa 10.4: 2.3-23.5 23.0: 6.5-76.2 6.3: 1.2-33.5 12.9: 3.2-35.8 La/Yb 116: 55-192 88: 54-115 243:92-483 179:83-267 LalNb 1.52: 1.17-2.11 1.27: 0.99-1.59 1.46: 0.91-1.86 1.82: 1.16-2.69 CeIY 22.1: 15.6-41.4 14.0: 11.4-19.1 27.6: 16.7-42.8 21.5: 14.5-30.0 (n) 4-17 21-30 27-35 28--31
Newlands NewElands Star SoverNorth
Rb/Sr 0.05: 0.02--{).OS 0.12: 0.10--{).13 O.OS: 0.05--{).1O 0.11: O.OS--{).13 'IxlNb 1.41: 0.79-2.00 3.61: 2.80--3.93 1.45: 1.12-1.65 8.06: 7.4S-8.91 BalSr 2.67: 0.67-4.79 1.29: 1.26-1.50 2.53: 1.22-3.68 1.54: 0.67-2.08 PblCe 0.05: @I-O.09 0.09: 0.08-0.11 0.06: 0.05-0.07 NbIU 28.7: 15.1-38.2 22.8: 17.7-29.0 23.8: 15.0--31.3 NbIY 11.7: 9.5-16.0 5.65: 5.37-6.34 7.76: 5.67-10.2 3.47: 3.19-4.14 KlRb 147: 110--174 243:232-262 187: 169-210 213: 158--271 'IxlHf 46.4: 35.4-58.9 33.5: 29.1-43.S 32.3 Nb/Ta 15.1: 12.8-21.1 14.7: 9.5--26.1 13.3: 11.2-16.7 KINb 59.7: 30.9-100.3 359:295-433 184: 167-212 496:334-603 KlBa 3.95: 0.72-13.1 21.0: 18.8--23.5 5.85: 3.88--7.72 17.1: 10.9-34.6 La/Yb 266: 185-382 167: 144-188 263: 185-338 129: 108-138 LalNb 1.48: 1.15-1.83 1.87: 1.78-2.00 1.40: 0.99-1.84 2.17: 1.86-2.50 CeIY 30.9: 25.3-37.3 19.3: 18.7-20.9 20.8: 17.8-27.1 14.0: 10.3-15.5 (n) 15-19 4 4-8 1-9
(n) = numberofsamples. Data sources: Swartruggens (Smithet al. I 985b. this work); Finsch (Fraser 1987. this work); Bellsbank (Tainton 1992. Kable et al. 1975. this work); Sover (Tainton 1992. this work); Newlands (Tainton 1992); New Elands (Smith et al. 1985b. this work); Star (this work); SoverNorth (Tainton 1992).
268
t 1000-
E 500-Q. Q. -~
N
CHAPTER 3
I I I
5 10 15
-- Hf (ppm) •
Figure 3.4. Zr versus Hf for orangeites. Data sources: Bellsbank (Tainton 1992, this work); Sover (Tainton 1992); Swartruggens (this work); Star (this work); Finsch (Fraser 1987).
Fissure relative to the Main and Bobbejaan dikes; this observation is significant if the Water Fissure rocks are evolved orangeites (Tainton 1992). Zr/Hf ratios of 24-50 are similar to those of a wide variety of mantle-derived rocks, including lamproite and kimberlite (Mitchell 1986, Mitchell and Bergman 1991). The coherent behavior ofZr and Hf on logarithmic plots of abundances (Figure 3.4) indicates that these elements are highly incompatible in orangeites and their ratios are unaffected by fractional crystallization or hybridization. Thus, Zr/Hf ratios may reflect those of the magma sources. Similar ratios in all orangeites suggest that these elements are hosted by only one phase in the mantle sources of the parental magmas.
Zr abundances of all orangeites are, on average, significantly less than those of olivine lamproites (Tables 3.7, 3.8). In particular, the low Zr contents of evolved orangeites provide a means of discriminating between these rock types and olivine lamproites of similar major element composition. There are no significant differences in the Zr (and Hf) contents of orangeites and kimberlites (Table 3.7).
3.5.2.2. Niobium and Tantalum
Nb and Ta are concentrated in the groundmass of orangeites where they substitute for Ti in late-stage rutile, Mn-ilmenite, and hollandite. Abundances are widely variable within and between intrusions. Finsch appears to be depleted in both elements relative to other unevolved orangeites. Average NblTa ratios of 11-14 are similar in all intrusions, although wide ranges are apparent within individual intrusions. Evolved orangeites are poor in Nb and Ta relative to unevolved types (Table 3.8, Figure 3.5). Significantly, the Water Fissure is low in Nb and Ta relative to unevolved Bellsbank orangeites (Kable et al. 1975), although Nb/Ta ratios (approximately 20) are similar (Figure 3.5).
GEOCHEMISTRY OF ORANGElTES
-E c. c. -,g Z
100
5
• BELLSBANK + SOVER • STAR o FINSCH
1992
• SWARTRUGGENS
10 15 20
To (ppm) •
269
Figure 3.5. Nb versus Ta for orangeites. Data sources: Bellsbank (Kable et al. 1975, Tainton 1992); Sover, Sover North, Pniel, and Postmasburg (Tainton 1992); Star and Swartruggens (this work); Finsch (Fraser 1987).
Figure 3.5 shows significant discrepancies exist between the data of Tainton (1992) and Kable et al. (1975) for the Bellsbank unevolved orangeites. Ta abundances given by Tainton (1992) and obtained by ICP-MS are much higher than the INAA data of Kable et al. (1975). Nb and Ta show only a weak correlation, and NblTa ratios vary considerably (Table 3.9). Similar trends are seen in Tainton's (1992) data for Sover. In contrast, orangeites from Finsch and Swartruggens, whose Ta content was determined by INNA, show greater coherence in their NblTa ratios, and they plot in Figure 3.5 close to the Nbrra ratio defined by the Bellsbank samples analyzed by Kable et al. (1975). Star orangeites resemble Sover in their Nb and Ta (lNAA) abundances. Further study is required to determine whether the differences noted above are real or due to inaccurate determination ofTa.
The majority of unevolved orangeites does not differ significantly in Nb and Ti contents relative to those ofkimberlites or olivine lamproites (Table 3.7, Figure 3.6). The contention of Smith et al. (1985b) that orangeites are poorer in Nb (120 ppm) than are kimberlites (210 ppm) is not supported by the larger data base now available (Table 3.7, Figure 3.6). Orangeites from Bellsbank are much richer in Nb than most low-Ti02 kimberlites. Only "off-craton" Namibian kimberlites (Spriggs 1988) have high Nb and low Ti02 contents. The spread of data in Figure 3.6 for Bellsbank and Sover is related to mixing with macrocrystal olivine (Tainton 1992).
Mineralogical observations suggest that the controls on Ti and Nb distribution are different in orangeites and kimberlites. Much of the Ti in kimberlites is bound in the abundant groundmass perovskite and spinels. Nb in kimberlites is carried primarily by
270
E 200 a. a.
.tl Z
100
o o
BELLSBA~"'''''7 ",. I
/ ISWR I I
•
., . , • 1 .. ;I , nl 'I. Y . "-... l :. , .. :.
• ? SOVER ..... '" .~ SOVER
/F ,.. ~
• I EO PNIELI W POSTMASBURG
1·0 2'0
• • • •
• • • •• •
NORTH
3·0
Ti02 wt. % •
CHAPTER 3
•
4'0
Figure 3.6. Nb versus Ti02 for orangeites and kimberlites. Data sources for orangeites as in Figure 3.5. Data for kimberlites from South Africa (Smith et al. 1985b, Clement 1982) and Namibia (Spriggs 1988). SWR = Swartruggens; F = Finsch; EO = evolved orangeites.
perovskites, although kimberlites rich in macrocrystal magnesian ilmenite may exhibit enhanced Nb contents, e.g., Premier (Kable et al. 1975, Mitchell 1986). Thus, increasing Nb is correlated with increasing Ti and reflects modal increases in groundmass titaniumrich oxides.
Figure 3.7 demonstrates that evolved orangeites cannot always be distinguished from unevolved types on the basis of Zr and Nb abundances, e.g., orangeites from Finsch and
300
t --. 200 E a. a.
.tl Z
100
o o
SOV:*,~
100 200 300 400 500 600
Zr (ppm) •
Figure 3.7. Nb versus Zr for orangeites. Data sources as in Figure 3.5.
GEOCHEMISTRY OF ORANGEITES
t 300
E 200
• KIMBERLITE
~ ... .J:l
Z 100
o o 100 200 300
271
400 500 600 700 800 900 1000
Zr (ppm) ~
Figure 3.8. Nb versus Zr for orangeites, kimberlites. and olivine lamproites. Data sources for orangeites as in Figure 3.5. Data for Ellendale and Prairie Creek lamproites from Jaques et al. (1986) and Fraser (1987). Kimberlite data from Clement (1982). Smith et al. (1985b). Spriggs (1988). and Kampata (1993).
Sover have low Nb contents. However, evolved varieties are characterized by low Nb
abundances «100 pmm) and widely varying Zr contents with relatively high (>3) ZrlNb ratios. Unevolved orangeites display a positive correlation with respect to increasing Nb and Zr contents. Individual localities have distinct, but overlapping, Nb and Zr contents and similar ZrINb ratios (Table 3.9).
Figure 3.8 indicates that kimberlites cannot be distinguished from orangeites on the basis of their Nb and Zr contents. Olivine lamproites may be easily distinguished from both kimberlites and orangeites on the basis oftheir much higher Zr contents. Sover North has Zr and Nb contents similar to the most Zr-poor olivine lamproites. Figure 3.8 indicates
that petrographically similar olivine-rich rocks, with ZrINb > 3 and Zr > 500 ppm, are more likely to be olivine lamproites than unevolved orangeites.
3.5.3. Thorium and Uranium Thorium and uranium are probably primarily concentrated in groundmass apatite
and perovskite. Table 3.7 and Figure 3.9 show that Th and U abundances vary widely within and between intrusions. Extremely high Th (630, 920 ppm) and U (15.0, 22.9ppm, Th/U >40) abundances have been reported by Fesq et al. (1975) for apatite-rich orangeites from the Bellsbank Main Fissure. On the basis of the limited data available, with the exception of Finsch, evolved orangeites appear to be poor in Th and U relative to unevolved types. Th/U ratios range from 3 to 9 (Figure 3.8). There is a notable lack of correlation between Th and U on logarithmic plots of abundances for Bellsbank, Sover,
and Sover North orangeites. Coherence is better for Swartruggens and Finsch. Whether the data spread is real and due to removal of U as the highly soluble uranyl ion during weathering (Paul et al. 1977, Fesq et al. 1975) or results from the compatible behavior
of U is unresolved. Th and U abundances and Th/U ratios of kimberlites are insufficiently well estab
lished to state conclusively that the Th/U ratios of orangeites (6-11) are greater than those
272
t 100
50 -E c.. c..
.£:.
I-- 10
5
• SWARTRUGGENS • SOVER + BELLSBANK o FINSCH
(ff} SOVER NORTH
5 10
U (ppm) ---.
Figure 3.9. Th versus U for orangeites. Data sources as in Figure 3.5.
CHAPTER 3
of kimberlites (3-7) as suggested by Gurney and Hobbs (1973). Paul et al. (1977) have shown that kimberlite Th/U ratios may exceed 10.
Olivine lamproites may have very high Th/u ratios (24) with no correlation evident between Th and U abundances (Jaques et al. 1986). It is not known whether the Th/U ratios of kimberlites and olivine lamproites reflect primary variations in Th and U abundances or are elevated because of U leaching.
3.5.4. Rare Earth Elements
Rare earth elements are concentrated in apatite, perovskite, and REE-bearing carbonates during the later stages of crystallization of orangeites (this work, Fesq et al. 1975, Mitchell and Reed 1988, Mitchell and Steele 1992). Orangeites typically have high total REE abundances (Table 3.10) and are characterized by extreme fractionation of the light REE relative to the heavy REE (Mitchell and Brunfelt 1975, Fesq et al. 1975).
Tables 3.9 and 3.10 indicate that LalYb ratios and REE abundances vary widely within and between intrusions. Total REE abundances typically exceed 500 ppm. Fesq et al. (1975) have reported atypical, extraordinarily high REE abundances (870 and 1120 ppm La, 1910 and 2080 ppm Ce) in apatite-rich orangeites from the BeIIsbank Main Fissure.
Representative chondrite normalized REE distribution patterns are given in Figures 3.10 to 3.12. Within a given intrusion, light REE are not fractionated relative to each other
GEOCHEMISTRY OF ORANGEITES 273
Table 3.10. Averages and Range of Rare Earth Element Abundances (ppm) in Orangeites
Swartruggens Finsch Bellsbank Sover
La 218: 145-316 62:41-100 252: 126-504 168:80-250 Ce 429:276-719 132: 82-235 464:236-871 324: 147-544 Pr 55:28-91 36: 17-55 Nd 158: 106-218 55:32-94 163:83-286 115: 53-178 Sm 20.0: 14.1-32.3 7.7: 4.8-12.8 17.0: 8.5-28.3 12.6: 5.9-18.0 Eu 5.03: 3.58-10.5 1.80: 1.14-3.10 3.93: 1.97-6.33 3.00: 1.50-4.62 Od 13.2: 6.7-19.3 9.51: 4.81-15.2 Tb 1.65: 0.86-3.31 0.52: 0.32-0.98 1.05: 0.56-1.59 0.82: 0.43-1.28 Dy 4.09: 1.57-6.33 3.09: 1.39-5.13 Ho 0.62: 0.25-1.00 0.52: 0.23-0.89 Er 1.51: 0.47-2.98 1.20: 0.48-2.33 Tm 0.17: 0.05-0.38 0.19: 0.43-1.90 Yb 2.07: 0.95-5.77 0.72: 0.37-1.11 1.13: 0.25-2.57 0.99: 0.43-1.90 Lu 0.14: 0.02-0.36 0.13: 0.05;0.34 (n) 15-21 22 21-35 28-30
New1ands New Elands Star Sover North
La 203: 140-265 247: 175-334 192: 164-220 149: 142-168 Ce 346:253-480 439:319-647 332:238-271 288:263-316 Pr 38:27-51 35:32-38 Nd 120: 84-156 174: 138-215 113: 107-132 Sm 11.3: 8.0-14.3 21.4: 17.3-24.8 15.0: 12.7-17.9 15.1: 14.1-16.2 Eu 2.69: 1.96-4.15 5.95: 4.33-6.78 3.16: 2.19-4.28 3.58: 3.04-4.14 Od 8.43: 6.08-14.1 10.5: 9.73-12.0 Tb 0.63: 0.45-0.87 1.36: 1.19-1.62 0.69: 0.06-1.08 1.07: 0.94-1.1 7 Dy 2.73: 2.08-3.63 4.08: 3.04-5.08 Ho 0.42: 0.40-0.51 0.66: 0.61-0.81 Er 0.97: 0.79-1.25 1.51: 1.34-1.94 Tm 0.13: 0.10-0.20 0.18: 0.15-0.23 Yb 0.76: 0.49-0.97 1. 72: 1.39-2.02 0.72: 0.26-1.19 1.17: 1.00-1.55 Lu 0.11: 0.08-0.13 0.14: 0.12-0.16 (n) 14-18 4-8 8 9
(n) = number of samples. Data sources: Swartruggens (Mitchell and Brunfelt 1975, this work); Finsch (Fraser 1987); Bellsbank (Tainton 1992, this work); Sover (Tainton 1992); New1ands (Tainton 1992, this work); New E1ands (Smith et al. 1985b, this work); Star (this work); Sover North (Tainton 1992).
and the patterns are subparallel. Divergence of the patterns, represented by upward inflections, are observed only with respect to the heavy REE. The divergences are attributable to either experimental errors (Tainton 1992) or to contamination of the rock with crustal material (Mitchell 1986).
Figures 3.10 and 3.11 show that unevolved orangeites have very similar REE distribution patterns and high LalYb ratios (typically 100-350). In contrast, evolved orangeites (Figure 3.12) are not as enriched in their absolute REE abundances and have lower LalYb ratios (80-150). Interestingly, the Finsch orangeites have very low LalYb ratios (20-115) and REE distribution patterns (Figure 3.12), similar to those of evolved orangeites.
No distribution pattern exhibits significant Eu or Ce anomalies. Fesq et al. (1975) have reported large negative Eu anomalies in samples from Bellsbank. Recent studies of
274
IJ.J ..... a: 102 o Z o ::z:: () ....... IJ.J ..... jjj (!) z <[ 10 a:: o
La Pr Ce Nd
• SOVER
o BELLSBANK
+ NEWLANDS
Eu Tb Ho Tm Lu Sm Gd Oy Er Yb
CHAPTER 3
Figure 3.10. Chondrite nonnalized REE distribution patterns for unevolved orangeites from Sover, Bellsbank, and Newlands. All data from Tainton (1992).
Bellsbank orangeites have not supported the existence of this anomaly (Tainton 1992, this work). Tainton (1992) considers that Fesq et al. (1975) analyzed strongly weathered material and the Eu anomaly resulted from preferential leaching of Eu2+ from the rocks by reducing ground waters. Mitchell and Brunfelt (1975) reported weak negative Eu anomalies in the Swartruggens orangeites. This anomaly may also may be related to weathering, as less-altered samples from Swartruggens analyzed in this work were found not to exhibit any Eu anomaly (Figure 3.11). It is concluded in this work and by Tainton (1992) that the REE distribution patterns of orangeites do not characteristically exhibit negative or positive Eu anomalies.
Surprisingly, the REE contents and distribution patterns of kimberlites are not well established. Although the basic features of the REE geochemistry are known, Mitchell (1986) has commented that many of the published data are subject to serious analytical errors, or analyses have been undertaken upon contaminated diatreme facies material and/or altered rocks. There have been no systematic studies of REE distributions in well-characterized suites of fresh uncontaminated kimberlite from a single intrusion.
Data for fresh kimberlites (Mitchell and Brunfelt 1975, Mitchell and Clement unpublished, Kampata 1993, Spriggs 1988) indicate that kimberlites are light REE
GEOCHEMISTRY OF ORANGEITES 275
103 0 .....
~ " • SWARTRUGGENS +~,,>, 0 NEW ELANDS
~.\ 0, + STAR \' \\ • ROBERTS VICTOR
IJJ ~ l-ll:: 102 +~\ 0 z 0 \ ~, ::c u
\(' ..... IJJ "
'-': " I- -.;::', IJJ -...;::', <.!)
+~ Z ~ 10 Il::
"0 0
+-+
La Pr Eu Tb Ho Tm Lu Ce Nd Sm Gd Oy Er Yb
Figure 3.11. Chondrite nonnalized REE distribution patterns for unevolved orangeites from Swartruggens, New Elands, Star, and Roberts Victor. All data this work.
enriched, REE distribution patterns are linear, Eu anomalies are absent, and REE abundances and LalYb ratios vary widely between intrusions.
Figure 3.13 indicates that kimberlites cannot be distinguished from orangeites on the basis of their Sm content and LalYb ratios. The claim by Mitchell and Brunfelt (1975) and Mitchell (1986) that micaceous kimberlites (i.e., orangeites) have higher LalYb ratios than kimberlites is not supported by the data in Figure 3.13.
Mitchell and Bergman (1991) have previously concluded that olivine iamproites cannot be distinguished from orangeites on the basis of their REE geochemistry, as LalYb ratios and REE distribution patterns are similar. This conclusion is supported by Figure 3.13.
Using a very limited data base, Mitchell (1986) suggested that kimberlites and orangeites could be distinguished on the basis of their SmiTh ratios. This conclusion is not supported by the more extensive data now available.
Smith et al. (1985b) have noted that a plot of P205 versus Ce content may serve to distinguish kimberlites from orangeites, although Figure 3.14, incorporating recent data, does not support this observation. Although most on-craton South African kimberlites have high P205 and low Ce contents relative to orangeites, there are no significant
w ~
a:: 102 0 Z 0 ::I: U ...... W ~ LiJ (!) z « 10 a: 0
La Pr Ce Nd
• SOVER NORTH
o PNIEL
+ POSTMASBURG
• FINSCH PK37
Eu Tb Sm Gd Dy
Ho Tm Lu Er Yb
Figure 3.12. Chondrite normalized REE distribution patterns for a Finsch orangeite (Fraser 1987) and evolved orangeites from Sover North. Pniel. and Postmasburg (Tainton 1992) .
• KIMBERLITE
t" • •
E Ii 20
E en POST; 10 M4SBURG
~ ~-.. ~-FINSCH
0
0 50 100 150 200 250 300
La / Yb ~
Figure 3.13. Sm versus La/Yb ratio for orangeites. kimberlites. and olivine lamproites. Data sources for orangeites as in Figure 3.5. Kimberlite data from Mitchell and Clement (unpublished). Mitchell and Brunfelt (1975). Spriggs (1988). and Kampata (1993). Olivine lamproite data for Ellendale. Argyle. and Prairie Creek from Jaques et al. (1986. 1989) and Fraser (l987).
276
GEOCHEMISTRY OF ORANGElTES
3'0
t -~
It)
o rr 1'0
• ON - CRATON KIMBERLITES o OFF -CRATON KIMBERLITES + UNEVOLVED ORANGEITES • SOVER NORTH • 0 ~ FINSCH
100
• •
200 300
o
400
Ce (ppm) •
277
+
+ +
• 500 600
Figure 3.14. P205 versus Ce for orangeites and kimberlites. Data for orangeites from Tainton (1992), Fraser (1987), Smith et al. (1985b) and this work. Data for kimberlites from Smith et al. (l985b), Spriggs (1988), Kampala (1993), and Mitchell and Clement (unpub.). SN = Sover North; KK = Kundelungu kimberlites.
differences for samples with low P and Ce contents. On-craton kimberlites from Kunde)ungu, and some Ce- and mica-rich South African kimberlites plot well within the "orangeite field," i.e., P205/Ce ratios <0.05, as do off-craton kimberlites from South Africa and Namibia.
3.5.5. Alkali Elements Rubidium is primarily hosted by phlogopite in all orangeites, and by amphibole and
potassium feldspar in evolved varieties. Wide variations exist in Rb contents within and between orangeites, e.g., Swartruggens (mean 191, range 146-273 ppm; Mitchell and Brunfelt 1975, Smith et al. 1985b, this work), Finsch (mean 137, range 43-187 ppm; Fraser 1987, this work), Bellsbank (mean 101, range 56-185 ppm; Tainton 1992, this work), Sover (mean 159, range 64-305 ppm; Tainton 1992), Newlands (mean 60, range 27-89 ppm; Tainton 1992), New Elands (mean 162, range 153-176 ppm; Smith et al. 1985a, this work), Star (mean 132, range 110-159 ppm; this work), Sover North (mean 60, range 134-202 ppm; Tainton 1992), Postmasburg (mean 234, range 212-256 ppm; Tainton 1992). Much of the variation reflects mixing with Rb-free olivine macrocrysts and variations in phlogopite to carbonate plus apatite ratios. The data indicate no significant differences in the Rb contents of unevolved and evolved orangeites. There is a strong positive correlation (Figure 3.15) between K and Rb, reflecting increases in modal phlogopite contents.
278
-E Q. Q. -
10000
• SaVER o BELLSBANK • SWARTRUGGENS + FINSCH
50 100 500
Rb (ppm) •
CHAPI'ER3
Figure 3.1S. K versus Rb for orangeites and olivine lamproites. Data sources for orangeites as in Figure 3.5. Ellendale and Prairie Creek data from Jaques et al. (1986) and Fraser (1987). respectively.
KlRb ratios (Table 3.9) also vary widely within and between intrusions. Some of the higher ratios. e.g., those greater than 300 at Finsch, may reflect preferential removal of Rb during weathering and/or deuteric alteration (Mitchell and Crocket 1971, Barrett and Berg 1975). There are no significant differences in the KlRb ratios of unevolved and evolved orangeites. KlRb ratios in the majority of orangeites lie between 100 and 250 and are thus typically lower than the crustal average of 250 (Figure 3.15).
Orangeites, because of their higher modal abundances of phlogopite, are enriched in Rb relative to kimberlites: e.g., Bultfontein, 72-111 ppm; Wesselton, 88-107 ppm; De Beers, 26-91 ppm; Jagersfontein, 51-107 ppm (Gurney and Berg 1969). Orangeites cannot be distinguished from kimberlites on the basis of their KlRb ratios (Gurney and Berg 1969, this work). KlRb ratios of kimberlites in the Kimberley group range from 90 to 238 (Gurney and Berg 1969).
Olivine lamproites from Ellendale (Jaques et al. 1986) and Prairie Creek (Fraser 1987) have higher Rb contents than orangeites of similar K content (Figure 3.15); consequently, KlRb ratios are less than those of orangeites. Note, that unlike orangeites, K and Rb in the Ellendale lamproites are not positively correlated (Figure 3.15).
Cesium is hosted primarily by phlogopite in all orangeites, but interpretation of Cs data is rendered difficult by the extreme mobility of this element during weathering and/or deuteric alteration (Fesq et al. 1975).
The few data available indicate that, in the Barkly West region, Cs contents of unevolved orangeites (Tainton 1992) are similar within and between intrusions, e.g., Sover (mean 9.51, range 3.29-19.96 ppm), Bellsbank (mean 9.53, range 4.79-18.33 ppm), and Newlands (mean 9.51, 3.29-19.96 ppm). Cs abundances show no correlation with K contents, suggesting that some of the variation may be due to alteration. Significantly lower Cs contents (3.0-4.4 ppm) for Bellsbank, which may also result from alteration, have been reported by Fesq et al. (1975). The Finsch orangeites have signifi-
GEOCHEMISTRY OF ORANGElTES 279
cantly lower Cs contents (mean 2.39, range 0.9-3.27 ppm; Fraser and Hawkesworth 1992) than the Barkly West orangeites. Roberts Victor orangeites have a mean Cs content of 6 ppm (range 5-8; Gurney and Berg 1969). Evolved orangeites from Sover North and Postmasburg average 4.09 ppm (1.21-8.91) and 4.07 ppm (3.01-4.64 ppm Cs), respectively. The wide range in Cs abundances found at Sover North shows no correlation with K content.
KlCs ratios of un evolved orangeites (Tainton 1992), e.g., Sover (mean 3357, range 1001-104140), Bellsbank (mean 1612, range 403-2734), Newlands (mean 3986, range 1892-64260) are lower than those of evolved orangeites, e.g., Sover North (mean 12867, range 5060-240130, Postmasburg (mean 11678, range 9998-14838). Finsch orangeites are anomalous with respect to other unevolved orangeites in having high KlCs ratios (mean 11000, range 6159-15400, Fraser and Hawkesworth 1992).
Insufficient data are available to state conclusively whether or not the Cs contents of kimberlites differ significantly from orangeites. Older spectrographic data, which are subject to serious analytical errors resulting in overestimation ofCs abundances, suggest 7-76 ppm Cs (KlCs = 118-4500) for Kimberley area kimberlites (Gurney and Berg 1969). Recent NAA data for the same kimberlites (Mitchell and Clement unpublished) give lower abundances (1.68-44.9 ppm Cs) and low K/Cs ratios (495-4440).
Few data exist for Cs in olivine lamproites. Cs abundances (0.7-44 ppm) and hence KlCs ratios (1900-25000) vary enormously for Ellendale olivine lamproites (Jaques et al.1986)
In summary, orangeites, olivine lamproites and kimberlites cannot be distinguished from each other on the basis of their Cs contents or KlCs ratios. Orangeites have greater Rb contents than kimberlites, although KlRb ratios are similar. Olivine lamproites are relatively rich in Rb and appear to have lower KlRb ratios than either kimberlites or orangeites.
3.5.6. Lead
Lead in orangeites is probably hosted primarily by phlogopite. Galena is present only as an insignificant accessory phase. The Pb contents of unevolved and evolved orangeites do not differ significantly, e.g., Sover (mean 17, range 6-31 ppm), Bellsbank (mean 27.8, range 14-50 ppm), Newlands (mean 18.8, range 4-32 ppm), New Elands (mean 35.5, range 30-41 ppm), Swartruggens (mean 35.7, range 22-40), Sover North (mean 13.8, range 14-23 ppm; Tainton 1992, Smith et al. 1985h). Finsch is anomalous in having relatively low Ph (mean 12.62, range 4.55-24.96) contents (Fraser and Hawkesworth 1992).
The limited data available indicate that kimherlites have widely varying Pb contents, e.g., Kimberley group, 2.3-21.9 ppm (Smith et al. 1985b); 10-37 ppm (Clement 1982); Namibia, 5.1-8.5 ppm (Spriggs 1988). Despite the wide range, Pb contents are typically less than 17.5 ppm (Figure 3.16). High Pb contents are found primarily in varieties modally enriched in mica.
Smith (1983) has noted that the Pb contents of orangeites are higher than those of kimberlites. A plot of Pb versus Ce (Figure 3.16) demonstrates that orangeites typically contain more than 15 ppm Pb and 200 ppm Ce, suggesting that this diagram, in most
280 CHAPTER 3
t 50 - +. KIMBERLITES ORANGEITES
40 - • SOVER NORTH .. FINSCH
+ + +
+ E + Q. 30 - + + +t. -i:t-+ + + Q. .. + SN + ~ + + ~
+
+ +
if 20 - .. + _~~ tr + + + tr / • + ~- -.-; -~."'-.-:~1!~-*-.t-:-:++++- +; - - - - -
10- .... ~ "'"+ 1e..... . · . ..., ... .; lr ... .;. •
1 1 1 III 1 1 1 1 1 1 100 200 300 400 500 600
Ce (ppm) ..
Figure 3.16. Pb versus Ce for orangeites and kimberlites. Data sources as in Figure 3.14. SN = Sover North.
instances, may be used to discriminate between the two rock types. Orangeites from Finsch are anomalous with respect to their Pb and Ce contents relative to other orangeites (Figure 3.16).
It is noted above that kimberlites are poor in Rb relative to orangeites; thus, plots of Pb versus Rb should also serve to discriminate between the two rock types. However, such plots will merely reflect the high modal abundances of phlogopite in orangeites relative to kimberlites and are thus unnecessary.
Olivine lamproites contain widely varying amounts of Pb (mean 51, range 17-124 ppm; Jaques et al. 1986) and do not appear to be significantly different from orangeites in their Pb content.
3.6. INTER-ELEMENT RELATIONSHIPS
3.6.1. Extended Incompatible Element Distribution Diagrams
Extended incompatible element distribution diagrams or "spidergrams" have become a sine qua non of geochemical studies of basaltic and related rocks. While such diagrams are entirely appropriate for interelement comparisons in rocks which approximate liquid compositions, their applicability to rocks whose compositions are controlled by crystal accumulation is open to question.
The relative order of elements in these diagrams reflects decreasing liquid--crystal distribution coefficients for a "normal" mantle source. This sequence gives smooth interelement distribution curves for oceanic basalts when trace element abundances are normalized to a primitive mantle composition. Clearly, magmas which are not derived from mineralogically similar sources will not give smooth distribution patterns on such plots. Thus, distribution patterns for rocks such as potassic lavas and melilitites characteristically exhibit negative anomalies for several trace elements when they are normal-
GEOCHEMISTRY OF ORANGEITES 281
ized to primitive mantle compositions. Interpretation of the geological meaning of such anomalies is subjective. Negative anomalies are commonly considered to reflect the presence of residual phases in the mantle which sequester the elements in question. However, the presence of these residual phases implies that the source mantle is unlikely to be primitive, therefore, the use of primitive mantle as a normalizing index must be inappropriate. Alternatively, the anomalies may be simply an indication that a given magma was derived from a mantle source very different from that of oceanic basalts. In the case of lamproites such a conclusion is not surprising, given that radiogenic isotopic studies indicate that lamproite parental magmas might be derived from an evolved mantle that has undergone several episodes of metasomatism.
Several studies of orangeites and related rocks (Tainton and McKenzie 1994, Fraser and Hawkesworth 1992, Tainton 1992, Sheppard and Taylor 1992, Jaques et al. 1989, Rock 1990) have used such extended incompatible element distribution diagrams in attempts to deduce the geochemical characteristics of the sources of parental magmas. In all of these studies it is claimed that incompatible element abundances, normalized to the composition of the primitive mantle, reflect the composition of the parent magma and not the effects of extended fractional crystallization and/or hybridization. Negative anomalies are interpreted as indicating the presence of residual phases in the mantle sources of the magmas. These sources are postulated to range in character from metasomatized evolved mantle to subducted oceanic crust; however, in all cases they are certainly not equivalent to the "normal" primitive peridotite mantle proposed as the source of oceanic basalts.
Figures 3.17-3.24 present incompatible trace element abundances for orangeites and related rocks normalized to the composition of primitive mantle (Sun and McDonough 1989). Bearing in mind the above comments, it is considered that inferences about the nature of the mantle sources from these diagrams mayor may not be realistic. Their principal value is for comparative purposes, as differences in the distribution patterns may reflect differing processes or sources involved in the generation of these magmas.
Figure 3.17 shows that unevolved orangeites from the Barkly West District give distribution patterns with prominent negative Rb, K, and Sr anomalies. Wide fluctuations in Cs abundances may reflect alteration as noted above. The figure clearly shows the enrichment of Bells bank orangeites in many incompatible elements relative to those from Sover. Orangeites from Swartruggens, New Elands, and Star exhibit similar distribution patterns (Figure 3.18) to those of the Barkly West orangeites, but differ in that the magnitude of the Rb and K anomalies is much less. The absence of Hf anomalies in these orangeites suggests that the minor Hf anomalies present in the Barkly West distribution patterns are an artifact of the analytical technique used to determine Hf. This conclusion is supported by the lack of a corresponding Zr anomaly in the Barkly West samples.
Figure 3.19 gives distribution patterns for evolved orangeites from Sover North and Postmasburg. The Sover North pattern is very irregular and differs from unevolved orangeites in having no significant negative Rb and K anomalies. Negative Sr and P and positive Zr and Hf anomalies are present. Postmasburg orangeites exhibit positive K and negative Sr, Th, and U anomalies. Figure 3.19 also shows the distribution pattern of the Finsch orangeites is similar to those of evolved orangeites in that they both lack Rb and K anomalies and have negative Sr and Th anomalies.
282
w ...J
10 3
I-Z <t :!:
w > ~ 10 2 :!: a:: a.. ..... w I-W
10 (!) z <t a:: 0
0 ORANGEITES
0--0 BELLSBANK - SOVER *--- ... NEWLANDS
Pb Rb Th K Nb Ce Nd Sm Zr Ti Yb Cs 80 U To La Sr P Hf Eu Y Lu
CHAPTER 3
Figure 3.17. Incompatible element distribution diagrams for unevolved orangeites from Bellsbank, Sover, and Newlands. All data from Tainton (1992).
103 w ...J I-z <t :!: w > 10 2 i= ~ a:: a.. ..... w I-w 10 (!) Z <t a:: 0
ORANGEITES
- SWARTRUGGENS 0--0 NEW ELANDS *--- .. STAR
Pb Rb Th K Nb Ce Nd Sm Zr Ti Yb Cs 80 U To La Sr P Hf Eu Y Lu
Figure 3.18. Incompatible element distribution diagrams for orangeites from Swartruggens, New Elands, and Star. All data this work.
GEOCHEMISTRY OF ORANGElTES l83
103 LLI ...J I-Z <I ~
LLI > 10 2 ~ ~ it: a.. ...... LLI I-i:ij 10 (!) Z <I 0:: 0
EVOLVED ORANGEITES
0--0 SOVER NORTH - FINSCH
*---... POSTMASBURG
Pb Rb Th K Nb Ce Nd Sm Zr Ti Yb Cs So U To La Sr P Hf Eu Y Lu
Figure 3.19. Incompatible element distribution diagrams for a Finsch orangeite (Fraser 1987) and evolved orangeites from Sover North and Postmasburg (Tainton 1992).
Figure 3.20 gives distribution patterns for "on-craton" kimberlites from the Kimberley area. These exhibit negative Rb and K and positive P and Zr anomalies but show no significant depletion in Sr. In contrast, "off-craton" kimberlites from Namibia exhibit significant negative Sr anomalies (Figure 3.21) in addition to Rb and K depletion. Figure 3.21 also shows that olivine melilitites from Namaqualand (South Africa) have virtually identical distribution patterns to those of Namibian kimberlites. This suggests that the processes controlling the distribution of trace elements during the partial melting episodes, which gave rise to the Namibian kimberlites and South African melilitites, were similar (see 4.8).
Figure 3.22 shows that distribution patterns for olivine lamproites from Ellendale have well-defined negative K, Rb, Sr, and P, and positive Zr and Hf anomalies. Similar patterns are evident for the Prairie Creek olivine lamproites. Note that the distribution patterns are very different from those of silica-rich lamproites (Figure 3.23), which have negative Th, D, Nb, Ta, and Sr anomalies, but lack K anomalies. The diagrams suggest that different processes have been involved in the formation of olivine and phlogopite lamproites, which reinforces the proposition of Mitchell and Bergman (1991) that olivine lamproites cannot be parental to phlogopite lamproites. Figure 3.24 demonstrates that the distribution patterns of other ultrapotassic volcanic rocks are similar to those of phlogopite lamproites but contain additional negative Ti anomalies.
In summary, incompatible element distribution patterns for orangeites suggest they formed neither by the same processes nor from the same sources as generate phlogopite
284 CHAPTER 3
LLJ 103
...J ~ Z <t ~
LLJ > 10 2 E ~ a: 11.. , LLJ t: 10 ...J 0:: LLJ CD ~ ~
KIMBER LITES
0--0 DE BEERS
*----* DUTOITSPAN - WESSELTON
Pb Rb Th K Nb Ce Nd sm Zr Ti Yb Cs Ba U To La sr P Hf Eu Y Lu
Figure 3.20. Incompatible element distribution diagrams for kimberlites from the Kimberley area. All data from Clement (1982) and Mitchell and Clement (unpublished).
- K35 BERSEBA RESERVE 2
0--0 GAMOEP MELILITITE
Pb Rb Th K Nb Ce Nd sm Zr Ti Yb Cs Ba U To La sr P Hf Eu Y Lu
Figure 3.21. Incompatible element distribution diagrams for an off-craton Namibian kimberlite (Spriggs 1988) and a South African melilitite (Rogers et al.1992).
GEOCHEMISTRY OF ORANGElTES
w > E ~
10 2 ~ Q.
...... w t-O a:: Q.
~ 10 <t ..J
W z :> ::J 0
, ~ I I '*, I I I I II
*
- ELLENDALE 4
0--0 ELLENDALE 9
*---",* PRAIRIE CREEK
Pb Rb Th K Nb Ce Nd Sm Zr Ti Yb Cs Ba U Ta La Sr P Hf Eu Y Lu
285
Figure 3.22. Incompatible element distribution diagrams for Ellendale (Jaques et aJ. 1986) and Prairie Creek (Fraser 1987) olivine lamproites.
W ..J I-Z <t ~
w ~ !:: ~ iE Q. ....... w I-0 a:: Q.
~ <t ..J
103
10 2
10
Cs Rb Th K Nb Ce P F Zr Ti Yb Pb So U To La Sr Nd Hf Sm Y Lu
Figure 3.23. Composite incompatible element distribution diagram for phlogopite lamproites (after Mitchell and Bergman 1991, Figure 7.29).
286
kI -oJ ~ Z < :t kI > ~ i ii: a.. .... en ~ < -oJ
~ a.. 0:
103
102
10
Ba K Nb Ce Nd Hf Sm Tb Rb Th To La Sr P Zr Ti
CHAPTER 3
Figure 3.24. Composite incompatible element distribution diagram for high-potassium Roman Province type (RPT) lavas from Monti Ernici and Vulsini. Italy (after PeccerilJo et al. 1988).
lamproites and other continental ultrapotassic volcanic rocks. The common presence of negative Ta-Nb-Ti anomalies in the latter indicate that their mantle sources have very different mineralogical characteristics from those of orangeites.
Orangeites, olivine lamproites, and kimberlites are all similar in having negative Rb and K anomalies, suggesting that similar minerals are retained in their mantle sources. The negative Sr anomaly found in many of these diverse rocks cannot have a common origin because of marked differences in the isotopic composition ofSr in each petrological clan (3.8.1).
The conventional explanation of the Rb and K anomalies is that residual phlogopite must remain in the mantle sources. The absence of K anomalies in evolved orangeites might imply they are formed by a greater degree of partial melting, resulting in the elimination of phlogopite as a residual phase. The Sr anomalies may be due either to the presence of a residual Sr-bearing phase, such as a phosphate, or to intrinsic depletion of the mantle in clinopyroxene, and hence Sr, by previous episodes of basaltic magma formation. Potassium richterite is another possible residual phase which might retain Sr, Rb, and K. Although Sr anomalies are present, there are no corresponding REE anomalies, suggesting that Sr and REE are hosted by different phases in the mantle sources of these magmas. The absence of Ta-Nb-Ti anomalies in orangeites, kimberlites, and olivine lamproites suggests that their mantle sources do not contain residual titanates.
3.6.2. Ce/Y and LafYb versus Zr/Nh
Plots of Ce/Y versus ZrlNb are commonly used to infer the degree of melting involved in the production of basaltic rocks from peridotitic sources. Figure 3.25 illustrates Ce/Y versus ZrlNb ratios for orangeites relative to the compositions of magmas
GEOCHEMISTRY OF ORANGEITES
>"'-
40
30
Q) 20 U
10
2 4
• SOVER NORTH
PNIEL / POSTMASBURG
6 8 10 12 14 16
Zr I Nb
287
Figure 3.25. Ce/Y ratio versus ZrlNb ratio for orangeites. Data sources as in Figure 3.5. B-N = BellsbankNewlands; ST = Star; SWR = Swartruggens. Solid curved line from (Tainton 1992) indicates the compositions of melts formed by various degrees (%) of equilibrium partial melting of a peridotite containing 1.4 ppm Ce, 3.45 ppm Zr, 8.51 ppm Zr, and 0.54 ppm Nb, i.e., a bulk earth composition.
formed by the partial melting of a peridotitic source. The plot shows quite clearly that
orangeites cannot be derived by single-stage partial melting of such a source (Tainton
1992). Note that evolved orangeites plot with higher ZrlNb and lower Ce/Y ratios than
unevolved orangeites, suggesting that the former are produced by greater degrees of
partial melting of a source mantle enriched in incompatible elements relative to "normal"
peridotite. Thus, Tainton (1992) interprets the trend of decreasing Ce/Y and increasing
ZrlNb from Bellsbank to Pniel as a partial melting trend. On this basis, Swartruggens
orangeites appear to represent the lowest degrees of partial melting of the orangeite source
mantle. Figure 3.26 shows that orangeites from different intrusions define a broad hyperbolic
trend of increasing LalYb ratio with decreasing Zr/Nb ratio, which, within a given intrusion, do not show any corresponding correlation. The conventional interpretation of the data is that the increasing LalYb ratios represent decreasing amounts of partial
melting. Thus, evolved orangeites could be formed by greater degrees of partial melting than unevolved varieties. Finsch is anomalous in having the geochemical characteristics of an evolved orangeite. Note that the degree of partial melting suggested by this plot is not the same as that deduced from Figure 3.25, e.g., Swartruggens on the basis of LalYb
ratios appears to represent a greater degree of melting than the Star and Barkly West
orangeites. An alternative explanation of this contradiction is that the Swartruggens and Star orangeites are derived from sources with compositions different from those of the Barkly West and Finsch orangeites. The LalYb and ZrlNb ratios are not compatible with derivation of the parental magmas from a simple peridotite mantle source unless the
degree of partial melting is vanishingly small («<1 %).
288 CHAPTER 3
9
8
r: PNIEL - POSTMASBURG
.Q
Z 5 ..... ... N
4
3
2
50 100 150 200 250 300
La / Yb •
Figure 3.26. ZrlNb ratio versus LalYb ratio for orangeites. Data sources as in Figure 3.5.
3.7. PERIDOTITE MIXING AND ASSIMILATION
The magnesium content and compatible trace element geochemistry of orangeites (3.3, 3.4) is dominated by the presence of ubiquitous macrocrystal olivines, considered to represent disaggregated mantle harzburgite or peridotite. Other phases derived from these contaminants appear to have been completely dissolved (Cr-diopside, enstatite) or mainly fractionated out (garnet, Cr-spinel). It is possible that the assimilated xenoliths also contained phlogopite, although the effects of phlogopite assimilation are extremely difficult to decipher from those of crystal fractionation in these mica-rich rocks (Fraser and Hawkesworth 1992); however, they are probably insignificant as this mineral is not abundant in common mantle peridotites «1 vol%). This conclusion will not apply if phlogopite is a major component of the mantle sources of orangeites and residual phlogopite (restite) is incorporated into partial melts.
Assuming orangeites are derived by small degrees of melting from a peridotite source, Fraser and Hawksworth (1992) calculated thatthe Ni and Cr contents of the Finsch orangeites are consistent with 5~60% entrainment of a peridotite containing 2000 ppm Ni and 2800 ppm Cr. A major problem with this conclusion is that it is based on the assumption of primary orangeite magmas having the same Ni and Cr contents (20~500
GEOCHEMISTRY OF ORANGEITES
2500 <> BELLSBANK - NEWLANDS • SOVER
-- 2000 E Q. Q.
z 1500
1000
99% AJE25
w • SOVER POSTMASBURG
NORTH
500 -+---....,...-----r---r----r--.I 0'075 0'100 0'125
Sm/Nd
0'150 0·175
289
Figure 3.27. Ni versus SmlNd ratio for orangeites. Solid curve is a hypothetical mixing curve calculated for the addition of a peridotite (AJE25) analyzed by Erlank et al. (1987) to a Bellsbank orangeite (after Tainton 1992).
ppm) as many other limited partial melts of normal mantle. The assumption is unlikely to be correct (see 4.2.3).
Using Ni versus SmlNd ratios (Figure 3.27), Tainton (1992) has shown that incompatible element ratios are insensitive to mixing of garnet lherzolite with orangeite provided that the amount of peridotite is less than 80 wt%. Orangeites from the Barkly West region are considered to have undergone 10-80% contamination, an amount in accord with estimates deduced in this work from the major element composition of these rocks (3.3.1).
The effects of peridotite mixing on incompatible element abundances and ratios are easier to estimate as normal peridotites have very low contents of these elements (<5 ppm Rb, <30 ppm Sr, <10 ppm Zr) even relative to the amounts found in contaminated orangeites. Fraser and Hawkesworth (1992) have concluded that peridotite entrainment principally dilutes the incompatible element contents of the initial melt. Hence, such melts must have had extremely high incompatible element contents and might have contained at least 100 ppm Nd, 1500 ppm Sr, and 400 ppm Zr, with a SmlNd ratio of 0.09 (Fraser and Hawkesworth 1992). However, these estimates are based upon the assumption that the initial melts, derived from normal mantle, contained 300-500 ppmNi and were mixed with peridotite containing about 2000 ppm Ni (Figure 3.28). Regardless of the veracity of the assumptions underlying these calculations, the conclusion has the merit ofshowing that even extremely small «0.5%) degrees of partial melting are incapable of producing orangeites from primitive mantle compositions (Fraser and Hawkesworth 1992).
Figure 3.29 shows that the LalNd versus SmlNd array formed by Finsch orangeites can be modeled as a mixture of a REE-enriched melt and entrained peridotite. Fraser and Hawkesworth (1992) consider that for small degrees of melting the SmlNd ratio of the melt is less than that of the source, and LalNd is greater. They conclude that the melt component must have been derived from peridotite with SmlNd ratios (0.25) different from those of the proposed entrained peridotite (SmlNd = 0.2).
290
3000
E 2000 a. a.
1000
o
-
-
-
005
PERIDOTITES --.. ,'" ..... _-- ;', .... 7----/ I I ',I /' I I 'I I v,,'''
". ",. I • FINSCH
/ KlM'rUTES MELT I . ,
I I I I 010 015
Sm/ Nd
020 025
CHAPTER 3
Figure 3.28. Ni versus SmlNd for Finsch orangeites. Solid line represents mixtures between peridotite and melts formed from primary mantle (after Fraser 1987).
Orangeites from Barkly West form an array plotting at a high angle to the mixing line defined for Finsch (Figure 3.29) by Fraser and Hawkesworth (1992), and are considered by Tainton (1992) to have been formed from melts with distinct incompatible element ratios. Orangeites from Star, New Elands, and Swartruggens define discrete fields different from those of Barkly West and Finsch on Figure 3.29. These data might suggest that different orangeites are derived from sources of different incompatible element content. Tainton (1992) considers that the correlation between La/Nd and Sm/Nd observed for Finsch (Figure 3.29) is due to the relatively low abundances of these elements in the Finsch parental magma. Hence, ratios were more susceptible to mixing processes, and initial La/Nd ratios were considerably higher than shown in Figure 3.29.
All of the above conclusions stem from the following assumptions: the composition of the assimilate is known; there is no significant crystal fractionation of the primary melts; these magmas are produced by small degrees of partial melting. The latter point is addressed further in Section 4.4.2.
As the magmas ascended from a minimum depth of 150 lan, there is ample opportunity for sampling a very wide range of mantle material, and contribution to the olivine macrocryst suite from many sources must occur. These may range from dunites to di verse metasomatized peridotites. Modeling of the amounts of incompatible elements added to the melt may not be as simple as proposed by Fraser and Hawkesworth (1992) and Tainton (1992), as such peridotites may have been cryptically or patently metasomatized (Dawson 1984). Thus, the mineralogical sites of incompatible elements in the assimilate may vary. Preferential extraction of intergranular constituents is clearly more probable than complete assimilation of refractory single crystals, and significant amounts of incompatible elements may be added to the melt without concomitant Ni addition.
Regarding crystal fractionation, some important physical and thermodynamic aspects of assimilation are not considered in geochemical models. If peridotite xenoliths are dissolved in small-volume magmas, heat is required to assimilate large quantities of
GEOCHEMISTRY OF ORANGEITES
"C Z
2·5
2·0
" 1'5
~
1'0
A SOVER NORTH • FINSCH D POSTMASBURG
Melt
Peridotite
0'5 -f----...----.---------r---'
0·05 0'10 0'15 0'20
Sm/Nd
291
Figure 3.29. La/Nd ratio versus SmlNd ratio for orangeites. Data sources as in Figure 3.5. Solid line (after Fraser 1987) represents mixtures between peridotite and melts formed from primary mantle.
orthopyroxene. This energy must be provided by crystallization from the magma of the primary liquidus phases stable at the P-T of assimilation. These phases may be subsequently removed (or concentrated) by flow differentiation and/or gravitational fractionation. The complete absence of orthopyroxene xenocrysts in orangeites indicates that the assimilation process must be complete, with the implication that magma ascent rate through the mantle must be relatively slow in order for this process to occur. Slow ascent rates will provide ample opportunity for crystal fractionation of both xenocrysts and primary minerals.
Finally, the small amounts of orangeite emplaced at high crustal levels suggests that if the magmas have indeed assimilated significant quantities of orthopyroxene, then initial magma volumes must have been large in order to provide the heat required for the partial assimilation of the postulated amounts of peridotite contaminant. Much of the initial magma, derivative melts, and many of the crystals precipitated during assimilation must be retained in the mantle.
If the above conclusions are correct, then orangeites are unlikely to represent unmodified primary melts. Alternatively, the orthopyroxene assimilation problem may be avoided if orangeite magmas are simply primary melts mechanically contaminated only by dunites. The presence of diamond and subcalcic chrome pyrope xenocrysts in orangeites is definitive evidence that disaggregation and/or assimilation of garnet-bearing ultramafic xenoliths has occurred (see 4.5.3).
292 CHAPTER 3
3.8. RADIOGENIC ISOTOPES
3.8.1. Strontium and Neodymium
Studies of the composition of radiogenic Sr and Nd by Smith (1983) provided the initial impetus for the confirmation of orangeites as a distinct magma type (1.1). This and all subsequent studies (Fraser et al. 1985, Fraser 1987, Fraser and Hawkesworth 1992, Tainton 1992, Clarke et al. 1991, Skinner et al. 1994) consider that the measured isotopic compositions of the bulk rocks must be very similar to those of their sources. This assumption is based upon the very high Sr and N d contents of the rocks and on calculations showing that either peridotite mixing or crustal contamination will have insignificant effects on Sr and Nd isotopic compositions. These conclusions are in accord with current opinion (McCulloch et al. 1983, Vollmer et al. 1984, Fraser et al. 1985, Nelson et al. 1986), which holds that the isotopic compositions of mantle-derived potassic magmas of high Sr and REE content reflect those of their mantle sources.
Interestingly, the first Sr isotopic studies of rocks from Swartruggens (Mitchell and Crocket 1971, Allsopp and Barrett 1975) established that they possessed high 87Sr/86Sr initial ratios. Unfortunately, the significance of these observations was not realized, as similar high initial ratios were contemporaneously determined in contaminated and altered archetypal kimberlites.
Despite the importance of Sr and Nd isotopic studies there are few published data for the Swartruggens (Smith 1983) and none for the Boshof, Winburg, and Kroonstad occurrences.
Measured and calculated initial 143Ndll44Nd ratios obtained in separate laboratories are commonly not directly comparable owing to the use of different standards and analytical techniques (Hawkes worth and van Calsteren 1984). In an attempt to overcome this problem, initial Nd isotopic ratios are commonly expressed as the deviation from the bulk earth chondritic value of 143Ndll44Nd at the time of formation (1) of the samples by the relation
[143Ndll44Nd sample initial ratio (1)
ENd = 143Ndll44Nd CHUR (1)
where CHUR is the isotopic composition at time (1) of a chondri tic uniform reservoir used to represent the SmlNd ratio and isotopic compositions of the bulk earth (DePaolo and Wasserberg 1976, O'Nions et al. 1979). Values of ENd of zero or near zero in mantle-derived rocks indicate undifferentiated primitive mantle sources in terms oftheir SmlNd ratios. Positive or negative values require that at least one episode of fractionation has increased or decreased the source SmlNd relative to the chondritic ratio. The ENd
values are particularly useful when attempting to compare the isotopic compositions of kimberlites, orangeites, and lamproites.
Table 3.11 lists the range of Sr and Nd isotopic compositions found in orangeites. Fortunately, all Nd isotopic determinations have been corrected for instrumental isotopic fractionation using the same 146Ndll44Nd ratio (0.7219) and are, thus, directly comparable.
GEOCHEMISTRY OF ORANGEITES 293
Table 3.11. Initial Sr, Nd, and Pb Isotopic Compositions of Orangeites 87Sr/86Sr 143Ndll44Nd ENd (N)
Finsch 0.70777-0.70983 0.51202-0.51217 -6.2 to-9.7 (23) Posbnasburg/Pniel 0.70718-0.70886 0.51190-0.51216 -7.6 to-l2.5 (7) SoverNorth 0.70713-0.70721 0.51180-0.51181 -13.1 to-l3.4 (3) Sover 0.70741-0.70776 0.51195-0.51196 -10.1 to-lO.4 (3) Bellsbank 0.70847-0.70896 0.51204-0.51206 -9.4 to-9.6 (3) Newlands 0.70764-0.70773 0.51202-0.51206 -9.3 to-10.3 (2) Prieska 0.70755-0.70868 0.51190-0.51208 (II)
NewElands 0.7074-0.7076 0.51208 (1-4)
Swartruggens 0.7090-0.7109 (4)
206pbP04Pb 207PbP04Pb 208pbP04Pb (N)
Finsch 17.69-18.24 15.44-15.58 37.48-38.23 (23) Posbnasburg/Pniel 17.298-17.539 15.453-15.517 35.967-37.681 (7)
Sover North 17.062-17.310 15.441-15.498 37.445-37.602 (3) Sover 17.365-17.428 15.465-15.529 37.442-37.658 (4) Bellsbank 17.440-17.542 15.487-15.537 37.540-37.736 (6) Newlands 17.479-17.550 15.514-15.538 37.657-37.674 (3) New Elands 17.21-17.26 15.47-15.48 (4) Swartruggens 17.63 15.62 (1)
Data sources: Finsch (Fraser 1987, Fraser and Hawkesworth 1992); PostmasburglPnie1, Sover North, Sover, Bellsbank, New1ands (Tainton 1992); Prieska (Skinner et al. 1994); New E1ands, Swartruggens (Smith 1983). (N) = number of samples. All Nd isotopic data are corrected for fractionation to l~dll''Nd = 0.7219.
Table 3.1 and Figure 3.30 show that within individual orangeite intrusions there is a limited but significant range in Sr and Nd isotopic composition not attributable to experimental error. Isotopic differences are also recognizable between groups. Thus, the Finsch orangeites contain relatively radiogenic Sr and Nd and appear to have been derived from sources that had slightly greater time-averaged Sm/Nd and Rb/Sr ratios than those of the Barkly West area. There are no correlations between the Sr and Nd isotopic compositions of the Finsch orangeites. Within the Barkly West group the evolved Sover North orangeites have the least enrichment in radiogenic Sr and greatest depletion in radiogenic Nd. In contrast, evolved rocks in the Postmasburg area are slightly enriched in radiogenic Nd relative to the Barkly West samples. The data suggest no simple fractional crystallization or partial melting relationship exists between evolved and unevolved orangeites. As a group, orangeites from the Barkly West-Postmasburg-Finsch areas define a weak correlation of decreasing 87Sr/86Sr with decreasing 143Ndll44Nd. Bellsbank orangeites show no intra-intrusion variation in 143Ndll44Nd, although Sr isotopic compositions vary widely. Tainton (1992) has suggested that this variation may be attributable to the introduction of ground water and carbonate country rock. Bonafide orangeites from the Prieska area have isotopic compositions overlapping (Figure 3.30) both the Finsch and Barkly West orangeites.
Figure 3.31 illustrates the Sr and Nd isotopic composition of orangeites relative to those of bulk earth and archetypal kimberlites, and shows clearly that each magma type must originate from compositionally distinct sources. Thus, Smith (1983) and subsequent
294
'"0 Z
'it
0-5122 -
0-5121 -
! 0-5120-....... '"0
f()Z 0-5119 -'it
0-5118 -
•
rpRlESKA7 - ;- - II I • • • • • rI. 1 I • .... 0.. 1 0 1 I 0 ~ ·0. 1
o 0 1 .1 o. • 0
I • I
•
L _______ ....J
o BELLSBANK a NEWLANDS • SOVER o PNIEL a PK35 ... SOVER NORTH a PK 36 t::. PK37
CHAPTER 3
- -7-5
f- -10-0
eNd
t- -12-5
_l-_----r __ -._---L...,..._F_I_NrSC_H_--. __ -. __ ,....t- -15-0 0-5117 I I I I I I I
0-7065 0-7075 0-7095
Figure 3.30. 143Ndll44Nd versus 87Sr;86Sr for orangeites from the Finsch (Fraser 1987) and Barldy WestPostmasburg region (Tainton 1992)_ Field for Prieska orangeites after Skinner et ai_ (1994).
workers have concluded that orangeites are derived from ancient sources with greater Rb/Sr and lesser SmlNd ratios than those of bulk earth. Kimberlites, on the other hand, appear to be derived from sources with compositions close to that of bulk earth or which have higher SmlNd and lower Rb/Sr ratios.
-0-5128 -
-~ 0,5126-
'it -! 0'5124-
" -~ 0'5122-
f() -
'it 0'5120--
0'5118 -
o
/+~ BULK EARTH 0
o KIMBERLITE • ORANGEITE o ANOMALOUS
PRIESKA
o
n.......n .' '-1...J:T"" ~" .: ...... \ ..
•• •• . .:.:,.'. ' . . . :. . '. '" •
I I I I I I I I 0'704 0'706 0'708 0'710
87Sr / 86Sr
Figure 3.31. 143Ndll44Nd versus 87Sr;86Sr for kimberlites (Smith 1983, Skinner et al. 1994),orangeites (Fraser 1987, Tainton 1992, Skinner et al. 1994), and anomalous Prieska samples (Skinner et al. 1994).
GEOCHEMISTRY OF ORANGElTES 295
The conventional interpretation of the isotopic data (Smith 1983) is that the orangeite sources were isolated for 1-2 billion years prior to the partial melting events which led to the formation of orangeite magmas. These sources are considered to be located in the nonconvecting continental lithospheric mantle. In contrast, kimberlites are believed to be derived from convecting asthenospheric mantle. Note that the data set in Figure 3.31 includes "on"- and "off'-craton kimberlites. These cannot be distinguished from each other on an isotopic basis. This observation indicates derivation of the magma from similar sources and the absence of contamination of "on" -craton kimberlites by ancient radiogenic Sr-enriched cratonic crust during emplacement.
Figure 3.31 also shows that some rocks from the Prieska area have isotopic signatures which are intermediate between those ofkimberlites and orangeites. Clarke et al. (1991)
and Skinner et al. (1994) refer to these rocks as "transitional kimberlites." They occur primarily in domain V of the Prieska region (Figure 1.12). Petrographically, they differ from bona fide orangeites in containing relatively abundant, coarse-grained, primary spinels, and perovskite. Skinner et al. (1994) and Skinner (1989) consider these features place the rocks as petrographically transitional between kimberlites and orangeites. However, the rocks have other characteristics, such as the presence of amphibole and sanidine, linking them to orangeites. Lacking detailed mineralogical investigations of these rocks, it would seem unwise and premature to recognize a new category of "transitional kimberlites" based upon the grain size of the ground mass minerals and only six determinations of isotopic composition. Further studies of these rocks are desirable.
8
-16
-24
PRAIRIE
MADUPITIC CO; LAMPROITE( LEUCITE I' 0
HILLS ~'\
G:J,G:J~ ,ORANGEITES
O PHLOGOPITE LAMPROITE _ SMOKY LEUCITE HILLS
BUTTE
0·700 0·705 0·710 0·715 0·720
875r/865r
Figure 3.32. Isotopic composition of Nd and Sr in orangeites. lamproites. kimberlites. and potassic volcanic rocks. Data sources as Figure 3.31 and Mitchell and Bergman (1991).
296 CHAPTER 3
As a group, orangeites exhibit limited 87Sr/86Sr (0.707-0.710) and 143Nd/I44Nd (0.5118-0.5122) compared to many other mantle-derived potassic volcanic rocks (Figure 3.32). They share with these magmas the trait of derivation from ancient sources, having lower Sm/Nd ratios and greater Rb/Sr ratios than bulk earth. Figure 3.32 shows that the majority of orangeite isotopic compositions plot on the West Kimberley/Murcia-Almeria isotopic trend, close to the composition of Ellendale olivine lamproites. Others (Sover North, Postmasburg, Sover, Bellsbank) plot between the West Kimberley trend and that defined by North American lamproites. These data suggest derivation of orangeites from sources of varying Sm/Nd and Rb/Sr ratios, intermediate in character between those of the low Rb/Sr and Sm/Nd North American lamproites and the high Rb/Sr, low Sm/Nd of the Australian lamproites.
Isotopic data may be intetpreted to suggest that intra- and inter-intrusion variations represent derivation from isotopically heterogeneous sources (Bergman 1987, Tainton 1992). In this case discrete domains of diverse Rb/Sr and Sm/Nd would have existed for long time periods (1-2 Ga). Partial melting of these regions, without mixing of the derivative magmas, is then required to explain the observed isotopic variation. While this process may be applicable to large lithospheric domains and regional inter-intrusion isotopic differences, it would seem less so to intra-intrusion variations, as considerable magma mixing must occur during their generation and emplacement.
An alternative explanation of the isotopic variations proposes that they are the products of mixing of two (or more) components of radically different isotopic composition (McCulloch et al. 1983, Vollmer et al. 1984, Mitchell et al. 1987). In this case the isotopic compositions do not necessarily represent those of their source rocks, as observed compositions are intetpreted as mixtures between depleted and enriched end-member components. The sources of these components may be found in the asthenosphere and the lithosphere, respectively. Hence, it is possible that the addition of asthenospheric material to the lithosphere triggers partial melting of ancient Rb- and light REE-enriched zones, leading to the formation of diverse potassic magmas containing hybridized Sr and Nd (see 4.5.2).
Isotopic variation within the Finsch orangeites is attributed by Fraser (1987) and Fraser and Hawkesworth (1992) to mixing of melts derived from enriched source regions with overlying depleted mantle. However, neither details of the isotopic mixing processes nor the cause of partial melting are specified.
3.8.2. Lead The isotopic composition of Pb in orangeites (Table 3.11) is unradiogenic with
respect to 206PbP04Pb ratios and plots to the left of the geochron and slightly below the Stacey and Kramers (1975) two-stage Pb growth curve in Figure 3.33. Each intrusion differs with respect to 206PbP04Pb-207PbP04Pb and 208PbP04Pb-206pb/204Pb ratios, and samples define distinct linear arrays on Figure 3.33. These arrays do not represent errorchrons, isochrons, or chords (anomalous Pb lines) to the growth curve.
Figures 3.33 and 3.34 also show that the isotopic composition of Pb in archetypal kimberlites is distinct from that of orangeites, implying that their parental magmas
GEOCHEMISTRY OF ORANGEITES
15'8
..c 15·6 a..
v o C\J
" 15-4 ..c a.. ,....
o C\J 15.2
15'0
14
W. KIMBERLEY 0 GAUSS BERG -0
LEUCITE HILLS
ORANGEITES /
PRAIRIE CREEK (f 15·60
N.~ ,\)~ c;;?
fit I....~ C; 15'50
<J;-~ SMOKY 15-45 o BUTTE
15-40
15 16 17 18
297
--_. -, .,...,.. . .,...,.. 4.l,. ·
• •• 6. • •
19 20
Figure 3.33. 207pbP04pb versus 206pbP04pb for orangeites (Smith 1983, Fraser 1987, Tainton 1992, Skinner et al. 1994), southern African kimberlites (Smith 1983), and W. Kimberley, Murcia-Almeria, Gaussberg, Leucite Hills, Prairie Creek, Smoky Butte, and Sisimiut lamproites (Nelson 1989, Alibert and Albarede 1988, Fraser 1987, Nelson et al. 1986, Fraser et al. 1985). Growth curve is from Stacey and Kramers (1975). Field of compositions of basalts from mid-oceanic ridges (MORB) and oceanic islands (OIB) is from Fraser (1987). Inset shows isotopic arrays relative to the growth curve (dashed line) for suites of orangeites from Finsch and the Barkly West region (after Tainton 1992). Symbols as in Figure 3.30.
originated from sources with very different U/Pb and Th/Pb ratios and Pb evolutionary histories.
Figure 3.33 shows that the only other potassic magmas which have Pb isotopic compositions similar to those of orangeites are lamproites from the Leucite Hills. The major problem in interpreting these Pb isotopic data is explaining the combination of unradiogenic 206PbP04Pb ratios with relatively radiogenic 207PbP04Pb ratios. Unfortu
nately, interpretation ofPb isotopic data is not unambiguous and subject to the prejudices of the interpreter. Hence, some geochemists favor mixing models over multistage growth
models, and vice versa. Common to most interpretations is the conclusion that orangeite
and lamproite Pb that plot to the left of the geochron have undergone evolution in a region
of low UlPb for a significant length of time.
Tainton (1992), Fraser (1987), and Fraser and Hawkesworth (1992) consider that the
intra-intrusion Pb isotopic arrays do not result from mixing of mantle Pb with lower
crustal Pb derived from granulites, as the latter does not have sufficiently high Pb contents and is highly radiogenic.
298
40~---------------------------,
39 .Q
a. ~ o N ....... 38 .Q a. ORANGEITES
CD o N
37
15
1,0 Go
SMOKY ~ PRAIRIE CREEK BUTTEU /
16 17 18 19 20
CHAPTER 3
Figure 3.34. 208pbl204pb versus 206pbl204pb for orangeites. kimberlites. and lamproites. Data sources as in Figure 3.33.
Fraser (1987) and Fraser and Hawkesworth (1992) have interpreted the Finsch Pb array as resulting from mixing of the magma with entrained peridotite having unradiogenic Pb isotopic compositions similar to those of diopsides separated from peridotites by Kramers (1977, 1979). However, several isotopic varieties ofPb exist in peridotites, and the mixing process cannot be as simple as that envisioned by Fraser (1987). Thus, bulk rock Pb isotopic compositions of peridotites cannot equal those of diopside, as coexisting garnet and enstatite may contain radiogenic Pb not in equilibrium with diopside (Gunther and Jagoutz 1991). Tainton (1992) has noted that because of very low contents of Pb in diopside «0.2-2 ppm), any bulk mixing model requires the addition of inordinately large amounts of peridotite to generate the observed Finsch Pb isotopic array. A further argument against this mixing model is that the Finsch Pb isotopic compositions do not correlate with the ratios of other incompatible elements.
Tainton (1992) also considers that the Pb isotopic arrays within individual orangeites represent mixing lines. To explain the Pb isotopic compositions a two-stage growth model is proposed, in which differences in the UlPb ratios of the magma sources during the second stage are required to explain the inter-intrusion isotopic differences. The Pb isotopic arrays are then generated by the mixing of these isotopically distinct melts with unradiogenic Pb derived from depleted peridotite. This process is not one of simple mixing, and Tainton (1992) suggests that it is "a more complex process involving partial re-equilibration of the lead isotopic composition of the melt with peridotite wall-rock" (Tainton 1992, p. 147). Specific details of the physicochemical nature of this process are not provided.
3.9. STABLE ISOTOPES
There have been few studies of the composition of hydrogen, carbon, and oxygen in orangeites, and most have concentrated upon carbonates while the stable isotopic com-
GEOCHEMISTRY OF ORANGEITES 299
position of other minerals is unknown. Initial studies were undertaken on a few isolated samples as part of general investigations ofthe isotopic compositions of kimberlite sensu lato and other mantle-derived carbonates. Sheppard and Dawson (1975) give 0180 = 10.06%0 and 013C = -7.06%0 for calcite from New Elands and 0180 = 11.43%0 and o\3C = -5.97%0 for dolomite from Sover. The isotopic composition ofH in the carbonate-free matrix of the Sover sample was oD = -98%0. Kobelski et af. (1979) give 0180 of 11.04%0 and 12.23%0 and o\3C of -7.61%0 and -6.81%0 for the Star and Roberts Victor orangeites, respectively. Too few samples were analyzed in these studies to determine whether any significant differences exist between the stable isotope composition of archetypal kimberlites and orangeites.
The work of Kirkley et al. (1989) represents the only detailed investigation of the stable isotopic composition of orangeites, and was limited to the determination of the whole rock isotopic composition of carbon and oxygen. Unfortunately, no attempt was made to analyze individual carbonates, thus creating a major difficulty in the interpretation of the data, as several varieties of carbonate are known to occur in the orangeites (this work, Tainton 1992) and kimberlites (Mitchell 1994b) investigated. The problem is especially significant for samples containing major modal amounts of dolomite and calcite. Thus, whole rock isotopic compositions are those of mixtures which reflect the relative proportions of these minerals. Contrary to the assertions of Kirkley et al. (1989), the data may have absolutely no petrogenetic significance, especially if the calcite and dolomite are not in isotopic equilibrium. This situation might arise if either mineral has formed during post-intrusion dolomitization or serpentinization. Unfortunately, Kirkley et al. (1989) did not undertake a complementary investigation of the carbonate mineralogy of the samples analyzed.
Figure 3.35 shows that though there is wide variation in the C isotopic compositions of orangeites, most samples fall within the limits established for primary mantle-derived carbon. In contrast, 0 isotopic compositions exhibit significant variation and are enriched in 180 relative to mantle carbonate values. Kirkley et al. (1989) note that the Finsch and Bellsbank orangeites which are emplaced in dolomitic country rocks are enriched in 180 and depleted in \3C relative to orangeites emplaced in other terrains (Figure 3.35). Orangeites from Swartruggens have similar isotopic compositions but are emplaced in lavas of the Pretoria series. However, the intrusions may have passed through dolomitic horizons at depth.
The data suggest that none of the 0 isotopic compositions represent those of the primary magmas. This conclusion is based upon H, C, and 0 isotopic studies of carbonatites and kimberlites (Deines 1989, Kobelski et aZ. 1979, Sheppard and Dawson 1975), which show that enrichment in 180 results from influx of meteoric water or re-equilibration of the carbonates with magmatic waters at low temperatures.
The above interpretation of the isotopic data is in direct contrast to the hypothesis presented by Kirkley et al. (1989)-that the isotopic compositions of Finsch, Bellsbank, and Swartruggens are due to the assimilation of dolomite enriched in 12C and 180. They suggest that the dissociation of dolomite produces C02 that mixes with mantle-derived C02, resulting in the crystallization of the 12C_ and 180-enriched ground mass carbonates. There is no petrographic or geochemical evidence to support the operation of this process, especially at Swartruggens, and the temperatures of the orangeite magmas suggested by
300
o
- 2 ---h~T"'7'''
co ~ -4
U If) -6 CO
-8
-10
5
o
CARBONATITES + +
co<?o o 0 00
o
• •• 0 00 0
I I I I
o 0 :, o
FIELD OF KIMBERLITES
10 15 20 25
Sl80SMOW
CHAPTER 3
Figure 3.3S. a13c versus al80 for orangeites (after Kirkley et al. 1989). Solid points are for orangeites emplaced in dolomitic country rocks. Open circles are orangeites emplaced in non-dolomitic terrains. Crosses are the isotopic composition of Chuniespoort dolomite. Compositional fields of carbonatites and kimberlites are from Deines and Gold (1973). Kobelski ef al. (1979). and Kirkley ef al. (1989).
Kirkley et al. (1989) are unrealistically high (1000-125()QC). Tainton (1992) has further noted that increasing amounts of deuteric dolomite at Bellsbank are accompanied by decreasing amounts of serpentine. These minerals are unlikely to be in isotopic equilibrium with each other or primary carbonates.
Kirkley et al. (1989) further claim that orangeites emplaced in dolomitic rocks have higher initial 87Sr/86Sr ratios than those emplaced in other rocks. While their observation appears to be correct, it does not follow that the high ratios are a consequence of the bulk assimilation of dolomite, as proposed to explain the e and 0 isotopic variations. Tainton (1992) has suggested that assimilation of the amounts of dolomite required by Kirkley et al. (1989) are petrologically unreasonable given the very low Sr contents of dolomite relative to those of orangeites. Ground-water alteration as proposed by Barrett and Berg (1975) provides a simpler and more realistic explanation of the high Sr isotopic ratios. In conclusion. the observed 0 isotopic compositions of these orangeites are considered to be due to exchange with l80-enriched ground waters derived from the dolomitic terrains rather than bulk assimilation of dolomite.
The e and 0 isotopic compositions of orangeites are not very different from those of archetypal kimberlites (Figure 3.35). The stable isotope composition of kimberlites has been reviewed by Deines (1989) and Mitchell (1986). Most investigators (Deines 1989, Kobelski et al. 1979, Sheppard and Dawson 1975) have concluded that the carbon in kimberlites is typical mantle-deri ved carbon and that the oxygen isotopic compositions have been modified by interactions with ground water.
There are, on the basis of the existing data, absolutely no grounds to support the hypothesis of Kirkley et al. (1989) that the mantle sources of orangeites are depleted in I3e relative to those of kimberlites. This conclusion is based upon the wide range in the isotopic composition of kimberlites, which completely overlap those of orangeites
GEOCHEMISTRY OF ORANGEITES 301
(Figure 3.35), and the above observation that there is no reason to believe that the whole rock C isotopic composition of orangeites represents that of the primary magma. Further, Kirkley et al. (1989) do not provide either calculations showing the distribution of carbon between carbon-bearing compounds during formation and evolution of the magma as a function of redox conditions, or information on the mineralogical control on whole rock isotopic compositions. Provision of these data is essential if the origins of the variations in the C and 0 isotopic compositions of orangeites are to be understood.
3.10. SUMMARY
Orangeites may be divided into unevolved and evolved types on the basis of their bulk rock major element geochemistry. Unevolved orangeites are ultrapotassic, peralkaline, perpotassic, ultrabasic rocks characterized by high and variable C02 and H20 contents. Bulk rock compositions are controlled by accumulation and fractionation of primary minerals and contamination with mantle-derived peridotites and cannot represent those of parental liquids. Evolved orangeites are richer in Si02 (41-47 wt%) than unevolved orangeites «40 wt% Si02) and are not characteristically peralkaline. Bulk rock compositions do not represent those of parental liquids as a consequence of hybridization with mantle-derived peridotites and crystal fractionation.
Orangeites are enriched in REE, Zr, Nb, Sr, Ba, Rb, and other incompatible elements, and Ni relative to common mantle-derived magmas. Enrichment in Ni (and MgO) reflects the presence of large amounts of xenocrystal olivine. Incompatible element abundances vary widely as a result of variations in the modal amounts of primary phlogopite, apatite, and carbonate and/or xenocrystal olivine. Thus, high phlogopite contents result in whole rock compositions characterized by high abundances of Rb and Pb, whereas apatite- and carbonate-rich rocks are relatively richer in Sr and REE. Sequestration of these trace elements in early crystallizing primary phases implies that they are not strictly incompatible elements during the crystallization of orangeites. The presence of xenocrystal olivine merely serves to dilute incompatible element abundances and has no effect upon their ratios or isotopic composition. All orangeites are strongly enriched in the light REE (typically La/Yb > 100). Chondrite normalized REE distribution patterns do not exhibit any Eu or Ce anomalies. Primitive mantle normalized extended incompatible element distribution diagrams display significant negative K, Rb, and Sr anomalies, but lack the negative Ta, Nb, and Ti anomalies characteristic of many lamproites and potassic rocks. Sr and Nd isotopic studies suggest that orangeites have been derived from ancient sources with lower SmlNd and higher Rb/Sr ratios than those of bulk earth. These sources are considered to be located in nonconvecting lithospheric mantle. Pb in orangeites is notably unradiogenic, suggesting evolution in a regime of low U/Pb ratio for an extended time. No simple relationships between evolved and unevolved orangeites can be recognized on the basis of trace element and isotopic studies. However, evolved orangeites appear to have lower REE and Nb contents, La/Yb ratios, Pb/Ce ratios, and higher ZrlNb ratios than unevolved types.
Unevolved C02-rich orangeites are geochemically unlike kimberlites and olivine lamproites with respect to their major and trace element geochemistry. Evolved orangeites are similar in their major element geochemistry to olivine and/or madupitic lamproites,
302 CHAPTER 3
but may be distinguished from these on the basis of the higher Ba and Zr contents of the latter.
Orangeites, in terms of their isotopic compositions, differ greatly from archetypal kimberlites and most other asthenosphere-derived basaltoid magmas. Isotopically they are similar to lamproites and other continental potassic volcanic rocks, implying derivation from similar metasomatically enriched lithospheric mantle sources.
The Finsch orangeites are anomalous in that they are petrographically similar to unevolved orangeites, whereas their geochemical signature is that of an evolved orangeite.