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Miner Petrol (2016) 110:333–360
DOI 10.1007/s00710-015-0390-6
ORIGINAL PAPER
Calcite and dolomite in intrusive carbonatites. I.
Textural variations
Anton R. Chakhmouradian 1 & Ekaterina P. Reguir 1 &
Anatoly N. Zaitsev 2,3
Received: 9 December 2014 / Accepted: 22 May 2015 / Published online: 27 June 2015
# Springer-Verlag Wien 2015
Abstract Carbonatites are nominally igneous rocks, whose
evolution commonly involves also a variety of postmagmatic
processes, including exsolution, subsolidus re-equilibration of
igneous mineral assemblages with fluids of different provenance, hydrothermal crystallization, recrystallization and tectonic mobilization. Petrogenetic interpretation of carbonatites
and assessment of their mineral potential are impossible without understanding the textural and compositional effects of
both magmatic and postmagmatic processes on the principal
constituents of these rocks. In the present work, we describe
the major (micro)textural characteristics of carbonatitic calcite
and dolomite in the context of magma evolution, fluid-rock
interaction, or deformation, and provide information on the
compositional variation of these minerals and its relation to
specific evolutionary processes.
Introduction
Carbonate minerals are the principal constituent of intrusive
carbonatites; their content ranges from 50 modal %, which is
accepted as a nominal threshold for this rock type (Le Maitre
Editorial handling: L. G. Gwalani
* Anton R. Chakhmouradian
[email protected]
1
Department of Geological Sciences, University of Manitoba,
Winnipeg, Manitoba R3T 2 N2, Canada
2
Department of Mineralogy, St. Petersburg State University, St.
Petersburg 199034, Russia
3
Department of Earth Sciences, Natural History Museum, Cromwell
Road, London SW7 5BD, UK
2002) to well over 90 % in some varieties interpreted as cumulates (e.g., Xu et al. 2007). Surprisingly, however, there
appears to have been no attempt in the literature to provide a
comprehensive analysis of textures exhibited by these minerals. Some studies addressed textural variations within a series of genetically related rocks (e.g., Platt and Woolley 1990;
Zaitsev 1996; Zaitsev et al. 2004, 2008; Chakhmouradian
et al. 2015a), i.e. on the scale of a single locality, but none
approached this task in a systematic fashion and on the basis
of multiple observations from many carbonatite occurrences.
Correct interpretation of carbonate textures is critical not only
to tackling purely petrogenetic problems (e.g., the relationship
between carbonatites and spatially associated silicate rocks),
but also for such practical tasks as orebody delineation, for
example. On many occasions in the past, misunderstanding of
certain petrographic characteristics of carbonatites has led to
misinterpretation of their origin and mineral potential. The
present contribution is a brief summary of some basic petrographic observations that will hopefully help our readers avoid
repeating these mistakes in their research or exploration
efforts.
Approximately 60 % of intrusive carbonatites worldwide
are predominantly calcitic; most of the remaining 40 % comprise members of the dolomite-ankerite series. The latter cover
a compositional range from nil to ~23 wt.% FeO or 70 mol.%
CaFe(CO3)2 (Table 1), i.e. approach the solubility limit
established for these minerals empirically (for detailed discussion, see Reeder and Dollase 1989). However, ankerite sensu
stricto [>18 wt.% FeO or 50 mol.% CaFe(CO3)2] appears to
be restricted to late-stage postmagmatic parageneses. The
high-pressure CaCO3 polymorph aragonite, occurring as
abundant phenocrysts in a calcitic matrix, has so far been
described only from Hajnáčka in Slovakia (Hurai et al.
2013). At lithospheric pressures, mantle-derived melts are
generally too hot for stable aragonite crystallization (cf.
334
A.R. Chakhmouradian et al.
Table 1 Composition of calcite and dolomite from carbonatites: selected data from the literature, showing the extent of element substitution, and from
the present work
wt.%
1
2
3
4
5
6
7
8
9
10
11
4.29
25.60
0.97
27.01
n.d.
3.98
21.52
n.d.
6.30
22.84
0.63
8.11
0.12
2.22
61.69
n.d.
57.34
n.d.
57.73
MgO
CaO
2.79
55.97
4.31
47.86
1.05
54.14
0.25
47.64
0.11
50.32
0.05
52.53
1.58
49.95
0.50
43.12
n.d.
53.17
MnO
0.07
0.32
0.30
5.69
5.36
3.98
0.44
4.14
FeO
SrO
0.28
0.62
2.37
0.47
0.74
0.52
1.75
0.92
0.77
1.10
0.28
1.41
4.99
0.03
0.24
13.12
BaO
Total
0.23
55.96
n.d.
55.33
n.d.
56.75
n.d.
56.25
0.04
57.70
0.10
58.35
n.d.
56.99
0.88
62.00
Formulae (calculated to one cation for calcite and two cations for dolomite)
Mg
Ca
0.064
0.924
0.107
0.851
0.026
0.955
0.006
0.877
0.003
0.900
0.001
0.926
0.039
0.886
0.013
0.788
–
0.927
0.226
0.968
0.052
1.043
Mn
0.001
0.004
0.004
0.083
0.076
0.055
0.006
0.060
–
0.119
0.657
Fe
0.004
0.033
0.010
0.025
0.011
0.004
0.069
0.003
–
0.674
0.245
Sr
Ba
0.006
0.001
0.005
–
0.005
–
0.009
–
0.010
–
0.013
0.001
–
–
0.130
0.006
0.059
0.014
0.013
–
0.003
–
wt.%
MgO
12
20.66
13
14.98
14
12.83
15
18.55
16
12.48
17
0.84
18
n.d.
19
0.47
20
0.36
21
n.d.
22
n.d.
CaO
MnO
FeO
27.34
1.17
0.39
29.23
0.51
9.52
29.65
0.67
12.14
29.23
0.40
4.83
31.06
0.40
11.01
49.81
1.31
1.61
56.13
0.16
0.07
54.81
0.26
0.83
55.60
0.06
0.22
55.03
0.01
0.05
55.42
0.02
0.03
SrO
BaO
Total
3.88
n.d
53.44
0.23
n.d.
54.47
0.18
n.d.
55.47
n.d.
n.d.
53.01
n.d.
n.d.
54.95
2.70
0.17
56.44
0.28
n.d.
56.64
0.85
0.07
57.29
0.23
n.d.
56.47
1.64
0.15
56.88
1.34
0.07
56.88
Formulae (calculated to one cation for calcite and two cations for dolomite)
Mg
0.968
0.719
0.620
0.873
0.606
Ca
0.920
1.007
1.030
0.989
1.083
Mn
0.031
0.014
0.018
0.011
0.011
Fe
0.010
0.256
0.329
0.127
0.300
Sr
Ba
wt.%
0.071
–
23
0.004
–
24
0.003
–
25
0.021
0.909
0.019
0.023
–
0.994
0.002
0.001
0.012
0.965
0.004
0.011
0.009
0.985
0.001
0.003
–
0.982
–
0.001
–
0.986
–
–
0.002
–
–
0.027
0.016
0.013
–
26
–
27
0.001
28
–
29
–
30
–
31
0.001
32
0.001
33
0.003
0.008
MgO
n.d.
n.d
23.14
18.88
16.21
CaO
54.88
55.98
27.62
29.57
29.22
MnO
0.15
n.d
0.34
0.52
0.62
FeO
n.d.
0.03
0.88
3.37
7.90
SrO
1.06
0.05
0.06
0.48
0.60
BaO
n.d.
n.d.
n.d.
n.d.
n.d.
Total
56.09
56.06
52.04
52.82
54.55
Formulae (calculated to one cation for calcite and two cations for dolomite)
Mg
–
–
1.059
0.888
0.768
Ca
0.988
1.000
0.908
1.000
0.994
Mn
0.002
–
0.009
0.014
0.017
Fe
–
–
0.023
0.089
0.210
Sr
0.010
–
0.001
0.009
0.011
Ba
–
–
–
–
–
wt.%
34
35
36
37
38
MgO
n.d.
n.d.
n.d.
0.17
0.34
20.15
29.97
0.76
1.50
0.46
n.d.
52.84
18.27
29.46
0.76
4.43
0.28
n.d.
53.20
n.d.
51.08
0.07
0.04
5.98
1.14
58.31
n.d.
53.60
0.10
0.02
3.99
0.25
57.96
n.d.
51.53
0.32
0.12
5.76
0.53
58.26
n.d.
52.93
0.28
0.12
4.48
0.30
58.11
0.934
0.999
0.020
0.039
0.008
–
39
1.01
0.861
0.997
0.020
0.117
0.005
–
40
0.32
–
0.932
0.001
–
0.059
0.008
41
15.28
–
0.958
0.001
–
0.039
0.002
42
18.01
–
0.934
0.005
0.002
0.056
0.003
43
20.70
–
0.949
0.004
0.002
0.043
0.002
44
7.62
CaO
MnO
54.72
n.d.
56.19
n.d
29.15
0.76
29.67
0.23
29.66
0.70
26.26
0.32
55.85
0.35
56.58
n.d.
56.12
n.d.
54.85
0.30
56.34
n.d.
Textural variations in carbonatites
335
Table 1 (continued)
FeO
0.19
n.d.
n.d.
0.13
0.02
n.d.
SrO
0.23
0.04
0.15
1.26
0.08
0.37
BaO
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
Total
56.62
56.62
56.27
56.71
56.78
56.10
Formulae (calculated to one cation for calcite and two cations for dolomite)
Mg
–
–
–
0.004
0.008
0.025
Ca
0.990
1.000
0.999
0.978
0.991
0.971
Mn
0.005
–
–
0.004
–
–
Fe
0.003
–
–
0.002
–
–
Sr
0.002
–
0.001
0.012
0.001
0.004
Ba
–
–
–
–
–
–
0.03
n.d.
n.d.
56.51
0.008
0.992
–
–
–
–
7.70
0.53
n.d.
53.42
0.742
1.017
0.021
0.210
0.010
–
4.40
n.d.
n.d.
52.31
0.961
0.990
0.018
0.025
0.006
–
0.94
0.31
n.d.
52.31
0.961
0.990
0.018
0.025
0.006
–
20.44
n.d.
n.d.
56.64
0.400
0.990
0.009
0.601
–
–
Analyses 1, 4, 7–9, 11 and 12 are from the literature; the rest were obtained in this study using wavelength-dispersive spectrometry with a Cameca SX
100 automated electron microprobe operated at 15 kV and 10 nA with a 10 μm beam; typical detection limits are 300–400 ppm for Mg, 200–300 ppm
for Ca, and 600–700 ppm for Mn, Fe, Sr and Ba
Locations and other details: (1) Mg-rich calcite associated with exsolved dolomite, Phalaborwa, South Africa (Dawson and Hinton 2003); (2,3) Mg-rich
calcite associated with exsolved dolomite, Goldray, Canada (Fig. 9a-c); (4) Mn-rich calcite, Khibiny, Russia (Zaitsev 1996); (5,6) coarse- and finegrained Mn-rich calcite, Bear Lodge, Wyoming (Fig. 6a); (7) Fe-rich calcite from Mud Tank, Australia (Currie et al. 1992); (8) Sr-rich calcite, SarnuDandali, India (Wall et al. 1993); (9) Sr-Ba-rich calcite, Murun, Russia (Konev et al. 1996); (10) most Fe-rich member of dolomite-ankerite-kutnohorite
series from carbonatites, Bearpaw Mts., Montana; (11) most Mn-rich member of dolomite-ankerite-kutnohorite series from carbonatites, Khibiny
(Zaitsev 1996); (12) Sr-rich dolomite, Khibiny (Zaitsev 1996); (13,14) inner and outer zones in a dolomite reaction mantle around a glimmerite xenolith,
Chipman Lake, Canada (Fig. 4g); (15,16) inner and outer zones in a dolomite reaction mantle around a glimmerite xenoliths, Aley, Canada (Fig. 4h);
(17,18) core and rim of a calcite phenocryst, Kontozero, Russia (Fig. 5a); (19,20) core and rim of a calcite crystal, Carb Lake, Canada (Fig. 5b); (21,22)
core and rim of a calcite lath, Kaiserstuhl, Germany (Fig. 5c)
Locations and other details: (23,24) core and rim of a calcite crystal, Afrikanda, Russia (Fig. 5d); (25,26) core and rim of a dolomite crystal, Chipman Lake
(Fig. 5e); (27) interstitial dolomite, Chipman Lake (Fig. 5e); (28,29) core and rim of a dolomite crystal, Aley (Fig. 5f); (30,31) core and rim of a calcite
crystal, Murun (Fig. 9f); (32–34) core, intermediate zone and rim core and rim of a calcite crystal, Bearpaw Mts., Montana (Fig. 9g); (35,36) core and rim of
a late-stage hydrothermal calcite crystal, Afrikanda (Fig. 12f); (37) primary calcite, Mountain Pass, California (Fig. 7c); (38,39) low-AZ and high-AZ zones
in late-stage hydrothermal calcite crystal, Mountain Pass (Fig. 12d); (40) hydrothermal spherulitic calcite, Iron Hill, Colorado (Fig. 12c); (41,42) primary
core and hydrothermal rim of a dolomite crystal, Aley (Fig. 12h); (43) primary dolomite core mantled by (44) hydrothermal ankerite, Aley (Fig. 12i)
Humphreys et al. 2010). Given the unusual crystallization
conditions of the Hajnáčka rock (interpreted to have formed
from overpressurized magma in the calcite stability field:
Hurai et al. 2013), and the lack of any petrographic or chemical evidence of aragonite-to-calcite conversion (Carlson and
Rosenfeld 1981; Theye and Seidel 1993; Keiter et al.
2008) at other localities, aragonite clearly has limited
significance for carbonatite petrogenesis. Although this
mineral can potentially develop at the expense of igneous calcite in collision zones, its preservation will require exhumation rates and fluid regime that are rarely
attained in the natural environment (Ghent et al. 1996;
Huang 2003). Other carbonate minerals that may locally
gain the rock-forming status include the magnesitesiderite-rhodochrosite series, (Mg,Fe,Mn)CO3 (Buckley
and Wooley 1990; Zaitsev 1996; Thompson et al.
2002; Zaitsev et al. 2004); kutnohorite, Ca(Mn,
Fe)(CO 3 ) 2 (Zaitsev 1996); fluorocarbonates of light
rare-earth elements (REE) and burbankite, (Na,Ca)3(Sr,
Ca,Ba,REE)3(CO3)5 (Zaitsev et al. 1998; Castor 2008;
Kynicky et al. 2013); strontianite, SrCO 3 , and
barytocalcite, CaBa(CO 3) 2 (Konev et al. 1996). The
present work focuses on calcite and dolomite because
these minerals account for the bulk of terrestrial igneous
carbonate material.
Phase and compositional relations of significance
to rock-forming carbonates in carbonatites
At ambient conditions, calcite forms only limited solid solutions with isostructural Mg, Mn and Fe2+ carbonates, and with
aragonite-type SrCO3, BaCO3 and PbCO3 (Chang 1971;
Chang and Brice 1972; Brice and Chang 1973; De Capitani
and Peters 1981). Binary and some ternary systems of relevance to carbonatites, such as CaCO3–MgCO3–FeCO3 or
CaCO3–MgCO3–SrCO3, are understood reasonably well
(Brice and Chang 1973; Anovitz and Essene 1987). Below
the solidus, these systems exhibit solvi separating CaCO3
from dolomite- or barytocalcite-type ordered carbonates of
intermediate composition. The solvi are asymmetric and their
steep side faces the carbonate phase that incorporates a smaller
cation—e.g., CaMg(CO3)2 along the CaCO3–MgCO3 join
(Fig. 3 in Anovitz and Essene 1987). Primary dolomitic melts
equilibrated with mantle peridotites (<63 wt.% CaCO3 at P>27 kbar) are likely to react with mantle rocks during their
ascent and evolve toward more calcic compositions (up to
89 wt.% CaCO3 in the carbonate component of the liquid)
through metasomatic wehrlitization of the conduit (Dalton
and Wood 1993; Kogarko et al. 1995; Wyllie and Lee 1998).
The topology of the system CaCO3–MgCO3 changes dramatically with decreasing pressure; the changes involve a shift of
336
A.R. Chakhmouradian et al.
In the presence of water, dolomite melts incongruently at T
approaching the solvus to yield a hydrous melt, periclase and
CO2 (Persikov and Bukhtiyarov 2013). The scarcity of periclase in carbonatites implies that Mg-rich magnetite or Mg
silicates (at higher SiO2 activities) precipitate in place of this
mineral in natural systems (Reguir et al. 2008, 2012).
The above-discussed phase relations have several important
implications for carbonatite petrogenesis. First of all, calcite
can precipitate as an early liquidus phase from primary
mantle-derived carbonate melts (Wyllie and Lee 1998; our
Fig. 1a). Secondly, calcite precipitated from such melts should
exsolve dolomite upon reaching the solvus, i.e. Mg-poor
cumulus calcite common in carbonatites (<0.5 wt.% MgO,
see below) is unlikely to have crystallized directly from
primitive carbonatitic magmas, and is probably derived from
more evolved melts. Degassing (reaction 1) and fractionation
of magnetite and forsterite, for example, can increase the Ca/
Mg ratio of the melt sufficiently to facilitate the formation of
near-end-member CaCO3. Thirdly, the field of high-T Mg-rich
calcite above the solvus is much wider than the dolomite field.
Indeed, both published and our own electron-microprobe analyses (Table 1) show that the Ca content of carbonatitic dolomite does not deviate from the ideal stoichiometric value by
more than 0.09 atoms per formula unit (apfu), and remains
within 0.04 apfu of that value in ~90 % of the data. This
amount of variation is equivalent to ~4 wt.% CaCO3, which
correlates with the width of the dolomite intersolvus field in
Fig. 1a. At the same time, the proportion of MgCO3 that can be
incorporated in calcite is much greater. Up to 2.8 wt.% MgO
(equivalent to 5.8 wt.% or 6 mol.% MgCO3) has been reported
in the literature for grains associated with exsolved dolomite at
Phalaborwa, South Africa (Dawson and Hinton 2003), and in
the present work, we measured up to 4.3 wt.% MgO (9.0 wt.%
or ~11 mol.% MgCO3) in a paragenetically similar sample
from Goldray, Canada (Table 1). From these considerations,
it is clear that exsolution of dolomite from Mg-rich calcite will
Fig. 1 Diagrams illustrating phase relations among carbonate minerals
most relevant to carbonatite petrogenesis. a CaCO3–MgCO3 at P=
10 kbar (after Byrnes and Wyllie 1981; Anovitz and Essene 1987);
dashed arrows show a possible cooling path of primary mantle-derived
carbonatitic melt (shaded field, after Wyllie and Lee 1998) precipitating
Mg-rich calcite that undergoes exsolution to Mg-poor calcite + dolomite;
note that the liquidus and solidus temperatures will depend strongly on
the proportion of alkalis and volatiles in the melt. Cal=calcite, Dol=
dolomite, L = liquid, Mgs = magnesite, Prc = periclase, V = vapor. b
CaCO3–SrCO3–BaCO3 at T=550 °C and P=10 kbar (Chang 1971), with
intersolvus boundaries omitted for clarity; the compositions of hypothetical Ca-Ba-Sr Bprotocarbonates^ from Murun (Russia) and Jogipatti
(India) are shown as stars and a triangle, respectively, and the compositions of Sr-Ba-rich calcite from Murun and Bearpaw Mts. (Montana) as
diamonds (Konev et al. 1996; this work). Note that some
Bprotocarbonate^ compositions plot near the high-T stability field of disordered ternary carbonates identified in experiments. Four-step exsolution of one representative Murun Bprotocarbonate^ (large star) is shown
schematically as dashed arrows I-IV, and its products as circles (Konev
et al. 1996)
the liquidus minimum to temperatures below the calcitedolomite solvus crest and from 62 wt.% CaCO3 at 30 kbar to
91 wt.% at CaCO3 at 5 kbar, and increasing importance of
dissociation reactions affecting the stability of Mg-rich carbonates at low pressures (Irving and Wyllie 1975). At P≤6 kbar,
dolomite is not stable above the solvus crest and in dry systems, dissociates to Mg-rich calcite, periclase and vapor:
CaMgðCO3 Þ2 ⟺ ð1 þ xÞ
ð1Þ
Ca0:99−0:81xþ0:32x2 Mg0:01þ0:81x−0:32x2 CO3
þ ð1−xÞMgO þ ð1−xÞCO2
Textural variations in carbonatites
be a more common phenomenon in carbonatites than exsolution of calcite from Ca-rich dolomite. In fact, we have not
observed any convincing examples of the latter texture type
in some 25 years of studying carbonatites.
Experimental and thermodynamic data (Goldsmith et al.
1962; Anovitz and Essene 1987; Reeder and Dollase 1989)
indicate that the addition of Fe should be expected to expand
the stability field of hypersolvus Mg-Fe-rich calcite and narrow the dolomite field until it disappears completely at
~70 mol.% CaFe(CO3)2 (equivalent to 29 wt.% FeCO3).
The maximum CaFe(CO3)2 content in the composition of natural dolomite-ankerite from carbonatites approaches this solubility limit (e.g., # 10 in Table 1). The addition of Sr has an
effect opposite to that of Fe: the calcite field shrinks, eventually giving way to aragonite-type Ca-Sr carbonates with low
levels of Mg (e.g., ≤ 3 mol.% MgCO3 at 650 °C and 5 kbar),
whereas the dolomite structure persists across the entire
CaMg(CO3)2–SrMg(CO3)2 series (Brice and Chang 1973).
At P=12–35 kbar, the system CaCO3–SrCO3–BaCO3, relevant to the so-called barium-strontium carbonatites (Konev
et al. 1996), comprises aragonite-type compounds forming the
witherite-strontianite and aragonite-strontianite series [(Ba,
Sr)CO 3 and (Ca,Sr)CO 3 , respectively], and a field of
barytocalcite-structured carbonates (Chang 1971). With decreasing P, these fields shrink, and the calcite structure is
adapted by low-Ba (≤ 3 mol.% BaCO3) Sr-bearing compositions near the CaCO3 corner, which are separated from the
aragonite-type (Ca,Sr)CO 3 series by a miscibility gap
(Fig. 1b). The solubility of Sr in calcite increases with decreasing P and increasing T, but does not exceed ~30 mol.% SrCO3
in experimental systems (Carlson 1980). Additionally, a region of disordered rhombohedral ternary carbonates appears
adjacent to the barytocalcite field at P≤13 kbar (Chang 1971).
These disordered phases approach in composition some of the
hypothetical Ca-Ba-Sr Bprotocarbonates^ described by Konev
et al. (1996) from Murun (Siberia) and Jogipatti (India), where
the precursor phase underwent multistep unmixing to calcite,
barytocalcite and strontianite (Fig. 1b).
Because the above-described phase relations depend strongly on pressure (e.g., compare data on the solubility of Sr in
calcite at P=0–15 kbar: Chang 1971; Brice and Chang 1973;
Carlson 1980), it is difficult to provide any estimates for the
extent of substitution of Mg, Mn, Fe, Sr and Ba in carbonatitic
calcite and dolomite. The maximum concentrations of these
elements reported for calcite in the literature and measured in
the present work (corresponding to ~11, 8, 7, 13 and 1 mol.%
respective end-members; Table 1) appear to be well within the
experimentally established solubility limits. Moreover, the majority of compositions contain much lower levels of these elements (up to 0.5 wt.% MgO, 1 wt.% MnO or FeO, 2.5 wt.%
SrO and 0.2 wt.% BaO). Most dolomite analyses show≤
0.6 wt.% SrO at Ca and Fe levels within the experimentally
determined range, as discussed above. Despite the existence of
337
synthetic Ca-Sr dolomites (Brice and Chang 1973), the Sr content of carbonatitic dolomite is significantly lower than in
cogenetic calcite (Dawson and Hinton 2003; Zaitsev et al.
2014; Chakhmouradian et al. 2015b). Because of the obvious
charge constraints, Na and REE can be incorporated in calcite
and dolomite to much lower levels than divalent cations; the
mechanisms of Na and REE uptake by these minerals are uncertain (for discussion, see Chakhmouradian et al. 2015b).
Interpretation of carbonatite textures: challenges
and limitations
Petrogenetic studies rely greatly on our ability to identify textural criteria indicative of a specific process and to differentiate among visually similar textures produced by different
mechanisms. This ability is diminished if the texture of interest has been affected or even obliterated by subsequent evolution of the rock. In this respect, carbonatites are commonly
treated similarly to other igneous rocks, which is a mistake.
For example, some strongly metamorphosed carbonatites
showing preferred orientation and modal layering have been
misinterpreted as products of magma flow and differentiation
(see discussion in Chakhmouradian et al. 2015a). In comparison to the majority of igneous rocks composed of silicate
minerals, carbonatites are readily susceptible to textural and
compositional re-equilibration (e.g., recrysallization and isotopic resetting, respectively), grain abrasion, fragmentation
and comminution, ductile deformation, dissolution and other
forms of chemical interaction with fluids even at relatively
low T and P. For example, fine-grained felsic rocks exhibit
textural evidence of semi-brittle flow at T generally in excess
of 600 °C and P≥10 kbar, whereas the brittle-ductile transition
in calcite marble of similar texture will occur at P<1 kbar even
at ambient T (Fredrich et al. 1989; Hirth and Tullis 1994;
Snoke et al. 2014). The rate of Sr diffusion in calcite is twothree orders of magnitude faster than in feldspars (Cherniak
1997), not even to mention that Ca-bearing rhombohedral
carbonates are far more soluble than most rock-forming silicates in natural fluids (e.g., Brantley 2008). As a result, only a
small percentage of carbonatites retain their original igneous
texture, which is particularly true of old carbonatites, and
those emplaced at a significant depth (> 3 km) or in a tectonically active environment. It is difficult to replace the word
Bold^ here with more precise temporal constraints because
even shallow crustal carbonate sediments younger than
10 Ma have been documented to show evidence of recrystallization (Andreasen and Delaney 2000). The propensity to
recrystallization will obviously depend on the size of carbonate grains, the rate of cooling, the availability and chemistry of
a porous fluid, and other parameters that will vary significantly from one occurrence to the next. Young rocks emplaced
extrusively or near the surface and unaffected by
338
postmagmatic processes exhibit a variety of readily recognizable textures. Their characterization is beyond the scope of the
present work; interested readers are addressed to the reviews
and case studies by Keller (1989), Mitchell (1997), Zaitsev
et al. (2008), Eby et al. (2009), Stoppa and Schiazza (2013).
Recognition and petrogenetic analysis of old and/or deformed
intrusive carbonatites generally requires detailed knowledge
of their trace-element and isotopic characteristics (e.g., Fig. 2).
Igneous carbonate textures in the context
of petrogenesis
With relatively few exceptions, carbonatite textures are a complex product of different processes involving the parental carbonate magma, its conduit and wall rocks (or fragmented and
altered material derived from these rocks), fluids and gases
derived from the magma or other sources, and changes in
stress regime that may occur both during and after the emplacement. Many carbonatites exhibit a juxtaposition of petrographic characteristics generated or imposed by different
geological forces at different stages in the evolution of the
rock. This complexity poses obvious problems for anyone
attempting to present an overview of carbonatitic textures
and to categorize them in familiar terms. In the present contribution, we focus on those aspects of carbonatite petrography that can be identified as igneous, metamorphic, etc. with a
fair degree of certainty, although some of the examples described below may have more than one plausible explanation.
Fig. 2 An example of
discrimination diagram that can
be used to distinguish texturally
similar carbonatites and
metasedimentary rocks from
collision zones (data from
Demény et al. 2004;
Chakhmouradian et al. 2008;
Moore et al. 2015; Xu et al. 2015;
Chakhmouradian et al. 2015b).
Note the similarity of deformation
textures in calcite carbonatite
from Eden Lake (Canada) and
marble from Miaoya (China),
shown in the insets
A.R. Chakhmouradian et al.
Igneous textures: background information
Carbonate melts can be produced via a variety of mechanisms
and be either primary or derivative (for details, see Lee and
Wyllie 1998a, b). A primary mantle-derived magma will be
dolomitic (Ca/Mg≈1–2.4 by weight), rich in alkalis and contain minor silica (Wallace and Green 1988; Dalton and Presnal
1998; Lee and Wyllie 1998a; Wyllie and Lee 1998; Dasgupta
et al. 2006; Litasov and Ohtani 2009). Although alkali carbonate liquids with (Na2O+K2O)≥MgO have also been proposed
as primary magmas in the transition zone and lower mantle
based on diamond-inclusion studies and low-SiO 2 ,
high-(Na2O+K2O) experiments (e.g., Kaminsky et al. 2009;
Litasov et al. 2013), the relevance of these data to the actual
carbonatite source regions remains to be demonstrated. Because the presently known examples of indisputably mantlederived carbonatites and experimental data are scarce, it is
difficult to ascertain (near-)liquidus phase equilibria relevant
to natural systems. Primary extrusive and shallow intrusive
varieties commonly contain macrocrysts or phenocrysts of
Fe-Ti-rich spinel-group minerals, ferromagnesian silicates
and apatite (Tappe et al. 2006; Chakhmouradian et al. 2009;
Eby et al. 2009), implying that carbonate minerals are not
necessarily the earliest phases on the liquidus. Although the
synthetic primary melt of Wallace and Green (1988) quenched
to dolomite + an unspecified Na carbonate, the addition of
fluxes (e.g., H2O, F, or Na2CO3) will lower the liquidus of
the dolomitic melt below the crest of the calcite-dolomite
solvus, causing crystallization of calcite as a primary carbonate (Harmer and Gittings 1997; Wyllie and Lee 1998).
Textural variations in carbonatites
BHybrid^ carbonate-silicate magmas may evolve to produce
Ca-rich liquids (up to 80 wt.% CaCO3) by immiscibility or
crystal fractionation, but in either case, will precipitate silicate
minerals before reaching the carbonate liquidus (Lee and
Wyllie 1998a, p. 500). The earliest carbonate phase to crystallize from these derivative melts will be calcite, dolomite,
nyerereite or gregoryite, depending largely on the Ca/Mg,
CO2/H2O and Na2O/CaO proportions in the melt (Otto and
Wyllie 1993; Mitchell and Kjarsgaard 2011). Precipitation of
calcite or alkali carbonates will produce enrichment in Mg (±
Fe) in the residual liquid, and facilitate crystallization of dolomite (Fig. 3a) and, in Na-rich systems, eitelite
Fig. 3 Igneous textures in
carbonatites (a, b images in backscattered electrons, BSE; c-g
photographs in cross-polarized
transmitted light, XPL; symbols
as in previous Figs.). a Early
precipitation of calcite,
pyrochlore (Pcl) and apatite (Ap)
in carbonatite evolving toward
Mg enrichment and precipitation
of dolomite; Schryburt Lake,
Canada (scale bar, SB 0.5 mm). b
Primary burbankite (Brb)—
eitelite (Eit)—calcite inclusion in
magnetite (Mgt) from calcite
carbonatite; Sallanlatvi, Russia
(SB 50 μm). c Calcite crystals
showing transition from a
spinifex-like texture along the
contact of a Miocene calcite
carbonatite dike (at left) to an
aggregate of randomly oriented
laths closer to its axial zone; the
groundmass consists of calcite,
apatite and magnetite;
Kaiserstuhl, Germany (SB
1.5 mm). d Rounded autoliths
(xenoliths?) of disaggregated
cumulate calcite in a Miocene
carbonatite dike; Kaiserstuhl (SB
1 mm). e Partially disaggregated
autoliths of cumulate calcite,
magnetite and apatite in
Pleistocene extrusive carbonatite
(crb); Kerimasi, Tanzania (SB
1 mm). f Primary (or weakly
recrystallized) polygonal texture
of Eocene carbonatite dike
composed of calcite and
phlogopite (Phl), note bending of
the phlogopite crystals indicating
their transport in a crystal mush;
Bear Lodge, Wyoming (SB
1 mm). g detail of (f); compare
these textures with Fig. 5–7d of
Buob (2003) (SB 0.5 mm)
339
[Na2Mg(CO3)2] or neighborite (NaMgF3) (Mitchell and
Kjarsgaard 2011). Carbonatitic magmas are commonly
enriched in Sr and Ba, and can incorporate as much as
20 wt.% P 2 O 5 at P > 20 kbar (Baker and Wyllie 1992;
Ryabchikov and Hamilton 1993). Although low-P experimental data for phosphate-, Sr- and Ba-rich carbonate
melts are not available, inclusion studies indicate that
primary burbankite-group phases, norsethite
[BaMg(CO3)2] and bradleyite [Na3Mg(PO4)(CO3)] may
also accompany calcite or dolomite in the natural environment (Fig. 3b; Zaitsev and Chakhmouradian 2002;
Zaitsev et al. 2004).
340
Igneous textures: key examples
Near-surface young carbonatite dikes grade from a feathery,
spinifex-like texture in their chilled margin to more-or-less
randomly oriented tabular crystals and then equigranular
fine-grained aggregates of anhedral grains further inward
(Fig. 3c). It is less clear what a typical plutonic carbonatite
unmodified by postmagmatic processes looks like because of
the propensity of calcite and dolomite to textural re-equilibration. From autoliths in young extrusive rocks, we can surmise
that, depending on the depth and rate of crystallization, freshly
crystallized material ranges from a compact aggregate of tabular crystals (Fig. 3d, cf. right part of Fig. 3c) to irregularly
Fig. 4 Igneous textures and
mineral interrelations in
carbonatites (a, c, e–h BSE
images; b XPL image; d photo in
plane-polarized transmitted light,
PPL). a Graphic intergrowths
between apatite and calcite in
dolomite carbonatite; Albany
Forks, Canada (SB 50 μm). b
Graphic intergrowth between
olivine (Ol) and calcite in calcite
carbonatite; Borden, Canada (SB
1 mm). c Cumulate aggregate of
apatite, diopside (Di), pyrochlore
and calcite; Firesand River,
Canada (SB 0.5 mm). d Cumulate
aggregate of apatite, magnetite,
phlogopite and dolomite; Aley,
Canada (SB 1 mm). e Primary
inclusions of calcite and
nyerereite (Nye) in cumulus
apatite; Guli, Russia (0.5 mm).
f Primary inclusions of eitelite,
dolomite (indistinguishable at this
contrast) and burbankite in
cumulus apatite and magnetite;
Albany Forks (SB 0.5 mm).
g Dolomite reaction mantles
around glimmerite xenoliths
(glm) entrained in calcite
carbonatite, Chipman Lake,
Canada (Ab=albite; SB 0.5 mm);
dolomite is zoned toward
Fe-Mn-rich compositions
(##13, 14 in Table 1). h Dolomite
reaction mantle around a
glimmerite xenolith (glm)
entrained in calcite carbonatite,
Aley (SB 0.5 mm); dolomite is
zoned toward Fe-rich
compositions at constant Mn
(##15, 16 in Table 1)
A.R. Chakhmouradian et al.
shaped grains of variable size either interlocked into a granular
fabric, or developed interstitially with respect to ferromagnesian silicates, apatite and magnetite (Fig. 3e-g). The latter
minerals commonly have a long crystallization span and
thus may also occur as inclusions in coarser-grained carbonate grains (Fig. 3e), groundmass constituents, or form
several grain populations showing variable textural relations with the rock-forming carbonates. Graphic intergrowths are rare; typically, the non-carbonate component
is apatite, olivine, or magnetite (Fig. 4a, b). These textures are interpreted as eutectic based on a careful analysis of the orientation and compositional variation of
their constituent minerals.
Textural variations in carbonatites
Plutonic carbonatites commonly grade into meso- to
melanocratic cumulate rocks spanning a wide range of modal
compositions dominated by apatite, silicates (major olivine,
phlogopite, diopside, nepheline, calcic amphibole or andradite
with accessory zircon and titanite), or oxides (major magnetite, pyrochlore, perovskite or ferrocolumbite with accessory
ilmenite, rutile, baddeleyite, zirconolite and calzirtite); pyrrhotite and pyrite are typical sulfide constituents
(Chakhmouradian and Zaitsev 2004; Chakhmouradian et al.
2015a; Mitchell 2015). In these rocks, primary carbonate minerals occur as interstitial grains (Fig. 4c, d) and ovoid or lobate
cogenetic inclusions trapped in apatite and oxide phases
(Fig. 4e, f). Assimilation of wall-rock material by carbonatitic
magmas is also common; it is manifested in the extensive
resorption of xenoliths and precipitation of silicate minerals
(phlogopite, diopside, feldspars, amphiboles, titanite, andradite, wollastonite, epidote, allanite, scapolite) in the
endocontact of carbonatite intrusions, including the development of comb-like encrustations in the selvage and reactioninduced mantles on silicate xenocrysts (e.g., Chakhmouradian
et al. 2008). From the standpoint of carbonate textures, assimilation of silica-poor ultramafic rocks is particularly interesting
because it can locally raise the activity of Mg in the melt
sufficiently to precipitate dolomite (Fig. 4g, h) and even members of the magnesite-siderite series. The potential significance of this process for the petrogenesis of
magnesiocarbonatites remains to be understood.
Slowly cooled Mg-rich magmas (such as those derived
directly from a carbonated mantle source) could be expected
to produce Mg-rich calcite before precipitating dolomite (see
above). Indeed, early calcite with up to 4.3 wt.% MgO
(~11 mol.% MgCO 3 ) has been documented in some
carbonatites (e.g., Goldray, Canada; ##2, 3 in Table 1); these
compositions approach the Mg solubility limit in calcite in the
realistic temperature and pressure range (Goldsmith and
Heard 1961). Given that the solubility of Mg decreases with
cooling, progressive crystallization of calcite is expected to
produce a primary zoning pattern involving a decrease in
Mg content rim-ward. For similar reasons, the content of other
cations, whose size precludes their unlimited substitution for
Ca2+, should be expected to decrease in primary calcite with
progressive crystallization. Indeed, this pattern is observed in
some shallow intrusive carbonatites unaffected by
postmagmatic re-equilibration (Fig. 5a-c; Thompson et al.
2002, p. 380). This pattern may sometimes involve manifold
changes in the content of substituent elements (e.g., Fig. 5a,
b). In other cases, where calcite zoning is too subtle to be
detectable in BSE images, cathodoluminescence (CL) imaging may be a more effective method of visualizing
intragranular compositional variations, particularly those involving Mn (Fig. 5d).
Zoning in dolomite is much easier to observe because of
the large differences in atomic mass between the two principal
341
substituent elements, Mg and Fe. In most cases, primary dolomite evolves toward Fe-rich compositions at constant or
increasing levels of Mn (Figs. 4g, h and 5e, f); this trend
may be expressed both on the scale of individual grains and
from early to late-crystallized carbonatites in the same series
(Platt and Woolley 1990; Zaitsev et al. 2004). Contemporaneous crystallization of magnetite, apatite or Fe-Mg silicates will
obviously affect the partitioning of substituent elements between carbonate minerals and their host melt, and potentially
result in inverse or more complex zoning patterns (for possible
examples, see Thompson et al. 2002; D’Orazio et al. 2007).
The textures described below are less straightforward; in
each case, we provide an alternative explanation of their origin
and arguments in support of our own interpretation. Many
carbonatites are extremely inequigranular rocks characterized
by a bimodal or more complex grain-size distribution. The
terms Bporphyritic^ or Bseriate^ generally imply that the observed variations in size are igneous in origin. Where these
variations are generated by other processes, terms such as
Bporphyroclastic^, Bporphyroblastic^, Bblastomylonitic^, etc.
should be used. It is oftentimes difficult to make that distinction with respect to carbonatites, which are susceptible to deformation and can develop large-scale grain-size variations in
response to deviatoric stress under relatively low T and P. The
texture shown in Fig. 6a is interpreted here as porphyritic
because there is no evidence of intense strain (see Textural
record of deformation in carbonatites) that could produce
comminution on this scale. Most phenocrysts in this sample
exhibit only a subtle undulatory extinction and are devoid of
twinning, whereas smaller grains in the groundmass lack any
discernible preferred orientation; both contain primary inclusions that are easily removed during deformation (e.g.,
burbankite). Has this rock experienced some textural re-equilibration? Yes, it most likely has, judging from the alignment of
the elongate phenocrysts and crosswise twinning in those oriented perpendicular to the direction of flow. However, we are
convinced that the observed grain-size variations are best
accounted for by earlier crystallization of the large crystals,
rather than by any postmagmatic process. This conclusion is
supported by differences in trace-element budget between the
fine- and coarse-grained carbonates at this locality (cf. Olinger
2012).
Fragmentation in intrusive carbonatites can result from
magma injection into fractured wall rocks at shallow crustal
levels, from hydraulic fracturing associated with fluid release
from the parental carbonatitic magma, or from
postemplacement deformation. In our experience, true intrusive breccias show ample evidence of reaction between the
carbonatite and clasts or xenocrysts entrained in it (e.g.,
sericitization of potassium feldspar, saussuritization of plagioclase, or uralitization of clinopyroxene). Reaction rims around
xenocrysts and comb-like or radiating fringes around xenoliths are common. Alteration can be pervasive or confined to
342
A.R. Chakhmouradian et al.
Fig. 5 Characteristic zoning patterns in calcite and dolomite from
carbonatites (a–c, e, f BSE images; d CL image). a Calcite phenocrysts
from a diatreme showing rim-ward depletion in Mg, Mn, Fe, Sr and Ba
(##17, 18 in Table 1); Kontozero, Russia (SB 50 μm). b Zoned crystals of
primary calcite (Cal1) showing rim-ward depletion in Mg, Mn, Fe, Sr and
Ba (##19, 20 in Table 1); both core and rim are replaced by late-stage Srpoor calcite (Cal2); Carb Lake, Canada (SB 0.5 mm). c Spinifex-like
calcite laths from a dike (see Fig. 3c); their zoning involves a rim-ward
decrease in Fe, Sr and Ba, but no change in Mn (##21, 22 in Table 1);
Kaiserstuhl (SB 50 μm). d Variations in the intensity of
cathodoluminescence due to a decrease in Mn from the core of primary
calcite crystals (Cal1) toward their rim (##23, 24 in Table 1); similar
changes are observed in secondary calcite (Cal2) replacing both core
and rim; Afrikanda, Russia (SB 0.5 mm). e Dolomite crystals zoned
toward higher Fe and Mn levels in the rim and interstitial material
(##25-27 in Table 1); Chipman Lake (SB 0.5 mm). f Dolomite crystals
zoned toward higher Fe at constant Mn levels (##28, 29 in Table 1); Aley
(SB 0.5 mm)
the peripheral parts of a xenoliths (Fig. 6b). Heterogeneous
nucleation of carbonate minerals around xenoliths can
also produce radiating aggregates projecting into the
carbonatite; in tabular intrusions, flow patterns developed tangentially with respect to clasts are common
(Figs. 4h and 6b).
Globular structures such as those shown in Fig. 6c and d
have been often attributed to immiscibility between silicate
and carbonate melts, which has been experimentally proven
a viable petrogenetic mechanism for Mg-poor
natrocarbonatites (~15 wt.% Na 2 CO 3 ) conjugate with
peralkaline nephelinites (Kjarsgaard and Peterson 1991; Lee
and Wyllie 1998a), but cannot explain predominantly calcitic
or dolomitic ocelli described in numerous alkaline volcanic
and hypabyssal rocks ranging in composition from nephelinite
to basalt, trachyte and lamprophyre (s.l.). Apart from their low
alkali content, the absence of early-precipitating silicate
phases in such ocelli should be treated as potentially indicating an origin unrelated to immiscibility (see above). There is a
number of alternative mechanisms that can explain the presence of globular (or less regularly shaped) carbonate segregations in silicate rocks, including: (1) early crystallization of
round calcite crystals (Lee et al. 1994); (2) disaggregation of
cumulate carbonate rocks genetically related to the silicate
magma, followed by resorption or abrasion of the carbonate
autoliths during their transport (cf. Fig. 3d); (3) fragmentation
of carbonate material (e.g., sedimentary wall rocks) by the
silicate magma, followed by recrystallization, resorption, or
abrasion of the carbonate xenoliths during their transport
(Azbej et al. 2006); (4) hydrothermal precipitation of carbonates in vesicles within a (partially) solidified silicate host
(Azbej et al. 2006; Fig. 6e, f); (5) mobilization of residual
Textural variations in carbonatites
343
Fig. 6 Miscellaneous carbonate
textures in igneous rocks (a–c, e
XPL images, SB 1 mm; d, f PPL
images, SB 0.5 mm). a
Porphyritic carbonatite; both
phenocrysts and matrix calcite
grains contain burbankite
inclusions and are interpreted as
primary; Bear Lodge. b
Sericitization (ser) of syenite
xenoliths (sye) at the contact with
calcite carbonatite; note calcite
growth patterns at the contact;
Tamazert, Morocco. c Calcite
globule in bergalite (feldspathoid
melilitite); Kaiserstuhl. d Calcite
globules with a selvage of Ca-K
zeolite (Zeo) separated by a
Bmeniscus^; vitrophyric combeite
nephelinite (Cmb=combeite;
Cpx=aegirine-augite; Ttn=
titanite), Oldoinyo Lengai,
Tanzania (cf. Figs. 3 and 4 of
Kjarsgaard and Peterson 1991). e
An amygdale (am) filled with
inward-projecting calcite crystals
which differ significantly in both
trace-element and isotopic
composition from primary calcite
in the matrix (mtx) and rounded
autoliths (aut), Kaiserstuhl (see
also Fig. 3d). f Irregularly shaped
zeolite and zeolite-calcite
aggregates in combeite
nephelinite, Oldoinyo Lengai
(Ne=nepheline). The textures
shown in c–f are products of
hydrothermal precipitation in
vesicles, not carbonate-silicate
immiscibility
liquids into gas vesicles during cooling (Cooper 1979); and
(6) pseudomorphization of equant silicate phenocrysts by latestage carbonates (see below). Clearly, it is not generally
possible to discriminate among these alternative mechanisms
based exclusively on petrographic and field evidence. A careful analysis of trace-element distributions and inclusions in
344
A.R. Chakhmouradian et al.
both carbonate segregations and their host rock is usually
applied for this purpose.
Some intrusive carbonatites exhibit poikilitic textures comprising relatively large (>0.5 mm across) oikocrysts of primary carbonate hosting numerous ovoid or lobate grains of calcite (in dolomite), dolomite (in calcite) and burbankite; typical
non-carbonate chadacrysts, where present, include apatite,
barite, monazite and fluorite (Figs. 5e and 7a, b). Very limited
data are available on these inclusion assemblages in the literature, where they are interpreted as either products of exsolution, or precipitates from a Na-rich immiscible fluid associated
with the parental carbonatitic magma (Platt and Woolley
1990; Faiziev et al. 1998; Tichomirowa et al. 2013). We consider both these interpretations problematic because (1) the
distribution of burbankite and other chadacrysts in the host
crystal is not controlled crystallographically, nor confined to
cleavage planes or fractures (see the next section); (2) strontianite (SrCO3), which is the most common product of calcite
breakdown, does not occur in these parageneses; (3) analogous inclusions are found in minerals, which cannot possibly
exsolve carbonates (e.g., Fig. 4e, f); (4) bulk analyses of dolomite grains containing burbankite inclusions give up to several thousand ppm Na and REE, which cannot be realistically
accommodated in the structure of this mineral
(Chakhmouradian et al. 2015b); and (5) subsolidus relations
in the system CaCO3-MgCO3 (see Phase and compositional
relations of significance to rock-forming carbonates in
carbonatites) suggest that exsolution of dolomite from calcite
Fig. 7 Poikilitic and contamination textures in carbonatites (a–c falsecolor BSE images, SB 50 μm; d-f BSE images, SB 0.5 mm). a Dolomite
oikocryst containing inclusions of calcite, burbankite, apatite and fluorite
(Fl); Chipman Lake. b Calcite oikocryst containing inclusions of dolomite, burbankite and barite (Bar); Aley. c Poikilitic calcite containing
inclusions of burbankite, celestine (Cls), monazite (Mnz) and bastnäsite
(Bst); Mountain Pass, California. d Autolith (xenolith?) of cumulate
carbonatite enclosed in aphanitic hypabyssal calciocarbonatite (at left);
the autolith comprises laths of Sr-rich, Mg-poor calcite (Cal1) and interstitial Mg-rich, Sr-poor calcite (Cal2); Kaiserstuhl. e Dolomite carbonatite
intruded by calcite carbonatite, Carb Lake; primary dolomite (Dol1) contains 7.5–11.2 wt.% FeO, 0.7–1.9 wt.% MnO, 0.1–0.3 wt.% SrO, and is
associated with potassic-fluoro-magnesio-arfvedsonite (Amp1); at the
contact with calcite, Dol1 is mantled by Fe-Mn-Sr-poor dolomite Dol2
(2.3–6.6, 0.3–0.8, 0–0.1 wt.% respective oxides), and Amp1 by fluororichterite (Amp2). f Magnetite-phlogopite-calcite carbonatite intersected
by a veinlet of richterite (Rct)—dolomite carbonatite, Aley; both calcite
and dolomite contain high levels of Sr, Ba, REE and are interpreted as
magmatic (note also the sharp contacts between the two rocks and a flow
pattern in the veinlet)
Textural variations in carbonatites
345
should be much more common than vice versa, whereas calcite with ovoid inclusions of dolomite is not as abundant as
poikilitic dolomite. Consequently, we consider these inclusions to be primary, i.e. syngenetic with the host carbonate
mineral. Although superficially similar poikilitic intergrowths
are produced by exsolution of initially homogeneous nonstoichiometric carbonates during cooling, they can be readily
identified as such on the basis of textural criteria (see the next
section). The round shape of carbonate inclusions, as the principal argument put forth by the proponents of their secondary
origin (Platt and Woolley 1990), is unconvincing because these minerals may (a) crystallize with this habit in response to
surface tension effects (Lee et al. 1994), or (b) attain this habit
due to their interaction with the host melt. Apatite, for example, is an early liquidus mineral in carbonatites, and yet it
typically occurs as oblong or equant grains with no recognizable crystal forms (e.g., Figs. 3a, e and 4c, e). Calcite
oikocrysts in postorogenic carbonatites commonly enclose
minute tabular grains of REE fluorocarbonates and celestine
(SrSO4), also interpreted as primary inclusions (Fig. 7c;
Chakhmouradian et al. 2008).
There is no doubt that some carbonatites comprise products
of crystallization of two (and, perhaps, more) different
magmas, which may or may not be genetically related to each
other, or even products of mechanical mixing (Le Bas et al.
2004; Tichomirowa et al. 2013). Several mechanisms can be
envisioned to produce such petrographically complex rocks
(Fig. 8): (1) magma mixing and subsequent crystallization of
the hybrid magma; (2) fragmentation of earlier-crystallized
carbonatites and transport of these fragments by a younger
magma; (3) infiltration of earlier-crystallized carbonatites by
a small volume of younger magma and its crystallization in
situ; (4) fragmentation and transport of solidified carbonatitic
material from different sources and its emplacement in the
form of breccia; and (5) tectonic mobilization of carbonatitic
material from different sources and its emplacement into the
same fissure by ductile flow. The latter two processes are
related to deformation and, hence, will be discussed below.
From the remaining three mechanisms, magma mixing
(Fig. 8a) is the most difficult one to recognize because it leaves
little textural evidence owing to the unique rheological and
diffusion characteristics of carbonate melts (Genge et al.
1995; Jones et al. 2013). The few available published studies
of mixing in carbonatites invoke microtextural complexity
and extreme trace-element or isotopic variations in earlycrystallizing minerals (such as pyrochlore, apatite and zircon)
to support this interpretation (Zurevinski and Mitchell 2004;
Chen and Simonetti 2012; Tichomirowa et al. 2013). Models
(2) and (3) both imply interaction of some earlier-solidified
material (host rock, xenoliths, xenocrysts or autoliths) with an
intruding melt (Fig. 8b, c) and, hence, are expected to produce
a modally and texturally heterogeneous rock featuring disequilibrium textures (e.g., resorbed crystals, reaction rims)
and appreciable variations in the composition of its constituent
minerals. Distinction between these two mechanisms generally cannot be done on the basis of thin-section petrography
alone, and requires some outcrop-scale observations that will
unequivocally establish the structural and volumetric relations
Fig. 8 Schematic diagram illustrating the possible mechanisms of
formation of Bhybrid^ carbonatitic rocks (for details, see text): a Mixing
of carbonatitic magmas; b magmatic erosion of carbonatite 1 by younger
carbonatite 2 (see Fig. 7d); c intrusion of carbonatite 1 by carbonatite 2
and in situ crystallization of the latter (see Fig. 7e); d fracturing and
brecciation of carbonatites 1 and 2 in the zone of brittle deformation
(see Figs. 13e, f); e tectonic mobilization of carbonatites 1 and 2 in the
zone of ductile deformation generating intimately mingled rocks such as
that in Fig. 16f
346
between the interacting entities. The examples shown in
Figs. 3d and 7d–f were interpreted using a combination of
field, drillcore, microscopic and geochemical evidence.
Textural record of the postmagmatic evolution
of carbonatites
Exsolution, Bexsolution^ and other types of subsolidus
chemical re-equilibration
Owing to the limited miscibility between CaCO3 and other
carbonate phases at low T (Fig. 1 and references therein), early
Fig. 9 Exsolution and carbonatefluid re-equilibration in
carbonatites (a PPL image; b–d, f,
g BSE images; e CL image;h
false-color BSE image). a–c
Calcite-dolomite exsolution,
Goldray (SB 1, 0.5 and 0.2 mm).
d Breakdown of Ca-Ba-Sr
Bprotocarbonate^ (Pcrb) to calcite
and barytocalcite (Bc) (Kh=
kukharenkoite); Murun (SB
50 μm). e Diffusion-induced
zoning in primary calcite
involving a decrease in Mn, Sr
and REE along grain boundaries
and fractures; Afrikanda (SB
0.5 mm); the Sr and REE released
from calcite are precipitated
interstitially as non-luminescent
strontianite and ancylite. f
Diffusion-induced zoning in
primary calcite (Cal1) involving a
decrease in Sr and Ba levels along
grain boundaries and fractures
(see ##30, 31 in Table 1); the
released Sr and Ba are deposited
as strontianite, minor
barytocalcite and unidentified BaCa-Sr carbonates (white specks)
in Cal2; Murun (SB 0.2 mm). g
Diffusion-induced zoning in
primary calcite (Cal1) involving
progressive depletion in Sr, Ba
and REE in Cal2 developed along
cleavage planes and grain
boundaries (##32–34 in Table 1);
Bearpaw Mts. (SB 0.5 mm). h
Detail of g showing Sr, Ba and
REE minerals deposited in
secondary calcite, including
strontianite (Str), burbankite,
ancylite (Anc) and barite;
fractures are traced with dashed
blue lines (SB 50 μm)
A.R. Chakhmouradian et al.
magmatic carbonates enriched in substituent elements can be
expected to Bunmix^ upon cooling. In reality, true exsolution
textures, generated by ionic diffusion in the solid state with no
changes to the bulk composition of the grain, are relatively
uncommon. Some examples, including unmixed Ca-Ba-Sr
Bprotocarbonates^ and Mg-rich calcite, are shown in
Fig. 9a–d. Other well-characterized exsolution textures in intrusive carbonatites include:
(1) Calcite-dolomite intergrowths from Siilinjärvi in Finland, Kovdor in Russia and Phalaborwa (Puustinen
1974; Zaitsev and Polezhaeva 1994; Dawson and Hinton
2003);
Textural variations in carbonatites
(2) Calcite-benstonite-barytocalcite(?) aggregates at
Jogipatti, India, developed at the expense of Ca-Ba-Sr
Bprotocarbonate^ similar to that in the Murun bariumstrontium carbonatites, but enriched in Mg (Konev et al.
1996; our Fig. 1);
(3) Lamellae of carbocernaite [(Ca,Na)(Sr,REE,Ba)(CO3)2]
in Sr-rich calcite at Sarnu-Dandali in India (Wall et al.
1993).
In true exsolution textures, lamellae are distributed within
the host crystal independently of fractures, cleavage planes
and grain boundaries, but may vary in size and density if the
precursor crystal was zoned. Smaller second- and third-order
inclusions (termed secondary and tertiary microstructures in
materials science literature) also occur in some samples
(Fig. 9c, d). Their formation does not require any compositional heterogeneities in the precursor carbonate and most
likely depends on the thermal history of the rock (e.g., Mitchell et al. 2004), including the rate of cooling and any subsequent thermal events, such as reheating during metamorphism. At present, meaningful interpretation of secondary
and tertiary microstructures is precluded by their small size
not amenable to quantitative analysis. The conditions, under
which primary lamellae and their host equilibrated, can be
estimated using model systems (Chang 1971; Carlson 1980;
Anovitz and Essene 1987), provided the composition of both
phases is reasonably simple (for some examples, see Wall
et al. 1993; Zaitsev and Polezhaeva 1994). The calculations
made in the previous studies for calcite range from 420 to
700 °C (Puustinen 1974; Wall et al. 1993; Zaitsev and
Polezhaeva 1994), i.e. are generally consistent with the lowP estimates of carbonatite solidus temperatures (Boettcher
et al. 1980; Andersen and Austrheim 1991; Veksler et al.
1998).
In many studies, the term Bexsolution^ is applied loosely to
any fluid-driven compositional changes in primary carbonates
(predominantly, calcite) that result in the diffusion and release
of specific elements from the precursor mineral and their sequestration in secondary phases. The latter are typically deposited as inclusions in the chemically re-equilibrated host
crystal along fluid passageways (fractures, cleavage planes,
etc.), as well as along grain boundaries (Fig. 9e–g). This pattern of distribution and the common presence of hydrous or
non-carbonate minerals in this paragenesis (commonly,
ancylite or barite: Fig. 9h) clearly indicate fluid involvement.
The inclusions are typically irregular in shape, do not exceed
30 μm in size, and comprise predominantly Ca-bearing strontianite; barytocalcite, burbankite and carbocernaite are present
as inclusions in calcite from some localities (Fig. 9h). As can
be expected, secondary calcite is depleted in Sr, Ba and REE
relative to its magmatic precursor (Fig. 9e–g). These processes
should not be confused with true exsolution for two reasons.
First of all, interaction of primary carbonates with a fluid
347
releases Sr and other cations into a fluid and, hence, may
potentially result in the development of hydrothermal raremetal mineralization away from, and not necessarily in any
spatial connection to, the source. One example is ancylitegroup minerals deposited in cavities in hydrothermally
reworked carbonatite at Afrikanda, Russia (Zaitsev and
Chakhmouradian 2002). Secondly, separation of strontianite
(or other phases) from a primary carbonate does not necessarily imply that the limit of Sr solubility has been reached. Instead, this process may be driven by kinetics (i.e., faster Sr
diffusion rate: Fisler and Cygan 1999) or external factors, such
as the relative activities of Sr2+ and Ca2+ in the fluid.
One important group of processes, that have far-reaching
implications for the dispersal and concentration of rare metals
in carbonatites but remain inadequately understood, is metasomatic replacement of primary magmatic minerals by secondary carbonates. Dolomitization of calcite (Fig. 10a) is the
best known process of that kind and can be recognized by the
presence of relict calcite grains in a network of fine-grained
dolomite veinlets, replacement of associated silicate and oxide
minerals (e.g., olivine by serpentine; phlogopite by tetraferriphlogopite or chlorite; clinopyroxene by alkali amphiboles; Fig. 10b–d), co-existence of fresh and heavily altered
mineral grains on a fine spatial scale, and changes in the isotopic composition (C, O and Sr) of carbonates (Kapustin
1987; Schürmann et al. 1997; Downes et al. 2012;
Chakhmouradian et al. 2015a). Metasomatic replacement of
primary dolomite is probably just as common, but is much
more difficult to trace (Chakhmouradian et al. 2015b). At
advanced stages of alteration, dolomite replaces not only the
rock-forming carbonates, but also ferromagnesian silicates
(olivine, amphiboles, phlogopite) and magnetite, commonly
in intimate association with other secondary minerals (chlorite, serpentine, quartz, goethite, rutile and Mg-Fe carbonates;
Fig. 10e–g). Because dolomite has a smaller capacity for Sr,
Ba and REE than its precursor calcite (Chakhmouradian et al.
2015b), these elements are often deposited as strontianite, barite, ancylite, REE fluorocarbonates and monazite confined to
fractures and grain boundaries. The extent and potential significance of this secondary rare-metal mineralization in
carbonatites strongly affected by dolomitization (e.g., Aley
in Canada) are yet to be ascertained. From the exploration
standpoint, dolomitization is also important because it is accompanied by the replacement of primary Nb minerals (predominantly, pyrochlore) by secondary phases, such as
fersmite (CaNb2O6), columbite (FeNb2O6), cation-deficient
pyrochlores, or Nb silicates (Voloshin et al. 1990; Subbotin
and Subbotina 2000; Melgarejo et al. 2012; Torró et al. 2012;
Chakhmouradian et al. 2015a). Clearly, these changes affect
not only the Nb and Ta grade, but also the amenability of ore
to beneficiation and metal recovery.
Other carbonate replacement processes documented in intrusive carbonatites include conversion of dolomite to calcite
348
A.R. Chakhmouradian et al.
Fig. 10 Dolomitization in
carbonatites (a PPL image, SB
1 mm; b–g BSE images, SB
0.2 mm). a Fresh phlogopitemagnetite-calcite carbonatite
(left) and products of its
reworking by Mg-rich fluids
(right), consisting of dolomite and
dolomite-rich pseudomorphs
(psd) after magnetite and
phlogopite; Aley. b Incipient
dolomitization of calcite
carbonatite; note complete
replacement of phlogopite by
dolomite and chlorite (Prs=
parisite); Aley. c Incipient dolomitization of calcite carbonatite;
note replacement of calcite by
dolomite and development of
acicular magnesioriebeckite
(Mrb) along fractures; Prairie
Lake, Canada. d Advanced dolomitization of calcite carbonatite;
Aley. e Euhedral dolomitechlorite pseudomorph after
phlogopite in dolomite
carbonatite; Aley; relative to primary dolomite (0.4–1.0 wt.%
SrO), the secondary variety lacks
detectable Sr. f Quartz (Qtz)—
dolomite pseudomorphs after
amphibole in dolomite
carbonatite; Aley; primary Fepoor, Sr-rich dolomite [3 mol.%
CaFe(CO3)2 and 0.5 wt.% SrO] is
mantled by hydrothermal ankerite
(Ank) [57 mol.% CaFe(CO3)2
and 0.2 wt.% SrO], whereas dolomite in the pseudomorphs contains 20 mol.% CaFe(CO3)2 and
no detectable Sr. g Rutile (Rt)—
quartz—dolomite pseudomorph
after magnetite in strongly
dolomitized dolomite carbonatite;
Aley; in contrast to (f), secondary
dolomite here is enriched in Fe
relative to the primary variety
(van der Veen 1965; see also Fig. 11a) and to Mg-Fe carbonates (Zaitsev et al. 2004; Fig. 11b, c). These reactions probably
involve incongruent dissolution and are accompanied by CO2
release, as indicated by the cavernous nature of the
metasomatized rock. Little is known about the hydrothermal
conditions at which the above-described transformations occur. Zaitsev et al. (2004) inferred a T of 250 °C at high Mg2+
activities (aCa2+/aMg2+ <2.3) and X(CO2)=0.4–0.6 for the
development of magnesite at Sallanlatvi (Russia), whereas
dolomitization of calcite carbonatites at Aley probably occurred at higher T (≥400 °C) in a comparable range of
X(CO2) (Chakhmouradian et al. 2015a). According to the
limited published data (Onuonga et al. 1997; Schürmann
et al. 1997; Demény et al. 2004; Chakhmouradian et al.
2015a), hydrothermally reworked carbonates exhibit a
positive shift in δ18O values indicative of crustal or atmospheric input, whereas their C isotopic signature could
remain unaffected at low CO2 levels in the fluid, become
Bdiluted^ by a 12 C-enriched juvenile component, or
change to higher δ13C values if the fluid equilibrated
with sedimentary carbonates. In terms of its timing with
respect to carbonatite emplacement, hydrothermal
overprinting can occur at the cooling stage and involve
fluids of carbonatitic provenance (e.g., Schürmann et al.
1997; Moore et al. 2015), or take place much later in
response to deformation and metamorphism unrelated to
Textural variations in carbonatites
349
Fig. 11 Characteristic
replacement textures in
carbonatites (a–c, e, f: BSE
images; d: XPL image). a
Calcitization of primary dolomite
(note precipitation of monazite in
calcite veinlets); Aley (SB
0.5 mm). b Replacement or
primary dolomite by magnesite;
Sallanlatvi (SB 0.2 mm). c
Replacement of primary dolomite
by zoned magnesite-siderite (Sd)
aggregates; Albany Forks (SB
50 μm). d Hexagonal prismatic
pseudomorphs after primary
burbankite (Bpsd) in calcite
carbonatite; Bear Lodge (SB
0.2 mm). e Detail of a cavernous
calcite-ancylite-strontianite
pseudomorph after burbankite;
Bear Lodge (SB 0.2 mm). f Detail
of calcite-strontianite-barite
pseudomorph after burbankite or
another Sr-Ca-Ba carbonate;
Mountain Pass (SB 0.5 mm)
this magmatism (e.g., Chakhmouradian et al. 2015a; Xu
et al. 2015).
Late-stage hydrothermal carbonates
At shallow crustal levels, interaction of carbonatites with
aqueous fluids will facilitate the dissolution and removal of
some carbonate material, enhance fracturing and porosity, and
lead to the development of distinctive mineral assemblages
confined to fractures, cavities and intergranular spaces. Precipitation of late-stage carbonates in this environment can be
triggered by CO2 release due to the opening of the fluid circulation system and depressurization (e.g., Coto et al. 2012;
reaction 2), or by an increase in pH during metasomatic reactions such as (3) (see also Fig. 10b, e):
Ca2þ þ 2HCO3 − ⇔CaCO3 þ CO2 þ H2 O
2KMg3 AlSi3 O10 ðOHÞ2 þ CaCO3 þ 2Hþ þ H2 O þ CO2 ⟺ Mg5 Al2 Si3 O10 ðOHÞ8 þ CaMgðCO3 Þ2 þ 3SiO2 þ 2Kþ
phlogopite
calcite
chlorite
dolomite
quartz
Late-stage calcite and dolomite can crystallize over a
wide range of conditions and, unless these conditions can
be constrained from mineral equilibria or fluid-inclusion
data, it may not be always possible to distinguish between hydrothermally deposited and supergene carbonates. For example, both varieties can exhibit depletion
in Mn and Fe owing to preferential partitioning of these
ð2Þ
ð3Þ
elements into secondary goethite and other hydroxide
phases in an oxidizing environment (Chakhmouradian
et al. 2015b). Stable-isotope studies do not necessarily
yield unambiguous interpretations either, because hydrothermal fluids can derive their C and O from a variety of
sources, including groundwater (e.g., Simonetti et al.
1995; Onuonga et al. 1997).
350
Hydrothermal processes in carbonatites are not merely of
academic interest. They affect the distribution and forms of
concentration of REE, Sr, Ba and F, and are capable of yielding orebodies of economic interest (Wall and Mariano 1996;
Zaitsev et al. 1998; Bühn et al. 2003; Ruberti et al. 2008;
Doroshkevich et al. 2009; Moore et al. 2015). One notable
example is decomposition of early magmatic burbankitegroup minerals to produce cavernous polymineralic pseudomorphs such as those shown in Fig. 11d. Based on the available geochemical evidence, this process involves carbonatitederived fluids of variable T and chemistry, which is reflected
in the relatively unmodified (i.e., mantle-like) isotopic composition of the pseudomorphs and their mineralogical diversity (Zaitsev et al. 1998, 2002; Moore et al. 2015). Calcite is a
ubiquitous component of these parageneses (Fig. 11e, f); its
low Sr content with respect to igneous calcite (e.g., 0.2–0.3 vs.
1.0–1.3 wt.% SrO at Mountain Pass, USA) attests to relatively
low deposition temperatures (Carlson 1980).
Late-stage hydrothermal calcite typically forms drusy clusters of rhombohedral or scalenohedral crystals projecting into
fractures and vugs, botryoidal aggregates and overgrowths on
dolomite (Figs. 6e and 12a–e). Commonly, these aggregates
contain a resorbed nucleus of earlier-crystallized material
(Fig. 12c) and exhibit zoning uncharacteristic of igneous carbonates, such as sectoral and highly periodic oscillatory patterns (Fig. 12d, f). Hydrothermal calcite is poor in Sr
(≤0.4 wt.% SrO), Fe, Mn and Ba (typically, below detection
of EMPA), but may contain appreciable levels of Mg (e.g.,
##38–40 in Table 1). It is often associated with dolomite,
strontianite, barite, fluorite, ancylite, REE fluorocarbonates,
chlorite, titanite, zeolites and quartz. Hydrothermal dolomite
occurs as rhombohedral crystals (commonly incorporating a
nucleus of early dolomite and showing intricate growth zoning) associated with calcite, siderite, chlorite, muscovite,
quartz, barite, strontianite, REE fluorocarbonates, fersmite
and other late-stage niobates (Fig. 12g, h). In comparison with
igneous dolomite, the hydrothermal variety is depleted in Mn,
Sr and Ba (cf. ## 41–44 in Table 1), but may show extreme
enrichment in Fe. In fact, ankerite-dominant compositions occur exclusively in this type of environment (##10 and 44;
Figs. 10f, 12i).
Textural record of deformation in carbonatites
Deformation of carbonate rocks: background information
Carbonatites are amenable to textural re-equilibration at low T
and P, particularly in the presence of a fluid, at time scales
readily reproducible in experiment (Terent’ev and Kunts
2001; De Bresser et al. 2005; Schultz et al. 2013). Even young
(<10 Ma) carbonate sediments from shallow crust have been
reported to show evidence of recrystallization (Andreasen and
A.R. Chakhmouradian et al.
Delaney 2000). Note that equant grain shapes and triple boundary junctions are not necessarily indicative of postmagmatic
textural re-equilibration (see, e.g., synthetic calcite in Figs. 5–
7d of Buob 2003), but coarse-grained carbonatites exhibiting
twinning and a polygonal mosaic texture (e.g., Fig. 13a) were
most certainly statically recrystallized.
Intrusive (nota bene!) bodies of non-igneous
metacarbonate rocks (e.g., marbles) have been known for over
170 years (Emmons 1842). It has since been established experimentally that carbonate rocks become ductile under relatively low confining pressures. Under stress, the depth of
brittle-ductile transition in these rocks will depend on the grain
size, porosity, content and connectivity of non-carbonate material, availability of fluids, and temperature (De Bresser et al.
2005; Paterson and Wong 2005; Renner et al. 2007; and references therein), and can be as shallow as ~1 km for a pure
material (Fredrich et al. 1989). A confining P of 0.5 kbar is
typically cited for the reference Carrara marble (Schubnel
et al. 2006), which is closely similar to fine-grained
anchimonomineralic (~98 % pure, with a mean grain size of
0.15 mm) calcite carbonatites. In a tectonically active environment, such as plate collision zones, carbonate rocks of any
kind (including igneous) can thus be readily mobilized to undergo ductile deformation, stress-induced flow and emplacement into rigid fractured silicate wall rocks in the form of ex
situ intrusions or Bextrusions^ (Roberts and Zwaan 2007;
Chakhmouradian et al. 2015a). Naturally, these processes will
overprint and, in some cases, obliterate many of the recognizable igneous characteristics of carbonatites, producing a variety of textures and structures indicative of cataclasis, ductile
deformation and dynamic recrystallization (Fig. 13).
Initial response to stress in metacarbonate rocks includes
mechanical twinning, undulatory extinction (Fig. 13b), and
elongation of carbonate grains obliquely with respect to the
shear zone boundary (SZB). With increasing strain, twin lamellae grow thicker, develop bent or lenticular shapes
(Fig. 13c), whereas grain boundaries become serrated (Pieri
et al. 2001; Barnhoorn et al. 2004). This intragranular deformation is accompanied by increasing irregularity of grain
boundaries and development of misoriented subgrains
(Fig. 13d). Already at γ=3 (at 500 °C for the Carrara marble)
and even lower shear strain levels at higher temperatures (γ=
1 at 730 °C), small grains begin to form interstitially by grainboundary bulging and subgrain rotation. This, in combination
with the changing character of twin boundaries and optical
extinction, points to dislocation creep as the principal mechanism of deformation in the ductile regime. As creep competes
with cracking near the brittle-ductile transition at low T, dislocation pile-ups at grain boundaries and twin intersections
yield micro-fractures (Schubnel et al. 2006; Fig. 13d). At elevated temperatures, progressive dynamic recrystallization produces what is known in metamorphic petrology as the Bcoreand-mantle (micro)structure^, i.e. elongate deformed
Textural variations in carbonatites
Fig. 12 Characteristic
morphology, zoning and
parageneses of hydrothermally
deposited calcite and dolomite in
carbonatites (a image in
secondary electrons; b, d, e, g–i
BSE images; c XPL image; f CL
image). a Rhombohedral crystals
of hydrothermal calcite from
silicocarbonatite; Afrikanda (SB
0.5 mm). b Rhombohedral
crystals of calcite mantled by
hydrothermal dolomite and Ferich chlorite; other associated
minerals include strontianite,
parisite, pyrite (Py) and sphalerite
(Sph); Bear Lodge (SB 0.5 mm).
c Interstratified drusy and
oscillatory-zoned botryoidal
calcite deposited on fragments of
apatite-dolomite carbonatite; Iron
Hill, Colorado (SB 1 mm). d
Drusy Mg-rich calcite showing
oscillatory and sectoral zoning
(##38, 39 in Table 1); Mountain
Pass (SB 0.5 mm). e
Rhombohedral dolomite with
interstitial calcite and fersmite
(Frs); Aley (SB 0.5 mm). f Zoning
in rhombohedral calcite involving
an increase in Sr and Mn from the
core outward; note sectoral zones
in the brightly luminescent rim;
Afrikanda (SB 0.2 mm). g
Rhombohedral dolomite
associated with siderite, chlorite,
muscovite (Ms) and barite;
Bearpaw Mts. (SB 50 μm). h
Rhombohedral dolomite (Dol2)
containing cores of primary
dolomite (Dol1) enclosed in
hydrothermal quartz; Dol1 is
enriched in Fe, Mn and Sr relative
to Dol2 (##41, 42 in Table 1);
Aley (SB 0.5 mm). i Rhombohedral ankerite containing cores of
primary dolomite; the dolomite is
enriched in Sr and Mn relative to
the ankerite (##43, 44 in Table 1);
Aley (SB 0.2 mm)
351
352
A.R. Chakhmouradian et al.
Fig. 13 Low-strain and
brecciation textures in
carbonatites (a–d, f XPL images;
e PPL image; SB 1 mm). a
Mosaic polygonal texture in
calcite carbonatite; Lackner Lake,
Canada. b Bent and
discontinuous twinning lamellae
in subtly elongate calcite crystals;
Aley; c Bent and segmented thick
twinning lamellae in calcite
overprinted by thin straight
lamellae produced during
decompression (Pieri et al. 2001);
Huayangchuan, China. d
Incipient deformation-induced
textural re-equilibration in the
brittle-ductile transition regime;
note grain-boundary bulging
(yellow arrows), misoriented
subgrains (red arrows) and microfractures (blue arrows); Bear
Lodge. e, f Breccia comprising
fragments from at least two
different carbonatite units,
distinguished on the basis of
dolomite compositions, and
unaltered albite xenocrysts
(indiscernible at this scale) (see
also Fig. 8d)
porphyroclasts set in a matrix of small nearly equant grains
generally devoid of twinning. The sheared rock is conspicuously foliated and characterized by a bimodal grain-size distribution. With increasing γ at a constant T, the size and abundance of porphyroclasts decrease, whereas their aspect ratio
increases, exceeding five at γ=5 and T=730 °C for the Carrara
marble (Pieri et al. 2001). The primary foliation (S1), defined
by the alignment of porphyroclast boundaries and their
{0001} planes, gradually gives way to a secondary preferred
orientation (S2) at 50–60° to the SZB owing to the alignment
of matrix grains. Complete conversion of the Carrara marble
to ultramylonite occurs by γ=10 at 730 °C, but requires much
higher levels of strain at lower temperatures (Barnhoorn et al.
2004). Published petrologic evidence indicates that such
levels are achievable even under lower greenschist-facies
metamorphic conditions (e.g., Bestmann et al. 2000). Although cataclasis and solution-transfer are believed to control
textural transformation of carbonate rocks under subgreenschist conditions, there is some petrographic evidence
for dynamic recrystallization in calcite at temperatures as
low as 150 °C (Kennedy and White 2001).
The presence of rigid components (i.e. crystals not as amenable to twinning and dislocation glide as carbonates) or a
fluid phase in the protolith has a dramatic effect on its strength
and response to deformation. For example, the addition of
quartz grains to experimental charges increases the threshold
strain marking the onset of steady-state flow in the carbonate
matrix and expedites the consumption of calcite
porphyroclasts (Dresen et al. 1998; Rybacki et al. 2003). Pore
fluids weaken the rock, particularly at T≤600 °C, by enhancing dilatation and promoting grain-boundary sliding, which
has a randomizing textural effect (Busch and van der Pluijm
1995; De Bresser et al. 2005). Another important, but poorly
understood from the standpoint of deformation, aspect of
fluid-rock interaction is element diffusion and mineral reactions (Burlini and Bruhn 2005) that will most certainly take
place during deformation of such complex rocks as
carbonatites (see Textural record of the postmagmatic
evolution of carbonatites).
Deformed carbonatites: selected examples
Intrusive carbonatites in extensional intraplate settings typically exhibit equigranular polygonal textures lacking any evidence of strain-driven changes (Fig. 13a), or less regular
low-strain fabrics where stress was accommodated through
Textural variations in carbonatites
high-density twinning and inragranular slip (Fig. 13b). There
are also many examples of carbonatites whose postmagmatic
evolution involved intense deformation. These primarily include intrusions emplaced either in collisional settings (e.g.,
Chakhmouradian et al. 2008; Xu et al. 2015), or along rifted
continental margins and overprinted by younger tectonic
events (e.g., Casquet et al. 2008; Chakhmouradian et al.
2015a). Deformation in the brittle regime produces fault breccias, which may incorporate material from several discrete
sources, including silicate wall rocks (Figs. 8d and 13d, e),
and can be distinguished from intrusive breccias by the absence of textural characteristics indicative of xenolith-magma
interaction (cf. Figs. 4h and 6b). At deeper levels in the crust,
fracturing becomes increasingly localized; strongly foliated
zones of comminuted material are intercalated with domains
showing little textural change (Fig. 14a) or predominantly
low-strain ductile deformation (Fig. 14b). The zones of
cataclasis serve as preferential passageway for fluids, facilitating localized metasomatic changes described in the previous
section (Fig. 14a, b).
Carbonatites deformed under high-strain low-T conditions,
but lacking any evidence of cataclasis, are uncommon. These
are competent foliated rocks composed of strongly deformed
grains showing high aspect ratios and twin densities
(Fig. 14c); also notable is their crystallographic alignment
with {0001} subparallel to the shear plane, i.e. similar to preferred orientations observed in experiment at high γ values
(see Fig. 12 of Pieri et al. 2001). Phlogopite and chlorite also
exhibit bending, undulatory extinction, conspicuous dimensional and crystallographic preferred orientation. Most other
minerals behave rigidly, i.e. fracture and rotate into the shear
plane (e.g., apatite in Fig. 14c).
At elevated temperatures, the onset of dynamic recrystallization is marked by the bulging and migration of grain boundaries, subgrain rotation (Fig. 15a) and the appearance of numerous small grains fringing porphyroclasts as Bcore-andmantle^ structures (Fig. 15b). High temperatures expedite dislocation creep, leading to highly irregular, interfingering
boundaries, amoeboid grain shapes and Bisland structures^
(Fig. 15c). In contrast to experiments on relatively homogeneous materials, recrystallization of carbonatites is accompanied not only by the removal of structural imperfections (i.e.,
twins and defects) from calcite and dolomite, but also by their
compositional Bpurification^. Poikilitic porphyroclasts are recrystallized to inclusion-free grains, whereas minerals present
in the protolith as chadacrysts are either dissolved and lost, or
precipitated in the fine-grained interstitial aggregate alongside
the major carbonate phase (Fig. 15d, e). These deformationinduced chemical reactions may have significant implications
for rare metal dispersal and concentration, but to the best of
our knowledge, have not been explored by previous workers
in any detail. Fully dynamically recrystallized carbonatites
(ultramylonites) are rare and comprise a fine aggregate of
353
Fig. 14 High-strain deformation textures in carbonatites; Aley (XPL
images, SB 1 mm). a Localized shearing in calcite carbonatite
facilitating crystallization of extremely fine-grained dolomite and
richterite. b Localized cataclasis of dolomite carbonatite; primary dolomite is partially replaced by calcite and goethite (Gt) in the cataclased
material; note preferred orientation of dolomite and apatite crystals. c
Elongate and intensely twinned calcite crystals with an aspect ratio of
~4.5 in strong dimensional and crystallographic preferred orientations
(the latter showing as parallel extinction in XPL)
newly formed carbonate grains intermixed with fragments of
associated rigid minerals (Fig. 15f). Depending on the proportion of non-carbonate material, a secondary preferred orientation reported in experiments (Pieri et al. 2001; Barnhoorn et al.
2004) may or may not be observable. Although the shear
sense and strain can theoretically be determined from petrographic observations (Pieri et al. 2001; Rybacki et al. 2003),
the applicability of experimental data to rocks is limited by the
(sometimes extreme) modal and structural heterogeneity of
carbonatitic protoliths (see below). In the absence of S1-S2
fabric, tension gashes filled with late-stage minerals
354
A.R. Chakhmouradian et al.
Fig. 15 Ductile deformation in
carbonatites (a-c, e-h: XPL
images; d: BSE image). a
Irregularly shaped calcite grains
showing evidence of grainboundary migration and subgrain
calving and rotation around a rigid magnesiohastingsite grain
(Mhs); Argor, Canada (SB
0.5 mm). b Dynamic recrystallization of highly strained coarse
dolomite porphyroclasts with an
aspect ratio of ~5 (SB 1.5 mm),
producing nearly equant
untwinned grains 40–180 μm
across shown in the inset (SB
0.2 mm); Upper Fir, Canada. c
Amoeboid calcite grains and
Bisland^ structures; Goldray (SB
1 mm). d Dynamic recrystallization of calcite containing primary
inclusions of dolomite and
burbankite (black and white, respectively), producing discrete
grains of inclusion-free calcite
(blue arrows) and dolomite
(yellow arrows); Aley (SB
0.5 mm). e BCore-and-mantle^
texture in dolomite carbonatite;
note undulatory extinction and the
abundance of primary calcite and
burbankite inclusions in
porphyroclasts and their absence
in the recrystallized material;
Chipman Lake (SB 1 mm). f
Completely recrystallized apatitedolomite carbonatite showing
primary and secondary foliations
(S1 and S2, respectively); Aley
(SB 1 mm). g Quartz-filled tension gashes in relict mediumgrained dolomite (Dol1)
surrounded by recrystallized dolomite (Dol2); note rhombohedral
dolomite (Dol3) lining the gashes;
the sense of strain is indicated by
yellow arrows; Aley (SB 1 mm).
h Phlogopite Bfish^ indicating the
sense of strain in calcite
carbonatite; Aley (SB 1 mm)
(Fig. 15g) and phlogopite Bfoliation fish^ (Fig. 15h) can be
useful shear-sense indicators.
One important aspect of carbonate rock deformation virtually neglected in the literature is separation of carbonate and
Textural variations in carbonatites
355
Fig. 16 Dynamic phase and
grain segregation in carbonatites
(a, b, e hand specimens, the coin
is~2 cm in diameter; c, d XPL
images; f: BSE image; SB 1 mm).
a Folded strongly foliated
carbonatite with segregated
apatite- and dolomite-rich cleavage domains; Aley. b Strongly
foliated dolomite carbonatite
showing dynamic grain
Bsorting^; Upper Fir. c Lenses of
recrystallized segregated apatite
and dolomite; Aley. d Grain-size
variations indicative of dynamic
segregation (small grains are<
0.3 mm in size and have a large
aspect ratio; larger grains are, on
average, 1.2 mm across and have
small aspect ratios); Aley. e Flowstrung and locally boudinaged
xenoliths (xen) of fenite (fen) in
an ex situ carbonatite dike; Aley. f
Foliated Bhybrid^ calcitedolomite carbonatite produced by
ductile confluence of material
(primary dolomite and calcite,
Cal1) from at least two different
sources; Aley; note a veinlet of
Sr-poor hydrothermal calcite
(Cal2) transecting the metamorphic fabric
non-carbonate phases, as well as different grain size fractions
during flow. It is reasonable to expect that particle segregation
processes analogous to those occurring in viscous fluids (e.g.,
flow differentiation in magmas and size sorting in pyroclastic
356
flows: Barrière 1976; Félix and Thomas 2004) also take place
in a moving mass of carbonate grains subjected to shear. Indeed, strongly deformed carbonatites often exhibit dramatic,
foliation-controlled variations in modal composition and grain
size at small spatial scales (Figs. 15g and 16a–d). The physics
of particle interaction and segregation is extremely complex
and multiparameteric. Calculations and simulations on simplified viscoplastic flows indicate that the efficiency of separation will depend on particle size, density, shape, shear gradient, matrix viscosity, flow channel geometry, and even such
less-obvious factors as the relative proximity of rigid particles
to one another and thermal gradient owing to frictional heating
(Tripathi and Khakhar 2011; Fan 2011; Chou et al. 2014).
Evidently, analysis of dynamic segregation in metamorphic
rocks is complicated by the sensitivity of some of the above
parameters (in particular, viscosity) to both T and P. Further
discussion of this topic is beyond the scope of the present
paper because the currently available models and experimental data are essentially restricted to gravity-driven binary granular and fluid flows at ambient T and P. It is also noteworthy
that some of the previously studied systems have not yielded
consistent results. The importance of future work in that direction, especially that focused on polymineralic rocks under
stress, is illustrated with Fig. 17, where a carbonatite intrusion
containing primary mineralization (e.g., pyrochlore, bastnäsite
or apatite) is schematically shown to undergo collision zonestyle deformation (Fig. 17a). Cumulate layers within the intrusion, where the mineralization is concentrated
(e.g., Mitchell 2015; our Fig. 4c, d), have a low carbonate
content but significantly higher density than the material separating them and, thus, will behave very differently under
stress. At depths below the brittle-ductile transition,
carbonatite will flow en masse (as opposed to spatially localized cataclastic and transitional forms of movement at
shallower levels), disrupting the pre-existing igneous structures. The ore-rich horizons will become ruffled, thinned, segmented and partially recrystallized into discontinuous lenticular units (Fig. 16c) of variable length and thickness
(Fig. 17b). Both concentration and dispersal of ore minerals
are expected to occur concurrently due to particle segregation
(see above) and flow-induced erosion of the cumulate layers,
respectively. Tectonically mobilized carbonatites may be
emplaced into fractured silicate wall rocks as dikes and entrain
boudinaged xenoliths (Fig. 16e), but lack evidence of contact
metasomatism or in situ crystallization (cf. Figs. 4h and 6b).
Our field observations indicate that stress-induced flow is capable of material transport over significant distances, i.e.
carbonatite Bextrusions^ can sometimes be found>100 m
away from their parental body and incorporate material from
two or more different carbonatite units (Figs. 8e and 16f).
Examples of mixed dikes comprising igneous and
metasedimentary material have also been reported in the literature (Le Bas et al. 2004). Clearly, the assessment of mineral
A.R. Chakhmouradian et al.
Fig. 17 Schematic illustration of the effects of deformation on the
homogeneity and distribution of primary mineralization in carbonatites.
a Carbonatite intrusion containing traceable ore horizons prior to
deformation. b Deformed carbonatite body with multiple ex situ
carbonatite dikes, large- and small-scale heterogeneities in carbonatite
composition and texture, and greatly modified ore distribution
potential in these cases will be much more challenging than
for undeformed deposits and require a good understanding of
the local tectonic framework and interrelations between the
observed geological structures and modal variations in the
carbonatite.
Acknowledgments This work was supported by the Natural Sciences
and Engineering Research Council of Canada (NSERC) and St. Petersburg State University, Russia (3.38.690.2013, including Geomodel Center). The instrumentation used for data collection was supported by the
NSERC. We would like to thank Taseko Mines Ltd. and Rare Element
Resources for providing access to their Aley and Bear Lodge properties
(respectively). Expert guidance of Jörg Keller at Kaiserstuhl, Jim Clark at
Bear Lodge, and Pete Modreski at Iron Hill is most gratefully acknowledged. Most of the samples examined in the present work were collected
by authors from outcrop and drill core, but some were loaned to us by the
Royal Ontario Museum (Toronto, Canada), Natural History Museum
(London, UK), or donated by Francis Ö. Dudás, Meghan A. Moore and
Alexey Rukhlov. We would also like to thank Lia N. Kogarko and Felix
V. Kaminsky for their constructive comments on the earlier version of this
paper, Johann G. Raith for his keen editorial eye, as well as Vincent
Vertolli and David Smith for arranging the museum loans.
Textural variations in carbonatites
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