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The Catalão I niobium deposit, central Brazil: Resources, geology and
Pedro Filipe de Oliveira Cordeiro
⁎, José Affonso Brod
, Matheus Palmieri
Claudinei Gouveia de Oliveira
, Elisa Soares Rocha Barbosa
, Roberto Ventura Santos
José Carlos Gaspar
, Luis Carlos Assis
Universidade de Brasília, Campus Darcy Ribeiro ICC Central, Instituto de Geociências, Brasília-DF, 70910-900 Brazil
Universidade Federal de Goiás, Campus Samambaia, Instituto de Estudos Sócio-Ambientais, Universidade Federal de Goiás, Goiânia-GO, 74001-970 Brazil
Anglo American Brazil LTDA, Avenida Interlândia 502, Setor Santa Genoveva, Goiânia-GO, 74672-360 Brazil
Received 24 March 2010
Received in revised form 23 June 2011
Accepted 24 June 2011
Available online 22 July 2011
The Catalão I alkaline–carbonatite–phoscorite complex contains both fresh rock and residual (weathering-
related) niobium mineralization. The fresh rock niobium deposit consists of two plug-shaped orebodies
named Mine II and East Area, respectively emplaced in carbonatite and phlogopitite. Together, these
orebodies contain 29 Mt at 1.22 wt.% Nb
(measured and indicated). In closer detail, the orebodies consist
of dike swarms of pyrochlore-bearing, olivine-free phoscorite-series rocks (nelsonite) that can be either
apatite-rich (P2 unit) or magnetite-rich (P3 unit). Dolomite carbonatite (DC) is intimately related with
nelsonite. Natropyrochlore and calciopyrochlore are the most abundant niobium phases in the fresh rock
deposit. Pyrochlore supergroup chemistry shows a compositional trend from Ca–Na dominant pyrochlores
toward Ba-enriched kenopyrochlore in fresh rock and the dominance of Ba-rich kenopyrochlore in the
residual deposit. Carbonates associated with Ba-, Sr-enriched pyrochlore show higher δ
expected for carbonates crystallizing from mantle-derived magmas. We interpret both the δ
pyrochlore chemistry variations from the original composition as evidence of interaction with low-
temperature ﬂuids which, albeit not responsible for the mineralization, modiﬁed its magmatic isotopic
features. The origin of the Catalão I niobium deposit is related to carbonatite magmatism but the process that
generated such niobium-rich rocks is still undetermined and might be related to crystal accumulation and/or
emplacement of a phosphate–iron-oxide magma.
© 2011 Elsevier B.V. All rights reserved.
Brazil is the largest niobium producer in the World due to mining
of residual deposits overlying the Araxá and Catalão I and II
carbonatite complexes. These deposits represent more than 85% of
the world's niobium supply. Although these complexes have been
mined for more than 30 years, data from the Araxá niobium deposit is
virtually unavailable and information on the Catalão I (Cordeiro et al.,
2010, 2011) and Catalão II (Palmieri, 2011) deposits was published
only recently. Not only general information is restricted but genetic
interpretation of these niobium deposits is limited to “weathering of
carbonatite related rocks”(Carvalho and Bressan, 1981; Gierth and
Cordeiro et al. (2010, 2011) studied the primary fresh ore and
determined that pyrochlore occurs mostly in apatite- and magnetite-rich
rocks that crosscut previous phoscorites and phlogopitites. According to
the classiﬁcation of Yegorov (1993) for olivine-poor member of the
phoscorite series these unusual rocks are named nelsonite. At Catalão I,
nelsonites are intimately associated with dolomite carbonatites and form
two main swarms of densely-packed thin dikes near the center of the
complex (East Area and Mine II orebodies). The direct relationship
between phoscorite-series rocks and niobium mineralization in fresh
rock has also been suggested in the Catalão II (Palmieri, 2011)andAraxá
niobium deposits (Nasraoui and Waerenborgh, 2001).
Although it is only the second largest niobium deposit in Brazil,
Catalão I is the best understood. Mining of the Catalão I residual deposit
started in 1976 with a reserve of 19 Mt at 1.08 wt.% Nb
(Hirano et al.,
1990; Rodrigues and Lima, 1984)andwasdiscontinuedin2001witha
remaining residual reserve of 9.65 Mt at 1.19 wt.% Nb
while mining focused on the Boa Vista mine in Catalão II. Recent
modeling of the fresh rock deposit indicate a unpublished resource of
21.8 Mt at 1.22 wt.% Nb
for the East Area orebody and 7.2 Mt at
1.23 wt.% Nb
for the Mine II, adding up to a total reserve of
approximately 29 Mt at 1.22 wt.% Nb
for the Catalão I complex.
In this paper we studied drill core samples from the fresh rock
Catalão I deposit collected between depths of 100 and 500 m. Our
Ore Geology Reviews 41 (2011) 112–121
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main aim is to describe the deposit and provide information on
pyrochlore chemistry in order to establish the main crystal chemistry
features and substitutions. We also compare Catalão I pyrochlore
chemical composition with that of Lueshe (Nasraoui and Bilal, 2000),
Oka (Gold et al., 1986; Zurevinski and Mitchell, 2004) and Sokli (Lee et
al., 2006) to contribute for a broader model of pyrochlore evolution in
carbonatite complexes. Finally, we address some points of signiﬁcance
for the bearing of magmatic processes in the origin of a phoscorite-
related niobium deposit.
2. Niobium deposits
Most commercial niobium is from carbonatite-related sources, but
minor production comes as a byproduct of tantalum and tin mining in
pegmatites. In Table 1 we compiled and updated data from Woolley
and Kjarsgaard (2008) on the World's niobium reserves. When
possible, we report only measured, indicated and historical reserves,
hence several resources listed in Table 1 are smaller compared to what
is found in the literature.
There are several carbonatite related niobium deposits worldwide,
comprising residual and/or fresh rock resources, but only the Boa
Vista (Catalão II), CBMM (Araxá) and Niobec (Saint Honoré) deposits
are currently in production (Fig. 1). The number of untapped niobium
deposits in Africa and the general lack of information on the Brazilian
underground resources is noteworthy. Detailed information on
Brazilian carbonatite-related deposits is given by Biondi (2005), but
an equivalent study of African niobium deposits is still to be made.
3. The Alto Paranaíba Igneous Province (APIP)
The APIP is a NW trending province of Late-Cretaceous alkaline
igneous rocks intruding Neoproterozoic rocks of the Brasília Belt,
between the NE border of the Paleozoic Paraná Basin and the SW border
of the Archean São Francisco Craton. The province origin is attributed to
the initialimpact of the Trindade Mantle Plumebeneath Central Brazil at
ca. 85 Ma. According to Gibson et al. (1995) and Thompson et al. (1998),
thinningof the lithosphereunder the BrasíliaBelt allowed mantleplume
heat to penetrate by conduction and advection causing melting of
readily fusible, K-rich parts of the lithospheric mantle.
Xenoliths of perovskite-rich pyroxenite (bebedourite) and pyroxe-
nite in APIP kamafugite lavas and pyroclastics are analogous to
ultramaﬁc rocks occurring in the carbonatite complexes, thus providing
evidence of the intimate association between kamafugites and
Fig. 1. Grade-tonnage data showing metal grades (wt.% Nb) for carbonatite-related
niobium deposits (please refer to Table 1 for references). Stars are deposits in
production and circles represent resources.
Comparison of carbonatite-related niobium deposits (adapted from Woolley and Kjarsgaard, 2008).
Complex Country Status Style Association Resources Reserve Mt Nb
% Nb% Main references
St-Honoré Canada Active mine Primary Nb+REE Measured and indicated 32 0.56 0.39 www.iamgold.com (Resources 2009)
Araxá Brazil Active mine Residual Nb +Fe+P 462 2.48 1.73 Rodrigues and Lima (1984),
Hirano et al. (1990)
Catalão II Brazil Active mine Residual Nb Probable reserve 3.4 1.67 1.17 http://www.cbmm.com.br/ (conference
paper by Guimarães and Weiss)
Lueshe Congo Past producer Residual Nb 30 1.34 0.94 Deans (1966)
Sukulu Uganda Past producer Residual P+ Nb 230 0.25 0.17 Deans (1966); van Straaten (2002)
Oka Canada Past producer Primary Nb Measured, Indicated
and historical reserves
37.5 0.53 0.37 http://www.niocan.com/
(Technical Report February 10 2010)
Catalão I Brazil Past producer
Residual Nb+ Fe+ P 19 1.08 0.76 Rodrigues and Lima (1984), Hirano et al.
Catalão I Brazil Resource Primary Nb +Fe+P Measured and indicated 29 1.22 0.85 This paper
Araxá Brazil Resource Primary Nb+Fe+ P 940 1.6 1.12 Issa Filho et al. (1984)
Tapira Brazil Resource Residual Nb 166 0.73 0.51 Melo (1997)
Bonga Angola Resource Primary Nb 824 0.48 0.34 Pena (1989); Kamitani and Hirano
Bingo Congo Resource Residual Nb + P 13 3.3 2.31 Woolley (2001)
Mrima Kenya Resource Residual Nb+REE 75 0.7 0.49 Deans (1966); Notholt et al. (1990);
Pell (1966); Woolley (2001)
Ondurakorume Namibia Resource Primary P+Nb+REE 8 0.3 0.21 Verwoerd (1967, 1986); Woolley
Tanzania Resource Primary Nb+P 125 0.3 0.21 Deans (1966); Woolley (2001);
van Straaten (2002)
Aley Canada Resource Primary N+P +REE 20 0.7 0.49 Richardson and Birkett (1996)
Canada Resource Primary Ta+ Nb Indicated 23.1 1.14 0.80 www.commerceresources.com
(Technical Report June 20 2007)
Argor Canada Resource Primary Nb+P +Zr 62.5 0.52 0.36 Stockford (1972); Woolley (1987);
Martison Lake Canada Resource Residual P +Nb Measured and indicated 62.2 0.34 0.24 Woolley (1987);www.sedar.com
(Technical Report May 31 2007)
Canada Resource Primary Nb Inferred 49.9 0.43 0.30 www.sarissaresources.com
(Technical Report July 2009)
Seis Lagos Brazil Resource Residual Nb Measured and indicated 239 2.47 1.73 Justo and Souza (1986)
113P.F.O. Cordeiro et al. / Ore Geology Reviews 41 (2011) 112–121
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carbonatites in the APIP (Brod, 1999; Brod et al., 2000, 2001). Those
authors argued in favor of a common subcontinental lithospheric
mantle origin for kamafugites and the parental magma of APIP
complexes (phlogopite picrite). The temporal and spatial association
between these alkaline rocks deﬁnes a kamafugitic–carbonatitic
association in the APIP, similar to those occurring in Italy (Stoppa et
al., 1997; Stoppa and Cundari, 1995)andChina(Yang and Woolley,
APIP carbonatite complexes also host phosphate (Araxá, Catalão I
and II, Tapira and Salitre), titanium (Serra Negra, Salitre, Tapira and
Catalão I), and rare earth (Catalão I) deposits, as well as occurrences of
vermiculite, copper, barite and magnetite. Thus, the APIP is of great
economic interest and can provide key information for mineral
exploration of carbonatite-related deposits.
4. The Catalão I Carbonatite Complex
The Catalão I Complex (Fig. 2) is located in Central Brazil at 18°08′
S, 47°48′W, near the cities of Catalão and Ouvidor. The complex has
intruded quartzites and schists of the Late Proterozoic Araxá Group as
a vertical pipe with a diameter of ~6 km at surface, creating a dome-
like structure. The age of the intrusion is reported by Sonoki and Garda
(1988) as 85 ±6.9 Ma (K–Ar, phlogopite). The complex can be divided
into an outer zone dominated by phlogopitite and an inner zone
composed mostly of dolomite carbonatites and phoscorite-series
The outer zone comprises phlogopitites and rare dunites, pyroxe-
nites and bebedourites (perovskite-rich pyroxenites). Phlogopitite is
interpreted as the result of interaction of former ultramaﬁcrockswith
carbonatite ﬂuids (Brod et al., 2001). Ultramaﬁc relicts within
phlogopitite, which sometimes retain the original mineral assemblage
unaffected by ﬂuid alteration, are a very strong evidence for phlogopi-
tization. The dominance of phlogopitite over other rock types in the
outer zone attests to the extremely intense carbohydrothermal
alteration that occurred in the complex.
The inner zone is composed of magnetite–apatite-rich rocks and
carbonatite. The Catalão I fresh rock deposit is intimately related to
these rocks and can be divided into Mine II and East Area orebodies
(Fig. 3). Mine II is a roughly oval, pipe-like body, 200 m long and 100 m
wide, hosted mainly by dolomite carbonatite. East Area is an L-shaped
orebody, 400 m long, 200 m wide hosted by phlogopitite. Both
orebodies are open at depth and deep drilling conﬁrmed their
extension at until a depth of at least 800 m.
Fig. 4 shows the general pipe-like geometry of East Area and Mine
II orebodies. Despite their shape, the orebodies do not represent
single, homogeneous pyrochlore-bearing magnetite–apatite rocks,
but rather dike swarms up to 2 m wide and plugs up to 10 m wide.
The main Nb-mineral within the orebodies is pyrochlore. Aside
from pyrochlore modal content, ore grades are also controlled by
Fig. 2. Geological sketch of the Catalão I Complex. The fresh rock niobium deposit occurs in the center of the complex comprising the nelsonite unit.
Adapted from Brod et al. (2004).
Fig. 3. Combination of an Ikonos image showing the roughly circular Mine II open pit
and a 3-D model of the Mine II and East Area orebodies.
114 P.F.O. Cordeiro et al. / Ore Geology Reviews 41 (2011) 112–121
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frequency and width of nelsonite dikes and can be largely diluted by
barren wallrocks (Fig. 5A, phlogopitite). Because of their dike-like,
plug-like and vein-like shape, these terms are used in a generally
The occurrence of magnetite–apatite rich rocks, named phoscorite,
in carbonatite complexes was reported by several authors (Krasnova
et al., 2004 and references therein) and due to their rarity,
nomenclature remains problematic. A discussion on phoscorite series
rocks is provided in Krasnova et al. (2004). We favor an adapted
version of the nomenclature of Yegorov (1993) as used by Cordeiro et
al. (2010, 2011). Therefore, phoscorite is an olivine-, phlogopite-,
apatite- and magnetite-bearing rock and nelsonite is a phlogopite-,
apatite- and magnetite-bearing rock.
In Catalão I, the phoscorite-series can be divided into two stages,
according to mineral chemistry and modal mineralogy. Early-stage
phoscorites are grouped under the P1 unit (Fig. 5B). Their main
characteristics are a) breccia structure; b) emplacement as small plugs
and dikes; c)no obvious direct relationship with carbonatite; d) olivine
occurs as altered to minute tetra-ferriphlogopite, indicating interaction
with carbohydrothermal ﬂuids; and e) pyrochlore is rare (although this
stage is an important source of apatite for the Catalão I residual
phosphate deposit). Late-stage P2 and P3 units (Fig. 5B and C) are
nelsonites and represent the bulk of the fresh rock niobium mineral-
ization. Nelsonites can be distinguished from early-stage phoscorites by
a) emplacement as dikes and small plugs; b) occurrence of internal
pockets of dolomite carbonatite; c) no visible evidence of carbohy-
drothermal alteration; d) absence of olivine; e) abundant pyrochlore,
reaching up to 50 vol.% in some samples.
Dolomite carbonatite is abundant, but upto 15 m wide plugs and up
to 2 m wide dikes of calcite carbonatite occur. Phoscorite is intensely
crosscutby dolomite carbonatite dikes, whichoriginates the breccia-like
aspect of these rocks. Widespread alteration of olivine crystals within
phoscorite to tetra-ferriphlogopite suggests that the inner zone was also
affected by carbohydrothermal ﬂuids. Nelsonite, on the other hand,
shows no sign of metasomatic alteration, indicating that its emplace-
ment occurred later, after the widespread alteration event.
Carbonatites, particularly dolomite carbonatites, dikes and plugs
are widespread in Catalão I and are especially abundant within P1.
One particular set of dolomite carbonatite, designated here DC, is
intimately related to P2 and P3 and may occur within them as
centimetric to metric pockets as well as dikes and plugs. DC can be
easily discriminated from earlier generations of dolomite carbonatites
by the absence of olivine and presence of pyrochlore and ilmenite.
4.1. Primary ore
Primary (fresh) rock ore in Catalão I is represented by nelsonite
dikes, but subordinate DC dikes with more than 1% modal pyrochlore
occur. P2 nelsonite is apatite-rich and its essential silicate phases are
tetra-ferriphlogopite crystals with phlogopite cores. Apatite is
prismatic, frequently zoned with cores surrounded by a ﬂuid
inclusions-rich rim. Magnetite is interstitial and may contain very
thin (ca. b0.01 mm) ilmenite lamellae.
P3 is magnetite-rich (apatite/magnetite b0.8 vol.%) and its
essential silicate phase is tetra-ferriphlogopite. In contrast to P2,
aluminous phlogopite cores are virtually absent. Apatite is prismatic
to rounded, but also occurs as aggregates of anhedral crystals, usually
associated with massive anhedral magnetite clusters. Magnetite
forms interstitial masses and may reach up to 71 vol.%.
Dolomite carbonatite (DC) occurs as pockets within nelsonites and
also as independent dikes and veins. Although other dolomite
carbonatite phases occur in the complex, the variety genetically
related to nelsonites crosscuts all rock types. DC dikes are believed to
Fig. 4. Schematic model of the fresh rock niobium ore, where apatite nelsonite P2, magnetite nelsonite P3, and dolomite carbonatite DC crosscut phlogopitite. The detail shows the
common textural feature of DC pockets.
Fig. 5. Main rock types of the Catalão I Nb-deposit. A. Phlogopitite with relicts of the
original ultramaﬁc rock cut by a magnetite nelsonite dike (P3) with dolomite
carbonatite (DC) pockets. B. Coarse-grained phoscorite (P1), cut by P3 dikes with DC
pockets. C. Equigranular apatite nelsonite (P2) with DC pockets.
115P.F.O. Cordeiro et al. / Ore Geology Reviews 41 (2011) 112–121
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represent the product of extraction of carbonatite from the nelsonite
crystallizing assemblage (Cordeiro, 2009). The pyrochlore content in
DC varies, but it is hardly more than 5 vol.%.
Pyrochlore from P2 and P3 nelsonites are texturally similar and
generally ﬁne-grained. They occur as anhedral to subhedral brownish
or yellowish crystals often showing optical zoning (Fig. 6A, B). DC
pyrochlore is often euhedral to subhedral and may occur as inclusions
in ilmenite, together with betaﬁte and columbite, and in magnetite
(Cordeiro, 2009). It is medium- to ﬁne-grained, often optically zoned
(Fig. 6C). Aggregates of pyrochlore and apatite occur within DC
5. Pyrochlore chemistry
Pyrochlore composition was determined by WDS using a CAMECA
SX-50 electron microprobe at the University of Brasília. The analytical
conditions were beam diameter 2 μm, 20 kV, 20 nA and two minute
count times. Detection limits varied between 0.01 and 0.05 wt.%,
except Nb and Ta (0.07%) and La, Ce and Y (0.2%).
The pyrochlore general formula is A
et al., 2010; Lumpkin and Ewing, 1995). The Asite is occupied by large
anions suchas As, Ba, Bi, Ca, Cs, K, Mg, Mn, Na, Pb, REE, Sb,Sr, Th, U and Y.
The Bsite comprises smaller and highly charged cations such as Nb, Ta,
Ti, Zr, Fe
, Al and Si (Zurevinski and Mitchell, 2004)andrarelyW
(Caprilli et al., 2006). The Yand Xanions can be O, OH and F. Vacancies
arecommonintheAand Ysites. In this paper, pyrochlore has been
calculated to produce a total of 2 cations in the Bsite (Wall et al., 1996).
Pyrochlore classiﬁcation is originally described by Hogarth (1977)
but an up to date CNMNC-IMA-approved nomenclature was published
by Atencio et al. (2010). The new nomenclature is composed of two
preﬁxes and a root name based on the content of Y,Aand Bsites. The Y
site content (cation, anion, H
O or vacancy) determines the ﬁrst preﬁx
and the A site content refers to the second preﬁx. The dominant
element in the Bsite determines the root name: pyrochlore (Nb),
microlite (Ta), roméite (Sb), betaﬁte (Ti) and elsmoreite (W).
The abundance of Nb over other Bsite elements classiﬁes Catalão I
pyrochlore supergroup minerals within the pyrochlore group (Fig. 7).
Pyrochlore representative compositions are shown in Table 2.Data
published by Fava (2001) indicates that more than 95% of all Catalão I
fresh rock pyrochlore exceeds 0.5 apfu and therefore should have the
preﬁxﬂuor. However, we haven't analyzed ﬂuorine and Atencio et al.
(2010) suggest preﬁxes should be droped in face of lack of data to avoid
misclassiﬁcation. Therefore the ﬁrst preﬁx won't be used in this paper.
Fig. 6. Texturalcharacteristicsof nelsonitespyrochlore.A. P2 nelsonite withsubhedral, brownto orange pyrochlore.B. P3 nelsonitewith anhedral tosubhedral brownto orange pyrochlore.
C. Sector zoning in pyrochlore from DC. D. P2 aggregates within DC, crossed polars. (Mag = magnetite; Apt = apatite; TFP = tetra-ferriphlogopite; Carb= carbonate; Pcl = piroclore).
Fig. 7. Triangular Nb–Ti–Ta pyrochlore classiﬁcation scheme (Atencio et al., 2010;
Hogarth, 1977, 1989) showing fresh rock pyrochlore as black circles. Outlines for
pyrochlore compositions from the Catalão I residual deposit (square pattern, Fava,
2001), Oka (gray, Gold et al., 1986; Zurevinski and Mitchell, 2004), Sokli (solid black
outline; Lee et al., 2004, 2006) and Salitre (dotted black outline, Barbosa, 2009) are
shown for comparison. BET = betaﬁte, PCL = piroclore, MCL = microlite.
116 P.F.O. Cordeiro et al. / Ore Geology Reviews 41 (2011) 112–121
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content varies from 50 to 70 wt.%. The average TiO
ranges from 3 to 5 wt.%, but may reach up to 17 wt.% in natropyrochlore
inclusions in ilmenite from DC. Most analysis show low Ta
from b1 wt.% to a maximum of 2 wt.% in one grain from P2 and in
pyrochlore crystals within DC ilmenite. ZrO
reach up to 5 and
3 wt.% respectively.
Pyrochlore from fresh rock nelsonite has Ca and Na as the main A
site elements, ranging up to 19 and 8 wt.%, respectively. Therefore
calciopyrochlore dominates in the Catalão I Nb deposit followed by
natropyrochlore. Ba is one of the most common substitutes for both
Na and Ca in this site and BaO content reaches 18 wt.% whereas SrO
may reach 7 wt.%. The sum of the analyzed rare earth (La+Ce+Y)
oxides varies from 3.5 to 6 wt.%. ThO
is up to 6 wt.% but its average
content is b2 wt.%. UO
is up to 4 wt.%, averaging b1 wt.%. FeO may
reach 3 wt.%, and MnO is always below 1 wt.%. Several analyses
indicate the occurrence of kenopyrochlore (zero-valent-dominant
pyrochlore) in the Catalão I Nb fresh rock deposit but according to
data from Fava (2001) they are dominant in the residual ore.
Several studies tried to explain the evolution of pyrochlore
composition throughout magmatic evolution (Chakhmouradian and
Williams, 2004; Hogarth et al., 2000; Knudsen, 1989). During the early
stages of carbonatite magmatism Nb and Ta are probably transported
as phosphate and ﬂuorine complexes, which might explain the
common correlation between the occurrence of apatite and pyrochlore
(Hogarth et al., 2000; Knudsen, 1989). Knudsen (1989) argued that
during the carbonatitic magmatism Nb is more soluble than Ta, which
could explain the occurrence of Ta-rich pyrochlore in primitive
magmas and Nb-rich, Ta-poor pyrochlore in more evolved, late stage
ones. Hogarth et al. (2000) concluded that the normal path of
evolution of pyrochlore in carbonatites is one of progressive
enrichment in Na, Ca and Nb and depletion in Ta, Th, REE, Ti and U. A
more detailed evolution scheme is proposed by Chakhmouradian and
Williams (2004) where Th-enriched Ca–Na dominant pyrochlore
evolves toward Ba–Sr-rich compositions in calcite–dolomite carbona-
tites from Kola carbonatite complexes.
Pyrochlore from the Catalão I fresh rock deposit seems to ﬁt well
into the proposed scheme, but early Ta–Th–U enriched pyrochlore
phases are not present. Pyrochlore analyses of several stages of Sokli
phoscorites (Lee et al., 2004; 2006) conﬁrm the trend of early Ta–U–Th
pyrochlore toward evolved Ca–Na compositions. Thus, Ca–Na pyro-
chlore in P2 and P3 nelsonites and related DC dolomite carbonatites
can be interpreted as belonging to a more evolved phase, similar to the
late-stage D5 dolomite carbonatite phase in the Sokli complex.
5.1. Chemical evolution of pyrochlore
The range of P2 and P3 pyrochlore chemical compositions overlaps
widely and we could not ﬁnd an applicable chemical criterion to
Representative compositions of pyrochlore group minerals from the Catalão I primary niobium deposit (b.d. = below detection limit; calcio = calciopyrochlore; keno =
kenopyrochlore; natro = natropyrochlore).
056-2 178-1 183-
Type calcio calcio calcio calcio calcio calcio calcio calcio calcio calcio calcio calcio keno keno keno keno keno natro natro natro
Unit P2 P2 P2 DC P2 P3 P2 P3 P3 P3 P3 DC P3 DC P3 P3 DC DC P2 DC
62.66 59.26 55.76 59.93 61.68 55.58 55.26 64.26 63.14 63.39 63.39 68.71 59.99 50.10 62.17 52.26 63.96 72.56 63.76 52.85
0.15 b.d b.d. b.d. 0.33 0.70 0.16 b.d. b.d. b.d. b.d. 0.28 0.81 0.77 0.57 0.80 0.81 1.61 0.37 0.92
b.d. b.d. 0.16 0.12 0.61 0.57 b.d. b.d. 0.04 b.d. b.d. b.d. 1.20 2.93 b.d. 0.61 1.10 b.d. b.d. b.d.
3.52 4.64 5.59 6.15 3.15 4.16 3.67 3.91 4.71 4.35 4.35 2.04 3.16 5.27 4.87 2.37 1.26 0.78 4.20 17.35
0.17 2.05 0.90 2.13 0.13 3.95 0.26 1.78 1.65 0.53 0.53 0.53 b.d. 2.44 0.32 3.20 0.75 b.d. 0.94 0.09
0.36 b.d. b.d. 1.02 0.59 2.35 0.19 0.14 b.d. b.d. b.d. b.d. 0.77 1.01 1.17 3.72 0.12 b.d. 0.82 b.d.
1.08 3.39 2.13 2.04 1.09 2.69 1.44 1.12 1.13 2.40 2.40 0.22 0.41 2.15 4.66 4.94 0.74 0.19 1.72 b.d.
1.21 0.39 0.75 0.62 1.14 0.62 0.87 0.95 0.68 0.71 0.71 1.63 1.30 0.32 1.13 0.42 0.92 0.37 0.96 0.35
4.20 2.90 2.85 2.37 3.42 2.92 3.09 2.47 2.00 2.72 2.72 3.26 3.37 2.91 4.09 3.04 3.54 0.73 2.68 0.24
0.44 0.57 0.49 0.55 0.32 0.26 0.34 0.45 0.46 0.53 0.53 0.57 0.26 0.40 0.48 0.20 0.68 0.39 0.55 0.20
FeO 0.14 0.50 0.70 0.40 0.94 1.93 0.31 0.86 0.46 0.20 0.20 0.18 0.69 4.47 0.40 1.49 0.77 0.18 0.16 4.22
MnO b.d. b.d. b.d. b.d. b.d. 0.36 b.d. b.d. b.d. b.d. b.d. 2.00 0.11 b.d. b.d. 0.08 b.d. b.d. b.d. 0.90
CaO 9.02 14.31 14.46 16.14 7.46 8.53 14.50 13.11 15.74 11.86 11.86 9.87 5.11 2.80 8.51 3.34 0.12 11.32 12.05 10.98
BaO 0.35 b.d. 0.24 b.d. 4.89 3.67 0.13 b.d. b.d. b.d. b.d. b.d. 11.03 14.61 2.81 12.24 15.20 0.18 b.d. b.d.
SrO 2.78 0.69 1.06 1.17 3.95 3.48 2.41 2.29 1.59 2.35 2.35 2.50 4.65 2.23 2.03 3.56 0.75 4.61 2.08 3.01
O 5.59 4.23 3.83 4.71 4.09 2.52 2.96 5.94 6.09 5.93 5.93 7.31 0.34 0.77 1.16 0.38 1.29 7.83 6.46 7.75
Total 91.67 92.97 89.07 97.41 93.87 94.28 85.59 97.32 97.81 95.04 95.04 97.17 93.25 93.18 94.43 92.65 92.01 100.84 96.78 98.92
Structural formulae calculated based on ∑B-site elements = 2
Nb 1.82 1.71 1.68 1.65 1.80 1.62 1.79 1.77 1.73 1.78 1.78 1.88 1.75 1.46 1.75 1.70 1.83 1.94 1.77 1.28
Ta b.d. b.d. b.d. b.d. 0.01 0.01 b.d. b.d. b.d. b.d. b.d. 0.01 0.01 0.01 0.01 0.02 0.01 0.03 0.01 0.01
Si b.d. b.d. 0.01 0.01 0.04 0.04 b.d. b.d. b.d. b.d. b.d. b.d. 0.08 0.19 b.d. 0.04 0.07 b.d. b.d. b.d.
Ti 0.17 0.22 0.28 0.28 0.15 0.20 0.20 0.18 0.22 0.20 0.20 0.09 0.15 0.26 0.23 0.13 0.06 0.04 0.19 0.70
Zr 0.01 0.06 0.03 0.06 b.d. 0.13 0.01 0.05 0.05 0.02 0.02 0.02 b.d. 0.08 0.01 0.11 0.02 b.d. 0.03 b.d.
2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00
U 0.01 b.d. b.d. 0.01 0.01 0.03 b.d. b.d. b.d. b.d. b.d. b.d. 0.01 0.01 0.02 0.06 b.d. b.d. 0.01 b.d.
Th 0.02 0.05 0.03 0.03 0.02 0.04 0.02 0.02 0.02 0.03 0.03 b.d. 0.01 0.03 0.07 0.08 0.01 b.d. 0.02 b.d.
La 0.03 0.01 0.02 0.01 0.03 0.01 0.02 0.02 0.02 0.02 0.02 0.04 0.03 0.01 0.03 0.01 0.02 0.01 0.02 0.01
Ce 0.10 0.07 0.07 0.05 0.08 0.07 0.08 0.06 0.04 0.06 0.06 0.07 0.08 0.07 0.09 0.08 0.08 0.02 0.06 b.d.
Y 0.02 0.02 0.02 0.02 0.01 0.01 0.01 0.01 0.01 0.02 0.02 0.02 0.01 0.01 0.02 0.01 0.02 0.01 0.02 0.01
Fe2 0.01 0.03 0.04 0.02 0.05 0.10 0.02 0.04 0.02 0.01 0.01 0.01 0.04 0.24 0.02 0.09 0.04 0.01 0.01 0.19
Mn b.d. b.d. b.d. b.d. b.d. 0.02 b.d. b.d. b.d. b.d. b.d. b.d. 0.01 b.d. b.d. 0.01 b.d. b.d. b.d. 0.04
Ca 0.62 0.98 1.03 1.05 0.52 0.59 1.11 0.86 1.02 0.79 0.79 0.64 0.35 0.19 0.57 0.26 0.01 0.72 0.79 0.63
Ba 0.01 b.d. 0.01 b.d. 0.12 0.09 b.d. b.d. b.d. b.d. b.d. b.d. 0.28 0.37 0.07 0.35 0.38 b.d. b.d. b.d.
Sr 0.10 0.03 0.04 0.04 0.15 0.13 0.10 0.08 0.06 0.08 0.08 0.09 0.17 0.08 0.07 0.15 0.03 0.16 0.07 0.09
Na 0.70 0.52 0.50 0.55 0.51 0.32 0.41 0.70 0.72 0.71 0.71 0.86 0.04 0.10 0.14 0.05 0.16 0.90 0.77 0.81
1.60 1.70 1.75 1.79 1.50 1.42 1.79 1.79 1.91 1.73 1.73 1.73 1.03 1.12 1.09 1.14 0.75 1.83 1.78 1.78
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discriminate pyrochloresfrom the two units. According to Cordeiro et al.
(2010) other minerals from P2 and P3, such as phlogopite and apatite,
aren't discernible from each other. This suggests only minor chemical
differences in the magmas that produced the two units. Therefore, the
compositional spread seenin our data might be related to factors such as
a) zoning (Chakhmouradian and Mitchell, 2002; Hogarth et al., 2000);
b) hydrothermal alteration (Chakhmouradian and Mitchell, 1998;
Geisler et al., 2004); c) weathering (Lumpkin and Ewing, 1995; Wall
et al., 1996). Chakhmouradian and Zaitsev (1999) point out that several
types of pyrochlore may be found in the same complex or even within
the same facies. Hence, in order to address the compositional variation
we need a classiﬁcation criterion other than lithology.
Lumpkin and Ewing (1995) argued that Asite large cations such as
K, Ba and Sr can be useful in the identiﬁcation of pyrochlore chemical
variation because their occurrence is related to the host rock
alteration. Accordingly, we adopted a division based on the variation
of the Asite content, allowing us to discriminate between three
pyrochlore groups (Fig. 8): a) calciopyrochlore; b) natropyrochlore;
and c) kenopyrochlore.
The occurrence of vacancy in pyrochlore from the fresh rock deposit
is an important feature of its evolution. Vacancy, sometimes accompa-
nied by Ba-enrichment, was attributed to alteration at Sokli (Lee et al.,
2006), to hydrothermal overprint at Oka (Zurevinski and Mitchell,
2004) and Lueshe (Nasraoui and Bilal, 2000) and to both oscillatory
zoning and alteration in the Bingo carbonatite (Williams et al., 1997).
The trend from calciopyrochlore and natropyrochlore toward
kenopyrochlore illustrated in Fig. 8 is related to the exchange of Ba for
Ca+Na, and consequent vacancy, in the Asite. Similar trends can be
found in Lueshe (Nasraoui and Bilal, 2000) and Bingo (Williams et al.,
1997) pyrochlore, described as product of weathering, and in Kola
carbonatites pyrochlore (Chakhmouradian and Williams, 2004), as
derived from supergene or low-temperature hydrothermal alteration.
An additional trend from calciopyrochlore toward natropyrochlore
can be considered. Despite considerable scatter, calciopyrochlore,
natropyrochlore and the ﬁelds of Oka and Salitre fresh rock pyrochlore
show an overall alignment to the 1:1 line in a Na vs. Ca diagram. This
trend is even more marked in crystal core composition from the
Catalão I weathered pyrochlore deposit (Fava, 2001). Taking into
account that most natropyrochlore are inclusions in ilmenites from
the last stage of magmatic evolution in the deposit (DC unit) and that
ilmenite is one of the last minerals to crystallize, natropyrochlore
formed at such stage would represent the most evolved pyrochlore
composition. Therefore, we interpret that the negative correlation of
Ca and Na represent the evolution from earlier calciopyrochlore
toward a late stage natropyrochlore.
5.2. Comparison with pyrochlore from the residual deposit
Pyrochlore chemistry from the Catalão I fresh rock and residual
deposits shows no clear differences in the Bsite, but some important
substitutions occur the Asite. Fava (2001) described the mineralogical
characteristics of pyrochlore from the residual deposit developed over
Catalão I nelsonites and carbonatites and concluded that weathering
induced substitutions in the Asiteandoriginatedbariopyrochlore,
renamed here as “Ba-enriched”kenopyrochlore according to the
nomenclature of Atencio et al. (2010). On the other hand, Catalão I fresh
rock and residual Ba-enriched kenopyrochlore are different from each
other. Fresh rock crystals show a negative Sr–Ca correlation that leads
toward Sr-enriched kenopyrochlore and the same correlation occurs in
the residual deposit pyrochlore crystal cores (Fava, 2001). However, the
majority of pyrochlore in the residual deposit is Ba-enriched kenopyro-
chlore that lack a negative Sr–Ca correlation. These features suggest that
different processes originated fresh rock kenopyrochlore and the residual
We suggest that the chemical shift from calciopyrochlore and
natropyrochlore toward kenopyrochlore in the Catalão I fresh rock
deposit is due to interaction with hydrothermal ﬂuids that also carried
Sr. Later weathering-related ﬂuids originated the residual deposit Ba-
enriched kenopyrochlore by depleting pyrochlore from Ca and Na.
With weathering progression even Ba is eventually leached from
pyrochlore leading to its destruction and consequent formation of
secondary Nb-enriched minerals in the soil (Wall et al., 1999).
6. Carbon and oxygen isotopes
Carbon and oxygen isotopes from DC pockets dolomite (Table 3)
were analyzed to establish a correlation with the pyrochlore chemistry
(Fig. 9). Carbonates were extracted from pockets with a manual
tungsten-carbide drill to avoid interference from different carbonate
generations or contamination with externalsources. Oxygen and carbon
Fig. 8. Ternary plots of Ca, Na and Asite vacancy. Compositional pyrochlore ﬁelds of other deposits are shown for comparison. Data sources as in Fig. 7, plus the Bingo ﬁeld from
Williams et al. (1997).
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isotope data were obtained reacting carbonate samples with 100%
at 72 °C, using a Gas Bench II System connected to a Delta V
Advantage gas-source mass spectrometer at the University of Brasília.
Results are expressed in delta notation, relative to the PDB(carbon) and
SMOW (oxygen) standards.
Data from Table 3 show that gray fresh dolomite (093 G1, 056) has
a mantle-like carbon- and oxygen-isotope signature, interpreted as
magmatic, whereas white brittle dolomite in the same samples has
values than gray calcite and is interpreted as affected
by low temperature H
O-rich ﬂuids probably of meteoric origin. The
same ﬂuids are likely to have altered pyrochlore, leaching Ca and Na
and leaving vacancy while partially replacing them with Ba. This
hypothesis is supported by comparison between calciopyrochlore or
natropyrochlore cores with Ba-enriched kenopyrochlore rims, sug-
gesting ﬂuid interaction. This would be consistent with general
alteration models (Wall et al., 1999).
7. Genetic implications
The dike-like emplacement of nelsonites and its relationship with
DC pockets is an important feature of the Catalão I Nb deposit,
suggesting a magmatic origin for the Nb-ore. Cordeiro et al. (2011)
showed that DC pockets within nelsonites have mantle-like C- and O-
isotope signatures and suggested Rayleigh fractionation and mag-
matic degassing as important processes for the evolution of such
rocks. The authors' results have also shown that metasomatism and
weathering played a role in the carbon and oxygen isotopic variations
of DC carbonates, albeit unrelated to the formation of pyrochlore.
Accordingly, we interpret the occurrence of pyrochlore in nelsonites
and dolomite carbonatite as an igneous process.
The genesis of late-stage phoscorite-series rocks, and therefore of
the Catalão I fresh rock niobium deposit, is still a matter of
controversy. Krasnova et al. (2004) argues in favor of AFC and/or
liquid immiscibility in generating phoscorites. Lee et al. (2004)
described chemical discrepancies between Sokli carbonatites and
related phoscorites as evidence for immiscibility that generated both a
carbonatite and a phoscorite melt. However, the authors point out
that experimental evidence for such process is still lacking.
The best evidence we could ﬁnd for the occurrence of phoscorite
melts is given by Panina and Motorina (2008).Theystudiedmelt
inclusions from the Krestovskii carbonatite complex, in the Maimecha–
Kotui province, Russia, and suggested a carbonatite immiscibility event
that originated alkali-rich phosphate melts. Evidence of Fe–P–Ti-rich
melts exists in carbonatite-unrelated settings such as the andesitic
Antauta subvolcanic complex in Peru (Clark and Kontak, 2004). The
authors describe Fe-rich melt inclusions that are interpreted as derived
from nelsonite-like magma, indicating that such unusual magmas may
indeed occur naturally.
Formation of cumulates is another possible mechanism in the
generation of apatite–magnetite rich rocks. Mitchell (2005) argues
that potential niobium ore rocks in carbonatites do not represent
liquid compositions nor reﬂect the Nb content of the parental magma.
Based on melt inclusion data, Veksler et al. (1998) argue that crystal
fractionation resulted in the formation of calcite carbonatites, which
evolved to forsterite–apatite–magnetite–phlogopite carbonatites
with subordinate phoscorite cumulates and dolomite carbonatites.
According to Downes et al. (2005) the Kola Alkaline Province
phoscorite–carbonatite rocks series are the result of complex
differentiation of an extremely phosphorous and iron enriched
carbonate–silicate melt. They also favor the formation of cumulates
as a mean of generating phoscorites.
Representative analysis of carbon and oxygen isotopes of carbonates from pyrochlore-bearing DC pockets.
Sample 056 056E 93 093G1 178G1 178G2 192G1
Type Carbonatite Carbonatite DC pocket in P2 DC pocket in P2 DC pocket in P2 DC pocket in P2 DC pocket in P2
−5.53 −5.86 −5.38 −5.53 −5.85 −5.16 −6.14
11.80 20.23 19.99 10.42 15.92 11.06 8.59
Fig. 9. Comparison between pyrochlore composition and carbon–oxygen isotope signatures of carbonates within the same pocket. Note that pyrochlore rims from samples 093 and
056 have systematically higher Asite vacancies than corresponding cores. In sample 056, the core is calciopyrochlore and the rim is Ba-enriched kenopyrochlore, while samples 192B
and 178 show only a slight Ba-enrichment and little vacancy. Carbon and oxygen isotopes show that samples with kenopyrochlore rims have wider variation in the δ
while less altered samples preserve the original composition. Stable isotope ﬁelds are from Cordeiro et al. (2011).
119P.F.O. Cordeiro et al. / Ore Geology Reviews 41 (2011) 112–121
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At this stage we are unable to determine which of the described
mechanisms was involved in the formation of the Catalão I fresh rock
niobium deposit. Despite the occurrence of phosphate melts within
carbonatite complexes, crystal accumulation is likely to be involved in
the generation of nelsonites. DC pockets within nelsonite dikes are
sometimes interconnected and low viscosity carbonatite lava could
ﬂow in the open space. Apatite, pyrochlore, magnetite and phlogopite
(i.e. nelsonite) could crystallize in situ until all open spaces were ﬁlled
and ﬂow would stop. Such mechanism could explain both the
occurrence of DC pockets and the related pyrochlore-bearing
Whereas magmatic controls were vital in the formation of the
fresh rock niobium deposit, weathering played an important role in its
enrichment, consequently forming the residual deposit. All rocks
within the Catalão I Complex are easily weathered compared to the
country-rocks (fenites and quartzites). The dome-like structure
prevents erosion and allows the establishment of very thick soil
cover over the alkaline rocks. On average, soil depth is 80 m, but
reaches at least 150 m over phoscorite-series rocks. A similar pattern
occurs at Seblyavr in the Kola Alkaline Province, Russia (Balaganskaya
et al., 2007) where a weathering crust up to 200 m deep is developed
over phoscorite-series rocks in the intrusion core. We believe that the
abundance of fractured easily-weathered carbonatites, either as dikes
cutting early-stage phoscorites or as DC pockets within nelsonites,
contributed to the development of such deep soils and the generation
of the residual deposit.
1) The pipe-like niobium orebodies at Catalão I consist of dike
swarms of late-stage phoscorite-series rocks (nelsonites) that cut
previous phlogopitite and carbonatite.
2) Weathering of such rocks originated the residual deposit, where
leaching of carbonates induced a residual concentration of
pyrochlore and other weathering-resistant phases.
3) Catalão I phocorite-series rocks can be divided into phoscorites (P1),
and the niobiumores apatite nelsonite (P2) and magnetite nelsonite
(P3). The mineralization can be classiﬁed as Nb (+Fe+P) on the
grounds of high modal content of apatite and magnetite. Dolomite
carbonatites (DC) associated with nelsonites are a subordinate
source of pyrochlore but their grades are comparatively low, hardly
above 0.3 wt.% Nb
4) The dominance of calciopyrochlore over other Nb-bearing phases
in the fresh rock deposit and its chemical variability are
independent of lithology. This indicates that pyrochlore formation
chemical conditions were similar in P2, P3, and DC.
5) Substitution of Na–Ca for Ba in the fresh rock pyrochlore structure,
leading to the formation of Ba-enriched kenopyrochlore, and the
signature of the associated carbonates suggest
interaction with hydrothermal ﬂuids. These ﬂuids affected
nelsonites but had no role in the formation of the fresh rock
niobium deposit itself.
6) We could not uniquely constrain the nelsonite formation process in
Catalão I although it is clear that they are genetically related to
carbonatite magmatism. Possible alternatives for the formation and
evolution of these rocks are: (a) crystallization from a phoscorite
magma (Lee et al., 2004, 2006); (b) crystal accumulation from a
carbonatite magma (Veksler et al., 1998); and (c) crystal accumu-
lation from a carbonated-silicate magma (Downes et al., 2005).
We are indebted to Anton Chakhmouradian, Nigel Cook and two
anonymous reviewers for their helpful review of the original
manuscript. This work was supported by the Brazilian Council for
Research and Technological Development (CNPQ), through grants to
the ﬁrst author, JAB and ESRB, as well as by Mineração Catalão and the
Anglo American Brazil Exploration Division. The University of Brasília
is gratefully acknowledged for ﬁeldwork support and access to
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