In order to know an object. I must not know its external but all its internal qualities
Ludwig Wittgenstein (1922) Tractatus Logico·Phiiosophicus
MINERALOGY OF ORANGEITES
The mineralogy of orangeites has not been extensively studied as data for these rocks have typically been incorporated with studies of kimberlites (sensu lato). Thus, previous studies did not attempt to delineate mineral assemblages and compositional trends which could be used to discriminate between each group of rocks. This chapter is based upon many new analyses of minerals in archetypal kimberlites and orangeites. The study of the latter is far from definitive as detailed studies have been made only of the major silicate and oxide minerals and not all occurrences of orangeites have been examined. The minor and accessory minerals require much further investigation, and the data presented here should be regarded merely as a reconnaissance study.
The information presented in this chapter summarizes all that is currently known about the mineralogy of orangeites and compares these data with the parageneses and compositions of similar minerals in kimberlites and lamproites. The minerals, with the exception of olivine, are discussed in their approximate sequence of crystallization. Particular emphasis is given to mica, as this mineral dominates the mode of most orangeites and exhibits extensive compositional variation.
Xenocrysts, including diamond, are not discussed, as this work is concerned only with primary minerals. Megacrysts belonging to the Cr-poor suite, which are characteristic of archetypal kimberlites, are discussed primarily in Sections 1.5 and 4.6.2, as they are only rarely found in orangeite (see 1.8.8,2.2.3).
The locations of the orangeites referred to in the text are given in Figure 1.3. Individual occurrences are described in Section 1.8.
2.1. MICA
2.1.1. Paragenesis
Mica is ubiquitous in orangeites. Typically, primary mica forms closely packed mosaics of tabular-to-square cross section, pale-brown, weakly pleochroic microphenocrysts (0.01-0.3 mm). Modal abundances (30-90 vol %) vary widely within and between intrusions and primarily reflect variations in the macrocrystal olivine content. Flow alignment is present in many occurrences. Microphenocrysts may be zoned from colorless or pale brown cores, to margins exhibiting stronger brown or red-brown pleochroism. Zoning may be continuous or discrete. Many of the microphenocrysts are
mantled by groundmass, red, reversely pleochroic tetraferriphlogopite. Some microphenocrysts consist of discrete cores containing abundant vermiform fluid inclusions and colorless, inclusion-free mantles. Glomeroporphyritic aggregates of microphenocrysts are common.
Groundmass micas occur as small (0.02-0.1 mm) strongly pleochroic brown-to-redbrown subhedral crystals. Many comprise the outermost mantles or rims of microphenocrysts, although discrete subhedral-to-euhedral crystals are commonly found throughout the mesostasis. The outer reversely pleochroic tetraferriphlogopite-rich margins of the crystals, which are immersed in the mesostasis, may be euhedral and fresh or ragged and resorbed. Unlike the microphenocrysts, groundmass micas typically contain inclusions of spinel, apatite, and diopside.
Large (0.2-0.5 mm) poikilitic plates of groundmass mica are found only in evolved orangeites, e.g., Sover North, Besterskraal. These are strongly pleochroic from yellowish red to deep red-brown and may be simply twinned. Inclusions of earlier-formed crystals
AI
I EAST *----- ---- -~SID
I
~I 0
~-:::----x~ / o~ ~ ~ 0
/ -/ 1 00 0
0
70 .~ ~,. •• I 0
()) PHL .~ •• 00
~ 75 *,---... :. ~~~ .. ~--.-- -if - .... ANN lao \ · .. it I ~ AI , • I. 0 -
\ ••• • C!>/I-• • • 0 0
\ •••• • ilin
\ ~" ••• I PHL \ .. : 1 0.1 EAST
• .: I Mg L-.......... __ ---" F 1~--------"7 I TFP TFA eT • MICROPHENOCRYSTS \
AND GROUNDMASS • " ~. \ .,.
T .:J FP • •• ., • • Mg t..=.===r:===:::;:====:;:::==~--* --:-• ....;.....,/ ----./..----...,/~---I.~ FeT
5 10 15 20 25 30 35 40 45
Figure 2.1. Compositions (atomsl22 oxygens) of micas from orangeites plotted in the ternary system AI-MgFer. Total Fe expressed as Fe2+. EAST = "eastonite." SID = siderophyllite. PHL = phlogopite. ANN = annite. TFP = tetraferriphlogopite. TFA = tetraferriannite.
MINERAWGY OF ORANGEITES 93
Table 2.1. Representative Compositions of the Cores of Microphenocrystal Phlogopites of the Extreme AI-Depletion Trenda
aFeOr = total Fe expressed as FeO; n.a. = not analyzed; 1-2, Sover Mine; 3-4, Lace; 5-6, New Elands; 7-8, Swartruggens level 6; 9-10, Finsch F4 and F7 intrusions, respectively. All data this work.
are characteristic. In these rocks mica also occurs as reaction mantles around olivine macrocrysts.
The common macrocrysts (>0.5 mm) are irregular, rounded, distorted, and weaklypleochroic pale-brown phlogopites which are optically identical to the microphenocrysts. These macrocrysts are interpreted (2.1.7) as having crystallized from orangeitic magmas prior to the emplacement of their current host and are thus considered to be cogenetic with the microphenocrysts.
Other, less common, cryptogenic macrocrysts include strongly-pleochroic very darkbrown or dark-green biotites. These are commonly mantled by pale-brown phlogopite which is optically and compositionally identical to phlogopite microphenocrysts and macrocrysts.
Detailed descriptions of mica paragenesis in orangeites may be found in Smith et ai, (1978), Skinner and Scott (1979), Mitchell and Meyer (1989a), and Tainton (1992). The
94 CBAPTER2
Table 2.2. Representative Compositions of Groundmass Micas of the Extreme AI-Depletion Trenda
"FeOr = total Fe expressed as FeO; n.d. = not detected; n.a. = not analyzed. 1-2, Finsch internal dike F4;~, Saltpeterpan; 5-7, Swanruggens level 6; 8, Lace; 9, Star; 10, New Elands. All data this work.
paragenesis of mica is discussed further in Section 1.10 and illustrated in Figures 1.26-1.54.
2.1.2. Composition of Primary Mica
Although of a reconnaissance nature, the study of nine random samples of orangeite by Smith et al. (1978) served to establish the overall character of the micas present. The principal findings were the recognition of relatively rare dark-colored Fe-rich mica (termed type I) of unknown origin, and common microphenocrystal micas zoned from phlogopite cores (termed type II) to tetraferriphlogopite margins. Subsequent reconnaissance studies of the Bellsbank dikes by Boctor and Boyd (1982), and detailed studies by Skinner and Scott (1979) and Mitchell and Meyer (1989a) of the Swartruggens and New
MINERALOGY OF ORANGElTES 9S
Table 2.3. Representative Compositions of Micas from Besterskraal and Sover NorthQ
"FeOT = total Fe expressed as FeO; 1-3, groundmass poikilitic mica, Besterskraal; 4-6, groundmass poikilitic mica, Sover North; 7-8 and 9-10, cores and rims of microphenocrystaJ mica, Sover North.
Elands occurrences, respectively, confirmed and amplified the observations of Smith et at. (1978). Recently, Tainton (1992) presented compositional data for micas in orangeites from the Barkly West region.
Discussion below of the compositional variation of mica in orangeites is based, in part, upon these earlier studies, but principally upon approximately 1000 new analyses of mica obtained during the preparation of this monograph. In this work, Smith et at. 's (1978) type I micas are termed AI-biotite macrocrysts, and their type II micas are referred to as primary phlogopite macrocrysts and microphenocrysts.
Figure 2.1 illustrates the major element compositional variation exhibited by mica in orangeites. The data plotted are intended to be representative of the principal varieties of mica present and not to reflect their relative abundances. The figure demonstrates that primary microphenocrystal and groundmass micas are members of the phlogopitetetraferriphlogopite series. Compositions plot close to, and parallel to, this Fe2+-free compositional join, as there is only minor solid solution toward biotite. Solid solution
96 CHAPTER 2
Table 2.4. Representative Compositions of Micas Exhibiting the Moderate AI-Depletion Trenda
aFeOr = total Fe expressed as FeO; n.d. = not detected; 1-2,3-4, cores and rims of microphenocrystal micas, Voorspoed; 5--6, 7-8, cores and rims of microphenocrystal micas, Postmasburg PK37; 9-10, groundmass micas, Postmasburg PK36. All data this work.
toward "eastonite," siderophyllite, tetraferriannite, and annite is not significant. Further discussion of the solid solutions present is given in Section 2.1.8.
Aluminous micas are rarely present in orangeites as microxenoliths (see 2.1.3) or
dark-colored macrocrysts (see 2.1.4). The former are aluminous phlogopites, and the latter
aluminous biotites which are intermediate members of a solid solution between phlogopite, "eastonite," and siderophyllite (Figure 2.1).
Orangeite micas exhibit extensive compositional variation with respect to their AI,
Fe, Ti, and Mg contents (Tables 2.1-2.4). Other elements show only minor variation and are not present in significant quantities (see 2.1.2.3). Consequently, mica compositional variation is best illustrated by plots of Ah03 versus Ti02 or FeaT.
MINERAWGY OF ORANGEITES
14
13
12
II
t 10
9
~ 8 -~ 7 ..,
o 6 ..!:' <t
5
4
3
2
LEVEL :3-4
• PHENOCRYSTS o GROUNDMASS
2 3 4
SWARTRUGGENS LEVEL 6
2 3 4
Ti02 wt. %
LEVEL 7
0
.f
1"\ . :. ~ . ., .
o
o
o
00 0 0
0
0
2
~
97
0
0
3 4
Figure 2.2. AI203 versus Ti02 compositional variation of micas from the Swartruggens orangeites. Samples were collected from different levels in the mine and represent distinct intrusions in this multiple dike system.
2.1.2.1. Ah03-TI02 Variation
Tables 2.1, 2.3, and 2.4, together with Figures 2.2 to 2.7, show that microphenocryst cores typically contain from 13.5 to 9.0 wt % Ah03 and 1 to 3 wt % Ti02. Intra- and inter-intrusion compositional differences are not significant. Thus, the geographically and temporally isolated Swartruggens dikes contain microphenocrysts (Figure 2.2) of similar composition to those in the younger Sover, Lace, and New Elands dikes (Figure 2.3). Micas in samples from different levels of the Swartruggens Mine are identical in composition (Figure 2.2). It is unlikely that these samples were derived from the same dike, given the complexity of the multiple-dike system at Swartruggens 0.8.6), suggesting that individual dikes cannot be distinguished on the basis of mica composition. Similar
98
16
15
14
13
12
t II 10
~ 9 ..: ~
8 If)
0 ~ 7
« 6
5
4
3
2
NEW ELANDS
[] []
o AI-BIOTITE XENOCRYSTS
• MICROPHENOCRYSTS
o GROUNDMASS
o
2 3 4 5 6
LACE
o o
MICROPHENOCRYSTS
+ (03)
• (07)
GROUNDMASS
o 0 (02)
00 o
2 :3 4 5
Ti02 wt. % ~
,
CHAPTER 2
SOVER
X MICROXENOLITHS
o AI-BIOTITE XENOCRYSTS
• MICROPHENOCRYSTS
[] GROUNDMASS
2 3 4 5
Figure 2.3. Al203 versus 1102 compositional variation of micas from the New Elands. Lace. and Sover Mine orangeites.
conclusions were reached by Skinner and Scott (1979) on the basis of fewer data.
Similarly, the different intrusions within the Finsch pipe cannot be distinguished on the
basis of their microphenocrystal mica compositions (Figure 2.4). Microphenocrysts in the Sydney-on-Vaal dike (Figure 2.4) are relatively rich in Ti02 (2.4-4.9 wt %) and Ah03 (11.2-14.2 wt %).
Mitchell and Meyer (1989a) have demonstrated that pale-brown microphenocrysts
rich in fluid inclusions and found in the New Elands dikes, are slightly richer in Ti02
(1.0-3.5 wt %) and FeOT (4.0-7.0 wt %, mg = 0.85-0.91) than coexisting colorless
microphenocrysts (Ti02 = 0.15-2.5 wt %, FeOT = 2.4-6.5 wt %, mg = 0.87-0.96). Tables 2.2 and 2.4 together with Figures 2.2 to 2.4 demonstrate that, in general,
groundmass micas are very poor in Al relative to microphenocrystal micas. Zonation and
MINERALOGY OF ORANGElTES 99
FINSCH
14
? 8
12
t 10
~ ..: 8 + ~
It)
0 ~
6 j <t
4 0 + • 0 F2
+0 • F4
2 • + F4E
• F4ID 0 F7
Figure 2.4. Al203 versus Ti02 compositional • variation of phenocrystal and groundmass micas 0
from the Finsch orangeite and phenocrystal micas l' from the Sydney-on-Vaal orangeite. F-numbers 2 3 4 5 represent petrographically discrete intrusions at
Ti02 Wt. % ---. Finsch.
mantling trends of microphenocrystal mica typically involve depletion in Al at constant or decreasing Ti contents. The outermost rims of microphenocrysts have compositions similar to those of the most AI-depleted groundmass micas. Individual intrusions differ only with respect to the degree of Al depletion and Ti content.
Similar zonation trends (not illustrated) are found in the Bellsbank (North Blow, Southern Extension), Newlands, Roberts Victor, Star, and Saltpeterpan occurrences.
Ti-rich groundmass micas have been recognized only in highly evolved rocks from Besterskraal (this work) and Sover North (this work, Tainton 1992). These are poikilitic groundmass micas containing 4.5-9.0 wt % Ti02 (Table 2.3, Figure 2.5). Individual
100
15
14
13
12
"
7
6
5
4
3
2
CHAPTER 2
SOVER NORTH • PHENOCRYST CORES AND RIMS
\ \
~++
LAMPROITE GROUNDMASS MICAS
o Leucite Hills
1- -"I W. Kimberley I.- _...J
I L_ I ---_I
• ZONATION TREND
2 3 4 5 6 7 8 9 10
Figure 2.5. AI203 versus Ti02 compositional variation of micas from the Sover North and Besterskraal evolved orangei tes. Data for phenocryst cores and rims, and for poikilitic groundmass micas, are from different samples of the Sover North intrusion. Fields delineating compositional variation of lamproite micas from Mitchell and Bergman (199\).
MINERALOGY OF ORANGElTES
14
13
12
II
t 10
If) 9 0
'" « 8
7
6
5
4
3
POSTMASBURG
CORE-RIM I GROUNDMASS • 0
8ESTERSKRAAL~ SOVER. NORTH
2 3 4 5 6 7 8
101
9
Figure 2.6. Al203 versus Ti02 compositional variation of micas from the Postmasburg sanidine-bearing orangeite. Also shown are compositional fields for poikilitic groundmass micas from Sover North and Besterskraal.
crystals are not zoned with respect to their Ti content, although distinct intergrain compositional variation is evident.
The Sover North intrusion also contains rocks which are petrographically similar to typical un evolved orangeites, such as occur at New Elands or Sover Mine. Microphenocrystal micas in these rocks have compositions (Table 2.3) identical to those in other orangeites and are zoned toward rims which are moderately depleted in Al at constant Ti contents (Figure 2.5, Table 2.3). Rocks containing micas intermediate in composition
102 CHAPTER 2
14
13
12
i II
10
It) 9 0
t\I
« 8
7
6
5
4
3
VOORSPOED K 1/110 • KI/III +
ZONATION TREND
BESTERSKRAAL~ SOVER NORTH
2 3 4 5 6 7 8 9
Figure 2.7. AIz03 versus Ti02 compositional variation of micas from two distinct intrusions in the sanidinebearing Voorspoed orangeite.
between the Ti-rich groundmass variety and the Ti-poor microphenocrystal type have not yet been found at Sover North. However, a compositional trend of decreasing Al and increasing Ti is present in micas from the Postmasburg and Voorspoed sanidine-bearing orangeites (Table 2.4 analyses 4 and 5; Figures 2.6 and 2.7). The evolutionary trend of these micas is toward the compositions of the Ti-rich poikilitic groundmass micas from Sover North and Besterskraal, suggesting the latter micas might have originated by the extensive fractional crystallization ofVoorspoed-type precursor magmas.
14
13
12
f II
10
ae 9
~
8 ~
10
0 7
~
« 6 5 4 3 2
3
LE
VE
L 3
-4
PH
EN
OC
RY
ST
S
LE
VE
L
3 -
4
4 5
6 7
8 9
7 P
HE
NO
CR
YS
TS
SW
AR
TR
UG
GE
NS
(L
EV
EL
6
)
• P
HE
NO
CR
YS
TS
o
PH
EN
OC
RY
ST
R
IMS
I
GR
OU
ND
MA
SS
---
ZO
NA
TIO
N
TR
EN
D
7 R
IMS
I G
RO
UN
D M
AS
S
o o
10
II
12
13
14
15
16
17
18
19
20
21
2
2
23
2
4
25
--
FeO
T
wt.
%
Fig
ure
2.8.
AI 2
03
vers
us F
eOT
com
posi
tion
al v
aria
tion
of m
icas
fro
m t
he S
war
trug
gens
ora
ngei
tes.
~ ~ g ~ ~ ~ - e
104 CHAPTER 2
2.1.2.2. Ah03-FeOr Variation
Orangeite micas are strongly zoned with respect to their FeOr (total Fe expressed as FeO) content. Increasing Fe is accompanied by Al depletion which may be either very pronounced (13 to <1 wt % Ah03; Table 2.2) or relatively weak (13->8 wt % Ah03; Table 2.4).
Figures 2.8-2.11 illustrate the FeOT enrichment and extreme Al depletion trend as observed in the Swartruggens, Lace, New Elands, and Sover Mine occurrences, respectively. The cores of microphenocrystal phlogopites in these examples are of relatively uniform composition with respect to their Alz03 (10.0-13.5 wt %) and FeOT (2-6 wt %) contents. Individual occurrences differ with regard to the range in FeOr in the cores and the extent and nature of the weak zoning present.
13 +
12 •
II
r: ~
8
•
4
LACE PHENOCRYSTS
+ (03) 0
2 • (07)
GROUNDMASS o (02) o 0
0
2 3 4 5 6 7 8 9 10 II 12 13 14
FeOr Wt. % ~
Figure 2.9. Al203 versus FeOT compositional variation of micas from the Lace orangeite.
MINERAWGY OF ORANGElTES 105
t ~ 0
.: ~
'" 0 N
<i
15
14
13
12
II
10
9
8
7
6
5
4
3
2
NEW ELANDS
o AI - BIOTITE MACROCRYSTS
o Fe - POOR MACROCRYSTS
• PHENOCRYSTS o GROUNDMASS
----- ZONING TREND
o
o
o
2 3 4 5 6 7 8 9 10 II 12 13 14 15 16 17 18 19
FeOT wt. % •
Figure 2.10. A1203 versus FeOr compositional variation of micas from the New Elands orangeite.
Figure 2.12 illustrates the complexity of zonation found in microphenocryst cores in the New Elands orangeite. Within this intrusion, microphenocrysts exhibit normal and reverse zoning with respect to Fe and increasing or decreasing Al content. Micas of different zonation type are commonly juxtaposed within a single specimen. Figure 2.13 shows that similar complex zoning patterns exist with regard to Ti and Fe content.
Different intrusions in the Swartruggens dike system cannot be distinguished on the basis of the composition of microphenocryst cores (Figure 2.8). Similarly, the Finsch F4 and F7 intrusions contain microphenocrystal micas of similar composition (Figure 2.14). Groundmass poikilitic plates in Finsch F2 are rich in Fe and poor in Al relative to the phenocrysts.
Increasing intensity of red pleochroism in microphenocryst rims and groundmass plates reflects increasing Fe and decreasing Al contents. The core-rimlgroundmass compositional trends, illustrated in Figures 2.8-2.11 and 2.14, culminate in the formation of groundmass poikilitic micas containing very little Ah03 «2.0 wt %). Importantly, the extreme Fe enrichment is not accompanied by a concomitant depletion in MgO (Table 2.1), and thus the mg numbers of even the most Fe-rich micas remain greater than 0.75.
106
15
14
13
1 12
II
10
~ 0 9
.: • 8 If)
0 .J:J 7 ~
6
5
4
3
2
SOVER MINE
c AI - BIOTITE MACROCRYSTS X MICROXENOLITHS
o Fe - POOR MACROCRYSTS
• PHENOCRYSTS
o GROUNDMASS
-- ZONING TREND
CHAPTER 2
o C§ 0 o
2 3 4 5 6 7 8 9 10 II 12 13 14 15 16 17 18 19
FeOT wt. % ~
Figure 2.11. AI203 versus FeOr compositional variation of micas from the Sover Mine orangeite.
This observation is interpreted to indicate that the AI-poor micas are tetraferriphlogopites exhibiting only minor solid solution toward biotite (see 2.1.2).
Individual intrusions differ regarding the extent of Fe enrichment attained in the groundmass mica, e.g., Swartruggens (level 6, 17-25 wt % FeOT; level 7, 18-20 wt % FeOT; levels 3 and 4, 14.5-16.0wt % FeOT; see Figure 2.8), SoverMine (15.0-18.5 wt % FeOT; see Figure 2.11), and Saltpeterpan (l5.~16.0 wt % FeOT). As all of these intrusions contain microphenocrysts of similar composition, this observation is interpreted to indicate that the post-intrusion crystallization history of each occurrence was slightly different.
Figures 2.15-2.17 illustrate the extent of FeOT enrichment for the moderate-to-weak AI-depletion trend observed in the Postmasburg, Voorspoed, and Sover North occurrences, respectively. Microphenocryst cores have compositions (Table 2.4) similar to those of microphenocrysts in orangeites whose micas evolve along the extreme AI-depletion trend. Compositional trends in each intrusion differ with respect to the degree of Al depletion and Fe enrichment. A wide range of compositions is found even
13'0
1 ". If)
q. 11'5 «
M
M
I"' 0 I----N-EW--E-LA-N-O-S-""
M - Fe - POOR MACROCRYSTS • • 0
10'5 CORE RIM ZONATION TRENO
2'0 2'5 3'0 3'5 4'0 4'5 5'0 5'5 6'0
FeOT wt. % ~
Figure 2.12. Al203 versus FeOr compositional variation of macrocrystal and microphenocrystal micas from the New Elands orangeites.
3 NEW ELANDS • CORE TO MANTLE
t ZONATION TREND • •
• ~ • ...: ~ 2 N • 0
l- • •
3 4 5 6 7 8 9
FeOr Wt. % •
Figure 2.13. Ti02 versus FeOr compositional variation of microphenocrystai micas from the New Elands orangeites.
107
188 CHAPTER 2
ca' FINSCH
r 12
10
~ 8 + -~
If) 6 0
~ « (;) F2 • F4 +
4 + F4E • F4ID 0 F7
2 ~ CORE TO RIM ZONATION TREND
2 4 6 8 10 12 14 16
FeOT wt. %
Figure 2.14. A1203 versus FeOr compositional variation of micas from the Hnsch orangeites. F-numbers refer to petrographically distinct intrusions.
r 14
13
12 ~ 0
.; II ~
If) 10 0 ~ « 9
8
3 4 5 6
POSTMASBURG
CORE TO RIM ZONATION K37 MICROPHENOCRYSTS
GROUNDMASS K36
7 8 9 10 II 12 13 14 15 16 17 18 19
FeOT Wt. % ~
Figure 2.1S. Al203 versus FeOr compositional variation of micas from the Postmasburg sanidine-bearing orangeites. K37 and K36 are different intrusions in the Postmasburg area.
MINERAWGY OF ORANGElTES
t 13
12
ae ...: II ~ fC)
010 ~ «
9
8
3 4 5 6 7 8 9
VOORSPOED MICROPHENOCRYSTS
CORE TO RIM ZONATION
10 II 12 13
FeOT wt. % •
109
Figure 2.16. Ah03 versus FeOr compositional variation of micas from the Voorspoed sanidine-bearing orangeites.
within a single specimen, e.g., Sover North (Figure 2.17). The extent of Fe enrichment is similar to that found in the AI-poor tetraferriphlogopites: however, this iron enrichment is accompanied by magnesium depletion (Table 2.4), representing solid solution toward biotite and tetraferriphlogopite (see 2.1.8)
Figure 2.17 shows that groundmass poikilitic plates in the Sover North evolved rocks are moderately AI depleted, but not strongly enriched, in Fe (Table 2.3). Individual crystals are typically not zoned, although inter-grain compositional variation is evident. Similar compositions have been reported by Tainton (1992). Groundmass micas of these compositions (Table 2.3, Figure 2.17) also occur in the Besterskraal orangeite.
Figure 2.17 also shows that the rims of some microphenocrysts in the Sover North microporphyritic rocks are of identical composition, with respect to Fe, but not Ti (Figure 2.5), to the poikilitic groundmass micas in the evolved rocks.
2.1.2.3. Macrocrysts Versus Microphenocrysts
Weakly pleochroic phlogopite micas, classified as macrocrysts on a textural basis, cannot in most instances be distinguished from typical microphenocrysts on the basis of their AI, Fe, and n contents (Figure 2.18). Macrocrysts may be zonation- and/or mantlefree, or mantled by mica having greater (Figure 2.12) or lesser FeOr than the cores.
Macrocrysts may represent either relatively large primary/cognate micas or xenocrysts derived by the fragmentation of upper mantle-derived xenoliths. Unfortunately, the compositions of many mantle-derived (Delaney et al. 1980, Dawson and Smith 1977) and microphenocrystal micas are similar, and it is not possible to distinguish unambiguously the origin of the macrocrysts on the basis of their major or minor element compositions.
14
[22l PHENOCRYSTS LEUCITE HILLS
13
12
t II
10
* 9
..: ~ 8 If)
0 N 7
<i 6
5
4
3 SOVER NORTH . PHENOCRYST CORES AND RIMS 2
+ GROUNDMASS
BESTERSKRAAL
2 3 4 5 6 7 8 9 10 II 12 13 14 15 16
Figure 2.17. AI203 versus FeOT compositional variation of micas from the Sover North and Besterskraal orangeites. Data for phenocrystal and poikilitic groundmass micas from Sover North are from different samples. Fields delineating lamproite mica compositional variation from Mitchell and Bergman (199\).
Figure 2.18. Ah03 versus Ti02 and FeOT compositional variation of macrocrysts and microphenocrystal micas from the Swartruggens orangeites.
110
MINERALOGY OF ORANGEITES 111
Relatively Fe-poor «3.0 wt % FeOT), Cr-rich (0.5-1.2 wt % Cr203) macrocrysts, which may represent xenocrysts of lherzolite-derived primary mica, appear to be absent or very rare in orangeites. When present they comprise much less than 1 % of the mica population. Such micas are present in the New Elands (Figure 2.10) and Sover Mine (Figure 2.11) mica populations.
Unfortunately, it is not possible to distinguish between relatively Fe-rich mantlederived xenocrysts and primary microphenocrysts on the basis oftheircomposition. This observation suggests that some of the macrocryst population may actually be xenocrysts. The problem of identification of xenocrystal and cognate mica is the same as that encountered in the characterization of the olivine population (see 2.3.2).
2.1.2.4. Minor Elements Of the minor elements occurring in orangeite micas only Cr, Ba, F, and rarely Na,
are present in significant amounts. Ni contents are uniformly low, being typically less than 0.2 wt % NiO, and commonly less than 0.1 wt % NiO.
The Cr content of orangeite mica varies widely. Representative Cr contents are given in Table 2.5. Cr203 contents greater than 2.0 wt % have not been found. Typically, Cr contents decrease with increasing Fe content and groundmass tetraferriphlogopites contain very low levels of Cr.
Compositional variations within macrocryst and microphenocryst populations are complex and no simple trends are evident. Macrocrysts may be Cr rich (>1.0 Cr203) or Cr poor «0.5 wt % Cr203). Commonly, the cores ofCr-poor macrocrysts are mantled by Cr-rich (1.0-1.5 wt % Cr203) mica. Typically, adjacent microphenocrysts differ widely with respect to their Cr contents. Figures 2.19-2.21 illustrate Cr-Ti variation in coexisting microphenocrysts from the Lace orangeite (this work) and the New Elands and StarBurns dikes (Mitchell and Meyer 1989a). These figures show that Cr may decrease or increase as Ti decreases or increases.
Microphenocrystal micas characteristically contain less than 1.0 wt % BaO and commonly have less than 0.5 wt % BaO (Table 2.5). Groundmass micas exhibit a wide range in Ba content and are typically enriched in Ba relative to microphenocrysts. Ba contents do not exceed 2.0 wt % BaO, and only groundmass micas from Besterskraal have consistently high inter-grain BaO contents. High Ba contents are not associated with increased Al contents, demonstrating the absence of solid solution toward kinoshitalite (see 2.1.9).
The majority of microphenocrystal and groundmass micas contain less than 0.5 wt % Na20 (Table 2.5). Relatively high Na contents (Table 2.5) are found only in the poikilitic groundmass micas from Sover North and Besterskraal. Na enrichment is positively correlated with Ba enrichment in these occurrences only. In contrast, groundmass micas from Lace have high Ba contents but are poor in Na (Table 2.5).
Little is known about the F content of orangeite micas. Data obtained in this study show that orangeite micas typically contain less than 1.0 wt % F. In the Swartruggens dikes, microphenocrysts contain from 0.61 to 0.14 wt % F and are zoned toward decrea<;ing F contents from core to margin. Groundmass tetraferriphlogopites contain 0.38-0.12 wt % F. Diverse samples of Lace orangeite show no consistent zoning trend and individual coexisting microphenocrysts may be zoned toward increasing or decreas-
112 CHAPTER 2
Table 2.5. Cr203, BaO and Na20 Contents (wt %) of Microphenocrystal and Groundrnass
Figure 2.20. Cf203 versus Ti02 compositional variation of micas from the New Elands orangeite (after Mitchell and Meyer 1989a).
113
114 CHAPTER 2
STAR DIKES 1·5 • BURNS
t •• WYNANDSFONTEIN
0 NEW STAR ..
~ ..: 1·0 ~
rt)
0 (\J ~
U
0·5
0·5 1·0 1·5 2·0 2·5
Ti02 wt. % •
Figure 2.21. CT203 versus Ti02 compositional variation of micas from the Star dike system.
ing F from core to margin. F contents range from 0.5 to 1.09 wt %. Some groundmass micas contain no detectable F, while others exhibit an inter-grain range of 1.5 to 0.13 wt % F. Microphenocrystal micas in the Sover Mine orangeite contain from 0.53 to 0.14 wt % F and are zoned toward decreasing F from core to margin. Groundmass tetraferriphlogopites contain from 0.45 to 0.11 wt % F. At Bellsbank Southern Extension, macrocrysts are slightly richer in F (0.6-0.3 wt %) than microphenocrysts (0.5-0.20 wt % F). However, groundmass tetraferriphlogopites encompass this range and contain 0.6-0.2 wt % F. Data for Besterskraal and Sover North micas are not available.
2.1.2.5. Trace Elements
The trace element content of orangeite micas has not been studied in detail. Smith et al. (1979) reported that microphenocrystal micas contain 90-820 ppm Rb20 and 470-5800 ppm BaO. KlRb ratios range from 112 to 1039, with the majority being
MINERALOGY OF ORANGEITES 115
100-300 (average 251). K/Ba ratios range from 16 to 199 and average 74. Allsopp and Barrett (1975) report wide variations in the Rb (202-952 ppm) and Sr (55-281 ppm) content of microphenocrystal micas from Roberts Victor and Swartruggens. K/Rb ratios range from 100 to 160.
2.1.3. Aluminous Mica-Microxenoliths
Micas which are enriched in AI, relative to the majority of most macrocrystal and microphenocrystal micas, occur as microxenoliths at Swartruggens. These micas contain 17.5-20.0 wt % Ah03 and are aluminous phlogopites exhibiting solid solution toward "eastonite" (Figure 2.1). Each microxenolith contains micas which are distinct with respect to their Ti and Fe contents (Table 2.6, Figure 2.22). The micas exhibit wide ranges in Cr203 (0.21-3.3 wt %) and Ti02 (0.1-1.2 wt %) contents within and between xenoliths. Single macrocrysts of these compositions, which might be derived by the disaggregation of microxenoliths, have not yet been found at Swartruggens.
Table 2.6. Representative Compositions of Aluminous Phlogopitea
QFeOr = total Fe expressed as FeO; n.a. = not analyzed. 1-5. aluminous phlogopites from microxenoliths. Swartruggens (this work); 6, secondary phlogopite from harzburgite (Delaney et al. 1980); 7-8, aluminous phlogopites from microxenoliths, Leucile Hills (Barton and van Bergen 1981, Mitchell and Bergman 1991, respectively).
116 CHAPTER 2
20
... MICROXENOLlTHS ...
t (j)
18 @ AI-BIOTITE
MACROCRYSTS
~ J
...: 16 ~
~ SWARTRUGGENS 0 ~ <[
14
12 MICROPHENOCRYSTS
2 4 5 6 7 8 9 10 II
-- FeOT wt. % • Figure 2.2l. Compositional variation of aluminous phlogopites from three microxenoliths in the Swartruggens orangeites.
The relationship of the microxenoliths to their host is unclear, but they are considered to be more likely genetically related to orangeite than being xenoliths of mantle material. This hypothesis is based upon two observations:
Primary phlogopites in mantle-derived Iherzolites are relatively poor in Ah03 (12.4-14.5 wt %; Delaney et al. 1980). Most secondary micas in lherzolites and harzburgites also have relatively low Al contents, with the exception of a single mica from a harzburgite which is similar in composition to some of the Swartruggens Fe-poor microxenolithic micas (Table 2.6, anal. 6). Micas from MARID xenoliths contain only 8-13 wt % Ah03 (Dawson and Smith 1977). Thus, the low Al content of most mantlederived micas implies that derivation of the microxenoliths from such a source is unlikely.
Phlogopites of similar paragenesis and composition (Table 2.6, Figure 2.22) have been described by Barton and van Bergen (1981) and Mitchell and Bergman (1991) from the Hallock Butte, South Table Mountain, and Hatcher Mesa lamproites of the Leucite Hills. These micas form a compositional continuum with phenocrystal micas oflower Al content, and are thus interpreted to be high-pressure phenocrysts.
MINERALOGY OF ORANGEITES 117
The wide range in Fe contents (Table 2.6) of the microxenolith micas suggests that orangeite magmas might undergo considerable high-pressure differentiation if the microxenoliths found at Swartruggens have a similar origin to that proposed for the Leucite Hills mica-rich xenoliths.
2.1.4. Aluminous Biotite Macrocrysts
Representative compositions of AI- and Fe-rich mica macrocrysts are given in Table 2.7. Figure 2.1 shows that the micas are aluminous biotites whose compositions may be regarded as solid solutions between phlogopite and siderophyllite. Although individual macrocrysts are of uniform composition there is considerable intergrain compositional variation (Figures 2.1, 2.3, 2.10, 2.11). Macrocrysts are markedly enriched in Fe and Al (Figures 2.1, 2.10, 2.11) and slightly in Ti (Figure 2.3) relative to microphenocrysts. In many instances the macrocrysts are mantled by pale-colored mica identical in composi-
Table 2.7. Representative Compositions of Aluminous Biotite Xenocrystsa
aFeOr = total Fe calculated as FeO; n.a. = not analyzed; n.d. = not determined. 1-2, Swartruggens; 3-4, Sover Mine; 5, Bellsbank; 6, New Elands (l..{') this work); 7, Saltpeterpan; 8, Monteleo (Smith et al. 1978); 9, Postmasburg K35; 10, Newlands (Tainton 1992).
118 CHAPTER 2
tion to microphenocrystal phlogopite (Figures 2.3, 2.1 0, 2.11; this work, Smith et al. 1978) and rarely by tetraferriphlogopite (Figure 2.10).
Macrocrysts exhibit wide intergrain compositional variation with respect to Ti and Ba contents, e.g., New Elands (1.71-4.28 wt % Ti02, 0.4-0.6 wt % BaO), Swartruggens (0.44-3.1 wt % Ti02, 0.05-0.44 wt % BaO), Sover Mine (3.1-3.90 wt % Ti02, 0.07-0.38 wt % BaO). Cr203 and Na20 contents are typically below 0.1 and 0.3 wt %, respectively. The few data available (this work) suggest that the macrocrysts are F poor, e.g., Sover Mine, <0.1-0.8 wt % F (11 analyses); Swartruggens, 0.1-0.3 wt % F (4 analyses).
Little is known of the trace element content of these micas. Smith et al. (1979) have determined that they contain 240-{)30 ppm Rb20 (average of 5 analyses = 470 ppm) and 1030-4320 ppm BaO (average of 5 analyses = 2670 ppm). KlRb ratios range from 137 to 348 (average 209) and KlBa from 20 to 88 (average 42).
The origins of the AI-biotite macrocrysts are unknown. Smith et al. (1978), while recognizing that the micas may be simply xenocrysts, suggest derivation from the "intrusion of another magma just prior to intrusion of the kimberlite." This magma was suggested to be carbonatitic. However, Smith et al. (1978) did not present any evidence to support this hypothesis, other than the observation that some micas in carbonatites are of similar composition to the aluminous biotites.
Mitchell and Meyer (1989a), emphasizing the constant association of the macrocrysts with orangeites, and the absence of polymineralic mantle-derived microxenoliths containing mica of similar composition, suggested that the micas are unlikely to be simply xenocrysts. The Al-biotites found in the Swartruggens Male lamprophyre (2.1.5) are Al poor relative to the macrocrysts. Hence, derivation from such a lamprophyric source is considered unlikely.
The absence of similar macrocrysts in archetypal kimberlites emplaced in the Kaapvaal craton (and elsewhere) suggests the Al-biotites have a genetic affinity with orangeite magmas or relationship to their source regions.
Determining the origin of the macrocrysts is analogous to the problem of the origin of green Fe-rich pyroxenes in other alkaline magmas (see reviews by Duda and Schminke 1985, Bedard et al. 1988, Dobosi and Fodor 1992). Despite intensive investigation the origins of these pyroxenes remain ambiguous. They are commonly interpreted, largely on subjective criteria, to be high-pressure phenocrysts derived from differentiated batches of magma.
2.1.5. Micas from the Swartmggens Male Lamprophyre The cores of microphenocrystal micas in the Swartruggens Male lamprophyric dike
have similar Ah03 contents to microphenocryst cores in the Swartruggens and other orangeites. The micas differ in being relatively richer in Ti02 and zoned primarily with respect to FeOT content (Table 2.8). The continuous zonation trend is one of increasing FeOT, decreasing MgO, and slightly decreasing Ah03 at essentially constant Ti02 content (Figures 2.23, 2.24). The compositional trend is similar to that found for the Postmasburg and Sover North micas (Figure 2.24) and represents an evolutionary trend from phlogopite toward AI-biotite. The absence of tetraferriphlogopite in these rocks suggests that
MINERAWGY OF ORANGEITES 119
Table 2.8. Representative Compositions of Micas in the Swartruggens Male LamprophyreQ
95.00 94.86 94.76 94.38 94.82 94.63 Structural formulae based on 22 oxygens.
Si 5.654 5.593 5.616 5.612 5.575 5.634 Al 2.061 2.104 2.113 1.955 2.152 1.798 Ti 0.406 0.449 0.410 0.489 0.447 0.484 Cr 0.002 0.004 Fe 1.036 1.397 0.967 2.149 0.891 2.605 Mn 0.005 0.011 0.010 0.033 0.006 0.037 Mg 4.753 4.332 4.799 3.661 4.815 3.411 Na 0.031 0.041 0.D28 0.051 0.023 0.052 K 1.864 1.882 1.874 1.849 1.877 1.826 Ba 0.406 0.028 0.022 0.040 0.030 0.042 Ni
mg 0.821 0.756 0.832 0.630 0.844 0.567
aFeOr = total Fe expressed as FeO; n.d. = not detected; C = core; R = rim.
they have no simple genetic relationship to the geographically associated Swartruggens orangeites.
The micas are very poor in Cr203 «0.1 wt %) and NiO «0.05 wt %). BaD ranges from 0.25 to 1.1 wt %, and F from 0.7 to 1.6 wt %. No consistent zonation trend is evident, and individual crystals may exhibit trends of increasing or decreasing Ba and F from core to margin. With respect to the minor elements only the low Cr contents of these micas distinguishes them from typical orangeite microphenocrystal micas.
2.1.6. Summary of Mica Compositional Variation
Microphenocrystal phlogopites exhibit a limited range in composition with respect to their AI, Fe, and Ti contents. Intra- and inter-intrusion composition differences are not significant. These observations are interpreted to suggest that all of the microphenocrystal micas are derived from compositionally similar parental magmas.
120 CHAPTER 2
15
1 SWARTRUGGENS
14 MALE LAMPROPHYRE
~ 13 • .. ..: ~
rt) 12 a ~ I&J « l-
II I&J • (!) z ~ a:
10 0
9
2 3 4 5 6
Ti02 wt. %
Figure 2.23. Al203 versus Ti02 compositional variation of microphenocrystal micas from the Swartruggens Male lamprophyre dike.
The coexistence of microphenocrysts that are normally or reversely zoned with respect to Ti, Cr, and Fe, demonstrates that the microphenocrystal mica assemblage could not have crystallized in situ. The complex zoning and mantling patterns documented in this work and by Mitchell and Meyer (1989a) can be best explained by assuming that the majority of the microphenocrystal micas crystallized from several, slightly compositionally different, batches of parent magma. Microphenocrysts (and some macrocrysts) crystallizing from each batch of magma were subsequently mixed, extracted from a common magma chamber, and emplaced as a heterogeneous hybrid assemblage.
Although crystallization in a magma chamber is invoked here to explain the complex zonation observed in the micas, it should be realized that orangeites do not appear to form large long-lived continuously replenished fractionating magma chambers of the type associated with some minettes or alkaline basaltic magmas.
An important consequence of the above hypothesis is that the microphenocrysts must be a transported assemblage and their present high modal concentrations are the result of
MINERALOGY OF ORANGEITES 121
14
13
12 0
1 II .... -10
9
~ 8
..: • If)
7
0 ~ 6 «
5
4
3
SWARTRUGGENS 2 MALE LAMPROPHYRE
CORE TO RIM . • 0 ZONATION TRENDS
3 4 5 6 7 8 9 10 II 12 13 14 15 16 17 18 19 20
FeOT wt. % .. Figure 2.24. AI203 versus FeOT compositional variation of microphenocrystal micas from the Swartruggens Male lamprophyre dike.
differentiation processes. It follows from this conclusion that the whole rock major element compositions of orangeites are far removed from those of their parental magmas (Mitchell and Meyer 1989a; see 3.2).
Macrocrysts of similar composition to microphenocrysts are merely larger crystals derived from the same sources as the microphenocrysts. Rare AI-rich phlogopite microxenoliths may represent high-pressure cumulates. Fe-poor mantle-derived mica xenocrysts are relatively rare. Aluminous biotite macrocrysts are cryptogenic, but the constant association with orangeite suggests a cognate relationship may be possible.
Subsequent to mixing and emplacement, each batch of crystal-laden magma underwent crystallization in a particular pressure-temperature-oxygen fugacity regime. Differences in these intensive parameters resulted in different compositional trends developing in the groundmass micas. These and micas occurring as mantles upon microphenocrysts, macrocrysts, and xenocrysts are considered to have crystallized in situ. Two principal trends are evident:
122 CHAPTER 2
1. A tetraJerriphlogopite trend, characterized by extreme Al depletion coupled with Fe enrichment at relatively constant Mg contents. Ti may decrease slightly or remain constant.
2. A biotite trend, characterized by Fe enrichment and Mg depletion accompanied by moderate Al depletion. Ti may decrease or increase slightly.
Groundmass micas following the biotite trend may be more common in occurrences which differentiate to felsic residua, i.e., Postmasburg, Voorspoed. However, the association is not universal, as micas from Makganyene follow the tetraferriphlogopite trend and those from Sover North follow a range of compositional trends between the two extremes.
The origins of the trends are undoubtedly related to the oxygen fugacity of the magma. The tetraferriphlogopite trend may represent mica evolution under relatively oxidizing conditions with the bulk of the iron being present as Fe3+. The biotite trend may represent crystallization under relatively reducing conditions, with iron being distributed between Fe2+ and Fe3+. However, other factors, such as the peralkalinity of the magma, playa role in determining the oxidation ratio of minerals crystallizing from the magma. Unfortunately, the relative role of each factor cannot as yet be assessed.
Differences in oxygen fugacity may be caused by extensive crystallization, contamination, and/or interaction of orangeite magmas with ground waters. Unfortunately, discussion of the origins of fugacity differences must necessarily remain speculative, as the nature of orangeite magmas and the physical conditions prevailing during their crystallization are unknown.
2.1.7. Solid Solutions in Orangeite Mica
The nature of the solid solutions present in orangeite micas is difficult to assess as most compositions have been determined by electron microprobe analysis. Water and halogens are typically not determined; consequently, structural formulae are calculated on the basis of 22 oxygens/formula unit. The actual Fe3+lFe2+ ratios of the micas are unknown, but in some instances may be approximately evaluated by recalculation of their compositions on a stoichiometric basis. However, calculation of ferric and ferrous iron contents by the methods suggested by Droop (1987) is not always successful, as the micas are commonly partially altered and/or contain less than 14 tetrahedral and octahedral cations/formula unit. In many instances unrealistically high ferric iron contents result from this recalculation procedure.
The majority of primary phenocrystal and groundmass micas contain insufficient Si and Al to occupy all of the available tetrahedral sites. Consequently, Alvi is absent and, within the overall compositional variation Si increases as Aliv decreases. Si commonly exceeds the ideal value of six atoms of Si/formula unit (Figure 2.25) in the most evolved micas. In common with lamproite micas, which exhibit similar compositional trends (see 2.1.9), it is suggested that the Al deficiency is. undoubtedly a direct reflection of the peralkalinity of the parent magma. The tetrahedral site deficiency increases with differentiation (Table 2.9) and is principally remedied by the entry of Fe3+ to this site. For many micas there is insufficient Fe3+, Ti, and Cr present to fill the tetrahedral sites (Table 2.9). In such instances the remaining deficiency may be eliminated by the entry ofMg into the tetrahedral site. The existence of Miv has been proposed by Robert (1981) on the basis
MINERALOGY OF ORANGElTES
2'0
1'0
0.5 + SWARTRUGGENS - FINSCH - LACE - BELLS BANK
• SOVER NORTH
o BESTERSKRAAL
• VOORSPOED - + + + POSTMASBURG - MAKGANYENE
5'5 6'0
Si •
123
6'5
Figure 2.25. Aliv versus Si (atomsl22 oxygens) for orangeite micas. Field of lamproite micas from Mitchell and Bergman (1991).
of the synthesis of MgiV-mica by Tateyama et al. (1974). The presence of excess octahedral-site cations in the recalculated structural formulae (Table 2.9) would seem to support this hypothesis.
Figure 2.26 illustrates the variation of Al with MgI(Mg + Fer) in orangeite micas and shows clearly that micas evolve from phlogopite-rich phenocrysts to eithertetraferriphlogopites (trend 1) or AI-deficient biotites (trend 2).
124 CHAPTER 2
Table 2.9. Representative Compositions of Swartruggens Groundmass Micaso
·Pt!203 and PeO calculated by the method of Droop (1987) on the basis of 14 tetrahedral (T) and octahedral (0) cationsl22 oxygens. CAT = cation total.
Figures 2.1 and 2.26, and the above discussion demonstrate that the majority of orangeite micas are members of a solid solution series between phlogopite [K2Mg6Si6Ah02o(OH)4], annite [K2F~+Si602o(OHh] and tetraferriphlogopite [K2Mg6-Si6Fe~+02o(OH)4]. Trend 1 is primarily a solid solution between phlogopite and tetraferriphlogopite, while trend 2 involves a three component solid solution between phlogopite, annite and tetraferriphlogopite.
Some phenocrysts and macrocrysts are slightly enriched in Al and contain more than two atoms AI/formula unit (Figure 2.26). Such micas may be considered as exhibiting limited solid solution toward a hypothetical "eastonite" molecule [K2MgsA1SisAh02o(OH)4]. (Livi and Veblin 1987 have shown that "eastonite" micas are actually a mixture of serpentine and phlogopite.) Significant amounts of this molecule
Figure 1.26. AI (atoms/22 oxygens) versus Mg/(Mg + FeT) for orangeite micas. FeT = total Fe expressed as Fe2+. PHL = phlogopite. See text for explanation of evolutionary trends I and 2.
are present in aluminous micas in microxenoliths (Table 2.6). These micas have no tetrahedral site deficiency and contain Alvi.
The Ti-rich micas from Besterskraal and Sover North plot on the tetraferriphlogopite trends in Figures 2.25 and 2.26, suggesting that they have significant Fe3+ contents. However, their ferric and ferrous iron contents cannot be evaluated as these micas contain less than 14 tetrahedral and octahedral cations/formula unit. This cation deficiency is probably due to the presence of vacancies in the octahedral sites introduced by the
126 CHAPTER 2
presence of Tivi. Hence, these micas are best regarded as solid solutions between phlogopite, tetraferriphlogopite, and Ti-octahedral site-deficient phlogopite [K2M~oriSi602o(OH)4] .
2.1.S. Mica in Kimberlites
The compositional characteristics of primary micas in archetypal kimberlites are insufficiently known because previous studies typically did not distinguish between these rocks and orangeites. Hence, the phlogopite-tetraferriphlogopite evolutionary trend, shown in Section 2.1.7 to be characteristic of orangeites, was considered to be typical of kimberlites in general (Mitchell 1986). This work demonstrates that this conclusion is incorrect.
16 0 o PHENOCRYSTS - WESSEL TON KIMBERLITE
"SECONDARY - TEXTURED MICA- XENOLITHS 15 0
i 0
+ 0
14
~ .,; 13 ~
If)
0 ~ 12 «
II
10
2 3 4
Ti02 wt. %
Figure 2.27. Ah03 versus Ti02 compositional variation of macrocrystal and megacrystal micas from orangeites (this work) and kimberlites (Shee 1985. Dawson and Smith 1975. this work) compared with that of primary and secondary textured micas in mantle-derived xenoliths (Dawson and Smith 1975. 1977. Carswell 1975. Delaney et al. 1980).
MINERALOGY OF ORANGEITES 127
2.1.8.1. Macrocrysts
Many micas in archetypal kimberlites are cryptogenic macrocrysts or megacrysts whose composition is relatively well-characterized (Mitchell 1986, Farmer and Boettcher 1981, Dawson 1980, Dawson and Smith 1975). These micas are aluminous phlogopites (10-15 wt % Ah03) which exhibit a wide range in Ti02 (0.5-5.5 wt %) and Cr203 (0.1-2.0 wt %). Individual micas are homogeneous, although there is wide intergrain compositional variation.
Dawson and Smith (1975) proposed that the origins of megacrystal micas in South African kimberlites may be deduced on the basis of their Cr, Ti, and Fe contents. Megacrysts relatively rich in Cr203 (>0.5 wt %) and poor in Ti02 «0.6 wt %) and FeO «3.7 wt %) are considered to be xenocrysts derived from phlogopite-bearing lherzolites. Megacrysts relatively poor in Cr203 «0.5 wt %) and rich in Ti02 (0.6-2.0 wt %) and FeO (>3.75 wt %) are interpreted to be phenocrysts. However, note that such empirical studies might not universally valid, as the compositions of micas in the upper mantle in other regions might be very different from those of the control group of primary and secondary micas, initially used by Dawson and Smith (1975) to establish the composition of mantle-derived xenocrystal mica.
16
t ~ 14
..: ~
10 0 ~ 12 <{
10
o PHENOCRYSTSWESSELTON KIMBERLITE
2 4
MEGA I MACROCRYSTS • ORANGEITE
o + SOMERSET ISLAND KIMBERLITES
• WESSELTON KIMBERLITE
6
FeOr Wt. % •
8
Figure 2.28. Ah03 versus FeOT compositional variation of macrocrystal and megacrystal micas from orangeites and kimberlites compared with that of primary and secondary textured micas in mantle-derived xenoliths. Data sources as in Figure 2.27.
128 CHAPfER2
MEGA! MACROCRYSTS
t o SOMERSET ISLAND
2'0 ~ KIMBERLITES
0 SECONDARY - TEXTURED
1'5 0 MICA - XENOLITHS 0
~ 0 0 <gc9
..: ~
0 0
If) 0
0 1'0 0 N 0 ~ 0 0 0
0 0
0
OO~ 0'5 o 0
0
2 3 4 5
Ti02 wt. %
Figure 2.29. Cn03 versus n02 compositional variation of macrocrystal and megacrystal micas from orangeites and kimberlites compared with that of primary and secondary textured micas in mantle-derived xenoliths. Data sources as in Figure 2.27.
Recent studies (Mitchell 1986) of the mica macrocryst population in Somerset Island kimberlites have shown that. within a given kimberlite province. individual kimberlites are characterized by compositionally distinctive macrocryst suites. This observation suggests that many of the macrocrysts may indeed by cognate.
Macrocrystal micas in kimberlites are very similar in major (Figures 2.27 and 2.28) and minor (Figure 2.29) element composition to the least-evolved macrocrystlmicrophenocrysts in orangeites. Thus macrocrystal mica compositions are considered to be oflittle use in discriminating between kimberlites and orangeites.
2.1.8.2. Primary Micas
Few studies of the compositional variation of bonafide primary micas in kimberlites have been undertaken and a characteristic evolutionary trend has not previously been recognized. This work summarizes existing data and incorporates new compositional data (approximately 350 analyses) for groundmass primary micas from phlogopite-bearing archetypal kimberlites from Somerset Island. Colorado-Wyoming. Siberia, China, and
MINERALOGY OF ORANGElTES 129
Guinea. Because of the diversity of mica populations, individual occurrences are de-scribed separately prior to summarizing the compositional trends.
2.1.B.2.a. Somerset Island Kimberlites. The Tunraq kimberlite has previously been termed a micaceous kimberlite by Mitchell (1979). It contains large (1-5 mm) flow-aligned, deformed phlogopite macrocrysts (Ah03 = 12.0-13.8 wt %, Ti02 = 1.6-5.5 wt %, FeOT = 4.4-6.2 wt %, mg = 0.87-0.92) but lacks groundmass micas. No compo-sitional differences exist between micas in the different facies of this kimberlite. The micas are unevolved, with respect to their Fe and Al contents, and are interpreted (this work) to represent phenocrysts which have crystallized at depth. These have been concentrated prior to intrusion.
In addition to rare phlogopite macrocrysts, the Elwin Bay kimberlite (Mitchell 1978a) contains common late-stage groundmass colorless micas (this work). These are not distributed uniformly throughout the groundmass but occur, in association with calcite, as discrete patches or segregations of interlocking small prisms and plates. Micas
Table 2.10. Representative Compositions of Groundmass Micas from the Elwin Bay and Jos Kimberlitesa
aFeOr = total Fe expressed as FeO; n.d. = not detected. CaO and NiO not detectable by electron microprobe. Compositions 1-5 Elwin Bay, 6-10 Jos. (All data this work.)
130 CHAPTER 2
within each segregation consist of B a-free, AI-rich cores which are strongly zoned toward Ba- and AI-rich margins. The intra- and intergrain zonation is irregular, and the texture is interpreted to suggest reaction between previously formed Ba-free groundmass micas and late-stage Ba-rich fluids. The micas contain <0.1-13.4 wt % BaO, 5.8-1004 wt % K20, <0.05-1.2 wt % Ti02, and 2-5 wt % FeOT (Table 2.10). The compositional variation is shown in Figures 2.30 and 2.31. BaO and Ah03 (13.3-19.7 wt %) contents vary inversely with K20 and Si02 (42.1-29.7 wt %) indicating considerable solid solution from phlogopite toward the brittle mica kinoshitalite [BaMg3AhShOlO(OHh]. The coupled substitution (K+,Si4+ = Ba2+,AI3+) is illustrated in Figure 2.32 by the 1: 1 correspondence between K and Ba. The micas are best regarded as barian phlogopites, as the number of Ba atomsl11 oxygens does not exceed 0.5.
21 0
. ELWIN BAY 20 0 JOS . ' . .. . .
I ZONATION TREND
19
18
ae - 17 ~
10 16 0
(\J
« 15
. LEUCITE . HILLS 14 PRIMITIVE
MICAS r-\---
13 I
2 3 4 5
Ti02 wt. % •
Figure 2.30. A1203 versus Ti02 compositional variation of groundmass micas in the Elwin Bay and Jos kimberlites. Somerset Island (this work). Field of Leucite Hills lamproite primitive micas from Mitchell and Bergman (1991).
FeOT wt. % • Figure 2.31. AI203 versus FeOr compositional variation of groundmass micas in the Elwin Bay and Jos kimberlites. Somerset Island (this work). Field of Leucite Hills lamproite primitive micas from Mitchell and Bergman (1991).
Ba-free or Ba-poor groundmass micas are rich in Ah03 (> 13 wt %) and poor in Cr203 «0.05) and Ti02 relative to macrocrystal micas «13 wt % Ah03, >1 wt % Ti02, >0.5 wt % Cr203, Figures 2.27-2.29).
Three varieties of mica occur in the Jos kimberlite (Mitchell and Meyer 1980). Type A micas are yellow-brown, low-Ti02 (<2.0 wt %), low-Cr203 «0.2 wt %) phlogopites [Mg/(Mg + Fer) = 0.93-0.87]. Type B phlogopites [Mg/(Mg + Fer) = 0.93-0.84] are darker in color and richer in Ti02 (2.5-5.6 wt %). Significant intergrain compositional variation is found with respect to Cr and Ti contents. There is a compositional gap between type A and B micas with respect to Ti02.
1/1 0'4 E o -c - 0'2 • JOS
+ ELWIN BAY
0'2 0·4 0 ·6 0·8 1'0
Ba (atoms / II oxygens ) ~
Figure 2.32. Ba versus K (atomic) for groundmass micas from the Elwin Bay and los kimberIites. Somerset Island.
Figure 2.33. Ba-rich micas belonging to the phlogopite-kinoshitalite serie8 occurring as discrete mantles on earlier-formed groundmass Ba-poor phlogopites. Jos kimberlite. Somerset Island. Backscattered electron image (500x).
132
MINERALOGY OF ORANGEITES 133
Type C micas are colorless, rich in Ah03 (15.0-21.8 wt %), and poor in Ti02 «2 wt %), Cr203 «0.5 wt %), and FeOT «4.5 wt %) relative to type A and type B micas. Type C micas exhibit continuous compositional zonation with respect to Ba (Table 2.10, Figure 2.32) and range from Ba-bearing «0.1 Ba/afu) phlogopite through barian phlogopite (0.1-0.5 Ba/afu) to potassian kinoshitalite (>0.5 Ba/afu). Ba enrichment exceeds that found in Elwin Bay micas and is in accord with the more-evolved character of the Jos calcite kimberlite relative to the Elwin Bay monticellite kimberlite. Increasing Ba contents are coupled with increasing Al and decreasing Si, Fe, and Ti contents (Figures 2.30 and 2.31). Hence, Mg/(Mg+FeT) ratios increase as the compositions evolve toward kinoshitalite. Previously, type C micas were regarded, incorrectly, as chloritized "eastonitic" micas (Mitchell and Meyer 1980, Mitchell 1986).
Type A micas may be mantled by type B or C micas. Type B micas have discrete type C mantles. Type C micas also occur as poikilitic groundmass plates and small prismatic crystals. Type C micas, replacing earlier formed micas, may occur as diffuse irregular patches at crystal margins and along cleavages and fractures or as discrete epitaxial mantles (Figure 2.33). Several growth periods of Ba-poor and Ba-rich mica have
Table 2.11. Representative Compositions of Groundmass Micas from the Iron Mountain Kimberlitesa
aFeOr = total Fe expressed as FeO; n.d. = not detected; Cao and NiO not detectable by electron microprobe. Compositions 1-3 Iron Mountain #7, 4-10 Iron Mountain #24 (All data this work.)
134 CHAPrER2
Figure 2.34. Groundmass laths of phlogopite in Iron Mountain (Wyoming) kimberlite showing diffuse patchy replacement by Ba-rich mica. Backscattered electron image (800x).
t 1·0
(/) 0'8 C Q) 01 >. )(
0 '6 0
....... (/) 0'4 E 0 +-0
0 ·2 ~
" '.~ .. ~1-'"
"' , " " " "
• IRON MOUNTAIN
m CHICKEN PARK
0 '2 0'4
" " " " "
0·6
" "
0 '8
" " " 1·0
Ba ( atoms / II oxygens) •
Figure 2.35. Ba versus K (atomic) for groundmass micas from the Iron Mountain (Wyoming) and Chicken Park (Colorado) kimberlites.
MINERAWGY OF ORANGElTES
18 - . \ . .. ~ . - •
17 - Oo} •• • o -:.
- 0 .... , 0 • · , • IRON MOUNTAIN - · t·~ 0 CHICKEN PARK
- •
16
• • • 15 - • • • t • • j
..: ~
14
If) 13 o ~ «
12
II
10
9
-- •
-- • - •
• -
• - •
-
--
--
LEUCITE HILLS PRIMITIVE MICAS
r----I-, I I
• l \
I I
1 I I .. I • • 0 • • I • I • 0 ,. 0
I 0 0
~----_Oi •
ORANGEITE 0
PRIMITIVE COO 0 MICAS 0
•
I I 2 4
Ti02 wt. % •
135
Figure 2.36. Al203 versus Ti02 compositional variation of groundmass micas from the Iron Mountain (Wyoming) and Chicken Park (Colorado) kimberlites. Field of compositions of Leucite Hills primitive lamproite micas from Mitchell and Bergman (199\).
136 CHAPTER 2
occurred, as inferred from the common presence of oscillatory-zoned crystals. Mica types A and B have not crystallized in situ and are interpreted as transported hybrid microphenocrystal assemblages.
2.1.B.2.h. C%rado-»yoming Kimberlites. The Iron Mountain kimberlites of Wyoming (Smith 1977, McCallum et al. 1975) contain abundant poikilitic plates (Figure 2.34) of colorless groundmass micas that are continuously, irregularly-zoned from Ba-poor cores to Ba-rich margins (Table 2.11, Figure 2.35). The compositional variation is from Ba-bearing phlogopite to barian phlogopite. Increasing BaO (1.2-11.3 wt %) is accompanied by increasing Ah03 (11.2-18.0 wt %). Ti02 and FeOT contents initially decline then increase slightly as Al increases (Figures 2.36 and 2.37). Individual intrusions
18
17
f 16
15
~ ..: 14 ~
",
o 13 ~ c:[
12
II
10
9
·.tI • • e • .. 0-:. •• •
••• • ti •
• :''0 • o .. : .. . -, • • • •
•• • • •
•
• IRON MOUNTAIN
o CHICKEN PARK
r-------,
'I~'·. . 0 PRIMITIVE :\. ~ LEUCITE
HILLS
MICAS I ~ 0 I. • 1·0 I • 0 0
: ORANGEITE 0 0 ~ L________ PRIMITIVE 0 MICAS
2 3 4 5 6 7 8 9
o 00 0
o
10
Figure 2.37. AI203 versus FeOT compositional variation of groundmass micas from the Iron Mountain (Wyoming) and Chicken Park (Colorado) kimberlites. Composition field of Leucite Hills primitive lamproite micas from Mitchell and Bergman (1991).
MINERAWGY OF ORANGEITES 137
Table 2.12. Representative Compositions of Groundmass Micas from the Chicken Park Kimberlitea
"FeOT = total Fe expressed as FeO; n.d.= not detected. CaO and NiO not detectable by electron microprobe. Compositions 1-2 Chicken Park I; 3-8 Chicken Park 3; Compositions 6-7 are tetraferriphlogopites. (All data this work.)
differ with respect to the degree of Ba enrichment, e.g., Iron Mountain 7 micas contain only 1.2-3.3 wt % BaO, while Iron Mountain 25 micas have 1.2-11.3 wt % BaO.
The Chicken Park kimberlite, Colorado (McCallum 1989), contains abundant strongly pleochroic brown, poikilitic groundmass phlogopites. The majority of these micas (Table 2.12) are rich in FeOT (3.0-10.1 wt %) and Ti02 (0.6-3.7 wt %) relative to Iron Mountain (Figures 2.36 and 2.37) and other kimberlite groundmass micas. In some instances narrow mantles of Fe-poor tetraferriphlogopite (FeOT = 10.7-14.8 wt%, Ti02 < 1.0 wt %, BaO <0.5 wt %) are developed upon these micas. The groundmass phlogopites are BaO poor «0.2-2.0 wt %) with those richest in AI having the lowest Ba and Fe contents (Table 2.12). Chicken Park 1 contains micas that are BaO poor «0.5 wt %) relative to those from Chicken Park 3 (BaO> 1.0 wt %).
2.1.B.2.c. Guinea Kimberlites. The Antochka and Bounoudou kimberlites of the Guinea kimberlite province (Bardet 1974) contain abundant colorless and pale yellow laths, and poikilitic plates of groundmass mica, respectively.
138 CHAPTER 2
Table 2.13. Representative Comparisons of Groundmass Micas from Guinea Kimberlitesa
DFeOr= total Fe expressed as FeO; n.d. = not detected. CaO not detectable by electron microprobe. Compositions 1-4 Antochka; 5-10 Bounoudou. Compositions 5 and 7, and 6--8 are cores and rims, respectively. All data this work.
Microphenocrystal micas in the Antochka kimberlite (Table 2.13) contain 12.0-14.1 wt % A}z03, 6.4-8.5 wt % FeOT, and 0.2-2.1 wt % BaO. Rarely, they exhibit thin rims of AI-poor tetraferriphlogopite-Iike mica (Table 2.13, anal. 4). Individual grains are homogeneous, although considerable intergrain compositional variation is evident (Figures 2.38 and 2.39).
Groundmass plates and laths of mica in the Bounoudou kimberlites (Table 2.13, Figures 2.38 and 2.39) are, in contrast, rich in Ah03 (15.3-18.4 wt %) and BaO (0.1-5.7 wt %) (Figure 2.40). FeOT contents (3.~.8 wt %) are relatively low. The micas are continuously irregularly zoned with respect to their Ba contents (Figure 2.41). Although the majority of the cores are poor in Ba and Al relative to crystal margins, the reverse situation may also be found. The overall compositional trend is one of increasing Ba, AI, and Fe coupled with decreasing Ti and Si (Table 2.13).
2.1.B.2.d. Chinese Kimberlites (Shandong Province). The Xi-Yu and Shengli (a.k.a. Changma) kimberlites of the Shandong province (Dobbs et al. 1994) contain abundant
MINERAWGY OF ORANGEITES 139
- . • • • • • 18 - • • • • •
- '.' • • • .. " • 17 -
- • • • • 16 -
•
i - • BOUNOUDOU •
0 ANTOCHKA • 15 -
• -ae 14 - 0 . ~ 0 ~ .' - r--O--~ • It) I OJ
0 13 -~ I 0 « - I
0 I LEUCITE I 0 HILLS 12 -I PRIMITIVE I .... MICAS -I I
II - I 0 I
- L... ______
10 - ORANGEITE .... PRIMITIVE 0
MICAS Figure 2.38. AI203 versus Ti02 -compositional variation of ground· mass micas from the Bounoudou and 9 -Antochka kimberlites (Guinea). Com- I I I I positional field of Leucite Hills primi- 2 3 4 tive lamproite micas from Mitchell
Ti02 Wt. % .. and Bergman (1991).
140 CHAPl'ER2
• • • • 18 .. • • • ••• • • • • • • 17 • • • •
•
1 o· • • • BOUNOUDOU
16 o ANTOCHKA
• • 15
• ~
14 0 • .: •• 0 ~ r------, • 0
I 0
If) 13 0 LEUCITE C\I <i HILLS
12 PRIMITIVE MICAS
ORANGEITE 0
II PRIMITIVE 0 L _______
MICAS
10 0
9
2 3 4 5 6 7 8
FeOT wt. %
Figure 2.39. AI203 versus FeOr compositional variation of groundmass micas from the Bounoudou and Antochka micas (Guinea). Compositional field of Leucite Hills primitive lamproite micas from Mitchell and Bergman (199\).
colorless poikilitic groundmass micas. These are strongly continuously or complexly oscillatory zoned (Figure 2.42), Ah03-rich (14.9-18.0 wt %), barlan (1.0-12.20 wt % BaO) phlogopites (Table 2.14, Figures 2.40, 2.43, 2.44). Fe and Ti contents do not vary significantly with Al and Ba contents. Tetraferriphlogopites are absent. Zonation patterns appear to be random and the cores of coexisting crystals mayor may not consist ofBa-rich phlogopite (Figure 2.42).
2.1.B.2.e. Namibia. In Namibian kimberlites, ground mass phlogopite forms small «0.1 mm) colorless equant grains which include previously formed spinels and
Figure 2.40. Ba versus K (atomic) for groundmass micas from Guinean, Chinese, and Namibian kimberlites.
perovskites (Spriggs 1988). The micas are Ah03-rich (14.0-18.4 wt %), Ti02-poor (0.8-1.6 wt %), barian (1.3-11.5 wt % BaO) phlogopites (Figures 2.40, 2.43, 2.44). Spriggs (1988) noted that all of the crystals were continuously-zoned from Ba-rich cores to Ba-poor rims and suggested that the onset of groundmass mica crystallization resulted in rapid depletion of Ba in residual liquids, due to very high crystal-liquid distribution coefficients for Ba. However, evidence from other kimberlites (this work) suggests that there is no such simple trend of Ba depletion with crystallization. Spriggs (1988) was the first to document the existence of Ba-rich groundmass micas in kimberlites, although he did not recognize the solid solution trend toward kinoshitalite.
2.1.8.2.f. Orroroo. The majority of micas found in the Orroroo kimberlite, South Australia (Scott Smith et al. 1984) are poikilitic colorless AI-rich phlogopites mantled by discrete thin rims of red tetraferriphlogopite. The aluminous micas are weakly zoned with Ti02 (1-2 wt %), BaO (0.8-3.9 wt %) and Ah03 (15.4-13.2 wt %) decreasing from core to rim. The micas contain 0.9-2.2 wt % F. Their compositional variation is illustrated in
Figure 2.41. Complex compositional zoning in groundmass micas from the Bounoudou kimberlite (Guinea). Light gray regions are enriched in barium. Backscattered electron image (500x).
Figure 2.42. Extremely complex oscillatory zoning in groundmass micas from the Xi-Yu kimberlite (China). Light gray regions are enriched in barium. Backscattered electron image (500x).
142
MINERALOGY OF ORANGElTES 143
Table 2.14. Representative Compositions of Groundmass Micas from the Xi-Yu and Shengli KimberlitesQ
aFeOr = total Fe expressed as FeO; n.d. = not detected. Cao not detectable by electron microprobe. Compositions 1-5, Xi-Yu; 6-10, Shengli. (All data this work.)
Figures 2.45-2.47. Tetraferriphlogopite mantles are poor in F (<0.3 wt %), Ah03 «1 wt %), and Ti02 «0.1 wt %) and rich in FeOr (15.0 wt %) relative to Al-phlogopites.
2.1.B.2.g. Benfontein. The highly evolved Benfontein calcite kimberlite (Dawson and Hawthorne 1973) contains colorless poikilitic groundmass plates of al uminous micas that are very rich in BaO (Mitchell 1994b; Ah03 = 16.8-21.0 wt %; BaO = 15.9-23.8 wt %). These micas contain limited amounts ofK20 (2.1-4.8 wt %) and are best described as potassian kinoshitalite. Individual crystals are relatively uniform in composition, although considerable intergrain compositional variation is present (Mitchell 1994b; Figures 2.45-2.47). Macrocrystal and tetraferriphlogopite micas are not present.
2.1.B.2.h. Kirkland Lake Province. Micas in the kimberlites of the Kirkland Lake area of Ontario (Brummer et al. 1992) have been briefly described by Smith et al. (1978) and Arima et al. (1986). The kimberlite occurring as a dike in the Upper Canada (gold) Mine contains dark brown xenocrystal aluminous biotites similar to those in the southern African orangeites. These are partially resorbed and exhibit a wide range in FeOT (11-21
144 CHAPTER 2
- • CHINA -++-
II + NAMIBIA 18 - • I)
•• • • • - ,+ +
I •
17 -"i
.+ • +
-•
~ ++ + 16 -
., + - ++
~ • It) - + + ++
0 • .,i\I <{ 15 - + +
• •• • -
14 -LEUCITE HILLS
PRIMITIVE MICAS - ORANGEITE r---------., Figure 2.43. AI203 versus Ti02
PRIMITIVE I I compositional variation in ground-MICAS..,. l I
mass micas from the Xi-Yu and I 13 - I I Shengli kimberlites (China) and Na-
I I I I I I mibian kimberlites (Spriggs 1988).
I 2 3 Compositional field of Leucite Hills
Ti02 wt. % • primitive iamproite micas from Mitchell and Bergman (1991).
wt %) and Ti02 (3--6 wt %) content. These micas are mantled by phlogopites similar in composition to macrocrysts and microphenocrysts in kimberlite and orangeite. Tetraferriphlogopite and barian phlogopite groundmass micas are apparently absent.
The Nickila Lake kimberlite contains phlogopite macrocrysts exhibiting a limited range in composition (<2 wt % BaO, <1.5 wt % Cr203. 1-4 wt % Ti02. 5-12 wt % FeOT). Micas in volcaniclastic rocks are poorer in Ba. Fe. and TI and richer in Cr than micas in hypabyssal facies rocks. Tetraferriphlogopites and aluminous biotites are absent.
Arirna et al. (1986) do not tabulate sufficient data to allow determination of Al-TI and AI-Fe evolutionary trends. Although described as micaceous kimberlites by Arima et al. (1986). the apparent absence of the tetraferriphlogopite trend indicates that the rocks are unlikely to be derived from orangeite-type magmas. It is suggested here that the rocks are archetypal kimberlites which have been modally enriched in rnacrocrystaVphenocrystal
MINERAWGY OF ORANGElTES
~
.." o t\I
18
16
<{ 15
14
13
-
-
-
-
-
-
-
-
-
-
-
-
* • • • • ••• • • • • ,
+ • , + + •
+ • • + +
t •• + • • + •
+ + + + •
+ • • • • • CHINA
+ NAMIBIA
LEUCITE HILLS PRIMITIVE MICAS
r--------------, I I
~ : I I I I I I
2 3 4
FeOT wt. % ..
145
•
+
I
5
Figure 2.44. AI203 versus FeOr compositional variation in groundmass micas from the Xi-Yu and Shengli kimberlites (China) and Namibian kimberlites (Spriggs 1988). Compositional field of Leucite Hills primitive lamproite micas from Mitchell and Bergman (1991).
mica. Aluminous biotites are considered by Arima et al. (1986) to be crustally-derived xenocrysts (however, see 2.1.4).
2.1.8.2.i. Siberian Kimberlites. The Zagodochnaya kimberlite (Egorov et al. 1991) contains large (5 mm) macrocrystal micas and smaller (2-3 mm) "ground mass micas." Insufficient textural information is provided for the groundmass micas, although they appear to be microphenocrysts. The macrocrysts are poor in Ti02 (<0.5 wt %) and Cr203 «0.2 wt %) relative to the "groundmass" micas (2-4 wt % Ti02, 0.8-1.2 wt % Cr203). The Ah03 (13-14 wt %) and FeOr (3.8-4.7 wt %) contents of both types of mica are similar (Table 2.15). The ground mass micas are weakly-zoned toward increasing Ti and
146 CHAPTER 2
21 -t
~ + BENFONTEIN M MAYENG K KOIDU
20 * • ORROROO l U UDACHNAYA + Y YUBILEINAYA I Z ZAGODOCHNAYA ,.,
Y Y 19 "1- Y
t -1 \ I
18 -i Y + + Y ...J + Y
~ I Y U 17 -* ..: U ~ K - K If) J' 0 16 -~
M~ <t - U
• \ U
• K K .}I 15 - ... ~'" • M - • U K Y
LEUCITE HILLS K U U Z ~M 14 - PRIMITIVE· Y Y
Z MICAS· K U JK - r\---U--1 J
Z J
• I 1 U 13 - J
ulUz ORANGEITE U PRIMITIVE K - U
MICAS U
I 1 I 2 3 4
Ti02 wt. % • Figure 2.45. Ah03 versus Ti02 compositional variation of groundmass micas from diverse southern African. Australian and Russian kimberlites. Data sources Benfontein (Mitchell 1994b). Mayeng (Apter et al. 1984). Orrorroo (Scott Smith etal. 1984). Udachnaya (Egorov etal.1991. this work). Yubileinaya (this work). Zagodochnaya (Egorov etal. 1991). Compositional field of Leucite Hills primitive lamproite micas from Mitchell and Bergman (1991).
MINERALOGY OF ORANGEITES 147
21 + + BENFONTEIN M
+ MAYENG
K KOIDU
~ • ORROROO
20 U UDACHNAYA Y YUBILEINAYA Z ZAGODOCHNAYA
+ Y 19 + + Y y +
+ + + -;.
18 :tt' + y +
+ y ~ ++ y - 17 -tI- +
yU
~ + U K
K It) U U
0 16 K C\I M KK <{ K U
U K M U K
15 • K
• M
• U
Z uU
14 • K Z y y
• y U UK r------ ZI ZY
Y
I LEUCITE HILLS I. 13 U I PRIMITIVE K I MICAS I
2 3 4 5 6 7
FeOT wt. % • Figure 2.46. AI203 versus FeOT compositional variation of groundmass micas from diverse southern African, Australian, and Russian kimberlites. Data sources as in Figure 2.45.
148 CHAPTER 2
Table 2.15. Representative Compositions of Mica in the Zagodochnaya and Yubileinaya Kimberlites, Russiaa
aFeOr = total Fe expressed as FeO; n.d. = not detected; n.a. = not analyzed. Compositions 1-5, Zagodochnaya (Egorov el 01. 1991); 6-10. YubiJeinaya (this work); I, macrocryst; 2-3, microphenocrysts; 4-5, core and rim microphenocryst; 6-7, core and rim groundmass micas; 8-10, groundmass micas.
Fe from core-to-margin. Figures 2.45 and 2.46 show that Zagodochnaya micas are relatively unevolved, Micas belonging to the tetraferriphlogopite trend are absent. The mica compositional data indicate that this kimberlite is an archetypal kimberlite which is modally enriched in macrocrystallphenocrystal mica.
The Yubileinaya kimberlite contains small, anhedral, light brown microphenocrystal micas that have been replaced at their margins by Ba-bearing mica (this work). BaO contents are typically from 0.1 to 5 wt %, although higher levels may be rarely found (Table 2.15, anals. 7 and 10). Figures 2.45 and 2.46 show that many of the micas are relatively unevolved. Increases in AI are accompanied by a decrease in Ti and increase in Fe. Groundmass poikilitic micas and tetraferriphlogopites are absent in the suite of samples examined in this work.
MINERAWGY OFORANGEITES
~ 1'0
-
....... CI)
0'6
E 0'4 o -c - 0'2
0'2
+ BENFONTEIN
• ORROROO
• YUBILEINAYA
~ Bo - MICAS IN
0'4
CARBONATITES, LEUCITITES , NEPHELINITES
0'6 0'8 1'0
Ba( atoms / II oxygens) ~
149
Figure 1.47. Ba versus K (atomic) for groundmass micas from the Benfontein (Mitchell 1 994b), Orroroo (Scott Smith et al. 1984), and Yubileinaya (this work) kimberlites. Compositional field of Ba-rich micas in carbonatites, leucitites, and nephelinites from Mitchell (1994b).
Micas in the groundmass of the East Udachnaya kimberlite (Egorov et al. 1991, this work) are anhedral, light-brown microphenocrysts. Many of these micas are similar in composition (Table 2.16, Figures 2.45 and 2.46) to unevol ved macrocrystal micas in other kimberlites. They are considered to be cognate earlier-fonning phases which have not crystallized in situ. The margins of the crystals are slightly enriched in BaO. Colorless groundmass poikilitic plates ofbarian phlogopite are absent. Phlogopite in mica-pyroxenite clasts from the same locality (Egorov etat. 1986) is of a similar composition to the macrocrystal micas (Table 2.16, anal. 10).
The West Udachnaya kimberlites contain groundmass micas identical in their paragenesis, and of similar composition (Table 2.16, Figures 2.45 and 2.46) to those of East Udachnaya. They differ in being relatively poor in BaO (0.3-0.9 wt %) and richer in FeOT (5.1-1.0wt%).
2.1.8.2.j. Aries. The Aries kimberlite (Western Australia) contains abundant phenocrystal and microphenocrystal mica (Edwards et al. 1992), although poikilitic groundmass plates appear to be absent. Phlogopite compositional zoning is weak, but complex, and differs in each portion of the intrusion. The overall composition is similar to that of unevolved macrocrystal and phenocrystal phlogopites from kimberlites, orangeites, and lamproites (Figure 2.48). Complex normal and reverse zoning trends indicate that magma mixing played a significant role in the development of the mica assemblage. Some of the
150 CHAPTER 2
Table 2.16. Representative Compositions of Micas from the Udachnaya KimberiiteO
,'!FeOT = total Fe expressed as FeO;n.d. = not detected; n.a.= not analyzed. Compositions 1-4. East Udachnaya (this work); 5-8. West Udachnaya (this work); 9. average composition of three "microphenocrysts." East Udachnaya (Egorov et al. 1991); 10. phlogopitein mica pyroxenite clast. East Udachnaya (Egorov et al. 1986).
more evolved micas contain small quantities of tetrahedral Fe3+, indicating the presence of minor amounts of the tetraferriphlogopite molecule. However, micas exhibiting the extreme Al depletion, characteristic of orangeites, are not present as the majority of micas are aluminous and plot in the Fe-Mg-AI ternary system (Figure 2.49) on the AI-rich side of the phlogopite-biotite join. Insufficient compositional zoning is present to state unequivocably (Edwards et al. 1992) that the Aries intrusion has mineralogical affinities to orangeites. The micas contain only 0.6-1.5 wt % BaD, and Ba-rich micas are absent.
2.1.B.2.k. Koidu. The Koidu kimberlite, Sierre Leone (Grantham and Allen 1960), contains reversely-pleochroic Fe-rich, Cr- and Ti-poor mica macrocrysts interpreted by Tompkins and Haggerty (1984) to be xenocrysts. These are mantled by Ti- and Cr-bearing phlogopites. Microphenocrystal or groundmass micas are similar in composition to the rims of cores and are, in tum, rarely mantled by thin rims of reversely-pleochroic, low-AI
+ core c intermediate zone • second intermediate zone o rim • groundmoss grain
/ zoning trend
Figure 2.48. Compositional variation of micas in the Aries kimberlites, Australia (after Edwards et al. 1992). ZI-Z4 Central Lobe; Sc, Sr, Mc, and Mr are cores and rims of grains from lithologies M and S of the North Lobe. K = trends of "kimberlite" mica compositions from Mitchell (I 986).
tetraferriphlogopite-like micas. Unfortunately, mica compositions are not tabulated by Tompkins and Haggerty (1984). Data obtained during the course of this work show that microphenocrysts contain 14.0-16.6 wt % Ah03, 2.5-3.4 wt % Ti02, and 5.0-7.5 wt % FeOT (Figures 2.45 and 2.46). The margins of the crystals rarely exhibit thin mantles of Ah03-poor «1.0 wt %) tetraferriphlogopite.
2.1.8.2.1. Mayeng. The Mayeng sills, South Africa (Apter et al. 1984), in addition to pale-colored phlogopite macrocrysts, contain two types of groundmass mica. All textural varieties of the kimberlite contain an early-forming, dark-colored mica which forms euhedral equant crystals. Following crystallization of this mica individual sills evolved separately and produced different assemblages of late-stage micas as mantles upon preexisting dark micas. Two of the sills thus crystallized pale-colored micas which poikilitically enclose spinels, while a third crystallized dark-red, reversely-pleochroic phlogopite.
The early dark micas contain 4-6 wt % Ti02, 0-1 wt % Cf203, 4.5-5.3 wt % FeOT, and 13.3-15.6 wt % Ah03. The pale micas are also rich in Ah03 (13.8-15.8 wt %), but are poorer in Ti02 (2.5-2.9 wt %) and Cr203 (0.04-D.12 wt %). Low analytical totals for these micas may reflect the presence of Ba not determined by Apter et al. (1984). The
152
A
5
c
AI
15 25 15
AI D
, , ++ + , , , , + , \ , ,
, \ + \ , \ \+ \ + ANTOCHKA \
, • BOUNOUDOU \
AI
,"P '"P '---r~r--r-r---7':":":''''' FeT '---r---.,,..--r--r-""';'!.!:.!:+ FeT Mg ~ 15 25 Mg 5 [5 25
E AI
50
/ ~o, 70 .ttl % .,
li.OfJooo 0 8 \-~·t~~---
\ '·.'f' \
• ARIES • ORROROO
x BENFONTEIN o KOIDU \ ,
F
\ TFP '---r---:r---r---:r--+.:....:..... FeT
Mg 5 15 25 Mg 5
AI
70
PHL ,----- - --\
\ , • UDACHNAYA WEST o UDACHNAYA EAST
YUBILEINAYA
+ ZAGODOCHNAYA \ \ TFP
15 25
CHAPfER2
Figure 2.49. Compositions of micas from diverse kimberlites plotted in the ternary system Al-Mg-FeT (atomsl22 oxygens). Data sources given in the text. FeT = total Fe expressed as Fe2+. EAST = "eastonite," PHL = phlogopite, TFP = tetraferriphlogopite.
MINERAWGY OF ORANGElTES 153
evolutionary trend appears, on the basis of the limited data tabulated, to be one of increasing Al and Fe with decreasing Ti (Figures 2.45 and 2.46) and is thus similar to that found in other archetypal kimberlites. Tetraferriphlogopites have low Ah03 «2 wt %), Ti02 (0.3-0.7 wt %), and Cr203 «0.1 wt %), and moderate FeOT (10-13.3 wt %) contents. They are relatively poor in FeOT compared to similar micas in orangeites.
Apter et al. (1984) consider the pale micas to be representative of the final stages of crystallization of this kimberlite but provide no explanation as to why tetraferriphlogopite forms in one textural variety. However, it is suggested here that the differences are due to different late-stage liquid compositions and/or conditions of crystallization subsequent to emplacement.
2.1.B.2.m. Fayette County. The Fayette County (Pennsylvania) kimberlite dike contains microphencrystal and groundmass mica (Hunter et al. 1984). The microphenocrysts are zoned toward margins whose compositions overlap those of the groundmass micas. All micas are relatively Ah03 rich (approx. 14.0 wt %). No systematic variations in Al content are evident, although there is a change in Cr, Ti, and Fe with paragenesis. Hence, phenocryst cores contain 1.3-1.8 wt % Cr203, 4.5-5.0 wt % Ti02, and 4.6-5.5 wt % FeOT. The rims contain 0.1-1.0 wt % Cf203, 3.6-4.5 wt % Ti02, and 4.9-6.0 wt % FeOT. The groundmass micas contain <0.2 wt % Cr203, 1.7-4.5 wt % Ti02, and 4.9-6.0 wt % FeOT. The overall trend of compositional evolution is one of decreasing Cr, Ti, and increasing Fe from core, via rim, to groundmass. Tetraferriphlogopite is absent.
2.1.B.2.0. Blue Ball. The Blue Ball (Arkansas) "kimberlite" (Sal pas et al. 1986, Neal and Taylor 1989) contains phenocrystal and groundmass phlogopites, in addition to micas which occur as coronas about olivine phenocrysts. The micas range in composition from phlogopite to tetraferriphlogopite. Two compositionally distinct varieties of phenocrysts are recognized on the basis of their mg numbers (Figure 2.50) or Ti02 «1.0 wt % or 1-1.5 wt %) and BaO «0.5 or >0.5 wt %) contents. Phenocrysts are mantled by micas exhibiting a very wide range in composition with respect to their Al and Fe contents. Some rim compositions are identical to those of micas defined as groundmass mica. Evolutionary trends from core to rim in the phenocrysts are of decreasing AI, Ba, and Ti with increasing Fe and octahedral site deficiency. The composition of the groundmass micas encompasses the entire range of phenocryst core-to-rim compositional variation with respect to Fe but not Al (Figure 2.50). Some of the groundmass micas may be enriched in Ti02 (up to 3 wt %) relative to the phenocrysts, yet others have compositions identical to those of the phenocrysts. The data suggest that all of the micas are of one generation/paragenesis, thus, many of the small micas interpreted by Neal and Taylor (1989) as ground mass mica may be in reality microphenocrysts. The mica assemblage has similarities to that of orangeites in being a complex transported assemblage.
Figure 2.51 shows the evolutionary trend of the Blue Ball micas is identical to that of orangeite micas. This trend, the absence of micas belonging to the phlogopite-kinoshitalite series and other mineralogical features, e.g., spinels ranging in composition from chromite through Ti-magnetite to magnetite, coupled with the absence of magnesian ulvospinel, suggests that the Blue Ball intrusion is not an archetypal kimberlite. Samples examined during the course of this work showed the rocks to consist of serpentinized
154
2'0
I- 1'0 <J
0'5
40 50 60
Mg ~ 70 80
CHAPTER 2
90
Figure 2.50. Tetrahedral site deficiency (~T) versus mg number of micas from the Blue Ball "kimberlite," Arkansas (after Neal and Taylor 1989). Stars represent the earliest phenocrysts. Solid circles represent the cores of the second group of phenocrysts. Open circles represent phenocryst rims. Tie lines join core and rims of the same phenocryst. The inset demonstrates the evolutionary path of compositions .
t 12
10
..: ~
It) 6 o
C\I
Ci 4
2
•
ORANGEITE • PRIMITIVE MICAS
BLUE BALL ARKANSAS
•
3
Ti02 -wt. %--.
6 8 10 12 14 16 18
FeOT Wt. %
20
• 22
•
24
Figure 2.51. Compositional variation of micas from the Blue Ball "kimberlite," Arkansas (after Neal and Taylor 1989).
MINERALOGY OF ORANGElTES 155
olivine set in a matrix of mica, apatite, and carbonate. Also present as accessory phases are perovskite (poor in Sr, Nb, REE), rutile, Ca-REE phosphate, Ni-pyrite, djerfisherite, barite, and Nb-Mn ilmenite. The rocks have mineralogical affinities with both lamproites and orangeites, but without further mineralogical study defy exact classification. Their proximity to the Arkansas lamproite province (Mitchell and Bergman 1991) and possible similarity in age (? mid-Cretaceous) suggests that they may be a part of this province.
2.1.8.2.n. West Greenland. The compositional variation of micas in the West Greenland kimberlites (Scott 1981, Larsen and Rex 1992) has been insufficiently studied to allow determination of evolutionary trends. Data given by Scott (1981) show the macrocrystal and groundmass micas to be Ti-bearing phlogopite (Ah03 = 11.5-15.5 wt %, Ti02 = 0.4-4.72 wt %, FeOr = 6.1-10.21 wt %). Many groundmass crystals exhibit rims of tetraferriphlogopite. Barian micas appear to be absent.
2.1.8.3. Summary of Kimberlite Mica Compositional Variation
Macrocryst compositions are not significantly different from macrocrysts/microphenocrysts in orangeites. Macrocrysts and the cores of microphenocrystal (and/or groundmass) micas are relatively low in AI, compared with crystal margins and poikilitic groundmass micas. Two evolutionary trends may be discerned on the basis of the data given above. The dominant and characteristic one is Al enrichment. Figure 2.49A-C shows that mica compositions belonging to this trend plot away from the phlogopitebiotite join toward aluminous phlogopites and the hypothetical "eastonite" composition. Micas following this trend are commonly enriched in BaO, and late-stage poikilitic micas are colorless members of the phlogopite-kinoshitalite series. Late-stage Ba-rich micas are commonly depleted in FeOr relative to unevolved micas. Ti02 contents are low «4 wt %) and do not show any systematic trends. The high Ti02 (>5 wt %), Ba-rich, pink micas characteristic of the groundmass of leucitites and melilitoids are absent.
The second, less common, evolutionary trend is toward tetraferriphlogopite. This trend is illustrated in Figure 2.490 by mica mantles from the Antochka kimberlite. Importantly, these AI-poor micas are typically developed as discrete thin mantles upon cores of AI-rich groundmass micas. Tetraferriphlogopite formation does not occur in all facies of a given kimberlite and has not occurred in most archetypal kimberlites. The abrupt change from aluminous mica crystallization to tetraferriphlogopite indicates sudden, drastic changes in the redox conditions prevailing during the final stages of crystallization. Its formation may be associated with the addition of ground waters to the magma and/or rapid carbon dioxide loss.
In conclusion, archetypal kimberlites may easily be distinguished from orangeites on the basis of the different evolutionary trends of groundmass micas (Figures 2.49, 2.52-2.53). Although some kimberlites do contain groundmass tetraferriphlogopite, its presence cannot be regarded as characteristic. Tetraferriphlogopite development appears to be random and is principally confined to thin discrete mantles upon preexisting AI-rich phlogopites. This paragenesis is in marked contrast to the ubiquitous presence oftetraferriphlogopite, as either the outer margins of continuously-zoned groundmass micas or late-stage poikilitic plates in orangeites.
156
1 ~ .; ~
If)
0 N
<t
t ae ..: ~
II)
0 ~ «
20
18
16
14
12
10
8
6
4
2
20
18
16
14
12
10
8
6
4
2
2 4 6 8 10 12
Ti02 Wt. % •
CHAPfER2
Figure 2.52. Compositional (A1203 versus Ti02) evolutionary trends of micas from orangeites (this work).lamproites (Mitchell and Bergman 1991). minettes (Mitchell and Bergman 1991). and kimberli tes (this work).
FIELDS OF MICROPHENOCRYST COMPOSITIONS
ffllI ORANGEITE [] LAMPROITE
2 4 6 8 10 12 14 16 18
Figure 2.53. Compositional (A1203 versus FeOr) evolutionary trends of micas from orangeites (this work). lamproites (Mitchell and Bergman 1991). minettes (Mitchell and Bergman 1991). and kimberlites (this work).
MINERALOGY OF ORANGElTES 157
2.1.9. Mica in Lamproites
The compositional variation shown by micas in lamproites is well characterized and has been summarized by Mitchell and Bergman (1991).
Individuallamproite provinces contain phenocrystal micas of distinct composition with respect to their Ti, AI, and Fe contents. Within a lamproite province phenocrysts may exhibit relatively little compositional variation or vary considerably from vent-to-vent. Groundmass micas and mantles upon phenocrysts are typically depleted in Cr and AI, and enriched in Fe, Ti, Ba, and Na relative to phenocrysts, and thus may be considered more evolved than the latter.
Mitchell and Bergman (1991) have shown that compositional and zonation trends with respect to the AI, Ti, and Fe contents of mica may be used to assess the degree of evolution of the magma from which they crystallized. Thus, micas with high Ah03 (10.5-\3.5 wt %), 4-10 wt % Ti02, and 2-4 wt % FeOT typically occur as phenocrysts. Such micas are considered to be relatively primitive un evolved micas which may have formed in the magma prior to eruption. Micas with low Ah03 (0.5-10 wt %) and high FeOT (>4 wt %) form mantles upon phenocrysts and occur as groundmass poikilitic plates, and are interpreted to have formed at relatively low temperatures and pressures after eruption or emplacement of the magma.
Compositional evolutionary trends fall between two extremes (Figures 2.52-2.56). One is of slight-to-moderate Al depletion, coupled with increasing Ti and Fe and decreasing Mg, i.e., a trend reflecting Fe2+ increase, and representing evolution from
AI
MO TFA 5 10 15 20 30 35 40 45 50 55 60 65
F 2+ __ ...... eT •
Figure 2.54. ComP9sitional evolutionary trends (A-F) of micas from di verse lamproi tes plotted in the ternary system Al-Mg-FeT2+ (atomic). A = Hatcher Mesa, Leucite Hills; B = Middle Table Mountain, Leucite Hills; C = Pilot Butte and Badger's Teeth, Leucite Hills; D = Mount North and Rice Hill madupitic lamproites, West Kimberley; E = Mount North phlogopite lamproites, West Kimberley; F = Mount Gytha, West Kimberley. PHL = phlogopite, TFP = tetraferriphlogopite, TFA = tetraferriannite.
158 CHAPTER 2
20 FIELDS OF MICRO PHENOCRYST COMPOSITIONS m ORANGEITE
.-., L _J LAMPROITE
18 • MINETTE
16
f 14
12 ~
..: ~
10
II)
0 01
<i 8
6
4
2
2 4 6 8 10 12
Figure 2.SS. A1203 versus 1102 compositional variation of micas from minettes. Data sources given in Mitchell and Bergman (1991). Compositional trends for micas in orangeites, lamproites, and kimberlites from Figure 2.52.
titanian phlogopite toward titanian biotite. The other is a trend of strong Al depletion associated with increasing Ti and Fe at essentially constant Mg content, i.e., a trend reflecting increasing Fe3+, and representing evolution from titanian phlogopite toward titanian tetraferriphlogopite. The compositional trend exhibited by mica in a given lamproite may lie anywhere between these two extremes and reflects the local postemplacement crystallization environment with respect to redox conditions, water content, and cooling conditions.
Lamproite phlogopites as a group have insufficient Si and Al to fill the tetrahedral sites in the mica structure, consequently Alvi is absent from phenocrystal and groundmass
MINERALOGY OF ORANGElTES
ae 14 ..: • .., 12
o N «
8
2 4 6
159
8 10 12 14 16 18 20
FeOT wt. %
Figure 2.56. Al203 versus FeOr compositional variation of micas from minetles. WG = Wattle Gill. England; CEL = Celebes. Indonesia; DH = Dale Head. England; SC = Shaws Cove. Canada; NAV = Navajo. U.S.A.; COL = Colima, Mexico; LIN = Linhaisai. Indonesia; DEV = Devonshire. England; H = Holmead Farm. England; BOH-Aalkaline minettes. Bohemia; BOH-B otherminettes. Bohemia. Data sources given in Mitchell and Bergman (1991). LH-P = Compositional fields of Leucite Hills lamproite phenocrystal micas. Compositional trends for micas from kimberlites. orangeites. and lamproites from Figure 2.53.
micas. The tetrahedral site deficiency may be accommodated by the entry of Ti4+, Fe3+, or Mg2+ to this site. Solid solutions present are primarily between phlogopite-biotite, octahedral site-deficient titanian phlogopite [K2(Mg,Fe~TiDSi6Ah02o(OH~], and tetraferriphogopite-tetraferriannite.
Other characteristic features of lamproite mica compositions are the high F (1-7 wt %) contents and enrichment of Na in groundmass micas. The latter contain 0.5-1.8 wt % Na20 and may be regarded as sodian titanian tetraferriphlogopites.
Orangeite micas have some compositional similarities to lamproite micas in that similar evolutionary trends with respect to AI-Fe are present. Figures 2.53 and 2.54 show that orangeite micas define compositional trends which overlap with those oflamproites. The trends originate on AI-Fe diagrams at similar primitive unevolved phenocryst compositions. By analogy with lamproites it is suggested that which trend is followed must depend upon local post-emplacement crystallization conditions.
Significant differences exist with respect to AI-Ti compositional trends and the Ti contents of micas. Lamproite micas are typically enriched in Ti02 (1-12 wt %) relative to orangeite micas, and compositions evolve toward increasing Ti contents (Figure 2.52). In contrast, the majority of orangeite micas are relatively poor in Ti02 (1-3 wt %) and evolution toward tetraferriphlogpite is accompanied by decreasing or constant Ti contents (Figure 2.52) with decreasing AI.
160 CHAPTER 2
However, other orangeites, which have evolved toward sanidine and richterite-bearing residua, contain micas with Al-Ti evolutionary trends similar to those oflamproites. Thus, micas from Postmasburg and Voorspoed follow typicallamproite mica Al-Ti evolutionary trends (Figures 2.6, 2.7,2.52,2.53), and AI-poor groundmass micas in the Besterskraal and Sover North occurrences are identical in composition to groundmass micas in some lamproites (Figures 2.5, 2.17, 2.52, 2.53).
Orangeite micas may be distinguished from lamproite micas on the basis of their low Na (typically 0.5 wt % Na20; Table 2.5) and F «1.0 wt %) contents relative to lamproite micas. Further, the complex mantling and reverse zoning patterns found in the mica microphenocryst populations of orangeites are not present in lamproite micas.
2.1.10. Mica in Minettes
Micas in minettes form a complex-mantled, continuously- or reversely-zoned assemblage of phenocrysts (Bachinski and Simpson 1984, Jones and Smith 1983, 1985, Schulze etal. 1985, Bergman etal. 1988, Mitchell and Bergman 1991, Meyer etal. 1994). Phenocrysts show wide variations in composition within and between minette occurrences. The cores of phenocrysts are, in some instances, compositionally similar to groundmass micas. Many crystals commonly have oxidized and/or resorbed margins.
Mitchell and Bergman (1991) have summarized the compositional variation of minette micas and shown that the characteristic evolutionary trend is one of increasing Fe with slightly increasing or constant Al (Figure 2.55). The Ti content may increase or decrease slightly with respect to Al (Figure 2.56). Many minette micas are aluminous and do not exhibit any tetrahedral site deficiency. The presence of Alvi indicates solid
AI , +-----.!..,-----SID
/ ~ 70
~ PHL leo
10 50 60
Figure 2.57. Compositional field of micas from minettes plotted in the ternary system AI-Mg-FIlT2+ (atomic). Data sources as in Figure 2.56 plus Meyer et al. (1994). EAST = "eastonite," SID = siderophyllite. PHL = phlogopite. ANN = annite. TFP = tetraferriphlogopite. TFA = tetraferriannite.
MINERALOGY OF ORANGEITES 161
solution toward "eastonite"-siderophyllite end-member molecules (Figure 2.57). Tetraferriphlogopites have not been reported from minettes.
Figures 2.52, 2.53, and 2.57 show that unevolved minette phenocrysts overlap the compositions of unevolved microphenocrysts from orangeites and lamproites. However, the compositional trends diverge from these common unevolved micas, such that the minette micas are always Al rich relative to orangeite and lamproite micas of equivalent Fe content. The evolutionary trend is from phlogopite toward titanian aluminous biotite and represents solid solution between phlogopite-biotite and the "eastonite"-siderophyllite molecules.
Minor elements in minette micas exhibit very wide ranges in concentration (F = 0.2-5.0 wt %, BaO = 0-2 wt %, Cr203 = 0-1 wt %) and cannot be used as discriminators between magma types.
In summary, orangeite micas may be easily distinguished from minette micas on the basis of their Al-Ti and AI-Fe evolutionary trends.
2.1.11. Mica in Ultramafic Lamprophyres
Discussion of the composition of micas in ultramafic lamprophyres is required because of the petrographic similarity of some olivine phlogopite lamprophyres to orangeites. Unfortunately, there have been few detailed studies of the compositional variation of micas in ultramafic lamprophyres, and most works tabulate only a few random compositions. Rock (1986, 1990) reviewed these rocks and noted that micas in them exhibit an extremely wide range in composition with respect to their AI, Fe Ti, and Ba contents. However, no characteristic evolutionary trends were identified. Tetraferriphlogopites occur in the groundmass of some examples (Rock 1990).
Table 2.17 and Figures 2.58-2.60 illustrate the compositional variation of mica in some ultramafic lamprophyres belonging to the alnoite-polzenite (or melnoite; Mitchell 1994c) suite. These rocks are characterized by widely varying modal amounts of olivine,
Table 2.17. Representative Compositions of Mica from Ultramafic Lamprophyresa
"FeOT = total Fe expressed as FeO; n.d. = not detected. CompoSitions 1-2, alntiite, Como, Quebec; 3-4, alntiite, type locality Alnti, Sweden; 5-6, polzenite, Polzen, Czech Republic; 7-8, alntiite, N. Hutson, Kentucky; 9-10, alntiite, Haystack Butte, Montana. (All data this work.)
162 CHAPTER 2
t 20 @PI B
• ~!'
18
~ 16 .,; ~ 14 If)
0 12 ..!I « 10
8
2 3 4 5 2 3 4 5 6
Ti02 wt. % • Figure 2.58. Compositional variation (AI203-Ti02) of micas in ultramafic lamprophyres. PI, P2, and P3 = diverse polzenites, Polzen, Czech Republic; HB = Haystack Butte, Montana; A-OPM = diverse olivine phlogopite alnoites, Alno complex, Sweden; A-TL = alnOite type locality, Alno, Sweden. All data this work, except Vestfold (Delor and Rock (1991) and Karinya (Muller et al. 1993). LHPM and OPM = compositional fields of microphenocrysts from Leucite Hills 1amproites (Mitchell and Bergman 1991) and orangeites (this work), respectively. K, 0, and L are compositional trends of micas from kimberlites, orangeites, and lamproites, respectively.
t 20
18
CIt! 16 ..:
~ It) 14
0
"""'''''''' ~I
Ki:~A~ \ . HB ------,
FEN
~12 <I LHPM
10
8
3 4 5 6 789 3 4 5 6 7 8 9 10 II
FeOr wt. %
Figure 2.59. Compositional variation (A1203-FeOT) of micas in ultramafic lamprophyres. Abbreviations and data sources as in Figure 2.58.
MINERALOGY OF ORANGEITES 163
AI AI
55 \
/
Mg
Figure 2.60. Compositional fields of micas from ultramafic lamprophyres plotted in the ternary system AI-Mg-Fer (atomic). MO = Monticellitovaya; other abbreviations as in Figure 2.58. Also shown are compositional evolutionary trends for groundmass micas in Somerset Island (SIK), Chinese (CK), Iron Mountain (1M), and Benfontein (B) kimberlites. EAST = "eastonite," PHL = phlogopite, TFP = tetraferriphlogopite. Fer = total Fe expressed as Fe2+.
phlogopite, spinel, perovskite, apatite, monticellite, and calcite. Melilite and clinopyroxene mayor may not be present. Although each occurrence is characterized by mica of a particular composition, their overall compositional evolution is typically toward increasing Ti and Fe, at constant or slightly increasing Al contents. All of the micas are rich in AI, and solid solutions are primarily between phlogopite-annite and "eastonite"siderophyllite. These aluminous micas commonly coexist with aluminous pyroxenes (2.2.7).
Compositional trends are unlike those observed in either orangeites or lamproites (Figures 2.58-2.60) but similar to those found in micas from minettes (Figures 2.55-2.57). The compositions of the relatively unevolved micas overlap with those of unevolved groundmass micas in archetypal kimberlites. In common with the latter they may in some occurrences, e.g., Haystack Butte (Wendlandt 1977), Polzen (this work), have significant Ba contents. These micas may be distinguished from Ba-rich micas in kimberlites on the basis of their evolutionary trends toward increasing Ti and Fe (Figures 2.58-2.60).
The diamond-bearing ultramafic lamprophyres from Bulljah Pool, Western Australia (Hamilton and Rock 1990) contain micas which are members of the titanian phlogopite-
164 CHAPTER 2
AI AI
I I
/ PHLOGO- I BIOTITES PITES I
t::. WO MT. BUNDEY IFI o PHENOCRYSTS
I •• t::. GROUNDMASS
• BULLJAH 10
Mg TFA
Fe2+ 90 80 60 50 40 30 20 10 T
-- Fe2+~ T
Figure 2.61. Compositions of micas from the Mount Bundy (Sheppard and Taylor 1992) and Bulljah Pool (Hamilton and Rock 1990) lamprophyres plotted in the ternary system AI-Mg-Fer2+ (atomic). Abbreviations as in Figure 2.60.
titanian tetraferriphlopite series with minor solid solution toward tetraferriannite and/or annite (Figure 2.61). They contain 6.4-11.7 wt % Ah03, 2-3 wt % Ti02 (rarely reaching 7 wt %), and <0.1 wt % Cr203. BaO contents are mostly <1 wt % but may reach 2.3 wt %. Figure 2.62 shows that the micas are slightly evolved relative to primitive micas in orangeites and lamproites. Unfortunately, Hamilton and Rock (1990) do not provide detailed phenocryst-groundmass compositional trends. Figure 2.62 suggests that the compositional trend is toward increasing Ti with decreasing AI, i.e., a lamproitic trend rather than an orangeite trend.
Olivine-mica-sanidine lamprophyre dikes from Mt. Bundey, Australia (Sheppard and Taylor 1992), have mineralogical affinities to evolved orangeites and lamproites. The micas in these rocks are AI-rich members of the phlogopite--biotite-siderophyllite series (Figure 2.61) which exhibit a wide range in Ti02 (1-8 wt %) and BaO (0.2-3.5 wt %). Figure 2.62 shows that they are richer in Al than primitive micas from orangeites and lamproites. Their compositional evolution toward increasing Ti (and Fe), at constant Al contents, serves to distinguish them from the latter rocks.
Micas in the Bow Hilllamprophyre dikes (Fielding and Jaques 1989) show a wide range in composition (Figure 2.62). Rare Ti02-rich (6 wt %) biotite cores of phenocrysts are mantled by AI-rich phlogopites (11.5-16 wt % Ah03) of variable Ti02 (1-2 wt %)
contents. These micas are richer in F (>1 wt %) and poorer in Crz03 «0.05 wt %) than the biotite (0.5 wt % F, >0.1 Cr203), and many are in turn mantled by narrow rims of Ti02-poor «0.5 wt %) tetraferriphlogopite. The overall zonation trend has similarities to that of orangeites rather than lamproites (Figure 2.62). Micas differ from those of orangeites in being AI-rich and coexisting with andradite--melanite garnet.
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166 CllAPTE1l2
2.2. CLINOPYROXENE
2.2.1. Paragenesis
Clinopyroxene is the only primary pyroxene found in orangeites. Modes vary widely and may range from trace quantities to as much as 26 vol % (Wagner 1914, Skinner and Clement 1979, Fraser 1987).
In unevolved orangeites, diopside typically forms small (mainly 0.1 to 0.5 mm and rarely up to 1.5 mm in length) microphenocrysts of colorless prismatic crystals, some of which are simply twinned. The diopsides in many instances are corroded and embayed and have, evidently, undergone resorption during transportation or crystallization of the mesostasis. The microphenocrysts are, rarely, continuously weakly-zoned to pale-yellow rims and/or mantled by pale-green thin rims of titanian aegirine. Strong continuous zoning, complex oscillatory zoning, sector zoning, and multiple epitaxial mantling are not characteristic of orangeite diopsides (this work, Mitchell and Meyer 1989a, Dawson eta/. 1977, Wagner 1914).
Clinopyroxene is most abundant in evolved orangeites. In these rocks it forms interlocking aggregates of relatively large resorption-free blocky crystals. Unlike the resorbed diopsides of unevolved orangeites these pyroxenes appear to have crystallized in situ. Thin green mantles of titanian aegirine are characteristic (Figure 1.48). Similar pyroxenes are found along fractures and cleavages. Commonly, pyroxenes are poikilitically included in groundmass mica and amphibole (this work, Tainton 1992).
Unevolved orangeite fields differ with respect to their pyroxene contents. In the Barkly West field, diopside is not a characteristic primary groundmass mineral (Tainton 1992, Bosch 1971). Here, diopside occurs as xenocrysts derived from country rock Ventersdorp basalts and in reaction assemblages adjacent to intrusion margins and xenoliths (Tainton 1992). In contrast diopsides are common in orangeites from the Boshof, Winberg. and Kroonstad fields, Swartruggens and Finsch (this work, Fraser 1987, Dawson et al. 1977, Wagner 1914).
2.2.2. Composition
Reconnaissance studies of the composition of groundmass c1inopyroxenes in orangeites have been undertaken by Dawson et al. (1977), Skinner and Scott (1979), Robey (1981), and Mitchell and Meyer (1989a). These studies demonstrated that the pyroxene is typically diopside of restricted compositional range and that significant inter-intrusion differences are not apparent.
The summary of the compositional character of diopside in orangeites presented below is based on over 250 new analyses of pyroxene obtained during the preparation of this monograph. Representative compositions of microphenocrystal pyroxene are given in Tables 2.18 and 2.19. These tables and Figures 2.63-2.65 confirm that clinopyroxenes in orangeites typically exhibit only limited compositional variation.
2.2.2.1. Diopside
The majority of the phenocrystal pyroxenes are Fe-poor diopsides exhibiting very little solid solution toward hedenbergite or aegirine (Table 2.18). The pyroxenes are poor
MINERAWGY OF ORANGEITES 167
Wt%
Si02
Ti02
AI20 3
Cr203 FeOr MnO MgO CaO Na20
Si AI Ti Cr Fe3+
Fe2+
Mn Mg Ca Na
Table 2.1S. Representative Compositions of Iron-Poor Clinopyroxenesa
</1. Sover North; 2, Besterskraal; 3, Sover Mine; 4, New Elands; 5, Swartruggens; 6, Silvery Home; 7, Blaauwbosch; 8, Makganyene; 9, Lace; 10, Voorspoed. FeOT = total iron calculated as FeO. Fe20J and FeO calculated by the method of Droop (1987). End member molecules calculated assuming all Na20 is present as aegirine.
in Ti02 (typically 0.1-2.0 wt%), Ah03 (0.02-1.3 wt %), and Cr203 (typically 0-1.0 wt %). Intra- and inter-intrusion compositional variation is very limited and most pyroxenes are not zoned.
Figures 2.63 and 2.64 illustrate the compositional range of pyroxenes in representative intrusions containing zonation-free pyroxenes. The figure demonstrates that inter-intrusion compositional variation is not significant. The majority of pyroxenes are diopsides containing between 2 and 6 atomic % Fe and more than 47 atomic % Ca. Mg-rich augite is found only in pyroxenes from Blaauwbosch. Note that the older
168 CHAPTER 2
Table 2.19. Representative Compositions of Relatively Iron Rich Pyroxenesa
al_5. Sover North; 6--9. Postmasburg; 10. Voorspoed. FeOr = total iron expressed as FeO. Fe203 and FeO calculated by the method of Droop (1987). End-member molecules calculated assuming all Na20 is present as aegirine. CATS = CaAlSiAlO6. n.d. = not detected.
Swartruggens intrusion does not contain pyroxenes significantly different to those in the younger orangeite intrusions.
Figure 2.65 shows that microphenocrystal pyroxenes from different facies of the Finsch intrusion are very similar in composition. Pyroxenes found in the different dikes of the Swartruggens system are also identical in their composition. These observations suggest that pyroxene compositions are of little use in determining the sequence of intrusion of different batches of orangeite magma in complex bodies.
MINERALOGY OF ORANGEITES
Figure 2.63. Compositions of microphenocrystal clinopyroxenes in diverse orangeites plotted in the ternary system Ca-Mg-Fe (atomic). All data this work except Silvery Home (Robey 1981).
46
48
48
169
- Fe
-SWARTRUGGENS
~_.L. .. 'r .,. I I I I ~
-- - LACE
- -- --", , --I.
VOORSPOED -
-- SILVERY HOME
Fe
The low Al content of the diopsides commonly results in there being insufficient Al to fill the tetrahedral sites in the pyroxene structure; thus, Si + Al is commonly less than 2.0 afu. The tetrahedral site deficiency of 0.0 15-0.035 afu may be remedied by the entry of Fe3+ to this site, as typically Fe3+ is greater than, and Ti less than, the site deficiency. The low Al and Cr contents also typically result in Na being greater than Al + Cr, i.e., the pyroxenes contain an aegirine component rather than kosmochlor or jadeite.
Significant compositional zonation is found only in pyroxenes in the more evolved varieties of orangeite, e.g., Sover North, Postmasburg, Makganyene, and is toward Fe and Na enrichment at the margins of the crystals (Table 2.19, Figure 2.66). This trend culminates in the formation of aegirine-rich mantles (see below).
49 1 48
47 () 46
(j 45
/ 44 43
••
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•• .. I ,
•
••
• •
-Fe ..
• • • ·t' Co ,~ ~ •
D • • • • • •
Mg Fe
BLAAUWBOSCH
NEW ELANDS
Figure 2.M. Compositions of microphenocrystal c1inopyroxenes from the Blaauwbosch (Scott Smith and Skinner unpublished) and New Elands (this work) orangeites plotted in the ternary system Ca-Mg-Fe (atomic).
2
Mg Fe
Fe • 345 6
••• •
7
FINSCH
F2
F4
F7
Figure 2.65. Compositions of microphenocrystal clinopyroxenes in different intrusions within the Finsch orangeite (this work) plotted in the ternary system Ca-Mg-Fe (atomic).
170
MINERAWGY OF ORANGElTES
46
•••• • • • .~. . ~ .. . \ ..
• •
l~l! . .... .. ••• •
• •
171
Fe
• •
POSTMASBURG
MAKGANYENE
SOVER NORTH
Figure 2.66. Compositions of clinopyroxenes from the Postmasburg (this work), Makganyene (this work), and Sover North (this work, Tainton 1992) orangeites plotted in the ternary system Ca-Mg-Fe (atomic).
2.2.2.2. TItanian Aegirine
Figure 2.66 illustrates the pyroxene compositional variation found in the Postmasburg, Makganyene, and Sover North intrusions. Here, pyroxenes are zoned from Fe-poor cores, identical in composition to the unzoned diopsides found in other intrusions, to margins which are richer in Fe and Na. This compositional trend reflects aegirine enrichment (Table 2.19) and culminates in the formation of discrete aegirine mantles (Figure 2.67). Similar compositional trends have been reported by Tainton (1992) from the Pniel occurrence, although aegirines in these rocks are relatively poor in Ti02 «1 wt%).
Table 2.20 gives representative compositions of mantling and groundmass aegirinerich pyroxenes. TheirTi02 contents range from 0.3 to 6.6 wt %. The more Ti-rich varieties may be regarded as titanian aegirine. Recalculation of the compositions on a stoichiometric basis indicates low-to-zero Fe3+ and high Fe2+ contents. Calculation of potential end-member pyroxene molecules suggests that they are members of an unusual quaternary solid solution between Na(Fe2+o.s,Tio.s}Sh06, Na(Mgo.s,Tio.s}Sh06, diopside, and aegirine (Table 2.20). In these pyroxenes all Fe2+ is considered to be combined with Na
172
Oi
90 -# I -,
Ae
80 1 Ae
~!7oi ~ ~ 60 Oi Hd
0\0
~ . .$ 50
Hd
/ 40 ) • POSTMASBURG + SOVER NORTH o MAKGANYENE • VOORSPOEO
10
co. • • o
o + :a+ .ut+ o~· .. ~·~
20 30 Hd
CHAPTER 2
Figure 2.67. Compositions of evolved aegirine-rich clinopyroxenes from the Posttnasburg, Sover North, Makganyene, and Voorspoed orangeites plotted (mol %) in the ternary system diopside (Di)-hedenbergite (Hd)-aegirine (Ae). All data this work.
in the Na(Fe2+o.5,Tio.5)Sh06 molecule. Consequently, the pyroxenes are considered not to contain the hedenbergite molecule. This recalculation scheme is considered the most plausible as it is the only one which results in the number of cations not assigned to a pyroxene molecule being less than 10% of the total cations present, i.e., <0.4 afu. Other recalculation schemes, including the hedenbergite molecule, result in a significant excess of unassigned Na and Ti.
2.2.2.3. MilWr Elements
Table 2.21 summarizes the minor element variation exhibited by microphenocrystal clinopyroxenes. Each intrusion is characterized by pyroxenes of similar compositional range, and only minor inter-intrusional differences exist, e.g., Voorspoed and Lace contain pyroxenes which are slightly lower in Ti, AI, and Cr, relative to other intrusions. The Swartruggens pyroxenes are notably very low in AI.
MINERALOGY OFORANGEITES 173
Table 2.20. Representative Compositions of Aegirine and Titanian AegirineQ
°1_2, Voorspoed; 3, Makganyene; 4-5, Sover Nonh; 6-7, Postmasburg; 8-9, Pniel. Fe:z03 and FeO calculated by the method of Droop (1987). NaFeTIPX = NaFeo.sTIo.sSh06; NaMnTIPX = NaMno.sTIo.sSi:z06;NaMgTIPX = NaMgo.sTIo.sSh06-
Figures 2.68 and 2.69 illustrate Ti versus Al contents. Well-defined compositional trends are absent. A weak correlation between increasing Ti and Al is evident at New Elands and Voorspoed, but typically Ti is independent of AI. Zoning, when present, is very weak and follows a trend of increasing Ti and AI. The Ti and Al contents are typically <0.07 and <0.05 (commonly <0.03) afu, respectively. Figure 2.69 shows that pyroxenes from the Finsch F4 intrusion fall into two groups. Those from the low-AI group have significantly lower Al contents than pyroxenes in the F2 and F7 intrusions, while those from the group with relatively high Al content are identical in composition to pyroxenes from other Finsch intrusions. Pyroxenes from the various dikes occurring at Swartruggens are identical in their minor element compositions.
Cr contents vary widely (Table 2.21) and show no correlation with Al or Ti. Cr is not detectable by electron microprobe in the more-evolved relatively Fe-rich pyroxenes.
The trace element contents of microphenocrystal pyroxenes have not been determined.
174 CBAPTER2
Table 2.21. Minor Element Content of Microphenocrystal Clinopyroxenea
"Data sources: I, this work; 2, Scott Smith and Skinner (pers. comm.); 3, Mitchell and Meyer (1989b); 4, Dawson et al. (1977); 5, Robey (1981); 6, Tainton (1992). N = number of analyses. n.d. = not detected.
0'06 -
-0'05 -
t 0'04 ~ -
I- 0'03-
I 0'02~ -
0'01-
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0'005
+
• SWARTRUGGENS
• BLAAUWBOSCH + NEW ELAN OS
+ •
• + •
I I I I I I I I I I I I I I I I I I
0'010 0'015 0'020
Figure 2.68. Al versus Ti (atomic) compositional variation of clinopyroxenes from the Swartruggens (this work), Blaauwbosch (Scott Smith and Skinner unpublished), and New Elands (this work) orangeites.
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7
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Fig
ure
2.69
. A
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itch
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176 CHAPTER 2
2.2.3. Pyroxenes in the Swartruggens Male Lamprophyre
This dike contains up to 40 vol %, euhedral-to-acicular (0.4 mm long, length-to-breadth ratios 8: 1) diopsides (Dawson et al. 1977). The pyroxenes are strongly-zoned toward Fe and Na-enriched margins (Table 2.22). Figure 2.70 shows that the relatively unevolved pyroxenes, constituting the bulk of the population, have compositions similar to those of unzoned pyroxenes in the Swartruggens orangeites (Figure 2.63). Figure 2.71 shows that the Male lamprophyre pyroxenes exhibit a zoning trend toward enrichment in
Table 2.22. Representative Compositions of Pyroxenes in the Swartruggens Male Lamprophyre Dikea
aFeOr = total Fe calculated as FeO. Ft:203 and FeO calculated by the method of Droop (1987). End-member molecules calculated assuming all Na20 is present as aegirine. CATS = CaAlSiAl06.
MINERALOGY OF ORANGEITES 177
Co -- Fe •
5 10 15 20
• I • Fe •
• •
L Hg Fe
•
Hg
Figure 2.70. Compositions of clinopyroxenes in the Swartruggens Male lamprophyre plotted in the ternary system Ca-Mg-Fe.
hedenbergite. The trend is similar to those found in some orangeites but is displaced from the main trend of evolution of orangeite pyroxenes.
The Male lamprophyre pyroxenes are not significantly enriched in minor elements (Table 2.21) and are weakly-zoned toward increasing Ti and Al from core to margin. They contain significantly more Al than Swartruggens orangeite pyroxenes (Table 2.18).
2.2.4. Megacrystal Pyroxenes
Unlike archetypal kimberlites, orangeites do not characteristically contain a suite of subca1cic diopside megaimacrocrysts. Chrome-poor clinopyroxene macrocrysts have been found in heavy mineral concentrates from Wielpan South and Silvery Home (Skinneret al. 1994) and have also been noted from Lace by Bell and Read (pers. comm.).
~ 0\0
~. 20
~
/ 10
• ~1
If! "I
• ~~I
~~ /. / ,: .
", . / •
•• I
Ae
~ Di Hd
•
DiL----T~~,-----r----,------~
10 20 30 40
Mol. % Hd ---.
Hd
Figure 2.71. Compositions of clinopyroxenes in the Swartruggens Male lamprophyre plotted (mol %) in the ternary system diopside (Di)-hedenbergite (Hd)-aegirine (Ae). Main orangeite trend from Figure 2.67.
178 CHAPTER 2
Many of the Cr-rich megacrysts reported from the Barldy West dikes (Hops et al. 1992) are probably derived from eclogites (Moore and Gurney 1991). Compositional data for megacrystal pyroxenes have not been published.
2.2.5. Comparison with Pyroxenes in Kimberlites
Bona fide primary pyroxenes are considered not to occur in archetypal hypabyssal kimberlites. Pyroxenes commonly occur as reaction mantles around crustal xenoliths, and such pseudoprimary pyroxenes crystallize only from magmas which have been contaminated by such xenolithic material. Their mode of formation is similar to that advocated by Scott Smith et al. (1983) for pseudoprimary pectolite. The composition of the very fine grained pyroxenes occurring in groundmass of diatreme facies rocks has not been determined.
Pseudoprimary phenocrystal and/or groundmass pyroxenes from the Premier type 3 (Scott and Skinner 1979, six analyses) and Orroroo (Scott Smith et al. 1984, three analyses) hypabyssal kimberlites are similar in their major element composition (Table 2.23, Figure 2.72) to the least Fe-rich pyroxenes in orangeites. They differ in being relatively poor in AI, Ti, and Cr (Table 2.23). The few data for pyroxenes (Table 2.23, Figure 2.72) from the Schuller pipe (Scott and Skinner 1979, three analyses) in contrast are relatively Fe rich and variable in their minor element content relative to orangeite pyroxenes. Diopside (mg = 0.92) in the Aries kimberlite (Edwards et al. 1992) is very poor in Ah03, Cr203, and Na20 «0.1 wt %) and Ti02 (0.15 wt %). These small (5-u0 ~m) pyroxenes of irregular habit are intergrown with serpentine and talc and are probably secondary in origin.
Microphenocrystal pyroxenes in the Zagodochnaya kimberlite (Egorov et al. 1991; Table 2.23) are low in Ah03 (O.3-Do4 wt %), Ti02 (0.1 wt %), Cr203 (0.6-1.1 wt %), and FeO (2.4-304 wt %). Other AI- and Cr-rich pyroxenes in this kimberlite appear to be derived from disaggregated upper mantle xenoliths.
Low-Cr, AI-poor pyroxenes in mica-rich xenoliths occurring in the East Udachnaya kimberlite (Egorov et al. 1986) are of similar composition to those of orangeites and
Table 2.23. Compositions of Pyroxenes in KimberJitea
"FeO,- = total Fe expressed as FeO; n.d. = not detected. Compositions 1-2, Premier (Scott and Skinner 1979); 3-4 ,Orroroo (Scott Smith el al. 1984); 5-6, Schuller (Scott and Skinner 1979); 7, Zagodochnaya (Egorov el al. 1991); 8, Udachnaya (Egorov el al. 1986).
MINERALOGY OF ORANGElTES 179
Co -Fe~
-+-L-L-L-L-L-L-L-L--L--L--L-L-L-L-L-I~ Fe
40 LAMPROITES
Mg
Figure 2.72. Compositions of clinopyroxenes from orangeites compared with those of clinopyroxenes in kimberlites (Orroroo, Scott Smith et al. 1984; Premier and Schuller, Scott and Skinner 1979), minettes (Agathla, Jones and Smith 1983; Devon, Jones and Smith 1985; Colima, Luhr and Carmichael1981; Allen and Cannichael 1984; Wallace and Carmichael 1989; Linhaisai, Bergman et al. 1988), and lamproites (Mitchell and Bergman 1991).
lamproites (Table 2.23). The xenoliths are petrographically unlike orangeites and appear to be phlogopite pyroxenite cumulates.
Pyroxenes are relatively common in the atypical, phlogopite kimberlites from West Greenland (Scott 1981). These are very different to orangeite pyroxenes, being relatively rich in Ti02 (0.9-2.8 wt %, Scott 1981; 0.1-4.6 wt%, this work) and FeOr (4.0--6.1 wt %, Scott 1981; 1.7-8.2 wt %, this work). Their high Alz03 contents (0.9-4.5 wt %, Scott 1981; 0.3-10.1 wt %, this work) suggests an affinity with ultramafic lamprophyres such as alnoites or aillikites (see 2.2.6).
2.2.6. Comparisons with Pyroxenes in Lamproites
The major element composition of microphenocrystal pyroxenes in lamproites (Mitchell and Bergman 1991) is identical to that of orangeite pyroxenes (Table 2.24, Figure 2.72). Lamproite pyroxenes typically are not zoned, although examples from the Leucite Hills, in rare instances, may be zoned to Fe-rich margins and exhibit acmitic mantles. However, Fe enrichment does not exceed 10 wt % FeOr and titanian aegirine is not found.
Figures 2.69 and 2.73 show that most lamproite microphenocrystal pyroxenes have very low «0.03 afu) and Ti «0.07 afu) contents. The majority of orangeite pyroxenes are of identical composition, although the compositional range for Al extends to 0.05 afu.
180 CHAPTER 2
Table 2.24. Representative Compositions of Pyroxenes from Lamproitesa
"FeOr = total Fe expressed as FeO; n.d. = not detected. Compositions 1-2, West Kimberley (Mitchell and Bergman 1991); 3-4, Leucite Hills (Mitchell and Bergman 1991); 5-6, Smoky Butte (Mitchell el al. 1987); 7-8, Prairie Creek (Scott Smith et al. 1984); 9--10, Kapamba (Scott Smith el al. 1989).
Only the Kapamba lamproite pyroxenes exceed these Al contents (Figure 2.73). In common with orangeite pyroxene there is insufficient Al present to occupy all of the tetrahedral sites in the structure.
2.2.7. Comparisons with Pyroxenes in Ultramafic Lamprophyres
Ultramafic lamprophyres typically contain highly aluminous pyroxenes, e.g., Wauboukigou Province (Illinois-Kentucky), 1.0-2.0 wt %; Alna, 0.5-10.0 wt %; Oka,
0·10 ORANGEITES
LAMPROITES , ., ", ,
t 0'08 , ", ., .,
", .,
", , 0'06
......
0'04 ROMAN PROVINCE
0'02 --- --- ---- ---
0
0 0'10 0'20 0'30 0'40
-- AI • Figure 2.73. AI versus Ti (atomic) compositional fields of clinopyroxenes from orangeites (this work), di verse lamproites (Mitchell and Bergman 1991), Kapamba 1amproites (Scott Smith et al. 1989), Roman Province lavas and Ugandan kamafugites (Mitchell and Bergman 1991).
MINERALOGY OF ORANGElTES 181
6.3-12.1 wt % Ah03 (Mitchell, unpublished data). Ti02 contents (wt %) are also commonly high, e.g., Wauboukigou, 0.8-S.2 wt %; Alno, 0.9-S.9; Como, 1-2; Oka, 2.O--S.4; Haystack Butte, 1.6-2.5; Polzen, I.S-S.l (Mitchell unpublished data). The combination of high Al coupled with high Ti indicates that these pyroxenes are members of the solid solution series between CaMgSh06 (diopside )-CaAlAISi06 (Ca-Tschermak's pyroxene) and CaTiAh06. Such pyroxenes typically contain Alvi and are thus unlike pyroxenes found in orangeites or lamproites.
The diamond-bearing Bulljah Pool ultramafic lamprophyres contain pyroxenes with relatively low Ah03 «1.0 wt %) and high Ti02 (1-2 wt %) contents (Hamilton and Rock 1990). Late-stage aegirine-augite and aegirine are present in "alteration assemblages" in these rocks. Unfortunately, compositional data are not provided by Hamilton and Rock (1990), and they cannot be compared with late-stage titanian aegirines found in some orangeites.
Primitive olivine-mica-sanidine lamprophyre dikes from Mt. Bundey, Australia (Sheppard and Taylor 1992), have petrographic affinities with lamproites and evolved orangeites. Pyroxenes in these rocks differ from orangeite pyroxenes in being oscillatoryzoned, in containing exsolved orthopyroxene and having high-Ah03 (1.3-3.2 wt %) contents.
Primary pyroxenes in the Bow Hill lamprophyres (Fielding and Jaques 1989) are not significantly different from those of orangeites.
2.2.S. Comparisons with Pyroxenes from Minettes
Figure 2.72 shows that pyroxenes from orangeites and minettes (Luhr and Carmichael 1981, Jones and Smith 1983, 1985, Allen and Carmichael 1984, LeCheminant etal. 1987, Bergman etal. 1988, O'Brien etal. 1988, Wallace and Carmichael 1989) cannot be distinguished on the basis of their major element composition. However, pyroxenes in minettes are commonly enriched in Al relative those of orangeites (Figure 2.73). Although some pyroxenes of low Ah03 content «1.0 wt %) may be found in minettes, the majority contain from 1.0 to 8.0 wt % Ah03. Their compositional evolution toward increasing AI, from core to margin, serves to distinguish them from orangeite pyroxenes. In addition, many pyroxenes in minettes are complexly-zoned and mantled (LeCheminant et al. 1987, O'Brien et al. 1988, MacDonald et al. 1992).
2.3. OLIVINE
2.3.1. Paragenesis
Olivine is a ubiquitous constituent of diatreme and hypabyssal facies orangeites. Olivines in diatreme facies rocks have not been extensively studied as only macrocrysts appear to be present. These are characteristically completely pseudomorphed by serpentine.
In hypabyssal facies orangeites, olivine is typically present as discrete grains which exhibit a wide range in size and modal abundance. The larger grains are typically rounded, whereas the smaller crystals may be either rounded or subhedral to euhedral. These observations suggest that, in common with archetypal kimberlites, two groups of olivine
182 CHAPrER2
are present: rounded macrocrysts which are primarily xenocrysts, and subhedraVeuhedral phenocrysts/microphenocrysts. Rounded microcrysts are considered to be a part of the xenocrystal suite.
It is not always possible to distinguish accurately between olivine xenocrysts and phenocrysts. Skinner (1989) has suggested that all grains which exhibit defonnation features (undulose, extinction, kink banding, recrystallization) or contain inciusions of other lherzolitic-suite minerals are xenocrysts. All strain-free crystals exhibiting planar crystal faces are believed to be phenocrysts. However, there is no a priori reason why all xenocrysts should be strained, and Mitchell (1986) considers that deformation-free rounded macrocrysts may be either xenocrysts or resorbed phenocrysts. Unfortunately, there are no simple compositional parameters which might permit discrimination between these two groups. Moore (1988) has noted that olivines in orangeites from Newlands, Sover, and Bellsbank very rarely show any evidence of undulose extinction.
In orangeites, rounded macrocrysts may constitute up to 50 vol % of the rock (Bosch 1971, Skinner and Clement 1979, Fraser 1987). However, modal abundances vary widely within and between intrusions as a consequence of flow differentiation (Tainton 1992). On average some orangeites, i.e., Bellsbank, Sover, appear to be enriched in macrocrystal olivine relative to others, e.g., New Elands, Swartruggens. Evolved orangeites are relatively poor in olivine. Macrocrysts may be fresh, partially or completely serpentinized, and/or carbonatized.
Some macrocrysts in the Bellsbank, Blaauwbosch, Roberts Victor, and Sanddrift orangeites exhibit pronounced overgrowths of optically and compositionally distinctive olivine which may be up to 0.5 mm in width (Skinner 1989, Tainton 1992). These overgrowths may be nonnally- or reversely-zoned relative to the cores (see 2.3.2).
Phenocrysts are considered to be euhedral-to-subhedral crystals, commonly 0.1 to 2 mm in longest dimension, although the majority are less than 0.5 mm (Skinner 1989). Abundances vary widely within and between intrusions, i.e., from 8 to 23 vol % at Finsch (Fraser 1987). The ratios of xenocrysts to phenocrysts typically range between 1: 1 and 5: 1 according to Skinner (1989), although he notes that higher values may also be found. The extreme variation in xenocryst/phenocryst ratios is exemplified by the Bellsbank intrusions. Here, Bosch (1971) did not observe any phenocrysts, whereas Tainton (1992), using the recognition criteria of Skinner (1989), reports that they are common. The differences may be real or simply reflect the difficulties in distinguishing the two populations.
Phenocrysts may be fresh, partially or completely altered to serpentine, and/or replaced by carbonate. Chromite is the only inclusion recognized in these olivines by Skinner (1989).
Skinnner (1989) has noted that many ofthe larger phenocrysts in the Blaauwbosch, Roberts Victor, and Sanddrift orangeites exhibit olivine overgrowths similar to those observed occurring on coexisting xenocrysts.
Olivines in evolved orangeites exhibit many of the characteristics described above. They differ in containing macrocrysts and "phenocrysts" which are mantled by parallel aggregates (dog's tooth habit) of prismatic olivine. Examples have been described from Postmasburg, Sover North (SN1), and the Bellsbank West Fissures by Tainton (1992) and the Prieska region by Skinner et al. (1994). Olivines of this morphology are considered
MINERALOGY OF ORANGElTES 183
to be either primary growth features resulting from the rapid crystallization of primary olivine on macrocrystal nuclei (Mitchell and Bergman 1991) or an imposed morphology produced during crystallographically controlled resorption of macrocrysts (Scott Smith et al. 1989). Olivine of this habit is common in olivine lamproites (Mitchell and Bergman 1991) but has not been observed in archetypal kimberlites.
Macrocrystal olivine mantled by phlogopite, optically and compositionally similar to groundmass phlogopite is common in evolved orangeites from Pniel, Sover North, and Besterskraal (this work, Tainton 1992). Such reaction mantles are not observed in unevolved orangeites.
2.3.2. Composition
Compositional data for phenocrystal (Table 2.25) and macrocrystal olivines have been provided by Skinner (1989), Moore (1988), Fraser (1987), and Tainton (1992) and for macrocrystal olivines by Mitchell and Meyer (1989a). Studies of olivine compositional variation are hindered by the common alteration and pseudomorphing of olivines. Commonly, small microphenocrystal (groundmass) olivines are completely pseudomorphed by serpentine and/or calcite. Alteration at crystal margins is a particular impediment to determining the compositional evolution of olivines. As a consequence the majority of the compositional data available are for the cores of crystals.
Skinner (1989) and Moore (1988), in general studies of orangeite olivines, have shown that, although indi vidual crystals are of uniform composition, there is considerable intergrain compositional variation. Typically, phenocrysts and xenocrysts of different composition are juxtaposed, suggesting that the assemblage results from the mixing of olivines derived from several sources. Figure 2.74 shows that the cores of the phenocrysts and xenocrysts studied by Skinner (1989) range in mg from 0.87 to 0.95 [mg = Mg/(Fer + Mg)]. However, the mode for the xenocrysts lies at a slightly more magnesian composition (mg = 0.93) than that of the phenocrysts (mg = 0.91-0.92) and is similar to the mean composition of olivine in peridotite xenoliths. Similar conclusions may be drawn from
Table 2.25. Representative Compositions of Olivine Phenocrystsa
"FeOr = total Fe expressed as FeO; n.d. = not detectable. 1, Bellsbank (Tainton 1992); 2, Sover (Tainton 1992); 3-4, Finsch (Scott Smith pers. comm.); 5, Roberts Victor (this work).
Figure 2.74. Histograms of olivine compositions. (A, B) Phenocrysts in diverse orangeites (data from Skinner 1989. Tainton 1992, respectively). (C) Phenocrysts in the Finsch orangeite (Fraser (1987). (0, E) Xenocrysts and macrocrysts in diverse orangeites (data from Skinner 1989, Tainton 1992, respectively). (F) Macrocrysts in the Finsch orangeite (Fraser 1987). Dashed histogram in D-Frepresents compositions ofolivines in peridotite xenoliths (Clement 1982).
the limited data presented by Tainton (1992) and Moore (1988) for the Postmasburg, Pniel. Sover, Bellsbank, and Newlands orangeites and by Fraser (1987) for olivines in the Finsch occurrence (Figure 2.74).
Skinner (1989) has noted that overgrowths on xenocrysts or phenocrysts are of similar composition in different orangeites. They are weakly-zoned from mg = 0.91 adjacent to the core to mg = 0.90 at the margin. Overgrowths are of similar character regardless of the composition of the core. Thus, reverse or normal compositional mantling is present (Figure 2.75).
Nickel is the only minor element present in the olivines in significant quantities (approx. 0.25~.5 wt % NiO; Table 2.25). There are typically no differences evident in the Ni contents of the cores of phenocrysts and xenocrystslmacrocrysts. Considerable intergrain compositional variation is present, although individual crystals are typically of uniform composition.
Skinner (1989) demonstrated that significant systematic variations in NiO content are present in overgrowths. Figure 2.75 shows that NiO contents, which may be relatively higher than those of the cores, rapidly decrease with decreasing mg toward the overgrowth margin. Clearly, the crystallization of overgrowth olivines, together with phenocrysts of similar composition, rapidly depletes the parental magma in Ni. Moore (1988) has noted that the rims of microphenocrystal olivines at Bellsbank and Newlands are depleted in NiO relative to the cores. Calcium contents (Table 2.25) range from not detectable by electron microprobe to as high as 0.l3 wt % CaO at the margins of phenocrysts, regardless of the presence or absence of overgrowths. The latter may be slightly enriched in CaO relative to the cores, but distinct core-mantle compositional discontinuities are not present (Figure 2.75). MnO contents range from approximately 0.01 to 0.2 wt %. The margins of crystals may be slightly enriched in Mn relative to the cores.
2.3.3. Comparisons with Olivines in Kimberlites
The paragenesis and composition of olivine in archetypal kimberlites has been reviewed in detail by Mitchell (1986). Olivine is the most common and most characteristic mineral of kimberlite and occurs as rounded, mantle-derived macrocrysts and euhedral-tosubhedral microphenocrysts. The macrocrysts are of two varieties, xenocrysts derived from disaggregated lherzolites and harzburgites, and rare relatively Fe-rich olivines of cryptogenic origin which are members of the discrete nodule suite.
186 CHAPTER 2
Olivine macrocrysts/xenocrysts in kimberlites cover a wide range in composition from approximately F084 to F095. Individual crystals are homogeneous and macrocrysts of very different composition may occur adjacent to each other. Regardless of composition, macrocrysts may possess thin mantles of olivine compositionally similar to euhedral groundmass olivine (F087-89). Macrocryst compositions are similar to those of olivines in mantle-derived lherzolites and harzburgites. Olivines of this paragenesis are identical in composition and origin to the macrocrystal olivine suite of orangeites.
Microphenocrystal (or groundmass) olivines in kimberlites on a worldwide basis also exhibit a wide range in composition (F087-93). The margins of crystals commonly have narrow rims of olivine which may be Fe-rich or poor, relative to the core. Regardless of whether the zoning is normal or reversed, mantle compositions converge upon F087-89.
The range in core compositions and the uniform composition of the mantles has been interpreted by MitcheU (1986) to imply that the groundmass olivine population is the result of mixing of several batches of magma containing phenocrysts (and macrocrysts) of slightly different composition. The normal and reversed zoned mantles found upon the microphenocrysts (and macrocrysts) reflect attempts by the crystals to equilibrate with the final hybrid magma precipitating olivines in the compositional range F087-90.
KIMBER LITES 40 ORANGEITES
30
N
20
10
87 91 95
Mg / ( Mg + Fe~+) •
Figure 2.76. Histograms comparing the compositions of phenocrystaJ oli vines in kimberlites (Shee 1986, Clement 1982) and orangeites (Skinner 1989).
MINERAWGY OF ORANGEITES 187
Figure 2.76 shows that microphenocrystal olivines in South African kimberlites (Clement 1982, Shee 1985) have a similar range in composition to phenocrysts in orangeites (Skinner 1989). A significant difference is that the mode for orangeites (F091-93) lies at more magnesian compositions than that of the majority of kimberlite olivines (F088-90). Unfortunately, the data base is limited and the modes are not statistically significant. Nevertheless, Skinner (1989) has interpreted the data to suggest that orangeites crystallize microphenocrystal olivines that are slightly more magnesian than those of kimberlites. No significant differences are evident between Ni, Ca, and Mn contents of orangeite and kimberlite phenocrysts.
The group of relatively Fe-rich (F078-88) olivine macrocrysts found in some kimberlites, e.g., Monastery (Gurney et al. 1979), Letseng-la-terae (Dawson et al. 1981), has not yet been recognized in orangeites or in many other archetypal kimberlites. Many of these macrocrysts contain inclusions of Cr-poor, Ti-pyrope or magnesian ilmenite and are clearly a part of the megacryst or discrete nodule suite. The absence of such Fe-rich olivines in orangeites is not unexpected given that the minerals of the megacryst suite are not characteristic of orangeites.
In summary, orangeites and kimberlites are essentially identical with respect to the paragenesis and composition of olivine. Orangeites share with kimberlites the characteristic that, despite the ubiquity of olivine, it exhibits extremely limited compositional variation. Primary olivines richer in iron than Foss have never been formed. The limited range in composition is in marked contrast to that of olivine in many other ultrabasic mantle-derived rocks, e.g., melilitites, alnoites, which display much greater ranges in olivine composition during their crystallization. Hence, olivine compositions cannot be used to assess the degree of evolution of different batches of orangeite magma in composite dike systems or multiple intrusions.
Juxtaposition of macrocrystal and phenocrystal olivines of different core compositions, which have attempted to equilibrate with the magma crystallizing compositionally distinct groundmass microphenocrystal olivines, demonstrates that the olivine assemblage in both orangeites and kimberlites has multiple sources. This hybrid assemblage is formed by the mixing of magmas containing compositionally distinct suites of macrocrysts and phenocrysts. The apparently slightly more magnesian composition of microphenocrystal olivines in orangeites, relative to those in kimberlites, may imply that their parental magmas were slightly more magnesian than kimberlite-forming magmas. This may reflect source characteristics and/or the degree of assimilation of mantle-derived xenocrystal material.
Recognizing that orangeites are hybrid rocks and the majority of the olivine macrocrysts are xenocrysts has important consequences with regard to geochemical studies of orangeites. Clearly, whole rock abundances of major elements and compatible trace elements (Ni, Co, Sc) hosted by olivine will not reflect those of the parental magma (see 3.3.2).
2.3.4. Comparisons with Olivines in Lamproites
Although olivine is a common constituent of lamproites, it is neither ubiquitous nor a characteristic mineral. Olivine contents are highly variable and reach maximum modal
188 CHAPTER 2
amounts (30-40 vol %) only in olivine lamproites. Two texturaly distinct varieties of olivine are found in most olivine-bearing lamproites: anhedral macrocrysts or xenocrysts and subhedral-to-euhedral phenocrysts. Olivines exhibiting "dog's tooth" habit are common in lamproites. The compositions of the xenocrystal olivines are identical to similar olivines in both orangeites and kimberlite and are considered to be derived by the disaggregation of similar mantle-derived ultramafic xenoliths. Phenocrysts in lamproites have similar compositions to those of orangeites, consequently, distinction between these rocks cannot be made on the basis of olivine compositions.
The textures of unevolved macrocrystal orangeites (Figures 1.26-1.44) are significantly different from those of olivine lamproites (see illustrations of the latter in Mitchell and Bergman 1991 and Jaques et at. 1986). Evolved orangeites, in contrast, have some similarities to olivine lamproites in containing dog's tooth habit olivine and macrocrysts mantled by phlogopite.
2.4. SPINEL
2.4.1. Paragenesis
Primary spinels occur as discrete euhedral-to-subhedral crystals. The majority of the crystals are small, from 0.01 to 0.02 mm. Rocks containing very small «0.00 1-0.Q1 mm) spinels are common, e.g., Swartruggens, New Elands. Typically, such spinels are anhedral and highly corroded. Large primary spinel crystals are very rare.
Although spinels are Ubiquitous, their modal abundances are low. Few accurately determined modes are available. Fraser (1987) determined that different facies of the Finsch orangeites contain from 1.9 to 9.0 vol % opaques (presumably predominantly spinel). Skinner and Clement (1979) give abundances from 1 to 6 vol % for examples from Sydney-on-Vaal, Swartruggens, and Star-Burnes orangeites. Many orangeites, e.g., New Elands, Star, are estimated to contain <1-3 vol % spinel. In such rocks K-Ba titanates (2.5) are common and appear to have replaced spinel as the principal late-stage primary Ti-bearing mineral.
Commonly, spinels are single-phase grains or composite crystals consisting of discrete cores and mantles. Cores may be deep-red and transparent or opaque. Mantles are characteristically opaque. Single-phase spinels typically occur as euhedral-to-subhedral crystals poikilitically enclosed by olivine, phlogopite, and tetraferriphlogopite and, rarely, diopside. Spinels enclosed by primary olivines are commonly euhedral, transparent, red crystals. Spinels classified as late-stage ground mass phases are commonly enclosed by late-stage tetraferriphlogopite, amphibole, sanidine, and carbonates.
Atoll-textured spinels (Figures 1.76, 1.79) are extremely rare. Tainton (1992) noted that rare, relatively large (>0.01 mm) euhedral spinels exhibit an atoll texture in samples from Sover-Doornkloof. However, such textures are not typical of the spinel population in these rocks.
Large (up to 0.1 mm) macrocrystal spinels are rarely encountered in thin sections, although they are common in heavy mineral concentrates. They may be mantled by thin rinds of opaque spinels which are compositionally similar to primary groundmass spinels.
MINERALOGY OF ORANGEITES 189
2.4.2. Composition
Few detailed investigations of the compositional variation exhibited by orangeite spinels have been undertaken. Significant compositional data have been provided only by Boctor and Boyd (1982), Mitchell and Meyer (1989a), and Tainton (1992). The discussion below is based on these data and new data for Lace, Bellsbank, Besterskraal, and Sover-Doomkloof obtained during the preparation of this work.
The compositions of orangeite spinels fall within the eight-component system MgCr204 (magnesiochromite)-FeCr204 (chromite)-MgAh04 (spinel)-FeAh04 (hercynite)-Mg2Ti04 (magnesian ulvospinel or qandilite)-Fe2Ti04 (ulvospinel)-MgF~04 (magnesioferrite)-Fe304. Manganese and zinc contents are typically low, and the MD2Ti04 (manganese ulvospinel), MnFe:!04 Gacobsite), MnCr204 (manganoan chromite), ZnAh04 (gahnite), and ZnCr204 (zincian chromite) end members are usually unimportant. Vanadium contents have not been determined in this or previous studies but are expected to be low given the reasonable weight percent oxide totals reported.
Spinel compositional variations are conventionally represented by plotting compositions in the "oxidized" and "reduced" spinel prisms (Irvine 1965, Haggerty 1976, Mitchell 1986). In the "reduced" spinel prism, total iron is calculated as FeO (Fe?+ or FeOT), whereas in the "oxidized" prism Fe3+ and Fe2+ are estimated from stoichiometry
Thble2.26. Representative Compositions of Spinels from the Bellsbank and Lace OrangeitesQ
"FeOr = total Fe calculated as FeO; n.d. = not detected. 1-5. BeIlsbank; 6-10. Lace. (All data this work except 3 from Boctor and Boyd 1982.)
190 CHAPTER 2
(Droop 1987) or by the recalculation procedure of Carmichael (1967b). Mitchell (1986) noted that the reduced prism is particularly useful for illustrating spinel compositions in that all of the major elements present in kimberlite (and orangeite) spinels are included in the prism. A disadvantage of this projection is that the prism fails to illustrate variations in Fe304 or MgFe204 end members which may be important in evolved spinels. Oxidized prisms have the disadvantage of failing to illustrate the presence of inverse spinels, as Ti is not included in the projection. Consequently, neither prism is entirely satisfactory for illustrating the complete compositional variation of spinels, which may range from magnesian ulvospinel-rnagnesiochromite~hromite solid solutions to ulvospinel-magnetite. Commonly, the reduced spinel prism, together with projections of the data onto the front and bottom prism faces, provides the best illustrations of compositional variation.
Representative compositions of primary spinels are given in Tables 2.26 and 2.27. These fall into two broad compositional groups:
Group A. Titanian (1-5 wt % Ti02) magnesiochrornite~hromite solid solutions characterized by very high Cr/(Cr + AI) ratios (typically >0.9) and moderate
Table 2.27. Representative Compositions of Spinels from the Besterskraal, Sover North, and Pniel Orangeitesa
"Fe(h = total Fe calculated as FeO. n.d. = not detected. 1--Q, Besterskraal (this work); I, euhedral spinel in olivine phenocryst; . 2-4,3-5, cores and mantles respectively; 6, euhedral groundmass spinel; 7-8, Sover North (TaintoD 1992); 9-10, Pniel (Tainton 1992).
MINERAWGY OF ORANGEITES 191
Figure 2.77. Representative compositions of spinels from the Bellsbank, Lace (this work) and Bums (Mitchell and Meyer 1989a) orangeites chosen to illustrate the compositional trends of orangeite spinels in the reduced spinel prism.
Fei+/(FeT2+ + Mg) ratios (0.4-0.7). Ah03 contents are low, typically <5 wt % and commonly <2 wt %.
Group B. Alumina-poor «1 wt % Ab03) titanian (5-10 wt % Ti02) magnesian « 10 wt % MgO) magnetites.
Red and/or opaque group A spinels typically form single-phase crystals. Opaque group B spinels occur as discrete single crystals and as mantles upon group A spinels.
Representative spinels from Bellsbank, Lace, and Burns are plotted in the reduced and oxidized spinel prisms, Figures 2.77 and 2.78 respectively, to illustrate the nature of evolutionary trends of spinel compositions in orangeites.
Group A spinels plot near the base of both prisms and evolve along the prism axis toward increasing Fer/(FeT + Mg) or Fe2+/(Fe2+ + Mg) ratios at approximately constant Cr/(Cr + AI) ratios and Ti contents. Group B spinels plot within the prism and near to the front rectangular face as a consequence of their high Cr/(Cr + AI) ratios, even though Cr contents are low «10 wt % Cr203). They evolve at approximately constant Fei+ /(Fei+ + Mg) ratios (0.8-0.9) or Fe2+ /(Fe2+ + Mg) ratios (0.7-0.8) toward the apices of the prisms. The trend reflects primarily increases in magnetite, magnesian-ulvospinel, aM ulvospinel contents.
192
O·A ........ CI'/, 0·E>
rCI'" "4/) 0.5
..........
CHAPTER 2
Figure 2.78. Representative compositions of spinels from the Bellsbank, Lace (this work) and Bums (Mitchell and Meyer 1989a) orangeites chosen to illustrate the compositional trends of orangeite spinels in the oxidized spinel prism.
These data demonstrate that orangeite spinels evolve from titanian magnesiochromite--chromite solid solutions toward magnesian ulvospinel-ulvospinel-magnetite solid solutions (Tables 2.26 and 2.27). This trend involves increases in Fe2+, Fe3+, and Ti with concomitant decreases in Cr and AI.
The entire compositional trend is not present in all orangeites. Zoning trends within individual crystals conform to the overall trend. The presence of discrete cores and mantles implies that spinel precipitation may not be continuous during post-intrusional crystallization. The compositional hiatus may result from a miscibility gap or be due to bulk compositional effects.
Figures 2.79 and 2.80 illustrate the compositions of spinels from orangeites, characterized by the phlogopite-tetraferriphlogopite mica composition trend (2.1.7), projected onto the bottom and front square faces of the reduced spinel prism. These data clearly show the differing compositions of group A (cores) and group B (mantle) spinels. Only minor inter-intrusion differences are evident. Thus, group B spinels from Lace do not exhibit the same degree of Ti and Fe3+ enrichment as spinels from Burns or Bellsbank. (This may be an artifact due to a paucity of data.) Group B spinels from Newlands may have slightly lower Fe/(Fe + Mg) ratios than other orangeite spinels.
Figures 2.81 and 2.82 illustrate the compositions of spinels from the Sover Mine orangeite and amphibole-bearing orangeites exhibiting the phlogopite-biotite mica com-
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he r
educ
ed s
pine
l pr
ism
.
... ~
("'l ~ ....
MINERALOGY OF ORANGElTES 195
positional trend (2.1.7) from Sover North, Besterskraal, and Pniel. Mantled spinels in the Sover Mine intrusion are identical to other orangeite spinels. Discrete spinels from Sover North correspond mainly to group A spinels. Group B spinels are rare (two analyses; Tainton 1992) and have high Ti02 contents (12.2-13.4 wt %). Their compositions fall within the Sover Mine group B trend.
Besterskraal spinels occur as inclusions in primary olivine and as two-phase crystals. Spinels included in olivine are Fe poor [Fer2+/(Fer2+ + Mg) < 0.5] relative to the cores of discrete, mantled spinels [Fer2+/(Fer2+ + Mg) = 0.55-0.75]. Mantles (group B) do not show extreme enrichment in Ti or Fe3+, although they may have slightly higher Fer2+/(Fer2++ Mg) ratios (>0.9) relative to other orangeite spinels. Only low-Ti group A spinels are present at Pniel.
Figure 2.82 suggests that group A spinels from amphibole-bearing evolved orangeites have slightly lower Ti/(Ti + Cr + AI) ratios than other orangeite group A spinels and that magnetite-rich group B spinels are not present. Further detailed investigations of these spinel populations are required.
The composition of macrocrystal spinels has been insufficiently documented. Mitchell and Meyer (1989a) have shown that spinel macrocrysts from New Elands are Ti02-poor «0.5 wt %) aluminous magnesian chromites exhibiting a wide range in their Cr203 (43.7-62.8 wt %) and Ah03 (6.9-22.7 wt %) contents. Those with high Cr/(Cr + AI) ratios form a continuation of the orangeite group A spinel compositional trend (Figure 2.79).
Mitchell (1986) has noted that Ti-, Fe3+ -poor spinels from kimberlites, mantlederived lherzolites and harzburgites, and mid-ocean ridge basalts, all plot on the base of the reduced spinel prism and have very similar compositions. These spinels have no compositional characteristics permitting their assignment to a particular paragenesis. The New Elands spinel macrocrysts have similar compositions to these Ti-poor spinels; consequently their origins cannot be unambiguously determined.
2.4.3. Comparisons with Kimberlite Spinels
The compositional trends of spinels in kimberlites are well documented. Previous studies did not distinguish between archetypal kimberlites and orangeites, although it was recognized that mica-rich kimberlites did contain spinels different in composition from those in monticellite kimberlites. Mitchell (1986) has shown that three main groups of spinels, each defining a distinct compositional trend in spinel prisms (Figure 2.83) are present:
1. The macrocrystal or aluminous magnesian chromite trend 2. Magmatic trend 1 or the magnesian ulvospinel trend 3. Magmatic trend 2 or the titanomagnetite trend
A fourth group of AI-rich ground mass spinels forming the pleonaste reaction trend is known only from a few kimberlites. These spinels have a multiplicity of origins (Mitchell 1986) and are not considered further in this work.
Spinels belonging to the AMC trend are Ti02 poor «2 wt %) and compositions plot on the base of spinel prisms (Figure 2.83). They range in composition from magnesian
1% CHAPTER 2
Figure 2.83. Representative compositions of spinels from kimberlites plotted in the reduced spinel prism. See text for an explanation of trends I and 2.
aluminous chromite (MAC) to aluminous magnesian chromite (AMC). Individual crystals are typically homogeneous, consequently it is difficult to establish compositional trends. However, Ti-bearing magnesian chromites, similar in composition to the more Cr-rich macrocrysts, mantle the latter spinels and occur as primary groundmass phases. This observation suggests that some AMC macrocrysts may be cognate.
MAC-AMC spinels are similar in composition to spinels occurring in a wide variety of ultrabasic and basic rocks. They do not exhibit any compositional or textural features that permit conclusive recognition of their origin. Thus, Mitchell (1986) suggests that Cr-rich members of the AMC trend are cognate and that AI-rich members are xenocrysts. In contrast, Shee (1984) regards all of the macrocrysts as xenocrysts.
Magmatic trend 1 is now recognized as the characteristic spinel compositional trend of archetypal kimberlites (Mitchell 1986, Mitchell and Bergman 1991). In the reduced spinel prism (Figure 2.83a) the trend is across the prism from the base near the MgCr204-FeCr204join [commonly Cr/(Cr+ AI) = 0.80-0.95, Fei+/(FeT2+ + Mg) = 0.4-0.6] toward the rear rectangular face [Le., decreasing Cr/(Cr + AI) ratios] and upward toward the Mg2Ti04-Fe2Ti04 apex. Spinel evolution is from titanian magnesian aluminous chromite (TIMAC) or titanian magnesian chromite (TMC) containing 1-12 wt % Ti02 toward members of the magnesian ulvospinel-ulvospinel-magnetite (MUM) series and is a trend of increasing Ti, Fe3+/Fe2+, and total Fe, and decreasing Cr at approximately constant F~2+/(F~2+ + Mg) ratios. Alumina may decrease, increase, or remain constant
MINERAWGY OFORANGElTES 197
over this trend. MnO contents are typically low «1 wt %), but may increase slightly in the more evolved spinels. The trend culminates with the formation of Ti- and Mg-free magnetite. The presence ofTi- and Mg-rich spinels (12-23 wt % Ti02, 12-20 wt % MgO) containing substantial proportions of the magnesian ulvospinel molecule (20-40 mol.% Mg2Ti04) is the hallmark of this compositional trend. Archetypal kimberlites are the only igneous rocks yet known to contain spinels rich in magnesian ulvospinel. Representative compositions can be found in Mitchell (1986), Shee (1984), and Pasteris (1983). The entire magmatic trend 1 spinel assemblage is not found in all kimberlites and individual occurrences may exhibit only a portion of the trend. Numerous examples of the trend are described in detail by Mitchell (1986).
Magmatic trend 2 is characterized by spinels ranging in composition from AMC through TMC and titanian chromite (TC) to members of the ulvospinel-magnetite (USP-MT) series. In the reduced spinel prism (Figure 2.83) the trend is initially along the axis ofthe prism toward increasing Fer2+/(Fei+ + Mg) ratios at relatively constant, but low, Ti02 contents and high Cr/(Cr + AI) ratios (>0.85), followed by a rapid increase in Ti at high Fer2+/(Fer2+ + Mg) ratios (>0.8) toward the Fe2Ti04 apex. MnO enrichment (> 1 wt %) occurs in the more evolved spinels. The trend is characterized by rapid MgO depletion, and spinels rich in magnesian ulvospinel (>20 mol %) are not formed. All of the spinels from trend 2 are poor in Al relative to those from trend 1.
Magmatic trend 2 is uncommon in archetypal kimberlites and insufficiently characterized. It has been recognized only from the Tunraq (Mitchell 1979), Zagodochnaya (Rozova et al. 1982), and Koidu (Tompkins and Haggerty 1984) kimberlites. The
Figure 2.84. Compositional trends of spinels from kimberlites (Tl and T2; Mitchell 1986. this work). orangeites (this work). and lamproites (L; Mitchell and Bergman 1991).
1'0
<t + 0'6 ... U
+ i= 0'4
...... i=
0'2
0'2 0'4 0'6 0'8 1'0
2+ (I:' 2+ ) -- FeT / leT + Mg --.
198 CHAPTER 2
common feature of these kimberlites is that they appear to have crystalized abundant phlogopite prior to the formation of the bulk of the ground mass spinels. In the case of the Tunraq kimberlite, spinels belonging to both magmatic trends are present in different facies of the intrusion. However, only the facies rich in macrocrystal phlogopite contains spinels belonging to trend 2.
Very rare spinels belonging to this trend have also been noted in the Elwin Bay (Mitchell 1978a), De Beers Peripheral (Pasteris 1980), and Marushkaya (Rozova et al. 1982) kimberlites.
Figures 2.77 and 2.84 show that orangeite spinels exhibit compositional trends identical to those found in kimberlite magmatic trend 2 (Figure 2.83). Spinels belonging to magmatic trend 1 are not present in orangeites.
Spinels in orangeites are not characterized by the presence of significant quantities (>20 mol %) of the magnesian ulvQspinel molecule (Tables 2.26 and 2.27). Spinels from orangeites containing 10-20 mol % Mg2Ti04 are Cr-rich members of the Mg2Ti04-FeCr204-Fe304 series rather than the Cr-poor Mg2Ti04-Fe2Ti04-Fe304 series characteristic of kimberlites.
2.4.4. Spinel Compositional Variation in Lamproites and Lamprophyres
Spinels of high Crf(Cr + AI) ratios (>0.9) very similar to trend 2 spinels, are also found in lamproites (Mitchell and Bergman 1991). In these rocks spinels evolve from
o.'~
O·A ......... C"/~ 0.6
C,. '1-4/) 0.8
..........
• PRAIRIE CREE K
o ELLENDALE
Figore 2.8S. Representative compositions of spinels from lamproiles plotted in the reduced spinel prism (after Mitchell and Bergman 1991).
MINERALOGY OF ORANGElTES 199
Ti-poor aluminous magnesiochromites, through titanian aluminous magnesiochromites, to AI-poor titanian magnesian chromite, and magnesian titaniferous magnetite. The evolutionary trend is initially along the axis of the spinel prism (Figures 2.85) with increasing Fer2+/(Fer2+ + Mg) ratios. Subsequently, the spinels evolve with increasing Ti/(Ti + Cr + AI) at constant Fer2+/(Fer2+ + Mg).
Mitchell and Bergman (1991) have noted that kimberlite spinel trend 2 and the lamproite spinel trend are Cr-rich variants of a common spinel trend found in a wide variety of rocks that include basalt, alnoites, melilitoids, minettes, and other lampro-
1'0
0·9
0·8
- 0·7
« + 0'6
~
() 0·5 + ~ 0'4 -"'-t- 0'3
0'2
O· I
• ULTRAMAFIC LAMPROPHYRES
o KIMBERLITE TREND 2
+ MINETTES • ORANGEITES
o - - - - - - --j
o
2 4 5 6 7 8 9
Cr I( Cr + AI) 10
Figure 2.86. Ti/(Ti + Cr + AI) versus Cr/(Cr + AI) for spinels from kimberlites (Mitchell 1986, Shee 1984. Pasteris 1980),lamproites (Mitchell and Bergman 1991), orangeites (this work, Tainton 1992, Boctor and Boyd 1982), basalts (Ridley 1977), ultramafic lamprophyres (this work) from Oka, Pol zen, Como, Alno, and Haystack Butte (HB), and minettes (Jones and Smith 1985).
200 CHAPfER2
phyres. The trends are similar in that their evolution is from Ti-poor, Cr-rich spinels toward ulvospinel-magnetite solid solutions. Compositional trends originate at the base of the prisms and evolve along the axes toward the ulvospinel or magnetite apices. The trends differ principally with respect to the Cr/(Cr + AI) ratios of the least-evolved members of the series (Mitchell 1986). Cr-rich spinels in alnoites and polzenites with Cr/(Cr + AI) ratios <0.85 are mantled by aluminous Cr-poor or Cr-free spinels (Figure 2.86). Spinels from different petrological provinces differ with respect to the composition of the least-evolved spinel. All follow an evolutionary trend ofCrdepletion coupled with Ti and Fe enrichment.
Spinel compositions, plotted in Ti/(Ti + Cr + AI) versus Cr/(Cr + AI) diagrams, clearly demonstrate these differences (Figure 2.86). Thus, spinel compositions with Cr/(Cr + AI) > 0.85 are considered to be indicative of orangeites, lamproites, and some varieties of kimberlite. Spinels with Cr/(Cr + AI) < 0.85 are characteristic of basaltoids, melilitoids (including alnoites and polzenites), and minettes. Detailed studies of the compositional variation shown by spinels in minettes have not yet been undertaken. The majority of minettes examined to date (Jones and Smith 1985, Bergman 1987, Peterson 1994, Lange and Carmichael 1991 , Allan and Carmichael 1984) contain ulvospinel-magnetite solid solutions as discrete crystals and as mantles upon earlier-formed MAC-AMC solid solutions.
Mitchell (1986) and Mitchell and Bergman (1991) have previously noted that kimberlite magmatic trend 2 is not diagnostic of a kimberlitic or lamproitic paragenesis. As orangeite spinels are of similar composition, it must be concluded that Cr-rich members [Cr/(Cr + AI) > 0.85] of spinel trends similar to kimberlite magmatic trend 2 are indicative of (1) kimberlites containing phlogopite macrocrysts, (2) lamproites, and (3) orangeites. Exact classification of rocks containing such spinels can only be made in conjunction with other mineralogical and geochemical criteria.
2.S. POTASSIUM BARIUM TITANATES
2.S.1. Hollandite Compounds belonging to the hollandite group have the general formula AI-2BI-2Ti6-7-
016, where A = K, Ba, Rb, Cs, Sr; B = Fe2+, Fe3+, AI, V, Ce, Ga, Sc, In, Ru; and C = Ti, Nb, Ge, Zr, Sn. They consist of paired chains of edge-sharing (B,C)06 octahedra extending along the crystallographic c-axis. The chains are linked by comer- and edge-sharing to form a framework containing continuous tunnels of approximately square cross section aligned along the c-axis. Each tunnel is bounded by four chains of paired octahedra. The tunnels are filled by A-site cations to varying degrees; thus hollandites typically display non-integral stoichiometries. Cation ordering within and between the tunnels results in incommensurate superlattice ordering (Bursill and Grzinic 1980, Pring and Jefferson 1983, Kesson and White 1986, Zandbergen et al. 1987).
2.5.1.1. Paragenesis
Hollandite group minerals are common, but not ubiquitous, late-stage groundmass phases in some orangeites. To date they have been recognized in the Star, Lace,
MINERALOGY OF ORANGElTES 201
Besterskraal, Sover North, and New Elands orangeites (Mitchell and Haggerty 1986, Mitchell and Meyer 1989b; this work). Hollandite does not occur in all samples from a given intrusion and typically occurs in rocks containing the most evolved micas. The factors controlling the presence of hollandite in some rocks but not others have not been established. The apparent absence in many orangeites may simply be a consequence of lack of investigation.
Hollandites typically occur as stellate clusters of subhedral prismatic crystals (Figure 2.87) enclosed within groundmass calcite segregations. Silicate-rich portions of the groundmass typically lack hollandite. Hollandites are opaque in thin section, with reddish internal reflections being observable only at the edges of very thin crystals. In reflected light the mineral is gray and exhibits medium reflection pleochroism and medium-tostrong anisotropy in tones of light-to-dark gray. Estimated white light reflectivities in air are approximately 12-15% and in oil immersion 5% (Mitchell and Haggerty 1986). BSE imagery shows that zoning may be weakly- or strongly-developed (Mitchell and Meyer 1989b).
2.5.1.2. Composition
Minerals belonging to the hollandite structural group form a complex series of solid solutions, the most important of which are between the K2(Mg,Fe2+)TbOI6-Ba(Mg,Fe2+)Ti7016 series and the K2(M3+hTi6016-Ba(M3+hTi6016 series where M3+ = Fe, Cr, V, Ce, AI. The existing nomenclature of these phases is unsatisfactory, as only four of the many possible end-member molecules have been defined as mineral species. These are: BaFe3+2Ti6016, termed "barian priderite" by Zhuravleva et al. (1978); mannardite, (Ba,H20)V2Ti6016 (Scott andPeatfield 1986, Szymanski 1986); redledgeite BaCr2Ti6016 (Gatehouse et al. 1986); and ankangite Bao.83 V 2.29Cro.osTis.83016 (Xiong et al. 1989).
Mannardite, redledgeite, and ankangite have been accepted by the International Mineralogical Association as valid mineral names. Currently, the potassian varieties of these minerals are referred to simply as potassium analogues, e.g., potassian redledgeite, rather than being given distinct names (Mitchell and Meyer 1989b). The status of ankangite as a distinct species is questionable. Wu et al. (1990) have shown this mineral to possess an incommensurately modulated structure, and, although compositionally similar to the V-hexatitanate, tenned mannardite, it is structurally similar to septetitanates such as priderite (see below). Hence, ankangite and mannardite may both be members of a Ba-V hexatitanate polysomatic series for which a multiplicity of names is undesirable.
Although priderite (Norrish 1951) is accepted as a valid name, it is not actually an end-member composition. Most priderites are intermediate members of the K2(Mg,Fe2+)ThOI6-Ba(Mg,Fe2+)ThOI6 series (Fe2+ » Mg) with some solid solution toward BaFe3+2Ti6016. The latter hexatitanate is an end-member molecule for which a new name is desirable as it is not a barian priderite. Ifpriderite is to be retained as a valid mineral species then the tenn must be redefined as an end-member molecule. Mitchell and Meyer (1989b) have suggested that the solid solutions based upon K-Ba septatitanates be tenned the priderite series, but did not redefine any of the end-member molecules as priderite.
Insufficient is known of the role of H20 in hollandites, as most natural hollandites are not analyzed for water. Scott and Peatfield (1986) and Szymanski (1986) consider
202 CHAPTER 2
Figure 2.87. Acicular hollandite in calcite and phlogopite matrix, New Elands. (A) Transmitted light, field of view 0.25 mm. CB) reflected light with partially crossed polars, field of view 0.1 mm.
MINERAWGY OF ORANGElTES 203
that mannardite and redledgeite contain one H20 molecule per formula unit (Le., approx. 2.1 wt % H20). However, as reasonable analytical totals are typically obtained for most electron microprobe analyses of hollandite in lamproites, it is assumed that the amounts of water present in most hoIlandites are very small.
Unfortunately, it is not possible to assess the FeO and Fe203 contents ofhoIlandites analyzed by electron microprobe. This is because the common nonstoichiometry of the compounds precludes estimation ofFe2+ and Fe3+ from the structural formula by standard methods. It should be clearly realized that natural hoIlandites contain both Fe2+ and Fe3+
and may even contain Ti3+ (Myhra et al. 1988). In this work total Fe is expressed as Fe203, and all structural fomulae quoted are based upon 16 oxygens in order that hoIlandites from diverse parageneses may be compared on the same basis. The actual oxygen content and number of cation vacancies may vary substantially, depending upon composition (Myhra et al. 1988, Kesson and White 1986).
Representative compositions of hoIlandites from Lace, Besterskraal, Sover North, Star, and New Elands are given in Tables 2.28 and 2.29. Individual crystals may be strongly-zoned with respect to their Ba, K, Nb, and V contents. Figure 2.88 demonstrates
Table 2.28. Representative Compositions of Hollandites from the Lace and Besterskraal OrangeitesO
A 1.016 1.051 1.012 1.151 1.057 1.206 1.476 1.489 1.528 B 1.279 1.170 0.996 0.988 1.230 1.458 0.874 1.103 0.943 C 6.563 6.608 6.797 6.812 6.589 6.381 6.976 6.821 6.897
"Total Fe expressed as Fe203; n.d. = not detected; n.a. = not analyzed. Compositions 1-3, Sover North (this work); 4-6, Star (Mitchell and Meyer 1989b); 7-9, New Elands (Mitchell and Haggerty 1986).
that each intrusion is characterized by hollandite of a particular composition. The majority of the data plot within or close to the quadrilateral of compositions defined by end-member K-Ba septe- and hexatitanates, suggesting that the solid solutions present are primarily between BaFe3+2Ti6016, BaFe2~i?OI6, and K2Fe2~i?OI6.
Figure 2.89 shows that hollandites differ with respect to their BaO and K20 contents. Those from Lace and Besterskral are Ba rich (Table 2.28) relative to hollandites from Sover North and Star (Table 2.29). Hollandites from New Elands are richest in K20 (Table 2.29). Figures 2.88 and 2.89 demonstrate that the solid solutions present, with the exception of New Elands hollandites, are dominated by Ba-septe- and hexatitanates. The New Elands hollandites are close to K2FeTh016 in composition.
Hollandites from each intrusion differ with respect to their Fe, V, and Nb contents. Figure 2.90 and Tables 2.28 and 2.29 show that hollandites from New Elands, Star, and Lace are enriched in V, relative to those from Sover North and Besterskraal. Samples with
MINERALOGY OF ORANGEITES 205
+ BESTERSKRAAL • SOVER NORTH • LACE o STAR / • NEW ELANDS
~ \
\ K!. TRISKAIDECATITANATES
MgO + M20 3 Ti02 + Nb205
Figure 2.88. Compositions of hollandites. K-triskaidecatitanates. and Ba-pentatitanates from diverse orangeites plotted in the ternary system (K20+BaO)--(Ti02+Nb20S)--(MgO+M203) (wt %). where M = Fe. Cr. V. and AI. Data sources: this work. Mitchell and Meyer (I989b). Mitchell and Haggerty (1986).
Fe3+j(Fe3+ + V) ratios greater than 0.5 may be regarded as mannardite-potassium mannardite solid solutions.
Hollandites typically contain from 0.5 to 3.0 wt % Nb20s, with those from Besterskraal being relatively enriched in NbzOs (2.0-6.8 wt %). Table 2.28 shows that as Nb contents increase, Fe contents decrease. There is no simple negative correlation between Ti and Nb, suggesting that Nb is accommodated at the B- and C-octahedral sites by a complex coupled substitution involving Nbs+, Fe3+, Ti4+, and lattice vacancies. The Nb end member of the solid solutions involved is as yet unidentified but is unlikely to be a hollandite group compound, as the host of barium niobium titanates in the iron-free BaO-Ti02-Nb20S system (Millet et at. 1987) does not belong to this structural group.
Mitchell and Haggerty (1986) have noted that hollandites from New Elands contain 0.7-1.7 wt % Ce203. Other hollandites analyzed during the preparation of this work were found to contain no detectable levels of Ce203 by electron microprobe analysis.
206
I
BoO
• LACE + BESTERSKRAAL • SaVER NORTH o STAR • NEW ELANOS
o 00 0
o o
\
. 'III Field of K-.. TRISKAIOECATITANATES :. saVER NORTH -
STAR - LACE
•
CHAPTER 2
\
Figure 2.89. Compositions of hollandites. K-triskaidecatitanates. and Ba-pentatitanates from diverse orangeites plotted in the ternary system BaO-{Ti02+Nb203+M203)-K20 (wt %), where M = Fe. Cr, V, and AI. Data sources: this work, Mitchell and Meyer (1989b). and Mitchell and Haggerty (1986).
MrnNERALOGYOFORANGEITES
t -c m + ~ -
O·g
0·8
0'7
0'6
0·5
0'4
0'3
0'2
0'1
• LACE o STAR + BESTERSKRAAL • SOVER NORTH • NEW ELANDS
o o
o 0 &
o
0'1 0'2 0'3
BaV2 TiS°H;
o
o
o
• •
o
o
o
•• • • • . :
o
• o •
• •
o
•
o
o o
207
•
•
•
Figure 2.90. Compositions ofhollandites (this work. Mitchell and Meyer 1989b, Mitchell and Haggerty 1986) from diverse orangeites plotted as Fe3+/(Fe3+ + V) versus KI(K + Ba) (atomic). Compositional field of other hollandites from Mitchell and Bergman (1991), Mitchell and Vladykin (1993), and Mitchell (1994c).
2.5.1.3. Comparison with Hollandites from Lamproites, Kimberlites, and Other Potassic Rocks
2.5.1.3.a. Lamproites. The priderite variety of hollandite is one of the typomorphic minerals of the lamproite clan (Mitchell and Bergman 1991). In these rocks it occurs as euhedral prismatic crystals which form after phenocrystal phlogopite and prior to groundmass phlogopite. Typically, strong compositional zoning is not present. Priderites crystallize contemporaneously with silicates and are not characteristically associated with calcite. A single occurrence of priderite within calcite globules has been reported by Jaques et al. (1989a) from the Argy~ olivine lamproite dikes. Priderites may be mantled by late-stage ilmenite or jeppeite.
208 CHAPTER 2
Table 2.30. Representative Compositions of Hollandites from Lamproitesa
"Total Fe expressed as Fe203. n.d. = not detected. Compositions 1-3, Mt. North, West Australia; 4-7, Endlich Hill, Leucite Hills, Wyoming; 8--10, Francis, Utah. (All data this work.)
Mitchell and Bergman (1991) have shown that each lamproite province is characterized by hollandite of a particular KlBa ratio. The solid solutions present in lamproite hollandites are similar to those of hollandites from orangeites. Hence, Table 2.30 and Figure 2.91 show that lamproite and orangeite hollandites do not differ significantly in their major element compositions.
Figure 2.92 shows that orangeite hollandites, with the exception of New Elands, are
typically richer in BaO than lamproite hollandites. BaO-rich hollandites from the Leucite Hills have compositions which overlap those of orangeite hollandites while those from
West Kimberley are significantly enriched in K20.
Significant differences exist with respect to the V and Crcontents of orangeite (Tables 2.28 and 2.29) and lamproite hollandites (Table 2.30). Mitchell and Bergman (1991) have
noted that the V203 contents of lamproite hollandites typically do not exceed 1 wt %,
although Jaques et at. (l989a) have reported priderite with 1.3-1.7 wt % V 203 from Argyle. Data obtained during the preparation of this work suggest that the V 203 contents of iamproite hollandite are unlikely to exceed 2.0 wt % (Table 2.30). In marked contrast to the very low Cr contents of orangeite hollandites, the Cr203 contents of lamproite
Figure 2.91. Comparison of the compositions of hollandites, K-triskaidecatitanates, and Ba-pentatitanates from orangeites (this work), Benfontein and Wesselton kimberlites (Mitchell1994c),lamproites (Mitchell and Bergman 1991, this work), and Murun ultrapotassic syenites (Mitchell and Vladykin 1993). P = Ba-pentatitanates, H = hollandites, T = K-triskaidecatitanates.
hollandite varies from <0.2-7 wt % and rarely reaches 11 wt % (Mitchell and Bergman 1991, this work).
Figure 2.93 illustrates how lamproite hollandites may be distinguished from those of orangeites on the basis of their Cr203 and V 203 contents. Note that because individual hollandites may have low Cr or V contents it is necessary to analyze a suite ofhollandites from a given locality before drawing any conclusions as to their magmatic affinity.
In summary, lamproite hollandites are very similar in composition to those in orangeites. The principal differences are that orangeite hollandites show solid solutions toward the mannardite--potassium mannardite series, while those from lamproites exhibit solid solutions toward the redledgeite--potassium redledgeite series. The hollandites differ significantly in their paragenesis, those in lamproite being relatively early-formed groundmass or microphenocrystal phases, whereas those in orangeites are very late stage groundmass phases intimately associated with carbonates and potassium triskaidecatitanates.
210
BoO \
Ba3TisO'3 / K20 ~ /
I \ • JEPPEITE / \ \ \ \
BoO \
6\ K~ Ti02
I / I / I / I
\ / \ / ,/ I
15 1- 1- 1- 10 ~ ,A~-r. ~-r.
,. IP III., ·0 ~"~,, ~ ~. ~.
~o "'0 .. .. '-.,-J
K - HOLLANDITES
85
CHAPTER 2
\ BENFONTEIN -
WESSELTON BaTi ,3 0 27
5 Ti02 + M2 0 3
+ Nb20,
Figure 2.92. Comparison of the compositions of hollandites. K-triskaidecatitanates. Ba-pentatitanates from orangeites. kimberlites,lamproites. and ultrapotassic syenites. Data sources and abbreviations as in Figure 2.91.
MINERALOGY OF ORANGElTES
t ': 1ORANGEITE/ -8-i' ' I '1-. I
i : j : / 1t)5~ ~ 4 -1 / ,."
3 J .·/ ",,,,,,, % ~ I ", ", ", ,. LAMPROITE
2 J-'" i O 00 - O~~o 800~'b°Q
g'lJ OC% 0
234567891011
Cr203 wt. % ) ~
Figure 2.93. enD3 versus V203 for hollandites from orangeites and iamproites. All data this work.
211
2.5.1.3.b. Kimberlites. Hollandites have only been described from the unusual. highly differentiated. Wesselton Water Tunnels and Benfontein calcite kimberlites (Mitchell 1994b). Here they occur in the groundmass of the rock as small (5 x 25 ~m) euhedral crystals intergrown with calcite and dolomite. Their typical absence from other less-evolved kimberlites. including segregation-textured hypabyssal varieties. indicates that hollandite is not a characteristic mineral of archetypal kimberlite. Its absence may be a consequence of the early crystallization of Ti-bearing spinels and perovskite and rapid depletion of the magma in Ti. prior to crystallization of the calcite-rich groundmass at temperatures which might permit K-Ba titanate crystallization. Consequently. Ba remains in the magma until sequestration in late stage phlogopite-kinoshitalite solid solutions.
Representative compositions of kimberlite hollandite are given in Table 2.31. Particularly noteworthy is the absence of potassium and chromium. Hollandites from Wesselton are poor in Nb (not detected) and rich in V203 (0.5-4.0 wt %) relative to those from Benfontein (1.0-5.7 wt % Nb205. V not detected). All are Cr203 poor. The minerals are principally members of the solid solution series between BaFe3+2Ti6016 and BaFe2+ Tb016 with minor solid solution toward mannardite and Nb-bearing hollandite and are thus very similar in composition to Ba-rich hollandites in orangeites (Figures 2.90-2.92).
2.5.1.3.c. Other Alkaline Rocks. Hollandites from the Kovdor carbonatite (Zhuravleva et al. 1978) occur in association with geikileite. zirkelite. and clinohumite. This hollandite. which probably formed by reaction of early crystallizing Ti minerals with residual carbonatite magma, is a Ba-septetitanate. Although lacking V, it is similar in
212 CHAPTER 2
Table 2.31. Representative Compositions of Hollandites from Kimberlites, Ultrapotassic Syenite, and Carbonatite Complexesa
"Total iron expressed as Fe203; n.d. = not detected. Compositions 1-3, Wesselton and 4-6, Benfontein Kimberlites (Mitchell 1994b); 7-8, Little Murun ultrapotassic syenite (Mitchell and Vladykin, 1993); 9, Kovdor complex (Zhuravleva et al. 1978); 10, Schryburt Lake complex, includes 1.38 WI % Ce:!03, 0.74 wI % CaO, 0.24 wI % srO (Platt 1994).
composition, Ko.06Bao.74f'eI.08Ti6.81016 (Table 2.31), but not paragenesis, to hollandites from orangeite.
Hollandites have been described from the Schryburt Lake carbonatite complex (Platt 1994). Here they occur enclosed within perovskite in a perovskite-spinel cumulate derived from an ultramafic lamprophyre. They are similar in composition (Table 2.31) to the Kovdor hollandite in being a Ba-septetitanate. They differ in containing significant amountsofCe203 (1.2-1.7 wt %), CaO (0.6-1.1 wt %), and Nb20S (0.5-0.7 wt %). They also differ in paragenesis, and Platt (1994) considers them to be an early primary phase and not a reaction product.
Hollandites from the Little Murun ultrapotassic complex have been described by Mitchell and Vladykin (1993). Here they occur in aegirine potassium feldspar syenites in association with wadeite, tausonite, sphene, Ti-magnetite, and barytolamprophyllite. They occur as small (<10 ~m) subhedral prisms mantling earlier-formed tausonite, Ti-magnetite, and K-triskaidecatitanate. The paragenesis is interpreted to indicate that the
MINERALOGY OF ORANGEITES 213
hollandites are the products of reaction between preexisting titanates and groundmassforming magma. The hollandites contain 0.2-2.0 wt % K20, 8.7-14.4 wt % Fe203, and 11.5-21.1 wt% BaO (Table 2.31) and are essentially solid solutions between BaFe2+
TbOl6 and BaFe3+ri6016. The Little Murun hollandites lack V, Cr, and Nb but are otherwise similar in
composition to orangeite hollandites (Figures 2.91 and 2.92). Particularly noteworthy is their association with K-triskaidecatitanate (2.5.2)
2.5.2. Potassium Triskaidecatitanate
The Lace, Sover North, and Star orangeites contain a Ba-free potassium titanate whose composition is regarded as K2Ti13027. The Sover North and Star occurrences have been previously described by Mitchell and Vladykin (1993) and Mitchell and Meyer (1989b) respectively.
At Lace this mineral occurs as stellate clusters of slender primatic crystals (Figure 2.94), at Sover North as anhedral crystals mantled by hollandite (Figure 2.95), and at Star as anhedral isolated grains. Optically, the mineral is very similar to hollandite, being opaque in thin section and gray in reflected light.
Figure 2.94. Prismatic potassium triskaidecatitanate (KT), Lace. Backscattered electron image. C = calcite, S = serpentine, M = mica.
214 CHAPTER 2
Figure 2.95. Potassium triskaidecatitanate (KT) mantled by hollandite (H), Sover North. Backscattered electron image.
Tables 2.32 and 2.33 give representative compositions of the mineral and demonstrate there are significant inter-intrusion compositional differences. Individual crystals may be homogeneous (Star) or strongly-zoned with respect to V and Fe (Lace). The mineral has significant V and Nb contents but lacks, or is poor in, K and Cr. There are negative correlations between Ti and Nb or V and Fe.
This titanate was previously regarded as potassium mannardite by Mitchell and Meyer (1989b); however, it has greater Ti02 (approx. 80 wt %) contents and lower K20 (8-10 wt %), Fe203 (1-8 wt %), and V203 (1-14 wt %) than K2Fe3+2Ti6016-K2V2Ti6016 solid solutions (K20 = 12.8-13.0 wt %, Fe203 = 21.8 wt %, V203 = 20.7 wt %, Ti02 = 65.4-66.3 wt %). Structural formulae calculated on the basis of 16 oxygens are not in accord with a hollandite structure. Figures 2.88 and 2.89 show that the mineral is compositionally distinct from hollandite and plots in Figure 2.89 close to ideal K2Ti\3027. Structural formulae calculated on the basis of 27 oxygens, with the exception of high-V examples, are in reasonable agreement with this composition.
Accordingly, the mineral is believed to have the general composition A2B\3027, where A = K, Ba and B = Ti, Nb, V, Fe, Cr. Deviations from the ideal composition are
MINERALOGY OF ORANGEITES 215
Table 2.32. Representative Compositions of K-V Triskaidecatitanates from the Lace and Sover North Orangeitesa
Nb 0.188 0.043 0.070 0.155 0.158 0.084 0.146 0.203 Ti 11.970 11.718 11.656 11.762 11.259 11.080 11.773 11.672 Cr 0.029 0.038 0.023 V 0.238 0.564 0.721 0.807 1.487 2.208 0.031 Fe 0.709 1.026 0.774 0.453 0.525 0.155 1.204 1.268 Mg 0.009 Ba 0.045 0.058 K 2.228 2.095 2.328 2.237 2.017 1.983 2.245 2.306
A 2.228 2.095 2.328 2.237 2.017 1.983 2.290 2.365 B 13.135 13.351 13.270 13.177 13.428 13.528 13.202 13.166
"Total Fe expressed as Fe:!03; n.d. = not detected. Compositions 1-6, Lace; 7-8, Saver North. (All data this work.)
undoubtedly due to nonstoichiometry resulting from the presence of elements occurring in two valence states and site vacancies created by complex B-site substitutional schemes. The compound is considered to be a new mineral. As X-ray diffraction data for this phase are unavailable, the mineral remains unnamed. In this work it is termed "potassium triskaidecatitanate" to reflect its composition.
The only other known occurrences ofK -triskaidecatitanate are from the Little Murun complex, Siberia (Mitchell and Vladykin 1993) and a lamproite from Sisimiut, Greenland (Scott 198 I).
In the Little Murun tausonite syenites it occurs as small anhedral grains mantled and replaced by Ba-rich hollandites. Prismatic crystals are absent, and it cannot be determined if the mineral is primary phase or a reaction product. The Murun material differs from that of orangeites in that it lacks Nb, V, and Cr and is essentially K2(Ti,Fe)13027 (Table 2.33).
The Sisimiut occurrence is described as priderite by Scott (1981); however, the Ti02 content (78.9-81.6 wt %) is too high for this mineral to be a hollandite (Table 2.33). In common with orangeite potassium triskaidecatitanate, the Sisimiut example contains appreciable Fez03 (9.5-9.6 wt %) and has very low BaO contents (0.3-0.8 wt %). This is the only known occurrence of this mineral in a lamproite.
216 CHAPTER 2
Table 2.33. Representative Compositions of K-V Triskaidecatitanates from Star Orangeite, Little Murun Syenite, and Sisimiut Lamproitea
Nb 0.141 0.428 0.584 Ti 11.814 11.111 11.079 12.456 11.769 11.631 13.000 Cr 0.029 0.181 0.222 0.064 0.074 V 0.221 0.689 0.481 Fe 1.027 0.567 0.580 0.696 1.525 1.578 Mg 0.441 0.401 0.289 Ba 0.019 0.058 K 2.105 2.273 2.296 2.088 2.118 2.399 2.000
A 2.105 2.273 2.296 2.088 2.137 2.457 B 13.273 13.367 13.235 13.152 13.358 13.283
"Total Fe expressed as Ftl203; n.d. = not detected; n.a. = not analyzed. 1-3, Starorangeite (Mitchell and Meyer 1989b); 4, Little Mumn syenite (Mitchell and V1adykin 1993); 5-6, Sisimiut lamproite (Scott 1981); 7, ideal composition ofK21i13027.
2.5.3. Barium Pentatitanate
The Sover North orangeite contains rare anhedral opaque crystals of an Fe-bearing barium titanate. A similar mineral has been reported by Mitchell and Vladykin (1993) from the Little Murun tausonite syenite. Both occur as overgrowths upon preexisting titanates. Table 2.34 indicates that both examples have compositions in accord with the general formula Ba(Ti,Fe)sOIl. The monoclinic compound BaTi50n has been synthesized by Tillmans (1969) and Ritter et al. (1986) but has not previously been reported as a mineral.
2.6. PEROVSKITE
2.6.1. Paragenesis
Perovskite is a common groundmass mineral in orangeite. Modal abundances vary from a few grains per thin section (<<0.1 vol %) to approximately 5 vol % (Skinner and Clement 1979; see also 2.6.3). Perovskite abundances range from trace amounts to 1.5 vol % (Fraser 1987) or 3.3 vol % (Clement 1982) in the Finsch orangeites. Commonly,
MINERALOGY OF ORANGEITES
Table 2.34. Representative Compositions of Barium Pentatitanate from the Sover North Orangeite and the Little
Nb 0.020 0.027 Ti 4.807 4.815 4.674 4.406 Cr 0.004 0.004 Fe3+ 0.230 0.208 0.454 0.827 Mn 0.004 Mg 0.031 Ba 0.975 0.945 0.961 0.930 K 0.019 0.019 0.017 0.035
Site occupancies
A 0.994 0.964 0.978 0.965 B 5.061 5.085 5.129 5.233
"Total Fe expressed as Ff!203. Compositions 1-2. Sover North; 3-4. Little Murun; 5. ideal BaTisOll.
217
perovskite is inhomogeneously distributed. Thus, some of the Swartruggens dikes appar
ently contain only trace amounts of perovskite, yet other contemporaneous dikes contain significant amounts (this work).
Most orangeite perovskites are very small (<0.01 mm) subhedral-to-rounded darkbrown crystals. The smallest are difficult to identify optically, as they are opaque and thus may be misidentified as spinels. Euhedral perovskites of cubic habit are only found as chadacrysts in groundmass mica. At Swartruggens euhedral perovskite has been found enclosed in wadeite (this work). The Besterskraal evolved orangeite contains deepred-brown groundmass plates of perovskite that poikilitically enclose altered prismatic silicate minerals (Figure 2.96).
Most orangeite perovskites have habits which indicate that they have been resorbed during the later stages of crystallization of the groundmass. Such perovskites lack ilmenite or rutile mantles, although some from Lace have apatite mantles. Those from Besterskraal are intergrown with apatite and barite.
Compositional data have previously been reported only for perovskites from Bellsbank (Boctor and Boyd 1982), New Elands (Mitchell and Meyer 1989a), and Sydney-on-Vaal (Mitchell and Reed 1988). Of these studies, only Boctor and Boyd (1982) presented full analyses of the mineral. Although new data are provided in this work for perovskites from Besterskraal, Bellsbank, Sover, and Sover North, much further work is required to characterize fully the compositional variation of orangeite perovskite.
Semiquantitative data obtained during the preparation of this work indicate that orangeite perovskite varies considerably in composition. Those present in the least
evolved orangeites contain fewer REE [<6 wt % total (REEh03] and srO «1 wt %) than those in the most evolved varieties [> 10 wt % (REEh03, >3 wt % SrO]. Perovskites from Sover and Bellsbank are rich in REE, but poor in srO «1 wt %).
Representative compositions of perovskite are given in Table 2.35. Sr- and REE-poor perovskites exhibit limited solid solution toward loparite [< 1 0 mol % (Nao.5,REEo.5)Ti03] and tausonite «1 mol % SrTi03) and may be termed "Ce-bearing perovskite" or
Table 2.35. Representative Compositions of Perovskite in OrangeitesO
"Total iron expressed as FeO; n.d. = not detected; n.a. = not analyzed. Compositions 1-2, Sover Mine; 3-4, Sover North; 5-7. Besterskraal; 8-9, Bellsbank Southern Extension; 10, Bellsbank. Data sources 1-9. this work; 10, Boctor and Boyd (1982) .
..!. o
• BESTERSKRAAL + SOVER NORTH • SOVER MINE
Figure 2.97. Compositions of perovskites (mol %) plotted in the ternary system perovskite-Ioparite-tausonite. LH and K are compositional fields ofperovskites from Leucite Hills lamproites (Mitchell and Steele 1992) and kimberlites (Mitchell 1986), respectively.
119
220
w I-a: 0 z 0 :I: U
..... 103
W I-!II:: en > 0 a: w a..
10 2
10
ORANGEITES • SYDNEY - ON - VAAL + BESTERSKRAAL o BELLSBANK • SaVER MINE
KIMBERLITES x BENFONTEIN
F:TI FIELD OF o HYPABYSSAL KIMBERLITES
La Ce Pr Nd SmEu Gd Tb Dy Ho Er TmYb
CHAPrER2
Figure 2.98. Chondrite normalized rare earth distribution patterns for perovskites from orangeites (this work. Mitchell and Reed 1988), Benfontein calcite kimberlite (Jones and Wyllie 1984), and diverse hypabyssal kimberlites (Mitchell and Reed 1988).
"perovskite" (sensu stricto). Figure 2.97 shows that Sr- and REE-rich perovskites from evolved orangeites may be termed "cerian strontian perovskite." The solid solutions present are primarily between perovskite (65-80 mol % CaTi03), loparite (19-27 mol %), and tausonite (5-10 mol %). Sr- and REE-poor perovskites from unevolved orangeites plot near the CaTi03 apex of this diagram. The evolutionary trend of composition is from CaTi03 toward loparite with slightly increasing tausonite contents.
MINERALOGY OF ORANGEITES 221
Figure 2.98 illustrates chondrite nonnalized REE distribution patterns for REE in orangeite perovskites. Unfortunately, the electron microprobe data given by Boctor and
Boyd (1982) have been shown by Jones and Wyllie (1984), and Mitchell and Reed (1988) to be erroneous. Consequently, these data are not plotted in Figure 2.98. As electron beam methods of analysis are unsatisfactory for the analysis of the heavy REE, a complete REE distribution pattern (obtained by ion microprobe) is available only for one perovskite from Sydney-on-Vaal (Mitchell and Reed 1988).
The available data nevertheless demonstrate the extraordinary enrichment of light REE in these perovskites as indicated by the very steep slopes of the REE distribution patterns. The LalYb ratio of the Sydney-on-Vaal specimen is 3145, and no Eu anomaly is present.
Other elements present in significant amounts include FeOT (1-4 wt %), Nb20S (0.6-3.6 wt %), Na20 (1.9-2.3 wt %), and Th02 «0.1-1.5 wt %).
2.6.3. Comparison with Perovskites from Kimberlite
Perovskite is a ubiquitous mineral in archetypal kimberlites. Modal abundances range from trace amounts to 15 vol % (Skinner and Clement 1979, Mitchell 1986, McCallum 1989), but can rise to major levels (>25 vol %) in occurrences such as the Benfontein Sills, where perovskite has been locally concentrated by differentiation processes.
The bulk of the perovskite occurs as discrete euhedral-to-subhedral or rounded crystals ranging in size from 0.01 to 0.2 mm, with the majority of unresorbed crystals being 0.5 to 0.1 mm. Perovskite also typically occurs as reaction mantles about magnesian ilmenite macrocrysts, rutile, and spinels. Perovskite is itself resorbed and mantled by rutile. For further details of perovskite paragenesis in kimberlites, see Mitchell (1986).
The principal differences in the perovskite paragenesis ofkimberlites and orangeites, first noted by Skinner (1989), are in the size and abundance of the crystals. Orangeites typically contain relatively few small perovskite grains, whereas kimberlites contain perovskites which are typically 2-5 times larger and 2-10 times more abundant.
Perovskites in kimberlites are similar in composition (Table 2.36) to those from unevolved orangeites in being essentially CaTi03 (>90mol %) and containing limited amounts of FeOT (1-2 wt %) and Nb20s (0.5-2.0 wt %). There is little variation in the composition of kimberlite perovskites with respect to either paragenesis (Mitchell 1986, McCallum 1989, Jones and Wyllie 1984) or the degree of differentiation of the magma. Sr-, Na-, and REE-rich perovskites do not occur even in the most-evolved kimberlites, e.g., Benfontein [5--6 wt % (REEh03, <1 wt % Na20, <0.5 wt % srO; Jones and Wyllie 1984, Boctor and Boyd 1981], Wesselton Sill [4.3-5.5 wt % (REEh03; Mitchell 1986]. Hence, the majority of kimberlite perovskites so far examined exhibit only very limited solid solution toward loparite «10 mol %) or tausonite «1 mol %) (Table 2.36, Figure 2.97). As perovskites from both orangeites and kimberlites are relatively poor in Na, there is an excess of Ce after formation of loparite in molecular calculation schemes such as used by Mitchell and Vladykin (1993). This excess Ce may be expressed as theCe2Th07 molecule, which typically amounts to <5 mol % (Table 2.36).
222 CHAPl'ER.2
Table 2.36. Representative Compositions of Perovskite in Kimberlites and LamproitesQ
"Total Fe expressed as FeO; n.d. = not detected; n.a. = not analyzed. Compositions 1-5, kimberlites; 6-9,lamproites; I, Frank Smith (Ibis work); 2, Chicken Park (McCallum 1989); 3, Benfontein (Jones and Wyllie 1984); 4, Premier (Jones and Wyllie 1984); 5. liqhobong (Boctor and Boyd 1980); 6-7. Middle Table Mountain. Leucite Hills (Mitchell and Steele 1992); 8. Walgidee Hills, West Australia (Ibis work); 9, Pilot BUlle, Leucite Hills (Carmichael 1967a).
Figure 2.99 shows that although the Sr and Ce contents of perovskites from kimberlites and unevolved orangeites are similar, there are no counterparts to the orangeite Srand REE-rich perovskites in kimberlites.
The single exception to the above observations is the Green Mountain "kimberlite" from which Boctor and Meyer (1979) report perovskites with 6.8-11.5 wt % Nh205 and 5.9-10.3 wt % (REEh03. These perovskites lack Na; thus, solid solutions with CaNh206 or Ca2Nb207 rather than loparite must be present. The anomalous composition of these perovskites, . relative to those from all other kimberlites, together with the unusual mineralogy of the rocks, i.e., abundant diopside and enstatite(?), suggest that the classification of this rock as a kimberlite should be re-appraised.
In common with orangeite perovskites those from kimberlites are enriched in the light rare earths. Chondrite normalized REE distribution patterns (Figure 2.98) have very steep slopes and no Eu anomalies. LalYb ratios of perovskite from hypabyssal kimberlites range from ] 828 to 3229 and are thus not significantly different from that of the
MINERALOGY OF ORANGElTES 223
7'0
t 6'0
ae 5'0
..: ~ 4'0
0 ... C/) 3'0
2'0
1'0
• BESTERSKRAAL + SOVER NORTH • SOVER o BELLSBANK
WALGIDEE a HILLS •• " "PILOT BUTTE
~ LEUCITE HILLS
/ PRAIRIE '" CREEK OLIVINE LAMPROITE
_-,------------,/ ELLENDALE
LI (/ .... , •.• ,,) 0."'" - .......... o ..... PEROVSKITE PYROXENITES , / .... ~" .' ...... '
.-~ :...,....-.-...... "_~.",, 00 · ... 0 KIMB R ~ANDtmOITESo ...... I I I I ~~_r_-._--r---r-__,_-_,.-_,---I
Figure 2.99. Ce203 versus srO (wt %) compositional variation of perovskites from orangeites (this work). lamproites (Mitchell and Steele 1992). kimberlites. alntiites. and perovskite pyroxenites (this work).
Sydney-on-Vaal orangeite perovskite. However, Figure 2.98 demonstrates that perovskites from orangeites are enriched in the light REE (La-Nd) relative to those from kimberlites. The figure also shows that kimberlite perovskites typically have LaN < CeN in contrast to those from orangeites where LaN> CeN. Although the data base is extremely limited, it appears that the heavy REE abundances are not significantly different.
2.6.4. Comparison with Lamproite Perovskite
Perovskite is not a ubiquitous mineral in lamproites. It is found only in olivine lamproites (Ellendale, Prairie Creek) and evolved madupitic lamproites such as occur at Pilot Butte and Middle Table Mountain (Leucite Hills). It typically forms small (<0.50 ~m) subhedral-to-euhedral crystals and is commonly poikilitically enclosed by groundmass mica (Mitchell and Bergman 1991). Poikilitic plates of perovskite, similar in character to those found in the Besterskraal orangeite, have been described from the Middle Table Mountain lamproite by Mitchell and Steele (1992).
The composition (especially the REE abundances) of perovskite in lamproites is inadequately characterized. Available data are summarized by Mitchell and Bergman (1991). Perovskite in Ellendale (West Australia) olivine lamproites varies widely in composition within and between intrusions and contains 3-12 wt % (REE)203, 0.2-1.5 wt % srO, 0.9-2.1 wt % FeO, and 0.6-1.3 wt % Nb20S (this work). Perovskites from the Walgidee Hills (West Australia) lamproite pegmatite and the Prairie Creek (Arkansas) madupitic lamproite contain 0.2-2.8 wt % srO and <2 wt % Ce203 (Mitchell and
224 CHAPTER 2
lLI I- 104 a:: 0 z 0 I U ....... lLI I-~ (J)
> 0
10 3 a:: lLI a..
LAMPROITES
: } MIDDLE TABLE MOUNTAIN
o WALGIDEE HILLS
[8l FIELD OF ORANGEITE COMPOSITIONS
La Ce Pr Nd Sm Eu Gd Tb Oy Ho Er
Figure 2.100. Chondrite normalized rare earth distribution patterns of perovskites from lamproites (Mitchell and Steele 1992, Mitchell and Reed 1988).
Bergman 1991, Figure 2.99). Perovskite from Pilot Butte and Middle Table Mountain are enriched in Sr and REE relative to all other lamproite perovskites. The purple poikilitic perovskites from Middle Table Mountain are remarkably similar in composition to Srand REE-rich perovskites from the Besterskraal and Sover North orangeites (Table 2.36, Figures 2.97 and 2.99).
Chondrite normalized REE distribution patterns for lamproite perovskite are illus
trated in Figure 2.100. Although reliable data for the heavy REE are not available, the
patterns indicate that these perovskites are strongly enriched in the light REE with LaN> CeN. The distribution patterns and REE abundances are similar to those of orangeite perovskite.
MINERAWGY OF ORANGEITES 225
2.7. PHOSPHATES
2.7.1. Apatite
2.7.1.1. Paragenesis
Apatite is a ubiquitous groundmass mineral with abundances ranging from trace amounts « 1 vol %) to 10 vol %. The distribution is not homogeneous within and between intrusions. Thus different facies of individual dikes at Swartruggens contain markedly varied abundances of apatite, and at Bellsbank the Bobbejaan dike contains significantly more apatite than the Main dike (Bosch 1971, Skinner and Clement 1979, Clement 1982, Fraser 1987, this work).
Apatite occurs primarily as euhedral elongated prisms ranging from 0.05 mm in width to 0.3 mm in length. Prisms of 0.05-0.1 mm in length are common. Apatite is commonly resorbed and may be replaced by calcite, Sr-Ba carbonate, or barite. In the Lace, Swartruggens, and Besterskraal orangeites, apatite occurs as large (0.1-0.3 mm) anhedral groundmass plates which poikilitically enclose previously formed spinels and perovskite. Inclusions of pyrite and magnetite are found in some apatites at Swartruggens. Apatite may occur as mantles upon perovskite or be complexly intergrown with daqingshanite and other REE phosphates. Apatite is not preferentially associated with groundmass carbonates and appears to have crystallized prior to and/or contemporaneously with late-stage mica. The radiating aggregates of acicular apatite, which are common in kimberlites, do not appear to be characteristic of orangeites.
2.7.1.2. Composition The composition of apatite in orangeites has not been previously studied. Repre
sentative compositions of apatites analyzed during the preparation of this work are given in Table 2.37. Fluorine and water contents were not determined, so it is not known whether the mineral is fluoro- or hydroxyapatite. X-ray spectra did not indicate the presence of chlorine.
Table 2.37 shows that apatites contain significant quantities of Si replacing P and that the principal compositional variation is with respect to their Sr, REE, and Ca contents. Prismatic apatites have the least Sr and REE contents and may be zoned from Sr-poor cores to Sr-rich margins. The Sr content varies between intrusions; thus, apatites from Swartruggens contain 3-6 wt % SrO, whereas Lace, Sover Mine, and Bellsbank apatites contain only 1-3 wt % srO. REE contents of all these early-formed apatites are low [<2(REEh03, commonly <1 wt %].
Groundmass poikilitic plates of apatite from Besterskraal contain from 2 to 22 wt % srO, 0.60 to 2.6 wt % Si02, <1 to 8 wt % (REEh03, and up to 1.5 wt % BaO (Table 2.38). Typically, Sr < Ca in these apatites, and they are rarely replaced by another discrete Sr-phosphate with Sr> Ca (see 2.7.4) and Sr-Ba carbonate. Besterskraal apatites are light REE rich with LaN> CeN.
2.7.1.3. Comparison with Kimberlite and Lamproite Apatite
2.7.1.3.a. Kimberlite. Apatite is a ubiquitous late-crystallizing groundmass mineral in kimberlites in quantities ranging from 1 to 10 vol%. Crystals are typically euhedral,
226 CHAPTER 2
Table 2.37. Representative Compositions of Euhedral ApatiteU
aTotal iron expressed as FeO; n.d. = not detected; n.a. = not analyzed. Composition 1-2, Lace; 3-4, Bellsbank; 5-6, Sover Mine; 7-8, Swartruggens. (All data this work.)
small (0.1-0.2 mm) prisms. Apatites are particularly abundant in calcite-rich portions of the groundmass where they commonly occur as radiating sprays of acicular crystals which have nucleated at the margins of calcite-serpentine segregations. The habit of such apatites suggests they formed during rapid quenching of the magma (Mitchell 1986, Clement 1982). Poikilitic groundmass plates of apatite have not been reported from archetypal kimberlites. Apatite is difficult to recognize optically where it is intimately intergrown with groundmass calcite. In many instances such apatites are completely pseudomorphed by calcite. For further details of apatite paragenesis see Mitchell (1986).
Very little is known of the composition of kimberlite apatites (Mitchell 1986). They are poor in SrO «1 wt %; Scott Smith et al. 1984, Exley and Smith 1982) and (REEh03 «1 wt %; Ilupin et al. 1971, Exley and Smith 1982, Mitchell 1984a). There is apparently no difference in their Sr and REE contents with respect to the degree of evolution of the host kimberlite. Jones and Wyllie (1984) have shown that REE abundances in perovskites are greater by an order of magnitude than those of coexisting apatite. This relationship is also observed in orangeites. Apatites from Benfontein and Premier have linear chondrite
Table 2.38. Representative Compositions of Poikilitic Apatite from Besterskraalu
a All data (this work) are SEMIEDS analyses summed to 100 wt % oxides. Total Fe expressed as FeO; n.d. = not detected.
MINERAWGY OF ORANGElTES 227
normalized distribution patterns in which LaN> CeN and Eu anomalies are absent (Jones and Wyllie 1984). Apatites in kimberlites contain significant quantities of Si02 (0.7-2.2 wt %; Scott Smith et.al. 1984, Exley and Smith 1982).
Kimberlite apatites differ from those in orangeites with respect to their paragenesis. In orangeites, apatites are principally microphenocrysts and late-stage groundmass crystals. Rapidly-quenched apatites associated with calcite segregations are not characteristic of orangeites. Late-stage crystallization as poikilitic plates is not observed in kimberlite. Orangeite apatites are typically richer in srO (> 1 wt %) than those from kimberlites. REE abundances, with the exception of the Besterskraal apatites, are similar.
2.7.1.3.h. Lamproite. Apatite is a ubiquitous mineral in lamproites, occurring as phenocrysts and microphenocrysts. Many apatites exhibit rounded and embayed habits. Resorption is not associated with replacement by calcite. Hollow-cored apatites which may have grown rapidly from supersaturated melts are common. Apatite crystallizes after phenocrystal phlogopite and is poikilitically enclosed by groundmass potassium richterite and sanidine. Crystallization is contemporaneous with priderite, wadeite, and perovskite (Mitchell and Bergman 1991).
Lamproite apatites are fluor-apatites (2-7 wt % F) characterized by high srO (typically 1-6 wt %, ranging up to 12 wt %, in some Leucite Hills examples). BaO contents vary widely. Those from the Leucite Hills typically contain 0.2-0.4 wt % BaO, with the exception of Sr-Ba apatites from Middle Table Mountain which contain up to 18 wt % Ba. The majority of apatites in the West Kimberley province contain 0.5-1.0 wt % BaO, although examples from the Walgidee Hills contain 2.1-12.3 wt % BaO (Mitchell 1986, Edgar 1989). Apatites with high Ba and Sr contents appear to be characteristic of lamproites. Although few reliable data exist, lamproite apatites appear to be poor in (REEh03 (<2 wt %; Kuehner et al. 1981, Carmichael 1967a). The single REE distribution pattern available is linear with LaN> CeN, a LalYb ratio of 90 and no Eu anomaly (Mason 1977). Si02 contents are low (typically <1 wt %; Mitchell 1986, Edgar 1989).
Apatites inlamproites do not differ significantly from those occurring in orangeites either in paragenesis or composition with the possible exception of their Si02 contents. However, further studies are required, as the composition of apatite in both rock types is insufficiently characterized.
2.7.2. Daqingshanite Orangeites from the Sover Mine contain anhedral 10-50 Ilm crystals of a REB-Sr
phosphate set in a matrix of calcite. Table 2.39 indicates that this mineral has a composition which corresponds to that of daqingshanite [(Sr,Ba)3REE(P04)(C03)3; Yingchen et al. 1983]. The Sover North example is very similar to daqingshanite-(Ce) from the Nkombwa carbonatite complex, Zamibia (Appleton et al. 1992) in being poor in BaO relative to barium daqingshanite-(Ce) from the type locality at the Bayan Obo iron ore-REE deposit, China (Yingchen et al. 1983). Although X-ray data are not available, the compositional data indicate that this Sr-REE carbonate is undoubtedly daqingshanite-(Ce). No other orangeites have been examined in sufficient detail to determine if
228 CHAPTER 2
Table 2.39. Representative Compositions of DaqingshaniteQ
Ca 0.322 0.356 0.122 0.723 Sr 2.669 2.614 2.694 1.655 3.000 Ba 0.103 0.169 0.199 0.685 Fe 0.030 0.033 La 0.225 0.262 0.419 0.318 Ce 0.474 0.522 0.498 0.407 1.000 Pr 0.035 0.047 0.034 0.027 Nd 0.095 0.104 ·0.068 0.062 P 0.868 0.776 0.987 1.086 1.000 Si 0.232 0.243
Site occupancies
A 3.124 3.172 3.005 3.063 3.000 B 0.829 0.935 0.864 0.814 1.000 C 1.100 1.019 0.987 1.086 1.000
"Total Feexpressedas FeO; n.a. = not analyzed. Compositions 1-2, Sover Mine (this work) expressed as 100 wt % oxides on a C02-free basis; 3, Nkombwa (Appleton et al. 1992); 4, Bayan Obo (Yingchen et al. 1983); 5--6, ideal composition of Sr3Ce(P04)(C03)3 expressed in terms of oxide wt % on a C02-free basis; 6, composition recalculated to 100 wt % oxides.
daqingshanite is a common mineral in these rocks. The abundance of Sr-Ba-REE carbonates and other Sr-REE-rich phosphates suggests that further investigations will bring to light new occurrences of this mineral. Daqingshanite has not been found in either lamproites or kimberlites.
2.7.3. Monazite
Orangeites from Bellsbank Southern Extension commonly contain oval 1-10 J..lm grains of a monazite-like, complex Th-Sr-REE phosphate, AB04, where A = REE, Th, Sr, Ca, and B = P, Si. The phase occurs as isolated discrete crystals set in an Fe-bearing dolomite groundmass, and as irregular complex intergrowths with Sr-poor apatite and Nb-rutile. Rarely, the phosphate is completely mantled by Nb-rutile. Isolated small grains of the same mineral are also common in the Lace orangeites.
MINERALOGY OF ORANGElTES 229
Table 2.40. Representative Compositions of Monazitea
"Total Fe expressed as FeO: n.d. = not detected: n.a. = not analyzed. All data (this work) SEMIEDS analyses summed to 100 wt % oxides. Compositions 1-4, Bellsbank: 5-6, Lace.
The mineral shows considerable intergrain compositional variation with respect to Th, Sr, and Ba (Table 2.40). Monazites from Lace are not as rich in Th and Ba as those from Bellsbank. Extremely high Th contents are rarely found in Bellsbank monazite.
Monazite is not present in kimberIites or lamproites (Mitchell 1986, Mitchell and Bergman 1991).
2.7.4. Sr-REE Phosphate At Besterskraal, an unidentified REE-poor Sr-phosphate with srO > CaO occurs as
discrete mantles upon relatively Sr-poor apatite. This mineral is intergrown with Sr-Ba carbonates and strontianite and may represent a reaction product formed between preexisting apatite and Sr-rich late-stage deuteric hydrothermal fluids. The mineral contains 26-37 wt % SrO, 20-36 wt % CaO, 1-5 wt % BaO, <4 wt % (REEh03, 31-34 wt % P205, and 0.8-1.3 wt % Si02.
2.S. AMPHIBOLES-POTASSIUM RICHTERITE
Amphiboles with the optical characteristics of potassium richterite have been recognized in evolved varieties of orangeite from Pniel (ErIank 1973, Tainton 1992), Sover North (Tainton 1992, this work), Besterskraal (Skinner pers comm., this work), Lace (this work), Makganyene (this work), and occurrences in the Prieska area (Skinner et al. 1994).
2.S.1. Paragenesis
Potassium richterite crystallizes as a late-stage groundmass mineral. At Pniel, and in the SN2 intrusion at Sover North, it forms anhedral tablets which may subpoikilitically enclose phlogopite and diopside. Pleochroism is from yellow-to-pinkish brown and crystals are zoned toward more strongly pleochroic rims. Thin overgrowths of greenbrown pleochroic amphibole may occur upon the potassium richterite or diopside (Tainton 1992). At Sover North, amphibole forms reaction coronas around olivine xenocrysts. Amphibole is rarely found as strongly-yellow-pink pleochroic anhedral
230 CHAPTER 2
crystals in the SN 1 intrusion at Sover North (Tainton 1992). In the Besterskraal orangeite, amphibole occurs as pink poikilitic plates intergrown with groundmass potassium feldspar (this work).
Skinner and Scott (1979) have tentatively suggested that altered elongated laths of a mineral from Swartruggens, which is regarded as an abundant (trace to 12 vol %) partially altered primary phase, may be "anthophyllite." The composition (Si02 = 52.5, Ah03 = 1.98, FeO = 5.4, MgO = 26.6, CaO = 0.6, Na20 = 0.2, K20 = 0.6, total = 88.1 wt %) is not in accord with that of anthophyllite and may represent that of a mixture of alteration products (serpentine, chlorite, brucite, etc.). It is considered here that anthophyllite is not present in orangeites, and the material described by Skinner and Scott (1979) represents pseudomorphs after an undetermined primary mineral.
2.8.2. Composition
The few compositional data available (Tainton 1992, this work) demonstrate that there is considerable intra- and inter-intrusion compositional variation (Figures 2.101 and
BESTERSKRAAL "' .... • '" \ '" \
PNIEL '" ' 0 '" ' , • SOVER NORTH
, 12 I .. ,
( , ZONATION TREND , , , , , , ,
10 , , , WEST
ae ,'KIMBERLEY , , - , ~ 8
, , , I- ,
0 I
If , /
6 : /-k-"\<V I , .~Q~ I " < Cj <Q~ , '\ I , \ I KAPAMBA , , ,
V 4 , , I , , , ,
2
2 3 4 5 6
Na20 wt. % • Figure 2.101. Na20 versus FeOT compositional variation ofamphiholes from orangeites. Data sources: Sover North (this work, Tainton 1992), Besterskraal (this work), Pniel (Tainton 1992). Compositional fields of amphiboles from lamproites from Mitchell and Bergman (1991).
MINERALOGY OF ORANGEITES
0'.
0'4
0'2
MARIO
1'0
UCITE HILLS
2'0 3'0
No I K 4'0
231
• BESTERSKRAAL o PNIEL • SOVER NORTH
5'0 6'0
Figure l.IOl. Na/K versus Tt (atomic) compositional variation of amphiboles from orangeites (this work, Tainton 1992). West Kimberley. Leucite Hills. Smoky Butte. and Francis lamproites (Mitchell and Bergman 1991). Kapamba lamproites (Scon Smith et al. 1989). and potassic lamprophyres from Kajan (Wagner 1986). Pendennis (Hall 1982). Bohemia (Nemcl! 1988). Yinniugou (Zhao et al. 1993). New South Wales (NSW) leucitites (Mitchell and Bergman 1991). and the Murun ultrapotassic complex (this work).
2.102). At Pniel, amphiboles are very low in Ti02 (0.1-1.8 wt %) ranging in composition from potassium richterite cores to magnesioarfvedsonite mantles (Table 2.41; Tainton 1992). Amphiboles at Sover North are primarily potassium richterites which exhibit a wide variation in Ti02 content (2.0-7.1 wt %) at approximately constant NaIK. ratios. Those richest in Ti02 (>0.25 Ti atoms/23 oxygens; Figure 2.102) are titanian potassium richterites and titanian potassium magnesiokatophorites (Table 2.42; Tainton 1992, this work). These amphiboles are mantled by Ti02-poor «1.5 wt %) magnesioarfvedsonite and arfvedsonite (Table 2.42, Tainton 1992). Amphiboles at Besterskraal show considerable intergrain compositional variation with respect to Ti02 (4.7-7.1 wt %) at approximately constant NaIK. ratios (Table 2.41, this work). The general evolutionary trend in all occurrences is one of an initial increase of Ti with Fe, followed by decreasing Ti and K and increasing Fe3+ with Na.
Amphiboles in the Pniel and Sover North orangeites are characterized by low Ah03 «0.7 wt %) contents, in contrast to those from Besterskraal which contain 0.8-1.6 wt % Ah03 (Tables 2.41 and 2.42), Richterites from Pniel contain sufficient Si and AI to occupy
232 CHAPTER 2
Table2Al. Representative Compositions of Amphiboles from Pniel and BesterskraalQ
aFeO and Fez03 calculated on the basis of stoichiometry (Droop 1987); FeOr = total Fe expressed as FeO; n.d. = not detected. Compositions 1-3, potassium richterite, Pnie1 (fainton 1992); 4-6, potassium magnesioarfvedsonite, Pniel (fainton 1992); 7-10, titanian potassium richterite and magnesiokatophorite, Besterskraa1 (this work).
all of the tetrahedral sites and thus contain octahedrally coordinated AI. Richterltes from Besterskraal and Sover North contain insufficient Si and Al to occupy all the available tetrahedral sites in the structure. This deficiency is probably remedied by entry ofTi to this site, as suggested for lamproite-derived amphiboles by Mitchell and Bergman (1991) and Thy et al. (1987). In contrast, potassium magnesioarfvedsonites characteristically contain octahedrally coordinated AI.
In many cases it is not possible to calculate the Fe2+ and Fe3+ contents of the amphiboles using the method of Droop (1987), as they are nonstoichiometric and contain less than 13 cations/23 oxygens. This, in some instances, is because the presence of Ti4+ creates vacancies in the structure (Mitchell and Bergman 1991). However, in all cases where estimation of oxidation state is possible, it appears that Fe2+lFe3+ ratios are much greater than unity (Tables 2.41 and 2.42).
MINERALOGY OF ORANGEITES 233
Table 2.42. Representative Compositions of Amphiboles from Sover NorthQ
Ca I 0.854 0.879 0.869 0.660 0.608 0.054 0.088 0.112 Na B 1.146 1.121 1.131 1.340 1.392 1.946 1.912 1.888
Na I 0.064 0.024 0.161 0.032 0.035 0.153 0.092 K A 0.886 0.883 0.855 0.856 0.878 1.057 1.009 1.006
"FeO and Ft!:!OJ calculated on the basis of stoichiometry (Droop 1987). FeOr = total Fe expressed as FeO. Compositions 1-3, potassium richterite (Tainton 1992); 4-5, titanian potassium richterite (this work); 6-8, potassium magnesioarfvedsonite (Tainton 1992).
Levels of BaO and Cr203 in amphiboles in orangeites are not detectable by standard electron microprobe analytical methods «0.1 wt %). Fluorine contents have not been determined.
2.8.3. Comparison with Potassium Richterite in Lamproite and Other Potassic Rocks
Titanian potassium richterite is one of the typomorphic minerals of lamproites (Mitchell and Bergman 1991). However, its presence is neither ubiquitous nor confined to lamproites. Importantly, primary amphiboles do not occur in archetypal kimberlites (Mitchell 1986).
In lamproites, potassium richterite occurs principally as groundmass poikilitic plates and is one of the last minerals to crystallize. Potassium richterite is commonly optically
234 CHAPTER 2
zoned, from pale-yellow-pink cores, through reddish-pink regions, to dark-red-brown margins. The amphiboles are commonly intergrown with groundmass plates of potassium feldspar. A second, less common paragenesis, is as small euhedral prisms lining vesicles in extrusive lamproites. These amphiboles are optically and compositionally similar to the groundmass poikilitic variety.
The paragenesis of poikilitic potassium richterite occurring in evolved orangeites is essentially identical to that of richterite in lamproites. Although at Pniel potassium richterite does not coexist with potassium feldspar, intergrowths with this mineral occur at Sover North and Besterskraal.
Mitchell and Bergman (1991) have shown that individuallamproite provinces are characterized by the presence of titanian potassium richterite and titanian potassium magnesiokatophorite of distinct composition with respect to their Fe and Na contents and Na/K ratios. Amphiboles from Murcia-Almeria, Kapamba, and Francis are Na-rich relative to those from the Leucite Hills, West Kimberley, Smoky Butte, and Prairie Creek (Figure 2.1 01). Compositional trends in all provinces are similar and characterized by increasing Fe and Na contents, commonly at nearly constant K contents. This trend represents evolution from FeOT-poor (2-3 wt %) titanian potassium richterite, through FeO~-rich (10-14 wt %) potassium richterite, to titanian potassium magnesioarfvedsonite. Lamproite amphiboles are Ah03 poor «1.5 wt %) and Ti02 rich (2-9 wt %). Substantial tetrahedral site deficiencies are typically present as a consequence of the low Al contents. Ferric iron-rich amphiboles are not characteristic of lamproites, and even rocks containing tetraferriphlogopite lack such amphiboles.
Figure 2.101 shows that amphiboles from orangeites have compositions similar to those of amphiboles in many lamproites. Compositional zonation and evolutionary trends are similar. Potassium richterites in orangeites have Na20 < K20 ratios greater than 0.8 (Pniel, 0.82-0.95, most <0.9; Sover North, 0.84-1.03; Besterskraal, 0.95-1.12) and in this respect are similar to amphiboles in some slightly evolved (6-10 wt % FeOT) amphiboles in the West Kimberley and Leucite Hills provinces. Tetrahedral site deficiencies are common to amphiboles from both rock types.
Figure 2.102 shows that richterites in orangeites differ principally from those in lamproites in terms of their Ti contents. All bona fide lamproites contain more than 0.25 atoms Ti/23 oxygens. Amphiboles from Sover North exhibit a wide range in Ti content at approximately constant Na/K ratios. Although the majority of these data do not fall within fields defined by lamproite amphiboles, some high Ti (>0.7 afu) examples of lamproitic character are present in examples with the highest Na/K ratios.
Richterites from Besterskraal have similar Ti contents and N a/K ratios to amphiboles from the Leucite Hills and Smoky Butte lamproites (Figure 2.102). Although they have Ti contents similar to amphiboles from Francis and Kampaba, they do not plot in the fields defined by these amphiboles as the latter have significantly lower K20 contents «4 wt %).
Amphiboles from Pniel are particularly low in Ti, and compositions plot far from the fields defined by lamproite amphiboles in Figure 2.102. This may be a consequence of the relatively unevolved character of their host rocks, if the trend defined by the Sover amphiboles is considered indicative of the evolutionary trend of orangeite amphiboles with respect to Ti and N a/K ratios. In this context it is significant that the Pniel amphiboles
MINERALOGY OF ORANGEITES 23S
have the lowest Na201K20 ratios and do not coexist with potassium feldspar. Hence, they and their host rocks may be considered as unevolved, relative to Sover North and Besterskraal orangeites.
Ti-poor magnesioarfvedsonites occur as mantles on richterite in both orangeites and lamproites. These amphiboles probably form as a result of reaction of preexisting amphiboles with late-stage Na- and Fe-rich deuteric fluids and may even be of subsolidus origin (Mitchell and Bergman 1991). Alkali amphiboles in lamproites tend to be richer in Ti02 (1-7 wt %; Mitchell and Bergman 1991) than those in orangeites «1 wt % Ti02), but are otherwise of similar composition.
In summary, amphiboles in orangeites and lamproites are similar in paragenesis and composition, the principal difference being that richterites in orangeites tend to be poorer in Ti02. On the basis of the existing data, there is no means of distinguishing richterites in orangeites from those in lamproites on the basis of their composition (see below).
Figure 2.102 also illustrates the compositions of some richteritic and magnesioarfvedsonitic amphiboles from peralkaline minettes, leucitites, and potassic ultramafic lamprophyres. Amphiboles from these occurrences typically have much higher Na/K ratios and lower Ti contents than amphiboles in lamproite ororangeite. Many have higher Ah03 contents (1-3 wt %) and coexist with nepheline and/or aluminous pyroxenes.
Mitchell and Bergman (1991) discussed the occurrence and paragenesis of richteritic amphiboles and concluded that, in some instances, amphiboles of similar composition may occur in different rock types because of a convergence of amphibole evolutionary trends in genetically unrelated magmas. Amphibole composition alone is not regarded as diagnostic of any particular magma type and must be considered in conjunction with the composition and paragenesis of associated minerals, if it is used in classifying a given rock. Thus, richterites in leucitites, while similar in some respects to those in some lamproites or minettes, differ in being richer in Al and occurring in association with nepheline and AI-rich pyroxenes.
2.9. POTASSIUM FELDSPAR
Potassium feldspar has been recognized as a late-stage groundmass mineral in orangeites from Besterskraal (Skinner pers. comm., this work), Sover North (Tainton 1992), Postmasburg (this work), Makganyene (this work), Voorspoed (this work), and the Prieska region (Brandewynkuil, Sweetput-Soutput, Droogfontein, Albertshoop, Nauga; Skinner et al. 1994). Typically the mineral occurs as colorless groundmass poikilitic plates and, apart from serpentine, is the last complex silicate phase to crystallize. The feldspar is typically altered to clay minerals and/or replaced by serpentine. Potassium feldspar forms pseudomorphs after (?) leucite in the Sover North and Postmasburg orangeites (Tainton 1992).
Table 2.43 gives representative compositions of primary potassium feldspar from the Sover North and Postmasburg occurrences (Tainton 1992, this work). Feldspars from Postmasburg contain <0.26 wt % Fe203, 0.15-0.4 wt % Na20, and 0.7-2.1 wt % BaO. Those from Sover North contain 0.3-1.7 wt % Fe203, <0.02--0.12 Na20, and 0.1-1.5 wt % BaO. Potassium feldspar pseudomorphing "leucite" (see 1.10) does not differ in composition from primary feldspar (Table 2.43; Tainton 1992).
236 CHAPTER 2
Table 2.43. Representative Compositions of Potassium Feldspar"
aTotal iron expressed as Fe20J; n.d. = not detected; n.a. = not analyzed. Compositions 1-3. groundmass sanidine, Sover North (this work); 4, sanidine pseudomorphing "Ieucite," Sover North (Tainton 1992); 5-7, groundmass sanidine, Postmasburg PK37 (this work).
Potassium feldspar is not found in archetypal kimberlites but is a common mineral in phlogopite and madupitic lamproites (Mitchell and Bergman 1991), where it occurs as euhedral-subhedral prismatic microphenocrysts, as groundmass poikilitic plates, and as pseudomorphs after leucite. Only the latter two parageneses are typical of potassium feldspar in orangeites.
Mitchell and Bergman (1991) have shown that sanidines in lamproites are characterized by low Na20 contents «2.5 wt %), widely varying Fe203 contents «0.1-5 wt %, typically >1 wt %), and 0.1-1.7 wt % BaO. Iron-rich feldspars belonging to the iron sanidine-sanidine series containing more than 1 wt % Fe203 (>4 mol % KFeSbOs), less than 1.0 wt % Na20, and negligible CaO «0.1 wt %) appear to occur exclusively in lamproitic parageneses. Each lamproite province is characterized by feldspar of a distinct composition with respect to Fe and Na content.
Figure 2.103 indicates that potassium feldspars from orangeites are similar in composition to groundmass sanidines from the West Kimberley and Kapamba lamproite provinces, but poor in Fe203 and Na20, relative to groundmass and microphenocrystal sanidines from other provinces. The BaO contents of both lamproite or orangeite feldspars
MINERALOGY OFORANGElTES
1'0
It)
02'0 N
If +
I 1'0 -+
o
LEUCITE HILLS
MURCIA - ALMERIA MADUPITIC
LAMPROITES ~ ,
FRANCIS
'SMOKY BUTTE
'----MURCIA -ALMERIA--PHLOGOPITE LAMPROITES
• POSTMASBURG - PK37 + SOVER NORTH
PRIMARY • SOVER NORTH
PSEUDOMORPHS
1'5
237
Figure 2.103. Na20 versus Fe203 (wt %) compositional variation of potassium feldspars from orangeites (this work, Tainton 1992) and lamproites (Mitchell and Bergman 1991, Scott Smith et al. 1989). WK = West Kimberley; K = Kampamba.
have been insufficiently characterized, but on the basis of the available data do not appear to be significantly different.
In summary, potassium feldspars in orangeites are similar in paragenesis and composition to some Fe- and Na-poor sanidines in lamproites. The iron-rich sanidines which are characteristic of lamproites have not been recognized in orangeites.
Potassium feldspars in the Swartruggens Male lamprophyre are similar to those in orangeites in having negligible Na20 contents and containing up to 1.0 wt % Fe203. They differ in that they are irregularly, continuously-zoned to Ba-rich varieties, which may contain up to 7 wt % BaO (this work).
2.10. ILMENITE
Macrocrystal magnesian ilmenite is not a characteristic mineral of orangeites (see 1.5). Groundmass ilmenite has been recognized in the Lace, Sover Mine, Besterskraal, Voorspoed, and Finsch orangeites (this work). In these occurrences, ilmenite occurs as small subhedral plates «25 /lm) and subhedral-to-euhedral (25 x 1-5 /lm) laths. The
238 CHAPTER 2
Table 2.44. Representative Compositions of llmeniteQ
aFeOr = total Fe expressed as FeO; n.d. = not detected. Compositions 1-4. Sover Mine (this work); 5-9, Lace (this work). Composition 8 contains 2.0 wt % Fe:!OJ, and 39.1 wt % FeO. and composition 9 contains 5.6 wt % Fe20J and 33.4 wt % FeO when recalculated on a stoichiometric basis.
irregular habit of the crystals suggests that original euhedral ilmenite laths became unstable subsequent to their crystallization and were resorbed during the later stages of crystallization of their parent magma. I1menites are not homogeneously distributed and not preferentially associated with any other groundmass phase. They may be found in contact with spinel and/or perovskite and included in apatite. The majority occur as discrete crystals in a carbonate mesostasis. Groundmass ilmenite does not occur in the Swartruggens orangeite or lamprophyre dikes.
Although few data are available, it appears that ilmenites from orangeites exhibit significant intra- and inter-intrusion compositional variation (Table 2.44). I1menites from one sample from Lace contain <0.1-3 wt % MgO, 1.2-1.9 wt % Nb20S, and 2.6-7.2 wt % MnO, whereas those in another contain <0.1-0.6 wt % MgO, 0.5-1.3 wt % Nb20S, and 17-19 wt % MnO (this work). I1menites from Sover Mine contain <0.1 wt % MgO and 8.4-10.2 wt % MnO and are relatively enriched in Nb20s (1.8-5.5 wt %; this work).
MINERALOGY OF ORANGElTES
MnTi03 60
t4;!~~t;] CARBONATITES
KIM8ERLITE GROUNOMASS
AND MEGACRYST
90
10 FeTi0 3
Figure 2.104. Compositions (mol %) of ilmenites from the Lace and Sover Mine orangeites (this work) plotted in the ternary system geikielite (MgTi03Hlmenite (FeTi03)-pyrophanite (MnTi03). (*) = Mg-rich ilmenites from Lace. Compositional fields for ilmenites from kimberlites (Mitchell 1986, Shee 1984), for carbonatites from Gaspar and Wyllie (1984) and Mitchell (l978b). Field P, ilmenites from the Premier calcite kimberlite dike (Gaspar and Wyllie 1984). Field R, outlined by dashed lines, from Wyatt (1979) for Premier kimberlites. Points joined by arrow are Mg-rich core and Mn-rich margin of an ilmenite from the Premier calcite kimberlite dike.
All the ilmenites contain negligible Ah03 «0.1 wt %) and have low Cr203 contents «0.1 wt % at Sover Mine; 0.2-0.6 wt % at Lace).
Table 2.44 and Figure 2.104 indicate that the ilmenites are primarily members of the pyrophanite (MnTi03Hlmenite (FeTi03) solid solution series. Attempts to calculate ferric iron contents on the basis of stoichiometry (Droop 1987) demonstrate that the majority of the ilmenites analyzed do not contain ferric iron and hence show no solid solution toward hematite. Maximum calculated Fe203 contents reach 5.6 wt % for an ilmenite from Lace. Because of the paucity of MgO, the majority of ilmenites examined exhibit no solid solution toward geikielite (MgTi03). The maximum geikielite content found was 11 mol % for a sample from Lace.
Manganese-rich ilmenites are not typically associated with either kimberlites or lamproites (see below), although they are commonly found in carbonatites (Figure 2.104).
240 CHAPTER 2
2.10.1. Comparison with Groundmass IImenites from Kimberlites
Groundmass ilmenite typically forms anhedral crystals. These are intergrown with spinels and/or perovskites, or occur as inclusions in olivines (Mitchell 1986, Shee 1984). The majority of groundmass ilmenites are probably microcrysts (see 1.2). Euhedral prisms are very rare and have been reported only from the Liqhobong (Boctor and Boyd 1980) and Lepelaneng (Haggerty 1975) kimberlites.
The few compositional data available (Shee 1985, 1984, Apter et al. 1984, Agee et al. 1982, Pasteris 1980, Boctor and Boyd 1980) indicate that kimberlite groundmass ilmenites are members of the ilmenite-geikielite series (50-90 mol % MgTi03) containing 5-10 mol % hematite (Figure 2.105). Cr203 contents are typically high (1-7 wt %), and MnO contents are characteristically low «1 wt % MnO).
Mn-bearing ilmenites are known from some kimberlites (Figure 2.104), and Mitchell (1986) has described a "manganese-enrichment" trend characterized by an increase in MnO (1-5 wt %) at the margins of Mg-ilmenite megacrysts. This paragenesis is clearly different from that of primary groundmass ilmenite and represents an attempt by megacrysts to equilibrate with carbonate-rich late-stage groundmass fluids (Haggerty et al. 1979, Boctor and Meyer 1979). Extreme examples of this trend are known from the Premier kimberlite where megacryst ilmenites are strongly, continuously zoned from MgO-rich cores to MnO-rich (up to 9 wt %) margins (Wyatt 1979, Gaspar and Wyllie 1984). Ilmenites in the late-stage carbonate-rich dikes at Premier (Gaspar and Wyllie 1984) contain similar amounts ofMnO (12-18 wt %) to those of groundmass ilmenites in orangeites, but may be easily distinguished from the latter on the basis of their higher MgO (1-5 wt %) and Fe203 (2-5 wt %) contents.
coexisting with other MnO-bearing (2.5-5.5 wt %) ilmenites, forming complex intergrowths with Mg-ilmenite, spinel, and perovskite in the De Beers kimberlite. This occurrence is not typical or characteristic of kimberlites.
The Nb contents of ground mass ilmenites in kimberlite are inadequately known. Ilmenite in the Premier calcite kimberlite contains 1.2-2.8 wt% Nb20S (Gaspar and Wyllie 1984).
Figure 2.10S. Compositional fields (mol %) of i1menites from orangeites [0] (this work). lamproites [L] (Mitchell and Bergman 1991). and groundmass ilmenites from kimberlites [K] (Shee 1984). plotted in the ternary system hematite (Fe203)-ilmenite (FeTi03)-geikielite (MgTi03).
MINERALOGY OF ORANGEITES 241
In summary, Mn-bearing ilmenites originating from kimberlites or orangeites may be distinguished from each other on the basis of their MgO, Cr203, Fe203, and MnO content and, to a lesser degree, on their paragenesis.
2.10.2. Comparison with I1menites in Lamproites
Ilmenite is not a characteristic mineral of most lamproites. It is known from the Jumilla (Murcia-Almeria), Sisco (Corsica), and Oscar, Mount North, Rice Hill, and other West Kimberley intrusions, where it occurs in the groundmass as subhedral prisms. Ilmenite is apparently absent from extrusive lamproites (Mitchell and Bergman 1991).
Lamproite ilmenites have significant MgO (1-7.5 wt %) and relatively low MnO (0.5-2 wt %) contents and are FeO-rich members of the ilmenite-geikilite series (>70 mol % FeTi03; Figure 2.1 05). Recalculation of the compositions shows that the ilmenites are poor in Fe203 «0.1-9 wt %) and typically contain less than 15 mol % hematite (Figure 2.1 05). Cr203 and Ah03 are typically less than 0.5 wt %.
The only lamproite-derived ilmenites known to contain significant amounts ofMnO (3-8 wt %) are found in carbonate in olivine lamproites from Argyle (West Australia). These are MgO poor «1 wt %) and contain up to 1.2 wt % Nb205 (Jaques et al. 1989a).
In summary, lamproite-derived ilmenites may be distinguished from those in orangeites on the basis of their MgO and MnO contents (Figures 2.104 and 2.105). In both rock types, ilmenite occurs as partially resorbed late-stage groundmass prisms.
2.11. RUTILE
Rutile is present as a trace accessory groundmass mineral in orangeites. It has been recognized in the Lace, Sover Mine, Bellsbank Southern Extension, and Swartruggens orangeites (this work). At Lace it occurs as small «50 !lm) anhedral crystals, which may be present as single grains, commonly with a "doughnut-like" habit, or intimately intergrown with anhedral ilmenite, monazite, or barite. At Bellsbank, rutile typically occurs as irregular intergrowths with anhedral apatite and monazite. Discrete rounded crystals occur at Sover Mine. Rutile is rare in the Swartruggens orangeites and occurs as irregular grains and/or very thin laths «1 lim) in segregations consisting principally of an unidentified Ca-Zr silicate, quartz, and apatite.
Table 2.45 shows that the rutiles contain significant amounts ofFe203 (0.5-1.2 wt % at Lace, 0.9-1.6 wt % at SoverMine, 1.8-4 wt % at Bellsbank) and Nb20s (1.1-1.9 wt %
Table 2.45. Representative Compositions of Rutilea
"Total Fe expressed as Fez03; n.d. = not detected. Compositions 1-3. Sover Mine; 4-6 Lace; 7, Bellsbank (this work).
242 CHAPTER 2
at Lace, 3.2-5.4 wt % at Sover Mine, 3-4.9 wt % at Bellsbank) in solid solution (this work). Note that coexisting groundmass i1menites at Sover Mine are also enriched in Nb relative to i1menites from Lace. Chromium contents are very low «0.1 wt % Cr203).
Comparison with Rutile in Kimberlites and Lamproites
Rutile occurs in six parageneses in kimberlites as (1) associated with microphenocrystal olivines, (2) a primary groundmass phase, (3) reaction mantles upon perovskite, (4) large discrete crystals (?macrocrysts) mantled by Mg-ilmenite, (5) graphic or symplectic rutile-silicate intergrowths, (6) lamellar or sigmoidal rutile-ilmenite intergrowths. Parageneses 5 and 6 are not found in orangeites and hence not discussed further in this work.
Needle-like and tabular crystals associated with olivine (Mitchell 1986, Skinner 1989, Pasteris 1980) are the most common paragenesis of rutile encountered in kimberlites. These may occur as inclusions confined to the margins of olivine microphenocrysts or be present as numerous small crystals in the groundmass adjacent to such olivines. In the latter case, rutile is commonly associated with spinels and/or perovskites in a "necklace-like" texture. These rutiles are high Cr203 (up to 2.5 wt %), low Fe203 «0.5 wt %) varieties. Nb contents are unknown but may be appreciable, i.e., 1-5 wt % Nb20s (Pasteris 1980).
Skinner (1989) initially noted that there are no counterparts in texture and composition to paragenesis 1 rutile in orangeites. Although chromite occurs as inclusions in orangeite olivine microphenocrysts, rutile, perovskite, and ilmenite are absent.
Primary groundmass rutile is known from the lower Benfontein sill (Mitchell 1994b, Boctor and Boyd 1981). Here it occurs as discrete twinned euhedral (1 (}-50 11m) crystals and anhedral-to-subhedral crystals intergrown with the corroded Ti-Mg magnetite or Ti-magnetite cores of atoll-textured spinels. These intergrowths are mantled by Mn-bearing, Mg-poor ilmenite, and magnetite. Both varieties of rutile may be replaced by parisite, ancylite, and ferroan dolomite. The rutile contains up to 3.3 wt % Nb20S and 1-3 wt % Fe203 (Mitchell 1994b). Orangeites do not contain rutiles belonging to this paragenesis.
Paragenesis 3 has not yet been observed in orangeites. However, in kimberlites complete replacement of perovskite leads to the development of discontinuous doughnut-like rings of rutile which are ultimately disaggregated and strewn throughout the groundmass. Such a process may account for the irregular rutile grains found at Lace. However, the absence of precursor perovskite':"ilmenite composite grains suggests this rutile may have other origins.
Anhedral fragments of paragenesis 4 rutile may occur in the groundmass of kimberlites. These may be distinguished from groundmass rutiles by their characteristic mantles ofMg-ilmenite, perovskite, and/or spinel. Anhedral rutile crystals in orangeites lack such mantles and are unlikely to be of similar origin.
Few data are available regarding the paragenesis and composition of rutile in lamproites, and the majority of lamproites studied lack rutile (Mitchell and Bergman 1991). Rutile is relatively common only in the anomalous Sisimiut lamproites. From these, Thy et at. (1987) describe rutile as a groundmass phase intergrown with ilmenite in amphibole lamproites and as the only opaque groundmass phase in leucite lamproites.
MINERAWGY OF ORANGEITES 243
Rutile from Sisimuit contains up to 2.2 wt % Cr203 and from 0.1-11 wt % Fe203 (Thy et al. 1987, Scott 1981). The Nb content of lamproite rutile is not known.
Secondary anatase is common in some lamproites but has not yet been recognized in orangeites.
In summary, rutile found in orangeites may be distinguished from that found in kimberlite on the basis of its paragenesis and Cr and Nb content. Rutile is not a characteristic or common mineral in lamproites.
2.12. ZIRCONIUM SILICATES
Zirconium silicates are relatively common as late-stage groundmass minerals in orangeites. Unfortunately, they have not been extensively studied and few paragenetic and compositional data are available.
2.12.1. Zircon
Zircon occurs in the Swartruggens, Lace, and Bellsbank Southern Extension orangeites as very small «10 ~m) irregular crystals set in a carbonate mesostasis (this work). At Lace, zircon occurs rarely as small patches (20-50 I!m) of very small «5I!m) euhedral, flow-aligned, prismatic crystals associated with subhedral wadeite crystals. Compositional data are not available.
2.12.2. Wadeite
A potassium zirconium silicate, provisionally considered to be wadeite (K2ZrSh09), has been identified by BSEJEDS methods in the Swartruggens and Lace orangeites (this work). At Swartruggens it occurs rarely as euhedral crystals (50 I!m) in orangeites from level 6 in the mine. These and associated unidentified Ca-Ti-Fe silicate (2.14), strontianite, and ancylite are set in a matrix of calcite. Small euhedral crystals of perovskite are poikilitically enclosed by the wadeite. At Lace, wadeite forms small subhedral crystals associated with zircon.
Table 2.46 gives representative compositions (WDS-EMPA) of Swartruggens wadeite and demonstrates that they are not significantly different from those of wadeites occurring in lamproites. Wadeites from both parageneses are essentially pure potassium zirconium silicates. With the exception of Ti02, other elements are not present in significant quantities. These data are considered to indicate that the K-Zr silicate from Swartruggens is probably wadeite. However, without X-ray diffraction data, the presence of kostylevite (monoclinic K2ZrSh09 . H20; Khomyakov et al. 1983a) or umbite (orthorhombic KzZrSh09 . H20; Khomyakov 1983b) cannot be ruled out.
Wadeite is considered by Mitchell and Bergman (1991) to be one of the characteristic minerals of lamproites. Although wadeite is not ubiquitous in lamproites, it is easily recognized by standard petrographic methods as it commonly forms large crystals occurring in significant modal amounts.
Wadeite is a rare mineral that, with the exception of lamproites, has been previously recognized only from two other parageneses, i.e., late-stage carbonatite-like rocks from the Khibina (Tikhonenkov et al. 1960) and Kovdor (Kapustin 1980) complexes and in
"Total Fe expressed as Fel03 or FeO; n.d. = not detected; n.a.= not analyzed. Compositions 1-5. wadeite; 1-2. Swartruggens orangeite (this work); 3. West Kimberley larnproite (Jacques et al. 1986); 4. Kovdor carbonatite (Kapustin 1980); 5. Little Murun tausonite syenite (Mitchell and Vladykin 1993); 6. Kimzeyitic gamet. New Elands (Mitchell and Meyer 1989a); 7. schorlomitic garnet. Burns (Mitchell and Meyer 1989a); 8, Kimzeyitic garnet. Wesselton (Mitchell I 994b); 9-10. unidentified Ca-Zr silicate. Skietkop (this work).
tausonite syenites from the Little Murun uItrapotassic complex (Mitchell and Vladykin 1993).
2.12.3. Zirconium-Bearing Gamet
Zirconium-rich garnets occur as brown, small (l00-500 J.lm) euhedral isotropic crystals in the groundmass of the New Elands orangeites (Mitchell and Meyer 1989a). The garnets lack reaction rims and resorption features and are thus considered to be primary phases which have crystallized in eqUilibrium with the calcite-rich groundmass.
Table 2.46 shows the garnets to have high Ti02 and Fe203 and low Ah03 contents, suggesting they are members of a solid solution series between schorlornite and kimzeyite. Ti-rich garnets with very low Zr02 contents found in the Bums orangeite (Mitchell and Meyer 1989a) represent the schorlornitic end member (Table 2.46). Schorlomitic garnets containing up to 32 wt % Ti02 have been reported in calcite-serpentine segregations in a Postmasburg area orangeite (Tainton 1992). Schorlomitic garnets are rarely found at Newlands (Tainton 1992).
Kimzeyitic garnets are not characteristic minerals of archetypal kimberlites (Mitchell 1986) and have never been found in lamproites (Mitchell and Bergman 1991). The only reported occurrence of Ti-Zr-rich garnet in a bona fide kimberlite is from a highly differentiated calcite kimberlite siIl at WesseIton (Mitchell 1994b). This garnet occurs as small (30-50 J.lm) anhedral groundmass crystals in samples lacking baddeleyite and calcium zirconate. The garnet is richer in Zr02 than the Zr-bearing garnets from the New Elands orangeite (Table 2.46).
Zr- and Ti-rich garnets are common in alnoites and ultramafic lamprophyres associated with carbonatite complexes (Platt and Mitchell 1979, Mitchell 1983). A very wide range of compositions has been reported within the series andradite-schorlornite-
MINERALOGY OF ORANGEITES 245
kirnzeyite. These encompass the compositions described above. Kirnzeyitic garnets from this paragenesis commonly coexist with monticellite and melilite.
2.12.4. Calcium Zirconium Silicate
An unidentified calcium zirconium silicate occurs rarely in the groundmass of the Swartruggens and Skietkop orangeites. At Swartruggens it forms ragged, irregular, small (1-5 !lm) crystals intergrown with rutile and apatite, set in a matrix of quartz. Compositional data are not available.
In the Skietkop occurrence the mineral forms granular aggregates of small (10-50 !lm) anhedral crystals within a calcite matrix. Semiquantitative analyses indicate 39.1-40.7 wt % Si02, 28.8-30.3 wt % Zr02, 15.1-16.4 wt % CaO, 9.1-10.9 wt % MgO, and 3.7-4.5 wt % FeOT. The mineral contains insufficient CaO and Si02 to be either Zr-Ti garnet or calcium catapleite, respectively.
2.13. CARBONATES
Carbonates are common late-stage groundmass minerals in orangeites with modes ranging from <1-13 vol % (Skinner and Clement 1979, Fraser 1987). Detailed investigations of their composition and paragenesis have not yet been undertaken.
2.13.1. Calcite
Calcite is the commonest carbonate found in orangeite. It typically occurs as anhedral patches of interlocking crystals intimately intergrown with groundmass serpentine. Less commonly, it forms the margins of calcite-serpentine segregations. Calcite may replace serpentinized olivine macrocrysts and phlogopite. Despite the ubiquity of calcite, compositional data have not been reported in previous studies of orangeite. Limited data obtained during this study show that it is typically pure CaC03 containing less than 0.2 wt % MgO, FeO, SrO, and MnO.
2.13.2. Dolomite
Dolomite is a relatively common groundmass constituent of the Bellsbank and Newlands orangeites (this work, Tainton 1992). The groundmass of the Main and Bobbejaan dikes consists of a serpentine-dolomite intergrowth. The dolomite forms poikilitic grains (0.6 mm). With increasing degrees of deuteric alteration, dolomite/serpentine ratios increase until all of the serpentine is replaced (Tainton 1992). In the Bobbejaan dike the matrix consists entirely of dolomite. Similar dolomite/serpentine relationships have been observed at Newlands, where fresh samples contain more serpentine than dolomite. The abundance of dolomite increases with increasing degrees of deuteric alteration (Tainton 1992). At Newlands, dolomite crystallization occurs before calcite (Tainton 1992).
The dolomites of the Bellsbank Southern Extension orangeite contain rounded patches of (?)exsolved norsethite (this work, see below). The dolomite is an Sr-Mn-bearing ferroan dolomite (Table 2.4 7). No other compositional data for dolomite in orangeites have been published.
246 CHAPTER 2
Table 2.47. Representative Compositions of Carbonatesa
aFeOr = total Fe expressed as FeO; n.d. = not detected. n.a. = not analyzed. Compositions 1-2 obtained by EMPAlWDS analysis, 3-9 by SEMIEDS methods and expressed as oxides summed to 100 wt % on a C02-free basis. Compositions 1-2, norsethite, Bellsbank; 3-5, ferroan dolomite, Bellsbank; 6-7, ancylite, Bellsbank, and Swartruggens, respectively; 8-9, strontianite and calcian strontianite, Swartruggens. (All data this work.)
2.13.3. Other Carbonates
Other carbonates present in orangeites include strontianite (SrC03), barian strontianite, witherite (BaC03), unidentified Ba-Ca carbonate (?barytocalcite), an Sr-rare earth carbonate (Table 2.47), considered to be ancylite [Sr(REE)(C03h(OH)· H20], and norsethite [BaMg(C03h].
The majority of these carbonates occur as irregular crystals within the groundmass and are intimately intergrown with calcite, dolomite, serpentine, and barite. Commonly, they replace earlier crystallizing minerals, e.g., barian strontianite replaces apatite at Lace; witherite replaces sanidine at Sover North and calcite at Sover Mine; ancylite replaces perovskite at Makganyene and Swartruggens. This petrographic evidence suggests that the majority of these carbonates are late-stage hydrothermalldeuteric minerals. Norsethite (Table 2.47) is common in the Bellsbank Southern Extension orangeite as rounded droplets within a ferroan dolomite matrix. The norsethite is considered to have exsolved from a complex dolomitic carbonate precursor as it is only found as inclusions within ferro an dolomite (this work).
Carbonates are major components of archetypal kimberlites. The parageneses of calcite and dolomite in these rocks (Mitchell 1986) are similar to those of orangeite. The principle differences are that kimberlites contain more carbonate, and calcite-serpentine segregations are common. Insufficient compositional data are available to permit useful comparisons of carbonate compositions in the two rock types. Surprisingly, detailed studies of the carbonate assemblages in kimberlites have not been undertaken; thus, it is not known whether or not Sr-, Ba-, and REE-rich carbonates are common. Ancylite and Ca-Ba carbonate have been reported in highly-evolved calcite kimberlites from the Benfontein sills (Mitchell 1994b). Preliminary studies of Somerset Island kimberlites have not revealed the presence of any Sr-Ba-REE carbonates.
MINERALOGY OF ORANGElTES 247
Mi tchell and Bergman (1991) have noted that carbonates are not characteristic groundmass minerals of lamproites. Groundmass calcite, when present, appears to result from post-emplacement secondary carbonatization. The majority of lamproitic carbonates occur in vesicles. Calcite, witherite, calcian strontianite, and strontian calcite have been noted in vesicles (Mitchell and Bergman 1991, Wagner and Velde 1986).
2.14. OTHER MINERALS
Barite is a common accessory mineral in most orangeites and occurs as irregularly shaped aggregates and veins of anhedral crystals throughout the groundmass or replacing earlier crystallizing minerals. Intergrowths of calcite and barite are common. Barite is considered to be a late-stage groundmass and deuteric hydrothermal mineral. Qualitative analysis indicates that many barites contain significant quantities of Sr.
The groundmass of the Swartruggens orangeites contains an unidentified euhedral purplish brown mineral occurring in matrices which include barite, calcite, and quartz. The habit is hexagonal, and the mineral exhibits first-order interference colors. The average (wt %) of four WDS-EMP analyses is 30.3% CaO, 26.2% Si02, 21.4% Ti02, 14.0% FeOr, 1.6% MgO, and 0.65% Na20 (94.5% total oxides). AI, Zr, and Ba were not detected.
Chalcopyrite and galena have been identified in the Besterskraal, Sover North, Lace, Bellsbank, and Swartruggens orangeites. Pyrite is common as interstitial late-crystallizing laths and plates at Voorspoed and Swartruggens. Fe-Ni sulfides are present in the groundmass at Postmasburg, Sover Mine, and Bellsbank. Ni-sulfides are common in serpentinized olivine macrocrysts in most orangeites. Small «111m) rounded crystals of FeAs2 (?loellingite) occur rarely in the groundmass of the Sover Mi ne orangeite.
Quartz is a common component of the mesostasis of the Swartruggens and the Sweetput-Souput orangeites. It is unclear whether the quartz is a late-stage primary phase or a secondary mineral. Ba- and Na-zeolites, replacing feldspars, occur in the groundmass of the Besterskraal and Swartruggens rocks.
2.15. SUMMARY
The above mineralogical studies demonstrate conclusively that the rocks here referred to as orangeites, previously termed "micaceous or group 2 kimberlites," have little mineralogical affinity with archetypal or "group 1" kimberlites. Notable differences exist with respect to the nature of the composition and evolutionary trends of mica compositions. Kimberlite micas are characterized by trends of Al enrichment, leading to the formation of barian aluminous phlogopite-:kinoshitalite solid solutions. In contrast, orangeite micas exhibit trends of Al depletion, leading to the formation of tetraferriphlogopite. Orangeites and kimberlites differ with respect to the nature of the primary silicate phases. Kimberlites crystallize monticellite and serpentine but not diopside, potassic amphibole, and potassium feldspar. The assemblages and compositional trends of iron and titanium oxides are different in the two rock types. Spinels rich in the magnesian ulvospinel molecule appear to be restricted to kimberlites. Spinels in orangeites are very
248 CHAPTER 2
Cr-rich and AI-poor relative to those in kimberlites. Perovskites in kimberlites are poor in REE and Sr relative to those in lamproites.
Although orangeites have few mineralogical characteristics in common with kimberlites, they have much closer affinities to lamproites. Micas follow similar evolutionary trends toward tetraferriphlogopite. The assemblages differ in that Ti-rich micas are not preSent in orangeites. Primary diopside, potassic amphibole, and spinel group minerals haVie similar compositions and evolutionary trends in both rock types. Orangeites and larriproites share the characteristic of containing primary AI-poor silicates which commonly contain insufficient Al to occupy all of the available tetrahedral lattice sites in the crystal structure. This evidence suggests that orangeites also crystallized from peralkaline parental magmas. Lamproites and orangeites differ primarily in the nature and composition of the accessory mineral assemblage. Hollandites in lamproites are principally Cr-bearing K-Ba septetitanates, whereas those in orangeites are V-bearing Ba-rich hexatitanates. K-triskaidecatitanates do not occur in lamproites. Orangeites contain a greater variety of Sr-, Ba-, and REE-bearing phosphates than most lamproites. The affinities with lamproites are discussed further in Section 4.7.2.
